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The M9.2 Sumatra–Andaman earthquake (SAE) occurred three months prior to the M8.7 Nias earthquake (NE). We propose that the NE was mechanically triggered by the SAE, and that poroelastic effects were a major component of this triggering. This study uses 3D ﬁnite element models (FEMs) of the Sumatra– Andaman subduction zone (SASZ) to predict the deformation, stress, and pore pressure ﬁelds of the SAE. The coseismic slip distribution for the SAE is calibrated to near-ﬁeld GPS data using FEM-generated Green's Functions and linear inverse methods. The calibrated FEM is then used to predict the postseismic poroelastic contribution to stress-triggering along the rupture surface of the NE, which is adjacent to the southern margin of the SAE. The coseismic deformation of the SAE, combined with the rheologic conﬁguration of the SASZ produces two transient ﬂuid ﬂow regimes having separate time constants. SAE coseismic pore pressures in the relatively shallow forearc and volcanic arc regions (within a few km depth) dissipate within one month after the SAE. However, pore pressures in the oceanic crust of the down-going slab persist several months after the SAE. Predictions suggest that the SAE initially induced MPa-scale negative pore pressure near the hypocenter of the NE. This pore pressure slowly recovered (increased) during the three-month interval separating the SAE and NE due to lateral migration of pore ﬂuids, driven by coseismic pressure gradients, within the subducting oceanic crust. Because pore pressure is a fundamental component of Coulomb stress, the MPa-scale increase in pore pressure signiﬁcantly decreased stability of the NE fault during the three-month interval after the SAE and prior to rupture of the NE. A complete analysis of stresstriggering due to the SAE must include a poroelastic component. Failure to include poroelastic mechanics will lead to an incomplete model that cannot account for the time interval between the SAE and NE. Our transient poroelastic model explains both the spatial and temporal characteristics of triggering of the NE by the SAE. © 2010 Elsevier B.V. All rights reserved.
Sumatra–Andaman and southern Nias segments (Fig. 1) (Ammon et al., 2005). The physical cause of this seismic barrier is likely due to a change in the seaﬂoor morphology, associated with the northern extent of a series of prominent offshore islands that are the result of the uplift of the seaﬂoor above the subducting oceanic ridge. The subduction of the extinct oceanic ridge may have caused a permanent kink in the subducting slab, which has been imaged by Pesicek et al. (2008). The seismic barrier also corresponds to a pronounced bend in the geometry of the oceanic trench at the toe of the megathrust (Fig. 1). Ridge subduction creates seismic barriers in numerous subduction zones worldwide (Gutscher et al., 1999; Ruff, 1996). The subduction of the Chile Rise at the southern terminus of the 1960 Mw = 9.5 megathrust earthquake is one example (Plafker and Savage, 1970). Seismic barriers may also correspond to changes in the coupling between the overriding plate and the subducting oceanic lithosphere at the megathrust due to the thinner sediment load above the subducting ridge. Seismic barriers thus correspond to a pronounced change in physical properties along the megathrust.
2006). 2005)..2 rupture (http://neic. Undrained or drained conditions imply either ﬂuid-ﬂux equals zero or pore pressure equals zero. 1999).usgs. illuminate the surface projection of the M9. poroelastic effects. 2007. Coulomb stress Stress-triggering is a mechanism for which a loading event. Nicobar Islands (NI). However. such that f′ is theoretically unbounded (−∞ b f′ b ∞) (Beeler et al. a problem domain including compliant fault zones no longer satisﬁes the homogeneous assumption required by standard analytical models for displacement due to an elastic dislocation (e. Seismotectonic setting of the SAE (adapted from Masterlark and Hughes (2008)). (1) is modiﬁed to σc = σs + f ′ σn ð2Þ The coseismic and postseismic deformation of the SAE changed the stress regime of the Sumatra–Andaman subduction zone (SASZ). Indo-Australian Plate (IAP).. Eq. 2006..65 and 0. Fault-slip from an earthquake induces relatively instantaneous incremental changes in stress and pore pressure in the near-ﬁeld region. King et al.290 K. Mignan et al. Harvard CMT Focal mechanisms are given for the 26 December 2004 (M9. However. Abbreviations are Andaman Islands (AI). Subarya et al. marks the boundary between rupture of the M9.. Great Sumatran Fault (GSF). The assumption that pore-ﬂuid pressure is proportional to fault-normal stress alone holds only if the fault zone is relatively compliant with respect to the surrounding materials (Cocco and Rice. 2006..g. preexisting fault. Hsu et al..gov). Alternatively. 2000). while afterslip. We assume that the state variables σ and P are incremental changes with respect to a reference state.. 2006. We construct poroelastic deformation models of the SAE to test the hypothesis that pore ﬂuid pressure near the hypocenter of the NE continually increased during the three-month interval separating the SAE and NE.2 and subsequent M8. where f′ is an apparent coefﬁcient of friction that is some unknown combination of material properties and transient ﬂuid-ﬂow conditions (Masterlark and Wang.H. such as slip along a fault.. .85 (Byerlee.. P is pore pressure.7 NE was triggered by the M9. Aftershock epicenters (yellow dots). 2000. 2002. 1978). Masterlark and Wang..7 events (seismic barrier). Sunda Plate (SP). 2003.ngdc. which thereby increases during the three-month interval between earthquakes and advanced the occurrence of the NE (Hsu et al. McCloskey et al. because Coulomb stress is often calculated for regions saturated with aftershocks along multiple faults (rather than a single fault). 2. which are known to signiﬁcantly inﬂuence Coulomb stress calculations over time periods consistent with the separation of the SAE and NE events (Beeler et al. Previous stress-triggering analyses of the SAE predict that the SAE increased the Coulomb stress near the hypocenter of the NE (Gahalaut and Kalpna.1. viscoelastic relaxation. changes in Coulomb stress are often calculated using the assumption that pore-pressure is proportional to the fault-normal stress rather than the mean-normal stress used in standard poroelastic theory (e. Masterlark and Wang. 2000). Masterlark. 2006a). The change in Coulomb stress (σc) for a given fault is σc = σs + f ðσn + P Þ ð1Þ Fig.g.. such as postseismic viscoelastic relaxation and afterslip. based on the proximity of these two great earthquakes in both space and time. 2005. Masterlark and Hughes. neither of these analyses included a transient mechanism to account for the threemonth interval separating the SAE and NE.L. 2000). negligible viscous relaxation) or long times (drained conditions. spanning 26 December 2004 through 28 March 2005. Hughes et al. 2003). The sharply truncated aftershock distribution.. 1994. 2000).noaa. 2000. Laboratory experiments on a variety of rocks indicate that the coefﬁcient of friction is robust and lies between 0.7) earthquakes. respectively. 1992). negligible deviatoric stresses in the viscous material) following a dislocation (Wang. 2006.. 2000. The rupture initiated on the southeast portion of the fault and propagated 1200 km northward.. Simeulue Island (SI). impose time-dependence on the Coulomb stress. However.. We propose that the M8.g. 2008). (2) leads to important prediction errors (Beeler et al. where σs is shear stress parallel to a speciﬁed slip vector. shown with a NE-trending dashed line that bisects Simeulue Island. / Earth and Planetary Science Letters 293 (2010) 289–299 2002). In this case. none of these previous analyses account for the pore pressure effects. Okada. and West Sumatra Fault (WSF). 1. Black triangles are near-ﬁeld GPS sites (Gahalaut et al. 2003. Masterlark and Wang.gov). Methods 2. changes the frictional stability of other faults in the near-ﬁeld region. Cocco and Rice. This problem can be extended to models that simulate distributions of material properties that do not include weak fault zones (Chlieh et al.2) and 28 March 2005 (M8.. King et al. Masterlark and Wang. and interseismic strain accumulation drive transient changes in stress and pore pressure after an earthquake has occurred. This increasing pore pressure translates to increasing Coulomb stress and thus predicts a systematic decrease in fault stability leading to the rupture of the NE. Positive values of the change in Coulomb stress indicate an increased tendency for the fault to slip and negative values indicate increased stability. Masterlark. 2000). Coulomb stress calculations allow us to quantify the changes in tendency for frictional slip to occur along a locked. Pollitz et al. and f is the coefﬁcient of friction (e. Others suggest transient mechanisms. σn is faultnormal stress (tension-positive). Eq.2 SAE. 1994. Because of these contradictory assumptions. Stein. The tectonic conﬁguration is modiﬁed from Bird (2003) and overlies a shaded relief image of global relief data (http://www. Static (vis-à-vis quasi-static) stress-triggering analyses of the causal relationship between earthquakes are applicable for either short times (undrained conditions. Burma Plate (BP).
