Reactivation of normal faults as high-angle reverse faults due to low frictional strength: Experimental data from the Moonlight Fault Zone, New Zealand

Journal of Structural Geology 105 (2017) 34–43

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Reactivation of normal faults as high-angle reverse faults due to low frictional strength: Experimental data from the Moonlight Fault Zone, New Zealand

T

S.A.F. Smitha,∗, T. Teseib, J.M. Scotta, C. Collettinib,c a

Geology Department, University of Otago, 9054, Dunedin, New Zealand Istituto Nazionale di Geofisica e Vulcanologia (INGV), 00143, Rome, Italy c Dipartimento di Scienze della Terra, Università di Roma, La Sapienza, Italy b

A R T I C L E I N F O

A B S T R A C T

Keywords: Friction High-angle reverse fault Reactivation Basin inversion Phyllosilicate Fault zone weakening

Large normal faults are frequently reactivated as high-angle reverse faults during basin inversion. Elevated fluid pressure is commonly invoked to explain high-angle reverse slip. Analogue and numerical modeling have demonstrated that frictional weakening may also promote high-angle reverse slip, but there are currently no frictional strength measurements available for fault rocks collected from large high-angle reverse faults. To test the hypothesis that frictional weakening could facilitate high-angle reverse slip, we performed single- and double-direct friction experiments on fault rocks collected from the Moonlight Fault Zone in New Zealand, a basin-bounding normal fault zone that was reactivated as a high-angle reverse fault (present-day dip angle 60°–75°). The fault core is exposed in quartzofeldspathic schists exhumed from c. 4–8 km depth and contains a < 20 m thick sequence of breccias, cataclasites and foliated cataclasites that are enriched in chlorite and muscovite. Friction experiments on water-saturated, intact samples of foliated cataclasite at room temperature and normal stresses up to 75 MPa yielded friction coefficients of 0.19 < μ < 0.25. On the assumption of horizontal maximum compressive stress, reactivation analysis indicates that a friction coefficient of < 0.25 will permit slip on high-angle reverse faults at hydrostatic (or even sub-hydrostatic) fluid pressures. Since foliated and phyllosilicate-rich fault rocks are common in large reactivated fault zones at basement depths, long-term frictional weakening is likely to act in concert with episodic build-ups of fluid pressure to promote high-angle reverse slip during basin inversion.

1. Introduction Basin-scale normal faults are often reactivated as high-angle reverse faults during compressional inversion (Turner and Williams, 2004 and references therein). This process is widely recognized as an important factor in the geometric and mechanical evolution of orogenic belts worldwide (e.g. Butler et al., 2006). Reactivation of normal faults has been well documented in petroleum basins, where seismic reflection profiles are combined with borehole measurements to reconstruct fault displacements during basin formation and subsequent compressional reactivation (e.g. Dore and Lundin, 1996; Reilly et al., 2015). Compilations of thrust and reverse fault earthquake ruptures from intracontinental regions show a prominent peak at fault dip angles of 45°–55°, indicating ongoing compressional reactivation of former normal faults at basement depths (Sibson and Xie, 1998). Mechanical models of high-angle reverse faulting have focused on

∗

Corresponding author. E-mail address: [email protected] (S.A.F. Smith).

https://doi.org/10.1016/j.jsg.2017.10.009 Received 4 July 2017; Received in revised form 26 October 2017; Accepted 26 October 2017 Available online 28 October 2017 0191-8141/ © 2017 Elsevier Ltd. All rights reserved.

the influence of elevated fluid pressure. For typical “Byerlee” (1978) values of frictional strength, reactivation of high-angle (dip angle > 60°) reverse faults is only possible if the fluid pressure approaches lithostatic or exceeds the (vertical) minimum principal stress (Sibson, 1985). If these conditions are not met, a new fault with a more optimal thrust orientation will form in preference to reactivation of the steeplydipping reverse fault. Evidence of elevated fluid pressure during highangle reverse faulting includes field observations of networks of hydraulic extension veins within and surrounding low-to moderate-displacement reverse faults (Sibson et al., 1988; Cox, 1995; Turner and Williams, 2004). Additionally, geophysical data from areas of active inversion include seismically low-velocity zones, reduced P- and S-wave velocities and anomalous bright spots in the middle to lower-crust in areas where recent reverse-fault ruptures have occurred (Sibson, 2009). These data have been interpreted to indicate regions of variably pressurized and interconnected aqueous fluid where active reverse faulting

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faulting also formed broad regional-scale folds including the Earnslaw Synform and Shotover Antiform (Fig. 1b) (Turnbull et al., 1975; Norris et al., 1978; Craw, 1985). In close proximity to the MFZ, belts of tightlyspaced chevron folds and kink bands formed in the footwall schists (Barry, 1966; Craw, 1985). Based on a study of conjugate kink bands adjacent to the MFZ, Barry (1966) inferred that the maximum principal stress (σ1) at the time of reverse faulting plunged shallowly (< 10°) to the ESE, at large angles (> 70°) to the MFZ (Fig. 1c). This is similar to the present-day regional orientation of σ1 in Otago derived from studies of microseismicity, which infer a sub-horizontal σ1 trending between c. 110°–120° (Fig. 1a) (Scholz et al., 1973; Norris and Carter, 1982; Reyners et al., 1983; Townend et al., 2012; Warren-Smith et al., 2017).

