Seismotectonics of a diffuse plate boundary: Observations off the Sumatra-Andaman trench K. Aderhold 1,2 and R. E. Abercrombie 1 1 Department of Earth and Environment, Boston University, Boston, Massachusetts, USA, 2 Now at Incorporated Research Institutions for Seismology, Washington, District of Columbia, USA Abstract The actively deforming Indo-Australian intraplate region off the Sumatra-Andaman trench hosted the largest strike-slip earthquake recorded by modern instruments, the 2012 M w 8.6 Wharton Basin earthquake, closely followed by a M w 8.2 aftershock. These two large events ruptured either parallel north-south trending faults or a series of north-south and nearly perpendicular east-west fault planes. No active east-west faults had been identified in the region prior to these earthquakes, and the seismic rupture for these two earthquakes extended past the 800°C isotherm for lithosphere of this age, deep into the oceanic mantle and possibly beyond the inferred transition to ductile failure. To investigate the seismic behavior of this region, we calculate moment tensors with teleseismic body waves for 6.0 ≤ M w ≤ 8.0 intraplate strike-slip earthquakes. The centroid depths are located throughout the seismogenic mantle and could extend through the oceanic crust, but are generally well constrained by the 600°C isotherm and do not appear to rupture beyond the 800°C isotherm. We conclude that while many earthquakes are consistent with a thermal limit to depth, large magnitude earthquakes may be able to rupture typically aseismic zones. We also perform finite-fault modeling for M w ≥ 7.0 earthquakes and find a slight preference for rupture on east-west oriented faults for the 2012 M w 7.2 and 2005 M w 7.2 earthquakes. This lends support for the presence of active east-west faults in this region, consistent with the majority of previously published models of the 2012 M8+ earthquakes. 1. Introduction The Indo-Australian plate is the most actively deforming intraplate oceanic lithosphere in the world and has prompted many studies, but it also remains one of the more perplexing tectonic regions. The uncertainty of our knowledge of this region was highlighted recently by the intraplate 11 April 2012 M w 8.6 earthquake and M w 8.2 aftershock 2 h later. They were unexpectedly large and deep [McGuire and Beroza, 2012], and most models show that they ruptured along previously unknown faults. These two earthquakes were observed to rupture along a complex arrangement of north-south (N-S) faults [Satriano et al., 2012] or along additional unidentified east-west (E-W) faults at lithospheric depths beyond the inferred limit to brittle failure [Wang et al., 2012; Duputel et al., 2012; Meng et al., 2012; Zhang et al., 2012; Ishii et al., 2013; Wei et al., 2013]. Since these events were so complex, we use moderate-sized earthquakes to understand the seismic behavior in this region, both in depth extent of seismic rupture and orientation of active faulting. The Indo-Australian plate appears to be breaking apart as a result of the collision of India against Asia in the north, along with ongoing subduction along the Sunda arc [Wiens et al., 1985]. This subduction caused an increase of seismic activity in the region following the 2004 M w 9.2 Great Sumatra-Andaman earthquake [Engdahl et al., 2007; Delescluse et al., 2012]. Wiseman and Bürgmann [2012] proposed that viscoelastic relaxa- tion of the asthenosphere in the wake of the 2004 and 2005 subduction earthquakes may have promoted failure on the conjugate fault planes of the 11 April 2012 earthquakes. Royer and Gordon [1997] divided this broad region of intraplate activity into three subplates: the Capricorn plate to the southwest, the Australian plate to the southeast, and the Indian plate to the north, all separated by wide, diffuse boundaries of deformation. The north-south trending Ninety East Ridge [Petroy and Wiens, 1989], an inactive hot spot track [Weis et al., 1993], separates the Indian Ocean into two regions with distinct deformation styles: the Central Indian Ocean to the west and the Wharton Basin to the east (Figure 1) [Cloetingh and Wortel, 1986; Sagar et al., 2013]. In the Central Indian Ocean the compression is accommodated by reactivating E-W striking normal faults with reverse motion [Chamot-Rooke et al., 1993; Montési and Zuber, 2003]; the pressure is also accommodated through long-wavelength folding [Krishna et al., 2001]. To the east of the Ninety East ADERHOLD AND ABERCROMBIE SEISMOTECTONICS OF A DIFFUSE BOUNDARY 3462 PUBLICATION S Journal of Geophysical Research: Solid Earth RESEARCH ARTICLE 10.1002/2015JB012721 Key Points: • Seismic slip extends from maximum expected depth to the crust in 45–75 Ma oceanic lithosphere • Large (M w > 7) strike-slip earthquakes can rupture multiple differently oriented faults • Rupture directivity modeling provides tentative support for active east-west oriented faults Supporting Information: • Supporting Information S1 Correspondence to: K. Aderhold, [email protected] Citation: Aderhold, K., and R. E. Abercrombie (2016), Seismotectonics of a diffuse plate boundary: Observations off the Sumatra-Andaman trench, J. Geophys. Res. Solid Earth, 121, 3462–3478, doi:10.1002/2015JB012721. Received 7 DEC 2015 Accepted 2 APR 2016 Accepted article online 7 APR 2016 Published online 7 MAY 2016 ©2016. American Geophysical Union. All Rights Reserved.
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Seismotectonics of a diffuse plate boundary: Observationsoff the Sumatra-Andaman trenchK. Aderhold1,2 and R. E. Abercrombie1
1Department of Earth and Environment, Boston University, Boston, Massachusetts, USA, 2Now at Incorporated ResearchInstitutions for Seismology, Washington, District of Columbia, USA
Abstract The actively deforming Indo-Australian intraplate region off the Sumatra-Andaman trench hostedthe largest strike-slip earthquake recorded bymodern instruments, the 2012Mw 8.6Wharton Basin earthquake,closely followed by a Mw 8.2 aftershock. These two large events ruptured either parallel north-south trendingfaults or a series of north-south and nearly perpendicular east-west fault planes. No active east-west faultshad been identified in the region prior to these earthquakes, and the seismic rupture for these two earthquakesextended past the 800°C isotherm for lithosphere of this age, deep into the oceanic mantle and possiblybeyond the inferred transition to ductile failure. To investigate the seismic behavior of this region, we calculatemoment tensors with teleseismic body waves for 6.0≤Mw≤ 8.0 intraplate strike-slip earthquakes. Thecentroid depths are located throughout the seismogenic mantle and could extend through the oceanic crust,but are generally well constrained by the 600°C isotherm and do not appear to rupture beyond the 800°Cisotherm. We conclude that while many earthquakes are consistent with a thermal limit to depth, largemagnitude earthquakes may be able to rupture typically aseismic zones. We also perform finite-fault modelingfor Mw≥ 7.0 earthquakes and find a slight preference for rupture on east-west oriented faults for the 2012Mw 7.2 and 2005Mw 7.2 earthquakes. This lends support for the presence of active east-west faults in this region,consistent with the majority of previously published models of the 2012 M8+ earthquakes.
1. Introduction
The Indo-Australian plate is the most actively deforming intraplate oceanic lithosphere in the world and hasprompted many studies, but it also remains one of the more perplexing tectonic regions. The uncertainty ofour knowledge of this region was highlighted recently by the intraplate 11 April 2012Mw 8.6 earthquake andMw 8.2 aftershock 2 h later. They were unexpectedly large and deep [McGuire and Beroza, 2012], and mostmodels show that they ruptured along previously unknown faults. These two earthquakes were observedto rupture along a complex arrangement of north-south (N-S) faults [Satriano et al., 2012] or along additionalunidentified east-west (E-W) faults at lithospheric depths beyond the inferred limit to brittle failure [Wanget al., 2012; Duputel et al., 2012; Meng et al., 2012; Zhang et al., 2012; Ishii et al., 2013; Wei et al., 2013]. Sincethese events were so complex, we use moderate-sized earthquakes to understand the seismic behavior inthis region, both in depth extent of seismic rupture and orientation of active faulting.
