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Review of major shale-dominated detachment and thrust characteristics inthe diagenetic zone: Part II, rock mechanics and microscopic scale

C.K. Morleya,⁎, C. von Hagkeb, R. Hansberryc, A. Collinsc, W. Kanitpanyacharoend, R. Kingc

a Petroleum Geophysics MSc Program, Department of Geological Sciences, Chiang Mai University, Chiang Mai, Thailandb Lehr-und Forschungsgebiet für Geologie – Endogene Dynamik (GED) Rheinisch-Westfälische Technische Hochschule, Aachen University, Germanyc Centre for Tectonics, Resources and Exploration (TRaX), School of Physical Sciences, University of Adelaide, SA 5005, Australiad Department of Geology, Faculty of Science, Chulalongkorn University, Bangkok, Thailand

A B S T R A C T

Characterising large shale thrust zone behaviour down to the diagenetic-metamorphic boundary is both simpleand complex. The task in critical taper, analogue and numerical models has been successfully approximatedusing simple material parameters. Yet the weakness of shales thrusts on a case-by-case basis show great varia-bility in key factors across a wide range of scales. Understanding of these variables and their relationships todifferent tectonic settings is patchy due to the different objectives of workers studying the various systems (e.g.seismic hazards, hydrocarbon exploration, structural geology research), and the types of available geological andgeophysical data. Early research recognized that megathrusts feature creeping and locked patches, which may bedescribed using the rate and state formalism. In addition to temperature, differential stress and strain rate, keycontrolling factors on thrust weakness are mineralogy, amount of weak phase present, structure localizationmechanisms, grain size and shapes, porosity, permeability and pore fluid pressure. The smectite-illite transitionhas been a focus of seismic hazard research due to its coincidence with the top of the seismogenic zone. Whilethere may be a relationship, other factors such as pore-fluid pressure variations or structure localization are alsoimportant. The critical taper model is a simple means of determining whether basic rock mechanics data (e.g.frictional strength, pore fluid pressure) is appropriate for natural wedges. For example, aseismic basal detach-ments of gravity driven systems (smaller critical taper wedges) appear to retain higher pore fluid pressures alongthe basal detachment than seismogenic basal detachments (higher critical taper wedges). In many continentaland deltaic fold and thrust belts high organic carbon content is a very important factor in shale weakness due to:1) the overall more ductile and well-cleaved nature of the shale when organic content is high, 2) the presence ofhigh overpressures due to maturation of organic material, and 3) metamorphism of carbon to low frictiongraphite. This multifaceted influence of organic carbon content is just one example of the diversity of potentialinfluences on shale thrust zone weakness, which enable shales to be weak despite considerable lateral deposi-tion-related and vertical burial-related changes in composition. Some key variations between different types offold and thrust belt (gravity driven; accretionary prism; Andean/Himalayan type) lie in systematic variations inclay mineralogy, magnitudes and origins of overpressures, seismic versus aseismic detachments, and differentstructural localization mechanisms. Further research is required to explore the viability of such distinctions andtheir impact on structural styles.

1. Introduction

Shale detachments may extend laterally for hundreds or eventhousands of kilometres, but their mechanical properties are defined ona microscopic scale. Consequently, to understand mechanics of claysand shale detachments, a multi-scale approach is required. We describethe basic detachment architecture and main structural elements of shalethrust zones in part 1 of this review (Morley et al., 2017 in press). This

second part addresses our understanding of shale detachment me-chanics mostly from the point of view of laboratory experiments andmicrostructures. Here we keep a very broad perspective on shale de-tachment behaviour in all tectonic settings, and consider the differencesand similarities between the various settings.

Detachments in shale-prone units are a very common feature at awide range of scales, due to both the mechanical properties of clayminerals, and their ability to trap or focus high pore fluid pressures (e.g.

http://dx.doi.org/10.1016/j.earscirev.2017.09.015Received 31 May 2017; Received in revised form 19 September 2017; Accepted 19 September 2017

⁎ Corresponding author.E-mail address: chrissmorley@gmail.com (C.K. Morley).

Earth-Science Reviews 176 (2018) 19–50

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Davis et al., 1983; Dahlen, 1990; Saffer and Tobin, 2011; Haines et al.,2013; Tesei et al., 2015). Classically, shales have been defined as rockscontaining more than ~40% clay minerals (Shaw and Weaver, 1965).We adopt this concept for shale detachments, and for brevity wheremudrocks or phyllosilicates are dominant we will refer to these as shaledetachments or thrusts in this review. However, we note that con-siderable lithological variations exist within such fault zones, not onlydue to heterogeneity of sedimentation or diagenetic processes but alsobecause of incorporation of surrounding rock fragments into the faultzone. Moreover, fault geometry and strength may be controlled byavailability of small percentages of clay minerals; for instance,< 3% ofweak shales may result in structures similar to those observed in shale-rich detachments dominated by cataclastic processes (Smeraglia et al.,2017a, 2017b). Consequently, we discuss features observed on suchshale-dominated faults that allow for insights on shale detachments. Asreviewed by Morley et al. (2011) shale detachments are the key sliphorizons in many structural settings: fold and thrust belts, accretionary

prisms, gravity driven systems ranging from major deltas to masstransport complexes and hybrid gravity-lithospheric-stress driven sys-tems. Although often we consider thrusts in shales to be simple through-going features in classic flat and ramp thin-skinned structures, we alsoneed to recognize that inversion of earlier normal faults, and mixedthick-and thin-skinned tectonics, can be important variants on thelarge-scale fault zone architecture, that impose important large-scaleheterogeneities on fault systems (e.g. Buchanan and Warburton, 1996;Tozer et al., 2002).

A better comprehension of detachment characteristics may impacteconomic decisions such as the hazards of drilling through detachmentzones to reach hydrocarbon traps, and exploitation of unconventionalshale hydrocarbon resources. Weak thrusts are associated with some ofthe most dramatic earthquakes reported (e.g. the Mw 7.6 1999 Chi-Chiearthquake, Taiwan (Yue et al., 2005), the Mw 7.3 1886 Charleston,South Carolina earthquake, Appalachians (Seeber and Armbruster,1981), the Mw 8.8 2010 Maule Earthquake (Moreno et al., 2010), and

1 cm

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Accretionary Prism

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Thrusts and lateral rampsFrontal thrust

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Fig. 1. External décollement of the Barbados accretionary prism, structural characteristics from seismic to microscopic scale, based on data from two cored sites (948, 949) of OceanDrilling Program Leg 156 (redrawn and modified from Maltman et al., 1997 and Labaume et al., 1997). A) Sketch across external accretionary prism, based on seismic sections. B)Detailed view of detachment zones as being composed of several high displacement strands, utilizing core-scale and microstructural data. C) Idealized sketch of the microstructures thatcontribute to the macroscopic appearance of scaly clays in cores, FN = fracture network, SF = S-fabric, C-S refers to the S-C deformation bands. D) Kinematic model for S-C banddevelopment in scaly-fabric zones. C = shear surface subparallel to bulk shear, C′= shear surfaces with extensional sense of motion with respect to bulk shear, S = foliation, (R) and (X)follow Riedel terminology.

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the 2011 Mw 9.0 Tohoku-Oki earthquake, Japan (Chester et al., 2013).On land and in marine conditions these thrusts, often affect highlypopulated areas, and form major geohazards. Due to their sealingproperties, they are targeted as nuclear waste disposal sites in somecountries (e.g. Madsen, 1998; Laurich et al., 2014). Consequently, adetailed knowledge of why shale thrusts and detachments are weak isimportant for addressing practical as well as fundamental scientific is-sues. Regarding the latter, these issues include the impact of the de-tachment on thrust sheet deformation style, the formation and dis-tribution of high strain zones within shales, fault zone properties toimprove numerical models of fold and thrust belt deformation, and thedistribution and importance of pore-pressure in thrust zones.

In this review, we first introduce the structures observed in majorshale thrust zones, focussing on observations at the microscopic scale(Section 2). Secondly we discuss the basics of rate-and-state friction, asshale thrust and detachment behaviour has been partly described usingfault frictional laws based on laboratory experiments or field studies(Marone, 1998; Haines et al., 2013; Tesei et al., 2015; Carpenter et al.,2016) (Section 3). In the fourth section of this review we focus on thethree main factors that can be considered as contributing to theweakness of shale thrust zones: 1) overpressured fluids, 2) bulk mi-neralogy characteristics, and 3) the efficiency of structure localizationduring deformation. These mechanisms can be complimentary and arenot mutually exclusive. We then discuss changes of the mechanismswith depth. In Section 5 important advances in our understanding offault zone mechanics in seismogenic regions as a consequence of dril-ling recently active fault zones is reviewed, before ending with a dis-cussion of how our understanding of shale zone weakness is applied tocritical taper models of fold and thrust belts.

2. Deformation processes and fault structures

2.1. Introduction

Major shale involved thrusts can cover areas up to the order of100,000's kms, and form the basal detachments to mass transportcomplexes, gravity-driven deltaic systems, the external zones of oro-genic belts, and accretionary prisms (see reviews in Morley et al., 2011,2017 in press). Given the diversity of these environments, their varia-tions in stratigraphy, mudrock composition, temperature/depth, degreeof lithification, structural position, pore fluid pressures, seismicity, fluidcomposition and origin no single model for a shale prone thrust zone is

applicable, although there are many features in common (see part 1,Morley et al., 2017 in press).

