# The effect of micas on the strength of experimental granitoid fault gouge

W. Zhan, A.R. Niemeijer, A. Berger, C.J. Spiers, F.E. Gfeller, M. Herwegh

► To cite this version:

W. Zhan, A.R. Niemeijer, A. Berger, C.J. Spiers, F.E. Gfeller, et al. The effect of micas on the strength of experimental granitoid fault gouge. Tectonophysics, 2025, 908, pp.230772. 10.1016/j.tecto.2025.230772. insu-05242094

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# The effect of micas on the strength of experimental granitoid fault gouge

W. Zhan<sup>a,c,*</sup>, A.R. Niemeijer<sup>b</sup>, A. Berger<sup>a</sup>, C.J. Spiers<sup>b</sup>, F.E. Gfeller<sup>a</sup>, M. Herwegh<sup>a</sup>

<figure><img src="image_3.png" /></figure>

## ARTICLE INFO

### Keywords:

- Deformation

- Fault gouge

- Mica

- Dissolution precipitation creep

- Mineral reaction

- Continental crust

## ABSTRACT

Sheet silicates play an important role in shaping crustal rheology and shear localization at shallow depths through their low strength. However, their effect at deeper levels (>3 km) remains unclear. We conducted hydrothermal ring shear experiments on three simulated gouges with similar quartz content but varying mica types and contents. Applied temperatures *T* ranged from 20 to 650 °C, with sliding velocities *V* between 0.03 and 1 µm/s, and an effective normal stress and pore water pressure of 100 MPa. Shear strains up to 30 were attained.

At 1 µm/s and 20 °C, granitoids exhibit a higher friction coefficient (μ = 0.81) than the muscovite-rich (μ = 0.47) and biotite-rich gouges (μ = 0.44). With increasing *T* at a fixed *V* of 1 µm/s, μ of granitoids decreases to 0.61 at 650 °C, whereas muscovite- and biotite-rich gouges strengthen (μ = 0.56–0.69) until *T* reaches 200–450 °C. At 650 °C, weakening is observed in granitoid and muscovite-rich gouges as *V* decreases, while biotite-rich gouge shows no change. Granitoids are weaker than the mica-rich gouges once *V* decreases to 0.03 µm/s and *T* reaches 650 °C. All gouges at 650 °C exhibit µm-wide principal slip zones, featuring microstructures consistent with dissolution-precipitation creep (DPC). These include truncated grain contacts, mineral precipitates, submicrometer grain size and low porosity. Muscovite breakdown and biotite formation occurs in muscovite-rich gouge as they weaken. Our results imply that mica enrichment in granitoid faults can lead to a weaker upper crust and stronger middle crusts as they weaken less than granitoids. The onset of DPC may trigger a frictional-viscous transition at greenschist facies depths.

## 1. Introduction

The lithological composition of the Earth's crust is heterogeneous. Granitoid rocks, composed mainly of quartz and feldspar, are typically considered as the predominant lithology of the crystalline part of upper-to-middle continental crust (Baumberger et al., 2022; Wedepohl, 1995; Wehrens et al., 2016). Nevertheless, sheet silicates can be found over a broad range of geological conditions in the granitoid system, from the shallow subsurface to the upper-to-middle crust, especially in domains with active fluid-rock interactions. Sheet silicates form through fluid-assisted processes including weathering, sedimentation, diagenesis, crystallization from melt, and metamorphic reactions, like feldspar replacement reactions (Goncalves et al., 2012; Wedepohl, 1995; Hickman et al., 1995; Velde, 1995; Wintsch et al., 1995).

Sheet silicates play an important role in shaping crustal rheology and localization of deformation, owing to their weak {001} grain interfaces and their capacity to develop foliations characterized by well-aligned basal planes. Fault core materials, retrieved from the actively creeping

San Andreas Fault at a depth of 2.7 km, are rich in sheet silicates and exhibit a friction coefficient of μ = 0.15 under in-situ conditions (Carpenter et al., 2011; Lockner et al., 2011), much lower compared to μ = 0.65–0.8 commonly observed across a broad range of rock types (typically referred to as Byerlee's law; Byerlee, 1978). This observation reveals a strong correlation between the presence of sheet silicates and the weakness of the San Andreas Fault, as evidenced by a lack of pronounced heat flow peak (Lachenbruch and Sass, 1980). At greater depths (e.g. ~6 km; Zoback et al., 1993), where direct heat flow or in-situ stress measurements are not feasible, the strength of the crust cannot be directly inferred. Therefore, strengths of fault rocks measured in laboratory tests under near in-situ conditions are used as proxies to evaluate the crustal rheology in the mid-lower levels. In the classical crustal rheology profile, the deformation behavior beyond the frictional-viscous transition is commonly represented by the viscous creep of quartz (Bürgmann and Dresen, 2008), or more recently of granitoid rock (Pec et al., 2016). However, at upper-to-middle crustal levels (greenschist facies condition), deformation has often been observed to localize

<footnote><sup>a</sup> Institute of Geological Sciences, University of Bern, Bern, Switzerland</footnote>

<footnote><sup>b</sup> Department of Earth Sciences, Utrecht University, Utrecht, the Netherlands</footnote>

<footnote><sup>c</sup> Institut des Sciences de la Terre d'Orléans, Univ Orléans, CNRS, BRGM, OSUC, Orléans, France</footnote>

<footnote>* Corresponding author at: Institute of Geological Sciences, University of Bern, Bern, Switzerland.
E-mail address: weijia.zhan@unibe.ch (W. Zhan).</footnote>
within fine-grain, mica-rich shear zones (Imber et al., 1997; Rennie et al., 2013; Wehrens et al., 2016). Given the considerable impact and abundance of sheet silicates, it is essential to examine the potentially significant impacts of mica on the rheology of upper-to-middle crust comprised primarily of granitoids but overprinted by fluid-rock interactions.

Many laboratory studies have shown that micas like muscovite and biotite, belonging to the most common sheet silicates across the upper-to-middle continental crust, are frictionally weak with friction coefficients of μ = 0.25–0.4 under room temperature and wet conditions (Lu and He, 2014; Mares and Kronenberg, 1993; Mariani et al., 2006; Scruggs and Tullis, 1998). Mica can deform viscously through gliding along basal planes and kinking in single crystals (Kronenberg et al., 1990; Shea and Kronenberg, 1992), and migration of nano-scale riplocation defects, which is a dilatant process driven by elastic strain energy release (Aslin et al., 2019). Additionally, the presence of mica-quartz contacts has been shown to accelerate the dissolution of quartz (Greene et al., 2009; Meyer et al., 2006 and references therein), thereby influencing the rate of dissolution-precipitation creep (DPC). DPC is a viscous process that is considered to trigger the frictional-viscous transition (Bos and Spiers, 2001; Zhan et al., 2024) and facilitate strength recovery over geological timescales (Keulen et al., 2008; Nakatani and

Scholz, 2004; Niemeijer and Spiers, 2006). Several laboratory experiments have been conducted on mica-rich rocks to investigate their deformation behavior across a broad range of temperatures (T) and strain rates (ϵ). However, many of these studies either used pure mica as the starting material, which is not representative of deeper crustal lithology (Kronenberg et al., 1990; Lu and He, 2014; Mariani et al., 2006), or the tested experimental conditions fell largely outside the viscous regime, resulting in predominantly frictional behavior (den Hartog et al., 2013).

The strength, microstructures, and the underlying mechanisms of granitoid gouge have been discussed in detail in Zhan et al., 2024. This study adds to the previously published work, with the aim to investigate the effect of sheet silicates on the strength of gouge-filled fault zones, which are primarily of granitoid composition. We conducted sliding experiments on simulated biotite- and muscovite-rich gouges to large displacements (~15 mm, shear strain up to 30). The relevant tests on granitoid gouges from Zhan et al. (2024) have now been incorporated into this study for the purpose of comparison. The experimental conditions were set to span the frictional-viscous transitional conditions of the granitoid system, encompassing temperatures (20–650 °C), sliding velocities (0.03–1 µm/s), and an effective normal stress and pore fluid pressure of 100 MPa (equal to a total normal stress of 200 MPa). After

<figure><img src="image_4.png" /><figcaption>**Fig. 1.** (a) Simplified geological map of the central Swiss Alps to show the collection locations of samples described in this study (modified after Goncalves et al., 2012). Outcrop photos and backscatter micrographs showing (b)-(c) the granitoid ultramylonites (sample A), (d)-(e) biotite-rich fault gouges (sample B), and (g)-(h) muscovite-rich fault gouges (sample C). Qtz = quartz; fsp = feldspar; bt = biotite; mus/ill = muscovite/illite; ep = epidote; chl = chlorite; apa = apatite.</figcaption></figure>
presenting and analyzing our results, we discuss the implications of our findings for extrapolating them to natural conditions in the upper-to-middle continental crust.

## 2. Materials and methods

### 2.1. Sample preparation and mineralogical characterization

We collected rock samples, designated as samples A, B and C, from fault/shear zones in the Aar Massif and the Tavetsch Massif in the central Swiss Alps (Fig. 1, see GPS coordinates in Table A.1 in Appendix A). These natural fault rocks were chosen to cover a wide range of mineralogical compositions that can be found across the upper-to-middle crustal fault zones. Sample A is a granitoid ultramylonite from a ductile shear zone within the granodiorite country rock, drilled from the underground Nagra Grimsel Test Site (GTS) (Fig. 1b-c; see also Zhan et al., 2024). Sample B is a fault gouge rich in biotite, collected from a brittle fault zone within a metabasic dyke in the GTS (Fig. 1d-e). Sample C is a fault gouge predominantly rich in muscovite, collected from an outcrop of a brittle fault zone that cuts through a greenschist metamorphic granitoid gneiss (Fig. 1f-g).

