The June 2000 Mw 7.9 earthquakes south of Sumatra: Deformation in

the India–Australia Plate

Rachel E. Abercrombie,1,2 Michael Antolik,1 and Goran Ekstrom1

Received 15 June 2001; revised 19 July 2002; accepted 23 October 2002; published 11 January 2003.

[1] Two large (Mw 7.9) earthquakes occurred on 4 and 18 June 2000, south of Sumatra,beneath the Indian Ocean. Both earthquakes were predominantly left-lateral strike-slip onvertical N-S trending faults that we interpret to be reactivated fracture zones. The 4 JuneEnggano earthquake occurred at the edge of the rupture area of the 1833 subductionearthquake. The first strike-slip subevent within the subducting plate triggered a thrustsubevent on the plate interface, which comprised at least 35% of the total moment andruptured SE away from the 1833 earthquake. The 18 June earthquake in the WhartonBasin is one of the largest shallow strike-slip faulting earthquakes ever recorded. A smallsecond subevent with reverse slip is required to fit the body waves. The orientation of bothsubevents in our preferred model is consistent with the current stress field in the region.Both the June 2000 earthquakes are consistent with recent models of distributeddeformation in the India–Australia composite plate. The occurrence of the Engganoearthquake implies that the stress field within the Indian plate continues to a depth of 50km in the subducting slab. The purely strike-slip source model of the Wharton Basinearthquake obtained by Robinson et al. [2001] matches the P waves very poorly and fitsthe S waves no better than our preferred model. The strike-slip subevents of bothearthquakes had few aftershocks and higher stress drops than the subduction thrustsubevent of the Enggano earthquake. This difference is consistent with previousobservations of oceanic and subduction earthquakes. INDEX TERMS: 7209 Seismology:

Earthquake dynamics and mechanics; 7215 Seismology: Earthquake parameters; 7230 Seismology: Seismicity

and seismotectonics

Citation: Abercrombie, R. E., M. Antolik, and G. Ekstrom, The June 2000 Mw 7.9 earthquakes south of Sumatra: Deformation in the

India–Australia Plate, J. Geophys. Res., 108(B1), 2018, doi:10.1029/2001JB000674, 2003.

1. Introduction

[2] On 4 June 2000 a large (Mw 7.9) earthquake occurrednear Enggano Island, off the coast of Sumatra, killing 90people and injuring 2000 more. It was felt as far away asSingapore where many high-rise buildings were shaken[Pan et al., 2001]. This earthquake was followed 2 weekslater by a second earthquake of similar magnitude, approx-imately 1000 km to the south, beneath the Indian Ocean.Both of these earthquakes were largely unexpected, and theaim of our study is to determine their rupture geometry andto investigate what they reveal about the complex seismo-tectonics of the Indian Ocean and western Indonesia.[3] The 4 June Enggano earthquake occurred within the

seismically active Sumatra subduction zone. The unusualaspect of this earthquake was its source mechanism. TheHarvard catalogue centroid moment tensor (CMT) mech-anism [Dziewonski and Woodhouse, 1983; Dziewonski et

al., 1981, 2001] is primarily strike-slip. The CMT solutionalso has a large, but poorly resolved, non-double-couplecomponent. The Indian plate is subducting beneath Suma-tra at an oblique angle [DeMets et al., 1994a], and thedeformation is divided between predominantly reversemotion at the trench (see Figure 1) [e.g., Fitch, 1972;Newcomb and McCann, 1987; McCaffrey et al., 2002] andright-lateral strike-slip on the Great Sumatra fault [e.g.,Sieh and Natawidjaja, 2000; Genrich et al., 2002]. Themore recently discovered Mentawai fault in the forearc[Diament et al., 1992] does not appear to be active [Genrichet al., 2000]. Sumatra is often cited as a classic example of aso-called slip-partitioned margin [Fitch, 1972]. The En-ggano earthquake was too deep (�50 km) to be on either theSumatra or the Mentawai faults. The epicenter of theEnggano earthquake is in a region of continuously highseismicity over the last few decades. It is also at thesoutheastern end of the rupture area of the great (M > 8)1833 subduction earthquake (Figure 2), as determined fromintensity reports [Newcomb and McCann, 1987] andsubmergence and uplift of corals [Zachariasen et al.,1999]. In the rupture area of the 1833 earthquake, the upperand lower subducting plates are thought to be fullyseismically coupled [Zachariasen et al., 2000], but thecoupling appears to decrease near the equator [e.g.,

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. B1, 2018, doi:10.1029/2001JB000674, 2003

1Department of Earth and Planetary Sciences, Harvard University,Cambridge, Massachusetts, USA.

2Now at Department of Earth Sciences, Boston University, Boston,Massachusetts., USA.

Copyright 2003 by the American Geophysical Union.0148-0227/03/2001JB000674$09.00

ESE 6 - 1

Prawirodirdjo et al., 1997]. South of Java, where thesubducting slab is older and thicker, subduction is nearlyaseismic and close to the direction of relative plate motion[e.g., Newcomb and McCann, 1987; Abercrombie et al.,2001].[4] The 18 June 2000 Wharton Basin earthquake is the

largest known earthquake in the central Indian Ocean, andit is one of the largest strike-slip earthquakes ever recor-ded. There are no previous earthquakes in either the Inter-national Seismological Centre (ISC) or NEIC catalogues(since 1970, complete above M � 5) within over 150 kmof the hypocenter of the 18 June earthquake. The twoearthquakes near its hypocenter in Figure 1 are bothaftershocks. Robinson et al. [2001] modeled the S wavesand concluded that this earthquake had an unusual rupturegeometry. Their preferred model involves simultaneousstrike-slip motion on two nearly perpendicular faults, withthe slip centroids separated by about 45 km. Such a ruptureprocess is unprecedented and therefore we investigate howwell it is resolved.