As a starting point for understanding many of the common char-acteristics of major shale prone thrust zones, the studies of the Barbadosaccretionary prism are significant (e.g. Maltman et al., 1997; Labaumeet al., 1997; Fig. 1). Unlithified cores in clay rich sediments across theactive and external basal detachment to the Barbados subduction zonewere recovered by the Ocean Drilling Program (e.g. Maltman et al.,1997; Labaume et al., 1997; boreholes 948, 949, Fig. 1). Around 500 mbelow the sea floor, the detachment zone is ~40 m thick, and hetero-geneous, comprising multiple alternating zones of fracture networks(cross-cutting, fine very low displacement planar cracks spaced5–10 mm apart), disrupted strata (commonly extended and boudi-naged), and sub-horizontal zones of scaly fabric (Labaume et al., 1997;Fig. 1B, C). Ninety-eight zones of scaly fabric, on average about 3 cmthick, occur across the ~40 m interval. Hence the high strain part of thedécollement is focussed on only about 8.8% of the décollement zone(Maltman et al., 1997).

Investigations from core-scale (or outcrop scale) to scanning elec-tron microscope (SEM) observations demonstrated the fractal-likenature of the scaly fabric (Labaume et al., 1997). At many differentscales sheared clay is wrapped round less deformed lenses or particles.The scaly clay zones contain a pervasive flattening fabric, which iscommon to other shear zones (e.g. Schultz, 1988; Housen et al., 1996;Housen and Kanamatsu, 2003; Yang et al., 2013). Such flattening canbe quantified from either strained objects within the shear zones (e.g.Fig. 2A) or from the anisotropy of magnetic susceptibility (AMS) of theshales, (e.g. Fig. 2B). Superimposed on the scaly fabric are distributedslip surfaces, and together the flattening fabric and the slip surfacesform S-C structures (Labaume et al., 1997; Fig. 1C, D). Typically, inboth poorly and well lithified mudrock thrust zones S-C fabrics arecommon, with shear strains imposed on a spaced foliation, or pervasivecleavage (Fig. 1D). The scaly fabrics in the Barbados décollement wereinterpreted by Labaume et al. (1997) to result from porosity collapse asclay particles rotated and aligned. Strain localization produces shearbands within the scaly fabric (Fig. 1D), but since individual shear zoneswill tend to lock (Hicher et al., 1994), the shear zone arrays are ex-pected to broaden with time (Maltman, 1997; Moore and Byrne, 1987),or propagate into the less deformed sediment lenses between shearzones. However, this broadening process does not proceed indefinitely,and ultimately displacement tends to become focussed on a few narrowprincipal displacement zones that are often in the order of centimetres

A B

Fig. 2. Examples of strain measurement from shale-prone thrust zones. A) Flinn diagram of competent boudinaged lensoids in thinly bedded Ordovician rocks, that are horses inf thePulaski thrust sheet. 73% of the measurements lie in the flattened ellipsoid field of the diagram. e1, e2 and e3 are the axes of principal extensions; a = ratio of elongations in e2 and e3directions; k = slope of line defining strain types. Redrawn from Schultz (1988). B) Flinn-type diagram illustrating the dominantly oblate shape of anisotropy of magnetic susceptibility(AMS) ellipsoids for wedge sediments (open circles, and underthrust sediments (solid circles). Across the décollement maximum AMS axes have a strong preferred NW-SE direction, belowthe décollement the maximum axes is sub-horizontal. AMS ellipsoids for the underthrust sediments are generally less oblate that those overlying the detachment. The magnetic fabricsmark an abrupt change in strain across the décollement (redrawn from Yang et al., 2013).

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wide.The Barbados example introduces a number of important themes

either mentioned above, or elsewhere in this review (e.g. how shearzones develop, the impact of porosity evolution on fluids and strainlocalization, cycles in pore fluid pressure and migration, the fractalnature of clay fabrics from core/outcrop to nano-scale). However, adetachment in a major orogenic belt such as the Osen-Røa thrust,Norway, can have more advanced state of compaction and diagenesis atthe onset of deformation than the Barbados example (see detailed de-scription in part 1, Morley et al., 2017 in press). The Osen-Røa de-tachment was probably formed at a depth of 2.5 km or less below thesurface (Bruton et al., 2010). The ~50 m thick Cambrian Alum ShaleFormation that forms the detachment was already compacted prior todeformation. The highest strain zone does not lie well within the shalesequence, but instead is localized near the base of the shales at themechanical contrast where they rest on crystalline Precambrian base-ment. The Alum Shale Formation displays a relatively low degree ofvein development, suggesting fluid flow through the detachment waslimited, possibly due to limited fluid source. Downwards-flatteningpervasive cleavage and closely spaced listric, and pseudo-duplex shearsurfaces are developed, short wavelength folding and pressure solutionseams are widespread, S-C fabrics are present, but not scaly clay tex-tures. The contrasting detachment zone styles in Norway and Barbadosare highlighted to emphasize that in the following sections the widevariety of factors described that impact shale-prone fault zone evolutionand behaviour are commonly present, but are not necessarily ubiqui-tous.

Structure localization to form weak shale detachments occurs al-ready before active shortening commences. Here we first discuss thecharacteristics of largely undeformed clays, and the earliest stages ofdeformation in clays and early diagenetic processes (Section 2.2). Wethen discuss formation of micro-fractures and their evolution towardsfully connected detachments (Section 2.3). This is followed by a de-scription of some well-known shale detachments (Section 2.4) beforeending this section with some general characteristics of observedstructures and processes (Section 2.5).

2.2. Formation of shale detachments – early structure localization in clays

Even before shale horizons become active detachments their pro-toliths (clay minerals and clay-rich rocks) feature macroscopic andmicroscopic heterogeneities, which are responsible for strain localiza-tion and are thus important for understanding their mechanic char-acteristics. It is possible to image clays and shales at nm resolutionusing Broad or Focused Ion Beam (BIB or FIB) polishing techniques andSEM microscopy in order to visualize heterogeneities such as pre-ferential alignment of minerals, pore space or mineralogical variations.Using these imaging techniques it has been shown that mineralogicaland grain size variations control microstructures and pore spaces oflargely undeformed clays (e.g. Hemes et al., 2013). A combination ofSEM analysis of BIB polished sections and Hg-injection porosimetry onsamples from fine- and coarse-grained parts of the Oligocene Boom ClayFormation, Belgium, has shown that different mineral phases havecharacteristic pore morphologies, pore size distributions, and differentinternal porosities (Hemes et al., 2013; Desbois et al., 2014; Klaveret al., 2015). Pore morphology is important for permeability, as well ascapillary processes and the sealing capacity of the shales (e.g. Desboiset al., 2009). Power-law distribution of pore sizes over six orders ofmagnitudes at the nm to μm scales support that microstructural char-acterization of the pore space may be used to infer larger-scale rockporosity and larger-scale structures (Hemes et al., 2015, 2016). Addi-tional heterogeneity in undeformed clays is introduced by preferentialalignment of the (001) lattice planes of sheet-like phyllosilicates duringsedimentation, compaction, and diagenesis (e.g. Ho et al., 1999; Aplinet al., 2006; Wenk et al., 2008). Similarly, early diagenetic re-crystallization processes lead to reorientation of clay platelets,

increased elastic and seismic anisotropy, and a loss of porosity andpermeability (e.g. Baker et al., 1993; Hornby et al., 1994; Sayers, 1994;Johansen et al., 2004; Draege et al., 2006; Bachrach, 2011). Poremorphology as well as preferred crystal orientations make clay horizonspreferred zones of strain localization.

As reviewed by Bennett et al. (1991) microfabric signatures in claysare determined by physicochemical, bioorganic, and burial diageneticprocesses. In this framework, mechanisms are defined as the specificenergy sources that drive microfabric development. Mechanisms are forinstance thermomechanical (e.g. modification of the fabric due to la-minar flow associated with temperature differences), electrochemical(e.g. chemical bonding of particles or electrostatic interactions that aresensitive to pH value), or biophysical (e.g. particle clustering due toorganisms) (Bennett et al., 1991). All these processes and associatedmechanisms may be acting at the same time but their relative im-portance for fabric evolution is variable (for instance bioorganic pro-cesses are most important in organic-rich sediments). However, as theydefine microstructure, they determine the physical and mechanicalproperties of clays.

Multiple studies, particularly in soil mechanics have addressed themechanic behaviour of different clays, including intact clay, remoldedclay, and clay treated with lime, cement or sand (e.g. Abdullah et al.,1997; Deng et al., 2012; Mesri and Vardhanabhuti, 2009; Walker andRaymond, 1968). These studies show that secondary deformation infine-grained rocks (i.e. volume change due to fabric evolution and soil-water interaction) is linearly related to compression characteristics (i.e.the relationship of void ratio or vertical strain to an increasing com-pressive stress, e.g. Raymond, 1966, Gregory et al., 2006). Thus, duringearly stages of deformation, porosity and porosity reduction exertsstrong control on formation of heterogeneities and later structure lo-calization. Porosity is influenced by compression and in claystones,different types of pore morphology can be characterized, in particularlarge jagged pores in strain shadows of clastic grains, high aspect ratiopores between similarly oriented phyllosilicate grains and crescent-shaped pores in saddle reefs of folded phyllosilicates (Desbois et al.,2009). Typically, porosity is enhanced in areas around irregular sur-faces of non-clay minerals within the clay matrix (e.g. Keller et al.,2013), and thus the spatial density of these non-clay minerals affectsclay mechanics. Unbound water and other fluids such as oil and gasoccupy pore spaces and exert a control on rock strength throughoverpressure, which is a function of both porosity and permeability. Wediscuss the influence of pore pressure, on detachment weakness inSection 4.1. Apart from porosity and preferential orientation of flakyminerals during deposition and early diagenesis, the amount of swellingclay minerals controls strength and deformation style: mudrocks be-come weaker as the percentage of swelling clays increases (Olgaardet al., 1997). Furthermore, in low effective pressure environmentslower percentages of swelling clays support more brittle deformation,whereas higher percentages support more ductile deformation (Olgaardet al., 1997). Generally, soft clays and mudstones as well as shalethrusts often show brittle-ductile transitional behaviour (e.g.Dehandschutter et al., 2005; Ingram and Urai, 1999; Takizawa andOgawa, 1999), which we will discuss in the following section.