We disaggregated the collected rock samples to produce non-cohesives gouge-like powders. Mineral compositions and grain sizes are summarized in Table 1. The procedures for sample preparation and compositional characterization are provided in subsequent paragraphs. Note that the term 'muscovite/illite' refers to the indistinguishable K-bearing sheet silicates due to the inseparable peaks of illite and muscovite in the X-ray Powder Diffraction (XRD) patterns. The three tested samples consist of comparable quartz contents, but they have different proportions of feldspar and sheet silicates. Sample A has a granitoid composition and represents a 'fresh' greenschist metamorphic mylonite of the primary felsic intrusive rocks, which is abundant in the upper-to-middle continental crust. In comparison to sample A, sample B has lower feldspar content (4 wt%) and higher biotite content (37 wt%). Energy-dispersive X-ray (EDS) analyses show that the Mg/(Mg + Fe) ratios in samples A, B and C are 0.25, 0.58 and 0.43, respectively. The EDS results reflect that the biotite in sample B exhibits a Mg-rich composition. Given that biotite is a group of sheet silicates that remains stable at high temperatures (for Fe-rich the temperature is >400 °C, for Mg-rich it is >500 °C) (Fleet et al., 2003) and the relatively high Mg/(Mg + Fe) ratio, we interpret the sample B to represent 'high-grade' metamorphic rocks. As for sample C, it consists of the highest proportions of muscovite/illite and chlorite among all samples tested. Since these sheet silicates are common in low-to-medium grade metamorphic rocks, sample C is considered to represent the 'low-grade' metamorphic rock. For clarity, in this study, samples A, B and C are referred to as granitoid, biotite-rich and muscovite-rich gouges, respectively.

In terms of sample disaggregation, we used a step-wise procedure in

<caption>**Table 1**</caption>

The mineral compositions and grain sizes of the starting materials (non-cohesives powders). See the grain size distribution of the three samples in Fig. A.1 in the Appendix A.

<table><thead><tr><th>Sample names</th><th>Mineral composition (wt%)</th><th>Grain size (µm) median ± standard deviation</th></tr></thead><tbody><tr><td>A</td><td>37 % quartz, 11 % orthoclase, 38 % albite, 8 % biotite, 6 % epidote, trace amount of pyrite</td><td>45 µm ± 47 µm</td></tr><tr><td>B</td><td>35 % quartz, 2 % orthoclase, 2 % albite, 37 % Mg-rich biotite, 21 % muscovite/illite, 3 % smectite, trace amounts of epidote, zircon, apatite and titanium oxides</td><td>28 µm ± 85 µm</td></tr><tr><td>C</td><td>38 % quartz, 1 % orthoclase, 4 % albite, 50 % muscovite/illite, 6 % chlorite, 1 % apatite, and trace amounts of iron oxides as well as biotite</td><td>10 µm ± 65 µm</td></tr></tbody></table>

order to maintain the original mineral proportions and grain sizes of the collected rock samples. The ultramylonite sample (sample A) was subjected to electrical fragmentation in our in-house SELFRAG Lab (high voltage pulse machine, https://www selfrag com/; Giese et al., 2010; Zwingmann et al., 2017) and milling (by hand and in an XRD-Mill McCrone corundum ball mill) to obtain powders with a grain size fraction smaller than 125 µm (see details in Zhan et al., 2024). The fault gouges of samples B and C were treated in an ultrasonic bath with ethanol for ~30 min. The remaining larger fragments were manually ground with a mortar and pestle, and subsequently wet sieved to obtain grain size fractions smaller than 300 µm for sample B, and smaller than 250 µm for sample C. All disaggregated materials were dried at ~40 °C for at least 48 h, and stored in sealed containers at room temperature and ambient humidity.

We determined the mineral compositions with XRD, using a PANalytical EMPYREAN diffractometer equipped with a PANalytical 1Der detector and a Cu-X-Ray tube (Cu Kα) running at 40 mA and 45 kV. For the determination of the main mineral content, 20 wt% corundum was added as an internal standard to quantify modal abundances. The sample-corundum mixtures were ground in a McCrone XRD-Mill and prepared in a sample holder for Bragg-Brentano geometry. For a clear differentiation of the clay-sized minerals (illite, chlorite, smectite, and kaolinite), NH₄OH(aq) 0.01 mol/L solution was added to the starting material, which was brought into suspension in an ultrasonic bath. Following this, the clay-size fraction (< 2 µm) was separated by allowing the suspension to settle in an Atterberg cylinder. Each sample was subjected to the following preparations: one set was air-dried, one was annealed at 550 °C for one hour, and one was immersed in ethylene glycol overnight. The phase identification was performed using PANalytical HighscorePlus 4.9 with the diffraction database ICDD4+. The mineral content was quantified using the Rietveld refinement method (Rietveld, 1969), using TOPAS-Academic v. 6 software (Coelho, 2018). The relative proportions of the four clay end-members were determined using the in-house software ClayFit3 on calculated clay mineral diffractograms created in NEWMOD2 (Mertens et al., 2016). Accessory minerals were identified using scanning electron microscopy.

### 2.2. Experimental techniques and data processing

We conducted all rock deformation experiments in a hydrothermal ring shear apparatus (see details in Zhan et al., 2024). In each experiment, ~0.6 g of simulated gouge powders were sandwiched between two René-41® superalloy pistons, resulting in a layer ~1.5 mm thick. We subsequently placed the sample-piston assembly inside a pressure vessel within an Instron loading frame, and gradually applied effective normal stress to the desired value using the Instron ram. The effective normal stress was transmitted directly to the sample through a pressure-compensated loading piston using the Instron ram. Simultaneously, we filled the pressure vessel with deionized water as pore fluid using a servo-controlled pump, and heated its internal assembly using a coiled thermocoax furnace. After equilibrating for at least 45 min, the rotary driving system was switched on to start the sliding test. During experiments, a broad range of signals, including displacement, torque, temperature, axial displacement and load, pore fluid pressure, and rotational velocity, were continuously recorded using an 18-bit A/D converter and an acquisition rate of 900 Hz. The signals were then averaged and logged at frequencies of 1–100 Hz, depending on velocity. Upon completion of the sliding tests, the temperature and pressure within the apparatus were reduced to room conditions within ~30 min. The sheared materials were then retrieved from the apparatus and dried overnight in an oven at ~50 °C. In some experiments, samples were maintained at the designated temperature and pressure for several hours after deformation, prior to unloading. The thermal hold duration is listed as t<sub>hold</sub> in Table 2. This scenario typically occurred in the evening, when unloading of sheared samples was impractical.

We performed experiments at temperature (T) in the range of
<caption>**Table 2**</caption>

List of experiments, conditions and key data. T=temperature; D=displacement; V=sliding velocity; τyd=shear stress at the yield point; τfin=shear stress at the end of the test. All experiments were conducted under an effective normal stress and pore fluid pressure of 100 MPa (a total normal stress of 200 MPa). Mean shear stress and its standard deviation (std) are calculated for repeated experiments.

