Precursory Stress Changes and Fault Dilation Lead to Fault Rupture: Insights from Discrete Element Simulations

1 Precursory stress changes and fault dilation lead to fault rupture: Insights 2 from Discrete Element Simulations 3 4 D. G. Blank1 and J. K. Morgan1

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6 1Rice University 6100 Main Street Houston, TX 77005 7 Corresponding author: David Blank ([email protected])

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12 Key Points: 13 • Earthquake precursors 14 • Physical controls on slow slip 15 • Earthquake stress transfer 16 Abstract 17 18 We use the discrete element method (DEM) to create numerical analogs to subduction 19 megathrusts with natural roughness and heterogeneous fault friction. Displacement boundary 20 conditions simulate tectonic loading, inducing fault slip. Intermittently, slip develops into 21 complex spatiotemporal rupture events that include foreshocks, mainshocks, and aftershocks. 22 Slow slip emerges prior to large events, potentially preparing the fault for subsequent seismic 23 slip. We probe the kinematics and stress evolution of the fault zone to gain insight into the 24 physical processes that govern these phenomena. Prolonged, localized differential stress drops 25 precede dynamic failure, a phenomenon we attribute to the gradual unlocking of contacts as the 26 fault dilates prior to rupture. Slip stability in our system appears to be governed primarily by 27 geometrical phenomena, which allow both slow and fast slip to take place at the same areas 28 along the fault. Similarities in slip behavior between simulated faults and real subduction zones 29 affirm that modeled physical processes are also at work in nature, however, some of these 30 processes may not be detectable in natural systems. 31 32 Introduction 33 34 Understanding deformation behavior that takes place prior to earthquake nucleation is 35 fundamental to our ability to assess seismic hazards, yet the physical controls that govern these 36 processes remain unknown. One avenue for constraining premonitory phenomena is the joint 37 use of geodetic and seismological methods. For example, recent studies suggest that slow slip 38 preceded the mainshock nucleation of the 2011 M9.0 Tohoku-Oki [Meng et al., 2011; Ito et al., 39 2013; Kato et al., 2012] and the 2014 M8.1 Iquique [Ruiz et al., 2014; Kato and Nakagawa, 40 2014] events, spurring increased rates of precursory foreshock activity. Slow slip earthquakes, 41 which are aseismic, might redistribute stresses along a fault zone, preparing it for large-scale 42 seismic slip. In that case, slow slip events could play an important role in earthquake early 43 warning. Unfortunately, the relationship between slow and fast slip within subduction zones is 44 still unclear; for example, inverted source mechanisms, which might be used to constrain stress 45 changes, have relatively coarse resolution and provide no indication of in situ stresses associated 46 with earthquake slip. 47 Experimental and numerical studies, however, provide another avenue toward the understanding 48 of the mechanisms responsible for fault slip processes, as they allow us to look inside well 49 constrained fault systems as they deform. Premonitory slow slip has been widely observed in the 50 laboratory prior to dynamic unstable rupture [Ohnaka and Kuwahara, 1990; Yamashita and 51 Ohnaka 1991; Dieterich, 1992; Ohnaka, 1992]. The transition from stable to unstable slip during 52 this precursory phase has been explained by rate dependent friction [e.g., Dieterich 1992] or slip 53 dependent strength [e.g., Ohnaka 1992]. Neither of these explanations, however, addresses the 54 physical properties or processes that control slip and its stability [Scott, 1996]. Recent frictional 55 sliding experiments [Harbord et al., 2017] have shown that fault roughness and local normal 56 stress may play a crucial role in slip stability and that geometrical heterogeneities must be 57 considered in slip stability analysis. Numerical models that include geometrical complexity 58 [Romanet et al., 2018] or fault roughness [Fournier and Morgan, 2012] demonstrate that both 59 fast and slow slip can emerge under uniform frictional boundary conditions. These observations 60 suggest that changes in fault geometry and contact distributions may be more important to 61 predicting fast versus slow slip, than is assumed fault rheology. 62 63 Given the important role of fault geometry, the stress evolution that accompanies precursory 64 phenomena might hold key insights into the physical processes that govern slip stability. 