The Indo-Australian plate appears to be breaking apart as a result of the collision of India against Asia in thenorth, along with ongoing subduction along the Sunda arc [Wiens et al., 1985]. This subduction caused anincrease of seismic activity in the region following the 2004 Mw 9.2 Great Sumatra-Andaman earthquake[Engdahl et al., 2007; Delescluse et al., 2012].Wiseman and Bürgmann [2012] proposed that viscoelastic relaxa-tion of the asthenosphere in the wake of the 2004 and 2005 subduction earthquakes may have promotedfailure on the conjugate fault planes of the 11 April 2012 earthquakes.
Royer and Gordon [1997] divided this broad region of intraplate activity into three subplates: the Capricorn plateto the southwest, the Australian plate to the southeast, and the Indian plate to the north, all separated by wide,diffuse boundaries of deformation. The north-south trending Ninety East Ridge [Petroy and Wiens, 1989], aninactive hot spot track [Weis et al., 1993], separates the Indian Ocean into two regions with distinct deformationstyles: the Central Indian Ocean to the west and the Wharton Basin to the east (Figure 1) [Cloetingh and Wortel,1986; Sagar et al., 2013]. In the Central Indian Ocean the compression is accommodated by reactivating E-Wstriking normal faults with reverse motion [Chamot-Rooke et al., 1993; Montési and Zuber, 2003]; the pressureis also accommodated through long-wavelength folding [Krishna et al., 2001]. To the east of the Ninety East
ADERHOLD AND ABERCROMBIE SEISMOTECTONICS OF A DIFFUSE BOUNDARY 3462
PUBLICATIONSJournal of Geophysical Research: Solid Earth
RESEARCH ARTICLE10.1002/2015JB012721
Key Points:• Seismic slip extends from maximumexpected depth to the crust in45–75 Ma oceanic lithosphere
• Large (Mw> 7) strike-slip earthquakescan rupture multiple differentlyoriented faults
• Rupture directivity modeling providestentative support for active east-westoriented faults
Supporting Information:• Supporting Information S1
Correspondence to:K. Aderhold,[email protected]
Citation:Aderhold, K., and R. E. Abercrombie(2016), Seismotectonics of a diffuseplate boundary: Observations off theSumatra-Andaman trench, J. Geophys.Res. Solid Earth, 121, 3462–3478,doi:10.1002/2015JB012721.
Received 7 DEC 2015Accepted 2 APR 2016Accepted article online 7 APR 2016Published online 7 MAY 2016
©2016. American Geophysical Union.All Rights Reserved.
Ridge, in theWharton Basin, the earthquakes are predominantly strike slip [Deplus et al., 1998; Abercrombie et al.,2003; Rajendran et al., 2011]. The strain field is optimally oriented to reactivate the ~15°N-S trending fossil frac-ture zones formed during spreading at the extinct Wharton Ridge [Delescluse and Chamot-Rooke, 2007; Ishiiet al., 2013; Andrade and Rajendran, 2014] and is assumed to host the numerous strike-slip earthquakesobserved here (Figure 1). These fractures zones extend thousands of kilometers and have been imaged to becontinuous under the obscuring sediments of the accretionary prism and Bengal Fan [Deplus et al., 1998;Matthews et al., 2011; Franke et al., 2008]. They offset lithosphere with ages from 40 to 85Ma in the southprogressing to over 100Ma to the north [Müller et al., 2008; Matthews et al., 2011].
Abercrombie et al. [2003] showed that the north-south fracture zones in theWharton Basin remain active as theyare subducted past the trench, and seismicity in the slab on the Investigator Fracture Zone was observed to adepth of 200 km [Fauzi et al., 1996; Lange et al., 2010]. These events can be damaging as in the case of the 2009earthquake beneath Padang, Sumatra [Wiseman et al., 2012]. It is not known how the subducting fracture zonesaffect the geometry of the subduction zone earthquakes themselves, such as the great 2004Mw 9.2 earthquakeand the 2005 Mw 8.6 Nias earthquake. Fracture zones provide a narrow zone of structural and compositionaldifferences in subducting oceanic lithosphere and can enhance hydration of the interface and overriding plate[Shillington et al., 2015]. The relief of fracture zones has been proposed to behave as a rupture asperity, with
Figure 1. (a) Bathymetry of the Central Indian Ocean and Wharton Basin. Background is GEBCO bathymetry. Solid black lines—plate boundaries [Bird, 2003], white dashed lines—fracture zones [Matthews et al., 2011], grey lines—1 km sedimentthickness contours [Whittaker et al., 2013], orange lines—2m slip contours of the 2004Mw 9.2 earthquake [Chlieh et al., 2007],pink lines—2m slip contours of the 2005 Mw 8.5 earthquake [Konca et al., 2007], turquoise lines—fault planes for finite-faultmodeling of the 2005 and 2012 earthquakes, black circles—centroid locations of earthquakes from Buchanan [1998], whitecircles—gCMT centroid locations of the earthquake modeled in this study, small grey circles—earthquakes from the NEICcatalog in 30 days following the 2012Mw 8.6 earthquake, small orange circles—earthquakes from the NEIC catalog in 2weeksfollowing the 2016Mw 7.8 earthquake, and grey and orange—moment tensors from gCMT catalog. (b) Lithospheric age of theCentral IndianOcean andWharton Basin. Background is age of lithosphere [Müller et al., 2008].White lines are same features inFigure 1a. Circles are same earthquake centroids as in Figure 1a. Double-couple moment tensors are from this study, and colordenotes the centroid depth.
Journal of Geophysical Research: Solid Earth 10.1002/2015JB012721
ADERHOLD AND ABERCROMBIE SEISMOTECTONICS OF A DIFFUSE BOUNDARY 3463
topographic heterogeneities affecting the coupling of the slab and the extent of subduction zone megathrustearthquakes [Müller and Landgrebe, 2012; Robinson et al., 2006; Scholz and Small, 1997]. Alternatively, the reliefof a subducted irregularity has been observed as creating an area of weak coupling on the subduction interface[Wang and Bilek, 2011; Sparkes et al., 2010; Mochizuki et al., 2008]. Regardless of how fracture zones are inter-preted, as either an inhibitor or an enhancer of rupture, numerous studies find them to influence subductionzones and their potentially damaging and tsunamigenic thrust earthquakes.