<table><thead><tr><th>Number</th><th>Sample</th><th>T (°C)</th><th>D (mm)</th><th>V (μm/s)</th><th>t<sub>hold</sub> (h)</th><th>τ<sub>yd</sub> (MPa)</th><th>τ<sub>fin</sub> (MPa)</th><th>τ<sub>fin</sub> - τ<sub>yd</sub> (MPa)</th><th>Mean shear stress ± standard deviation</th></tr></thead><tbody><tr><td colspan="10"><strong>Group I: standard-shear experiments at 1 μm/s</strong></td></tr><tr><td>u1038</td><td>A</td><td>20</td><td>15.98</td><td>1</td><td>10.0</td><td>68.97</td><td>81.11</td><td>12.14</td><td></td></tr><tr><td>u908</td><td>A</td><td>200</td><td>15.60</td><td>1</td><td>0.0</td><td>56.60</td><td>75.41</td><td>18.81</td><td></td></tr><tr><td>u909</td><td>A</td><td>450</td><td>15.10</td><td>1</td><td>0.1</td><td>57.79</td><td>73.04</td><td>15.25</td><td></td></tr><tr><td>u913</td><td>A</td><td>650</td><td>15.60</td><td>1</td><td>12.0</td><td>64.34</td><td>49.78</td><td>-14.56</td><td></td></tr><tr><td>u1154</td><td>A</td><td>650</td><td>16.25</td><td>1</td><td>0.0</td><td>59.61</td><td>67.83</td><td>8.22</td><td>τ<sub>yd</sub>: 61.99 ± 2.37</td></tr><tr><td>u1060</td><td>A</td><td>650</td><td>17.10</td><td>1</td><td>41.9</td><td>62.03</td><td>63.94</td><td>1.91</td><td>τ<sub>fin</sub>: 60.52 ± 9.5 (std/mean = 16 %)</td></tr><tr><td>u1058</td><td>B</td><td>20</td><td>17.10</td><td>1</td><td>0.1</td><td>34.52</td><td>43.55</td><td>9.03</td><td></td></tr><tr><td>u1054</td><td>B</td><td>200</td><td>17.10</td><td>1</td><td>12.0</td><td>33.08</td><td>45.69</td><td>12.61</td><td></td></tr><tr><td>u1053</td><td>B</td><td>450</td><td>17.10</td><td>1</td><td>7.6</td><td>39.09</td><td>55.61</td><td>16.52</td><td></td></tr><tr><td>u1101</td><td>B</td><td>650</td><td>16.54</td><td>1</td><td>13.5</td><td>42.67</td><td>54.61</td><td>11.94</td><td></td></tr><tr><td>u1141</td><td>C</td><td>20</td><td>16.10</td><td>1</td><td>14.7</td><td>36.64</td><td>46.59</td><td>9.95</td><td></td></tr><tr><td>u1095</td><td>C</td><td>200</td><td>16.89</td><td>1</td><td>0.0</td><td>42.95</td><td>66.23</td><td>23.28</td><td></td></tr><tr><td>u1094</td><td>C</td><td>450</td><td>15.71</td><td>1</td><td>0.2</td><td>48.95</td><td>68.59</td><td>19.55</td><td></td></tr><tr><td>u1099</td><td>C</td><td>650</td><td>16.69</td><td>1</td><td>0.8</td><td>58.26</td><td>68.36</td><td>10.10</td><td></td></tr><tr><td colspan="10"><strong>Group II: slow-shear experiments at 0.03-0.1 μm/s</strong></td></tr><tr><td>u910</td><td>A</td><td>650</td><td>19.80</td><td>0.1</td><td>9.6</td><td>60.09</td><td>37.27</td><td>-22.82</td><td>τ<sub>yd</sub>: 65.68 ± 6.77</td></tr><tr><td>u1150</td><td>A</td><td>650</td><td>13.10</td><td>0.1</td><td>0.0</td><td>73.21</td><td>49.70</td><td>-23.51</td><td>τ<sub>fin</sub>: 41.68 ± 6.22</td></tr><tr><td>u1175</td><td>A</td><td>650</td><td>15.49</td><td>0.1</td><td>2.0</td><td>63.73</td><td>43.08</td><td>-20.65</td><td>(std/mean = 15 %)</td></tr><tr><td>u1234</td><td>B</td><td>650</td><td>5.59</td><td>0.1</td><td>0.0</td><td>60.00</td><td>56.32</td><td>-3.68</td><td></td></tr><tr><td>u916</td><td>C</td><td>650</td><td>5.00</td><td>0.1</td><td>0.1</td><td>57.11</td><td>54.49</td><td>-2.62</td><td>τ<sub>yd</sub>: 74.03 ± 23.92</td></tr><tr><td>u1227</td><td>C</td><td>650</td><td>15.26</td><td>0.1</td><td>0.3</td><td>90.94</td><td>67.48</td><td>-23.46</td><td>τ<sub>fin</sub>: 60.99 ± 9.19</td></tr><tr><td>u1235</td><td>C</td><td>650</td><td>2.14</td><td>0.03</td><td>0.1</td><td>75.30</td><td>54.93</td><td>-20.37</td><td>(std/mean = 15 %)</td></tr></tbody></table>

20–650 °C, using velocity (V) in the range of 0.03–1 µm/s (Table 2). Experiments were subjected to an effective normal stress ($\sigma_{n}^{eff}$) of 100 MPa and pore fluid pressure ($P_{f}$) of 100 MPa, resulting in a total normal stress of 200 MPa ($\sigma_{n}^{eff} + P_{f}$). Experiments are subdivided into two groups based on the applied V: (I) standard-shear experiments, and (II)

slow-shear experiments. In the standard-shear experiments, gouges were sheared at a fixed velocity of 1 µm/s (a strain rate of ~$10^{-3}$ s⁻¹) and T ranging from 20 °C to 650 °C for ~15 mm displacement (a shear strain $\gamma$ up to 30). Based on the results from these experiments, we performed an additional set of slow-shear experiments at 0.03 or 0.1 µm/s (a strain rate of ~$3 \times 10^{-5}$ or $10^{-4}$ s⁻¹) and 650 °C to investigate the effect of

<figure><img src="image_5.png" /><figcaption>**Fig. 2.** Plots of shear stress (top row) and normalized layer thickness (bottom row) as a function of displacements for all samples tested under constant velocities of 1 µm/s but various temperatures. The three colors blue (a & d), orange (b & e) and purple (c & f) represent samples A, B and C, respectively. See Table 1 for a list of conditions applied in these experiments. Red labels in (a) show that yield-point ($\tau_{yd}$) and final shear stress ($\tau_{fin}$) are averaged measurements taken over a 1 mm displacement interval. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>
strain rate on gouge strength.

We calculated the shear stresses using externally measured torque data, and corrected them for pre-determined seal friction following methods described by Niemeijer et al. (2016). In each experiment, shear stresses at both the yield point ($\tau_{yd}$) and the final stage ($\tau_{fin}$) were determined by averaging measurements taken over a 1 mm displacement interval, beyond the yield point displacement value and the designated final displacement, respectively (Fig. 2a). The yield point is defined as the point at which the second derivative of the shear stress-displacement curve reaches its minimum. The evolution of gouge layer thickness during the sliding tests was determined from the position of the Instron loading ram. Due to minor material extrusion, the absolute layer thickness may vary substantially between different experiments. We normalized and rescaled the layer thickness with respect to their reference layer thickness (at start of shearing). This approach facilitates better comparison across different experiments by providing a consistent reference frame.

## 2.3. Microstructural analysis

After each experiment, we collected a ring-shaped gouge layer with dimensions of ~3 mm in width and ~ 0.5–0.7 mm in thickness. Deformed samples often broke apart into fragments along planes shal- lowly dipping in the direction of sliding. For microstructural in- vestigations, we air impregnated the fragments with epoxy resins, assuming their microstructures were representative of the entire sample. The impregnated samples were sectioned for scanning electron micro- scopy (SEM) imaging. Sectioning was performed in a plane normal to the shear plane and parallel to the shear direction, followed by a 10 nm thick carbon coating. Backscatter images were collected using a ZEISS EVO 50 SEM at 20 kV and 0.5–1.5 nA. Close-up images were obtained using a ZEISS Gemini FEG-SEM 450 operating at 5–10 kV and 100–300 pA. The un-impregnated fragments and loose powders were stored in airtight containers for XRD analyses to investigate possible changes in mineral composition due to deformation.

# 3. Results

## 3.1. Standard-shear experiments

### 3.1.1. Mechanical data

Fig. 2 shows the evolution of shear stress and normalized layer thickness as a function of displacement for all materials tested under constant low V (1 µm/s) and various T (20, 200, 450, 650 °C). Note that despite the high displacements, the majority of the experiments did not achieve a steady state in shear stress and layer thickness, with the exception of the shear stresses supported by samples B and C sheared at 650 °C. The experiment u913 involving sample A sheared at 650 °C was repeated, resulting in variable shear stress- and layer thickness-displacement curves. In general, the reproducibility of experiments is good as indicated by repeated experiments and the correlation between experiments with changed deformation conditions. However, there are some outliers in the obtained data, which have been thoroughly exam- ined in Zhan et al. (2024). For example, the observed high stress value in u1154 was attributed to loading-induced material extrusion, which re- sults in a small shear zone thickness and high shear strain rate. Fluctu- ations in the applied pore fluid press and temperature conditions can also contribute to the observed variation. Considering material extru- sion and experimental condition fluctuations, the standard deviation of the final shear stresses for the three tested samples remains within a reasonable range (with a maximum of ~16 % of the mean stress for sample A; see Table 2). Therefore, all experiments were retained for analysis in this study.

Despite some variability, all materials show a steep, near-linear in- crease in shear stress until reaching a yield point at ~1–2 mm displacement. This stage is followed by a continuous, gradual increase in

shear stress, occasionally accompanied by stick-slip events towards the end of the experiment. The only exception in the shear stress evolution is observed for sample A tested at 650 °C. In this experiment, the shear stress reaches its peak value shortly after the yield point, and then slowly decreases.

Regarding layer thickness evolution (Fig. 2d-f), all samples sheared at low T (20–450 °C) show a substantial decrease in layer thickness during the initial sliding stage. The rate of decrease of thickness slows down as displacement continues. Samples that are sheared at 650 °C typically show the least apparent compaction compared to those sheared at lower T. Only sample C exhibits minor dilation (thickening) starting from a displacement of ~10 mm. Overall, there is no systematic change in layer thickness with applied T across all tested samples.

To illustrate effects of temperature and mineral composition on the sample strength, we plot the yield-point shear stress ($\tau_{yd}$) and final shear stress ($\tau_{fin}$) as a function of T in Fig. 3. The average shear stress and its standard deviation are plotted for experiments that have been repeated. When sheared at relatively low T (20–450 °C), the granitoid sample A consistently shows higher yield and final shear stresses compared to the muscovite-rich sample C and the biotite-rich sample B. At 650 °C, sample A shows a marked decrease in $\tau_{fin}$ to ~61 MPa. In contrast, samples C and B exhibit increases in shear stresses with increasing T, and level off at 200 °C and 450 °C, respectively. The maximum shear stress for samples C and B are $\tau_{fin} \approx 69$ MPa and 56 MPa.