65 Numerical models allow for the simultaneous monitoring of stresses and kinematics, and 66 therefore serve as a useful tool to investigate potential earthquake precursors. Particle dynamics 67 models, such as the discrete element method (DEM), offer a unique approach, wherein particle 68 assemblages interact according to simplified contact laws, yielding complex kinematic and 69 mechanical behaviors. DEM simulations have proven effective in producing emergent 70 instabilities that resemble earthquakes, along with other slip behaviors [Scott, 1996; Morgan, 71 2004; Fournier and Morgan, 2012; Morgan, 2015], which can be directly correlated to the 72 associated stresses. In this study, we use DEM to simulate slip on a model megathrust fault with 73 variable friction along the interface, to investigate the occurrence of slow and fast slip and their 74 relationships. 75 76 77 Numerical methods and experimental setup 78 79 The discrete element method (DEM) [Cundall and Strack, 1979] is a particle based numerical 80 simulation tool, rooted in molecular dynamics. A full description of RICEBAL, the numerical 81 code implemented here, is presented elsewhere [Morgan et al., 2015]. For simulations of fault 82 slip, DEM offers several distinct advantages over other numerical methods. First, the 83 morphology of faults within the particle assemblage is inherently rough, defined by the discrete 84 particles that bound the fault. These rough surfaces result in emergent stress heterogeneities 85 during slip [Fournier and Morgan, 2012], so it is not necessary to impose stress variations along 86 fault surfaces. Moreover, the assemblage experiences changes in particle configuration in 87 response to deformation (e.g., contraction and dilation), which affects the bulk properties of our 88 simulated rock, resulting in a more natural mechanical response [e.g., Morgan and Boettcher, 89 1999; Morgan, 1999]. 90 91 Our system is designed to represent a simplified subduction zone, where an upper plate is pushed 92 over a lower plate (Figure 1). To achieve this setup, ~100,000 particles of two different sizes (60 93 and 80 m in radius) are randomly generated within an 80km x 20km numerical domain and 94 allowed to settle under gravity. Once settled, bonds are applied to the deepest 4km of the particle 95 assemblage (green), which corresponds to the underthrusting plate. The overriding accretionary 96 prism (black) is then trimmed to a prescribed wedge geometry by removing particles from its 97 surface. Bonds are then applied to the prism assemblage, which has a final height and width of 98 13km and 55km, respectively. The interface between the underthrusting plate and prism units 99 remains unbonded, to localize slip into this zone. The pre-defined surface between these two 100 cohesive units is analogous to a megathrust fault. The backwall of the prism is captured and 101 displaced at a constant rate to induce slip on the fault (Figure 1). 102 103 We introduce heterogeneous interparticle friction along the fault interface, in order to document 104 its effects on fault slip behavior. Interparticle friction is set to zero along the black-green 105 interfaces, and higher along the red-green interface. Although we do not seek to reproduce any 106 specific subduction zone, the prescribed high and low friction distributions along the fault 107 surface are based loosely on natural examples, specifically the inferred low friction at both deep 108 and shallow depths within subduction zones [Moore et al., 2001; Wech and Creager, 2011; 109 Keren and Kirkpatrick, 2016]. In this paper, we analyze results from a single simulation, in 110 which the interparticle friction along the strong interface is set to 0.40. After the model is 111 constructed, the fault is preconditioned [e.g., Fournier and Morgan, 2012] by displacing the 112 backwall until the full length of the fault has slipped, decoupling the wedge. Once this stage is 113 complete, the backwall velocity is set to 0.2 m/s, and slip is monitored over an interval of 960 114 meters of backwall displacement, during which we track particle displacements and the evolving 115 stress field along the fault. 116 117 Our ability to resolve discrete earthquakes in our model, as well as other slip behaviors, allows 118 us to compare simulated slip events to natural examples. We calculate fault slip velocity by 119 dividing the incremental fault displacements by the numerical time step. Slip velocity greater 120 than 3m/s is characterized here as earthquake slip, falling into the range of documented 121 earthquake slip velocities [e.g., Di Toro et al, 2004]. Earthquake slip events that are spatially 122 and temporally continuous are characterized here as individual “events”. Likewise, we refer to 123 continuous episodes of slow slip as “slow slip events”. 124 125 Simulation Results 126 127 The time evolution of slip velocity along the decollement is a key quantity of this dynamic 128 system, as it defines the slip behavior and its controls. Figure 2 shows representative sequences 129 of decollement slip and associated stress changes accompanying two major events, both of which 130 reveal the complex interplay of earthquake slip, creep, and slow slip through time and space, 131 driven by constant back wall displacement. 