On 11 April 2012 aMw 8.6 earthquake, the largest strike-slip earthquake ever recorded instrumentally recorded,ruptured in the Wharton Basin, followed by a Mw 8.2 aftershock 2 h later. The complex faulting geometryand unexpected depths of these two earthquakes challenged our understanding of deformation in theregion and the depth controls on oceanic earthquakes globally [McGuire and Beroza, 2012]. The aftershockdistribution, combined with back projections of high-frequency energy from these two events, revealed acomplex faulting pattern. Satriano et al. [2012] determined acceptable model fits to their data using a faultgeometry of only parallel N-S trending strike-slip faults, but all other studies included significant slip onadditional perpendicular faults. These studies used sets of three or more orthogonal faults oriented bothN-S along existing faults and E-W along previously unidentified faults that extended from the 94°Efracture zone to the Ninety East Ridge [e.g., Wang et al., 2012; Duputel et al., 2012; Meng et al., 2012;Zhang et al., 2012; Ishii et al., 2013; Wei et al., 2013]. The complexity of these earthquakes makes it difficultfor the fault planes to be unambiguously determined using back-projection and finite-fault inversiontechniques. The aftershocks of the 2012 events as located by the National Earthquake InformationCenter (NEIC) outline an intricate pattern supporting the existence of E-W faults (Figure 1). However, manyearthquake catalog locations in this region have significant error of over 20 km due to limited and non-uniform coverage by local and regional seismic stations [Pesicek et al., 2010]. Recent investigation ofthe subsurface faulting in the region finds no evidence for active E-W structures extending below thesediments [Geersen et al., 2015]. Carton et al. [2014] and Qin and Singh [2015] imaged faults extendingdeep into the mantle, but the dips are too shallow for them to have accommodated significant slip inthe great earthquakes in 2012.
The depths of oceanic earthquakes are thought to be limited by thermal structure;McKenzie et al. [2005] andJackson et al. [2008] demonstrated how robust earthquake centroids are limited by the 600°C isotherm,consistent with laboratory studies [Boettcher et al., 2007]. Global Centroid Moment Tensor catalog (gCMT)centroid depths for the 2012 Mw> 8 events were unusually deep (46 and 55 km) implying slip up to andbeyond the 800°C isotherm [Ekström et al., 2012], necessitating novel modes of failure [McGuire and Beroza,2012] (Figure 2). The expected depth of the 600°C isotherm is 35–50 km, determined from the age of theoceanic lithosphere in the rupture zone. Duputel et al. [2012] performed a moment tensor analysis of the very
Figure 2. Depth of earthquakes versus lithospheric age. Labeled isotherms are shown in black for a half-space coolingmodeland corresponding 400°, 600°, and 800° isotherms in blue for a plate-cooling model [McKenzie et al., 2005]. Slip depth extentsforMw ≤ 8.0 events are based on the seismicmoment of the point sourcemodel assuming circular rupture and constant stressdrop of 5MPa. Depth sensitivity testing results of the 21 May 2014 Bay of Bengal earthquake are on the right.
Journal of Geophysical Research: Solid Earth 10.1002/2015JB012721
ADERHOLD AND ABERCROMBIE SEISMOTECTONICS OF A DIFFUSE BOUNDARY 3464
long period waves and found centroid depthsof 30 km and 30–40 km for the Mw 8.6 and 8.2earthquakes, respectively. Deep slip in the2012 earthquakes (40–60 km) is supported byboth finite-fault inversion [Wei et al., 2013]and GPS modeling [Hill et al., 2015]. Fasterstrain rates may increase the depth of thebrittle to ductile transition but only by a fewkilometers [e.g., Rolandone et al., 2004]. Toexplain the exceptional depths, McGuireand Beroza [2012] proposed a slip-velocitystrengthening mechanism dependent onstrain rate using a fine-grained, viscous shearzone that would temporarily increase thestrength of the fault at depth within the limitedtemperature range of 600–800°C [Kelemen andHirth, 2007]. This mechanism could only takeplace during large magnitude earthquakes.
Here we investigate seismic behavior in theWharton Basin by modeling the thirteen6.0 ≤Mw ≤ 7.5 strike-slip earthquakes that haveoccurred since 1990 (Table 1). These earth-quakes are large enough for teleseismic datato be adequate to resolve the source processesbut usually not too large for complexity to limitour resolution. We focus on addressing thequestions raised by the 2012 Mw 8.6 and 8.2earthquakes: the orientation of active faultsand the depth distribution of seismic slip.We calculate moment tensors and centroiddepths by modeling teleseismic body waves.Additionally, for the best-recorded events, weinvestigate the fault geometry and spatialdistribution of slip by finite-fault inversion.
2. Moment Tensors and Centroidsfor Mw≤7.0 Earthquakes
We model teleseismic body waves for the 10strike-slip earthquakes of 6.0 ≤Mw ≤ 7.0 thatoccurred in the oceanic lithosphere along andto the east of the Ninety East Ridge in theWharton Basin and in the Bay of Bengal since1990 (Figure 1). We consider the earthquakein the Bay of Bengal separately because ofthe significant differences in local velocitystructure. We follow the methods outlined inAderhold and Abercrombie [2015], which arebased on the work of Maggi et al. [2000].Seismograms of P and S waves recorded atstations between 30° and 90° distance fromthe NEIC hypocenter for each earthquake areobtained from the Incorporated ResearchInstitutions for Seismology Data ManagementTa
ble
1.Earthq
uakesMod
eled
inTh
isStud
ya
Year
(DOY)
Dateb,Tim
e(Hyp
ocen
ter)
Latitud
eNEIC
Long
itude
NEIC
Mw
NEIC
h 0 NEIC
Latitud
egC
MT
Long
itude
gCMT
M0
gCMT
h 0gC
MT
CLV
DgC
MT
h 0 (km)
Strik
e(º)
Dip
(º)
Rake (º)
Mom
ent
(Nm)
Distance/Azimuth
NEICHyp
ocen
ter
togC
MTCen
troid
2005
(205
)20
05/07/24
15:42:06
7.92
092
.190
7.2
16.0
7.92
91.88
8.88
e26
125.5%
2011
672
169
8.7e
1934
.2/270
º
2010
(163
)20
10/06/12
19:26:50
7.88
191
.936
7.5
35.0
7.85
91.65
1.95
e27
33.1
4.5%
18/41
127/11
975
/64
176/98
1.1e
20/8.3e1
931
.7/264
º
2012
(010
)20
12/01/10
18:36:59
2.43
393
.210
7.2
19.0
2.59
92.98
7.55
e26
23.7
3.1%
1210
388
188
7.7e
1931
.0/304
º
2014
(141
)20
14/05/21
16:21:54
18.201
88.038
6.0
47.2
18.10
88.09
1.67
e25
57.6
1.2%
5332
082
180
1.3e
1812
.5/154
º
2012
(106
)20
12/04/15
05:57:40
2.58
190
.269
6.2
252.49
90.31
3.1e
2533
17.1%
3812
6935
42.8e
1811
.1/156
º
2009
(314
)20
09/11/10
02:48:46
8.08
291
.899
623
8.05
91.86
1.18
e25
19.9
0.8%
2811
877
161
1.0e
185.6/23
0º
2007
(277
)20
07/10/04
12:40:31
2.54
392
.903
6.2
352.47
92.83
2.74
e25
128.4%
1410
584
184
2.6e
1811
.5/225
º
2006
(109
)20
06/04/19
20:36:46
2.64
393
.226
6.2
172.70
93.22
2.55
e25
17.2
14.2%
1628
889
179
2.5e
186.4/35
4º
2005
(001
)20
05/01/01
06:25:44
5.09
992
.304
6.7
11.7
4.97
92.22
1.2e
2612
4.9%
1410
889
176
1.2e
1917
.1/213
º
1999
(333
)19
99/11/29
03:46:30
�1.275
89.043
6.4
10�1
.25
88.98
4.5e
2515
0.0%
822
8414
4.7e
187.5/29
2º
1999
(319
)19
99/11/15
05:42:43
�1.339
88.976
710
�1.21
88.89
3.3e
2615
5.0%
3613
8735
63.8e
1917
.2/326
º
1995
(312
)19
95/11/08
07:14:18
1.83
395
.05
6.9
332.00
94.77
2.65
e26
29.6
20.6%
2018
968
344
2.6e
1936
.2/301
º
1990
(288
)19
90/10/15
01:35:44
�2.211
92.249
6.8
32.2
�2.20
92.29
1.35
e26
237.9%
3228
584
164
1.4e
194.7/75
º
a Mom
entten
sorsfrom
pointsou
rcemod
elingarelistedalon
gwith
thegC
MTcentroidforthe
correspo
ndingearthq
uake.h
0iscentroidde
pthan
dCLV
Disthecompe
nsated
linearv
ectord
ipole.
