Earthquakes and strain in subhorizontal slabs

Michael R. Brudzinski1

Department of Geology and Geophysics, University of Wisconsin-Madison, Madison, Wisconsin, USA

Wang-Ping ChenDepartment of Geology, University of Illinois at Urbana-Champaign, Urbana, Illinois, USA

Received 6 October 2004; revised 2 March 2005; accepted 17 March 2005; published 6 August 2005.

[1] Using an extensive database of fault plane solutions and precise locations ofhypocenters, we show that the classic patterns of downdip extension (DDE) or downdipcompression (DDC) in subduction zones deteriorate when the dip of the slab is less thanabout 20�. This result is depth-independent, demonstrated by both intermediate-focus(depths from 70 to 300 km) and deep-focus (depths greater than 300 km) earthquakes. Theabsence of pattern in seismic strain in subhorizontal slabs also occurs locally over scales ofabout 10 km, as evident from a detailed analysis of a large (Mw 7.1) earthquake sequencebeneath Fiji. Following the paradigm that a uniform strain of DDE/DDC results fromsinking of the cold, dense slab as it encounters resistance from the highly viscous mantleat depth, breakdown of DDE/DDC in subhorizontal slabs reflects waning negativebuoyancy (‘‘slab pull’’) in the downdip direction. Our results place a constraint on themagnitude of slab pull that is required to dominate over localized sources of stress and toalign seismic strain release in dipping slabs. Under the condition of a vanishing slabpull, eliminating the only obvious source of regional stress, the abundance of earthquakesin subhorizontal slabs indicates that a locally variable source of stress is both necessaryand sufficient to sustain the accumulation of elastic strain required to generateintermediate- and deep-focus seismicity. Evidence is growing that the process ofseismogenesis under high pressures, including localized sources of stress, is tied to thepresence of petrologic anomalies.

Citation: Brudzinski, M. R., and W.-P. Chen (2005), Earthquakes and strain in subhorizontal slabs, J. Geophys. Res., 110, B08303,

doi:10.1029/2004JB003470.

1. Introduction

[2] In the framework of plate tectonics, earthquakesdeeper than about 100 km occur in the cold interior ofsubducted lithosphere. Otherwise, deep earthquakes shouldnot be restricted to zones of recent convergence [e.g., Isacksand Molnar, 1969, 1971; Isacks et al., 1968; Wadati, 1927].The estimated limiting temperature for seismicity is approx-imately 600� to 800�C [Chen and Molnar, 1983; Molnar etal., 1979; Wiens and Stein, 1983]. As such, inclined zonesof seismicity (the Wadati-Benioff zones, WBZ), connectingshallow earthquakes near the trench with earthquakes deepin the mantle, represent traces of active, sinking lithosphere.On a global scale, for WBZ that reach depths of about300 km or less, fault plane solutions of earthquakes typi-cally show a pattern of downdip extension (DDE) where theaxes of maximum extension (T axes) closely follow the dipof slab as revealed by the distribution of seismicity. For

WBZ that penetrate toward the bottom of the mantletransition zone (down to depths of almost 700 km), faultplane solutions reveal a dominant pattern of downdipcompression (DDC) where the axes of maximum compres-sion (P axes) are subparallel to the dip of the slab (Figure 1a)[e.g., Isacks and Molnar, 1969, 1971].[3] The current paradigm is that these patterns of uniform

strain reflect internal deformation of subducted lithosphere[e.g., Apperson and Frohlich, 1987; Isacks and Molnar,1971; Vassiliou et al., 1984]. Specifically, the cold, strongslab acts as a stress guide: A uniform strain field develops inthe cold slab as it sinks because of negative buoyancy(‘‘slab pull’’) and encounters resistance from a highlyviscous mantle at greater depths. DDE and DDC arepatterns of internal deformation—manifestations of the slabeither being extended under its own weight or beingcompressed by mantle resistance against slab pull, respec-tively. In both cases, slab pull is the instigating factor thatprovides a natural driving mechanism for continual accu-mulation of elastic strain in the slab that is necessary tocause deep earthquakes.[4] However, acting alone, density differences that drive

subduction seem neither necessary nor sufficient to causedeep earthquakes. An observational basis for this reasoning

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 110, B08303, doi:10.1029/2004JB003470, 2005

1Also at Department of Geology, Miami University, Oxford, Ohio,USA.

Copyright 2005 by the American Geophysical Union.0148-0227/05/2004JB003470$09.00

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comes from the abundance of deep seismicity in zones that aresubhorizontal [e.g., Araujo and Suarez, 1994; Barazangi andIsacks, 1976; Chen and Brudzinski, 2001; Hamburger andIsacks, 1987; Hasegawa and Sacks, 1981; Okal and Kirby,1998]. Since the downdip component of slab pull is approx-imately proportional to the sine of slab dip, such a drivingforce is greatly diminished in subhorizontal slabs, indicatingthat slab pull is not necessary to cause earthquakes.[5] Insofar as earthquakes suddenly release accumulated

elastic strain, the occurrence of abundant seismicity insubhorizontal slabs indicate that a localized source of stressis both necessary and sufficient—in the absence of aregional stress field, locally generated stress and this stressalone must provide the source of strain accumulation. Thenecessity of localized stress also comes from experimentalrock mechanics. Unlike conditions associated with shallowearthquakes, confining pressure is so large at depths greaterthan approximately 100 km that the shear stress necessary todrive frictional sliding is expected to exceed the long-termductile strength of geologic materials [e.g., Scholz, 1990].This perspective suggests that regional tectonic stress (i.e.,slab pull) alone is insufficient to initiate deep earthquakes;some additional mechanism(s) are needed to trigger thesudden release of stored elastic strain.[6] Prominent patterns of DDE or DDC in most subduc-

tion zones indicate that regional, tectonic stress dominatesin these cases, making it difficult to isolate factors thattrigger deep earthquakes. In this study, we seek limitingconditions under which slabs cease to exhibit a prevalenceof DDE or DDC, and we investigate the interplay ofmultiple factors in seismogenesis.