and x3 are equivalent Cartesian coordinates x. 2008. 2003). A matrix of GFs for the entire suite of m nodepairs is assembled by implementing an algorithm that systematically generates the unit dislocation and welding conﬁgurations over the rupture. seismogenic data (Ammon et al. / Earth and Planetary Science Letters 293 (2010) 289–299 291 2. for displacement due to slip along a fault. we sweep this two-dimensional cross section along the curving strike of the Sunda trench to produce a three-dimensional model (Fig.000. The far-ﬁeld lateral boundaries and base of the problem domain are zero displacement.simulia. we specify a distribution of rheologic properties over the partitioned problem domain (Fig. A vector of Green's Functions (GFs). The total longitudinal equivalent strain is: ε = εe + εf and dεf η = Aσd dt ð5Þ where εe is the elastic strain.000 octahedral ﬁnite elements having trilinear interpolation basis functions and 1. 2005. We use the generalpurpose FEM code Abaqus (http://www. A slip event induces relatively instantaneous incremental changes in stress and pore pressure and coseismic deformation is thus undrained. This tessellation is validated by Masterlark and Hughes (2008). and d is a 1 × n column vector of three-component displacements and/or displacement derivatives of the GPS site positions that can be time- . The subscript i spans orthogonal direction components 1.. Seismicity data (Engdahl et al. Melosh and Raefsky. and uz. While forward models allow us to predict deformation caused by fault-slip. x2. and (5) by setting P = 0 and assuming steady-state conditions. Viscoelastic behavior in the mantle is simulated by imposing an additional stress-dependent creep relationship.2.. where G is an matrix of GFs. 2007) constrain the geometry of the subducting slab. and µf is the pore-ﬂuid viscosity. and 3 and the subscript k implies summation over these three components. This conﬁguration implicitly assumes the geologic structure is constant along the trench. P) over a 3D problem domain partitioned into elastic and poroelastic regions. the governing equations for poroelastic materials are G∇ ui + 2 2. the location of the rupture is constrained by seismicity (Fig. north. Masterlark and Hughes. The relationship is equivalent to a Maxwell material for η = 1 and A is half of the inverse of the linear viscosity (Turcotte and Schubert. Second. 2005). Masterlark and Hughes. is calculated by predicting the displacement of GPS site positions caused by a unit dislocation for a given node-pair while simultaneously welding the remaining node-pairs. The governing equations for an elastic material are recovered from Eqs. Quasi-static fault-slip can be simulated with an FEM as the dislocation of a node-pair. 1980. 2008. 1981. The 3D ﬁnite element model (FEM) presented in this study is designed to simulate coseismic and poroelastic postseismic deformation of the SAE. εf is the strain due to viscous ﬂow. 2007. x1. The characteristic dimension for elements is a few kilometers near the fault and generally increases with distance from the fault. respectively. Kopp et al.H. Sε is the constrained storage coefﬁcient. α is the Biot–Willis coefﬁcient. The viscoelastic rheology is speciﬁed for the mantle only in a separate FEM used to predict longterm postseismic deformation discussed later. εkk = Σ∂uk/∂xk is the volumetric strain. The fault-slip of the SAE occurs along the interface separating the subducting slab. Fourth. 1). u1. However.. Smith. with all time derivatives equal to zero. (3). A is a constant that can be augmented to account for temperature dependence. The elastic and poroelastic properties are taken from compilations of laboratory experiments (Turcotte and Schubert. and z (east. we calibrate the slip distribution via least-squares inverse methods that account for the distribution of material properties within the 3D problem domain of the FEM. 2003. F is a body force per unit volume. The top of the problem domain is a stress-free surface. we specify boundary conditions and impose fault-slip. we can use geodetic or seismic data. κ is the permeability coefﬁcient. This conﬁguration is similar to that of Masterlark and Hughes (2008) and includes both lateral and vertical rheologic variations that correspond to the regionalscale geologic structure of a subduction zone. ν is Poisson's ratio (drained). FEM conﬁguration Construction of the SAE FEM involves a series of steps. 1980.. respectively. P = 0) and undrained (short time. implemented via kinematic constraint equations (Masterlark. Third. uy.. 2000). The conﬁgurations of the volcanic arc and backarc basin regions are based on geologic maps and cross-sections of the SASZ (Barber et al. 2). Kopp and Kukowski. and u3 are equivalent to ux. 2003). Similarly. The curved surface of the rupture comprises an assembly of node-pairs along an internal boundary of the FEM problem domain. 2000). having undrained values of Poisson's ratio substituted into the poroelastic portions of the problem domain. The elastic properties are in accord with seismic tomography and gravity data (Kieckhefer et al. This is a useful result because m separate FEM calculations are required to assemble the matrix of GFs and the computation time for an elastic FEM is substantially lower than that of a coupled poroelastic FEM. Kopp et al. Hughes and Masterlark. consisting of depleted mantle capped by mid-oceanic ridge basalt (MORB). m is a vector of dislocations. 2002. Deformation model Deformation models provide the linkage between the observed surface deformation and the source of the deformation—the fault-slip at depth. 2002) and enriched mantle wedge (Kieckhefer et al. 2005). 2). executes the FEM. and extracts the predicted displacements caused by the dislocation of each node-pair. inverse models estimate the distribution of faultslip. In this formulation.. Therefore. 2). The forward solution for elastic deformation due to a distribution of dislocating node-pairs is Gm = d ð6Þ G ∂2 uk ∂P =α −Fi ð1−2νÞ ∂xi ∂xk ∂xi ð3Þ α ∂εkk ∂P κ 2 ∇ P+Q + Sε = μf ∂t ∂t ð4Þ where G is the shear modulus. For the SAE. by substituting drained and undrained values of Poisson's ratios. 2002). and σd is the deviatoric stress. 2. In both cases. 2002.. The subducting slab and overriding plate are welded together along the intersection of the fault and the trench. 2002. To delimit the geometry of the rupture surface. based on observed deformation and pre-supposed deformation models. Expressed in index notation.. Stein and Okal. The governing equations for elastic materials are sufﬁcient to describe limiting cases of drained (long time. no ﬂuid ﬂow) static deformation. and Q is a ﬂuid source term deﬁned as volume of ﬂuid per unit bulk volume per unit time (Wang. while simultaneously accounting for the known geologic structure of the subduction zone.. (3) and (4) (Wang. The FEM is driven by a coseismic slip distribution calibrated to near-ﬁeld GPS data.com) to solve for displacement (u) and coupled displacement and pore-ﬂuid pressure (u. 1974).000 degrees of freedom. a model and some assumptions are required.3. a description of transient poroelastic deformation requires both Eqs. (4). u2. 2003. respectively. Kopp and Kukowski. and the overriding forearc (Kopp and Kukowski. 2008). The tessellated problem domain comprises about 340.L. First.K. 2000). Wang. Hughes et al. and vertical). Kopp et al. y. the (undrained) coseismic deformation is calculated using the simpliﬁed elastic governing equations. we design a trench-normal slice through the SASZ (Fig. respectively. Rather than use a published slip distribution for the SAE. and results from previous modeling studies (Chlieh et al.