Publishe
dearthq
uake
source
parameterscanbe
foun
din
Tables
S2an
dS3
inthesupp
ortin
ginform
ation.
bDates
areform
attedas
year/m
onth/day.
Journal of Geophysical Research: Solid Earth 10.1002/2015JB012721
ADERHOLD AND ABERCROMBIE SEISMOTECTONICS OF A DIFFUSE BOUNDARY 3465
Center (IRIS DMC). We pick phases first using the TauP arrival times [Crotwell et al., 1999]. Each arrival is thenhandpicked with careful attention paid to the polarity of the close to nodal P waves as well as the pP and sPdepth phases. Depth phases of oceanic strike-slip earthquakes recorded at teleseismic distances are oftenincorrectly picked as the first arriving P waves since the takeoff angles are much closer to the nodal planesrelative to dip-slip earthquakes of comparable magnitude [Schramm and Stein, 2009]. A poorly picked arrivalcan greatly impact the determination of the moment tensor and centroid depth of an earthquake.
The length of the record that we include in themodel is between 30 and 90 s depending on themagnitude ofthe earthquake and the presence of resonating late arrivals from structural complexity. Seismograms aredownsampled to one sample per second. We rotate records of the S waves on the two horizontal compo-nents to produce the transverse SH waves and downweight them in the models due to their large amplituderelative to the P waves.
We use the program MT5 [McCaffrey et al., 1991] to determine double-couple moment tensors andcentroid depths. All depths are reported with respect to the seafloor. We investigate the uncertaintyof our models by performing a grid search in 3 km increments around centroid depth and note howmechanism parameters of strike, dip, and rake are affected. Best fit point source models for allearthquakes are available in Figures S4–S16 in the supporting information. We also investigate whethersignificant rupture directivity can identify which nodal plane hosted the earthquake. We model possibledirectivity for the largest events by fixing a horizontal line source along each nodal plane and solving forthe best fitting source time function and depth at range of rupture velocities. Using a line source did notsignificantly improve the fit for any of these earthquakes, so we do not observe resolvable directivityand cannot distinguish the fault plane.
2.1. Wharton Basin and Ninety East Ridge
For the nine Mw ≤ 7.0 earthquakes in the Wharton Basin and along the Ninety East Ridge, we use a one-dimensional velocity model at the source with a water layer based on global bathymetry between 2and 5 km, a crustal layer of 7 km (VP of 6.5 km/s) and a mantle half-space (VP of 8.1 km/s). Thinner crust(3.5–4.5 km) has been observed by Singh et al. [2011] in the basin east of the Ninety East Ridge, butsubstituting thin crust has a minimal effect on our modeling. The preferred centroid depths would shal-low by less than 1 km if we replace several kilometers of slower crust with faster mantle in our velocitymodel. A 1-D source structure velocity model is appropriate as the teleseismic waves arrive at near-vertical incidence.
For earthquakes near the trench, we consider the additional complexity that comes from the dipping seafloorinterface. Wiens [1989] modeled the effect that this interface can have by producing significantly different Pwaveforms than simple horizontal layers due to the varying takeoff angles for later arriving water phases. Thedipping seafloor affects strike-slip earthquake seismograms more than those of thrust or normal earthquakesbecause most ray paths are taking off very near to a nodal plane. The P wave water reverberations producedby dipping layers can be mistaken for source and velocity structure complexity. The Swaves do not enter thewater layer and are unaffected by the dipping seafloor so we use them to distinguish between complexity inthe waveforms caused by source and structure. We do not explicitly model the dipping seafloor, but for theearthquakes near the trench we only model complexity in the Pwaves with additional source parameters if itis also present in the SH transverse waves. The effect of dipping sea floor and water multiples is considered inmore detail in the text and in Figures S1 and S2 in the supporting information.
The earthquake centroids that we calculate are distributed throughout the seismogenic zone, with thedeepest edging the 600°C isotherm as calculated using a half-space cooling model (Figure 2). We use anambient mantle temperature of 1350°C and a thermal diffusivity of 1e6m2/s in our half-space cooling model.There is a trade-off between increasingly shallow dips and shallow centroid depth for some events. Depthalso trades off with the source time function; possible depth phases are fit instead by a longer andmore com-plex source. The minimum depths of the two shallowest events, (1999 Mw 6.4 and 2005 Mw 6.7), are not wellconstrained, with little difference in variance from the best fitting depths (8 and 14 km, respectively) to theshallowest tested depth of 2 km. We report the centroids of these two earthquakes as amaximum depth, withslip possible through the oceanic crust. This would be expected assuming a circular fault and constant stressdrop of 5MPa where theMw 6.4 would have a rupture radius of 7.4 km and theMw 6.7 would have a radius of
Journal of Geophysical Research: Solid Earth 10.1002/2015JB012721
ADERHOLD AND ABERCROMBIE SEISMOTECTONICS OF A DIFFUSE BOUNDARY 3466
10.1 km. The depths of the seven other events show clear minima in Figure S3 in the supporting information.The error bars are determined by grid search, based on the depth range where the variance increases by 3%or less over the best fit solution and also where there is no clear visible decline in fit to small phases. Theuncertainty in centroid depth on average is ± 3 km. This is similar to the uncertainty in centroid depth foundby Buchanan [1998] for earthquakes primarily in the Central Indian Ocean Basin. We obtain consistent resultsfor the only earthquake in both studies, 1990 Mw 6.7 (32 and 37 km with overlapping error bars), lendingconfidence that these methods and results are sound.
The earthquake centroids do not show a direct positive relationship between older lithosphere and deepercentroids nor do they follow a relationship of the largest earthquakes having the deepest centroids.However, the deepest earthquake centroids, the 1999 Mw 6.9, the 1990 Mw 6.7, the 2012 Mw 6.3, and the2014 Mw 6.1, lie along the 600°C isotherm (Figure 2).
2.2. The 2014 Mw 6.0 Bay of Bengal Earthquake
The 2014 Mw 6.0 earthquake in the Bay of Bengal is much farther north than the other events, on theedge of the thickest part of the Bengal Fan. Sediments in this region vary between 5 and 16 km, withsteep slopes on the edges where active deposition is occurring, supplied by the Ganges andBrahmaputra rivers [Curray et al., 2002]. The source structure is very different from the Wharton Basinto the south, and the teleseismic P waves show additional phase arrivals that we associate with the thicksediments and dipping seafloor bathymetry. We use a velocity model at the source consisting of a 2 km waterlayer, an averaged 10km layer of sediments (VP of 3 km/s), and a 7 km oceanic crust (VP of 6.5 km/s) over a man-tle half-space (VP of 8.1 km/s) [Curray, 1991]. Our preferred centroid depth 53km (±3 km) is the deepest of theevents that we study, consistent with it rupturing within the oldest lithosphere (Figure 2). Varying the velocitystructure here would only change the depth by a few kilometers, similar to varying the crustal thickness for theearthquakes in the Wharton Basin.