is occurring. Compressional inversion of sedimentary basins represents an optimal situation for overpressure generation due to the combined effects of porosity reduction (Turner and Williams, 2004) and increases in mean stress coupled with a decrease in overall fluid storage capacity (Sibson, 1995). However, fluid pressures near or exceeding the lithostatic load are episodic and are likely to be relaxed by faulting and fracturing (Townend and Zoback, 2000). Analogue and numerical models have shown that frictional weakening of high-angle reverse faults can also result in reactivation. For example, Marques and Nogueira (2008) carried out sandbox experiments with a pre-existing weak fault dipping at up to 70° and showed that high-angle reverse faulting in their experiments accommodated up to 30% shortening before a new thrust was initiated. Comparable results are obtained in numerical experiments with variable amounts of fault weakening (Ruh and Verges, 2017). Despite these results pointing towards frictional weakening as a potentially important process during high-angle reverse faulting, no frictional strength measurements have been made on fault rock materials collected from large-displacement reverse faults. To test the hypothesis that low frictional strength could contribute to reactivation and slip along high-angle reverse faults, we carried out frictional sliding experiments on fault rock materials from the core of the Moonlight Fault Zone (MFZ), a basin-scale inverted normal fault in the South Island of New Zealand. The MFZ has been uplifted and deeply exhumed in the last 5 Ma, and therefore presents an opportunity to quantify the weakening processes that were active in a basin-scale reverse fault at basement depths. Alder et al. (2016) documented the structure and fault rock assemblages of the MFZ. We review and expand on the work of Alder et al. (2016) before presenting our frictional strength measurements that are analyzed in the context of simple twodimensional frictional reactivation models.

3. Structure and fault rock assemblages in the Moonlight Fault Zone Alder et al. (2016) described the structure and fault rock assemblages of the MFZ based on detailed field and microstructural analysis carried out in five creek sections that cut across the hanging-wall, fault core and footwall (Fig. 1b). In the studied exposures, the MFZ dips between 60 and 75° to the west or NW (Fig. 1b and c). Host rock mineralogy (see below) indicates peak metamorphism at chlorite-zone lower greenschist facies at around 300 °C and 4–5 kbar (Mortimer, 2000). Estimates of reverse offset along the MFZ are in the range of 3–5 km, with exhumation of exposed fault rocks likely in the range of 4–8 km, increasing north towards the Alpine Fault (Alder et al., 2016). The creek sections that were studied by Alder et al. (2016) (Fig. 1b) therefore reveal structures and fault rock assemblages that are representative of a large-displacement, high-angle reverse fault exhumed from mid-to upper-crustal depths that reactivated a basin-scale normal fault zone. Though the structure and fault rock sequences of the MFZ vary along strike as a function of host rock composition and other factors (Alder et al., 2016), key characteristics are represented in the Matukituki Valley section where samples were collected for experiments (Figs. 1c and 2). Below, we summarize the observations of Alder et al. (2016) from the Matukituki Valley and expand on aspects of fault rock microstructure that are relevant to interpretation of the new experimental data. Movements along the MFZ in the Matukituki Valley juxtaposed strongly foliated and lineated metabasite “greenschists” in the hangingwall against folded “greyschists” in the footwall (Figs. 1c and 2). The hanging-wall greenschists consist of chlorite, epidote, albite and titanite with minor stilpnomelane, actinolite, calcite, quartz, biotite, apatite and chalcopyrite. Footwall greyschists consist of muscovite, quartz, albite, chlorite, and calcite with minor epidote, titanite, clinozoisite and rutile. The hanging-wall greenschists are laced with pseudotachylytes (solidified frictional melts) in a damage zone exceeding 500 m wide (Fig. 1c). Pseudotachylyte fault veins are mainly subparallel to the steeply-west dipping greenschist foliation (Fig. 1c) and individually accommodate up to tens of centimetres of slip. The core of the MFZ in the Matukituki Valley contains a sequence of breccias, cataclasites and foliated cataclasites up to a total of 10 m wide (also observed in other creek sections along strike; Figs. 1c and 2a). This sequence is fully developed in the footwall greyschists, whereas the more massive and competent hanging-wall greenschists contain only a relatively thin (< 1 m) layer of foliated cataclasite (Fig. 2a). Overall, rocks in the fault core show a textural transition towards the main fault trace (Fig. 2a) that is characterized by: 1) an increase in the degree of brecciation and cataclasis and a progressive decrease in grain size, 2) an increase in the abundance of matrix phyllosilicates (mainly chlorite in the hangingwall and muscovite in the footwall), and 3) the progressive development of a steeply west-dipping, fault-parallel foliation. The foliation is defined by an alignment of platy phyllosilicate grains and elongate quartz-albite clasts, as well as the preferred orientation of seams of relatively insoluble minerals (Fig. 2; see below).

2. The Moonlight Fault Zone: a case study of normal fault reactivation The Moonlight Fault Zone (MFZ) is a regionally important fault in the South Island of New Zealand (Fig. 1a). At the present exposure level the fault mainly cuts the Otago Schist, an extensive metamorphic belt composed of rocks from the Caples, Rakaia and Aspiring tectonostratigraphic terranes (Fig. 1b). These dominantly quartzofeldspathic accretionary terranes were amalgamated during Jurassic – Cretaceous orogenesis and preserve regional-scale isoclinal folds and large nappe structures (Craw, 1985). Metamorphic facies vary across the Otago Schist, ranging from prehnite - pumpellyite facies to upper greenschist facies (Mortimer, 2000). The highest grade assemblages within the Otago schist (garnet-biotite-albite zone greenschist) experienced conditions of 8–10 kbar and 350–400 °C (Mortimer, 2000). Packages of Oligocene sediments up to 450 m thick, known locally as the Bobs Cove Beds (Fig. 1b), are also found as fault-bound slivers and wedges along the main trace of the MFZ (Fig. 1b; Turnbull et al., 1975). In the Late Eocene-Oligocene, regional extension occurred across the South Island of New Zealand and generated rapidly subsiding sedimentary basins (Turnbull et al., 1975; Norris et al., 1978). Norris et al. (1978) suggested that the MFZ initiated during the Early Oligocene as a large basin-bounding normal fault zone down-throwing basement schist to the west and allowing for deposition of the Bobs Cove Beds. Some of the sediments that form the Bobs Cove Beds were derived from the upthrown block of schist in the footwall of the MFZ (Turnbull et al., 1975). During the Miocene the MFZ was reactivated as a high-angle reverse fault as much of the South Island experienced broadly NW-SE directed compression due to oblique convergence along the Australian – Pacific plate boundary (Norris et al., 1990). By the Late Miocene, the sedimentary fills in many of the previously-formed extensional basins had been folded and uplifted above sea level. This phase of reactivation caused dismemberment and infaulting of the Bobs Cove Beds. Reverse 35