[5] The Indian Ocean is the most seismically active oce-anic plate interior on Earth [Bergman and Solomon, 1985].The distributed deformation is thought to be a consequenceof the collision of India with Eurasia impeding theapproximately NNE motion of the India–Australia platein the west and continuing subduction along the Sunda Arcbeneath Java and Sumatra. Seismicity is distributed alongand around the Ninety East (90�E) ridge, a prominentfeature (Figure 1) that is thought to be an old hot spot tracefollowing a fossil transform fault [e.g., Stein and Okal,1978]. To the west of the 90�E Ridge, long wavelengthfolds trend E-W [e.g., Weissel et al., 1980], and similarlyoriented faults have been identified as currently active[Chamot-Rooke et al., 1993; Bull and Scrutton, 1990]. Theseismicity is characterized by reverse faulting earthquakesstriking E-W and strike-slip faulting earthquakes strikingNW-SE (or NE-SW) (see Figure 1). To the east of the 90�Eridge, however, the earthquakes occur on N-S (or E-W)striking strike-slip faults and a few NE-SW striking reversefaults. The area to the east of the 90�E ridge, including the

Figure 1. Location map of the 4 June (Enggano) and 18 June (Wharton Basin) earthquakes. TheHarvard CMTs are joined to their NEIC epicenters. The Harvard catalogue CMTs for earthquakes Mw �6.5 are shown in midgray. In the Indian Ocean (where the seismicity rate is lower than along the Sumatraand Java trenches), all mechanisms in the Harvard catalogue are shown (pale gray). All mechanisms areplotted as lower hemisphere projections throughout the paper. The arrow indicates the motion of theIndia–Australian plate relative to the Eurasian plate [DeMets et al., 1994b]. The bathymetry is shaded inthis and in subsequent maps [from Smith and Sandwell, 1997].

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Wharton Basin, is cut by many N-S trending fossiltransform faults, including the Investigator Fracture Zone,clearly seen in the bathymetry (Figure 1). Recent seismicprofiles in the Wharton Basin show evidence that these N-Sfaults are currently being reactivated as left-lateral strike-slip faults [Deplus et al., 1998]. These various observationsimply a change in the orientation of the maximumcompressive stress from N-S west of the 90�E ridge, tomore NW-SE in the east [e.g., Cloetingh and Wortel, 1986;Kreemer et al., 2000].[6] Studies of plate motion in the Indian Ocean region

cannot simultaneously fit data along the three mid-ocean

ridge systems that meet at the triple junction betweenIndia–Australian, Antarctic, and African plates [Minsterand Jordan, 1978; Gordon et al., 1987]. These observationshave led to proposals that the India–Australia plate shouldbe considered as two or more separate plates, divided by adistributed zone of deformation. InitiallyMinster and Jordan[1978] andOkal and Stein [1978] proposed a two-platemodeldivided by a broad N-S zone of distributed deformationparallel to the 90�E ridge. Wiens et al. [1985], Gordon et al.[1990], and DeMets et al. [1994b] preferred a broad equa-torial plate boundary striking roughly E-W. The WhartonBasin earthquake occurred about 5� south of the distributed

Figure 2. (a) Map showing the seismotectonic setting of the Enggano earthquake sequence. TheHarvard CMT mechanism for the 4 June main shock is joined to its NEIC epicenter (black star) and CMTcentroid (white star). The mechanisms of earthquakes in the Harvard catalogue Mw � 6.5 are plotted ingray at their centroid locations joined to their NEIC epicenters. All seismicity in the ISC catalogue from1973–1998, relocated as described in the text, is plotted (small gray circles). All the NEIC locatedaftershocks from June to October 2000 are plotted as black circles. Note how well the aftershocksequence matches the location of the area of high background seismicity. The approximate rupture area ofthe 1833 subduction zone earthquake [Newcomb and McCann, 1987; Prawirodirdjo et al., 1997] ismarked (thick gray box), and the area shown in (b) is outlined by the black rectangle. The bathymetry isshaded as in Figure 1. Enggano Island is where the maximum damage was reported [Pan et al., 2000]. (b)Close-up map of the Enggano earthquake sequence. The mechanisms of all earthquakes in the HarvardCMT catalogue before June 2000 (pale gray) and those of the earthquakes occurring in June–October2000 (dark gray) are plotted at their relocated epicenters. The mechanisms of two large (M6.8 and M7.3)earthquakes that occurred in January and February 2001, after this work was completed, are labeled 01.All earthquakes located by the NEIC from June to October are shown as black circles.

ABERCROMBIE ET AL.: THE JUNE 2000 INDIAN OCEAN EARTHQUAKES ESE 6 - 3

zone of deformation proposed by Gordon et al. [1990](Figure 3). More recently, a three-plate model has beenproposed to explain the region of complex deformation in thecentral Indian Ocean (Figure 3) [Royer and Gordon, 1997].[7] The two large earthquakes that occurred on 4 and 18

June 2000 can provide critical constraints on the nature ofpresent-day deformation in the complex region of the IndianOcean. We perform waveform modeling and relocate after-shocks to determine the fault planes of the two earthquakes.We then interpret our results in the context of the regionalseismotectonics. We also consider what we can learn aboutthe deformation of oceanic lithosphere from the sourceproperties of the two earthquake sequences.

2. 4 June 2000 Enggano Earthquake Sequence

2.1. Broadband Body Wave Modeling of the MainShock

[8] We model teleseismically recorded broadband bodywaves to investigate the rupture geometry of the Enggano

earthquake. The seismograms are relatively long and com-plex (Figure 4), even for an earthquake of this size.We use allthe available recordings of both P and SH waves at GlobalSeismic Network (GSN) stations. We correct for theinstrument responses to obtain displacement seismogramsfiltered between 1 and 100 s. We then invert for the momenttensor, source time function, depth, and rupture directivityusing the method developed by Ekstrom [1989]. We correctfor attenuation using a t* (travel time divided by average Q)of 0.6 s for P waves and 3.0 s for S waves and a layeredvelocity model over a half-space, with a 23 km thick crust(essentially PREM) and 2 km of water. We pick P and Sarrivals for the inversion, but fix the source at the NEICepicenter. Initially, we invert only P waves and predict the Swaves to assist in the correct picking of the S wave arrivals.The method allows us to include the CMT solution as aconstraint in the body wave inversion, but we do not usethis option. If we allow the moment tensor to include a non-double-couple component, we can fit the waveforms with asingle mechanism similar to that of the CMT solution.