C.N. Rao et al. [2015] used regional waveforms to invert for a more complex structure reflective of the regionand favored a centroid depth for the 2014 Bay of Bengal earthquake between 50 and 54 km, a range thatagrees with our preferred depth. The SCARDEC solution preferred a depth between 44 and 53 km, and theNEIC body wave moment tensor solution has a depth of 47 km, both similar to our preferred depth. Singhet al. [2015] used P waves to obtain a depth range of 60–85 km. The shallower end of this depth rangeoverlaps with our models, taking into account the differences in velocity structure. We attempt to fit the seis-mograms of this earthquake using the greater centroid depths, but we cannot fit the SH waves or pP depthphases that were not considered in Singh et al. [2015].
The oceanic lithosphere at this location is about 110Ma, with relatively thin crust confirmed by high P wavevelocities [C.N. Rao et al., 2015;Mitra et al., 2011; Radhakrishna et al., 2010; Brune et al., 1992]. The 600°C isothermfrom a simple half-space coolingmodel is at a depth of 50 km for 110Ma oceanic lithosphere (Figure 2), but it isalso comparable to a plate-cooling model that takes into account temperature-dependent thermal parameterswith plate thickness of 106 km and mantle potential temperature of 1315°C [McKenzie et al., 2005; Craig et al.,2014]. The centroid of this earthquake, therefore, appears to be at the brittle-ductile transition within theuncertainties in velocity structure, waveform modeling, and thermal structure.
3. Mw≥7.0 Earthquakes
We begin by performing the same point source analysis as above for the three largest earthquakes to obtaincentroid depths and orientation of the nodal planes. The largest earthquake (2010 Mw 7.5) has complexrupture that makes it difficult to determine fault planes unambiguously, but the nodal planes and centroiddepths of the 2005 and 2012 Mw 7.2 earthquakes are sufficiently well resolved to investigate the spatialdistribution of slip using finite-fault inversion. We follow the approach of Aderhold and Abercrombie [2015]and Antolik et al. [2000, 2004, 2006]. Green’s functions are calculated through amodified reflectivity algorithm[Saikia, 1994] using the same oceanic source structure as in the point source models, the IASP91 model forthe mantle [Kennett and Engdahl, 1991] and the QL6 model to account for attenuation [Durek and Ekström,1996]. The linearized inversion method we use minimizes seismic moment, constrains slip to be positive,and is spatially smoothed. Both nodal planes are tested with a range of rupture velocities and spatialdimensions to acquire reliable constraints.
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3.1. The 2010 Mw 7.5 Nicobar Earthquake
The 2010 earthquake had no recorded foreshocks but was followed by 45 earthquakes (M4.0–5.5) within200 km and 30 days of the main shock. The NEIC locations of these events cluster in a N-S striking linearformation proximal to a fossilized fracture zone [Matthews et al., 2011; Rajendran et al., 2011] in Figure 3.Five of the largest earthquakes (Mw 4.9–5.5) have strike-slip mechanisms in the gCMT catalog. The deepestaftershock (mb 5.2) in the NEIC catalog has a hypocenter at a depth of 42.8 km.
The first-motion polarities at stations recording the 2010 main shock require a mechanism with a steeper dipthan the gCMTmoment tensor and one nodal plane with a strike that matches that of the nearby N-S fracturezone as it bends into the trench (Figure 1). A single fault is not enough to represent the rupture of this earth-quake. Stations on either side of the 30° azimuth (TLY and SSE) have different polarities of first motions, yetless than 5 s later they have waveforms very similar to one another (Figure 4). This implies that these stationsare in different quadrants for the first strike-slip subevent and the same quadrant for a later subevent. Themost notable feature is the pervasive negative phase arriving ~25 s after the initial P picks observed at nearlyall stations. We proceed to model the earthquake by constraining the first subevent to the first-motion solu-tion and inverting for the orientation of a second subevent, leaving the depths of both free. The best fittingsolutions are a combination of rupture on a strike slip and a thrust fault.
The 2010 earthquake is located near the Sunda trench. If the second subevent occurred on the slabinterface, it should have a similar mechanism to past seismicity on the slab interface. The nearest earth-quake in the gCMT catalog after 1990 with a thrust mechanism is the 8 January 2007, Mw 6.1 earthquakewith a centroid just under 75 km from the centroid of the 2010 earthquake. The 32° dip of the momenttensor is steeper than the slab dip in this area of 13–15° and the strike is 317°, which is rotated from theslab strike of 333° [Hayes et al., 2012]. Fixing a second subevent with this mechanism improves the fit toseismograms over a single event solution. However, a better fit is achieved by allowing the secondmechanism to be free to rotate counterclockwise off the strike of the subduction zone by ~43°.
We perform a grid search for the orientation of the second subevent, but solutions would not convergeand the orientation could not be well resolved. We therefore adopt a compromise second subeventmechanism that fits the seismograms; this mechanism is ~34° rotated off the expected strike and at a~27° dip for a better fit to the slab interface of the subduction zone at this location. Fixing the first sube-vent on the first-motion mechanism (127°/75°/176°) and the second subevent on the compromisemechanism (119°/64°/98°) results in nearly equal moment on the two subfaults, each hosting slipequivalent to a ~ Mw 7.25 earthquake. The strike-slip subevent prefers a centroid depth 23 km shallowerthan the thrust subevent, with the 29.5 km average centroid depth of the two subevents in agreementwith the gCMT centroid depth of 33.1 km. The depth and duration of both subevents are difficult to con-strain due to the necessary overlapping source time functions, with the depth phases of the first sube-vent masked by the initial phases of the second subevent.
The complexity and uncertainty in the fault orientations of this earthquake render it unsuitable for finite-faultinversion. Instead, we perform forwardmodeling to test our two subevent interpretations. Placing the secondsubevent to be 25 km in each cardinal direction from the first event, the most improvement (3.3%) isachieved in the northeast direction. With depth remaining free, the first subevent remains at about 20 kmand the second subevent varies between 33 and 43 km. The slab interface in this area should be about20 km deep (Figure 1). The second event would have to be about 60 km to the northeast for a depth to slabinterface of 35 km. Placing the subevent at larger distances than 25 km results in a worse fit. Pesicek et al.[2010] identified a westward bias in catalog locations in this region that could account for the discrepancyand would push the locations of both events farther west where the subevent could occur on the deeperslab interface.
An expected half duration of 13.1 s and a centroid time minus hypocenter time of 9.9 s implies that this was arelatively fast earthquake, perhaps due to simultaneous rupture of two fault planes [Duputel et al., 2013]. Thisbehavior is most similar to the observations of the 18 June 2000Mw 7.9 Wharton Basin strike-slip earthquakes,which was fit well with an initial strike-slip subevent and followed by synchronous rupture on a second thrustsubevent [Abercrombie et al., 2003]. The 4 June 2000 Mw 7.9 Enggano was also fit well with a second thrustsubevent, but it had a longer duration than the initial strike-slip subevent, allowing it to be better resolved
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[Abercrombie et al., 2003]. Our modeling indicates that the highly oblique subduction where both the 2005and 2010 earthquakes occurred is undergoing complex deformation, with active strike-slip faults in thedowngoing slab and a thrust fault consistent with slip on the interface of the downgoing plate.
Figure 3. Best fit model of (left column) theMw 7.2 earthquake on 24 July 2005, (middle column) theMw 7.5 earthquake on12 June 2010, and (right column) theMw 7.2 earthquake on 10 January 2012. Pwavemechanism is on top with the SHwavemechanism underneath. Recorded seismograms are solid lines and synthetics are dashed. The letter next to the stationname corresponds to the location on the focal sphere. First motions are determined by handpicked P wave polarities atstations between 30° and 90°. The second subevent of the 2010 earthquake is a compromise mechanism between thenearest thrust event, the slab orientation, and the best fit free mechanism.