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Fig. 1. Simplified tectonic and geological setting of the Moonlight Fault Zone (MFZ). a) The MFZ lies sub-parallel to the Alpine Fault on a regional scale and is close to the western margin of the active Otago reverse-fault province (Litchfield and Norris, 2000). The present-day SHmax orientation in Otago and nearby areas is approximately sub-horizontal and between 110° and 120°. b) Map of the trace of the MFZ between Queenstown and the Matukituki Valley showing the location of the Aspiring, Rakaia and Caples Terranes, and fragments of Oligocene “Bobs Cove Beds” that were originally deposited in a large fault-bound basin. The five creek section labeled 1–5 are those studied in detail by Alder et al. (2016). The stereoplots show representative orientations of the main trace of the MFZ in the various creek sections. c) Summary of field observations from the Matukituki Valley. The core of the MFZ contains the fault rocks sequences shown in Fig. 2, from which the experimental samples were collected.

(Fig. 2b and c). Clasts in both types of foliated cataclasite are mainly quartz and albite (derived from the host rocks) and were deformed by brittle fracturing (Fig. 2). Fractures in quartz-albite clasts often contain authigenic chlorite and muscovite (Fig. 2b,e). Within the matrix of the foliated cataclasites, seams of relatively insoluble minerals including rutile, titanite and Fe-oxides preferentially developed sub-parallel to the steeply-west dipping main fault trace (Fig. 2a,b,d). The seams often truncate the upper and lower edges of quartz and albite clasts (Fig. 2b,e). Authigenic growth of phyllosilicates (in fractures and “strain shadow” regions) and concentration of insoluble minerals along dissolution seams indicate that pressure solution accompanied cataclasis and frictional sliding as a significant deformation mechanism in the matrix of the foliated cataclasites.

Scanning-Electron Microscope (SEM) and X-Ray Diffraction (XRD) observations made by Alder et al. (2016) indicate that the hangingwall-derived foliated cataclasites are dominated (> 95 wt%) by chlorite as the matrix phyllosilicate phase (Fig. 2a–c). The footwall-derived foliated cataclasites are dominated by muscovite (50–80 wt%) with subordinate chlorite (usually < 20 wt% but locally up to 50 wt%; Fig. 2a,d,e). The matrix of the foliated cataclasites consists of interconnected networks of phyllosilicates (Fig. 2b–e). In the footwall muscovite-rich samples, the phyllosilicate layers are relatively pure and wrap around isolated clasts of quartz-albite (Fig. 2d and e). In contrast, the matrix of the chlorite-rich foliated cataclasites contains abundant second phases (including titanite, magnetite, Fe-oxides and epidote) that are dispersed along platy grain boundaries and foliation surfaces such that the overall connectivity of the chlorite layers is disrupted 36

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(caption on next page)

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Fig. 2. Summary of microstructural observations from the fault rock sequences preserved in the Matukituki Valley section of the MFZ (location in Fig. 1). a) Thin section scans of samples collected at distances of < 10 m from the main fault trace, showing the progressive textural and mineralogical changes that occur as the main fault trace is approached. Experimental samples were derived from the hanging-wall (green) and footwall (grey) foliated cataclasites at distances of < 1 m from the main fault trace. b) Optical microscope image of the green chlorite-rich foliated cataclasites showing large clasts of quartz-albite, an interconnected matrix of chlorite and an abundance of disseminated second phases including titanite, magnetite, Fe-oxides and epidote. Some of the second phases are concentrated (by pressure solution) along the margin of the large quartz-albite clast. c) Backscatter SEM image highlighting the abundance of second phases in the chlorite-rich matrix of the hanging-wall foliated cataclasites. Fe-oxides and titanite are widely dispersed between platy chlorite grains. d) Optical microscope image of the grey muscovite-rich foliated cataclasites shows an interconnected network of muscovite surrounding small clasts of quartz-albite. Colour variations in the matrix represent different modal proportions of muscovite and chlorite. e) Backscatter SEM image of the relatively pure muscovite (and chlorite) matrix of the footwall-derived foliated cataclasites. A thin seam of Fe-oxides lines the outer edge of the large quartz-albite clast. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

Fig. 3. Experimental techniques. Photos of cut faces of the a) chlorite-rich (hanging-wall) foliated cataclasites and b) muscovite-rich (footwall) foliated cataclasites. c) Intact sample wafers were cut from the phyllosilicate-rich portions of the foliated cataclasites, but clasts of quartz-albite were inevitably included. The metal side block contains teeth orthogonal to the experimental slip direction. d) Cartoon of single-direct experimental geometry showing location of sample wafer. Inset shows double-direct geometry used in one experiment (see Table 1).

4. Frictional strength of foliated cataclasites from the core of the Moonlight Fault Zone

et al., 2009a) with dimensions of 4 × 4 cm (Fig. 3c). The initial thickness of the wafers was between 1.0 and 1.4 cm. The wafers were cut from areas dominated by phyllosilicate-rich matrix, although quartz-albite clasts up to c. 1 cm in size were inevitably included in the wafers (Fig. 3c). The wafers were sandwiched between a steel side block and a steel central block in “single-direct” configuration (Fig. 3d). Constant normal stress (σn) was applied during the experiments by the horizontal piston acting in load-feedback control mode (Fig. 3d). Deformation of the sample wafers (shear stress, τ) was accomplished by moving the vertical piston downwards in displacement-feedback control (Fig. 3d). The vertical piston was advanced at a velocity of 10 μm/s. Both the side and central blocks were lined with teeth orthogonal to the slip direction to prevent shearing becoming localized along the sampleblock interfaces (Fig. 3c). The central block and the support for the side block were lubricated with MoS2 (frictional strength, μ, < 0.01) where they come in to contact with the loading frame (Fig. 3d). One experiment was performed in “double-direct” configuration using two sample