Figure 2. (continued)

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When we constrain the source to be a double couple,however, we cannot fit the waveforms with a singlemechanism. A close look at the P wave onsets shows thatthey require a strike-slip solution (Figure 4). For example,note the change from positive first motion in the NE (e.g.,TLY round to PMG) to negative in the SE (CAN, SNZO, andVNDA). The later phases (�25–75 s after the P onset) arethe same polarity in the NE and SE, implying a change ofmechanism. We thus try inverting for two separate mechan-isms. We fix the first mechanism at strike-slip solutionsconsistent with the polarity and amplitude of the firstmotions (which provide tight constraints of ±2�) (Figure 4)and invert for the mechanism and source parameters of thesecond subevent. We vary the onset time of the secondsubevent and the duration of both subevents. Initialinversions demonstrate that the second subevent has to bepredominantly reverse faulting, striking approximately NW-

SE. When we find a solution that fits the P waves well, wecan use the S waves predicted by the P wave model to assistin picking the S arrivals in the data. We can then include the Swaves in the inversion. They provide sufficient additionalconstraints to enable us to invert for all the source parametersof both subevents and obtain a stable solution.[9] We also consider the possibility that the second,

reverse faulting subevent is simply an artifact produced byusing a simple 1-D velocity model to match waveforms thatwere excited in a dipping structure. Wiens [1989] showedthat dipping interfaces between sediments, crust and mantlehave a relatively small effect on waveforms, but thatreverberations in a deep water layer above a dippingseafloor produce late arriving pulses, which are not presentin a horizontal geometry and that can easily be mistaken forsource complexity. Such water multiples are only observedin the P waves, and the SH waves are unaffected. Both P

Figure 3. Distribution of seismicity in the Indian Ocean. All Harvard catalogue CMT solutions(midgray) for intraplate earthquakes are plotted, but only those with M � 6.5 are shown along plateboundaries (thin black lines). The 4 and 18 June earthquakes are shown in dark gray, and mechanisms forearthquakes prior to 1976 are light gray [Bergman and Solomon, 1985; Petroy and Wiens, 1989]. Thethick gray line approximately outlines the area of diffuse deformation defined by Gordon et al. [1990].The pale shaded regions are the diffuse ‘‘plate boundaries’’ proposed by Royer and Gordon [1997]between their Capricorn, Australian, and Indian ‘‘subplates.’’ Also shown are their proposed poles ofrotation and 95% confidence ellipses (C-I represents the pole for Capricorn, India, etc.). The darker graytriangular region is their locus of their triple junction, and the arrows indicate directions of maximumcompression (black) and extension (gray). Note that left-lateral slip on N-S faults is consistent with thisdeformation pattern east of the 90�E ridge in the region of the June 2000 earthquakes.

ABERCROMBIE ET AL.: THE JUNE 2000 INDIAN OCEAN EARTHQUAKES ESE 6 - 5

Figure 4. Broadband body wave modeling of the Enggano (4 June) earthquake, showing the observedP and SH seismograms (solid) and synthetics (dash) for our preferred model. The vertical lines on eachseismogram indicate the time window used in the inversion. (Those seismograms without such lines werenot used to constrain the inversion and are simply forward modeled because of their epicentral distanceand high noise level or to balance the azimuthal distribution in the inversion.) The vertical arrows are thepicked onset times, and the numbers to the right of the seismograms are the maximum amplitudes inmicrons. The stations are located on the focal mechanisms of the first of the two subevents, and themoment rate function of the first subevent is shown beneath the P wave mechanism. Both subeventmechanisms are shown at the bottom. The strike, dip, and rake of the two mechanisms are 194�, 82� and�1�, and 330�, 37� and 72�, respectively. The arrows on the focal mechanisms indicate the direction ofrupture propagation (at velocity VR). The moment rate functions of the two subevents are also shown,separately and combined.

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and SH seismograms of the Enggano earthquake include thesecond subevent. Also, Wiens [1989] found that suchreverberations are insignificant for earthquakes at therelatively large depth and shallow water layer of theEnggano earthquake. Complex reverberations in the dippingsedimentary layers could produce both complex P and SHwaveforms, but they should exhibit strong azimuthalvariation not observed in the waveforms of the Engganoearthquake. In addition, no such phases are observed in theP and SH seismograms of the largest aftershocks. Therefore,we conclude that the reverse faulting subevent is real andnot an artifact of an oversimplified model velocity structure.It is probable that structural heterogeneity is contributing tothe smaller-scale complexity of the waveforms, especiallythe later arrivals, and so we do not try a more complicatedsource model that attempts to match every detail.[10] We estimate the uncertainties in our preferred 2

subevent solution using the measured RMS misfit betweenobserved and synthetic seismograms and also noticeabledeterioration in the visual fit, especially the polarity of thefirst motions. We find that there is a relatively broad rangeof solutions that can fit the data adequately, with littlechange in RMS. The first motions constrain the mechanismof subevent 1 to within about 2�, but the SH waves areslightly better fit by a mechanism rotated about 10�anticlockwise. This discrepancy may represent a curvedfault plane, but we do not have enough resolution toconfirm this. The depth of subevent 1 can vary ±10 km withno significant change in RMS. The first subevent is a singlepulse of slip lasting about 30 s and includes up to two-thirdsof the moment. The second subevent starts 13 s after thefirst; the onset is well constrained by arrivals in theseismograms not generated by the strike-slip source. Thesecond subevent involves three well-resolved pulses spreadout over almost 100 s. The fourth small pulse at �100 s(Figure 4) is poorly resolved; shortening the solution toexclude this fourth pulse has a negligible effect on the RMSerror, whereas shortening it to only the first two pulses (50 s)increases the RMS error by over 10%. The depth of thesecond subevent can vary by ±10 km with negligible effecton the waveform fits. The west dipping plane of themechanism is the better constrained. The orientation of thenull axis can vary by about 15� and the strike and dip ofthe east dipping plane between 300� and 315� and 28� and40�, respectively, without significant increase in RMS. Themoment of the second subevent varies from about a third toa half of the total, depending on its orientation. Bothsubevents show significant directivity. The first subeventruptured to the north, which implies that the NS plane is thefault plane. The second subevent ruptured to the SE, alongthe strike of the subduction zone. Fixing either source as apoint source increases the RMS error by about 5% andresults in an obvious decrease in visual fit. We thus believethe directivity in both subevents to be resolvable and therupture propagation velocity to be �2 km/s. The sum of themoment tensors of the two subevents is close to the Harvardcatalogue CMT solution derived from long period body andmantle waves.