2.3. Fracturing and formation of a through-going detachment

During formation of a through-going detachment horizon, bothductile shear bands as well as brittle fracturing occurs. In Boom Clay,microstructural analysis suggests ductile deformation occurs in zones oflow strain, whereas brittle structures dominate in high strain zones thatare also characterized by higher strain rates (Dehandschutter et al.,2005). Fracture formation in rocks and their evolution have been stu-died for almost a century (Griffith, 1921), but the inherent complexityis not yet fully understood (e.g. Virgo et al., 2014). Initially isolatedcracks develop that grow until they are connected and form a through-going fault system. At the microscopic scale, micro-cracks can be

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observed, which are planar features in isotropic material, but theycommonly form along anisotropies such as grain boundaries. The de-velopment of a fracture in an elastic medium as described by Griffith(1921) is given by:

=σ 2Eγπac (1)

where σc is the critical stress that must be overcome at the tip of ani-sotropy, E is the Young's modulus, γ is the surface energy necessary tocreate a new surface, and a is the initial length of the anisotropy(Ougier-Simonin et al., 2016). To predict stress at the tip of a fracture,the stress intensity factor has been developed in fracture mechanics,which has been formulated for homogeneous, linear elastic materials.However, in geological applications, fracture propagation through dif-ferent mediums is complex, and their direction is influenced by het-erogeneities such as fluid inclusions, rock fragments, or adjacent frac-tures that may grow simultaneously. Additionally, for shear fractures ithas been shown that failure is partly time-dependent due to pore waterdiffusion to fracture tip, viscoelastic deformation of the material, andon longer timescales on changes of mechanical properties, for exampledue to weathering or hydrothermal alteration (Palmer and Rice, 1973).Due to this time dependence and high non-uniformity of the stress fieldat all scales, mathematical formulation of fracture propagation is ex-tremely difficult, and hence only limited studies exist (e.g. Patrício andMattheij, 2010; Virgo et al., 2014). However, several mechanisms havebeen recognized as influencing the degree of microfracture develop-ment, in particular mineral and organic carbon content, layering,overpressure, thermal shrinkage and dehydration as well as tectonicfactors (see Ougier-Simonin et al., 2016 for a review). These micro-fractures will grow into larger features and ongoing slip along the faultwill produce a zone of fractured rocks with variable amounts of sur-rounding material being incorporated into the slip horizons. Strainaccumulation will result in development of increasing amounts of finegrained material. These typically lens-shaped microlithons are sur-rounded by zones of ductile shearing along well-aligned particles(Fig. 3). Generally, faults evolve from tip-process zone microfracturingto segment linkage by relay branching and lens formation, to splay-faulting and re-connection forming side-wall rip-outs (Wibberley et al.,2008). Thin shear zones are already present in clay rich sediments atthe early stages of deformation (e.g. Labaume et al., 1997;Dehandschutter et al., 2005; Laurich et al., 2017). The development ofbrittle and ductile structures in clay shear zones at the earliest stages ofdeformation depends on an interplay between compaction and con-solidation forces in combination with tectonic forcing. This causes thesuccessive development of compacting shear bands, hybrid shear frac-tures and mode-1 fractures (Labaume et al., 1997; Dehandschutteret al., 2005). Grain size reduction, particle alignment, pore collapse andcrystal plasticity in clays in combination with pressure solution reac-tions will result in formation of polished shear fractures (slickensides),that interconnect during progressive deformation (see Laurich, 2015and Laurich et al., 2017 for recent reviews). These surfaces accom-modate most of the strain in shale thrust zones. Different stages of crackgrowth from isolated to connected may be associated with fault creep(see discussion on rate-and-state laws below).

In the Barbados Accretionary Prism, incipient structures associatedwith the basal detachment have been identified in a high porosityradiolarian mudstone some 5 km from the active deformation front(Behrman et al., 1988). Structures evolved from incipient scaly clayswith a flattening fabric, to scaly fabrics that act as discrete microshearsand with both mud and calcite filled veins. The Barbados example issimilar to findings from the Opalinus Clay, Switzerland, where scalyclay first increases due to an increase in thin shear zone intensity andlater progressively evolves to clay gouge in restraining bends of theshear zones (Laurich et al., 2014). Calcite veins may form at releasingbends (Laurich et al., 2014). In the Millaris Fault, Pyrenees,

geochemical mass balancing and microstructural analysis shows duringearly stages of deformation up to 50% of calcite can be lost from thehost rock, and phyllosilicates are enriched and reoriented (Lacroixet al., 2015). These findings show that already during early stages ofdeformation the rocks are severely altered and become weaker.

2.4. Natural examples of structures in shale detachments

Recently, the development and mechanisms of crystallographicpreferred orientation (CPO) or fabric can be effectively quantified ef-fectively though high-energy synchrotron X-ray diffraction (SXRD) andsynchrotron X-ray microtomography (SXMT) techniques. For example,Callovo-Oxfordian shale and Opalinus clay samples retrieved fromboreholes in the radioactive waste disposal sites in France andSwitzerland were characterized for fabrics and elastic properties (Wenket al., 2008). The Callovo-Oxfordian shale shows a random orientationof calcite grains and a relatively weak orientation of clay minerals.Conversely, Opalinus clay samples from Benken and Mont Terri showstrong alignment of calcite and clay minerals e.g. illite-smectite, kao-linite, and chlorite. This intrinsic contribution to elastic anisotropy isconsistent with macroscopic physical properties where elastic aniso-tropy is caused both by the orientation distribution of crystallites andhigh-aspect-ratio pores. In addition, Posidonia Shale collected in theHils Syncline from Germany was characterized for the influence of claycontent, burial depth, and thermal history (Kanitpanyacharoen, 2012).The samples used in this study experienced different local temperaturesduring burial and uplifting, as established by the maturity of kerogen,but their constituent clay minerals show similar degrees of fabrics.Results from SXRD measurements imply that the difference in localthermal history, which significantly affects the maturity of kerogen, atmost marginally influences fabrics of clays, as the alignment of clayswas established early in the history.

The 3D microstructures in Kimmeridge Shale (North Sea, UK) andBarnett Shale (Gulf of Mexico, USA) have been quantified to a resolu-tion of ~1 μm using the synchrotron X-ray microtomography (SXMT)technique (Kanitpanyacharoen et al., 2013). Measurements were doneat different synchrotron facilities to characterize 3D microstructures,explore resolution limitations and develop satisfactory procedures fordata quantification of fine-grained rocks. The 3D segmentation imagesshow that low density features, including pores, fractures, and kerogen,are mostly anisotropic and oriented parallel to the bedding plane. Smallpores are generally dispersed, whereas some large fractures andkerogen patches have irregular shapes and remain aligned horizontally.A comparison of fabrics in phyllosilicate-rich rocks such as SAFOD faultgouges, Kimmeridge shale, and schist from the Alps has also been in-vestigated in Wenk et al. (2010). The results demonstrate that phyllo-silicates show large fabric variations in various environments, wheredifferent mechanisms produce the rock microfabrics. Fault gouge fab-rics are quite weak and asymmetric due to heterogeneous deformationwith randomization, as well as dissolution-precipitation reactions.Shale fabrics have a stronger degree of fabric orientation and this is dueto sedimentation and compaction (Fig. 3). The strongest fabrics wereobserved in metamorphic schists and developed by deformation as wellas recrystallization in a stress field. Laboratory experiments testing therole of fabric show if weak minerals such as talc are critically alignedand form a through-going layer, only ~4% of the weak mineral arerequired for weakening of the fault zone (Niemeijer et al., 2010).

Laurich et al. (2014, 2017) analysed microstructures and deforma-tion mechanisms of the Opalinus Clay in the Main Fault (Fig. 4) exposedin the Mont Terri Rock Laboratory in Switzerland. BIB polishing of thesamples revealed microstructures at the nano-scale. Microlithons in theshear zone show power law distribution of size and are surrounded bydistributed ductile shear zones. This extreme strain partitioning pos-sibly is related to intense softening during deformation, caused byporosity collapse, particle alignment, and formation of nanoparticles(Laurich et al., 2017). In the fault zone, slickensides associated with

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μm-thin shear zones, clay gouge, as well as calcite veins and scaly clayaggregates could be distinguished (Laurich et al., 2017; Fig. 4). Basedon their microstructural observations, the authors argue scaly clayevolved by progressive development of an anastomosing network ofmicro-shear zones (Laurich et al., 2014). The highly variable shearstrain is continuous-discontinuous in such shear zones and affectsrupture nucleation and seismic style (Fagereng and Sibson, 2010).Progressive formation of weak fault rocks was discussed by Maltmanet al. (1997) for the Barbados accretionary prism, and has also beenreported from the San Andreas Fault (Schleicher et al., 2010). Eventhough the latter fault occurs in a strike-slip environment, the generalevolution of fault rock is comparable to thrust fault settings, and theresulting microstructures are similar. For instance, Schleicher et al.(2010) show development of through-going shear planes together withdevelopment of nano-clay coatings along the fault surfaces associatedwith decrease of fault frictional strength, similar to the anastomisingmicro-shear zones discussed above. In the fold-thrust belt of the

Pyrenees, microstructural analysis of metre-thick shear zones withinclay-rich sediments shows strong foliation due to preferentially orientedphyllosilicates, pervasive pressure solution, and multiple generations ofshear and extension, as witnessed by deformed calcite, chlorite, andquartz bearing veins (Lacroix et al., 2011). Shallow crustal thrust faultsin the Apennines show similar microstructures, in particular S-CC′tectonites, polished principal slip zones with cataclasites and ultra-cataclasites and foliated (YPR) cataclasites (Tesei et al., 2013).