### 3.1.2. Mineralogical changes

To identify any possible syn-deformational changes in mineralogy, a limited amount (< 0.3 g) of gouge material was collected after each experiment and measured on top of a single crystal silicon (Si) plate to obtain a refined X-ray diffraction pattern. In the untreated patterns, the peak intensity (y-axis) and background (e.g. increased intensity towards low angles) are dependent on the amount of measured material and the instrument, and thus do not reflect any change in the modal abundance. Nevertheless, changes in the relative peak heights are indicative of changes in their relative abundance, and a visual broadening of peak shape demonstrates a decrease in crystal size and an increase in the number of crystal defects. Each XRD pattern was normalized to its highest peak intensity (Fig. 4a-c) in order to obtain the relative peak height. The background that is defined by a best-fit polynomial baseline, was subtracted for comparison across experiments (Fig. 4d).

All samples exhibit minor MoS₂ impurities (visible peak at 2θ = 13°) due to the antifriction lubricant (Molykote®), which was used to reduce the wall friction between sample and confining rings. As shown in Fig. 4a-c, the majority of the mineral peaks that were identifiable in the starting material remain discernible in materials that underwent deformation under various T. The exception to this is observed in sample C, where the peak of chlorite (2θ ≈ 6.5°, 12.5° and 25°) in the starting material becomes indistinguishable for those deformed at 200–650 °C (Fig. 4c). If the starting material contains certain amount of biotite (i.e. samples A and B in Fig. 4d), the biotite peak at 2θ ≈ 8.7° has substan- tially lower intensity once undergoing deformation at 200 °C, then it exhibits higher intensity as T increases to 650 °C. In the muscovite-rich sample C, the muscovite peak at 2θ = 8.9° shows a reduction in intensity at 200 °C, followed by an increase in peak intensity and peak broadening as T increases. At 650 °C, the observed peak broadening at 2θ = 8.9° appears as an increase of peak intensity at 2θ ≈ 8.7–8.8°, which is close to the referenced biotite peak position.

### 3.1.3. Microstructures

The microstructures of samples A, B and C sheared at a rate of V = 1 µm/s and various T (20–650 °C) are shown in Figs. 5, 6 and 7, respec- tively. The microstructures of granitoid sample A are characterized by planar fabrics, include Y-fractures, R-fractures and P-foliations (nomenclature after Logan et al., 1992; Passchier and Trouw, 1996). Fractures are traced along the former micro-shear bands, which opened
<figure><img src="image_6.png" /><figcaption>Fig. 3. (a) Plot of yield-point shear stress ($\tau_{yd}$) as a function of temperature, at a constant velocity of 1 µm/s. (b) Plot of final shear stress ($\tau_{fin}$) as a function of temperature. Each data point is derived from a single experiment. In the case of the experiment being repeated (i.e. sample A, 650 °C, 1 µm/s), the standard deviation is indicated in the graph using error bars.</figcaption></figure>

<figure><img src="image_7.png" /><figcaption>Fig. 4. (a-c) Normalized X-ray diffraction patterns of oriented gouge materials measured on a single crystal Si plate, showing changes in mineral composition before and after deformation at $V = 1$ µm/s and $T = 200$ °C, 450 °C and 650 °C. Note that MoS$_2$ is the impurity from antifriction lubricant (Molykote®) in the sample assembly. (d) Close up of XRD patterns after normalization and background correction. The background is defined by a best-fit polynomial baseline. The powder XRD patterns of pure Mg-rich biotite and muscovite from the online database (American Mineralogist Crystal Structure Database: https://rruff.geo.arizona.edu/AMS/amcsd.php) are included as references.</figcaption></figure>

during the unloading stage of the experiments due to their inherent weakness. They are either parallel to the shear plane as Y-fractures, or inclined as R-fractures. P-foliations are defined by alternating layers of aligned biotite and sigmoidal aggregates of fine-grained orthoclase in the P-orientation (Fig. 5a). Fabric elements characterizing the micro-structures of the mica-rich samples B and C include: (1) porphyroclasts,

defined by single quartz or quartz aggregates surrounded by a finer grained, polymineralic matrix (Fig. 7c); (2) fabric elements resembling S-foliations, C- and C'- shear bands, respectively (Chester and Logan, 1987; Passchier and Trouw, 1996). The C-, C'- fabrics are termed as shear bands instead of fractures because they do not consistently split and open along the bands upon unloading as in sample A. The geometrical
<figure><img src="image_8.png" /><figcaption>**Fig. 5.** (a)-(d) Mosaics of backscatter images of sample A, showing the main structure developing in sample sheared at a constant velocity of 1 µm/s and temperatures of 20–650 °C (modified from Zhan et al., 2024). The shear sense is sinistral for all samples. Sketches in the bottom left of each subfigure summarize the geometrical features of fractures and foliations. Close-up views include samples sheared at (e) 20 °C and (f, g, h) 650 °C (e and g modified from Zhan et al., 2024). (h) presents an SEM micrograph of a freshly fractured sample surface.</figcaption></figure>

orientations of S-C-C' fabrics are similar to the P-foliations, Y-fractures and R-fractures described in sample A, respectively.

In general, the granitoid sample A is characterized by a matrix-supported microstructure under all tested *T*, where competent clasts are embedded within a large volume of finer-grained matrix (*<* 1 µm). At 20 °C and 200 °C, the entire gouge layer exhibits pervasive fracture arrays and incipient P-foliations (Fig. 5a-b). In addition, grain size reduction in all mineral phases and densification are observed within a planar slip zone close to the sample-piston interface. This domain often constitutes ~40 % (~195 µm) of the total width of whole gouge layer (~500 µm), where fractured angular clasts (diameter of ~2 µm) are embedded in a large volume of fine-grained and porous matrix (diameter < 1 µm; Fig. 5e). At 450 °C and 650 °C (Fig. 5c-d), the total width of the principal slip zone (PSZ) comprises only ~5 % (~42 µm) of the entire gouge layer width (~893 µm). The differences in grain size and porosity between the less deformed domain and PSZs are pronounced. Only a few fracture arrays are observed to persist, whereas P-foliations orientated subparallel to the shear plane are pervasive. Notably, in the sample sheared at 650 °C, we observe microstructural features indicative of dissolution and precipitation processes. These include truncated grain-to-grain contacts, submicron pores trapped within crystals and along grain boundaries, mineral overgrowth, and small, idiomorphic precipitates (Fig. 5f-h). Relatively large quartz, albite and epidote grains are embedded in a dense and ultrafine-grained matrix with extreme grain size reduction. Overall, the sheared sample A transitions from a porous and fractured microstructure at low *T*, to a compacted and highly

strain-localized microstructure at high *T*.

The evolution of microstructures with increasing *T* remains closely similar between samples B (Fig. 6) and C (Fig. 7), with the degree of mica grain interconnectivity and foliation intensity increasing in both cases. At 20 °C, gouge layers show widespread fractures and no discernible development of PSZs. Porphyroclasts are composed of single large quartz grains (~60 µm in diameter), interlayered between incipient S-C fabrics (Figs. 6a & 7a). Mica grains (~2 µm) are aligned and oriented at 30° or subparallel to the shear plane, resulting in incipient S-foliations and C-shears observed at a larger scale. Closer inspection reveals that quartz exhibited angular shapes while elongated, platy mica grains are characterized by frayed and jagged grain boundaries (Figs. 6e & 7e). Small pores preferentially develop at the edge of mica-mica contacts. At 200 °C and 450 °C, the originally large quartz grains (~30–80 µm in Fig. 7a) are fragmented to smaller sizes, forming sigmoidal-shaped porphyroclasts comprising of a small number of grains (~30 µm) surrounded by finer ones (*<*10 µm) (Figs. 6b-c & 7b-c). Compared to low *T*, S-C fabrics observed in the 200–450 °C experiments are better developed, particularly the S-foliations.

At 650 °C, strain is localized into two PSZs, each situated adjacent to the upper and lower sample-piston boundaries (Figs. 6d & 7d). Together, these PSZs span ~230 µm in total width, constituting roughly 40 % of the total gouge layer width. The localization of strain is indicated by (1) intensified reduction in grain size and porosity (Fig. 6f), (2) S-foliations dipping at a less steep angle to the shear plane, and (3) the development of C'-shears, which are considered as large shear strain structures
<figure><img src="image_9.png" /><figcaption>**Fig. 6.** (a)-(d) Mosaics of backscatter images of sample B sheared at a velocity of 1 µm/s and temperatures of 20–650 °C. Sketches summarize the geometrical features of S-C-C' fabric and porphyroclast system. (e) Close-up views of the sample sheared at 20 °C, showing pervasive microcracking and moderate alignment of mica grains along S-foliation orientation. (f)-(g) Close-up views of the strain localized domain of the sample sheared at 650 °C, showing remnant pores within quartz porphyroclasts, a local truncated quartz-quartz contact and a slightly elongated quartz grain (outlined in white). Mica grains show bending, kinking and grain edge-edge contacts.</figcaption></figure>

(Fig. 6d; Passchier and Trouw, 1996). Within these PSZs, quartz porphyroclasts have sigmoidal shapes and are sandwiched between well-aligned micas. Aligned mica grains form S-foliations dipping at ~30° to the shear plane, similar to those sheared at 200–450 °C. The porphyroclasts consist of angular to sub-angular quartz aggregates, varying in grain size from a few nanometers to ~3 µm, with nanopores often observed between the quartz-quartz contacts (Fig. 6f). Elongated quartz grains and truncated quartz-quartz contacts are observed sparsely distributed within the PSZs (Figs. 6g & 7g). Internal C-shears contain fine-grained, dense, and polymineralic mixtures (Fig. 7g), resembling the S-foliations. Submicron quartz grains are sparsely distributed among a mica-dominant matrix, resulting in a more homogeneous mixture of phases compared to those sheared at lower T. Flaky mica grains (~1 µm) are tightly compacted, and are bent and kinked when located nearby

quartz aggregates (Figs. 6g & 7f).