2. Data and Analysis

[7] We survey all major subduction zones to identifyregions where the configuration of slab is well definedand approximately planar. In doing so, we follow criteria

Figure 1. Principal axes of seismic strain compared withorientation of subducted slabs. (a) Traditional analysis ofplotting the plunge of P/T axes as a function of slab dip. Thedashed line along the diagonal for a perfect correlationbetween the two parameters is close to the result of leastsquares regression (r = 0.96, standard deviation 7.7�) whenV1 is excluded. However, this plot does not includeimportant information regarding the scatter of P/T axes.(b) Standard deviation of Fisher distribution for P/T axes(sP/T) as a function of slab dip. Notice the stepwise increasein sP/T when slab dip is less than 15�–20�. (c) Plotsummarizing the relative orientation between the slab andthe mean of the P/T axes (a) for each of the 25 cases. Underequal-area projection, the downdip direction (d) is at thecenter of the plot, with the equatorial plane of the plotmarking the slab (dotted line). Difference between a and d isexpressed by angles y and q, where y represents thedeviation of a normal to the slab and q reflects azimuthaldifference between a and d measured in the slab plane(insets). To avoid clutter and to account for the large rangein slab dip, standard deviations of P/T axes are plottedseparately in Figure 1b. See color version of this figure inthe HTML.

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established in previous work [Apperson and Frohlich,1987; Burbach and Frohlich, 1986] to avoid regions withcomplex geometry or multiple zones of seismicity. Weimprove the estimate of slab geometry by using leastsquares fit to the distribution of precisely relocated hypo-centers of large- to moderate-sized (mb � 5) earthquakesbetween 1964 and 2002 [Engdahl et al., 1998; Engdahland Villasenor, 2002]. In regions of particular interest, wesupplement this data set with precisely determined focaldepths that take full account of regional variations inseismic wave speeds [Brudzinski and Chen, 2000, 2003](Appendices A and B).1 To avoid possible bias fromearthquakes along the plate interface, we restrict ouranalysis to earthquakes deeper than 70 km. In most cases,the resulting distribution of relocated hypocenters is denseenough to distinguish neighboring planar segments of theWBZ, if differences in their dips exceed 10�.[8] We then examine orientations of principal strain in

each segment using a comprehensive compilation of faultplane solutions mainly determined from waveform inver-sion [Dziewonski et al., 1981; Huang et al., 1997; Lay et al.,1989; Skomurski et al., 1999; Tavera and Buforn, 2001;Wiens et al., 1993]. We also used waveform inversion tostudy a large (Mw 7.1) earthquake sequence within subhor-izontal slab beneath Fiji (Appendix A in the auxiliarymaterial), and included a supplemental compilation ofwell-determined fault plane solutions from first motions[Billington, 1980; Isacks et al., 1969; Okal and Kirby,1998; Richter, 1979].

[9] We characterize the alignment and the clustering ofP/T axes by their mean and standard deviation (sP/T) usingFisher statistics appropriate for spherical data [Fisher et al.,1987]. To ensure statistically meaningful results, we excludesegments that have fewer than ten fault plane solutions. Wethen evaluate the prevalence of DDC or DDE by comparingstatistical characteristics of P/T axes with the geometry ofslab (Figure 1). Such a quantitative analysis reveals impor-tant new features that are now attainable because of the largesize of our final data set, including about 1,000 fault planesolutions and 2,500 relocated hypocenters.[10] In all, we identify 25 planar segments that cover ten

major subduction zones: Izu-Bonin, Japan, Java, Kurile-Kamchatka, New-Britain, Philippines, Ryukyu, SouthAmerica, Tonga, and Vanuatu (Table 1). Among them,Izu-Bonin did not appear in an earlier compilation ofApperson and Frohlich [1987], but now there are ampledata to include this zone in our analysis.

3. Global Survey

[11] Figure 1a summarizes the result of traditional anal-ysis between mean plunges of P/T axes and dip angles ofslabs. Except for case V1, the outcome seems to simplycorroborate classic patterns of DDC/DDE, with tight corre-lation between the two parameters (r = 0.96, standarddeviation 7.7�). However, such an analysis is incompleteand even misleading in two regards.[12] First, Figure 1a considers only the plunge of the

mean P/T axis, so the result obscures azimuthal differencebetween the direction of true slab dip and azimuth of themean P/T axis. An alternative is to analyze the orientation ofthe mean P/T axis, a, with respect to the downdip direction