(7) using second-order Tikhonov regularization to damp the null space of the data kernel (Aster et al. Similarly. we pre-multiply Eq. The 3D FEM is constructed by sweeping the 2D proﬁle along the curvature of the Sunda Trench. (a) Conceptual model. Second. Inverse methods We apply linear inverse methods to calibrate the slip distribution of the SAE. 2) Imposes positive thrust and right-lateral strike-slip components. along-strike and down-dip boundaries are set to Dirichlet (null) boundary conditions and the up-dip boundary is set to Neumann speci ﬁ cations (∂m /∂x = 0) (Fig. The FEM comprises about 340. (b) FEM design and conﬁguration. this matrix of FEMgenerated GFs is readily calculated for inverse analyses of deformation data for dislocations embedded in an arbitrary domain (Masterlark. Each patch comprises four node-pairs sharing slip characteristics. dependent. The down-dip and strike-slip sub-matrices of L are independent of one another but share the boundary condition speciﬁcations. Each coefﬁcient Gij represents the contribution to the displacement of dj due to unit dislocation of node-pair mi. 2006) and span northern Sumatra and the Nicobar and Andaman Islands (Table 1). This 2D proﬁle illustrates the geologic structure of the subduction zone. and 4) Accounts for the relative uncertainties of the GPS data.. we reconﬁgure Eq.H. simultaneously 1) Estimates the slip distribution that minimizes misﬁt to GPS data.000 elements. indicated by the distribution of aftershocks (Fig. into a 25 (alongstrike) × 7 (down-dip) grid of quadrilateral slip patches. 2. 2003. 2.. The top of the problem domain is a free-surface. 2006. (6) to account for the relative uncertainties of the data WGm = Wd = Gw m = dw ð7Þ where W is a diagonal data weighting matrix constructed from reported GPS measurement uncertainties. (Table 1).. These data are compiled from previous studies (Gahalaut et al. 2005) ðGw Gw + β L LÞm = Gw dw and L = T 2 T T  Ldd 0 0 Lss  ð8Þ where L is a 2m × 2m matrix of coefﬁcients for the ﬁnite difference approximation of the Laplacian operator for ∇2m = 0 over the 2D rupture surface. (6) into a forward model that when inverted.4. We limit our study to near-ﬁeld deformation data because the relative data importance (Menke. 1989) of far-ﬁeld GPS sites (more than a fault-width from the rupture) is insigniﬁcant compared to that of GPS sites within the surface projection of the rupture for a megathrust event (Hutton et al. G = (GddGss) and has dimensions of 2m × n. We partition the curved rupture surface. 2001). Wii = 1/σi. For both down-dip (dd) and strike-slip (ss) dislocations. Most importantly. 1).292 K. Conceptual model and FEM conﬁguration (adapted from Masterlark and Hughes (2008)). We then recast Eq.L.. based on observed near-ﬁeld displacement data from 34 GPS sites in the near-ﬁeld region. Hughes et al. 3) Damps spurious solution oscillations. 3a) (Wang and . The exploded view reveals the likeness of the FEM to the conceptual model. the dislocation vector has dimensions 2 m and m = (mddmss)T. 2008). Lateral and bottom boundaries are zero displacement. Subarya et al. / Earth and Planetary Science Letters 293 (2010) 289–299 Fig. Thus. FEMs permit us to simulate variable dislocations along fault surfaces embedded in a problem domain partitioned into the 3D rheologic conﬁguration expected for the SASZ. First. Masterlark and Hughes.
0916 0.0666 0.0397 0.48000 5.69600 9.1421 East 0.08665 2.3763 − 2.9320 UGRH 92.0264 0.6823 NIND 98.0354 − 0..0200 0.0611 − 0.0053 − 0.43378 5.27800 13.03600 7.0100 0.8209 − 2.7200 − 2.0200 0.0869 0.6200 − 1.7600 − 5.0127 − 0.7634 − 1.0200 0.9999 LHOK 97.0200 0.60702 4.0500 0.1211 0.6760 HBAY 92.5690 CARN 92.4873 LANG 97. Fault-slip is concentrated along the up-dip portion of the rupture that is west of northern Sumatra.1038 0.7719 − 2.0631 0.0198 − 0.3993 − 0.0830 0. subject to positivity constraints (Menke.0675 − 1. 1982). 2006) bm12 98. °E Lat.7109 Up − 0.0405 0.63100 12.0500 0.0100 0.0669 0.0572 (Gahalaut et al.8188 PIDI 95. based on the diagonal elements of the parameter resolution matrix.0100 0.2716 K515 95. The least-squares solution to Eq. In the absence of these constraints. 3.3420 − 0.3500 − 2.03600 11.0535 − 0.3654 R175 95.0129 − 0.0462 − 0.9100 − 3.9333 PISU 99.14524 5. These positivity constraints are different from Masterlark and Hughes (2008) where solutions were allowed to contain both thrust and normal components.44756 2. 1989).8040 TERE 93.1717 − 0.1140 − 2.6600 − 0.2700 − 5.1800 − 0. 2005). reﬂecting some aspect of the FEM that fails to adequately represent some unknown complexity in that region of the SASZ.0240 0.0065 − 0.0041 − 4.0238 0.0649 0.0700 0.3600 0.4143 − 2.3681 − 0. solutions for m will include more oscillatory distributions having both normal and left-lateral slip regions that are not compatible with the focal mechanism (Fig.0807 0.52419 2.0805 − 0.0300 − 1.1240 KARD 93.5075 k504 95.0200 0.0608 0. GCV results suggest β2 = 0.0882 0.0662 0.1200 0.3600 − 3.5710 0.9557 − 0.0200 0.5490 MERO 93.0400 0.0200 0. (9).0500 0.0100 0.0988 − 0.68602 2. we can use the trade-off curve for roughness versus misﬁt (Gubbins. / Earth and Planetary Science Letters 293 (2010) 289–299 Table 1 Near-ﬁeld GPS data.6010 − 0.0411 0.24116 5.0200 0.1188 0.0761 − 1.0414 0.0100 0.91856 2.1000 − 1.0326 0. 