132 133 Figure 2a shows a sequence of slip that is initiated by three temporally and spatially separated 134 bursts of earthquake slip (A, B, C), embedded within a broad zone of slow slip. The combined 135 slip from this set of foreshocks mobilizes the entire length of the fault and results in 4.8 meters of 136 movement at the toe. In the wake of this foreshock sequence, background displacement rates 137 decrease, although temporally perturbed by small locking and unlocking events that propagate 138 from the back wall. Interestingly, a localized zone of enhanced creep emerges ~27 km from the 139 back wall (D), and persists in the same location for about 12m of back wall displacement. This 140 enhanced creep is characterized by peak velocities of about 0.1m/s, slightly higher than 141 background slip rates in the surrounding areas. The enduring localized creep through D leads 142 directly into the onset of a complex earthquake rupture. This large displacement event, referred 143 to collectively as “event 1” (dashed box in Figure 2a) begins with an impulsive, discrete burst of 144 earthquake slip (red), which then propagates updip towards the toe at a relatively constant rate of 145 0.5 km/s (E). The rupture front then splits into a bidirectionally propagating rupture (F), which 146 reruptures downdip segments of the fault that have already undergone slip. During this phase, 147 fault slip breaks through the toe, accumulating a total of 600 meters of displacement, thus 148 completing the mainshock. The aftershocks begin with two separate events, which nucleate 149 independently and coalesce at (G). Several more aftershock events take place both up dip (H, J) 150 and down dip (I). 151 152 Pronounced stress changes are expected to accompany megathrust slip, and this is confirmed in 153 the time evolution of modeled changes in differential stress along the decollement (Figure 154 2b). To quantify stress change, we calculate average differential stresses within 500x500m 155 domains that span the fault, and difference these values incrementally. The fault zone 156 experiences a mix of both stress rise (red) and drop (blue) through time and space. The series of 157 foreshocks defined in Figure 2a (A, B, C) modify stresses in a complex manner along the fault. 158 The hypocenters of these foreshocks experience immediate stress drops, whereas the lateral 159 migration of slow slip from their hypocenters correlate with stress rises, followed by oscillating 160 stress changes. More broadly, however, all of the foreshocks transfer stress updip; this is most 161 evident for the first foreshock (A), which shows only stress increases to the right of about 32 km. 162 This corresponds to a marked decrease in slip rate (Figure 2a), showing that differential stress 163 rises as the fault locks up. 164 165 An intriguing pattern that emerges for all three foreshocks is the persistent precursory decreases 166 in differential stresses at the future nucleation points of impending events. The same thing is 167 evident immediately following the foreshock sequence, where a gradual decrease in differential 168 stress corresponds to the eventual nucleation point of event 1, and correlates with the localized 169 fault creep (D) identified in Figure 2a. This sustained precursory stress drop is embedded within 170 a zone of broad, backwall-induced stress rise. Updip of this precursory stress drop, a secondary 171 localized stress drop emerges independently. Instead of a broad zone of uniform stress drop on 172 the rupture patch, as is often assumed in models, we observe a complex redistribution of stresses. 173 174 To gain insight into the evolution of absolute stresses leading into earthquake rupture, we track 175 total differential stress at the eventual nucleation point of event 1. As shown in Figure 2c, two 176 discrete stress rises (2.0 and 6.8 MPa) are felt at the event 1 nucleation point, correlated in time 177 with foreshock events (A) and (B). After the second stress rise, the nucleation patch slowly 178 releases 4.0MPa, correlated with the precursory localized creep (D). The nucleation point then 179 experiences a rapid 83.45MPa stress drop, associated with the onset of seismic slip. This part of 180 the fault then regains 63.2MPa amid the coseismic slip and postseismic creep. The nucleation 181 point stress drops recorded here are not averaged over the length of the slip patch, and are thus 182 higher than predicted earthquake stress drops. 183 184 Figure 2d shows a contrasting decollement slip sequence, where localized slow slip events 185 appear to play a more direct role in the rupture initiation. These three zones of localized slow 186 slip (K, L, and N) accelerate above background creep without precipitating earthquakes. 