## 3.2. Slow-shear experiments

To investigate the effect of strain rate on mineral reactions inferred from XRD data (Fig. 4), we conducted additional slow-shear experiments with V = 0.03–0.1 µm/s and T = 650 °C (Table 2; one test for sample B and three tests for sample C). The experiments on sample A sheared at similar rates from Zhan et al. (2024) are included for comparison. The experiments of samples A and C at 650 °C were repeated, showing variation in the shear stress-displacement curve. As mentioned in Section 3.1, the observed variation of sample A has been thoroughly examined in Zhan et al. (2024). For sample C, experiment u1227 consistently exhibits shear stresses that are ~15 MPa higher than that
<figure><img src="image_10.png" /><figcaption>**Fig. 7.** (a)-(d) Mosaics of backscatter images of sample C sheared at a velocity of 1 µm/s and temperatures of 20–650 °C. Close-up views include (e) sample sheared at 20 °C, (f) S-foliation as well as (g) C-shear of the strain-localized domain of sample sheared at 650 °C. White arrows point to the key microstructural features described in the main text.</figcaption></figure>

for experiment u916. This is due to slightly different experimental conditions (e.g. pore fluid pressure), sample-to-sample compositional variations, and deformation apparatus maintenance between experiments (Text A.2 and Fig. A.2 in the Appendix A).

The strength of sample B sheared at 0.1 µm/s and 650 °C, remains comparable at $\tau = \sim 56$ MPa to that observed at 1 µm/s. In contrast, the

strengths of both samples A and C decrease with decreasing V (Figs. 8 vs. 3). Shear stresses of samples A and C exhibit a steep increase in the first 0.5–1 mm displacement, followed by continuous weakening until a steady state is reached at ~5 mm displacement (experiments u1227 and u910; Table 2), or until the termination of experiments (experiments u1235, u916, u1175). At $V = 0.1$ µm/s, sample C displays a higher final
<figure><img src="image_11.png" /><figcaption>**Fig. 8.** (a) Plot of shear stress as a function of displacement for samples A, B and C sheared at ultraslow rates of *V* = 0.03–0.1 µm/s and *T* = 650 °C. Micrographs of sample A showing (b) local reactions forming calcium-rich domains and (c) PSZ with newly precipitated iron-magnesium-rich silicates (e.g. amphibole). (e)-(f) Micrographs of sample C showing reactions within the PSZ, forming biotite (bt). Newly formed phases are labeled in red. Chemical composition analyses of newly formed minerals are shown in Fig. A.3 in the Appendix A. The grain size distribution of quartz/albite grains within the PSZs for samples A and C are shown in (d) and (g), respectively. The apparent diameters are area-weighted diameters of granular particles. The dashed line represents the median diameter, while the dotted line represents the mode diameter. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

shear stress compared to sample A.

At $V = 0.1$ µm/s and 650 °C, the deformation of both samples A and C is preferentially localized within very thin, dense and fine-grained PSZs. The widths of the PSZs constitute 3–6 % (~20–30 µm) of the whole gouge layer in sample A (~141–208 µm), and ~3 % (~20 µm) of the whole layer in sample C (~667 µm). The less deformed domains outside the PSZs are characterized by typical brittle/cataclastic features, while newly formed phases are observed within the PSZs in both samples. EDS analyses (Fig. A.3 in Appendix A) show that the local albite-orthoclase

contacts observed in sample A become enriched in calcium (Fig. 8b), along with submicron iron-magnesium-rich silicates (e.g. amphibole or biotite) distributed at grain-grain boundaries (Fig. 8c). Quantitative microstructural analyses show that the median and mode area-weighted diameters of quartz/albite in sample A are ~480 nm (Fig. 8d). In sample C, the breakdown of muscovite/illite and the formation of biotite are observed within PSZs (Figs. 8e-f and A.3 in Appendix A). Quartz/albite (~200 nm–2 µm) and muscovite/illite (~200–400 nm) exhibit tabular grain boundaries. Newly formed biotite grains have an anisotropic
needle shape and are small, with ~250 nm in length. No evidence of bending of biotite needles is observed. Quartz/albite grains in sample C have a median diameter of 310 nm and a mode diameter of 220 nm (Fig. 8g).

# 4. Discussion

## 4.1. Comparison with previous studies

To demonstrate the consistency of our mechanical data with previous studies, we calculate the apparent friction coefficients ($\mu = \tau/\sigma_n^{eff}$), assuming that cohesion is negligible (following Niemeijer et al., 2016). The symbol $\tau$ represents shear stress in MPa, derived at the end of experiments (shear strain of 30). $\sigma_n^{eff}$ is the effective normal stress in MPa, equivalent to normal stress minus pore fluid pressure. The comparison of original shear stress and their detailed experimental conditions are shown in Fig. A.4 and Table A.2 in the Appendix A, respectively.

At room temperature, the friction coefficients decrease in the following order (Fig. 9a): granitoid sample A ($\mu = 0.81$), muscovite-rich sample C ($\mu = 0.47$) and biotite-rich sample B ($\mu = 0.44$). Pure mica gouges exhibit the lowest friction coefficients with $\mu = 0.27–0.38$ (den Hartog et al., 2013; Lu and He, 2014; Mariani et al., 2006) compared to the mica-bearing polymineralic mixtures (samples A, B and C). The three tested gouges of this study have comparable amounts of quartz, with sample A having a high feldspar content and a low biotite content, while samples B and C contain low feldspar content and high sheet silicates content (Section 3.1). This indicates that strength decreases as the proportion of sheet silicates increases, which is consistent with findings from previous studies (Ashman and Faulkner, 2023; Tokle et al., 2023; Lupini et al., 1981; Moore and Lockner, 2011; Tembe et al., 2010). As the proportion of sheet silicates increases, the microstructure transitions from a qtz/fsp grain-supported (Fig. 5e) to a sheet silicate foliation-supported fabric (Fig. 7e). This observation provides further support to the hypothesis that the strength reduction is mainly controlled by the microstructural evolution of the gouge layer, such as the development of interconnected weak layers, e.g. foliations (Crawford et al., 2008; Handy, 1990; Lupini et al., 1981; Niemeijer, 2018; Tembe et al., 2010).

At elevated temperature and still low velocity (1 µm/s), wet mica-rich gouge materials (samples B and C) exhibit a slight but consistent increase in friction coefficients (25–47 %) with increasing $T$, reaching a plateau at $T = 200–450$ °C (Fig. 9a). This T-dependent strengthening

behavior has been reported by many previous studies focusing on pure sheet silicate end-member gouges, such as muscovite (den Hartog et al., 2013; Mariani et al., 2006; Van Diggelen et al., 2010), biotite (Lu and He, 2014), chlorite (Belzer and French, 2022; Okamoto et al., 2019), serpentine (Moore et al., 1997; Takahashi et al., 2011) and illite (den Hartog et al., 2012a, 2012b). The microstructural evolution of mica-rich gouges in our study exhibits several common features with these previous studies: (1) deformation becomes more localized with increasing $T$, manifested as PSZs of several microns width or as foliations with inclined orientation; (2) PSZs typically exhibit extreme grain size reduction and compaction; and (3) the overall porosity of the gouge layer decreases substantially, with larger pores that are fewer in number, and replaced by submicron pores developed at the hinges of kink bands in micas. These observations consistently suggest that T-dependent strengthening is primarily attributed to the microstructural changes within the gouge mixture, such as enhanced strain localization and elimination of larger pores. A detailed discussion on the possible mechanisms behind this T-sensitive behavior is provided in Section 4.2.