3b).7730 GOVI 92.0669 0.1054 0.3900 − 1.1472 SIPA 99.1585 MART 98.0400 0.72953 1.6599 − 0.6000 0.9830 PBLR 92.0228 0.0411 − 1.0700 − 2.71287 − 0. (9) for a given regularization parameter (Aster et al.30200 8. This slip distribution is rather rough and includes an unrealistic slip maximum of ≥ 100 m along the northeast edge of the rupture.5622 R171 95.H.0270 EAST 93.K. 1). Rm (Aster et al.1144 0.1600 − 1.0855 Up 293 0.1427 0.0434 0.0270 − 0.0100 − 0.0586 − 0.7100 − 0.0600 0.0733 0.0253 0.9600 − 0.0699 0.0617 0.42753 5.8537 − 2. 4).1087 1σ (m) North 0.3600 − 0..9449 d962 97.0218 − 1.0100 0.1426 − 3.0246 − 0.0890 TIGA 98.9340 2. The positivity constraints require that solutions contain combinations of thrust and right-lateral strike-slip.0426 0.64900 11.0500 0.5183 R174 95.0277 0.10263 2. Hughes et al..0143 − 0.0332 0.1312 − 0. 2005) Rm = ðGw Gw + β L LÞ T 2 T −1 Gw Gw : T ð11Þ ð10Þ The small patch of signiﬁcant slip at the northern edge of the rupture (Fig.0100 0.0825 − 0.1745 Displacement (m) North − 0.9500 − 3.9600 − 0.0236 0.1000 − 2. 2006) ABAY 93.0388 0.1228 − 0.33080 2.0470 LONG 92.9100 − 4.0200 0.64259 1.0277 0.0300 Anderson.0899 0.0908 13.4800 − 0.L.0100 0.0478 0.0671 0.0119 0. The generalized cross validation method (GCV) provides a means for selecting an optimal regularization parameter by minimizing the functional V(β) V ðβÞ = n‖Gw mβ −dw ‖2 2 2 T −1 T 2 Trace½I−ðGT G  w Gw + β L LÞ where I is a 2m × 2m identity matrix and mβ is the solution of Eq.9500 − 1. even though these unconstrained solutions ﬁt the GPS data better than their constrained counterparts. This pattern is reasonably well resolved (Fig.21600 12.2435 K505 95.9600 − 2. (8) is (Aster et al.0841 0.0637 0. 2005) m = ðGw Gw + β L LÞ T 2 T −1 Gw d w : T ð9Þ We solve Eq.4349 − 2.0558 0.1900 − 2.7455 − 1.0100 0.0490 0.0859 0. 2004) and a priori fault-slip constraints to identify a slip distribution that simultaneously minimizes roughness and misﬁt and has a maximum slip magnitude of about 30 m (Fig.0100 0.4465 D972 96.0500 0. 4b).1034 0.17800 10.2030 R176 95. Resolving this issue is the subject of ongoing analyses and beyond the scope of this study. Results We wish to select a solution that gives a balance of misﬁt and smoothness. Site Lon.00400 − 3.67586 5. while sweeping through β parameter space to ﬁnd optimal solutions for m.0600 0.51400 7. °N East (Subarya et al.0400 0.5779 − 0.5100 − 1.0882 0.0100 0.1079 0. Alternatively.0100 0.8600 − 3.0420 0.6500 − 2.0355 − 0.4546 − 0.1386 0.9700 − 4. this misﬁt near the northern edge of the SAE rupture .1217 0.5410 CAMP 93.1057 0.0600 − 1.95996 4. 3c) in accord with other studies.9000 − 3.4900 0.1448 − 0. The regularization parameter β controls the tradeoff between minimizing misﬁt and satisfying the Laplacian operator.1027 − 0.0311 0.0989 − 0.0890 − 0.17441 3.2266 − 0.5838 − 0.1100 − 2.0400 0.0305 0.2190 − 0.006 is statistically the best solution (Fig.0597 0.2600 − 1.0765 − 0.0649 0.0200 0.7506 PAND 98.2031 − 2.7210 PASG 92.0100 0.84193 5. A band of lower magnitude fault-slip occurs sub-parallel to the Sunda Trench and beneath the Nicobar and Andaman Islands (Fig. 4a) is probably a numerical artifact.22500 8. Other investigators report similar problems of GCV results producing rough solutions and having slip magnitudes that are much too high (Freymueller et al.3221 − 2.0873 0.0200 0.0452 2.0230 0.0200 0.8400 − 2..8500 − 1.3877 R173 95.37600 12. Nonetheless.0418 0.7100 − 2.2161 0.56851 4. 1994). This is our preferred solution..0100 0.0458 0..6245 Jahe 98. as discussed below.
occurs at the knee of the curve. Inverse methods. Hughes et al. . 2007). 2000. The oceanic crust capping the subducting slab is also poroelastic to a depth of 50 km. the maximum limit of rupture depth (Hyndman. (c) Vertical deformation.. and backarc sedimen- tary basin are poroelastic materials (Fig. 3. 2009). A Laplacian operator is applied via ﬁnite-difference methods to smooth the estimated slip distribution. Fisher. concentrated west of northern Sumatra. 1998) and seismological observations (Audet et al. (a) Smoothing. 2006). geodetic based analyses (Chlieh et al. This solution produces a good balance of misﬁt and roughness. 4a) is applied to a 3D FEM to predict the transient postseismic poroelastic deformation.. Fig.L. (b) Parameter resolution and prediction misﬁt. The optimal solution (β2 = 0... Grilli et al. 2003). 2007) and tsunami-genesis results (Fujii and Satake. Ioualalen et al. The top (free surface) and lateral boundaries of the poroelastic materials are noﬂow boundaries. and limited by the time interval between the two earthquakes (Fig. is consistent with seismological estimates (Ammon et al.006). 2007). (b) GCV and maximum slip versus β.03.H. However. The pattern of shallow fault-slip. (c) Tradeoff curve for roughness versus misﬁt. 2007. 2007. Rhie et al. A value of β2 = 0. the focus of this study.294 K. 5). 2007) place the majority of fault-slip near the Nicobar Islands.. (a) Estimated fault-slip distribution and deformation predictions. 2). In this FEM. Our calibrated slip distribution (Fig. volcanic arc. forearc. 2005. / Earth and Planetary Science Letters 293 (2010) 289–299 Fig.. includes an unrealistic fault-slip maximum of more than 100 m. our preferred solution. The poroelastic properties of the oceanic crust are constrained by laboratory permeability experiments (Wang. 4. does not inﬂuence predictions for the southern portion of the SAE. The permeability of the poroelastic materials is κ = 10− 16 m2. The predicted hinge line delimits zero vertical deformation and agrees with ﬁeld observations (Meltzner et al. Predicted slip and deformation of the SAE (adapted from Masterlark and Hughes (2008)). as well as a maximum fault-slip of about 30 m.. an estimate for the bulk permeability of oceanic crust (Masterlark. according to GCV. the accretionary wedge.