Table 1. Summary of Global Survey

Region SegmentRange of Depths,

kmDip of Slab,

degNumber ofHypocenters

Plunge of StrainAxes, deg

Number of FaultPlane Solutions P or T Axis

Nazca N1 72–190 9 ± 4 111 10 ± 24 50 TNazca N1 546–643 75 ± 15 22 79 ± 10 14 PNazca N2 70–134 5 ± 4 80 �4 ± 29 34 TNazca N3 70–281 25 ± 5 280 25 ± 17 136 TNazca N4 71–182 7 ± 4 97 �2 ± 29 37 TNazca N4 569–618 54 ± 20 15 50 ± 14 11 PNazca N5 72–172 24 ± 15 40 39 ± 15 22 TTonga T1 330–628 13 ± 9 37 29 ± 27 21 PTonga T2 74–286 49 ± 4 110 54 ± 19 57 PTonga T2 479–664 52 ± 5 463 53 ± 15 212 PVanuatu V1 593–664 3 ± 4 44 81 ± 30 17 PVanuatu V2 70–288 69 ± 8 203 80 ± 19 103 TVanuatu V3 75–266 47 ± 8 98 57 ± 18 43 TVanuatu V4 76–291 61 ± 8 80 54 ± 20 29 TNew-Britain B1 82–197 64 ± 8 26 67 ± 18 12 PNew-Britain B2 102–207 72 ± 8 19 73 ± 16 11 PIzu-Bonin I1 306–432 35 ± 6 23 35 ± 14 10 PIzu-Bonin I2 338–533 49 ± 10 59 47 ± 16 40 PIzu-Bonin I3 330–565 81 ± 5 31 67 ± 18 30 PJapan J1 333–581 23 ± 7 37 30 ± 12 24 PKurile-Kamchatka K1 399–650 61 ± 6 19 61 ± 14 16 PKurile-Kamchatka K2 320–640 43 ± 7 15 52 ± 14 10 PPhilippine P1 300–675 53 ± 7 85 58 ± 19 30 PRyukyu R1 86–208 48 ± 9 20 54 ± 13 12 TRyukyu R2 79–238 50 ± 5 70 50 ± 20 24 PSumatra-Sunda S1 70–213 39 ± 6 63 42 ± 20 20 TSumatra-Sunda S2 402–645 87 ± 4 70 82 ± 18 39 PSumatra-Sunda S3 80–281 62 ± 4 180 73 ± 18 66 T

1Auxiliary material is available at ftp://ftp.agu.org/apend/jb/2004JB003470.

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of each slab, d, in three dimensions. Figure 1c shows theresults in an equal-area projection, where d and each planarsegment of slab are projected as the center and the equato-rial plane, respectively. Difference between a and d is nowreadily expressed by two parameters, angles y and q, wherey represents the deviation of a normal to the slab and qreflects azimuthal difference, measured in the slab plane,between a and d (Figure 1c, insets).[13] It becomes apparent that both cases V1 and T1 are

clear exceptions to the classic patterns of DDC/DDE(Figure 1c); yet case T1 entirely escapes detection inFigure 1a. The reason is twofold: (1) the large angulardeviation between a and d (close to 60�) lies mainly in q,so a is approximately in plane with the slab; and (2) the dipangle of the slab is small (�13�). As such, the apparentangle of slab dip, measured along the azimuth of a, differsonly slightly from the true angle of dip being considered inFigure 1a.[14] Second, Figures 1a and 1c do not consider the full

statistical characteristics of the large amount of data nowavailable. In particular, a new feature becomes salientwhen the scatter of P/T axes, not just the mean of P/Taxes, is compared with slab dips that span the whole rangefrom 3� to 87� (Figure 1b). For moderate- and steep-dipping slabs (>20�), the scatter in P/T axes is small, withan average sP/T of about 17�, and this value remainsconstant over a very large range of slab dip (20� to 90�).In contrast, when slab dip drops below a threshold of 15�–20�, there is a sharp increase in the scatter of P/T axes,reaching up to 30� in sP/T (Figure 1b). Among a total of25 cases, the pattern of DDE/DDC breaks down in thegroup of five subhorizontal slabs (N1, N2, N4, T1 andV1). The anomalous nature of segments N1, N2 and N4would escape detection, if the large scatter of P/T axes isnot considered (Figures 1a and 1c).[15] In general, configurations of slabs are well deter-

mined by seismicity, with standard deviations of less than10� in slab dips. So the threshold relationship between slabdip and sP/T is not an artifact due to uncertainties in the dip ofsubhorizontal slabs. For instance, in case V1 (Vanuatu), sP/Tfor the mean plunge of P axes is 30�, over seven times largerthan the standard deviation of slab dip (4�). In subhorizontalslabs (dip < 15�–20�), apparently P/T axes do not cluster andthe pattern of DDC/DDE strain breaks down (Figure 1b).Notice that this observation is independent of depth, holdingtrue for both intermediate- and deep-focus earthquakes,indicating a pressure-independent process that is controlledmainly by geometry.[16] The onset of clear DDE/DDC when the slab dip

reaches above 20� corroborates the paradigm that slab pullis the main cause of uniform strain in dipping slabs [e.g.,Apperson and Frohlich, 1987; Isacks and Molnar, 1969,1971; Vassiliou et al., 1984]. More important, this observa-tion places a quantitative constraint on the magnitude ofslab pull that can overcome localized sources of stress andalign seismic strain release in dipping slabs: at least onethird (sine of 20�) of the slab pull is required along thedowndip direction.[17] Meanwhile, abundant seismicity in subhorizontal

slabs indicates that a downdip component of slab pull,the driving force of subduction, is not a necessary condi-tion for deep earthquakes. In the absence of a regional

stress field, localized sources of stress must be able togenerate large intermediate- and deep-focus earthquakes bythemselves, revealing that localized processes are necessaryand sufficient for seismogenesis. In the next two sections,we discuss the five cases of subhorizontal slabs in somedetail. When applicable, we also contrast subhorizontalslabs with neighboring segments of the WBZ whereDDE/DDC prevails.

4. Subhorizontal Slabs at Intermediate Depths:Peru and Argentina

[18] There are three well-known segments of subhori-zontal slab along the Nazca subduction zone beneathSouth America, two in Peru and one in northwesternArgentina and Chile (boxes N1, N2 and N4 in Figure 2)[e.g., Araujo and Suarez, 1994; Barazangi and Isacks,1976; Chen et al., 2001; Hasegawa and Sacks, 1981;Smalley and Isacks, 1987]. Even the smallest of the threecovers a significant area of approximately 400 � 400 km2

(N2, in southern Peru). In all three regions, a subhorizontalWBZ extends landward for several hundred kilometers butremains at depths of less than 150 km or so. These regionsare generally devoid of Holocene volcanoes, in markedcontrast to the prevalent, active magmatic arc aboveneighboring regions of moderate slab dip (Figure 2)[e.g., Barazangi and Isacks, 1976; Pilger, 1981; Sacks,1977].[19] Pronounced lateral variations in slab dip along the