4.1. Experimental methods To measure the frictional strength of foliated cataclasites from the core of the MFZ, eight experiments were performed in the BRAVA biaxial rock deformation apparatus (Fig. 3; Collettini et al., 2014) at the National Institute of Geophysics and Volcanology (INGV) in Rome, Italy (Table 1). Experiments were performed at room temperature under water-saturated conditions. Three experiments were performed on chlorite-rich foliated cataclasites derived from hanging-wall greenschists (Figs. 2 and 3a) and five experiments were performed on muscovite-rich foliated cataclasites derived from footwall greyschists (Figs. 2 and 3b; Table 1). To test the strength of the natural fault rock fabric, friction experiments were performed on rectangular sample wafers (e.g. Collettini

Table 1 Experiments performed on footwall and hanging-wall foliated cataclasites from the Moonlight Fault Zone. Experiment

Geometry

Foliated Cataclasite

Initial thickness (mm)

Final thickness (mm)

Displacement (mm)

Normal Stresses (MPa)

i239sds1Moon5102030 i240sds1Moon53050 i249sds1Moon5102030 i250sds1Moon53050 i251sds1Moon30 i252sds1Moon30 i253s1Moon5075 i254sds11Moon30

single-direct single-direct single-direct single-direct single-direct single-direct double-direct single-direct

Footwall Footwall Hanging-wall Hanging-wall Hanging-wall Footwall Footwall Footwall

10.5697 13.628 13.948 13.972 13.039 9.948 – 10.136

3.7038 3.1382 6.0839 5.9646 4.4131 3.8418 – 7.18

15.91 14.937 16.019 14.543 9.2032 9.8184 12.648 2.0314

5,10,20,30,40,10 5,30,50 5,10,20,30,40 5,30,50 30 30 50,75 30

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Fig. 4. Experimental results. a) Example of experiment i249 performed in single-direct geometry at normal stresses between 5 and 40 MPa. Data show measured shear stress and normal stress and calculated layer thickness. The steady-state shear stress was measured at each normal stress step as an average of the data in the red boxes. b) Summary of shear stress vs. normal stress for all experiments. A linear best-fit to the data yields a friction coefficient of 0.19 for the footwall muscovite-rich foliated cataclasites and 0.24 for the hanging-wall chloriterich foliated cataclasites. The field of “Byerlee” friction with 0.6 < μ < 0.85 is shown for reference. c) Backscatter SEM image of deformed hanging-wall sample i250. Inset shows location of figure parts c and d. The sample is cut by a series of R1 Riedel shears with displacements of up to 1 mm. d) Detail of sample margin where grooves are preserved shows that localization did not occur along a discrete boundary shear. e) Cataclasis of quartz and albite grains, as well as other non-phyllosilicate phases, occurred during the experiments, especially along Riedel shears. Red star highlights an aggregate of chlorite and muscovite that was reworked and rotated during the experiment. f) Grains of chlorite show folding and buckling as well as layer delamination. Red star highlights another aggregate of chlorite and muscovite. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

allowed to saturate for approximately 1 h, during which time a slight decrease in wafer thickness due to compaction was recorded. After compaction had stopped, normal stress was increased to the desired value and shearing was initiated in the sample. Forces were measured

wafers sandwiched between side blocks and the central block (Fig. 3d inset). After preparation of the sample assembly, a small normal load was applied to the sample (c. 1 MPa) and water was added to a flexible plastic membrane surrounding the sample (Fig. 3d). Samples were 39

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normal stress (without confining pressure) on seawater-saturated samples. It is also slightly lower than values (μ = 0.27–0.32) reported by Ikari et al. (2009). Moore and Lockner (2015) determined that the frictional strength of chlorite is a function of composition and report that Fe-rich chlorite (< 21.6 wt% FeO in their experiments) is relatively strong (μ in the range of 0.26–0.36) compared to Mg- and Si-rich chlorite (μ in the range of 0.16–0.22). Electron microprobe analysis of chlorite grains from hanging-wall foliated cataclasites in the MFZ shows that the chlorite is Fe-rich and has FeO contents of up to 31 wt% (Supplementary Materials). Our measured frictional strengths are slightly lower than the Fe-rich chlorites of Moore and Lockner (2015). Our friction coefficient for muscovite-rich foliated cataclasites (μ = 0.19; Fig. 5b) is notably lower than all published data collected in water-present conditions, with the exception of the single-crystal measurement of Kawai et al. (2015). Ikari et al. (2011) also reported a friction coefficient of 0.16 for muscovite schist deformed in room-humidity conditions. The muscovite-rich foliated cataclasites from the MFZ show thorough-going networks of phyllosilicates that are composed of muscovite with variable amounts of chlorite and only minor amounts of non-phyllosilicate material (Fig. 2d and e). One possibility to explain our relatively low value of muscovite friction is that sliding in our experiments occurred mainly along well-aligned muscovite (and chlorite) basal (001) planes within the sample matrix, so that the frictional strength approaches that of single crystals. The weakening effect of a through-going foliation has previously been demonstrated in several studies (Collettini et al., 2009a; Ikari et al., 2011; Carpenter et al., 2012; Tesei et al., 2012, 2015) that measured significantly lower (up to 65%) friction coefficients for intact, phyllosilicate-rich samples compared to their powdered equivalents, consistent with our results for muscovite. Our frictional strength measurements for chlorite are only slightly lower than published values (excluding the Mg-rich chlorites of Moore and Lockner (2015)), raising the question as to why the weakening effect of a through-going foliation is not as pronounced as in the muscovite-rich samples. The natural chlorite-rich foliated cataclasites from the MFZ show an interconnected matrix (Fig. 2b and c) similar to the muscovite-rich foliated cataclasites. However, unlike the muscoviterich samples, the chlorite matrix is loaded with secondary phases (e.g. titanite, magnetite, Fe-oxides and epidote) that are disseminated along the platy grain boundaries of chlorite (Fig. 2b and c). Optical and SEM images indicate that such non-phyllosilicate phases may locally comprise up to 20–30 wt% of the matrix. The influence of these disseminated secondary phases during experimental shearing is not known, but given their abundance it seems possible that disruption of platy chlorite grains and grain-to-grain contacts occurred during the experiments. If this is true, the weakening effect of an interconnected chlorite-rich foliation could be reduced.