The increase in pore pressure (2. but prior to the NE.1 MPa and 0.5 2. 5. (1)) and thus the systematic decrease in fault stability following the SAE. A bulk permeability above 1 × 10− 16 m2 cannot account for the 90 day time interval between earthquakes. for which it is customary to assign postseismic deformation mechanisms according to characteristic time constants of observed deformation epochs (Paul et al.3 4.. 1988). 1988). is two orders of magnitude greater than the minimal threshold. as time progresses ﬂuids ﬂow and re-equilibrate in response to the coseismic pore pressure gradients in the region of decreased pore pressure near the hypocentral location of the NE (the response).7 7.6 0. we deduce that the pore pressures range from 105 to 107. Christensen and Ramananantoandro. Hughes et al.7 Whole model (MPa) Max 4. Table 2 Pore pressure maxima and minima. and bulk permeability below 1 × 10− 17 m2 is not geologically reasonable for cold. along the rupture surface of the NE.. These predictions suggest that the coseismic pore pressure distribution leads to a more persistent ﬂow regime near the seismic barrier.4 − 4. First.. Furthermore. pore pressures range from − 0.. the sensitivity to the slip distribution was tested using an average slip of 15.6 − 1.g. Including this permeability instead of a no ﬂow boundary would not signiﬁcantly change our results. This has implications for studies of postseismic deformation. 6a).7 to 4. Our model simulates the relatively instantaneous coseismic poroelastic deformation and subsequent transient poroelastic deformation over the 90-day time interval between the SAE and NE. permeability. 2002. and thus Coulomb stress. 2009). We constrain these assumptions based on geological and seismological evidence for the SASZ (Barber et al. 2008). The absolute changes in pore pressure far from the seismic barrier are 4. Additionally. slip distribution. stress.. the oceanic crust juxtaposed to the seismic barrier experiences an increase in Coulomb stress (McCloskey et al. From these analyses. 6b). the range of pore pressure in the subducting oceanic crust increases when considering the whole model (Table 2). / Earth and Planetary Science Letters 293 (2010) 289–299 295 Bulk permeability for the oceanic crust above 1 × 10− 16 m2 cannot account for the time delay between earthquakes due to rapid pore pressure re-equilibration. 2003..25 m for each fault patch. as well as through the seismic barrier to Fig.. believed to trigger slip.7 to 0. which is geologically reasonable (Audet et al. The poroelastic model allows for pore pressure recovery following the SAE.6 MPa in the deep forearc (about 10 km depth) and − 2. brittle subducting oceanic crust (Fisher. Fisher. We envision triggering of the SAE and NE in both space and time as a two step (impulse and response) process. Following the coseismic time step. even though the speciﬁed permeability is constant for all poroelastic materials in the FEM. 6a) due to the coseismic rupture of the SAE (the impulse). Christensen and Ramananantoandro.0 MPa) southeast of the seismic barrier. this variation in recovery time is due to the down-dip variation in slip over the rupture surface combined with the geometric conﬁguration of the poroelastic oceanic crust near the down-dip limit of the rupture being “sandwiched” between relatively impermeable mantle of the underlying slab and overlying mantle wedge. Kopp and Kukowski. Poroelastic stress-triggering due to the SAE. 1998. Additionally. The coseismic distributions of deformation. 1998). Pesicek et al.4 − 0. Following the SAE.1 MPa in the subducting oceanic crust near the seismic barrier (Table 2 and Fig.. and pore pressure are initial conditions for the postseismic poroelastic deformation model. 2006b). At the conclusion of the model (two days before the NE). ﬂuids migrate from regions of high pore pressure to regions of low pore pressure.. Audet et al.7 Min − 1. Location Time step Near seismic barrier (MPa) Max Forearc (4 °N/94 °E) Forearc (4 °N/94 °E) Oceanic crust (2 °N/97 °E) Oceanic crust (2 °N/97 °E) Beginning End Beginning End 1. Various bulk permeabilities were examined for the subducting oceanic crust.4 MPa in the deep forearc and from − 0. Kopp et al.7 MPa for corresponding time steps. That is.4 to 1. (2009) determined the interface of the seismogenic zone between the subducting oceanic crust and overlying plate to have a permeability of 5 × 10− 25 to 5 × 10− 22 m2 which is much lower than the permeability of the actual oceanic crust.7 MPa for the coseismic (initial conditions) and 90-day time step. 1998.7 MPa in the subducting oceanic crust near the seismic barrier (Fig. Since the speciﬁed permeability and pore ﬂuid boundary conditions are constant for all poroelastic materials in the FEM. Furthermore. We ran multiple analyses of the FEM using varying permeability and rheology parameters and slip distributions. 6c).4 4. the pore pressure recovery in the subducting oceanic crust takes several months.7 − 0.4 − 1. Pollitz et al. A bulk permeability below 1 × 10− 17 m2 has not been shown to be geologically reasonable (Wang.1 0. . 2005. translates to systematic increases in Coulomb stress (Eq. We utilize an impermeable overlying plate on the time scale of the poroelastic model. Second. 10 kPa (Toda et al..H.0 MPa during the 3-month interval separating the SAE and NE (Fig. Discussion Although we focus on pore pressure changes near the NE hypocenter. The pore pressures in the shallow forearc and volcanic arc recover relatively quickly within about a month of the coseismic rupture (Fig.0 Min − 3. the coseismic slip introduces signiﬁcant pore pressure changes in the subducting oceanic crust south of the SAE rupture. This lateral migration of ﬂuids within the subducting oceanic crust occurs along the rupture zone of the SAE. the seismological observations of the Cascadian subduction zone indicate that the permeability contrast between the permeable oceanic crust and the overlying mantle wedge is more signiﬁcant (1 to 4 orders of magnitude) than the actual bulk permeability assigned (Audet et al. the geometric conﬁguration of rheologic properties and boundary conditions can introduce multiple “apparent” deformation time constants even if the parameters that control the time dependence (e. the pore pressure changes in the subducting oceanic crust north of Sumatra recover more quickly than the maxima and minima near the seismic barrier. and near the hypocentral location of the NE. 2009) due to the fact that the poroelastic effects occur within days to months. The predicted postseismic increase in pore pressure near the NE hypocenter. The uncertainties within these initial conditions rest on the assumptions of rheology. respectively.1 0.7 and − 0. permeability) are constant.5 − 2. as well as laterally (along-strike) within the subducting oceanic crust.5 to 0. and boundary conditions of the FEM. In this ﬂow regime. 5). ﬂuids migrate both up-dip and down-dip to the edge of the seismogenic zone. We found that the postseismic poroelastic results of the FEM were robust and did not change our conclusions. 2000. Absolute changes in pore pressure near the seismic barrier are − 2.K. the pore pressure near the hypocentral location of the NE increased by 2. That is. 2005) and decrease in pore pressure (Fig. pore pressures range from − 1. 2007. In contrast.L.
Red solid or dashed line represents location of the seismic barrier in the subducting oceanic crust. Top row of images includes forearc poroelastic deformation. In the bottom row of images the forearc has been stripped away to view the subducting oceanic crust.H. (b) Middle time segment (45 days after coseismic dynamic rupture). (c) Last time segment (two days before Nias earthquake). Hughes et al. 6. Pore pressure recovery following the SAE. / Earth and Planetary Science Letters 293 (2010) 289–299 Fig. . (a) First time segment (coseismic poroelastic deformation).296 K.L. Teal star represents NE hypocentral location.