187 Interestingly, Event N propagates laterally at rates comparable to fast earthquake rupture, 188 whereas event K does not propagate laterally and manifests as an “orb” of slow slip. Event L 189 first propagates laterally, but eventually evolves into a localized zone of creep (~29km along the 190 fault). The fault undergoes six episodes of locking and unlocking near the back wall (Interval 191 M), which interrupt the otherwise broad low rate creep across much of the fault zone. However, 192 a slightly elevated rate of creep is observed persistently (O), eventually leading into the initiation 193 of wholesale fault slip. However, slow slip event N first reactivates the creeping segments, and 194 is followed shortly by onset of earthquake slip (P). Nucleation of event 2 takes place at the same 195 part of the fault that had creeped locally during slow slip event L. 196 197 Figure 2e shows differential stress changes that again correspond to decollement slip. Slow slip 198 events L and N share the characteristic stress field perturbation of those caused by A, B and 199 C. They both experience hypocentral stress drops with stress rise propagating outwards with 200 slow slip, and are followed by oscillating stress changes. By comparison, slow slip event K 201 exhibits a relatively simple stress evolution. The maximum magnitude of stress drop corresponds 202 to maximum slip velocity, and the tips of the actively slipping patch of slow slip event K 203 experience stress rises. Similar to the precursory stress drops associated with event 1, event 2 204 shows two simultaneous localized zones of stress drop, apparently triggered by Event L. In this 205 case, the rate of localized stress drop gradually decreases, becomes reinvigorated by event N, and 206 ultimately runs away into earthquake stress drop. Periodic stick slip events near the back wall 207 (interval M) also perturb the stress field, and disrupt the otherwise constant backwall induced 208 stress rise. Event 2 correlates with general updip stress rise, particularly between ~40-50 km 209 along the fault, in contrast to the near-toe stress drop seen during event 1. This highlights the 210 differences in dynamic evolution that accompanies an earthquake that displaces the toe and one 211 that does not. 212 213 Similar to the nucleation point of event 1, the nucleation point of event 2 undergoes a gradual 214 stress drop prior to the onset of dynamic rupture (Figure 2f). The rate of stress change decreases 215 during this phase, as stresses on the asperity apparently stabilize with time. As event N 216 propagates across the nucleation point, a sudden 2.7MPa stress drop takes place, shortly followed 217 by a larger, earthquake related stress drop of 17.0MPa, less than a quarter of the nucleation stress 218 drop of event 1. The fault patch ultimately recovers more stress (17.36MPa) than it lost during 219 the earthquake amid post seismic slow slip, and does so more rapidly than event 1. 220 221 Discussion 222 223 Our models show that simulated fault slip results in a wide range of slip behaviors, including 224 earthquakes that develop distinct spatiotemporal patterns (e.g., foreshocks, mainshocks, and 225 aftershocks), as wells as localized slow slip events. Intriguingly, we find abundant evidence for 226 precursory phenomena that take place prior to the onset of dynamic failure, such as periods of 227 slow slip that ultimately feed into earthquake slip. We see persistent co-located stress drops in 228 association with this type of precursory slow slip. Furthermore, the initiation of these 229 phenomena sometimes coincides with the occurrence of foreshocks, suggesting that foreshocks 230 can trigger long duration nucleation processes that ultimately precipitate mainshocks. In the 231 following discussion, we offer insights into the physical mechanisms that might be directing the 232 phenomena we see. 233 234 Our simulated earthquakes have highly complex source processes, which we attribute to 235 heterogeneities of our modeled fault (similar to many natural faults). The modeled 236 heterogeneities stem primarily from the initial and evolving particle configurations, which 237 introduce roughness to the fault [e.g., Fournier and Morgan, 2012], as well as imposed 238 differences in along-fault frictional strength and gradients in overburden stress. Event 1 (Figure 239 2a) is an excellent example of emergent source complexity, and our results resolve this in detail. 240 The event commences with a period of localized slow slip (D), which does not propagate 241 laterally. After this precursory slip, the mainshock nucleates at depth, propagates updip, 242 displaces the toe, and ultimately re-ruptures at depth again. Technically, this re-rupture (G) is an 243 aftershock that occurs shortly after the mainshock. The rupture front propagates at various rates 244 throughout the duration of the event. The size and complexity of Event 1 is likely due to the 245 failure of multiple neighboring asperities during rupture, as well as temporary locking and 246 unlocking as they drag past one another during slip. In our system, asperities are local geometric 247 protrusions along the fault that are capable of withstanding higher shear stress before failure. 