2.2. Aftershocks and Background Seismicity

[11] The 4 June Enggano earthquake had an activeaftershock sequence, including 51 earthquakes M � 5 in

the first month (Figure 2). The epicentral distribution isclose to the pattern of background seismicity in this activeregion. This seismicity pattern has been stable over the lastfew decades, with about 8 earthquakes M � 5 each year inthe area of Figure 2b. The aftershock distribution is alsoconsistent with our preferred rupture geometry of theEnggano earthquake. Most aftershocks are located to theSE of the main shock epicenter, extending 150 km alongthe subduction interface, coinciding with the inferredrupture area of the second subevent. The aftershocks alsoextend over 50 km north of the main shock epicenter, alongthe preferred fault plane of the first subevent. The HarvardCMT mechanisms of the larger aftershocks exhibit consid-erable diversity (Figure 2). Some are consistent withreverse slip on the plate interface, and the others are mostlystrike-slip with nodal planes in various orientations. Toconfirm the reliability of the CMT mechanisms, we modelthe body waves of the three largest aftershocks with verydifferent CMT mechanisms. Our results are consistent withthe CMT catalogue solutions, and thus we think that thecatalogue mechanisms are reliable and that the diversity isreal.[12] In order to investigate further the rupture geometry of

the sequence we relocate the aftershocks recorded by theNEIC from June to October 2000, using the hypocentroiddecomposition method of relative relocation developed byJordan and Sverdrup [1981] and Bergman and Solomon[1990]. We use direct and reflected P and S phases reportedby the NEIC in the inversions. The largest earthquakes areusually the best recorded, so we start with those aftershockswith M � 5. We discard the ones with the smallest numberof recorded phases to enable the inversion to converge. Wethen add the smaller earthquakes with the largest number ofreported phases, resulting in a catalogue of 64 earthquakeswith hypocentral relocation uncertainties of less than 15 km(all but 8 are under 10 km), and these are plotted in Figure 5.The epicentral distribution differs insignificantly from thatof the NEIC, but our relocations have well resolved depths,whereas all but the three deepest aftershocks are fixed at 33km depth by the NEIC.[13] We then apply the same relocation technique to the

earthquakes recorded by the ISC between 1973 and 1998along the Sumatran subduction zone. We divide the regionfrom 0�S to 7�S into overlapping rectangles, allowingcomparison of hypocenter relocations derived for earth-quakes included in more than one inversion. In this relo-cation, we use direct and reflected P and S phasesreported by the ISC. Our relocated catalogue includes manyearthquakes that have larger location errors than theaftershock sequence, with depth errors of up to 25 km.These relocations significantly improve our resolution of thesubducting slab compared to the original ISC locations,which show only a very diffuse dipping structure over 100km thick. Our relocations are sufficiently accurate toprovide the context for the Enggano earthquake sequencethat is necessary here. Further work to improve theselocations and investigate the spatiotemporal patterns ofseismicity along the Sunda Arc is planned.[14] The main shock is included in the relocation of the

aftershock sequence and has a hypocentral depth of 38 ± 6km (95% confidence ellipse), shallower than the broadbandcentroid depths but within the error bounds. It is plotted

ABERCROMBIE ET AL.: THE JUNE 2000 INDIAN OCEAN EARTHQUAKES ESE 6 - 7

separately in Figure 5b. The relative depths of the hypo-center and centroids suggest downward rupture. The bodywave inversion results predict more horizontal rupturedirections, but the resolution of the data cannot rule out acomponent of downward rupture.[15] The rupture propagation in the second subevent is

�2 km/s to the SE, and the rupture lasted about 90 s. Thus,the halfway point is approximately 100 km SE of thehypocenter, and it is at this point that we plot the mecha-nism of the second subevent in Figure 5. This location doesnot imply any preference for the actual slip distribution, butis simply chosen to represent the preferred geometry fromthe broadband modeling.[16] The maximum depth of the aftershocks decreases to

the SE (Figure 5, cross section A), and several aftershocksare at similar depths to the centroid of subevent 1. Thethickness of the subducting slab defined by our relocatedbackground seismicity is poorly constrained, due to the

relatively large uncertainties in the locations, but subevent 2and the related thrust aftershocks are consistent with slip atthe slab interface (Figure 5, cross section B). The dips of thenodal planes of these earthquakes are also in good agree-ment with the dip of the subducting oceanic lithosphere.Subevent 1 is deeper and, within the error range canconfidently be considered within the subducting slab. Crosssection C (Figure 5) is through the main shock hypocenter,approximately perpendicular to our preferred N-S faultplane for subevent 1. Three aftershocks lying to the southof subevent 1 (Figure 5a) have similar strike-slip mecha-nisms to subevent 1 and so are consistent with our inter-pretation of a N-S fault plane within the subducting slab.

2.3. Summary

[17] Our preferred source model for the Enggano earth-quake consists of two separate subevents. The first (�65% ofthe moment) involves strike-slip motion on a near vertical,

Figure 5. The geometry of the Enggano earthquake and aftershock sequence. (a) Map showing therelocated seismicity and locations of cross sections A, B, and C. The relocated earthquakes in the ISCcatalogue (1973–1998) are plotted as gray circles and the relocated aftershocks from June to October2000 as black circles. The hypocenter of the main shock is shown in gray with a thick black outline.Aftershocks with CMT mechanisms are plotted at their relocated hypocenters (best double-couplecomponent only). The main shock subevents have thick outlines. The first subevent of the Engganoearthquake is plotted at the epicenter. The second subevent is plotted at the approximate center of itsrupture area. (b) Cross sections with front projections of the focal mechanisms with symbols as in (a).Note that the first subevent of the main shock is plotted at the depth determined in the body wavemodeling, deeper than the relocated hypocenter. A is a face-on view of the slab. B shows the subductingplate. Subevent 2 and many aftershocks are consistent with thrust faulting at the plate interface. C is thealignment of the �N-S preferred fault plane of the main shock with three aftershocks.