Therefore. The horizontal deformation is maximum (a few tens-ofcentimeters) where the coseismic nodal plane intersects the land surface.. interseismic strain accumulation continues during all stages of the earthquake cycle.H. The stress released by the coseismic fault-slip propagates into the region surrounding the fault. However. this region of predicted maximum horizontal poroelastic deformation is offshore and no GPS sites are available to verify the predictions. 2005). with respect to the coseismic deformation ﬁeld. (b) Viscoelastic deformation. which has many more adjustable parameters). the character of poroelastic deformation may be contrary to some displacement observations (e. Colors and arrows represent vertical and horizontal deformation. Paul et al. The predicted increase in pore pressure (2. . In fact. and poroelastic relaxation. Alternatively.. substantial poroelastic uplift is predicted for the Nicobar Islands. particularly for the vertical deformation of the Nicobar Islands and the islands west of northern Sumatra (Fig.L. the poroelastic contribution to postseismic Coulomb stress changes near the NE hypocenter is much greater than that of the corresponding viscoelastic contribution. If we then rule out poroelastic deformation altogether in favor of some other mechanism (say afterslip. The response of the near-ﬁeld region to this stress depends on the rheologic partitioning. This addresses the bias inherent to interpretations that are based on a single mechanism and neglect the other contributions. contributions from all three postseismic deformation mechanisms are expected to be signiﬁcant. Hughes et al. poroelastic effects may correspond to increasing changes in Coulomb stress due to slow re-equilibration and lateral migration of pore ﬂuids within the subducting oceanic crust. 7). (a) Poroelastic deformation. except during the relatively instantaneous coseismic slip. Comparatively speaking. A similar dominance of viscoelastic (compared to poroelastic) postseismic deformation was predicted for other subduction zone earthquakes (Masterlark et al. (1)) that are two orders of magnitude greater than corresponding Coulomb stress increases predicted for viscoelastic models of postseismic deformation (Pollitz et al. 2. respectively. (2006) to account for this same deformation does not require any viscoelastic relaxation. Additionally. Furthermore. the predicted poroelastic deformation is expected to be a signiﬁcant contributor to the observed GPS measurements.. / Earth and Planetary Science Letters 293 (2010) 289–299 297 the south of the SAE rupture near the NE hypocenter. 7. Unfortunately. 2006a).0 MPa) for the NE hypocentral region directly correlates to positive changes in Coulomb stress (Eq. 2007. There are three mechanisms (all of which have been demonstrated in laboratory and ﬁeld measurements) that contribute to postseismic deformation — afterslip.. the direction of displacements differs for the different deformation mechanisms. due to the distribution of coseismic slip and geometric conﬁguration of rheologic properties. Lateral migration of ﬂuids demonstrated here has also been proposed to account for migration of slow slip events in subduction zones elsewhere (Melbourne et al. ﬂuids are Fig. we are assuming that either ﬂuids are not present in the crust or that poroelastic behavior is insigniﬁcant. the poroelastic deformation will most certainly not explain a large portion of the observed postseismic deformation. the afterslip distribution proposed by Hashimoto et al. Compared to afterslip and viscoelastic relaxation. Considering the sheer size of the SAE. Arrows are not plotted for predicted horizontal displacements less than 1 cm. Thus... the timing of postseismic poroelastic relaxation near the NE hypocenter is consistent with the three-month interval separating the SAE and NE. Magnitudes of viscoelastic deformation are about 5 times greater than those of poroelastic relaxation. We propose treating the predicted poroelastic deformation as a correction to postseismic deformation data. Such studies based on a single postseismic deformation mechanism introduce unknown bias and suppress the reliability of interpretations and predictions. as is customary for predicted interseismic strain accumulation. Poroelastic deformation is not considered because the other two mechanisms dominate the measured near-ﬁeld postseismic deformation. This suggests that postseismic poroelastic effects are important. The slip model of Masterlark and Hughes (2008) drives the FEM shown in Fig. Nonetheless. the expected magnitude of viscoelastic deformation is expected to be ﬁve times greater than that of the poroelastic deformation for the SAE (Fig. 7). However. However.K. In contrast to the dominance of viscoelastic postseismic deformation. For example. particularly for earthquake stress-triggering analyses. Viscoelastic relaxation is calculated for a period of 10 years following the SAE (μ = 1018 Pa s).g. the rheologic (viscosity) structure proposed by Pollitz et al. Paul et al. viscoelastic relaxation. the predicted coseismic distribution of pore pressure for the SAE and timing of postseismic poroelastic relaxation produce Coulomb stress changes of sufﬁcient magnitude to account for both the spatial and temporal proximity of the SAE and NE. 2001). Predicted poroelastic and viscoelastic deformation of the SAE. Measurable poroelastic deformation is primarily limited to the surface projection of the rupture. (2006a) to account for postseismic deformation of the SAE does not require any afterslip. Previous studies calibrate the afterslip and viscoelastic relaxation parameters to near-ﬁeld GPS measurements (Chlieh et al. 2007). 2007). Postseismic poroelastic deformation is complete several months after the SAE..
2002). the pore pressure southeast of the seismic barrier and near the hypocenter of the NE slowly.... R. Geophys. 107.. V. Corrections to “Pore pressure and poroelasticity effects in Coulomb stress analysis of earthquake interactions. Y. Ear Plan Space 58. Bull.M. Watts. Rev.. J. C. P. Sánchez. Moore. 637–658.. and references therein) and their presence and signiﬁcance in an accretionary prism are clear (e.. Geophysical Journal International 146. 2006. K. 2000.. K. V. Sieh.. D.C. Fisher. E. J. Bull.T. K.. 2007. 76–78. M. Simpson. isotropic..J. Memoir no.. Moore et al. J. (1)) (e. but systematically. S. Soc. Masterlark and Wang. 2007. Wald.H. Res. 76–82. 112.15 Sumatra–Andaman earthquake of 2004.. M. Ji.M. Res. even if other postseismic deformation mechanisms dominate the deformation signal. Masterlark and Wang.. 2006.. T. Academic licensing and technical support for Abaqus software is provided by Simulia Inc.W.. poroelastic contributions are signiﬁcant for explaining the delayed triggering of the NE due to the SAE.. S192–S207. J. Grilli. J. Hsu. 2002. 863–889. 108. Hutton. Ioualalen. Masterlark. Gahalaut et al... (Eds. 2005). Soc.3 Landers earthquakes) (Bosl and Nur. 97. triggered. Villaseñor... S. 36. Science 312. stress.0 Colima-Jalisco earthquake. Gary Hooks Endowed Geology Fund. S. H. B.. Soc.L. Res.. Coseismic slip and afterslip of the great Mw 9.. Polet. 97. 97. Let. Resources and Tectonic Evolution.. Grilli. Stock. M... which account for the distribution of material properties of the SASZ. J.. R. 2007). M... L. Bull.. Nur and Walder. Gahalaut. J. Malavieille. Geo. The Co-seismic slip distribution of the Landers earthquake. This suggests that transient pore pressure contributes signiﬁcantly to the spatial and temporal proximity of these two events. The poroelastic structure of the SASZ produces two ﬂow regimes having two separate time constants.. Seismol. Tsunami source of the 2004 Sumatra–Andaman earthquake inferred from tide gauge and satellite data.. J. King. Geology. 