2.5. Deformation processes and structures occurring in shale-rich faultzones

Grain-scale microstructures form as a response to deformationprocesses and thus allow identification of these processes (Passchierand Trouw, 2005). In this section we consider the types of micro-structures that develop as shale thrusts evolve within the diageneticzone. The most dominant processes in shale detachments are cataclasis

Fig. 3. Typical cataclasite in thin section. The samples weretaken from the shale dominated detachment between car-bonate sequences in the Khao Kwang Fold-and-Thrust Belt,Thailand. A: S-C fabric in shale dominated lithology. Clastssurrounded by thin shear zones are reworked calcite veinfragments. Pressure solution dominates the fabric. B: cata-clasite close to a shear zone (right side). Shear zone featuresvery fine grained clasts and strong foliation within clay-dominated (dark) regions. Further away from the shearzone (going left), clast size increases, few veins roughlyperpendicular to foliation; most large clasts are reworkedvein material, rarely quartz grains are present. Left half,new veins form in at least three directions, partly followingshear zone foliation. During progressive shearing, theseveins are chopped up and form clasts. C: shear surface atfootwall contact of the shale dominated shear zone. Thefault surface is underlain by fractured clast, which mightindicate seismic rupture (e.g. Smeraglia et al., 2017b). D:overview of cataclastic structure. Central part shows ana-stomosing deformation bands, where most strain has ac-cumulated. Surrounding the zone of high strain are parti-cles of various sizes that have been broken down duringdeformation. Veins crosscut the shear zone and are partlyreworked during later deformation phases, as witnessed byvein clasts in the shear zone.

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(particularly when shale-rich fault zones contain granular materialssuch as quartz or calcite), dissolution-precipitation reactions, frictional-sliding along phyllosilicate-foliae, interlayer delamination, dislocations,nanofaults, clay-rich layer bending, interlayer fissuring as well as phasetransformations (Passchier and Trouw, 2005). We discuss phase trans-formations when reviewing depth-dependent changes in shale thrustzone behaviour. The different deformation processes at play result incomplex deformation structures observed in shale-dominated faultzones and tectonic mélange, Typical structures observed in shale de-tachments at the meso- to macroscopic scale are thin shear zones, scalyclay, cleavage duplexes, S-CC′ shears, duplexes, principal displacementzones, folds, veins, footwall thickening, short-wavelength domes andsynclines and mushwads. We discussed these structures part 1 of thisreview (Morley et al., 2017 in press), here we focus primarily on mi-croscopic structures.

Cataclasis is a process involving grain size reduction and so, in thecontext of shale-involved thrusts, will predominantly occur at the baseof relatively strong lithologies (e.g. limestone, sandstone, crystallinerocks) being moved over a shale-rich footwall (a classic example is thePulaski Thrust, Appalachians see review in Schultz, 1988; also see part1 of this review, Morley et al., 2017 in press), or where layers or lensesof such lithologies are interbedded within a shale-prone thrust zone(e.g. Byrne, 1984; Schultz, 1988). With progressive deformation com-petent lenses within shales tend to be extensively affected by cataclasis,fractures, and flattening (Schultz, 1988; Fig. 2A). Cataclasis can alsooccur within shales where veins (typically quartz or calcite) that tra-verse shale shear zones become dismembered and in some cases re-duced to grain-size objects embedded within the shear zone (Fig. 3; seepart 1, Morley et al., 2017, in press). More brittle, cemented mudrocksare likely to deform by cataclasis, and secondary crystal plastic me-chanisms, which can be observed down to the nanometer scale (Desboiset al., 2017). Fig. 5 shows a conceptual figure of how microstructuresdeveloped in triaxially deformed, relatively brittle Callovian-OxfordianClay (Desbois et al., 2017): intergranular cracking commences in bothclay mineral and non-clay minerals, progressively fragmentation be-comes more intense within a defined shear zone (Fig. 5A–C), shearingdrives cataclastic flow and opening of pore spaces (Fig. 5D), until thedamage zone becomes sealed due to shear, pore collapse and formationof a clay gouge (Fig. 5E).

Pressure solution creep is a dominant mechanism for cleavage for-mation (G. H. Davis and Reynolds, 1996), and it has been shown in

many orogens to contribute significantly to permanent shorteningstrain (e.g. Duebendorfer et al., 1998), and is an important process oncreeping megathrusts (Fagereng and den Hartog, 2017). A key in-gredient for pressure solution creep is heterogeneity that can resultfrom distributed deformation, variable temperature, mineralogy, oramount of water. Several different flow laws of deformation by pressuresolution exist (Rutter and Elliott, 1976). The key parameter to distin-guish between these different creep laws is the relationship betweenstrain rate and mass transfer distance (Fig. 6). Pressure solution creepexperiments show creep rate is proportional to stress (e.g. Dysthe et al.,2003; Spiers et al., 1990), and from pressure solution creep features(e.g. stylolites) inferences can be made about stress orientation, stressmagnitude, the solubility of the solid in solution and mass transferdistance. However, quantifying the amount of pressure solution creepmay be difficult, particularly for open systems (Fig. 7).

Typically in shale thrust zones extensive evidence for pressure so-lution can be seen in hand specimen, particularly where limestonelayers are present, or calcite veins cross-cut shales (as discussed in Part1, Morley et al., 2017 in press). In hand specimen fairly widely spaced(milimetric to centimetric spacing) stylolitic pressure solution seamsare the most obvious features. While in thin section more planar, andanastomosing, closely spaced seams are present in high-strain zones(Fig. 7), and commonly it becomes apparent that foliation seen in handspecimen or at low resolution in thin section (Fig. 7A) is related topressure solution seams at higher resolution (Fig. 7B). Often pressuresolution appears to be occurring in a closed system, with thin vein ar-rays at the microscopic scale, and larger veins macroscopically accu-mulating the dissolved minerals from the adjacent pressure solutioncleavage (Fig. 7). For a discussion on the strain associated with cleavageformation see Part 1 (Morley et al., 2017 in press).

A full review of the dynamics of pressure solution creep is beyondthe scope of this study; readers interested in more details on pressuresolution creep are referred to the excellent review by Gratier et al.(2013). The importance of pressure solution processes in fault rocks isdemonstrated in fault zones in carbonate-rich sequences in the Apen-nines described by Viti et al. (2014). Their multiscale study involvingoutcrop to microstructural and TEM analysis of the fault rocks showsthat pressure solution was responsible for substantial changes in bulk-rock composition and mechanic properties of the rock. Continuous clay-rich horizons with low friction were formed as a consequence of pres-sure solution, and are enriched with authigenic smectitic clays.

3 cm

Scaly clay

Highly polished slickenside

with nano-size clay

particles

Calcite and celestite veins

and calcite enriched patches

thin, non-porous

shear zone

(

Shortening is accommodated predominantly by frictional sliding ratherthan cataclastic processes. Typical microstructures are stylolitic seamsarranged in S-CC′ structures, where stylolite thickness decreases withincreasing phyllosilicate content (Viti et al., 2014).

Processes such as frictional sliding or interlayer-delamination pos-sibly in combination with pressure solution will eventually lead toformation of clay gouge. Clay gouge is an incohesive fault rock withvery weak to random fabric and< 30% of visible rock fragments(Sibson, 1977). Numerous studies have focussed on the description andclassification of clay gouge texture, both in experiment and outcrop(e.g. Cladouhos, 1999b; Haines et al., 2013; Logan et al., 1992; Rutteret al., 1986; Snoke et al., 1998). Structures observed in experiment andin outcrop or at microscopic scale are comparable, and most prominentfeatures are preferred alignment of clay minerals (P-foliation), Y- andRiedel shears; classification may be based on particle size, angularityand orientation of rock fragments, or presence or absence, inter-connectivity and shape of synthetic and antithetic (R, R′) Riedel shearsas well as P- and Y-structures. Y-shears are interpreted to accommodatemost displacement (Logan and Rauenzahn, 1987). Clay gouge mayfeature a different mineralogy as compared to the surrounding rock dueto clay mineral transformations at temperatures between 50 and 180 °C(Buatier et al., 2012; Haines and van der Pluijm, 2012). Brittle de-formation processes are important for clay gouge formation, howeverthe term is generally used without mechanistic connotations (Vrolijkand van der Pluijm, 1999). Clay gouge may originate from clay smear,where clay present in the wall rock is incorporated into a fault zone byinjection, abrasion, or shearing (van der Zee and Urai, 2005), or me-chanical alteration of wall rock that may result in neoformation of clay

minerals (Cladouhos, 1999a; Holland et al., 2006; Laurich, 2015). Thestrength of clay gouge is a function of shear strain (Haines et al., 2013;Holland et al., 2006; Mizoguchi et al., 2009). In friction experiments onillite-rich fault gouge it has been shown that frictional behaviourchanges with temperature: whereas deformation at low and high tem-peratures (i.e. 150–250 °C and 400–500 °C respectively) is character-ized by velocity strengthening, at intermediate temperatures(250–400 °C) velocity-weakening is observed (den Hartog and Spiers,2014; den Hartog et al., 2012). This correlates with microstructuralchanges: at low temperatures deformation is predominantly controlledby granular flow, medium and high temperatures are characterized bystress corrosion cracking of quartz grains plus possible pressure solution(den Hartog et al., 2012). Laboratory experiments on different claygouges show permeability decreases dramatically with shearing, butchlorite gouge is more permeable after shearing as compared to mon-tmorillionite and illite gouge (Ikari et al., 2009a, 2009b). However,relatively little is known about crystallographic and shape preferredorientations (CPO or fabrics and SPO) of phyllosilicates in shales, whichwe will discuss in the following section.