with stainless steel load cells (+-0.03 kN) and displacements were measured with LDVT's (+-0.1 μm) attached to each piston. In each experiment, the sample wafer was sheared at the lowest desired normal stress until sliding at “steady-state” shear stress was noted, typically after 3–4 mm of displacement (Fig. 4a). Shearing was then stopped, the normal stress was increased, and then shearing was initiated again. This was repeated at several different normal stresses until a maximum displacement of c. 16 mm was reached (Table 1). At each normal stress, the steady-state shear stress was measured as an average value of data towards the end of each normal stress step (i.e. in the red boxes in Fig. 4a). At the highest displacements, progressive slip hardening was often observed (Fig. 4a). In such cases, the steady-state shear stress was measured as an average value of data in the first half of the normal stress step (i.e. red box at 40 MPa in Fig. 4a). 4.2. Results Fig. 4b summarizes mechanical data for all experiments performed on the footwall muscovite-rich (black data) and hanging-wall chloriterich (green data) foliated cataclasites over a range of normal stresses between 5 and 75 MPa. The shear stress measured during steady-state sliding shows a linear relationship with normal stress consistent with a Mohr-Coulomb pressure-dependent failure envelope. The frictional strength, μ, calculated as the slope of the best-fit lines in shear stress vs. normal stress plots, is c. 0.19 for footwall foliated cataclasites and 0.24 for hanging-wall foliated cataclasites. Deformed wafers from the hanging-wall and footwall samples show similar microstructures. Strain was localized within regularly-spaced R1-type Reidel shears (white arrows in Fig. 4c; Logan et al., 1979) that have displacements of up to c. 1 mm and deflect the sample foliation (Fig. 4c). Samples preserving an imprint of the teeth from the steel blocks show that strain was not localized to boundary shears at the margins of the sample wafers (Fig. 4d). In the bulk of the wafers, deformation was broadly distributed and occurred by two main mechanisms: 1) quartz-albite and other resistant (e.g. titanite, oxides) grains were intensely fractured, particularly close to Reidel shears where fractured grain aggregates were smeared along the Reidel shear surfaces (Fig. 4e); 2) matrix phyllosilicates (Fig. 4e and f) were folded and buckled, probably resulting from frictional sliding within platy layers accompanied by crystal delamination. The matrix also contains subrounded aggregates of phyllosilicates (marked by red stars in Fig. 4e and f) that were reworked and rotated during shearing. 5. Discussion 5.1. Comparison with published friction data

5.2. Frictional strength and reactivation of high-angle reverse faults

Fig. 5 shows a compilation of published friction coefficients for chlorite (Fig. 5a) and muscovite (Fig. 5b) deformed under water-present conditions over a wide range of effective normal stresses, confining pressures, grain sizes and loading rates. Published data in Fig. 5 were acquired at room temperature unless indicated otherwise, and all experiments were performed with pure gouges (i.e. 100 wt% chlorite or muscovite) except Ikari et al. (2009) who used a powdered schist containing 46 wt% chlorite. The study on muscovite by Kawai et al. (2015) includes data on both powders and single-crystals deformed under water-saturated conditions at a normal stress of 25 MPa (without confining pressure). There is substantial variation in the frictional strength of chlorite and muscovite between the different studies (Fig. 5), which suggests the need to more carefully control and report water contents, mineral compositions and other experimental parameters to allow for a reliable comparison between datasets (Moore and Lockner, 2007). Our measured friction coefficient for chlorite-rich foliated cataclasites (μ = 0.24; Fig. 5a) is slightly lower than the value (μ = 0.26) of Brown et al. (2003) who conducted direct-shear tests at 5–40 MPa

Our field, microstructural and experimental results demonstrate that the core of the MFZ is dominated by a sequence of foliated phyllosilicate-rich cataclasites that are frictionally weak. The foliated cataclasites are regionally-significant in that they are present in all studied exposures of the main fault trace (Alder et al., 2016). Available constraints indicate that the foliated cataclasites are interconnected within the fault core over distances of at least hundreds of metres, and probably up to several kilometres. The frictional and hydrological properties of the foliated cataclasites are therefore likely to have played an important role in accommodating displacement during high-angle reverse faulting. The continuous nature of the phyllosilicate-rich foliation is likely to result from pressure-solution, which is clearly an important deformation mechanism in the natural fault rock samples from the MFZ. Experiments designed to study the influence of pressure solution in phyllosilicate-bearing fault rocks have repeatedly shown a pronounced weakening effect compared to purely frictional strength (e.g. Bos et al., 40

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Fig. 5. Compilation of published friction coefficients for a) chlorite and b) muscovite. Experiments were performed under water-present conditions at a range of normal stresses, confining pressures, grain sizes and loading rates (Scruggs and Tullis, 1998; Morrow et al., 2000; Moore and Lockner, 2004; Mariani et al., 2006; Van Diggelen et al., 2010; Behnsen and Faulkner, 2012; den Hartog et al., 2013).

reasonable levels of differential stress if the fluid pressure is very high (i.e. λv is close to 1 or exceeds σ3). If these conditions are not met, a new fault with a more optimal orientation should instead form. However, the frictional strength of fault rock samples from the MFZ is sufficiently low that the reshear field broadens substantially and reactivation at dip angles of 60°–75° is permissible even at hydrostatic (λv = 0.4) fluid pressures (Fig. 6b). Field evidence in the form of fault-related hydraulic extension veins indicates that high fluid pressures are transiently developed within and surrounding low-to-moderate displacement high-angle reverse faults (Sibson et al., 1988; Cox, 1995; Turner and Williams, 2004). In regions of active inversion such as NE Honshu, there is also a range of geophysical data (summarized in Sibson (2009)) that can be interpreted as indicating an over-pressured middle to lower crust in regions where recent high-angle reverse fault ruptures have occurred. These lines of evidence suggest that episodically and heterogeneously developed fluid pressure is an important mechanism of fault zone weakening in areas of high-angle reverse faulting. Extensive vein networks are generally absent from exposures of the core of the MFZ. Though this cannot be used to discount the possibility that high fluid pressures developed during faulting, the experimental results suggest that low frictional strength is a plausible weakening mechanism in basin-scale reverse faults, and that it may act in concert with high fluid pressure to allow reverse fault reactivation at low differential stresses.