2001 M7.g. H. 2008. Fujimori. 2003.-A. Seismol. J. Geo.H. 127–139. Lay. T.T.H.. Bull. C.298 K. from GPS geodetic constraints. Galetzka. 97. We estimate the slip distribution for the SAE from near-ﬁeld GPS data using linear inverse methods and FEMgenerated Green's Functions. A.. Sumatra. C. W.K.. Hughes. J.R.-P. Bull... Sumatra. C.. Hughes et al. N. A complete explanation of stress-triggering initiated by the SAE must include poroelastic effects. D. 97. J. In particular. 85. Bostock. Masterlark. Milsom. Let. 2007.. 2008. and 1992 M7. W. Pore ﬂuid pressure..-J.. D. Ioualalen. Rice. J. London. N. Aftershocks and pore ﬂuid diffusion following the 1992 Landers earthquake..H. 2007.F. G. Choosakul. M.H.. Soc. Bull. S. EPSL 242... 1133–1139. present in the crust (e.). 143–182.g.I. D..R. C. Natawidjaja.. K. Helmberger. N.. Seismol. 2007. 2001. J.. H. Beeler.. H. Geoph. and pore pressure due to megathrust earthquakes in subduction zones having complex geometric conﬁgurations of rheologic properties. 935–953.. K. Ichinose.R. M. Time Series Analysis and Inverse Theory for Geophysicists.. M. Masterlark. N. Borchers.-Y. if pore pressure is an important contributor to analyses of stress-triggering... Seismol.. B.S. F.g. Tectonics 22. H. Kayal.J.. Pore pressure. 116. Audet. G. S43–S61..E. V. P. Parameter Estimation and Inverse Problems.. M. Furthermore. J.. 1998. 2005. 1978.L.C. Sci. Poroelastic coupling between the 1992 Landers and Big Bear earthquakes. Rupture process of the 2004 Sumatra–Andaman Earthquake. O. Ji.. 1994. Flueh. J. C.A.. Byerlee.R. Boll. decays rapidly (∼ 1 month) in the shallow forearc and volcanic arc of the overriding plate.. An updated digital model of plate boundaries. M. Fujii. Kieckhefer. A. Prawirodirdjo. Wang. D.. 2005. S. Gahalaut. 5..I. App. Poisson-solid. 1999. 181–186... Sladen. Crustal deformations associated with the great Sumatra–Andaman earthquake deduced from continuous GPS observation.. V. K... Geo. Hebert. 108. Permeability of the oceanic crust based on experimental studies of basalt permeability at elevated pressures. Lallemand. Suárez. Seismol. Y..R. Kopp. Das. Christensen. Wang. Nagarajan. 1976) and changes in pore pressure in this study and previous studies have been shown to contribute to changes in Coulomb stress at the same magnitude as normal and shear stresses (Eq. D. T. Pore pressure and poroelasticity effects in Coulomb stress analysis of earthquake interactions. Watts. New York. Bosl. Kukowski. J. K... N. Bilek. Soc.. 27. or aftershock? Cur. Res. 1980.. 2007. Shor Jr.. to the geometric conﬁguration of the poroelastic oceanic crust being “sandwiched” between relatively impermeable mantle of the underlying slab and overlying mantle wedge (Audet et al.. recovers during the three-month interval separating the SAE and NE. 35. J. It is well-known that a transient pore pressure pulse can trigger transient seismicity (Raleigh et al. A. 2004 Indian Ocean tsunami....K. Frictional afterslip following the 2005 Nias– Simeulue earthquake. Kirby. Cambridge University Press. S. Using earthquake source durations along the Sumatra–Andaman subduction system to examine fault-zone variations. H. The timing for pore pressure recovery is more sluggish (several months) in the oceanic crust of the down-going slab due. Conclusions We present a quantitative analysis of poroelastic deformation of the SAE and stress-triggering of the NE. Finite element model predictions of static deformation from dislocation sources in a subduction zone: sensitivities to homogeneous. EPSL 168. Am. Shi. A. 105. Y. 2003. J.D.. pp. S. 92. 414–428. 2009.K. J. Crow. S. J.. J. Seismol. Cocco. Kaewbanjak. 107. F. Kanamori. H. Res.. 84. 2003. NSF Geophysics award EAR-0911466..R. 1921–1926. in part.. Pollitz.. Avouac. A.. 2002. Tectonophysics 460. Friction of rocks. Asavanant. M..L. Gutscher. Mexico earthquake. 49. Static stress changes and the triggering of earthquakes.H. Thurber. 31. Geo. Sieh.. The FEM-based techniques presented here allow for simulating the evolution of coseismic and postseismic deformation. Galetzka... R. L. J. J. 25533–25542...G. DeShon. 2009). as demonstrated here. Nature 457. Lockner. Am. Source constraints and model simulation of the December 26. Stixrude for astute reviews. Barber. Bock. Res. T.K. N. Am.. We thank Fred F. Hjorleifsdottir. S86–S102.J. Acknowledgements This work is supported in part by NASA under award NNX060F10G. V. 646–659. Y. Tectonophysics 149. / Earth and Planetary Science Letters 293 (2010) 289–299 Banerjee. T. G-Cubed 4. Bock. Res.. J. Nur. Geo. Columbia University Press.. Hickman. J. T. Engdahl. Hyndman. 365–374.. Mexico. Gubbins. S.... Masterlark. N. and Coulomb failure.. Geo.J. then poroelastic deformation should not be neglected from postseismic deformation analyses. P. Geo. Soc. Lin. Aster. 1994. 615–626. 89. 2003. Res. 2004. R. C. Slip distribution for the 2004 Sumatra–Andaman earthquake constrained by both GPS data and tsunami run-up measurements... Bialas.-P.. J.C. Fukuda. and thus poroelastic deformation. Geo. Hashizime. R. P. Cambridge. P.. Science 308.L. Nagarajan.. Rice. Am. . Soc. J. Avouac. Constraints on 2004 Sumatra–Andaman earthquake rupture from GPS measurements in Andaman– Nicobar Islands. Bull. Masterlark.. 2000.. The seismogenic zone of subduction thrust faults: what we know and what we don't know. R. J. Am. Christensen. Kirby. Stein.. Thio.. Gahalaut. Chlieh. Catherine. Permeability within basaltic oceanic crust. Ramananantoandro. Pollitz for insightful comments and Thorne Lay and Lars P.. 2002. Kopp. Waterway Port Coastal Ocean Eng. B. Soc. 2007. Hjorleifsdottir.. Bird. Geo. Transient stress-coupling between the 1992 Landers and 1999 Hector Mine earthquakes. Segall. Ni. Peacock.G. 133.... 1990. 2007. J.. 28 March 2005 Sumatra earthquake: expected. Bull. 3647–3650.T. J. H... T. E. 15–40. Backstop geometry and accretionary mechanics of the Sunda margin.. Next generation of deformation models for the 2004 M9 Sumatra–Andaman earthquake. Teor.. Res... Chlieh. T. 2008. D. Kumar. Am. and half-space assumptions. Freymueller. Reichert. Y.F.R. R. 2002. G... A. Masterlark. Elsevier Academic Press. India..K. 2002.. Klaeschen. Crustal structure of the Java margin from seismic wide-angle and multichannel reﬂection data. This relatively rapid recovery in pore pressure may help to explain the timing and location of near-ﬁeld aftershock swarms (Fig. P. Simons.. Cocco. Song. Takemoto. Slip kinematics and dynamics during and after the 1995 October 9 Mw = 8. J. 2008. S.M.H. J. King. San Diego. 2000..T. Geo. S152–S173. Seismol. 2005.. References Ammon. In: Dixon. M. Hughes.... 1) (Piombo et al. DeMets... Dassault Systèmes. Satake. Thurber.. S62–S70. Prawirodirdjo.. 1470–1486. 1988. 84. 107.-R.S. 255–270.. Res.P. Am.. Coseismic slip distributions of the 26 December 2004 Sumatra–Andaman and 28 March 2005 Nias earthquakes from GPS static offsets. Seismol. Collet.A. Am. Kalpna. The estimated slip distribution then drives a forward model that simulates poroelastic processes induced by the SAE.. apparent friction. 1995 M8 Jalisco. M. and the W.. H. Bürgmann. 452–454. Hashimoto. Curray.. By extension. 2005. Poroelastic relaxation and aftershocks of the 2001 Bhuj earthquake. Robinson. Takiguchi. The Seismogenic Zone of Subduction Thrust Faults. G. Seismic evidence for overpressured subducted oceanic crust and megathrust fault sealing. 2003. Tectonic segmentation of the Northern Andean margin: impact of the Carnegie Ridge collision.. Seismic refraction studies of the Sunda Trench and forearc basin. Asavanant. Geo. P.. Modeling the 26th December 2004 Indian Ocean tsunami: case study of impact in Thailand.6 Bhuj earthquake. Geological Society. 2006. Gahalaut.