3. Fault strength and the rate and state formalism

3.1. Introduction

In this review we regard gravity driven systems as being aseismic, orat least absent of moderate and higher earthquakes (> Mw 5), anddiscuss extensive evidence for creep mechanisms operating alongshallow thrusts in orogenic belts. For accretionary prisms it has been

Quartz

Organic matter

Incipient

shearingShearing

Foliation of clay aggregates

opened space

Mica

Calciteσ1

30°- 45° 30°- 45°

30°- 45°

30°- 45°

30°- 45°

A B

C D

E

.. .

.

..

. .

.

Active fracture Inactive fracture

Fig. 5. Conceptual model of the progressive microstructuredevelopment in cemented Callovo-Oxfordian Clay duringtriaxial deformation. A, intergranular microcracking in-itiated at the boundaries of non-clay and clay minerals. B,fragmentation of original fabric by transgranular and in-tragranular microfracturing of non-clay minerals. C,Incipient shearing enhanced by phyllosilicate plasticity atmicrofracture boundaries initiates cataclastic flow of ori-ginal fabric's fragments. D, Ongoing shearing drives cata-clastic flow, opens up pore spaces. E, the damage zonebecomes sealed due to shear and pore collapse, forms claygouge. Redrawn from Desbois et al. (2017).

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inferred for many years that the toe regions were too weak for coseismicrupture, with the base of the aseismic zone being highly variable,ranging between about 5 and 15 km (e.g. Byrne et al., 1988; Hyndmanet al., 1997; Tobin et al., 2009). Given that this review is focused onshale-prone thrusts in the upper crust, a detailed discussion of seismicslip as one of the elements in upper crust displacement on weak de-tachments may seem redundant. However, as discussed in Section 5

there are two well documented, and drilled recently active seismogeniczones in shallow thrusts within shales (Chi-Chi earthquake, Taiwan;Tohoku-Oki earthquake, Nankai Trough, Japan) that indicate earth-quake ruptures can initiate deeper and propagate into shallower, weakzones (see detailed discussion in Section 5). According to Faulkner et al.(2011) this behaviour can be explained by rapid thermal pressurizationof pore fluids within the clay gouge, which make propagation ofearthquakes energetically favourable. This view of shallow fluid-rich,clay gouge forming the shallow fault zones maybe applicable to ac-cretionary prisms, but in continental fold and thrust belts the shalesforming the detachment can be much older than the orogenic belt andmuch better lithified, and therefore will behave differently from ac-cretionary prisms. Indeed Hubbard et al. (2015) investigated bothsubduction zones, and other orogenic fold and thrust belts and foundthat shallow detachments are capable of large coseismic slip events thateven rupture in the toe regions. While some events described inHubbard et al. (2015) are related to ruptures that initiated relativelydeep, and propagated updip, other events actually ruptured in thefrontal shallow part of the wedge (also see Li et al., 2016). Hubbardet al. (2015) conclude that this seismic slip behaviour is “common tomany (perhaps most) accretionary wedges and fold-and-thrust belts”.Consequently we consider it necessary in this review to consider seismicruptures alongside creep and slow slip behaviour for shale thrusts anddetachments, and also in Section 5 (seismicity and shale thrust zones) toconsider what evidence might exist for seismic slip in outcrops of an-cient fault zones.

3.2. Rate and state formalism and the seismic cycle

Fault strength in the upper crust is largely a function of effectivenormal stress and the coefficient of friction. Yet despite multiple stu-dies, the origin of fault weakness is not fully understood and is an on-going matter of debate (Scholz, 2000; Niemeijer et al., 2010; Teseiet al., 2015; King and Morley, 2017). For subduction zones, differencesin seismic coupling, which is an expression of weakness, have long beenknown (e.g. Jarrard, 1986). On a large scale, coupling seems to be re-lated to tectonic forces; subduction zones subjected to high normalforces show high coupling, and below a critical value of normal forcethere is low coupling (Scholz and Campos, 1995; Scholz and Campos,2012). Laboratory experiments and post-seismic studies of aftersliphave shown that the coefficient of fault friction varies proportionally tothe logarithm of sliding velocity (Dieterich, 1979; Hsu et al., 2006;Jolivet et al., 2013; Perfettini and Avouac, 2004). Observation fromlaboratory experiments reinforced using theoretical approaches haveresulted in development of the empirical constitutive formalism of rateand state friction (e.g. Dieterich, 1979; Marone, 1998; Ruina, 1983).This formalism is a widely applied and powerful tool to assess faultbehaviour, describe frictional sliding of rocks, and model geodetic dataover multiple seismic cycles (Jolivet et al., 2013; Perfettini and Avouac,2004; Stevens and Avouac, 2015). Many different forms of this form-alism exist. One version that fits well with laboratory observations isknown as the Dieterich-Ruina law (Scholz, 1998):

�⎜ ⎟= ⎡

⎣⎢+ ⎛

⎝⎞⎠

+ ⎛⎝

⎞⎠

⎤⎦⎥

τ a bσ μ ln VV

ln V θ00

0

(2)

The Dieterich-Ruina law (Eq. (2)) represents friction as the sum ofμ0 (steady state friction at V = V0), a second-order instantaneous de-pendence of friction on log slip velocity V over reference velocity V0,and another second order dependence of friction on the state variable θand critical slip distance ℒ (Beeler, 2009; Scholz, 1998); τ is shearstress, σ is effective normal stress. a and b are empirical materialparameters that depend on applied normal stress, sliding velocity, porefluid pressure and temperature, and define the stability regime of thefault (e.g. Marone, 1998). Where (a–b) ≥ 0, the rock's frictional re-sistance increases with increasing slip rate (rate- or velocity-

Fig. 6. Basic concepts of pressure solution at the grain scale (redrawn after Gratier et al.,2013). A: due to normal stress acting on the grain boundaries, material is dissolved andmay be either re-deposited locally (trapped fluid, material transport over distances of μmto dm), or regionally (free fluid, material transport over larger distances), depending onthe size of the closed system. B: initial sedimentary package before pressure solution. C:small closed system size, local dissolution and precipitation. Dominant process is diffu-sion. Note vertical shortening. D: large closed system, no local precipitation and includingfluid advection possibly over several km. E: oversaturated fluid flow results in mineralprecipitation. F: effect of under-saturated fluid flow including free face dissolution of thegrain boundaries.

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strengthening). In areas where (a–b) ≤ 0 the rock's frictional resistancedecreases with increasing slip rate (rate- or velocity weakening)(Scholz, 1998) (Fig. 8A).

A straightforward illustration of static, quasi-static and dynamicsliding along faults is the spring-slider model (e.g. Burridge andKnopoff, 1967; Fig. 8B), which has successfully been applied to shale-dominated megathrust earthquakes (e.g. Perfettini and Avouac, 2004;Hsu et al., 2006). A block sliding along an interface with a certainfriction coefficient is loaded at a constant velocity through a spring witha defined stiffness. Shear stress during frictional sliding is a function ofeffective normal stress, sliding velocity, and interface frictional para-meters, including (a–b) (Perfettini and Avouac, 2004). In natural sys-tems, this friction parameter may vary along fault strike for a variety ofreasons including: lateral changes in lithology and minerals involved inthe fault zone, variations in fluid-rock interaction, and the effects oftemperature and pressure on fault zone behaviour. These variations canresult in laterally variable mechanical behaviour of a fault zone andcreate regions of asperities on a fault zone that are more prone toseismic slip than other areas (Fig. 9). Using the spring-slider model,McLaskey and Kilgore (2013) show that heterogeneity of fault strengthcan cause slip transients within the nucleation zone of a larger stick-slipevent, explaining short-term foreshocks originating close its hypo-center. Spring-slider experiments indicate complex fault slip behaviouris related to transient frictional dynamics of faults associated withevolving damage zones, elevated pore pressures and structure locali-zation (Leeman et al., 2016). Moreover, not only fault strength, alsoloading conditions or stiffness of the loading system influence fault slipbehaviour, and long healing time followed by abrupt increase inloading rate stimulates seismic slip (McLaskey and Yamashita, 2017,

Scuderi et al., 2016). An interesting implication of these findings is thatthe same fault patch can release seismic energy in different ways,ranging from creep to slow earthquakes to dynamic failure (e.g. Obaraand Kato, 2016). This stresses the importance of transient effects onfault strength, effects that Hubbard et al. (2015) considered importantin the seismic slip history of shallow, weak (shale prone) detachmentfaults.

Based on the rate and state formalism, models can predict stages inthe seismic cycle, interseismic strain accumulation, or the degree ofseismic coupling and seismic rupture patterns (e.g. Ader et al., 2012;Chlieh et al., 2008; Fulton et al., 2010; Kaneko et al., 2010; Lovelessand Meade, 2011; Moreno et al., 2010; Stevens and Avouac, 2015;Thomas et al., 2014b). For active faults lacking shallow creepthroughout the interseismic period, it can be shown that the creepingsegment is characterized by velocity strengthening, whereas lockedfaults are velocity weakening (Barbot et al., 2012; Kaneko et al., 2013).Friction heterogeneities along fault strike provide an explanation for avariety of observed fault behaviours (Kaneko et al., 2010). However, asthe rate-and-state law is based on laboratory experiments conducted atlow slip velocities and slips on the order of centimetres, it does not takeinto account transient effects of dynamic weakening that may occurduring fast slip events, for instance the formation and shearing of asilica gel (Goldsby and Tullis, 2002, 2011), decarbonation (Han et al.,2007), frictional melting (e.g. Tsutsumi and Shimamoto, 1997), porefluid pressurization (e.g. Noda and Lapusta, 2010; Sibson, 1973), andelastohydrodynamic lubrication (Brodsky and Kanamori, 2001). Faultlubrication processes are independent of rock composition, and havebeen observed in different cohesive as well as non-cohesive rocks (DiToro et al., 2011).