Nazca subduction zone provide an ideal, natural setting forstudying the relationship between seismic strain and geom-etry of the slab. In Figures 2 and 3, we summarize resultsfrom this subduction zone in map view and cross sections,respectively. We show the actual distribution of P/T axes ineach cross section and summarize the distribution by itsmean and 2sP/T (Figure 3 insets, solid curves). For com-parison, we also plot the downdip direction (d, diamonds)and regions up to 33� away from d (Figure 3 insets, dashedcurves). The value of 33� is simply the typical value of2sP/T from the global survey of slabs with dips greaterthan 20� (Figure 1b). In other words, we consider a P/Taxis to be consistent with the pattern of DDC/DDE if theaxis plots within the bounds of dashed curves. Using thiscriterion, we also color-coded fault plane solutions tovisually identify earthquakes that are exceptions to DDC/DDE (Figure 3, red/black ‘‘beach balls’’).[20] Marked increases in the scatter of P/T axes associated

with subhorizontal slabs are apparent when one comparessegment N4 with its neighbors, segments N3 and N5(Figures 3c–3e). For segment N4, T axes loosely lie withinthe plane of the slab but do not cluster around the downdipdirection (Figure 3d). In contrast, T axes in segment N3align with the downdip direction of slab so tightly that thelimit of 2sP/T for the T axes (solid curves) coincidesremarkably well with the nominal definition of DDE basedon global survey (Figure 3c, dashed curves). A similar,albeit less dramatic contrast exists between neighboringsegments N4 and N5 which has a small number of availablefault plane solutions (22). Notice that the realignment of Taxes into the pattern of DDE occurs over small distances ofless than 100 km away from the subhorizontal segment N4(Figure 2)).

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Figure 2. Map summarizing results of five cases studied in western South America. Circles showepicenters of large- to moderate-sized (mb � 5) earthquakes between 1964 and 2002 (color codedaccording to focal depth), with triangles marking Holocene volcanoes [Simkin and Siebert, 1994].Contours are depths of the Wadati-Benioff zone (WBZ), defined by the inclined distribution ofearthquake hypocenters. Boxes indicate locations of cross sections shown in Figure 3. For each box, anequal-area projection of the lower hemisphere compares the plane of the slab (dashed arc, with a thickbar marking the downdip direction, d), the mean orientation of the P/T axes (a, diamond), and bounds of2sP/T (thin solid curves). A true pattern of DDE/DDC is present only if the scatter in P/T axes is small andvectors a and d are subparallel (such as cases N3 and N5). When abundant seismicity is present at depthsgreater than 300 km (cases N1 and N4), a separate plot is shown (marked ‘‘Deep’’). See color version ofthis figure in the HTML.

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[21] Downdip along segments N1 and N4, there is anatural break in seismicity between depths of about 200and 500 km (Figures 3a and 3d). In both segments, there isconsiderable scatter in T axes associated with the subhor-izontal WBZ above this break. However, the WBZ dipssteeply (>50�) below the break and the P axes are wellaligned into a clear pattern of DDC. Thus the relationshipbetween scatter of P/T axes and slab dip holds true forvariations in slab geometry both in the direction of dip andalong strike of the subduction zone.[22] These observations seem to work against the hypoth-

esis that fault reactivation is a necessary mechanism fordeep earthquakes [Christova and Scholz, 2003]. To accountfor the uniform pattern of DDE/DDC in dipping slabs, thehypothesis assumes that pre-existing faults in the oceaniclithosphere, originally formed with a regular orientation atthe outer rise prior to subduction, are reactivated duringdeep earthquakes [e.g., Jiao et al., 2000; Kirby et al.,1996a; Savage, 1969]. If fault reactivation is necessary forseismogenesis, wide scatter of fault plane solutions in flat-lying sections of the Nazca slab requires an incoherentnetwork of pre-existing faults (segments N1, N2 and N4).Yet highly coherent pre-existing faults must simultaneouslyexist in neighboring segments of slab that dip at about 25�(such as segments N3 and N5 (Figure 2)). Such a scenarioseems unlikely because it implies rapid, lateral variationsof tectonic stress along the outer rise, inconsistent withregular fault plane solutions seaward of the trench [e.g.,Christensen and Ruff, 1988; Forsyth, 1982]. Were we toaccept such an ad hoc scenario, it still fails to explaincontrasting patterns of poor alignment in T axes at shallow

Figure 3. Cross sections along western South America,showing how downdip clustering of P/T axes deterioratesappreciably when the dip of slab is less than about 20�(Figures 3a, 3b, and 3d at intermediate depths). The fivecross sections, N1–N5, are arranged by decreasing latitudes(from north to south), each taken along the direction of slabdip as determined by a least squares fit of seismicity to aplanar surface. In each plot, fault plane solutions (beachballs with size proportional to Mw) are equal-area projec-tions of the northern hemispheres of the focal spheres,where shaded quadrants represent compressional firstmotions, with P and T axes plotted as solid and opencircles, respectively. Crosses show hypocenters with nofault plane solution available. Using the same projection asthe fault plane solutions, each circular inset showsorientations of P/T axes for all fault plane solutionsenclosed in the associated box. Also shown are the meanof P/T axes (diamonds) and bounds of 2sP/T (solid curves);the latter is compared with a nominal criterion for DDE/DDC (dashed curves): the region within 33� from thedowndip direction of the slab, corresponding to the typicalvalue of 2sP/T from global survey of slabs with dips greaterthan 20� (Figure 1b). For events whose P/T axes fall outsidethe nominal criterion, we highlight the corresponding P/Taxes and fault plane solutions in red/black so that exceptionsto DDE/DDC are easily seen. Positions of Holocenevolcanoes and the trench axis are plotted as triangles andinverted triangles, respectively. See color version of thisfigure in the HTML.