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approximately N-S trending fault within the subducting slab,lasting �30 s. The rupture propagated toward the north fromthe hypocenter. This first subevent then triggered reverse slipon the subduction interface, rupturing toward the SE. Thissecond subevent consisted of at least three pulses of slip,lasting a total of �90 s. The locations of the aftershocksprovide further support for this interpretation. The orienta-tion of the strike-slip subevent would lead to increased staticshear stress on the plate interface to the SE, consistent withtriggering slip there in the second subevent.

3. 18 June Wharton Basin Earthquake

[18] We perform a similar analysis of the 18 June Whar-ton Basin earthquake, combining broadband body wavemodeling and aftershock relocation to investigate the rup-ture geometry.[19] Robinson et al. [2001] analyzed the Wharton Basin

earthquake and obtained an unusual rupture geometryinvolving slip on two perpendicular strike-slip faults. Theyrecalculated the long period moment tensor and determinedthat the large non-double-couple component in the HarvardCMT solution (35%) is not well resolved. We investigate theuncertainties in the Harvard CMT solution and also find thatthe best double couple solution (strike, dip, and rake are 168�,53�, and 11�, respectively) has an RMS error insignificantlyhigher than that of the solution in the catalogue. Thus the non-double-couple component cannot be resolved from the long

period waves alone. We do not use the Harvard CMT as aconstraint on the body wave modeling.

3.1. Broadband Body Wave Modeling

[20] We model the teleseismically recorded body wavesof the Wharton Basin earthquake following the sameprocedure described for the Enggano earthquake. We useboth P and S waves, unlike Robinson et al. [2001] who onlyattempt to model the S waves. The waveforms (Figure 6) aresignificantly shorter and less complex than those of theEnggano earthquake. We use a velocity model with 5 km ofwater and a 6 km crust over a half-space. We can model thebroadband body waves with a similar non-double-couplemechanism to the Harvard CMT. We then attempt P waveinversions constraining the mechanism to be a pure doublecouple. Like the Enggano earthquake the P wave firstmotions tightly constrain the first subevent to be strike-slip(Figure 6), with negative polarity onsets in the NW and SEand positive in the other two quadrants. Such a mechanismcan fit many of the seismograms but it cannot fit the Pwaves in the NW, nor the SH waves in the ENE. The Pwave onsets in the NW are negative polarity but a larger,positive polarity phase that arrives within 5 s after the onsetcannot be modeled with a pure strike-slip mechanism. Thisarrival is too early and too large to be any kind of waterreverberation. We thus try adding a second subevent andinvert for two subevents with different mechanisms, varyingthe time delay between them from 0 to 20 s.

Figure 5. (continued)

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[21] Our preferred model is shown in Figure 6. The sumof the two subevents is close to the Harvard CMT mech-anism, including the non-double-couple component. Againwe estimate the uncertainties in our solution using thecalculated RMS error, the visual fit to the waveforms, andthe consistency of results of inversions with different start-ing parameters and constraints. The mechanism of the firstsubevent is tightly constrained by the first motions to ±2�.Varying the depth of the first subevent by ±5 km increasesthe RMS error by 4%. The directivity is toward the northand the visual fit to the waveforms decreases significantly(the RMS error increases by 20%) if the solution is fixed asa point source. Thus the N-S plane is the likely fault plane.

The location of the Harvard catalogue centroid to the northof the hypocenter is consistent with this interpretation(Figure 7). The second subevent is less well resolved.Varying the depth by ±10 km gives only a 4% increase inRMS error. The mechanism is predominantly reverse, withthe null axis to the SW. Varying the strike by 10� increasesthe RMS error by 5%. The poorer constraints result from therelatively small size of this subevent (25% of the moment)and the fact that it occurs entirely within the ruptureduration of the larger, first subevent. Including directivityin the second subevent gives only a negligible improvementto the fit, and the orientation is not consistent from inversionto inversion. Thus we do not believe that the rupture

Figure 6. Broadband body wave modeling of the 18 June Wharton Basin earthquake. The layout andsymbols are the same as in Figure 3. A number of stations are omitted from the inversion to give a morebalanced azimuthal distribution. The strike, dip, and rake of the two mechanisms are 337�, 77�, and 17�and 206�, 40�, and 30�, respectively.

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direction can be resolved or used to distinguish the faultplane for this second subevent. The onset time and orienta-tion of the second subevent are constrained by the need tofit the P waves to the NW and SW. The orientation of thesecond subevent is consistent with that of the regional stressfield, which has maximum compression in the NW-SEdirection. It is also similar to that of a late aftershock thatoccurred in September 2001 (Figure 7), just to the NE of themain shock.

3.2. Relocation of Aftershocks

[22] The Wharton Basin earthquake had only 21 after-shocks located by the NEIC during 2000, of which twowere large enough to be included in the Harvard CMTcatalogue (Figure 7). The ISC and NEIC catalogues (since1970, complete above M � 5) contain no previous earth-quakes in the area of Figure 7. The depths of the aftershocksas reported by the NEIC are all fixed at 10 km. We attemptto relocate all of the aftershocks recorded by the NEICbetween June and September 2000, but had to discardpoorly recorded earthquakes to enable the inversion toconverge. Inverting for depth proved unstable, and so wefix all the depths at 10 km. The relocated epicenters of the

main shock and the twelve aftershocks with uncertainties<10 km are shown in Figure 7. The relocated epicenters areclose to their corresponding NEIC locations, but the NNW-SSE alignment is a little more tightly defined after reloca-tion. This agrees well with the preferred fault plane forsubevent 1. The other aftershocks could be interpreted asoff-fault events. Off-fault aftershocks are common follow-ing strike-slip earthquakes as a result of the static stresschanges induced by the main shock [e.g., Das and Scholz,1981]. Alternatively, those aftershocks to the NE could berelated to the second subevent. In September 2001, a M6reverse faulting earthquake occurred NE of the WhartonBasin earthquake, with a mechanism similar to that ofsubevent 2 (Figure 7).[23] The Wharton Basin earthquake occurred near to the