248 There are striking similarities between the rupture propagation behaviors produced in our models 249 and those observed in nature. For example, the 2011 Tohoku-Oki M9.0 event nucleated at depth, 250 propagated updip, displaced the trench, re-ruptured downdip, and did so with variable rupture 251 velocities throughout [Meng et al., 2011; Ito et al., 2013]. Due to the similarity in rupture 252 characteristics between our model and nature, we believe that fault roughness plays a central role 253 in the complexity of earthquake rupture on real faults. 254 255 To first order, the size and lateral extent of the precursory activity seem to correlate with the 256 lateral extent of the eventual mainshock. For example, event 1 in Figure 2a is preceded by a 257 foreshock sequence embedded within an enduring, migrating period of slow slip that displaces 258 the entire length of the fault (A, B, C). The mainshock that follows breaks through the zone that 259 was displaced during the foreshock sequence. By comparison, the three precursory slow events 260 shown in Figure 2d (K, L, N) span a smaller portion of the fault, and the corresponding 261 mainshock (Event 2) is constrained within this smaller zone. This suggests that seismic or 262 aseismic precursory slip sequences can reduce coupling prior to a large earthquake, preparing the 263 fault for a larger main event. Similar behavior has been inferred to have taken place prior to the 264 2011 Tohoku Oki event, based on the rate with which precursory seismic events propagated at 265 depth and by the presence of repeating earthquakes [Kato et al., 2014]. Here, we are actually 266 able to measure this type of slow slip and note its correlation with foreshock sequences. We 267 propose that the slow slip phenomena produced in our simulations are analogous to slow slip 268 events that precede large earthquakes in natural systems, supporting the hypothesis that slow slip 269 can be a precursor to large magnitude megathrust earthquakes. 270 271 As a further test of this hypothesis, we evaluate the stress evolution that drives fault deformation, 272 and find that the complex kinematics of fast and slow slip in our models are accompanied by 273 distinct and systematic changes in the stress field. For example, foreshock A (Figure 2b) is 274 followed by broad updip stress rise, associated with the propagation of slow slip. This updip 275 stress transfer loads the fault and which may precipitate the updip foreshocks B and C. 276 Foreshock stress transfer is clearly seen at the event 1 nucleation point in the sudden 2.0 and 6.8 277 MPa increases in stress state (Figure 2d), which are temporally correlated with foreshocks A and 278 B. Furthermore, we find that low magnitude, sustained precursory differential stress decreases 279 systematically take place at earthquake nucleation points prior to rupture. In the case of Event 1, 280 this decrease in stress also correlates with persistent localized precursory slow slip, which 281 eventually feeds into earthquake slip. The slow nucleation phase of event 1 begins as a direct 282 consequence of stress transfer associated with Event B (Figure 2c). Thus, in our model, 283 foreshock stress transfer is capable of triggering long-lasting nucleation processes that ultimately 284 culminate in large magnitude earthquakes. 285 286 It is commonly assumed that shear stress rise under constant normal stress leads to earthquake 287 failure [Scholz, 2002]. The sustained stress decreases leading up to nucleation show that the 288 very opposite occurs in our models. We propose a geometrical mechanism on a rough fault that 289 can account for both the gradual differential stress reduction and slow slip prior to seismic slip 290 (Figure 3a). In order for shear displacement to occur along the fault, particles must shift up and 291 over one another, which corresponds to a gradual reduction in fault parallel contact forces, 292 proportional to fault shear stress (Figure 3a). This slow particle climbing correlates with slow 293 slip. It is also associated with local dilation along the fault. Once the overriding particles reach 294 the apex, the resisting contact forces are at a minimum, and fast slip can occur. The onset of fast 295 slip coincides with collapse of the assemblage. Thus, slow slip instantaneously runs away into 296 fast slip, and with minimal resistance, the stress transfer is rapid enough to allow the rupture to 297 propagate. This mechanism can account for the gradual precursory reduction in shear stress, 298 accompanied by slow slip, and the sudden transition from slow to fast slip. 299 300 As a test of our geometrical hypothesis, we hone in on the region where rupture initiates in event 301 1. We map out particle configuration within this enlarged zone, as well as the volumetric strain 302 that accompanies it (Figure 3b and c). We see that the asperity controlling event 1 nucleation 303 (dashed box) is characterized by interlocked red and green particles that are slowly shifting up 304 and over one another, opening up a large void along the fault. Once the red particle domain 305 slides atop the green domain, the existing pore space rapidly collapses and slip instantaneously 306 transitions from slow to fast. This is confirmed in the volumetric strain