At 650 °C and $V = 0.03–1$ µm/s, the strength of muscovite-rich sample C decreases at a less pronounced rate with decreasing $V$ than that of the granitoid sample A (Fig. 9b). This rate-sensitive behavior of mica-rich gouges is consistent with pure muscovite, or muscovite-quartz gouges sheared at $T$ between 500 and 700 °C, $V \approx 1 \times 10^{-4}–10$ µm/s, $P_f = 80–103$ MPa and $\sigma_n^{eff} = 100–120$ MPa (Mariani et al., 2006; Niemeijer, 2018; Van Diggelen et al., 2010). At 650 °C and 0.1 µm/s, sample C exhibits muscovite breakdown and the formation of biotite (Fig. 8d-e). Similar reactions and simultaneous weakening have also been reported in nearly pure muscovite gouges sheared at 700 °C, $V < 1$ µm/s, and similar normal stresses and pore fluid pressures (Table A.2 in Appendix A) (Mariani et al., 2006; Van Diggelen et al., 2010). The compositions of the newly formed minerals (biotite in sample C) differ from those being present in the starting material (Fig. A.3). This fact, together with small grain sizes and idiomorphic grain shape (Fig. 8), clearly indicates neocrystallization and mineral reaction. This suggests that chemical reactions, particularly the breakdown of muscovite, may play an important role in the weakening seen at elevated $T$ towards low strain rates. Compared to the muscovite-rich gouge, the granitoid gouges (sample A) and the 20 % muscovite +80 % quartz mixture (Niemeijer, 2018), both exhibit a more abrupt decrease in strength as $V$ decreases. The microstructure in the former case points to pervasive DPC-related features mostly in quartz and feldspar (Fig. 5f-h, also in Zhan et al.,

<figure><img src="image_12.png" /><figcaption>**Fig. 9.** Comparison of friction coefficients obtained in this study with those from previous studies, illustrating (a) the temperature-dependent strengthening in mica-rich gouges at $V = 1$ µm/s, and (b) velocity-dependent weakening in samples A and C at $T = 650$ °C. The original shear stress data and detailed experimental conditions are presented in Fig. A.4 and Table A.2 in the Appendix A, respectively.</figcaption></figure>
2024), while the weakening in the latter case was interpreted by Niemeijer (2018) as due to frictional sliding over a weak foliation and DPC of intervening quartz clasts. Taken together, this implies that the quartz and/or feldspar grains may exhibit more pronounced weakening when compared to mica grains, probably due to faster kinetics of DPC in quartz and/or feldspar grains.

## 4.2. Temperature-, rate-sensitive behavior in mica-rich gouges

The mica-rich samples B and C consist mainly of two mineral phases, which are quartz and mica. Albite is also present in samples B and C, but in amounts less than 5 wt%. For the purpose of discussion, we will focus on the deformation processes affecting the volumetrically more important phase, not albite. Some tests on samples B and C were held at 650 °C for maximum 13.5 h (see Table 2). Since previous work showed that samples being held for zero and ~10 h exhibited similar microstructural features (Fig. 5e vs. 5f in Zhan et al., 2024), the microstructural features presented in Figs. 6–7 are representative of the deformation conditions, irrespective of the hold time.

At 20 °C and 1 µm/s, the angular, frayed and jagged grain boundaries (Figs. 6e & 7e) suggest that quartz and mica are both deformed by fragmentation and frictional granular flow, i.e. by cataclastic flow. At 650 °C, micas were observed to bend and kink, indicating a component of crystal plastic deformation. Grain size reduction of mica is likely caused by sequential operation of frictional processes followed by dissolution-precipitation creep (DPC). While the frictional processes can efficiently reduce the grain sizes, they typically produce a porous texture with angular-shaped grains embedded within a finer-grained matrix. This is not the representative microstructure observed here. Instead, DPC in a polymineralic system may further reduce the grain size through dissolving parent grains and precipitating new grains, resulting in a texture with smoothed grain boundaries and low porosity.

For rocks deformed solely by a viscous creep mechanism at high temperature and low stress, the constitutive equation describing flow law (Gleason and Tullis, 1995; Twiss and Moores, 1992), is generally expressed as $\dot{e} = A\sigma^n d^{-m}\exp(-Q/RT_K)$. Here, A is the pre-exponential constant, *n* the stress exponent, *d* the grain size, *m* the grain size exponent, *Q* the activation energy, *R* the gas constant, and *T*<sub>K</sub> the temperature in Kelvin. The power *n* represents the stress sensitivity of strain rate. Previous work (Zhan et al., 2024) shows that in a granitoid system (referred to as sample A in this study), a decrease of *n* from >10 at 20–450 °C to ~2.2 at 650 °C coincides with an increased contribution of DPC over cataclastic flow at high temperatures. A similar trend was observed by McDonnell et al. (1999) in a synthetic dunite system, where *n* value decreased from ~2.2 to 1.3 as temperature increases from 850 °C to 1000 °C. This was attributed to an increasing contribution of diffusion creep (typically *n*=1–3) over dislocation creep (*n*=3–5). Taken together, these observations suggest that a reduction in the measured stress exponent may indicate a progressive shift in the relative contributions of multiple deformation mechanisms. Here, by fitting the $\dot{e}$ to $\sigma$ obtained from tests run at 650 °C, the apparent stress exponent *n* for sample A is 2.2 ± 1.8, and for sample C is 6.8 ± 2.2 (Fig. 10). The calculation of *n* value assumes similar grain sizes across individual test, but the potential influence of grain size variation on the final *n* value remains uncertain. The *n*-value of 6.8 for sample C is lower than *n* > 10 (as estimated for cataclastic flow) and higher than 2.2 for viscous creep dominated deformation in sample A. This might indicate a larger contribution of cataclastic flow over DPC to the overall deformation in sample C compared to sample A. Alternatively, it may reflect that DPC is less efficient in sample C than in sample A. It is important to note, however, that the determination of *n* is sensitive to both the number of data points and the range of strain rates used in the fitting. Therefore, robust and consistent data fitting is essential for reliable interpretation of the relative contributions of different deformation processes.

It seems counter-intuitive that mica-rich gouges exhibit an increasing strength with *T* (Fig. 3), especially considering the greater contribution

<figure><img src="image_13.png" /><figcaption>**Fig. 10.** A log-log plot of strain rate as a function of differential stress for samples A and C, sheared at *T* = 20, 200, 450 and 650 °C. Strain rate ($\dot{e}$) = shear strain rate ($\dot{\gamma}$) / $\sqrt{3}$; differential stress ($\sigma$) = $\sqrt{3} \times$ final shear stress ($\tau_{fin}$) (Verberne et al., 2017). The shear strain rate $\dot{\gamma}$ is equal to the velocity (*V*) divided by the width of strain-localized domain (e.g. PSZ). The slope of the linear regression fitting line represents the stress exponent (*n*).</figcaption></figure>

of viscous creep inferred from the microstructural evolution as *T* increases (Figs. 6 & 7). In the case of the granitoid sample A, its strength decreases at *T* above 450 °C due to an increased DPC component (Fig. 2, see also Zhan et al., 2024). Therefore, the *T*-dependent strength increase in the mica-rich gouges must be related to the micas. Several microstructural aspects in mica can be listed considering this strength increase. First, the locally observed bending of mica grains at higher *T* makes frictional sliding along the weak {001} grain interfaces more difficult. Second, the formation of smaller-sized mica grains (~2 µm in Fig. 7g vs. <1 µm in Fig. 7e) may increase the number of edge interactions or collisions when mica slide over each other (den Hartog et al., 2020; Okamoto et al., 2019), resulting in a higher frictional resistance and thus an increase in shear stress. Third, the phase mixing of submicron quartz with the mica (Figs. 6f & 7g) may play a role through interrupting the development of interconnected weak mica layers (Handy, 1990; den Hartog et al., 2012a, 2012b), or by increasing the proportion of stronger quartz-mica contacts over weak mica-mica contacts. The frictional resistance between quartz-mica interfaces with $\mu \approx 0.3–0.6$ (den Hartog et al., 2013; Niemeijer, 2018) is generally considered to be higher than those between mica-mica interfaces with $\mu \approx 0.08–0.27$ (Homola et al., 1990; Kawai et al., 2015; Niemeijer, 2018). Phase mixing may also increase the amplitude of anastomosing foliation “waves” and hence dilatation observed at 650 °C (Fig. 2f). In gouge materials, dilatation can cause a higher shearing resistance (Ashman and Faulkner, 2023; den Hartog and Spiers, 2014).

Water films adsorbed on the grain surface can act as lubricants for frictional sliding (Behnsen and Faulkner, 2012; Israelachvili, 1997; Renard and Ortoleva, 1997). Previous studies have proposed that higher *T* may cause water film thinning, leading to more rigid grain structure and consequently higher frictional resistance (Belzer and French, 2022; Kubo and Katayama, 2015; Okamoto et al., 2019). However, the surface water film thickness depends on factors such as pore fluid salinity, pH and the normal stress concentration at the grain contact (Alcantar et al., 2003; Renard and Ortoleva, 1997). It is reasonable to speculate that the water saturated, high pore fluid pressure applied in our study inhibits the thinning of the water films at high *T*.

In addition to the *T*-sensitive behavior, rate-sensitive behavior is evident in sample C at *T* ≥ 650 °C and/or *V* ≤ 0.1 µm/s. Rate-sensitive behavior is accompanied by viscous processes and pervasive reactions. As *V* decreases, reactions that were mainly inferred by XRD pattern
changes (Fig. 4) become evident in the microstructures at 650 °C and V = 0.1 µm/s (Fig. 8f). Enhanced reaction is accompanied by a decrease in strength, from µ = 0.69 at V = 1 µm/s to µ = 0.61 at V = 0.1 µm/s (Fig. 9). Muscovite breakdown may soften the gouge by reducing the grain size of the original reactive grains, and producing fine-grained reaction products, e.g. biotite (Fig. 9f; Brodie and Rutter, 1985). Smaller grains can substantially enhance any grain-size-sensitive viscous processes (e.g. DPC) due to their faster dissolution via the larger reactive surface area, and faster diffusive mass transportation over shorter distances at the grain scale. Although muscovite is frictionally stronger (µ = 0.52–0.66 at 650 °C; den Hartog et al., 2013; Van Diggelen et al., 2010) than biotite (µ = 0.44 at 650 °C; Lu and He, 2014), newly formed biotite does not necessarily directly contribute to softening. This is because the biotite grains are observed to be randomly oriented (Fig. 8f) and are not yet aligned to form a weak foliation.