2000; Niemeijer and Spiers, 2005; Niemeijer et al., 2008; Fagereng and den Hartog, 2017). This is related to dissolution of relatively strong mineral phases and their replacement by weak and relatively insoluble phyllosilicates, a process that has been documented within the cores of many mid-to upper-crustal fault zones (e.g. Gratier and Gamond, 1990; Wintsch et al., 1995; Holdsworth, 2004; Jefferies et al., 2006; Imber et al., 2008; Bistacchi et al., 2012; Tesei et al., 2013; Wallis et al., 2015). Once the phyllosilicate-rich network becomes established, frictional sliding between the platy phyllosilicate grains can control fault strength (e.g. Collettini et al., 2009b). In our experiments the natural fault rock foliation remains relatively continuous for the duration of the tests. Our frictional strength measurements may therefore represent an “upper bound” for the long-term shear strength of the MFZ, in which the continuity of phyllosilicate-rich layers is maintained by pressuresolution processes. The implications of low frictional strength on the mechanics of highangle reverse faulting can be considered by carrying out two-dimensional reactivation analysis. Following Sibson (1985, 1989) the differential stress required for frictional reactivation of a cohesionless reverse fault can be found as: (σ1 - σ3) = μs[(tan θr + cot θr)/(1-μs tan θr)] ρgz(1-λv)

(1)

where μs is the coefficient of friction, θr is the fault dip angle, ρ is average rock density (density of schists surrounding MFZ is 2730 kg/m3 after Tenzer et al. (2011)), g is gravitational acceleration, z is depth and λv is the ratio of fluid to overburden pressure. Fig. 6 shows plots of the differential stress required for reactivation of a reverse fault at a depth of 5 km for different values of λv. Fig. 6a shows the case of μs = 0.6, at the lower end of Byerlee's (1978) range of rock friction, whereas Fig. 6b adopts a representative value of μs = 0.2 determined from our experiments. The observed dip range of the MFZ (60°–75°) is shown as the grey shaded region. Also shown as dashed lines is the shear strength of intact rock with a tensile strength of 10 MPa, calculated on the assumption of a coefficient of internal friction of 0.75 and a cohesive strength equivalent to twice the tensile strength (Sibson, 1989, 2009). The reshear field represents the region in which reactivation of a preexisting fault will occur in preference to shear failure of intact rock. Sibson (1985) has pointed out that reactivation of high-angle reverse faults possessing “Byerlee” friction (Fig. 6a) is only possible at

6. Conclusions The > 300 km long Moonlight Fault Zone was an Oligocene basinbounding normal fault that reactivated in the Miocene as a high-angle reverse fault (present-day dip angle 60°–75°). Friction experiments on intact water-saturated samples of chlorite- and muscovite-rich foliated cataclasites from the fault core yielded friction coefficients of 0.19–0.24. If the laboratory friction coefficients are representative of the shear strength of the Moonlight Fault Zone, reactivation analysis indicates that frictional weakening could have facilitated high-angle reverse faulting under hydrostatic fluid pressures. Our results indicate that frictional weakening can occur along high-angle reverse faults at basement depths and that it may operate together with transiently high fluid pressures to promote reverse fault reactivation during basin 41

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Pf >

3

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Fig. 6. Plots of differential stress (σ1- σ1) required for frictional reactivation of cohesionless faults at a depth of 5 km lying at angle θr to the maximum principal stress, for different values of the pore fluid factor λv (after Sibson, 1985, 2007, 1998). Grey shaded region is the observed dip range of the Moonlight Fault Zone (MFZ). Dashed lines approximate the shear strength of intact rock with a tensile strength of 10 MPa (see text for details). a) Results for a “Byerlee” friction coefficient of 0.6 and b) a representative friction coefficient of 0.2 from our laboratory measurements.

inversion.