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3.2. The 2005 Mw 7.2 NicobarEarthquake
The 2005 strike-slip earthquake occurredon 24 July 2005, less than a year afterthe 2004 Mw 9.2 Sumatra-Andamansubduction earthquake that ruptureda continuous 1300 km segment ofthe trench from 2°N and ending at14°N [Chlieh et al., 2007]. The NEIChypocenter of the 2005 strike-slipearthquake is within 50 km of the sub-duction zone trench and at the samelatitude as the northern portion ofthe subduction zone that experiencedup to 16m of slip during the 2004earthquake (Figure 1a). The 2005 and2010 earthquakes gCMT centroids areonly 26 km apart.
It is challenging to separate seismicityrelated to the 2005 Mw 7.2 earthquakefrom the ongoing and prolific after-shock sequence of the 2004 Mw 9.2earthquake, only 7months before.Within 30 days, and 100 km of theNEIC epicenter of the 2005 earth-quake, there were five earthquakesbefore (M4–4.9, the largest has astrike-slip mechanism in the gCMTcatalog) and 52 after (M3.7–5.6)(Figure 3). The depth of the seismicityfollowing the main shock rangedfrom 9.7 to 43 km. This sequence ofseismicity is more diffuse than thatfollowing the 2010 Nicobar earth-quake, but this is likely an artifact ofbeing superimposed on the after-shock sequence of the 2004 GreatSumatra-Andaman earthquake.
We calculate the best fitting momenttensor to determine the nodal planesto use for the finite-fault inversion ofthe 2005 earthquake. The effect of a dip-ping seafloor can be seen in the teleseis-mic P waves with the relatively largeamplitude later arriving negative andpositive phases that are not fit by thesynthetics at stations in the up-dipdirection like TAM (Figure 4). Includinga 3 km water layer produces phasearrivals with the correct polarity andtiming; however, the amplitudes aremuch smaller than in the recorded data.
Figure 4. Foreshock and aftershock activities for (top) 2005Mw 7.2, (middle)2010 Mw 7.5, and (bottom) 2012 Mw 7.2. Circles mark the 100 and 200 kmradius from the gCMT centroid for each main shock. Foreshocks (red) andaftershocks (yellow) from the ISC catalog are defined, respectively, asoccurring 30 days before and after the main shock. gCMT solutions areshown for the main shock (black) and the largest of the fore/aftershocks.White stars show the NEIC hypocenter of the main shocks. Solid black linesare plate boundaries [Bird, 2003], white dashed lines are fracture zones[Matthews et al., 2011], and solid grey lines are 2m slip contours of the 2004Mw 9.2 earthquake [Chlieh et al., 2007].
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The SH waves are well fit by a simple one source mechanism, leading us to believe that the complexityin the P waves is simply from the structure and would be fit if a dipping seafloor was included inthe model.
We test line source models to investigate directivity for the 2005 earthquake. A slight improvement in fit isachieved with a rupture velocity of 3 km/s at an azimuth of 297°. Since rupture directivity in the 2005earthquake can be observed using a line source, we use this best fit mechanism to define faults along thetwo nodal planes and perform finite-fault modeling (Figure 1). The fault planes are made purposely largeat first to test high rupture velocities, starting at 200 km long and 50 km wide with the hypocenter at thecenter of the fault. A rupture velocity of 2 km/s results in the best fits to seismograms, with the E-W fault planeperforming better than the N-S plane except at the lowest rupture velocity (Figure 5).
Figure 5. Finite-fault model of (left) theMw 7.2 earthquake on 24 July 2005 and (right) theMw 7.2 earthquake on 10 January2012. Slip distribution for the best fit, constrained models are on top. The seismograms of P and SH waves for each modelare on the bottom with synthetics in red and data in black. Inset plots at the bottom show bilateral directivity testing andunilateral directivity testing for the 2005 and 2012 earthquakes.
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We reduce the fault plane for the 2005 earthquake to determine if the fit is affected by removing all but thelargest slip asperity. The constrained fault plane is 105 km in length and 35 km wide with the top of the faultextending to 5 km below the seafloor and a dip of 72°. The hypocenter is positioned at 20 km depth from theseafloor and centered horizontally. Slip is predominantly unilateral to the west and very compact (Figure 5).This agrees with the gCMT catalog that reported a similar half duration and centroid time minus hypocentertime, implying that this earthquake rupture is not unusually long for its magnitude. An average slip of 8.4m isdistributed over a fault area of 330 km2 with a moment of 10.0e19Nm that is larger than the point sourcemodel. The depth distribution of slip extends primarily between 10 km and 30 km, with the peak at 20 km,close to the centroid depth of our point source modeling. Limiting the slip to occur at least 15 km belowthe seafloor does not affect the fit to data, leading us to conclude that surface rupture did not occur duringthe coseismic slip of this event (Figure 6). Restricting the slip to occur only above 20 km results in a worse fit,strengthening the argument for deep (>20 km) slip (Figure 6). Restricting the slip of this fault to strictlyunilateral rupture along each direction of the conjugate fault planes of the preferred mechanism results ina better fit with rupture to the north at 1 km/s but a significant preference of rupture to the west at all higherrupture velocities (Figure 5).
3.3. The 2012 Mw 7.2 Wharton Basin Earthquake
On 10 January 2012, three months before the great Mw 8.6 earthquake, a Mw 7.2 strike-slip earthquakeoccurred near a N-S trending fracture zone. Within 30 days and 100 km, there were no recorded earthquakesprior to the January 2012 event and 23 earthquakes (M4.0–5.2) afterward (Figure 3). The two largest after-shocks, one on 11 January 2012 and the other on 27 January 2012, both have strike-slip mechanisms inthe gCMT catalog but with a shallower dip than the main shock. The aftershock hypocenters are relativelydistributed in the area between two N-S fracture zones and do not fall on a single, well-defined fault plane.
The U.S. Geological Survey (USGS) preliminary finite-fault solution for the 2012 earthquake fits the later partof the seismograms well, but the first arriving phase polarity is not fit by stations from azimuths 103–149°(see Data and Resources). We try to improve on these fits with our preferred mechanism calculated on ournew phase picks. Our mechanism has a similar strike and dip to the gCMT moment tensor and a relativelylong duration of ~25 s (Figure 4). One nodal plane of this mechanism aligns with the strike of the N-S
(a)
(b)
Figure 6. Fault area sensitivity testing for (a) 2005 and (b) 2012 earthquakes. All numbers correspond to the variancereduction for each model.
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oriented fracture zones (Figure 1). The quantitative best fit point source model has a minimum in variancewith shallow slip along this fault above 20 km. However, the beginning of waveforms on some stations tothe southeast like CTAO was fit much better with a centroid depth of 25 km. Here we prefer a centroiddepth of 12 km to reconcile correct polarity and overall fit. Though the fit to seismograms is not perfect,our model shows an improvement over the preliminary USGS model by fitting the polarity of the firstarriving phases at stations to the southeast. All of these modeled depths are well above the 600°C.