### 4.3. Strength reversal at 650 °C: granitoid gouge is weaker than mica-rich gouges

At 20–450 °C and 1 µm/s, the granitoid sample A consistently exhibits higher strength compared to mica-rich samples B and C. This trend is reversed at 650 °C and V ≤ 0.1 µm/s, where granitoid gouge becomes much weaker (Fig. 9). The following discussion mainly addresses the unexpected strength reversal between the samples A and C.

At 650 °C and 0.1 µm/s, the quartz grains within the PSZs in granitoid and muscovite-rich gouge are at least partly deforming by DPC, as indicated by their smooth grain boundaries and mineral reactions (Fig. 8). Our initial hypothesis was that the observed strength reversal between the granitoid and muscovite-rich gouges could be attributed to a faster DPC in the granitoid system. Given that DPC is highly sensitive to grain size variations, we quantitatively analyzed the grain sizes of quartz and/or albite grains (their colour contrast in backscatter micrographs is very subtle). Results show that the qtz/ab grains in the muscovite-rich gouge (a total of ~42 wt%) are ~1.5 times smaller in median size when compared to those in the granitoid gouge (a total of ~75 wt% qtz/ab) (Figs. 8d & 8g). This suggests that the DPC rate of qtz/ab in the muscovite-rich system could be 1.5 to ~3.4 times faster than in the granitoid system, depending on rate-controlling processes (rate $\dot{\epsilon}$ ∝ size $d^{-m}$, with m = −1 for dissolution and precipitation-controlled DPC, or −3 for diffusion-controlled DPC) (Niemeijer et al., 2002; Spiers et al., 2003). Considering the volume fraction of qtz/ab in the samples, we infer that the strain rate contribution by DPC should still be faster in the muscovite-rich gouge than the granitoid gouge by a factor in the range of 0.84 to 1.9. This implies that the muscovite-rich gouge should be similar in strength to (or weaker than) than the granitoid gouge at 650 °C and 0.1 µm/s, which it is not consistent with our measurement. This observation suggests that the remaining minerals within the granitoid gouge, namely 11 wt% orthoclase, 8 wt% biotite and 6 wt% epidote, may play a more important role in determining the apparent lower strength of the whole gouge layer at 650 °C. In addition, DPC rate is also highly sensitive to the thickness of PSZ, which in turn affects the strain rate. Independently, the pore fluid chemistry also influences the DPC rate. The development of PSZ leads to a thinner shear zone, resulting in a faster strain rate that favors frictional processes over DPC. As samples A and C have similar PSZ widths (~3–6% of the entire gouge layer, section 3.2), their shear strain rates are most probably similar. This observation rules out the effect of PSZ widths on DPC rates, thus redirecting the focus towards the pore fluid chemistry. Previous work of Niemeijer and Spiers (2002) reported a decrease of DPC rate (reflected by an increase of compaction rate) as muscovite being added to the quartz sandstone. They proposed that the dissolved Al³⁺ from muscovite may decrease the quartz solubility, therefore slowing down the rates of dissolution and precipitation of quartz. Their hypothesis assumes that the rate-controlling processes are dissolution and/or precipitation rather than diffusion. It may offer a possible explanation for the slower DPC rates of sample C observed in our study, but this hypothesis needs to be tested

through systematic experiments that are designed to confirm the rate-controlling processes.

### 4.4. Implications

To compare the differential stress required to reactivate a pre-existing mica-rich fault with a granitoid fault, we construct strength profiles for experimental fault gouges using data points from the experiments reported in this study (Table A.3 in Appendix A). We distinguish two types of extrapolations: (i) one for natural fault zones deformed at strain rates similar to our experiments, meaning relatively fast but still aseismic ($\dot{\epsilon}$ = $10^{-4}$–$10^{-3}$ s⁻¹; Fig. 11, upper part). (ii) The second extrapolation type deals with low natural strain rates ($\dot{\epsilon}$ = $10^{-14}$–$10^{-12}$ s⁻¹; see Fig. 11, lower part).

#### 4.4.1. Strength profile for experimental fault gouges at fast rates

The methodology for constructing strength profile is detailed in Zhan et al. (2024). The temperature used in the experiments is converted to depth using a geothermal gradient (25 °C/km), and the effective normal stress ($\sigma_n^{eff}$) is estimated by subtracting the hydrostatic pore fluid pressure ($P_f$ = 0.4 × lithostatic pressure) from the normal stress ($\sigma_n$, 27.5 MPa/km) expected at each depth. Differential stress ($\sigma$) is calculated by multiplying $\sigma_n^{eff}$ with the apparent friction coefficients ($\mu_{fin}$) of those experiments sheared at $\dot{\epsilon}$ of $10^{-4}$–$10^{-3}$ s⁻¹. The calculated strength profiles reflect the deformation processes active under the experimental conditions of this study, with $\dot{\epsilon}$ falling within the range $10^{-4}$–$10^{-3}$ s⁻¹ (or V = 0.1–1 µm/s), T between 20 and 650 °C and confining pressures $\sigma_3$ between ~50–80 MPa. Given that the ring shear apparatus does not actively apply confining pressure to the sheared gouge material, we calculated the passive $\sigma_3$ assuming a two-theta angle of 120° in the Mohr diagram (Text A.1 in Appendix A). The strength profile of mica schist (~75% biotite) is plotted for comparison, using a power law flow law with parameters of n = 31, Q = 98 kJ/mol, and A = 1.3e-67 MPa⁻¹ s⁻¹ (Shea and Kronenberg, 1992). This flow law was determined using mechanical data for mica schist sheared at 45° to its foliation, and the deformation behavior was interpreted to be semi-brittle, accommodated by a combination of brittle and crystal plasticity grain-scale processes. The flow stress calculated from this flow law is included as an end-member case to demonstrate the lower bound of strength for natural rocks with high mica content.

The strengths of synthetic biotite-rich and muscovite-rich gouges increase linearly at V = 1 µm/s with T from 20 °C up to 650 °C, characterized by µ = 0.37 and µ = 0.46, respectively (Fig. 11b). The linear increase in strength is a combined effect of the linear lithostatic/hydrostatic gradient, the linear geotherm, plus the nearly linear increase in µ with T. At T = 650 °C and V = 0.1 µm/s, the strength of muscovite-rich gouge decreases most likely due to reaction-driven grain size reduction hence accelerating grain-size sensitive DPC in mica grains (see Section 3.2.3). The mica-rich gouges are weaker than granitoid gouges (Fig. 11a) within T range of 20 °C to ~450 °C. Above this T, granitoid gouges are significantly weaker than mica-rich gouges, due to the onset of viscous creep, which is well constrained at experimental conditions of this study ($\dot{\epsilon}$ = $10^{-4}$–$10^{-3}$ s⁻¹; see also Zhan et al., 2024). However, at T as high as 650 °C in natural geodynamic settings (e.g. subduction and collision zones), the confining pressure will be much higher and affect mineral stabilities. For example, muscovite/illite breakdown (melt) may not occur at higher T in the presence of high confining pressure and excess water in the P-T diagram of muscovite-granite system (Huang and Wyllie, 1981). Our experimental results can therefore only be applied to natural conditions at T < 650 °C, low $\sigma_3$, and at $\dot{\epsilon}$ similar to the experimental ones (e.g. slow slip domains).

A major discrepancy between nature and experiments is that most aseismic deformation in nature occurs at strain rates of $\dot{\epsilon}$ = $10^{-14}$–$10^{-12}$ s⁻¹, many orders of magnitude slower than those in our experiments with $\dot{\epsilon}$ = $10^{-4}$–$10^{-3}$ s⁻¹. Our experiments highlight that the weakening
Strength profile for experimental fault gouges at fast rates (10⁻⁴~10⁻³ s⁻¹)

<figure><img src="image_14.png" /></figure>

Strength profile for fault rocks at slow natural rates (10⁻¹⁴~10⁻¹² s⁻¹)

<figure><img src="image_15.png" /><figcaption>**Fig. 11.** (a)-(b) Strength profiles constructed for synthetic fault gouges sheared at fast strain rates ($\dot{e} = 10^{-4}–10^{-3} s^{-1}$) on granitoid sample A, biotite-rich sample B, and muscovite-rich sample C. Each dot represents differential stress calculated using the measured shear stress (Table A.3 in the Appendix A). The curved envelop is manually drawn line to the extended data points. Strength profile for granitoid gouge sheared at 0.1 $\mu$m/s is redrawn from Zhan et al. (2024). The flow stresses calculated from mica schist flow law are plotted for comparison (Shea and Kronenberg, 1992) (see flow law parameters in the main text). (c)-(d) Conceptually extrapolated strength profiles for mica-poor and mica-rich granitoid rocks deformed at natural slow rates ($\dot{e} = 10^{-14}–10^{-12} s^{-1}$). Micrographs shown in Fig. 12 illustrate that natural granitoids may partly be altered and rich in mica content. DPC can induce weakening and strain localization in both mica-poor and mica-rich granitoid systems.</figcaption></figure>
in the granitoid and muscovite-rich gouges is triggered by enhanced grain-size-sensitive, temperature- and rate-dependent viscous creep, namely DPC. This has important implications for extending the strength profile for experimental gouges to natural conditions, because DPC competes better with frictional processes at lower strain rates, lower temperature, and finer grain sizes. A consequence of this is an onset of the frictional-viscous transition in granitoid and mica-rich gouges at lower crustal levels and temperatures as outlined in the next section.