interconnected talc networks and weakening of continental low-angle normal faults. Geology 37, 567–570. Cox, S.F., 1995. Faulting processes at high fluid pressures - an example of fault valve behavior from the Wattle Gully fault, Victoria, Australia. J. Geophys. Res. Solid Earth 100, 12841–12859. Craw, D., 1985. Structure of schist in the Mt aspiring region, northwestern Otago, NewZealand. N. Z. J. Geol. Geophys. 28, 55–75. den Hartog, S.A.M., Niemeijer, A.R., Spiers, C.J., 2013. Friction on subduction megathrust faults: beyond the illite-muscovite transition. Earth Planet. Sci. Lett. 373, 8–19. Dore, A.G., Lundin, E.R., 1996. Cenozoic compressional structures on the NE Atlantic margin: nature, origin and potential significance for hydrocarbon exploration. Pet. Geosci. 2, 299–311. Fagereng, A., den Hartog, S.A.M., 2017. Subduction megathrust creep governed by pressure solution and frictional-viscous flow. Nat. Geosci. 10, 51–57. Gratier, J.P., Gamond, J.F., 1990. Transition between seismic and aseismic deformation in the upper crust. Deformation mechanisms. Rheol. Tect. 54, 461–473. Holdsworth, R.E., 2004. Weak faults - rotten cores. Science 303, 181–182. Ikari, M.J., Niemeijer, A.R., Marone, C., 2011. The role of fault zone fabric and lithification state on frictional strength, constitutive behavior, and deformation microstructure. J. Geophys. Res. Solid Earth 116. Ikari, M.J., Saffer, D.M., Marone, C., 2009. Frictional and hydrologic properties of clayrich fault gouge. J. Geophys. Res. Solid Earth 114. http://dx.doi.org/10.1029/ 2008jb006089. Imber, J., Holdsworth, R.E., Smith, S.A.F., Jefferies, S.P., Collettini, C., 2008. Frictionalviscous flow, seismicity and the geology of weak faults: a review and future directions. In: Wibberley, C.A.J., Kurz, W., Imber, J., Holdsworth, R.E., Collettini, C. (Eds.), The Internal Structure of Fault Zones: Implications for Mechanical and Fluid Flow Properties. The Geological Society of London, London, pp. 151–173. Jefferies, S.P., Holdsworth, R.E., Wibberley, C.A.J., Shimamoto, T., Spiers, C.J., Niemeijer, A.R., Lloyd, G.E., 2006. The nature and importance of phyllonite development in crustal-scale fault cores: an example from the Median Tectonic Line, Japan. J. Struct. Geol. 28, 220–235. Kawai, K., Sakuma, H., Katayama, I., Tamura, K., 2015. Frictional characteristics of single and polycrystalline muscovite and influence of fluid chemistry. J. Geophys. Res. Solid Earth 120, 6209–6218. Litchfield, N.J., Norris, R.J., 2000. Holocene motion on the akatore fault, south Otago coast, New Zealand. N. Z. J. Geol. Geophys. 43, 405–418. Logan, J.M., Friedman, M., Higgs, N., Dengo, C., Shimamoto, T., 1979. Experimental studies of simulated gouge and their application to studies of natural fault zones. In: Proceedings of Conference VIII on Analysis of Actual Fault Zones in Bedrock, pp. 79–1239 U.S. Geological Survey Open-File Report. Mariani, E., Brodie, K.H., Rutter, E.H., 2006. Experimental deformation of muscovite shear zones at high temperatures under hydrothermal conditions and the strength of phyllosilicate-bearing faults in nature. J. Struct. Geol. 28, 1569–1587. Marques, F.O., Nogueira, C.R., 2008. Normal fault inversion by orthogonal compression: sandbox experiments with weak faults. J. Struct. Geol. 30, 761–766. Moore, D.E., Lockner, D.A., 2004. Crystallographic controls on the frictional behavior of dry and water-saturated sheet structure minerals. J. Geophys. Res. Solid Earth 109, B03401. http://dx.doi.org/10.1029/2003JB002582. Moore, D.E., Lockner, D.A., 2007. Friction of the smectite clay montmorillonite: a review

Acknowledgements This work was funded by the Marsden Fund Council (project UOO1417 to Smith) administered by the Royal Society of New Zealand. Matt Ikari is thanked for his review of the paper and Ian Alsop for his editorial handling. Appendix A. Supplementary data Supplementary data related to this article can be found at http://dx. doi.org/10.1016/j.jsg.2017.10.009. References Alder, S., Smith, S.A.F., Scott, J.M., 2016. Fault-zone structure and weakening processes in basin-scale reverse faults: the Moonlight Fault Zone, South Island, New Zealand. J. Struct. Geol. 91, 177–194. Barry, J., 1966. Structural Analysis in the Middle Shotover Valley, Northwest Otago. University of Otago. Behnsen, J., Faulkner, D.R., 2012. The effect of mineralogy and effective normal stress on frictional strength of sheet silicates. J. Struct. Geol. 42, 49–61. Bistacchi, A., Massironi, M., Menegon, L., Bolognesi, F., Donghi, V., 2012. On the nucleation of non-Andersonian faults along phyllosilicate-rich mylonite belts. In: Healy, D., Butler, R.W.H., Shipton, Z.K., Sibson, R.H. (Eds.), Faulting, Fracturing and Igneous Intrusion in the Earth's Crust. Geological Society of London, Special Publications, London, pp. 185–199. Bos, B., Peach, C.J., Spiers, C.J., 2000. Frictional-viscous flow of simulated fault gouge caused by the combined effects of phyllosilicates and pressure solution. Tectonophysics 327, 173–194. Brown, K.M., Kopf, A., Underwood, M.B., Weinberger, J.L., 2003. Compositional and fluid pressure controls on the state of stress on the Nankai subduction thrust: a weak plate boundary. Earth Planet. Sci. Lett. 214, 589–603. Butler, R.W.H., Tavarnelli, E., Grasso, M., 2006. Structural inheritance in mountain belts: an Alpine-Apennine perspective. J. Struct. Geol. 28, 1893–1908. Byerlee, J., 1978. Friction of rocks. Pure Appl. Geophys. 116, 615–626. Carpenter, B.M., Saffer, D.M., Marone, C., 2012. Frictional properties and sliding stability of the San Andreas fault from deep drill core. Geology 40, 759–762. Collettini, C., Di Stefano, G., Carpenter, B., Scarlato, P., Tesei, T., Mollo, S., Trippetta, F., Marone, C., Romeo, G., Chiaraluce, L., 2014. A novel and versatile apparatus for brittle rock deformation. Int. J. Rock Mech. Min. Sci. 66, 114–123. Collettini, C., Niemeijer, A., Viti, C., Marone, C., 2009a. Fault zone fabric and fault weakness. Nature 462, 907–910. Collettini, C., Viti, C., Smith, S.A.F., Holdsworth, R.E., 2009b. Development of

42

Journal of Structural Geology 105 (2017) 34–43

S.A.F. Smith et al.