panels, which show a 307 distinct and persistent zone of dilation (orange) at the nucleation point of the earthquake, 308 associated with the gradual unlocking of the asperity. At the instant of rupture initiation, the 309 particles collapse into the void space (blue), and zones that previously accommodated dilation, 310 now experience contraction. We speculate that, in the case of slow slip that does not run away 311 into fast slip, a similar phenomenon may occur, but characterized by lower relief fault roughness, 312 and lower corresponding dilation and contraction, thus minimizing the contractive response we 313 see during slip. Similar correlated changes in configuration and fault stress can be expected 314 along natural faults as well, as they are inherently rough (Figure 3d). Mated rock surfaces along 315 a decollement must first dilate in order to slip. Thus, this local region must undergo similar 316 stress changes to what we see in our models. This phenomenon might not be detectable 317 instrumentally, thus these precursory phenomena may not be evident on natural faults. 318 319 Our results corroborate the idea that geometry plays the dominant role in controlling fast vs. slow 320 slip, even in our simplified model that lacks pore fluid interactions. We note that if pore fluids 321 were included, they would tend to enhance the responses that we see. In particular, as pore 322 spaces open during precursory dilation, negative pore pressures could develop if fault 323 permeability is very low compared to the slip rates (i.e., undrained system). This would serve to 324 strengthen the fault locally, favoring slow slip rather than fast. Conversely, as pore space 325 collapses during earthquake initiation, fluids within the collapsing pore space rapidly become 326 overpressured, decreasing the effective normal stress, further weakening a zone already poised 327 for failure, and thus facilitating major earthquake rupture. Pore fluids may indeed play an 328 important role in the interplay between fast and slow slip, but our simulations suggest that 329 geometry may be the dominant control. 330 331 Finally, our models show that a wide array of realistic fault deformation behaviors can take place 332 on our simulated megathrust fault, characterized by fault roughness, dilatant and contractive 333 deformation, and emergent stress heterogeneities, all without invoking complex constitutive 334 models for friction and rheology. We argue that the observed phenomena instead can be 335 explained largely by simple geometrical mechanisms. Precursory geometric unlocking and 336 accompanying fault zone dilation (Figure 3) can account for the shift from slow to fast slip 337 during earthquake nucleation. We also show that this slow precursory process can be triggered 338 by stress perturbations associated with nearby foreshocks. Extrapolated to nature, this process 339 can explain the delay between “causative” and “triggered” events observed in nature (delayed 340 triggering). That is, stress transfer from foreshocks (or other far-field events) can initiate 341 geometrically controlled long-duration nucleation processes that ultimately precipitate 342 mainshocks. 343 344 Conclusions 345 346 Our numerical simulations produce phenomena whose behaviors are strikingly similar to those of 347 natural systems, and allow us to probe the kinematics and stress evolution that accompany them. 348 We show that small differential stress drops consistently precede dynamic failure, which we 349 attribute to the gradual unlocking of contacts, as particle packings dilate prior to rupture. We also 350 show that foreshocks have the capacity to transfer stress along the fault and load the nucleation 351 point of impending earthquakes, triggering this gradual nucleation process. This provides an 352 explanation for the naturally occurring delay between foreshocks and mainshocks. We propose 353 that the complexity in stress evolution and rupture propagation during earthquakes can arise from 354 the roughness of the fault and the temporary catch and release of local seismic asperities, 355 accompanied by fault dilation and contraction. A wide range of realistic slip behaviors is 356 produced in our models despite the lack of prescribed rate-state dependent friction or other 357 rheologic behaviors, suggesting that slip stability is controlled primarily by changes in fault zone 358 roughness and attendant mechanical responses. 359 360 References Cited 361 362 Atkinson, Barry Kean. "Subcritical crack growth in geological materials." Journal of 363 Geophysical Research: Solid Earth89.B6 (1984): 4077-4114. 364 365 Cundall, Peter A., and Otto DL Strack. "A discrete numerical model for granular 366 assemblies." geotechnique 29.1 (1979): 47-65. 367 368 Dieterich, James H. "Earthquake nucleation on faults with rate-and state-dependent 369 strength." 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458 Wech, A. G. & Creager, K. C. Automated detection and location of Cascadia tremor. Geophys. 459 Res. Lett. 35, L20302 (2008).