### 4.4.2. Implication for granitoid fault zones deformed at slow rates

The strength profile for experimental fault gouges (Fig. 11) highlights that increased mica content (or decreased feldspar content) leads to weakening at low *T* and fast *ϵ*. Given the widespread presence of micas (e.g. muscovite) across a broad range of geological conditions, relying solely on quartz or granitoid rheology to represent the crustal behavior may be oversimplified. In nature, micas may form through precipitation from fluids infiltration, or through feldspar replacement reaction in hydrothermally active fault zones (e.g. Wintsch et al., 1995 and references therein). To highlight the relevance of our results for

natural granitoid fault systems deforming at much slower rates, we compare the microstructural features between experimental gouges with natural granitoid fault rocks from the Aar Massif, which underwent distinct degrees of feldspar alteration.

At the Nagra Grimsel Test Site, two types of end-member granitoid fault rocks are discriminated based on their mica content: the mica-poor and the mica-rich end-members (Fig. 12). These compositions correspond well to those of granitoid sample A and muscovite-rich sample C in this study, respectively. In the mica-poor granitoid host rock, the degree of feldspar alteration is limited (Fig. 12, left column). The magmatic fabric and feldspar grains are generally preserved, with minor evidence of healed fractures (Fig. 12a). Quartz exhibits grain size reduction through dynamic recrystallization by subgrain rotation, owing to Alpine deformation under viscous conditions (Schneeberger et al., 2017; Wehrens et al., 2016). With progressive exhumation and cooling, rocks are characterized by embrittlement, grain size reduction, strain localization, and the formation of fault gouges/cataclasites characterized by two grain-size classes: angular clasts with *d* = 0.5–2 mm (Fig. 12b), and matrix with resolvable *d* up to 200 µm (Fig. 12c). The

<figure><img src="image_16.png" /><figcaption>**Fig. 12.** Micrographs showing the transition of weakly deformed natural granitoid host rocks into (viscous-) frictional gouges for mica-poor (left column) and mica-rich (right column) systems. (a) Mica-poor granitoid host rock (Central Aar granite) representing a magmatic fabric with slight deformation overprint (see text). (b) Mica-poor cataclastic gouge dissecting the mica-poor Central Aar granite host (plane light), with detail shown in (c). (d) Same as (a) but dissected by a fracture with white mica infill. (e) Completely altered K-feldspar grains (upper part), and a mixture of muscovite/illite and quartz/albite as alteration products. Detail of (e) is shown in (f). With the exception of image (b), all the other images (a, c-f) were taken under cross-polarized and plane light conditions.</figcaption></figure>
presence of submicron grains has been documented in other natural granitoid gouges (e.g. Keulen et al., 2007). These natural microfabrics correlate very well with the experimentally deformed granitoid gouges (sample A; Fig. 5). In comparison to our experimental results, we attribute the weakness of the natural granitoid fault rocks to the operation of the grain-size sensitive DPC occurring within the fine-grained domains (see Zhan et al., 2024 and discussion therein).

In the case of mica-rich rocks, two members are distinguished. In the first member, originally mica-poor host rocks are fractured and injected by fluid, which are subsequently filled by muscovite/illite precipitates (Fig. 12d). These precipitations are typical of low-temperature deformation stages in granitoid rocks on the retrograde cooling path (e.g. Schneeberger et al., 2017; Wehrens et al., 2016). The second member is characterized by extensively altered feldspar, replaced by muscovite and albite (Fig. 12e). Deformation tends to localize within these mica-rich domains, manifesting as either interconnected muscovite fault planes (Fig. 12d) or shear bands comprised of mus-ab-qtz mixtures (Fig. 12f). Dissolution-precipitation processes are evident from the microstructures. Such muscovite-rich fault rocks are commonly associated with low-temperature deformation in the altered granitoids, with their respective tectonites referred to as phyllonites (e.g. Hippertt, 1994; Wallis et al., 2015). Based on the mineral assemblages, the formation temperatures of phyllonites are often estimated to be between 350 and 400 °C (e.g. Maggi et al., 2012; Montemagni et al., 2023; Wallis et al., 2015). In addition, Berger et al. (2017) demonstrated that DPC in mica depends on grain sizes and the chemical equilibrium processes. Mineral compositions and structures adjust to new conditions of temperature and pressure. With additional chemical driving forces, DPC can become active in mica-rich tectonites at temperature as low as 200–300 °C.

The efficient operation of DPC may reduce the crustal strength and trigger the frictional-viscous transition in the mica-rich tectonites, similar to those in the mica-poor granitoids. Given the comparable low temperature limits required for the onset of DPC in mica-poor granitoids (180–320 °C; Zhan et al., 2024) and mica-rich phyllonites, both systems may exhibit a frictional-viscous transition at similar crustal depths (lower bound of greenschist metamorphic facies). The strength reversal observed in this experimental study (Section 4.3) is less likely to occur in natural, steadily creeping crustal faults, unless the local conditions (e.g. strain rate or confining pressure) are altered in favor of muscovite breakdown reaction due to deformation. A major problem remains to be answered: what is the efficiency of DPC at low rate and low temperature? To answer this question, additional information is needed to evaluate the kinetics of dissolution, diffusion, and precipitation processes.

## 5. Conclusions

We performed sliding experiments on biotite-rich, muscovite-rich and granitoid gouges under hydrothermal conditions and low displacement rates to simulate the geological environments in the upper-to-middle continental crust. Based on the mechanical data and microstructural observations, the following conclusions can be drawn:

1. Biotite- and muscovite-rich gouges are frictionally weaker ($\mu$ = 0.44–0.47) than granitoid gouges ($\mu$ = 0.81) at room temperature and velocity of 1 µm/s.

2. The two mica-rich gouges sheared at 1 µm/s exhibit strengthening with increasing temperature and reach a strength plateau at $T \ge 200–450$ °C.

3. Weakening in the muscovite-rich gouge occurs as velocity falls below $V \le 0.1$ µm/s at 650 °C, while in granitoid gouge, it is observed at $V \le 1$ µm/s and $T \ge 450$ °C.

4. Weakening is accompanied by strain localization into µm-wide, submicron-sized, and dense principal slip zones. These zones exhibit features such as truncated grain-grain contacts, nanopores, neorecrystallization, and mineral reactions (i.e. the formation of Ca-rich

feldspar and amphibole in granitoids, and muscovite breakdown plus biotite formation in muscovite-rich gouge). These features indicate a key role of dissolution-precipitation creep (DPC) as the underlying deformation process in addition to cataclastic flow.

5. The DPC rate in granitoids is faster than the muscovite gouges at temperatures >450 °C and strain rates of $10^{-4}–10^{-3}$ s$^{-1}$, as the granitoids become weaker. Given the similar amount and grain size of quartz in both gouges, the faster DPC in granitoid gouge may be due to the efficient DPC of less volumetrically dominant phases such as orthoclase, biotite and epidote. Additionally, previous study shows that the presence of muscovite may inhibit DPC by altering pore fluid chemistry through $Al^{3+}$ dissolution, thereby slowing down dissolution and precipitation rates.

6. DPC is promoted at higher temperature, lower strain rate and smaller grain size, and becomes important in both mica-poor and mica-rich granitoid systems. Applying our results to natural fault zones deformed at low strain rates ($10^{-14}–10^{-12}$ s$^{-1}$), the onset of DPC may trigger the frictional-viscous transition at depths corresponding to greenschist facies metamorphic conditions.

## CRediT authorship contribution statement

W. Zhan: Writing – review & editing, Writing – original draft, Visualization, Validation, Methodology, Investigation, Formal analysis, Data curation, Conceptualization. A.R. Niemeijer: Writing – review & editing, Supervision, Resources, Methodology, Data curation. A. Berger: Writing – review & editing, Visualization, Supervision, Resources, Conceptualization. C.J. Spiers: Writing – review & editing, Supervision. F.E. Gfeller: Writing – review & editing, Data curation. M. Herwegh: Writing – review & editing, Supervision, Resources, Project administration, Funding acquisition, Conceptualization.

## Declaration of generative AI and AI-assisted technologies in the writing process

During the preparation of this work the author(s) used chatGPT in order to improve the readability and language of the manuscript. After using this tool/service, the author(s) reviewed and edited the content as needed and take(s) full responsibility for the content of the publication.

## Funding sources

This work is funded by the Swiss National Science Foundation (SNSF) Grant 20002_192124.

## Declaration of competing interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

## Acknowledgements

We thank Andreas Kronenberg and Diane Moore for their constructive comments, which have helped improve the clarity and quality of this manuscript. We thank Thony van der Gon Netscher, Gerard Kuijpers, and Floris van Oort for technical supports of the ring shear apparatus, and Beatrice Frey for support in the FEG-SEM work.

## Appendix A. Supplementary data

Supplementary data to this article can be found online at https://doi.org/10.1016/j.tecto.2025.230772.
# Data availability

Data will be made available on request.

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