redistribution through fault-valve action. In: Buchanan, J.G., Buchanan, P.G. (Eds.), Geological Society Special Publication No. 88. Basin Inversion Geological Society of London, London. Sibson, R.H., 1998. Brittle failure mode plots for compressional and extensional tectonic regimes. J. Struct. Geol. 20, 655–660. Sibson, R.H., 2007. An episode of fault-valve behaviour during compressional inversion? The 2004 M(J)6.8 Mid-Niigata Prefecture, Japan, earthquake sequence. Earth Planet. Sci. Lett. 257, 188–199. Sibson, R.H., 2009. Rupturing in overpressured crust during compressional inversion - the case from NE Honshu, Japan. Tectonophysics 473, 404–416. Sibson, R.H., Robert, F., Poulsen, K.H., 1988. High-angle reverse faults, fluid-pressure cycling, and mesothermal gold-quartz deposits. Geology 16, 551–555. Sibson, R.H., Xie, G.Y., 1998. Dip range for intracontinental reverse fault ruptures: truth not stranger than friction? Bull. Seismol. Soc. Am. 88, 1014–1022. Tenzer, R., Sirguey, P., Rattenbury, M.S., Nicolson, J., 2011. A digital rock density map of New Zealand. Comput. Geosci. 37, 1181–1191. Tesei, T., Collettini, C., Carpenter, B.M., Viti, C., Marone, C., 2012. Frictional strength and healing behavior of phyllosilicate-rich faults. J. Geophys. Res. Solid Earth 117. Tesei, T., Collettini, C., Viti, C., Barchi, M.R., 2013. Fault architecture and deformation mechanisms in exhumed analogues of seismogenic carbonate-bearing thrusts. J. Struct. Geol. 55, 167–181. Tesei, T., Lacroix, B., Collettini, C., 2015. Fault strength in thin-skinned tectonic wedges across the smectite-illite transition: constraints from friction experiments and critical tapers. Geology 43, 923–926. Townend, J., Sherburn, S., Arnold, R., Boese, C., Woods, L., 2012. Three-dimensional variations in present-day tectonic stress along the Australia–Pacific plate boundary in New Zealand. Earth Planet. Sci. Lett. 353–354, 47–59. Townend, J., Zoback, M.D., 2000. How faulting keeps the crust strong. Geology 28, 399–402. Turnbull, I.M., Barry, J.M., Carter, R.M., Norris, R.J., 1975. The Bobs Cove Beds and their relationship to the Moonlight Fault zone. J. R. Soc. N. Z. 5, 355–394. Turner, J.P., Williams, G.A., 2004. Sedimentary basin inversion and intra-plate shortening. Earth-Sci. Rev. 65, 277–304. Van Diggelen, E.W.E., De Bresser, J.H.P., Peach, C.J., Spiers, C.J., 2010. High shear strain behaviour of synthetic muscovite fault gouges under hydrothermal conditions. J. Struct. Geol. 32, 1685–1700. Wallis, D., Lloyd, G.E., Phillips, R.J., Parsons, A.J., Walshaw, R.D., 2015. Low effective fault strength due to frictional-viscous flow in phyllonites, Karakoram Fault Zone, NW India. J. Struct. Geol. 77, 45–61. Warren-Smith, E., Lamb, S., Stern, T.A., 2017. Stress field and kinematics for diffuse microseismicity in a zone of continental transpression, South Island, New Zealand. J. Geophys. Res. Solid Earth 122, 2798–2811. Wintsch, R.P., Christoffersen, R., Kronenberg, A.K., 1995. Fluid-rock reaction weakening of fault zones. J. Geophys. Res. Solid Earth 100, 13021–13032.

and interpretation of data. In: Dixon, T., Moore, C. (Eds.), The Seismogenic Zone of Subduction Thrust Faults. Columbia University Press, pp. 317–345. Moore, D.E., Lockner, D.A., 2015. Correlation of Chlorite Frictional Strength with Composition. American Geophysical Union (AGU), San Francisco Fall Meeting, 2015, abstract #MR21C-2628. Morrow, C.A., Moore, D.E., Lockner, D.A., 2000. The effect of mineral bond strength and adsorbed water on fault gouge frictional strength. Geophys. Res. Lett. 27, 815–818. Mortimer, N., 2000. Metamorphic discontinuities in orogenic belts: example of the garnet–biotite–albite zone in the Otago Schist, New Zealand. Int. J. Earth Sci. 89, 295–306. Niemeijer, A.R., Spiers, C.J., 2005. Influence of phyllosilicates on fault strength in the brittle-ductile transition: insights from rock analogue experiments. In: Bruhn, R., Burlini, L. (Eds.), High Strain Zones: Structure and Physical Properties. The Geological Society of London, London, pp. 303–327. Niemeijer, A.R., Spiers, C.J., Peach, C.J., 2008. Frictional behaviour of simulated quartz fault gouges under hydrothermal conditions: results from ultra-high strain rotary shear experiments. Tectonophysics 460, 288–303. Norris, R.J., Carter, R.M., 1982. Fault-bounded blocks and their role in localising sedimentation and deformation adjacent to the alpine fault, southern New Zealand. Tectonophysics 87, 11–23. Norris, R.J., Carter, R.M., Turnbull, I.M., 1978. Cainozoic sedimentation in basins adjacent to a major continental transform boundary in southern New Zealand. J. Geol. Soc. 135, 191–205. Norris, R.J., Koons, P.O., Cooper, A.F., 1990. The obliquely-convergent plate boundary in the South Island of New Zealand: implications for ancient collision zones. J. Struct. Geol. 12, 715–725. Reilly, C., Nicol, A., Walsh, J.J., Seebeck, H., 2015. Evolution of faulting and plate boundary deformation in the Southern Taranaki Basin, New Zealand. Tectonophysics 651, 1–18. Reyners, M., Ingham, C.E., Ferris, B.G., 1983. Microseismicity of the upper Clutha valley, South Island, New Zealand. New Zeal J. Geol. Geop 26, 1–6. Ruh, J.B., Verges, J., 2017. Effects of reactivated extensional basement faults on structural evolution of fold-and-thrust belts: insights from numerical modelling applied to the Kopet Dagh Mountains. Tectonophysics. http://dx.doi.org/10.1016/j.tecto.2017. 05.020. Scholz, C.H., Rynn, J.M.W., Weed, R.W., Frohlich, C., 1973. Detailed seismicity of the Alpine Fault zone and fiordland region, New Zealand. Geol. Soc. Am. Bull. 84, 3297–3316. Scruggs, V.J., Tullis, T.E., 1998. Correlation between velocity dependence of friction and strain localization in large displacement experiments on feldspar, muscovite and biotite gouge. Tectonophysics 295, 15–40. Sibson, R.H., 1985. A note on fault reactivation. J. Struct. Geol. 7, 751–754. Sibson, R.H., 1989. High-angle reverse faulting in northern New Brunswick, Canada, and its implications for fluid pressure levels. J. Struct. Geol. 11, 873–877. Sibson, R.H., 1995. Selective fault reactivation during basin inversion: potential for fluid

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