The source time function of the 2012 point source model is longer than that of the 2005 event, so we extendour initial fault plane to allow for a longer duration rupture. Our starting geometry is a fault plane of 300 kmby 50 km. The E-W fault plane is preferred over the N-S fault plane and the best fit model at a rupture velocityof 2 km/s (Figure 5). Slip is distributed across much of the fault but with a larger amount of slip to the west ofthe hypocenter. Multiple phases in the first 20 s of the seismograms manifest as slip asperities along strike.The repetitive and evenly spaced vertical slip patches greater than 50 km from the hypocenter to the eastare likely from using spurious slip in the model to fit water phases. Since a water layer is not included inthe source structure, we reduce the fault plane along strike to omit this spurious slip. Our reduced faultgeometry does not have a significant impact on the fit to the earlier part of the seismograms.
The fault geometry of the constrained solution for this earthquake is 102 km long and 25 km wide extendingfrom the seafloor nearly vertically with a dip of 88°. The hypocenter is fit best at a depth of 15 km and is cen-tered in the fault plane along strike. The best solution has a rupture velocity of 2 km/s, resulting in a momentof 7.7e19Nm that is approximately the same as the point source model. Slip is bilateral with more slip towardthe west than the east (Figure 5). The slip is more distributed than the 2005 earthquake, with an average slipof 2.9m over an area of 750 km2. The longer rupture duration of the 2012 earthquake relative to the 2005earthquake is confirmed by the expected half duration of 9.6 s and a gCMT centroid time minus hypocentertime of 14.2 s, implying that this is a longer than average earthquake rupture duration. The slip distributionwith depth peaks at 10 km but with some slip concentrated on the bottom edge of the fault at ~25 km.Limiting slip to only 20 km in depth results in a larger amount of slip at the 20 km bottom edge of the faultplane but no reduction in fit (Figure 6). Allowing slip to only go in one direction at a time along each ofthe two nodal planes in turn shows a preference for the western fault plane (Figure 5). The maximum slipis less than that of the 2005 earthquake, but the fault area experiencing significant slip is larger resulting inthe same magnitude of 7.2.
4. Results and Discussion
We calculate point source moment tensors and centroid depths for 13 earthquakes along the Ninety EastRidge and in the Wharton Basin and Bay of Bengal; the depth distribution is shown in Figure 2. We investigatethe source processes of the three largest earthquakes in more detail, finding one to be a complex earthquakeincluding both strike-slip and thrust subevents (Figure 4) and calculating the preferred fault plane and slipdistribution of the other two that are modeled well with a single fault plane (Figures 5 and 6).
4.1. Depth Distribution and Extent of Seismic Slip
The centroid depths that we obtain are distributed throughout the predicted seismogenic depth range, withthe deepest clustering around the 600°C isotherm. Our centroids and finite-fault inversions show that thecrust is slipping in some of these large earthquakes, but the majority of the seismic slip is occurring withinthemantle. Significant slip from shallow crustal depths down to at least the 600°C isotherm shows that a largeportion of the depth range of these faults can be seismogenic. Whether there is any along-strike or downdipvariation in seismic coupling is completely unknown. For ~M6 earthquakes, the total depth range of large slipis likely to be little more than the error bars in Figure 2. The depth range of slip of the larger earthquakesis more extensive due to the larger area ruptured. The maximum depth of earthquakes increases with age,with both the October 1990 earthquake (32 km, 45–65Ma lithosphere) and the Bay of Bengal earthquake(52 km, 110Ma lithosphere) centroids being at or close to the expected depth of transition from brittle to duc-tile failure. The centroid of the 2012Mw 8.6 (38 km) is also at a similar temperature, but the greater size of thisearthquake means that it is likely that the slip extended further below the predicted brittle-ductile transition.TheMw 8.2 aftershock ruptures deepest into the mantle due to a combination of relatively young lithosphereand large centroid depth. This along with the fact that the 2012Mw 6.3 aftershock is the deepest event in theWharton Basin suggests that the depth of the brittle-ductile transition is increased by strain rate.
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We find that the depth of seismic rupture on oceanic strike-slip faults in the Wharton Basin is primarilythermally controlled and limited by the 600°C isotherm, consistent with the global observations ofMcKenzie et al. [2005] and Jackson et al. [2008] and observations of outer rise seismicity [Craig et al., 2014].
The depth of coseismic slip during large events is proposed to extend below the usual seismogenic limit dueto rupture into a velocity-strengthening lower layer [Shaw and Wesnousky, 2008]. This mechanism could bebehind the unexpectedly deep centroids of the Mw 8+ 11 April 2012 earthquakes (gCMT Catalog), as wellas the deep slip (50 km) modeled by Yue et al. [2012] and Wei et al. [2013]. The constraints on this deep slipare not ideal with teleseismic and regional waves alone due to the complex nature of the rupture and thetiming of the deepest slip in the middle of the wave train of earlier subevents. However, significant slip atdepths of at least 60 km is also required to fit geodetic data [Hill et al., 2015]. Duputel et al. [2012] obtainedcentroids more consistent with the expected thermal limit to brittle failure, but significant slip is still likelyto have extended deeper than the centroid and below the predicted brittle-ductile transition. A seismicreflection study imaged faults extending down to at least 45 km in this region, showing that deformationis possible to these expected depths [Carton et al., 2014], but the shallow dip of these faults is inconsistentwith the rupture of the great 2012 earthquakes.
The deepest centroid in the Wharton Basin was the Mw 6.3 aftershock that occurred 4 days after the Mw 8.6earthquake. It ruptured to a depth of at least 38 km in the same area where the final subevent of theMw 8.6 earthquake likely occurred [Duputel et al., 2012]. This significant depth could be driven by the deepen-ing of the brittle failure limit by postseismic slip and an increase in strain rate in the immediate aftermath ofthe larger earthquake [Rolandone et al., 2004].
Along oceanic transforms, slip is distributed throughout the thermally controlled seismogenic zone[Abercrombie and Ekström, 2001; Braunmiller and Nábělek, 2008]. We know that there is significant aseismicslip on oceanic transform faults by using the known slip rate and the cumulative seismic moment to deter-mine the expected coupled area [Brune, 1968; Boettcher and Jordan, 2004]. Intraplate earthquakes also appearto be distributed throughout the thermally controlled seismogenic zone, but as there are no good constraintson the slip rate of intraplate faulting, we do not know if significant slip is accommodated aseismically.
Our model places the slip of the 10 January 2012 earthquake shallower than most catalog depths. Due to thewell-constrained dip and our preferred shallow depth, we do not believe that this earthquake ruptured alongthe shallow dipping and deep faults imaged by Qin and Singh [2015]. This highlights the need for moreseismic imaging and studies of microseismicity in this region.
4.2. Active Faults in the Wharton Basin
Nine of the moment tensors have a nodal plane striking within 7° (average of 3°) of the mapped N-S trendingfracture zones and are consistent with left-lateral slip along these faults. There is a positive trend of higherstrike rotations with earthquakes located at higher latitudes, with the maximum rotation for the farthestnorth Bay of Bengal earthquake (Figure S17 in the supporting information). The Bay of Bengal earthquakecould be hosted on the same N-S trending fracture zones proposed for the southern events [G.S. Rao et al.,2015]; however, it is difficult to determine if the orientation of these faults changes beneath the thick sedi-ment layer. The remaining three events all occur near the trench, and their orientations are likely affectedby bending as the slab descends.