Investigator Fracture Zone but, although the fault orienta-tion is similar, the locations are too far to the west (�75 km)for the earthquake to have ruptured the ridge identified asthe Investigator Fracture Zone, and it must have occurred ona subparallel feature. Robinson et al. [2001] proposed that atrough just to the west of the Investigator Fracture Zone(fracture F, Figure 7) may have been the main fault plane.Accordingly, they shift the relocations of all the earthquakesin the sequence 50 km to the east. The orientation offracture F is more N-S than that of both the main shock faultplane and the alignment of the majority of the aftershockscontrary to their conclusion. In addition, a systematic shiftof 50 km is large, especially in a region where there is nostrong lateral heterogeneity in velocity structure as, forexample, in a subduction zone. Antolik et al. [2001] andSmith and Ekstrom [1996] performed location tests withearthquakes M5.5 and above and observed RMS misfits of12–14 km using the same velocity models as the NEIC andISC, and so it is unlikely that this entire set of earthquakescould be 50 km or more mislocated by the NEIC or ISC.Therefore, we do not prefer any particular feature in thebathymetry as the main shock fault plane.

3.3. Summary and Comparison With the Study ofRobinson et al. [2001]

[24] Our preferred model of the Wharton Basin earth-quake involves two subevents. The first (75% of themoment) is left-lateral strike-slip on a near vertical, almostNS plane, lasting just under 30 s. Our model for the firstsubevent is very similar in moment, slip orientation, anddimension to that of Robinson et al. [2001]. The rupturedirectivity and hypocenter, centroid and aftershock locationssupport this interpretation and also imply rupture to thenorth. The fault plane is subparallel to the InvestigatorFracture Zone, but too far away to have occurred on thisfracture. The second subevent is much smaller andrelatively poorly resolved as it occurred contemporaneouslywith the first subevent. It was predominantly reverse slip,with the null axis dipping to the SW and a fault planestriking either SW-NE or E-W. We cannot resolve whichnodal plane is the fault plane.[25] Our preferred model for the second subevent differs

significantly from that of Robinson et al. [2001]. The mainreason for the difference between the two models is that wemodel both the P and S waves, whereas they only considerthe S waves. To compare the two models, we fix ourinversion parameters to their preferred solution and obtain

Figure 7. The geometry of the Wharton Basin earthquakeand aftershock sequence. The epicenters of all earthquakesfrom June to October 2000 located by the NEIC (plus signs)and those relocated (stars), i.e., each star has a correspond-ing plus. The main shock epicenter is shown by the largestof each of these symbols, with a thickened white outline.The mechanisms of the main shock and two aftershocks inthe Harvard CMT catalogue are shown joined to theircentroid and relocated epicenters, respectively. The twosubevent mechanisms derived from the body wave model-ing of the main shock are also joined to the Harvard CMTcentroid location (white circle). A late aftershock occurredon 7 September 2001, and the Harvard CMT solution isplotted joined to the NEIC epicenter. The InvestigatorFracture Zone (IFZ) and Fracture Zone F [Robinson et al.,2001] are identified with dashed lines. Bathymetry isshaded as in Figure 1.

ABERCROMBIE ET AL.: THE JUNE 2000 INDIAN OCEAN EARTHQUAKES ESE 6 - 11

the model shown in Figure 8. We use their mechanisms,depths, rupture velocities, and temporal and spatial separa-tion between the two subevents. We invert only for momentrate functions. The moment of the first subevent is close totheir published model, but that of the second is smaller. Ourpreferred model is able to fit both P and S waves from allquadrants (Figure 6). The model of Robinson et al. [2001]does not fit the S waves any better than ours (Figure 8). Thepoor fit to stations SSE and YSS (Figure 8) would perhapsbe improved if the second subevent were larger, as in theirpublished model. The location of these stations near the SHnode for both subevent mechanisms suggests that the poorfit is more likely the result of incorrect model sourceorientation. More importantly, the model of Robinson et al.[2001] does not fit the P waves to the NW and SW.Robinson et al. [2001] claimed that the P waves cannot bemodeled because of reverberations in the water layer. We donot see any indication of such reverberations in the part ofthe waveforms that we model. Modeling strike-slip earth-quakes in the Atlantic Ocean with and without a water layermade a negligible difference to the shape of the synthetics,suggesting that such reverberations are small for strike-slipmechanisms in relatively flat-lying strata [Abercrombie andEkstrom, 2001]. The P phase that Robinson et al. [2001]failed to model is the phase with the largest amplituderecorded at stations to the NWand starts only about 5 s afterthe onset. It is the occurrence of this phase that correspondsto the onset of our second subevent and requires it to have aprimarily reverse slip orientation. Simply increasing themoment of the second subevent in Robinson et al.’s modelcannot produce this phase, even if it did start significantlybefore the 12 s delay time that they propose. The relativelypoor fit of the rupture geometry proposed by Robinson et al.[2001] to the larger data set used here suggests that theirresults for the location, timing and orientation of slip are notwell resolved.