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depths versus clear downdip compression at greater depths(Figures 3a and 3d). So fault reactivation is likely not anecessary ingredient for generating earthquakes at depth.

5. Subhorizontal Slabs in the Transition Zoneof the Mantle: Tonga and Vanuatu

[23] The southwestern Pacific is a region of extremelyfast subduction where abundant seismicity occurs deep inthe transition zone of the mantle. A case in point is thenorthern Fiji basin, bounded to the east and west by theTonga and Vanuatu (New Hebrides) trenches, respectively(Figure 4). Subduction of old, cold slab exceeds 200 mm/yralong most of the Tonga trench, resulting in half of theworld’s deep earthquakes [e.g., Frohlich, 1989].[24] Beneath the northern Fiji basin, there are two deep

swaths of ‘‘outboard’’ earthquakes (Figure 4, circles): sub-horizontal bands of earthquakes that extend up to onethousand kilometers away from the tip of well-definedWBZ (Figure 4, contours) [e.g., Barazangi et al., 1973;Chen and Brudzinski, 2001; Hamburger and Isacks, 1987;Okal and Kirby, 1998]. Beneath the island of Fiji, anextensive patch of outboard earthquakes partially overliesthe WBZ associated with the Tonga trench to the east. Theother swath of outboard earthquakes occurs to the east of theVanuatu WBZ which ends at a depth of only about 350 km(Figure 4).

[25] Since neither swath of outboard earthquakes hasobvious connection with active subduction [e.g., Barazangiet al., 1973; Chen and Brudzinski, 2001; Hamburger andIsacks, 1987; Okal, 2001; Okal and Kirby, 1998], thedowndip direction of slab is undefined. Nonetheless, it isapparent that there is no preferred orientation in fault planesolutions (Figure 5) [e.g., Chen and Brudzinski, 2001; Okaland Kirby, 1998]. A lack of pattern in fault plane solutionsin segments T1 and V1 is in stark contrast to well-aligned Paxes along the downdip direction in the steep-dipping WBZalong the Tonga and Vanuatu subduction zones (Figure 4)[e.g., Apperson and Frohlich, 1987; Chen and Brudzinski,2001]. Except for obvious differences in focal depths,spatially varying seismic strain in segments T1 and V1resembles that in subhorizontal slabs of the Nazca subduc-tion zone (particularly segment N2 (Figure 3b)).[26] Beneath the northern Fiji basin, large outboard earth-

quakes are common. For instance, a large event (Mw 7.1)recently occurred on 15 July 2004 near 17.73�S and178.51�Wat a depth of about 580 km (Harvard CMT catalog,http://www.seismology.harvard.edu/CMTsearch.html).A particularly interesting sequence of outboard events beganon 26 May 1986. The large main shock (Mw 7.1) comprisesthree distinct subevents, followed by moderate-sized after-shocks that continued for about two weeks afterward(Appendix A and Tables A1 andA2 in the auxiliary material).A detailed analysis of this earthquake sequence shows that

Figure 4. Map summarizing results in the southwestern Pacific. The general layout is the same as thatin Figure 2. In this case, only hypocenters of outboard earthquakes are plotted, all of which are deeperthan 300 km and extend up to 1,000 km from the trenches. Contours mark depths to the active WBZ,where seismic strain is predominantly downdip compression [Chen and Brudzinski, 2001]. See colorversion of this figure in the HTML.

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spatially variable seismic strain in subhorizontal slabs occurboth in regional (100 to 1000 km) and local scales (about10 km), as the largest subevents and aftershocks all havedistinct fault plane solutions (Figures 5 and A1 (seeauxiliary material)) [Skomurski et al., 1999]. Consideringwaning slab pull in subhorizontal slabs, these observa-tions reinforce the notion that local sources of stress areinvolved in generating large earthquakes at great depths.

6. Discussion

[27] The call for locally variable stress in the making ofdeep earthquakes is consistent with the longstandingrealization that a localized perturbation in stress is neces-sary to overcome high confining pressure that wouldotherwise prohibit sudden release of strain in the form ofearthquakes [e.g., Scholz, 1990]. Mechanisms that havebeen demonstrated by laboratory experiments to inducesudden release of stored elastic strain involve local pro-duction of small amounts of a new phase. Such a phasebehaves mechanically as a low-viscosity fluid, either a truefluid such as water or a pseudofluid such as a superplasticaggregate of particles [e.g., Green and Marone, 2002]. Theexact mechanism that produces deep earthquakes, whichmust include a localized source of stress, has been thesubject of continual research (see reviews by Green andHouston [1995] and Frohlich [2005]). Lately there isgrowing evidence that seismogenesis is tied to petrologicanomalies that accompany the source region of deepearthquakes.

6.1. Deep-Focus Earthquakes in the TransitionZone of the Mantle

[28] For the large patch of outboard earthquakes beneathFiji, there is a comprehensive set of in situ evidence for acorresponding petrologic anomaly. There are two keypoints. First, the source region of outboard earthquakes isseismically anisotropic [Chen and Brudzinski, 2003], aphenomenon rarely observed in the transition zone of themantle [Savage, 1999]. Second, high VP and VS, expected toaccompany low temperature in the source region of seis-micity, are indeed present in peripheral regions surroundingthe outboard earthquakes [Brudzinski and Chen, 2000,2003]. However, within the anisotropic zone of earthquakesitself, VP and VS are both low, requiring a petrologicanomaly that counteracts the effect of low temperature.[29] Within the transition zone of the mantle, candidates

for the petrologic anomaly that will simultaneously reduceVP and VS, produce seismic anisotropy, and trigger earth-quakes include metastable phases such as olivine (a phase),volatiles in the form of hydrous phases, or partial meltinduced by volatiles [e.g., Faul et al., 1994; Green andBurnley, 1989; Green and Houston, 1995; Kirby, 1995;Kirby et al., 1991; Meade and Jeanloz, 1991]. To sustainseismic anisotropy, shape-preferred orientation of hydrousphases or melt pockets requires a regional stress pattern andtherefore is inconsistent with incoherent orientations of faultplane solution in segments T1 and V1. In contrast, lattice-preferred orientation of metastable olivine remains a viablecandidate as petrofabric of olivine grains is a common causeof strong seismic anisotropy in the mantle lithosphere

Figure 5. Cross sections summarizing results of subhorizontally distributed outboard earthquakes nearVanuatu (V1) and Tonga (T1). In both cases, P axes of outboard earthquakes show large scatter. Layout isthe same as that of Figure 3, with fault plane solutions that are associated with the large sequence of26 May 1986 marked (see Appendix A in the auxiliary material). See color version of this figure inthe HTML.