The 11 April 2012 earthquakes were the largest oceanic intraplate events ever recorded. With greaterresearch capabilities made possible by large seismometer arrays such as Hi-net in Japan and new techni-ques developed around them such as back projection [Ishii et al., 2005] as well as enhanced aftershockdetection from the greater station coverage, different groups were able to make observations of theseevents that suggest a complex strike-slip rupture pattern along both N-S fracture zones and approximatelyorthogonally orientated E-W faults [Wang et al., 2012; Duputel et al., 2012; Meng et al., 2012; Zhang et al.,2012; Ishii et al., 2013; Wei et al., 2013; Yadav et al., 2013]. Along the Ninety East Ridge, seismic reflectionprofiles reveal a change in faulting from transpressional WNW-ESE faults (0–5°N) to thrust/strike-slipWNW-ESE faults (5–8°S) [Sagar et al., 2013]. This changes again to the south, implying that the northernportion of the Ninety East Ridge is controlled by the motion of India-Australia plates and the southern iscontrolled by motion of the India-Capricorn plates [Sagar et al., 2013]. In a detailed survey of the southernpart of the Wharton Basin around 93°E, E-W-oriented oceanic fabric associated with the fossilized ridge
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was mapped [Deplus et al., 1998]. Despite this, prior to the 11 April 2012 earthquakes strike-slip rupture inthe Wharton Basin was assumed to follow the prominent N-S trending fracture zones.
Finite-fault inversion of the 10 January 2012 foreshock to the 11 April 2012 events has a better fit for bilateralrupture along an E-W fault, with more slip to the west than the east. This would match the strike of the north-ern linear E-W aftershock sequence of the 2012 events (grey dots in Figure 1), corresponding to the fault thathosted the first subevent of the Mw 8.6 event [Meng et al., 2012; Yue et al., 2012; Duputel et al., 2012; Zhanget al., 2012; Ishii et al., 2013;Wei et al., 2013]. This study shows that slip is not limited to the E-W fracture zonesin the Wharton Basin only in large magnitude events, but that rupture is also hosted on E-W faults duringsmaller magnitude events.
The 2005 and 2012 earthquakes both prefer westward rupture; however, the directivity is not stronglyobserved and north-south rupture cannot be ruled out. Our preferred models of these events havevery different characteristics otherwise. The slip model of the 2005 earthquake has compact unilateralslip concentrated deeper at 20 km depth, extending 24 km to the west. The model of the 2012 earth-quake has dispersed bilateral slip along the entire 100 km length of the fault, concentrated slightlyshallower at 10 km. Despite the differences in slip behavior, the models of these two events had thesame magnitude of 7.2.
The preferred model of the 2010 Mw 7.5 earthquake is comprised of two subevents, the first is a strike slipand the second is a thrust with a mechanism consistent with the subduction zone interface. Both sube-vents were of equal moment, each equivalent to a Mw 7.3 earthquake. The greater depth of the reversesubevent is consistent with the global distribution observed by Craig et al. [2014] for outer rise earthquakesrelated to plate bending. The combination of strike-slip and reverse subevents would be similar to thetwo Mw 7.9 2000 earthquakes in this region, where a strike-slip subevent initiated rupture along anadjacent thrust fault [Abercrombie et al., 2003]. The 4 June 2000 Enggano earthquake ruptured a strike-slipfault for at least 30 s, with a second reverse fault activated 13 s into the main rupture. In the case of theEnggano earthquake the total duration was ~100 s and rupture extended to a depth of 57 km, muchlonger and deeper than the smaller magnitude 2010 earthquake in this study. Aftershocks of the 2010earthquake cluster closely on the strike of the N-S fracture zone that coincides with the centroid location,extending for a distance of ~100 km and consisting of both strike-slip and normal faulting events. Thisdistribution suggests that the earthquake probably ruptured the N-S fracture zone before proceedingto rupture a thrust fault either coincident with the subduction zone interface or at a greater depth drivenby plate bending.
On 2 March 2016 a Mw 7.8 earthquake ruptured a fault near the 94° Fracture Zone while this paper was inrevision. This earthquake had a short (≤50 km) and compact rupture for its size, evidenced by rapid finite-faultmodeling (see Data and Resources) and preliminary back projection [Trabant et al., 2012]. It is difficult todiscern directivity of a rupture of this length, and the rapid finite-fault modeling was unable to identify thefault plane. The aftershocks (mb ≤ 5.5) are primarily located to the north of the main shock hypocenter, butuncertainties are large (Figure 1). The occurrence of this earthquake demonstrates that routine proceduresare unable to determine the fault planes of earthquakes in this region. Further work of the kind performedhere is needed to resolve the source characteristics and active structures within the enigmatic Wharton Basin.
5. Conclusions
We model strike-slip earthquakes of magnitude 6 ≤Mw ≤ 8 in the Indian Ocean to investigate the depth dis-tribution of slip and orientation of active faults in this region of complex deformation. We determine double-couplemoment tensors and centroid depths for 13 earthquakes and resolve the slip distribution of two of thelargest events.
1. Seismic slip extends throughout the seismogenic zone, rupturing both the crust and mantle. Earthquakecentroid depths near to the expected brittle-ductile transition imply that seismic slip may extend into theusually ductile zone during large earthquakes. None of the earthquakes had slip that extended to suchhigh isotherms as the great earthquakes of 11 April 2012.
2. The 2014 Bay of Bengal earthquake is consistent with slip near the brittle-ductile transition within oceaniclithosphere of 110Ma, despite having the greatest centroid depth.
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3. The 2010 Mw 7.5 earthquake in the Wharton Basin ruptured along two faults, initiating along a strike-slipfault in the subducting oceanic lithosphere and triggering slip on a thrust fault, consistent with thesubduction interface.
4. All of the earthquakes except the 2014 Bay of Bengal earthquake have a nodal plane between 0° and22° (average of 6°) off of the N-S trending fossil fracture zones. The best fit finite-fault models for the2005 and 2012 Mw 7.2 earthquakes both involved slip on the E-W-oriented nodal plane. There wasonly a slight difference in fit between the two nodal planes. Although this provides some supportfor the existence of active E-W faults in the region, stronger evidence would be preferred. Morestudies to image the faults are needed to confirm the existence of active E-W faulting in this location.However, teleseismic observations remain the only viable option for characterizing the majority ofoceanic strike-slip faults in the world.
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AcknowledgmentsThis work was supported by theNational Science Foundation throughaward OCE 0850503. We thank M.Antolik for sharing his inversion codeand three anonymous reviewers fortheir valuable feedback. The facilities ofIRIS Data Services and specifically theIRIS Data Management Center wereused for access to waveforms used inthis study. IRIS Data Services are fundedthrough the Seismological Facilities forthe Advancement of Geoscience andEarthScope (SAGE) Proposal of theNational Science Foundation underCooperative Agreement EAR-1261681(last accessed February 2015). Some ofthe figures were made using theGeneric Mapping Tools (GMT) version4.5.8 [www.soest.hawaii.edu/gmt;Wessel and Smith, 1998] (last accessedJune 2014). Seismic Analysis Code (SAC)was used during data processing[Goldstein and Snoke, 2012] (lastaccessed September 2010). The GlobalCentroid Moment Tensor Project data-base was also used [www.globalcmt.org] (last accessed March 2016).Bathymetry data came from theGEBCO_08 Grid version 20100927[www.gebco.net] (last accessed June2014). Preliminary finite-fault models forthe 10 January 2012 earthquake and the2 March 2016 earthquake by G. Hayeswere accessed online from ANSSComprehensive Catalog [www.comcat.cr.usgs.gov] (last accessed March 2016).Aftershock locations and magnitudeswere accessed from the FDSN throughthe IRIS DMC FetchEvent tool (seiscode.iris.washington.edu/projects/ws-fetch-scripts) (last accessed March 2016).Earthquakes were provided at inconve-nient times by Earth (last accessedMarch 2016).
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