Fig. 7. Examples of pressure solution at the grain scale. A:mixed shale-limestone detachment. Sub-vertical foliationdue to pressure solution. Intense veining and dissolution ofthe veins along horizontal pressure solution seams suggestssmall system size and alternating or coinciding solution andprecipitation. B: vertical and sub-vertical pressure solutionseams. Whereas calcite is dissolved and largely transportedout of the area, quartz grains remain intact. This results inconcentration of quartz and clay minerals in the area, in-fluencing rock strength. C: sub-horizontal anastomosingpressure solution seams. Vein sets are partly parallel topressure solution, but also at low and high angles. Note inthe left a sub-vertical vein has been partly dissolved, and iscrosscut by another vein, sub-parallel to pressure solutionseams. Coarse grains are quartz and calcite, roughly at samequantities. A–C from the Khao-Kwang fold-and-thrust belt,Thailand.

Fig. 8. Rate and state-variable friction law. A: frictionalresponse to a suddenly imposed e-fold increase and thendecrease in sliding velocity as observed in laboratory ex-periments (Scholz, 1998). Initially, friction abruptly in-creases by magnitude a, which is followed by an evolvingdecrease of friction of magnitude b. (a–b) is the materialdependent frictional stability parameter defining the sta-bility regime of the fault. B: spring-slider model illustratingresponse of a fault to a temporal stress perturbation(Perfettini and Avouac, 2004). The block is loaded at aconstant velocity V0. k is stiffness of the spring, τf is shearstress during frictional sliding. Despite its simplicity, thismodel can be successfully used to derive the analytical

expression of transient slip along seismogenic faults.

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To incorporate the physics of co- inter- and post-seismic phases ofearthquakes including fast slip into rate-and-state formalism, quasi-dynamic and fully-dynamic models may be used (e.g. Kaneko et al.,2010; Lapusta and Liu, 2009; Rice, 1993; Thomas et al., 2014a).Nevertheless, even with fully dynamic models, some observationscannot be fully explained in the rate and state framework, e.g. tremor,slow earthquakes, or slip transients on all timescales (e.g. Schwartz andRokosky, 2007). Furthermore, inherited crustal damage around a faultmay influence earthquake slip (Cappa et al., 2014). Causes for this maybe that full representation of all fault and rock heterogeneities occur-ring at various temporal and special scales is difficult, and these dy-namic effects must be simplified to some extent (Lapusta et al., 2000;Thomas et al., 2014a). However, using the rate and state formalism it ispossible to determine whether a fault segment is creeping or locked(e.g. Barbot et al., 2012), and whether creep is associated with weakfaults.

Creeping faults were recognized in the 1960′s, and since thencreeping fault segments have been described for compressional, ex-tensional and strike slip settings around the globe, for instance alongthe San Andreas Fault, USA (de Michele et al., 2011; Galehouse andLienkaemper, 2003; Steinbrugge et al., 1960), the Asal Rift, Djibouti(Doubre and Peltzer, 2007), the Haiyuan fault, China (Jolivet et al.,2013), the Tainan tableland and the Longitudinal Valley Fault, both inTaiwan (Huang et al., 2009; Thomas et al., 2014b), the North AnatolianFault, Turkey (Kaneko et al., 2013), and the Philippine Fault(Duquesnoy et al., 1994), among many others. Laboratory experimentsshow that rocks creep and fail statically at stresses significantly lowerthan their short term failure strength (e.g. Kranz, 1979; Lockner, 1998).

The classic experiments on fracturing of natural rocks were per-formed on Westerly Granite (e.g. Yukutake, 1989; Lockner et al., 1991).Even though granite has a very different rheology compared to shales,the basic finding is applicable: fracturing in heterogeneous mediacannot be described as propagation of a single crack. Instead, de-formation history, grain-scale heterogeneity and interaction betweenmultiple cracks growing at the same time control crack growth througha zone of distributed damage at its tip (e.g. Lockner et al., 1991; Huanget al., 1991, Lyakhovsky et al., 1997). Three characteristic creep phasescan be determined: 1) primary, transient creep characterized by time-dependent strain rate, associated with formation of micro-cracks; 2)secondary creep at a constant strain rate where cracks grow but do not

interact, and 3) tertiary, accelerating creep culminating in failure due tointeraction of cracks (Brantut et al., 2012; Lockner, 1998). These stu-dies emphasize that rock strength is time dependent, the process isreferred to as ‘brittle creep’ (Brantut et al., 2012). Laboratory experi-ments allow for empirical exponential functions to be established,analogous to the rate and state formalism, that describe the relationshipbetween creep strain rate and stress deficit, where the state variablerelates to inelastic axial strain (Brantut et al., 2014). For instance, la-boratory creep experiments from shale gas reservoirs show the con-stitutive relation between strain and time is best described by a power-law function of time (Sone and Zoback, 2014). The transition from sub-critical cracking to pressure solution dominated creep occurs with in-creasing depth and decreasing strain rates (Brantut et al., 2012).Keeping the rate and state formalism and creep in mind, we will discussin the following sections the three main causes that have been putforward to explain fault weakness: overpressure, mineralogy, and effi-ciency of structure localization.

4. Causes for detachment weakness

In sedimentary basins salt and shale are the most common detach-ment-forming lithologies, because they are considerably weaker thanother lithologies (Fig. 10). Using a representative strain rate of10−14 s−1, Fig. 10 shows the tensile and compressional strengths ofdifferent sedimentary rock types and how their strength within thebrittle regime increases with depth (confining pressure). The dasheddry limestone line towards the base of the figure shows decreasingstrength with depth due to the onset of ductile processes. The creep andfrictional strengths of wet and dry salt are much lower than other se-diments, with the exception of very shallowly-buried (less than a fewhundred metres) water-saturated muds, where the water content com-monly exceeds 50%. Fig. 10 illustrates that the presence of water,particularly overpressured water (λ = pore fluid pressure ratio) con-siderably weakens shale (and other lithologies).

It is well established that clay minerals, particularly smectite, can beconsiderably weaker than most rocks that follow the Byerlee frictionlaw (Fig. 11, e.g. Morrow et al., 1992; Marone, 1998; Haines et al.,2013; Tesei et al., 2015; Carpenter et al., 2016), and that pore fluidpressures are commonly high around shale detachments (Davis et al.,1983; Dahlen, 1990; Morley et al., 2011). On a case-by-case basis the

Seism

ogen

ic zo

ne

Accretionary wedge

Trench

Uppe

r Plat

e

Aseismic (stable)

Downdip limit

Updip limit

Subduc

ting pla

te

Asperity (unaltered patch)

strong coupling, quick release non-asperity (extensively hydrated patch)

weak coupling, and slow release

Str

ess/s

tra

in

Str

ess/s

tra

in

Inte

r-seism

ic

Co

-se

ism

ic

Post-seismic

emiTemiT

Fig. 9. Schematic model of the plate interface at a sub-duction zone. Fluid supply to the plate interface promoteshydration reactions, forming clay minerals, which influ-ence the frictional properties of the subducting plate. Clay-rich regions could of the interface may be weakly coupled,and act as non-asperities. Unaltered patches remainstrongly coupled, and may give rise to megathrust earth-quakes. The schematic models of the seismic cycles showco-seismic slip generated by the unhydrated strongly cou-pled patches, and post-seismic slip controlled by theweakly coupled hydrated patches. Redrawn fromKatayama et al. (2015).

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key factors that causes shale thrust zone and detachments weaknessremains a matter of debate (e.g. Tesei et al., 2015; Suppe, 2007; Kingand Morley, 2017), and goes beyond just considering bulk variations inmineralogy and pore fluid pressure to include strain localization pro-cesses (e.g. Rutter et al., 2013) and the dynamics of earthquakes (e.g.Wang and Hu, 2006).

As reviewed above, different processes and mechanisms control themicrofabrics of clays, and formation of heterogeneity and structurelocalization occurs already during earliest deposition and continuesduring diagenesis and subsequent deformation events. This microfabricresults in areas of higher and lesser porosity and permeability, influ-encing pore pressures. Pore pressures are also related to stratigraphicthickness and mineralogy. Likewise, transient effects related to tectonicforcing influence detachment behaviour. With the close relations inmind, in the following sections we discuss known causes for detach-ment weakness, in particular overpressure, mineralogy, structure lo-calization, and depth dependent changes.

4.1. Overpressure

While the detachment weakness can be entirely due to materialproperties, overpressure is commonly invoked as a means of attainingthe narrow taper of critical taper wedges observed in many natural foldand thrust belt examples (e.g. Dahlen, 1990; Davis et al., 1983; Bilottiand Shaw, 2005; Suppe, 2007). In saturated systems, fluids support

some of the applied load, lowering the total stress on the grain frame-work. A possibility to describe the influence of pore pressure onstrength is the law of effective stress:

′ = −σ σ χu (3)

where σ′ is effective stress, σ is total stress, u is pore pressure, and χ isthe effective pressure coefficient, accounting for partial transfer of porepressure to the granular framework (e.g. Cuss et al., 2012). χ is variablefor different shales, and depends upon the distribution of clay minerals(e.g. Kwon et al., 2001). Water and gas injection tests in different ar-gillaceous rocks have shown that even though many case studies showrocks follow the law of effective stress, a single bulk χ-value may not berepresentative due to localized deformation and pore-pressure varia-tions at small scale (Cuss et al., 2012).