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[Savage, 1999]. Furthermore, seismic anisotropy is often theresult of relic petrofabric, so its association with outboardearthquakes does not require the prevalence of present-dayregional strain [Chen and Brudzinski, 2003]. Two additionalconsiderations also favor the presence of metastable olivine.First, if partial melt is a primary component, the amount ofreduction in VS should be much greater than that in VP,inconsistent with comparable values of 2–3% observed forboth [Brudzinski and Chen, 2003]. Second, so much low-density hydrous phases would be required to reduce VP andVS by 2–3% that the slab would become buoyant through-out the mantle and unable to subduct [Brudzinski and Chen,2000].[30] A subtle but intriguing point is that cold, metastable

olivine in a subducting slab is buoyant in the transition zonebut not in the upper mantle. So as a cold slab descends intothe transition zone and leaves the stability field of a phase,metastable olivine would make the leading edge of slabbuoyant [e.g., Bina, 1996; Marton et al., 1999]. This addedbuoyancy would have two effects on the subducting slab.First, buoyancy from metastable olivine resists subductionin the transition zone, adding to the effect of viscousresistance from the surrounding mantle that promotesDDC. Second, metastable olivine probably influences thetrajectory of deep penetrating slab toward a flat-lyingposture, because the buoyant tip will rise but never floatabove the 410-km discontinuity where warmer olivine ofthe ambient mantle is also in a phase. So buoyancy of theslab is a self-limiting process confined to the transitionzone, offering a natural explanation for the presence oflarge-scale, subhorizontal slabs in the transition zone.Moreover, a buoyant slab favors retention of subductedmaterial in the transition zone and hinders slab penetrationinto the lower mantle, a key issue pertaining to the currentdebate of mantle convection and mixing [Chen andBrudzinski, 2001].[31] Numerical modeling showed that because of a large

reduction in volume, high differential stresses are producedin slabs when metastable olivine transforms to ringwoodite(g or spinel phase) [e.g., Devaux et al., 2000; Goto et al.,1985; Kirby et al., 1996b]. On a local scale, such transfor-mation produces a superplastic aggregate with localizedstress perturbations and causes rapid faulting in laboratoryexperiments (transformational faulting), emulating deepseismicity [e.g., Green and Burnley, 1989; Kirby, 1987].Thus a localized source of stress seems to be part of theseismogenic process due to a petrologic anomaly which byitself is both necessary and sufficient to generate deep-focusearthquakes. We suspect that in the absence of a dominant,regional stress field such as slab pull, locally variablestresses during transformation of isolated pockets of meta-stable olivine may be a key cause of spatially varyingseismic strain in flat-lying slabs [e.g., Kirby et al., 1991;Okal and Kirby, 1998].

6.2. Intermediate-Focus Earthquakes in theUpper Mantle

[32] Recent thermal-petrologic models suggest thatwater may be important in the earthquake-generatingprocess at intermediate depths, as hydrous minerals arepredicted to exist where intermediate-depth seismicity isabundant [Hacker et al., 2003]. However, in situ con-

straints on the source of seismicity have been sparse.Recently, Zhang et al. [2004] reported that zones ofanomalous Poisson ratios are associated with the doubleplanes of seismicity beneath Japan. They interpret that theanomalous regions correspond to dehydration of alteredoceanic crust and mantle lithosphere during subduction.These reactions, involving the release of water during thetransformation of mafic crust to eclogite and dehydrationof serpentine, may release as much as 2.5 wt% of water atdepths of 60–140 km.[33] Raised pore pressures due to dehydration and ensu-

ing embrittlement has long been proposed as a triggeringmechanism for earthquakes [e.g., Green and Houston,1995; Kirby, 1995; Meade and Jeanloz, 1991; Raleighand Paterson, 1965]. Recent laboratory experiments onthe dehydration of antigorite (of the serpentine group) upto 9 GPa have observed acoustic emissions and deforma-tion features that appear to be caused by dehydrationembrittlement under high–pore fluid pressures [Dobsonet al., 2002; Jung et al., 2004]. Furthermore, Jung et al.[2004] showed evidence for the separation of fluid fromsolid residual during the dehydration of antigorite atpressures of up to 6 GPa. This result seems to alleviatean obstacle for the embrittlement hypothesis below depthsof about 60–120 km, where serpentine goes through anegative volume change that would otherwise inhibitembrittlement.[34] At the moment, it is difficult to relate seismogenic

processes with those responsible for flat-lying slabs atintermediate depths. The latter has been attributed to sub-duction of oceanic plateaus or aseismic ridges where un-usually thick crust may cause the slab to become neutrallybuoyant [e.g., Gutscher et al., 1999, 2000; McGeary et al.,1985; Pilger, 1981]. This hypothesis appears to be incom-patible with earthquake-generating reactions involving eclo-gite, a material denser than the average mantle, as the slabwould surely sink once the thick crust turns into eclogite. Inaddition, it is unclear whether dehydration can explain thelocally varying sources of stress observed in flat slabs.While researchers have begun to examine stresses asso-ciated with dehydration embrittlement in the presence ofplate bending [Wang, 2002], how local stresses generatedby dehydration behave in the absence of a regional stressfield is unknown.