4. Results and Discussion

4.1. Seismotectonics

[26] Our preferred interpretation of both the 4 and 18 June2000 earthquakes is that they primarily involve left-lateralslip on N-S trending near-vertical faults (Figure 1). Althoughthe faults are not well known, the Wharton Basin is cut bymany parallel N-S fossil fracture zones. The 18 June earth-quake is close to (though not on) the Investigator FractureZone, which is the most prominent fracture zone east of the90�ERidge.Newcomb andMcCann [1987] cite evidence forfracture zones at 102.5� and 104.5�, based on age offsetsand sediment filled troughs. The westerly of these two is alikely candidate for the fault that ruptured in the 4 JuneEnggano earthquake. Previous seismicity east of the 90�ERidge suggests that some of these fossil fracture zones arecurrently active, and DePlus et al. [1998] find clearevidence for reactivation of N-S fractures in a left-lateralsense on seismic profiles through the Wharton Basin. Thetwo June earthquakes thus confirm that these N-S fracturesare active and capable of moving in large earthquakes.[27] The direction of strike-slip motion in both of these

earthquakes is consistent with recent deformation models ofthe Indian Ocean (Figure 3) [Royer and Gordon, 1997]. Thelocation of the Wharton Basin earthquake confirms that the

diffuse boundary extends over a wider region than initiallydefined by Gordon et al. [1990]. The orientation of theearthquakes is also consistent with the present-day orienta-tion of stress and strain in this part of the India–Australiaplate [Cloetingh and Wortel, 1986; Kreemer et al., 2000].The occurrence of the 4 June earthquake implies that thisdeviatoric stress field exists in the oceanic lithosphere thathas been subducted to a depth of at least 50 km. Thisinterpretation is consistent with a study of stress in theSunda arc by Slancova et al. [2000]. They used CMTsolutions of earthquakes prior to 1997 to identify domainswith similar stress orientations within the subducting slaband observed a domain with compression parallel to thetrench in the depth range 25–225 km along the length ofSumatra. They suggested that this might represent stressesin the oceanic plate continuing into the subduction zone.The continuation of the deviatoric stress field in the oceaniclithosphere into the subducting slab may also explain theseismicity on the subducting Investigator Fracture Zone,recorded by a local network [Fauzi et al., 1996]. A band ofearthquakes, consistent with faulting on a near vertical N-Splane follows the subducting Investigator fracture zone to adepth of about 150 km. Fauzi et al. [1996] noted that suchseismicity is unusual as most subducting fracture zones areaseismic. Their preferred explanation is that the seismicity isrelated to lateral compression in the slab caused by its localgeometry. They did not observe a tear in the slab, but wereunable to resolve a vertical offset in hypocenters of less than20 km.[28] The 4 June Enggano earthquake is quite similar to

the 1986 (Mw 7.7) Kermadec earthquake in many ways.Houston et al. [1993] showed that the Kermadec earthquakeinvolved primarily strike-slip faulting in the subducting slabat a depth of about 45 km. Thus it represents anotherexample of tearing of the subducting slab. The LouisvilleSeamount Chain is subducting just north of the location ofthe Kermadec earthquake and so stresses involved with thiscould have produced the tearing in 1986. The LouisvilleSeamount Chain follows the trace of the fossil Eltaninfracture zone, similar to our present understanding of the90�E Ridge [Herzer et al., 1987]. Thus it is possible thatboth the 4 June 2000 and the 1986 Kermadec earthquakesruptured fossil fracture zones in the subducting slab. The1986 Kermadec earthquake did not include reverse slip onthe plate interface, like that observed in the Engganoearthquake, but most of the aftershocks, including thelargest (Mw 6.1), were interplate thrusts [Houston et al.,1993]. The 1994 Kurile Islands earthquake (Mw 8.3) hasalso been interpreted as a tear within the subducting slab,but it did not trigger slip at the plate interface [Tanioka etal., 1995]. The more recent December 1999 Kodiak Islandearthquake was also intrapate with an oblique strike-slipmechanism, but aftershock relocations imply that it ruptureda steeply dipping fault plane parallel to the trench and wasdriven by downdip tension in the slab [Ratchkovski andHansen, 2001]. The Enggano earthquake is also consistentwith downdip tension in the slab, but our analysis impliesthat the trench-parallel nodal plane is not the fault plane.[29] The strike-slip subevent of the 4 June Enggano

earthquake triggered a second subevent that ruptured theplate interface to the SE. The hypocenter and strike-slipsubevent are near the southern extent of the great 1833

ESE 6 - 12 ABERCROMBIE ET AL.: THE JUNE 2000 INDIAN OCEAN EARTHQUAKES

rupture (Figure 3). The second subevent thus ruptured awayfrom the 1833 rupture zone. Simple static stress modelingsuggests that subevent 1 increased the stress on the plateinterface to the SE, in the region of the second subevent.The uncertainty in the distribution of slip in the subevents ofthe Enggano earthquake prevents us from calculating ameaningful model of the static stress change induced by

this earthquake in the rupture area of the 1833 earthquake.The slip on the plate interface would have increased theshear stress on the neighboring parts of the interface,including the 1833 rupture area. The aftershocks are all inan area of continued high activity over recent decades. Thedegree of seismic coupling in this region is not well known.Prawirodirdjo et al. [1997] show that Enggano Island is

Figure 8. Broadband body wave modeling of the 18 June Wharton Basin earthquake [after Robinson etal., 2001]. The layout and symbols are the same as in Figure 4, and the same stations are included in theinversion as in Figure 6. We fix the orientation, depth, and rupture velocities at their published values andinvert for the moment rate function. The strike, dip, and rake of the two mechanisms are 165�, 87�, and�2� and 75�, 82�, and 173�, respectively. The moment of the first subevent is close to the publishedmodel but that of the second subevent is smaller in our inversion.

ABERCROMBIE ET AL.: THE JUNE 2000 INDIAN OCEAN EARTHQUAKES ESE 6 - 13

moving with the forearc and the Indian plate, suggestingthat the subducting plate was coupled to the overlying platein this region prior to the 4 June earthquake. It will beinteresting to see what motions are observed in the regionafter the Enggano earthquake.