Pore fluid pressure influences crustal strength, and it has long beenknown that particularly in sedimentary basins pore fluid pressures mayexceed hydrostatic conditions at depth (e.g. Terzaghi, 1943, Hubbertand Rubey, 1959; see review in Osborne and Swarbrick, 1997). Basedon geochemical analysis and temperature anomalies, it has been shownthat in accretionary prisms, faults are the main locus of fluid flow(Moore et al., 1987; Gieskes et al., 1990; Fisher and Hounslow, 1990a,1990b), as fault-parallel permeability may be 10–10,000 time greaterthan the permeability of adjacent rocks (Moore et al., 1995). Thispermeability distribution is not straight forward, since shear zones canalso exhibit permeability which may be several orders of magnitudelower than the host rock due to increased alignment of flaky clay mi-nerals and shear induced pore space collapse (Maltman et al., 1997;Clennell et al., 1998). Moreover, post-deformational cementation maydecrease fault zone permeability (Clennell et al., 1998). Data frommultiple wells shows that below a certain depth, called the fluid-re-tention depth, overburden loading becomes partially supported by porefluids rather than rock compaction (Suppe, 2014). This modifies theclassic linearly increasing brittle strength-depth relationship (Brace andKohlstedt, 1980) (Fig. 10), where increased pore fluid pressures lead toconstant and low brittle strength below the fluid-retention depth(Suppe, 2014). Experimental data on sandstones with and withoutelevated pore pressure show excess pore pressures causes reduction ofbrittle strength, and enhances slip instability (Ougier-Simonin and Zhu,2013). Slow slip events seem to require a brittle-ductile transitionalregime and elevated pore pressures (Dragert et al., 2004). The Hubbertand Rubey (1959) model predicts high pore-pressure aids motion onregional detachments. Shale detachments in sedimentary basins arewidely associated with high pore fluid pressures (e.g. Dahlen, 1990;Fisher and Hounslow, 1990a, 1990b; Moore et al., 1995; Fisher et al.,1996; Fisher and Zwart, 1997; Maltman et al., 1997; Moore and Tobin,1997; Screaton et al., 1997; Cobbold et al., 2004, 2009; Bilotti andShaw, 2005; Morley, 2007; Saffer and Tobin, 2011). Notably in some

0

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4

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th (

km)

0 10 20 30 40 50-50 -40 -30 -20 -10

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Dry limeston

e and quart

zite

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Wet qu

artzite

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Dry sh

ale λ =

0

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Wet

shale

λ =

0.42 Wet shale λ = 0.42

Wet

sha

le λ

= 0

.82 Wet shale λ = 0.82

Dry salt

Wet salt

Fig. 10. Comparison of creep and frictional strength of wetand dry sedimentary rocks (from Jackson et al., 1994a,b).Creep strength is the deviatoric stress necessary to causesalt flow at a rate of 10−14 s−1. Pore pressure ratio (λ) = 0for dry rocks, 0.46 for hydrostatic pressure and 0.86 foroverpressure. Wet salt is so weak it virtually plots along the0 MPa line.

Fig. 11. Evolution of fault permeability and adjacent wall rock at similar depths duringthrust-associated loading (Moore et al., 1995). Wall rock permeability is higher than faultpermeability when effective confining stresses are high. Transiently raised pore pressuresresults in dilation and short-lived fault permeability higher than wall rock permeability.This high fault parallel permeability causes drainage of the excess pressures and sub-sequent fault sealing.

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examples of industry drilling overpressures are low to moderate in thesection (3–4 km thick) overlying a detachment and pore fluid pressureratio increases rapidly in the section (100's m thick) overlying the de-tachment (e.g. Erinmi-1X well, outer Niger Delta in Cobbold et al.,2009). Indirect evidence of widespread overpressured conditions asso-ciated with shale detachments is provided by the numerous mud vol-canoes and fluid escape features present in most deepwater fold andthrust belts, some of which require enormous quantities of fluid to drivethem (see review in Morley et al., 2011). The deep-seated nature ofmany fluid escape pipes and mud volcano feeders is shown by seismicreflection data that reveals fluid pipes rising from large thrust faults(e.g. Oregon margin, MacKay et al., 1992; Niger Delta, Cobbold et al.,2009; Baram Delta Province, Morley, 2009; Hikurang subduction zone,Barnes et al., 2010) and sea floor sampling of mud volcanoes, whichshows many have high heat flows and the fluids contain thermogenichydrocarbons (e.g. Deville et al., 2003, 2006; Dolan et al., 2004;Zielinski et al., 2007; Warren et al., 2011). Local negative-polarityseismic reflections of many accretionary prisms including the Oregonaccretionary prism indicate well-defined zones of high fluid pressurewithin proto-thrusts and frontal thrusts, which are bounded by discretepermeability barriers (Moore et al., 1995). Experimental results andnumerical models suggest the observed velocity reduction can be ex-plained with fluid pressures in the range of 86%–99% of lithostatic forthe in situ confining stress (Tobin et al., 1994).

A high fluid-pressure, dilatant zone of high porosity about 14 mthick characterizes the basal décollement in the northern BarbadosRidge (Shipley et al., 1994). In the Nankai accretionary prism, very low-frequency earthquakes and tremors are associated with elevated porepressures (Ito and Obara, 2006; Walter et al., 2013), and high porepressures coupled with strengthening by lithification focuses deforma-tion (Lewis and Byrne, 2003). Similarly, fluid injection may triggerseismic and aseismic slip, and multiple examples exist from all over theworld (e.g. Hubbert and Rubey, 1959; Zoback and Harjes, 1997). Fluidinjection helps identification of zones of high pore pressure (Cornetet al., 1997). Fluid injection during a geothermal drilling project inBasel, Switzerland stimulated several thousand earthquakes in 2006and 2007 (Bachmann et al., 2011; Deichmann and Ernst, 2009; Häringet al., 2008; Ripperger et al., 2009). The infamous 2012 Brawley Swarmin California was triggered by injection-induced aseismic slip (Weiet al., 2015). Controlled pressurization of low-permeability faults showsinduced seismicity is related to mechanical loading induced by openingof adjacent permeable sub-faults by a few microns (Guglielmi et al.,2008).

Within sedimentary basins there are a number of mechanisms thatcan generate overpressures at different depths, these include: 1) dis-equilibrium compaction in response to burial, 2) disequilibrium com-paction in response to horizontal compression-induced shortening, 3)thermo-chemical reactions (e.g. smectite to illite, kaolinite to illite,kerogen to oil and gas, and oil to gas transformations; see reviews inOsborne and Swarbrick, 1997 and Morley et al., 2014), and 4) per-meability decrease due to increased shear strain (Ikari et al., 2009a,2009b). Thermo-chemical reactions discussed here are generally lim-ited to temperatures< 150–160 °C (as reviewed by Morley et al.,2014). The maturation of hydrocarbons has been proposed as a sig-nificant cause of overpressure in shale detachments and thrusts atdepths around 3–6 km (Morley, 1992; Morley and Guerin, 1996;Cobbold et al., 2013; Zanella et al., 2014). For the Niger Delta thrustbelts Cobbold et al. (2009) suggested that the thrust front may track theonset of the hydrocarbon maturation widow, which creates elevatedpore pressures. Hence within a homogeneous, fine grained sequence thelocation of overpressured zones, rather than frictional properties of theclay minerals may dictate the location of the detachment. The rate ofadvance of the detachment may also be controlled by the kinetics ofmaturation (Cobbold et al., 2009). The increase of deformation in-tensity within the most organic-rich portion of the Marcellus Shale ledAydin and Engelder (2014) to conclude hydrocarbon generation was

important to the localization of thrusting to that part of the shale.However, such a mechanism clearly does not apply to all mudrockssince some are low in organic matter, or were either over-mature orimmature for hydrocarbon generation at their time of deformation.

To sustain overpressures in detachment zones that penetrate deep insedimentary basins and further into the crust other mechanisms foroverpressure must be sought. The widespread occurrence of high porefluid pressures is perhaps easiest to argue for in accretionary prismsettings, where large quantities of fluid are derived from the subductingslab (e.g. Berhmann, 1991; Barnes et al., 2010; Saffer and Tobin, 2011).For the Southern Alps of New Zealand, seismic P-wave velocity analysis,local tomography and magnetotelluric data suggest near lithostatic porefluid pressures extend to depths between 20 and 30 km in the crust(Stern et al., 2001; Eberhart-Phillips and Bannister, 2002; Wannamakeret al., 2002; and Scherwath et al., 2003). Prograde metamorphism ofclastic sedimentary rocks is probably responsible (Etheridge et al.,1984). The transition from shale to phyllite and schist may result innear lithostatic fluid pressures (Suppe, 2014). The widespread occur-rence of veins in shale detachments in accretionary prisms also in-dicates near-lithostatic fluid pressures (e.g. Etheridge et al., 1984;Berhmann, 1991; Needham, 2004; Fagereng et al., 2010; Fagereng,2011; Van Noten et al., 2011; Fagereng and Harris, 2014). Widespreadhigh pore fluid pressures (estimated pore pressure ratios generally be-tween 0.6 and 0.9) around a broad basal detachment zone in the NankaiTrough have been interpreted based on seismic reflection data (Tsujiet al., 2014). A study on the Nankai Trough based on laboratory mea-surements of core samples from the region, and numerical modelssuggests pore pressure ratios along the detachment exhibit pore pres-sure ratios between 0.68 and 0.77 (Skarbek and Saffer, 2009). In theforeland fold-and-thrust belt of the European Alps, Laurich et al. (2014)showed a strong porosity reduction for thin shear zones in scaly clay,which is a common feature alongside several large detachments (e.g.Chester et al., 2013). Formation of scaly-fabric zones is associated withflattening (Fig. 2). and porosity loss, which may drive overpressuredfluids into more limited pockets of remain pore space, or completelyexpel them from the fault zone (e.g. Labaume et al., 1997). Duringstages of relaxed shear stress dilation may occur associated with ingressfluids close to lithostatic pressure, and associated with this phase ex-tensional veins are formed (often along spaced foliations), and did notform during shear movement. In this kind of evolution Labaume et al.(1997) envisaged scaly-fabrics formed during stages of more moderateoverpressures (limited displacement strong flattening), while displace-ment on the boundaries of the scaly-fabric zones occurred duringphases of relatively high overpressure (higher displacement, but limitedcompaction within the décollement zone sediment).