6.3. Summary

[35] Most current proposals of earthquake triggers aredepth-dependent, engaging different mechanisms thatoperate either in the upper mantle or at greater depths ofthe transition zone. Our results from flat-lying slabshighlight the importance of local sources of stress regard-less of depth. Obviously, a continual supply of stress, localor regional, is necessary to cause recurring seismicity.Results summarized in Figure 1 imply that slab pull,through the action of the slab as a stress guide, overridesor at least influences the local source(s) of stress. Thisprocess is apparently independent of depth. The exactmechanisms through which regional and local sources ofstress interact to produce uniform DDE/DDC are presentlyunknown and would be a natural target for future work.Nevertheless, it is interesting to note that there seems to bea threshold phenomenon at work: once the slab dip

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exceeds about 20�, statistical characteristics of the P/T axesremain more or less constant.

7. Conclusions

[36] On the basis of a database of over 1,000 fault planesolutions and 2,500 precisely relocated hypocenters, wequantify statistic characteristics of DDE/DDC in subductedslabs. Over a wide range in slab dip (from about 23� to 87�),the scatter of P/T axes around the downdip direction is moreor less uniform, approximately 33� at the 95% confidencelevel. For shallow dipping (<20�) slabs, there is no obvious,preferred orientation of fault plane solutions. In these cases,deep earthquakes are probably not caused by reactivation ofpreexisting faults from a regional stress field, because auniform strain is absent and the downdip component of slabpull is diminishing, eliminating the otherwise dominatesource of regional stress. Instead, the occurrence of abun-dant seismicity in subhorizontal slabs indicates that aspatially varying, localized source of stress is both neces-sary and sufficient to generate earthquakes. Apparently thelocalized stress is masked by DDE/DDC when one third ormore of the slab pull is along the downdip direction. Thecall for locally variable stress as part of the seismogenicprocess is consistent with the longstanding notion that alocal perturbation in stress is necessary to overcome highconfining pressure that would otherwise inhibit suddenrelease of elastic strain. Evidence is growing that seismo-genesis at depth, including localized sources of stress, isrelated to petrologic anomalies.

[37] Acknowledgments. F. Skomurski assisted us with some of thework presented in the appendices. R. Allen and C. Thurber provided helpfulcomments that improved the manuscript. We also thank T. Atwater,H. Green, B. Hacker, S. Kirby, and P. Molnar for useful discussions.T. Lay kindly shared his compilation of fault plane solutions with us.C. Frohlich, H. Green, and an anonymous associate editor provided helpfulreviews. This work is supported by National Science Foundation (NSF)grants EAR98-04718 and EAR01-06786 (W.-P.C.). Any opinions, findings,and conclusions or recommendations expressed in this material are those ofthe authors and do not necessarily reflect those of the NSF.

ReferencesApperson, K. D., and C. Frohlich (1987), The relationship between Wadati-Benioff zone geometry and P, T and B axes of intermediate and deepfocus earthquakes, J. Geophys. Res., 92, 13,821–13,831.

Araujo, M., and G. Suarez (1994), Geometry and state of stress of thesubducted Nazca plate beneath central Chile and Argentina: Evidencefrom teleseismic data, Geophys. J. Int., 116, 283–303.

Barazangi, M., and B. Isacks (1976), Spatial distribution of earthquakes andsubduction of the Nazca plate beneath South America, Geology, 4, 686–692.

Barazangi, M., B. L. Isacks, J. Oliver, J. Dubois, and G. Pascal (1973),Descent of lithosphere beneath New Hebrides, Tonga-Fiji and New Zea-land: Evidence for detached slabs, Nature, 242, 98–101.

Billington, S. (1980), The morphology and tectonics of the subducted litho-sphere in the Tonga-Fiji-Kermadec region from seismicity and focalmechanism solutions, Ph.D. thesis, Cornell Univ., Ithaca, N. Y.

Bina, C. R. (1996), Phase transition buoyancy contributions to stresses insubducting lithosphere, Geophys. Res. Lett., 23, 3563–3566.

Brudzinski, M. R., and W.-P. Chen (2000), Variations of P wave speeds andoutboard earthquakes: Evidence for a petrologic anomaly in the mantletransition zone, J. Geophys. Res., 105, 21,661–21,682.

Brudzinski, M. R., and W. Chen (2003), A petrologic anomaly accompany-ing outboard earthquakes beneath Fiji-Tonga: Corresponding evidencefrom broadband P and S waveforms, J. Geophys. Res., 108(B6), 2299,doi:10.1029/2002JB002012.

Burbach, G. V., and C. Frohlich (1986), Intermediate and deep seismicityand lateral structure of subducted lithosphere in the circum-Pacific re-gion, Rev. Geophys., 24, 833–874.

Chen, P.-F., C. R. Bina, and E. A. Okal (2001), Variations in slab dip alongthe subducting Nazca Plate, as related to stress patterns and momentrelease of intermediate-depth seismicity and to surface volcanism, Geo-chem. Geophys. Geosyst., 2, doi:10.1029/2001GC000153.

Chen, W.-P., and M. R. Brudzinski (2001), Evidence for a large-scaleremnant of subducted lithosphere beneath Fiji, Science, 292, 2475–2479.

Chen, W., and M. R. Brudzinski (2003), Seismic anisotropy in the mantletransition zone beneath Fiji-Tonga, Geophys. Res. Lett., 30(13), 1682,doi:10.1029/2002GL016330.

Chen, W.-P., and P. Molnar (1983), Focal depths of intracontinental andintraplate earthquakes and their implications for the thermal andmechanical properties of the lithosphere, J. Geophys. Res., 88, 4183–4214.