460 Yamashita, Teruo, and Mitiyasu Ohnaka. "Nucleation process of unstable rupture in the brittle 461 regime: a theoretical approach based on experimentally inferred relations." Journal of 462 Geophysical Research: Solid Earth 96.B5 (1991): 8351-8367. 463 464 Zhao, Dapeng, et al. "Structural heterogeneity in the megathrust zone and mechanism of the 465 2011 Tohoku-oki earthquake (Mw 9.0)." Geophysical Research Letters 38.17 (2011). 466 467 Figure 1. Experimental setup. Upper plate (black and red) and lower plate (green) are bonded. 468 Unbonded interface between the upper and lower plates defines a pre-existing decollement with 469 varying friction. Backwall particles (dashed blue box) move inward at a constant velocity. 470 Particles at bottom of footwall (dashed orange box) are fixed at zero velocity. Interparticle 471 friction between black and green particles is set to 0.0, and between red and green particles to 472 0.40. 473 474 Figure 2. Slip behavior and corresponding changes in differential stress along the decollement 475 as a function of back wall displacement. (a) Fault slip velocity map for interval encompassing 476 complex rupture Event 1 (dashed box). Red zones identify slip velocities greater than 3 m/s, 477 designated here as earthquakes. Slip velocities between 0.2 and 3 m/s are plotted in yellow, and 478 represent slow slip. Creep from 0.0 to 0.2 m/s is plotted on the color scale. (b) Differential stress 479 change along decollement corresponding to slip velocity map in (a). (c) Absolute differential 480 stresses at the nucleation point of Event 1 (star). (d) Fault slip velocity map for interval 481 encompassing Event 2 (dashed box), including several precursory slow slip events (K and L). 482 Colors as in (a). (e) Differential stress change along decollement corresponding to slip velocity 483 map in (d). (f) Absolute differential stresses at the nucleation point of Event 2 (star). 484 485 Figure 3. (a) Schematic representation of the dilative rupture process, where the shear 486 component of stress relative to the normal component decreases. (b) Particle assemblage during 487 event 1 nucleation. Pore space expands prior to rupture at the nucleation point (dashed box) and 488 collapses as the earthquake begins. (c) Volumetric strain panels show dilation (warm colors) 489 leading to failure and contraction (cool colors) accompanying it. (d) Schematic geometrical 490 asperity along a megathrust fault that would experience similar unlocking and associated stress 491 changes to failure. 492 493 494 495 496

a) b) c) a)408 b) c) 408 I J I J 396 I J I J 396 H H 384 Event 1 H H 384 G G Event 1 G G +63.2 MPa 372 +63.2 MPa 372 F F 360 F F 360 E E E E Onset of dynamic rupture 348 Onset of dynamic rupture 348 -83.5 MPa -83.5 MPa 336 DD D Back wall displacement (m) 336 -4.0 MPa Back wall displacement (m) -4.0 MPa 324 324 C C B C +6.8 MPa 312 B B +6.8 MPa 312 B AA A +2.0 MPa 300 A +2.0 MPa 300 10 20 30 40 50 10 20 30 40 50 10 20Distance along fault (km)30 40 50 10 20 Distance along fault (km)30 40 50 d) Distance along fault (km) e) Distance along fault (km) f) d)912 e) f) 912 Event 2 Event 2 900 O P O 900 P Onset of dynamic rupture -17.0 MPa-17.0 MPa Onset of dynamic rupture Partial stress release 888 N N Partial stress release 888 N N -2.7 MPa-2.7 MPa

876876 O O 864864 MM MM

852852 Back wall displacement (m) Back wall displacement (m) 840840 L LL

828828 K KK 1010 2200 3030 4400 5050 10 2020 3030 4040 5050 Distance along fault (km)Distance along fault (km) Distance along fault (km)Distance along fault (km) 497 498 499 500 501 502