Elevated pore pressures along permeable faults must be constantlyor periodically replenished by fluid flow or they would drop throughdiffusion or advection. Fluid pressures are not stable but evolve as afunction of fault permeability. For the frontal Oregon accretionaryprism, permeability along the active area of faults exceeds permeabilityof adjacent sandy layers (i.e.> 10−13–10−14 m2), as they capture flowfrom sandy layers, whereas permeability across the fault is several or-ders of magnitude lower (10−16–10−17 m2) (Moore et al., 1995). Thesevalues are transient, and vary according to the stress path of the faultzone. This permeability anisotropy is related to the preferred orienta-tion of minerals and fracture fabrics of accreted rocks, as shown inexperimental data and drill-hole data from various accretionary prisms(Arch and Maltman, 1990; Behrmann et al., 1988; Moore and Vrolijk,1992). Experimental deformation of mud samples by Moore et al.(1995) show fault permeability evolution through time (Fig. 11): duringincreasing compactional loading faults decrease in permeability andeventually become aquitards. In Barbados, burial compaction followedtectonic thickening and caused pore volume loss of up to 70%, asso-ciated with defluidization (Speed, 1990). After rocks have becomeimpermeable, new permeability is created along fault and fracturenetworks (Moore and Vrolijk, 1992). This requires intermittent dilation

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along the faults, which may be related to transiently elevated porepressures (Moore et al., 1995), for instance due to hydraulic connectionto an overpressurized lower reservoir (Célérier et al., 1998). Fluidpumping through a shear zone involves fluid-filled pores that commu-nicate along grain boundaries during deformation, and granular-scalefluid pressure differences that drive fluid flow (Fusseis et al., 2009).Alternatively, the vertical stress magnitude may be decreased by ero-sion or a change in tectonic regime, which may result in dilatant frac-turing of the fault zone (Célérier et al., 1998). In deeper regions of theaccretionary prism crack-seal textures indicate episodic phases of veinopening and consequently great fluctuations of permeability and pres-sure through time (Byrne and Fisher, 1990; Moore et al., 1995). Rockdeformation experiments suggest a stress dependency related to thepermeability of fractures; once a fracture is established it can remainopen even without fluid overpressures (Gutierrez et al., 2000). Veinformation also alters the mechanical properties of the fault zone. Anincrease in stress will result in brittle deformation focussed on theveined fault zone, breaching its integrity and increasing its permeabilitywith respect to the surrounding rock (Clennell et al., 1998).

Application of the simple effective stress equation (σ′=σ−u), willbe insufficient when there is significant flow of fluid affecting the de-tachment, which gives rise to seepage force (Cobbold et al., 2001;Mourgues and Cobbold, 2003). Overpressure magnitudes vary within arock volume, and overpressured fluid will flow into regions of lowerfluid pressure when permeability permits. Flowing fluid exerts a see-page force (j) on the solid rock fabric related to fluid drag on the rockelements, and the pressure gradient causing the fluid flow. Seepageforce is j= i γw, where i = the hydraulic gradient, and γw is the unitweight of water (~1 g/cm3). Effective stress at a point should be cal-culated by considering the total stress plus the effects of fluid flow,before subtracting the fluid overpressure (Cobbold et al., 2001;Mourgues and Cobbold, 2003). Seepage forces act in the same directionas flow, hence in the case of vertical fluid flow seepage forces actvertically, but not horizontally. Seepage forces were invoked byCobbold and Rodrigues (2007) to explain the widespread occurrence ofbedding parallel mineralized veins in sedimentary basins.

Mourgues and Cobbold (2006a) reported the results of gravity-driven deformation in sand analogue models using air injection into thebase of the experiments to model the effects of overpressure. Theyfound that the types of detachment styles that arise from overpressuresand associated seepage forces, were thin flat shear bands very differentin appearance from analogue models that use silicone putty. Fault dipangle was a function of basal frictional resistance and fluid pressure. Inair injection experiments investigating critical taper Mourgues andCobbold (2006b) found that permeability of the wedge was more im-portant than the coefficients of friction in controlling wedge taper. Inhomogeneous sand deformed under high fluid pressure gradients, de-formation became diffuse, with a ductile appearance, whereas in multi-layered models detachments formed below low permeability layers,where sharp drops in shear strength were related to high overpressures(Cobbold et al., 2001; Mourgues and Cobbold, 2006b).

The effect of seepage forces on thrusts will be highly variable indifferent regions of large thrusts, basins and tectonic settings. In somesettings or stages of fault development there maybe little fluid flow, oreven the presence of hydrostatic pressures. For example the shale de-tachment forming the Osen-Røa thrust overlies crystalline Precambrianbasement (i.e. a very poor and impermeable fluid source), and the basaldetachment is very close to the interface (Morley, 1986, 1987), hence itis difficult to see how significant vertical fluid flow across the detach-ment could occur. Conversely dewatering from mass transport com-plexes may produce very significant, short-lived seepage forces. In othersettings, such as accretionary prisms fluid flow may be gently inclinedalong the basal detachment, rather than vertical due to strati-graphically-controlled permeability barriers.

Fluid pressures can also become elevated and affect the effectivestress during earthquake ruptures due to frictional heating. The degree

to which this process of thermal pressurization operates has been un-certain in the past, although today its significant role is more widelyappreciated (e.g. Sibson, 1973; Lachenbruch, 1980; Brodsky andKanamori, 2001; Bizzarri and Cocco, 2006; Noda and Lapusta, 2010).According to Sibson (2003) one meter of slip on a zone a few milli-meters thick would produce a temperature change larger than 1000 °C.Evidence for such temperatures would be the presence of pseudo-tachylyte, however they are relatively scarce in exhumed fault zones(Sibson, 2003). This scarcity may reflect a preservation issue, ratherthan a problem with the model. Drilling following the Chi-Chi earth-quake established the presence of pseudotachylyte in the earthquakerupture zone, but it was already 75% altered to smectite just a few yearsafter the earthquake occurred (Kuo et al., 2009). Numerical modelingby Bizzarri and Cocco (2006) suggested that temperature changes as-sociated with slip on a thin fault zone are in the order of 800 °C, andwould be sufficient to cause considerable melting. Bizzarri and Cocco(2006) considered that both thermal pressurization and flash heating ofcontact asperites along a fault are important mechanisms for droppingthe friction coefficient during an earthquake, and causes of dynamicfault weakening. The mechanisms are thought to be particularly sig-nificant at the late nucleation stage of an earthquake at, or beforeseismic slip speeds are reached (Schmitt et al., 2011). Fig. 12 illustratesthe dynamics of a fault-fluid system and encompasses the earthquakecycle (Sibson, 1990; Goodwin et al., 1999). High pore fluid-pressuresdecrease effective normal stress, facilitating faulting (Skempton, 1984).Fluid pressure then drops during faulting as fluids are pumped along,and expelled from the fault zone. Structure localization and post-seismic mineralization and diagenesis results in permeability decreaseand consequent fluid pressure build-up (Goodwin et al., 1999).

Even in wells where no elevated pore pressures are observed, it hasbeen shown that significant amounts of nano-porosity exist in clays-tones and shales (Desbois et al., 2014; Klaver et al., 2015), as well as infault gouge (Janssen et al., 2011). This porosity increases with in-creasing clay content and is possibly linked to maximum strain (Janssenet al., 2011).

There are, however, several potential concerns about invoking fluidpressures as the main explanation for detachment weakness. First, ex-posed accretionary prisms and subduction interfaces as well as the toeof accretionary prisms show locally high concentration of carbonateand quartz veins, and pockets of hydraulic breccia suggesting highlylocalized fluid flow (Maltman et al., 1993; Moore et al., 2007). It thusseems difficult to assume high pore pressures can be sustained over thewhole length of major thrust sheet. Second, it is unclear how high pore

Per

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Dilation & elevated pore pressure

Fig. 12. Dynamics of a fault-fluid system, encompassing the seismic cycle (Goodwinet al., 1999; Sibson, 1990). Raising the pore fluid pressure decreases the effective normalstress on the fault, facilitate fault rupture. Faulting results in drop of fluid pressure.Subsequently, diagenesis and mineral precipitation during the interseismic phase maydecrease permeability and increase fluid pressures again.

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pressures are maintained over the long-term movement of the thrustsheet, a process that is likely to lead to leakage of fluids (see review inAydin and Engelder, 2014). However, as discussed above there are awide range of causes of overpressure, some of which are transient andrelated to the earthquake cycle (e.g. thermal pressurization), conse-quently these very general arguments related to loss of overpressuredue to fluid leaka