4.2. Source Properties

[30] The two earthquakes considered in this study areamong the largest and best-recorded oceanic events. Theycan thus provide information about the nature of fracture inthe oceanic lithosphere. They are both mainly intraplatestrike-slip earthquakes, although the Enggano earthquakehas a substantial subduction component. We can estimatethe rupture areas of the various subevents, using thedurations and rupture velocities determined from the bodywave modeling and also from the aftershock areas. We canthen calculate the stress drops, following the study ofEshelby [1957]. Both earthquakes occurred in �70 Malithosphere and so can be assumed to have rupture widths of30–40 km (corresponding to a brittle–ductile transition at600–700�C) [Wiens and Stein, 1983; Abercrombie andEkstrom, 2001]. The length of rupture in the strike-slipsubevent of the Enggano earthquake was �50–100 km, andin the Wharton Basin earthquake, �80–120 km. Theserupture areas imply stress drops of 10–30 and 5–10 MPafor the earthquakes, respectively. Robinson et al. [2001] alsoestimated a relatively high stress drop for the WhartonBasin earthquake. These values are similar to the stress dropof 16 MPa estimated by Antolik et al. [2000] for the1998 Antarctic earthquake (Mw 8.1) which was the largestintraplate strike-slip earthquake recorded to date. The 1986Kermadec earthquake appears to have been similar to theEnggano earthquake and that had a stress drop (dominated bythe strike-slip component) of 10–20 MPa [Houston et al.,1993]. In contrast, assuming a downdip width of 50–70 kmand length of 150–200 km for the second, subductionsubevent of the Enggano earthquake gives a lower stress dropof only�1 MPa. This is significantly lower, despite the widerange of possible rupture areas considered in the calculationsto reflect the resolution of the parameters. Choy andBoatwright [1995] performed a global study of earthquakeapparent stress, an alternative estimate of stress drop. Theyfound that oceanic intraplate and oceanic transform strike-slip earthquakes have the highest apparent stress and that theapparent stress of subduction zone earthquakes is signifi-cantly lower. Thus our observations are consistent with thisglobal trend. In the case of the Enggano earthquake, the lowstress drop could be related to the degree of seismic coupling.The plate interface may be slipping in part aseismically. Theoverall rupture area of the earthquake, therefore, couldinclude large regions that slipped aseismically during andfollowing the seismic slip. Estimating the stress drop ofsubduction earthquakes in such regions from the totaldimensions and the seismic radiation could therefore producean underestimate of the stress drop.[31] The number of aftershocks is also different for the

strike-slip and plate interface subevents of the Engganoearthquake. Typically continental earthquakes have one after-shock within an order of magnitude of the main shock, 10within 2 orders of magnitude and so on (Bath’s Law, firstcited by Richter [1958]). Subduction zone earthquakes havesimilarly active aftershock sequences. Both the strike-slip

subevents of the Enggano and Wharton Basin earthquakeshave few aftershocks. The Enggano earthquake had onlythree strike-slip aftershocks within 2 orders of magnitudeand the Wharton Basin had only one. Houston et al. [1993]noted only one strike-slip aftershock within 2 orders ofmagnitude of the 1986 Kermadec earthquake, and theDecember 1999 Kodiak Island earthquake had only twosuch aftershocks. Other examples of oceanic strike-slipevents with few aftershocks include the 1994 Romanchetransform earthquake (Mw 7.1) which had no recorded af-tershocks at all (NEIC and ISC, complete above aboutM4.5) and the 1994 (Mw 6.5) earthquake on the Blancotransform which had only two aftershocks within 2 orders ofmagnitude [Dziak et al., 2000]. The 1998 Antarcticearthquake (Mw 8.1) had only 26 aftershocks recorded bythe ISC, of which only two were within 2 orders ofmagnitude of the main shock. A recent global study byBoettcher and Jordan [2001] also finds that oceanictransform earthquakes typically have few aftershocks. Bycontrast, both the Enggano and Kermadec earthquakes hadlarge numbers of plate interface aftershocks, suggesting thatthe subduction zone earthquakes are more like continentalearthquakes in their ability to trigger aftershocks.[32] Both the stress drop and number of aftershocks are of

interest as they are first-order observations as to the differ-ences in earthquake rupture in continental and oceaniclithosphere. Kanamori and Anderson [1975] observed adifference in stress drop between interplate and intraplateearthquakes and interpreted it as a dependence on strain rateand recurrence time. It is hard to define whether theseismicity in the Indian Ocean is strictly intraplate orinterplate, however, since it is considered part of a broadzone of deformation.

5. Conclusions

[33] The 4 June Enggano and 18 June Wharton Basinearthquakes were both primarily left-lateral strike-slip on N-S, vertical faults, one within the subducting slab and theother beneath the Indian Ocean. Both earthquakes had asecond, reverse slip subevent. The purely strike-slip ruptureson conjugate planes proposed for the Wharton Basin earth-quake [Robinson et al., 2001] are inconsistent with the Pwaves and so can be discounted. In the Enggano earthquakethe reverse slip subevent was a plate interface thrustcomprising at least 35% of the total moment and rupturingto the SE away from the rupture area of the great 1833earthquake.[34] The occurrence of these earthquakes indicates that

fossil fracture zones in the Indian Ocean are being reac-tivated in the current stress field as strike-slip faults and thatthey can rupture in large earthquakes. The stress fieldcausing this reactivation is present to depths of at least 50km in the subducting slab.[35] Both strike-slip subevents have high stress drops and

few aftershocks, consistent with previous observations oftransform and oceanic strike-slip earthquakes. In contrastthe subduction thrust subevent had an active aftershocksequence and a lower stress drop, typical of subductionearthquakes. These differences between earthquakes inoceanic and continental lithosphere and at subduction zonesare consistent and reliable. They should assist us in under-

ESE 6 - 14 ABERCROMBIE ET AL.: THE JUNE 2000 INDIAN OCEAN EARTHQUAKES

standing the state of stress in plates and the role of rheologyin governing the nature of earthquake rupture.

[36] Acknowledgments. We thank E. Bergman and D. Wiens forproviding us with the relative relocation code and for assisting us in usingit. We are grateful to K. Felzer and J. Rice for discussions on the static stressfield induced by these earthquakes and for discussions with K. Sieh and J.Milsom on the seismotectonics of Sumatra and with R. Scrutton on theIndian Ocean. Comments by D. Kilb, J. Milsom, and the AE J. Freymuellerand two anonymous reviewers greatly improved this manuscript. We usethe GMT software for many of the figures [Wessel and Smith, 1991]. Thiswork was funded by NSF award EAR-98-05172.

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�����������������������R. E. Abercrombie, Department of Earth Sciences, Boston University,

685 Commonwealth Avenue, Boston, MA 02215, USA. (rea@bu.edu)M. Antolik and G. Ekstrom, Department of Earth and Planetary Sciences,

Harvard University, Cambridge, MA, USA.

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