S. Academic Press. Hashimoto. C. D. C... Colorado. Miller. W.. vol. Plate-boundary deformation associated with the great Sumatra–Andaman earthquake. G.A. / Earth and Planetary Science Letters 293 (2010) 289–299 Masterlark... Schubert. Melosh. P..W.. 507–515. Meltzner.. Thesis.. Lacassin.F. Nature 402. Moore.J. Dieterich... W. Complex slab subduction beneath northern Sumatra. J.... 1982. 2008. Homogeneous vs heterogeneous subduction zone models: Coseismic and postseismic deformation.. A. 34. Anderson. Raefsky. Int..W. Princeton. Taira. Slip of the 2004 Sumatra– Andaman earthquake from joint inversion of long-period global seismic waveforms and GPS static offsets.R. 24543–24565... L. Res. Agnew. M. Stein. Geo. 113–127. S115–S127. Bilham.M. 291. DeShon. O.. C.). Bull.C. Res.C. S. Vigny.). G. McCaffrey. Res. Walder. J. Thurber... J. 1018–1040. Socquet.. Time-dependent strain accumulation and release at island arcs: Implications for the 1946 Nankido earthquake. Jr. K.. Pollitz. Geo. 1230–1237. (Eds..D... Science 191.I. R.. Avouac. 167. Choosakul. Nature 434. Lowry.J. D. R.B. Banerjee. shock: effect on aftershocks and future earthquake probabilities. In: National Research Council (Ed. N.3) and March 28. Chlieh.S. 33. 97. 397–420. J.H. Geo. Am. King... A.. 111... C.J. M. Geophysical Data Analysis: Discrete Inverse Theory.C.. Abu. 2005 (Mw = 8. 1974.. J. In: Bebout. The Role of Fluids in Crustal Processes. Bamphenyu. A. D. R. American Geophysical Union.. Stock.. K. MIT.. Bowman. Banerjee. Nalbant. Hudnut. 2006. H. 1976. Reasenberg. 1960.. R... Nature 440. 71. B. R. Washington... 299 Pollitz. pp. A.F.. Theory of Linear Poroelasticity: With Applications to Geomechanics. Bangs. G.A.F. 2006.. Healy. Wang. 2004 Great Sumatra–Andaman earthquake. Insight into the 2004 Sumatra–Andaman earthquake from GPS measurements in Southeast Asia. 32. Sieh...P. H. Res. Dmowska. N.. A.. J. G. Am. Burgmann..0–9.-P.. McCloskey..L. 46–51.F.9 Kobe. Piombo. Melbourne. Geo. Szeliga.. Bredehoeft. Inc. Seismol. Soc.L. Ph. T. M... Geo.S. N. 2000. Mechanism of the Chilean earthquakes of May 21 and 22. 1998. A. D. Martinelli. Widiyantoro. Internal deformation due to shear and tensile faults in a half-space. Sen.. Smalley Jr. Let. J.. D.. 1992. 2007. International Geophysical Series. Mignan.H. Abrams. Okada.7) earthquakes. G.. Let. 2007.F. H. Washington.J. E. 1970.. Rhie. Okal. Extent and duration of the 2003 Cascadia slow earthquake. Pangborn. M.H. H.S. S. G. Y. D. J. Introduction to Groundwater Modeling: Finite Difference and Finite Element Methods. R. Savage..H. Santillan... A simple and efﬁcient method for introducing faults into ﬁnite element computations. . Time-dependent hydraulics of the Earth's crust. H.. Meltzner. 1128–1131.. J. 1996. Smith. Tobin. R.. Japan. Avouac. Steacy. H.. 81. EPSL 244... 103.. Subarya. Cambridge University Press.. 45.F. S. W. 2005. Post-seismic ﬂuid ﬂow and Coulomb stress changes in a poroelastic medium. 162..J.D.L. B.. L. Bürgmann. Academic Press. Plafker.. Geodynamics: Applications of Continuum Physics to Geological Problems. Subarya. 4047–4050. C..-P.A. Satirapod.. Geo.E. Ruff. 639–654. National Academy Press. Natawidjaja.. Paul.K.. 1391–1400. An experiment in earthquake control at Rangely.. A. D... Geophysical Research Letters 28.. J. Natawidjaja. P. Indonesian earthquake: earthquake risk from co-seismic stress. Bull. 91–104.. R. S. Turcotte. J. Bürgmann.. 1990. E. Postseismic deformation of the Andaman Islands following the 26 December. F. Engdahl.. Nur. DeMets. 2005. T. Raleigh. R. Scholl.F. J.C. A. 2002. A. Dreger. S. Nature 434.. Pesicek. F.D.A. T. 605–609. Princeton University Press. V. M. Prawirodirdjo.H. C. Stress changes along the Sunda trench following the 26 December 2004 Sumatra– Andaman and 28 March 2005 Nias earthquakes. S.. Abidin.. 201–206. P.. Let. Seismic activity in the Sumatra–Java region prior to the December 26.. Ambrosius. 581–582.J. Bock. Sánchez... Res.. Int. 2005. 2006a.P. 2007. J...H. New York... Menke.C. M. Subduction: Top to Bottom. Soc. Geo.M.J. S.. Dragoni. E. D. R. Res. 1999. S. J. Kirby.. Stein.. Yoshida. Wang.. K. Sieh. 1001–1030. J. 2006b.T. C. Seismol. Post-seismic relaxation following the great 2004 Sumatra–Andaman earthquake on a compressible self-gravitating Earth. Wang.R.. Bulletin of the Seismological Society of America 82.. San Diego.H. Omar. Platt. H. 2005. Bull. 1981.R. Speed and size of the Sumatra earthquake.. second ed. 2004 (Mw = 9. Hughes et al.. Large earthquakes in subduction zones: segment interaction and recurrence times. J. pp. D.. 2001.M. J. Three dimensional splay fault geometry and implications for tsunami generation. Let. Am. Simons. 2006. 35. Toda... K. Uplift and subsidence associated with the great Aceh– Andaman earthquake of 2004. Geo. R. 1989. Stein. Stress transferred by the 1995 Mw = 6. Choosakul. 2005. Romanowicz. A. Geo. Science 318. Soc. The role of stress transfer in earthquake occurrence. A. Nature 436. Kuramoto..

References: V. 
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 V. 
 V. 
 V. 
 V.