Christensen, D. H., and L. J. Ruff (1988), Seismic coupling and outer riseearthquakes, J. Geophys. Res., 93, 13,421–13,444.

Christova, C., and C. H. Scholz (2003), Stresses in the Vanuatu subductingslab: A test of two hypotheses, Geophys. Res. Lett., 30(15), 1790,doi:10.1029/2003GL017701.

Devaux, J. P., L. Fleitout, G. Schubert, and C. Anderson (2000), Stresses ina subducting slab in the presence of a metastable olivine wedge,J. Geophys. Res., 105, 13,365–13,373.

Dobson, D. P., P. G. Meredith, and S. A. Boon (2002), Simulation ofsubduction zone seismicity by dehydration of serpentine, Science, 298,1407–1410.

Dziewonski, A. M., T.-A. Chou, and J. H. Woodhouse (1981), Determina-tion of earthquake source parameters from waveform data for studies ofglobal and regional seismicity, J. Geophys. Res., 86, 2825–2852.

Engdahl, E. R., and A. Villasenor (2002), Global seismicity: 1900–1999, inInternational Handbook of Earthquake and Engineering Seismology, partA, edited by W. H. K. Lee et al., chap. 41, pp. 665–690, Elsevier, NewYork.

Engdahl, E. R., R. D. van der Hilst, and R. P. Buland (1998), Globalteleseismic earthquake relocation with improved travel times and proce-dures for depth determination, Bull. Seismol. Soc. Am., 88, 722–743.

Faul, U. H., D. R. Toomey, and H. S. Waff (1994), Intergranular basalticmelt is distributed in thin, elongated inclusions, Geophys. Res. Lett., 21,29–32.

Fisher, N. I., T. Lewis, and B. J. J. Embleton (1987), Statistical Analysis ofSpherical Data, 329 pp., Cambridge Univ. Press, New York.

Forsyth, D. W. (1982), Determinations of focal depths of earthquakesassociated with the bending of oceanic plates at trenches, Phys. EarthPlanet. Inter., 28, 141–160.

Frohlich, C. (1989), The nature of deep-focus earthquakes, Annu. Rev.Earth Planet. Sci., 17, 227–254.

Frohlich, C. (2005), Deep Earthquakes, Cambridge Univ. Press, New York,in press.

Goto, K., H. Hamaguchi, and Z. Suzuki (1985), Earthquake generatingstresses in a descending slab, Tectonophysics, 112, 111–128.

Green, H. W., and P. C. Burnley (1989), A new self-organizing mechanismfor deep-focus earthquakes, Nature, 341, 733–737.

Green, H. W., and H. Houston (1995), The mechanics of deep earthquakes,Annu. Rev. Earth Planet. Sci., 23, 169–213.

Green, H. W., II, and C. Marone (2002), Instability of deformation, inPlastic Deformation of Minerals and Rocks, edited by H. R. Wenk etal., pp. 181–199, Mineral. Soc. Am., Washington, D. C.

Gutscher, M.-A., J.-L. Olivet, D. Aslanian, J.-P. Eissen, and R. Maury(1999), The ‘‘lost Inca Plateau’’: Cause of flat subduction beneath Peru?,Earth Planet. Sci. Lett., 171, 335–341.

Gutscher, M.-A., W. Spakman, H. Bijwaard, and E. R. Engdahl (2000),Geodynamics of flat subduction: Seismicity and tomographic constraintsfrom the Andean margin, Tectonics, 19, 814–833.

Hacker, B. R., S. M. Peacock, G. A. Abers, and S. D. Holloway (2003),Subduction factory: 2. Are intermediate-depth earthquakes in subductingslabs linked to metamorphic dehydration reactions?, J. Geophys. Res.,108(B1), 2030, doi:10.1029/2001JB001129.

Hamburger, M. W., and B. L. Isacks (1987), Deep earthquakes in thesouthwest Pacific: A tectonic interpretation, J. Geophys. Res., 92,13,841–13,854.

Hasegawa, A., and I. S. Sacks (1981), Subduction of the Nazca platebeneath Peru as determined from seismic observations, J. Geophys.Res., 86, 4971–4980.

Huang, W. C., E. A. Okal, G. Ekstrom, and M. P. Salganik (1997), Centroidmoment tensor solutions for deep earthquakes predating the digital era;the World-Wide Standardized Seismograph Network dataset (1962–1976), Phys. Earth Planet. Inter., 99, 121–129.

Isacks, B., and P. Molnar (1969), Mantle earthquake mechanisms and thesinking of the lithosphere, Nature, 233, 1121–1124.

Isacks, B., and P. Molnar (1971), Distribution of stresses in the descendinglithosphere from a global survey of focal-mechanism solutions of mantleearthquakes, Rev. Geophys., 9, 103–174.

B08303 BRUDZINSKI AND CHEN: EARTHQUAKES AND STRAIN

10 of 11

B08303

Isacks, B., J. Oliver, and L. R. Sykes (1968), Seismology and the newglobal tectonics, J. Geophys. Res., 73, 5855–5899.

Isacks, B., L. R. Sykes, and J. Oliver (1969), Focal mechanisms of deep andshallow earthquakes in the Tonga-Kermadec region and the tectonics ofisland arcs, Geol. Soc. Am. Bull., 80, 1443–1470.

Jiao, W. J., P. G. Silver, Y. W. Fei, and C. T. Prewitt (2000), Do intermedi-ate- and deep-focus earthquakes occur on preexisting weak zones? Anexamination of the Tonga subduction zone, J. Geophys. Res., 105,28,125–28,138.

Jung, H., H. W. Green, and L. F. Dobrzhinetskaya (2004), Intermediate-depth earthquake faulting by dehydration embrittlement with negativevolume change, Nature, 428, 545–549.

Kirby, S. (1987), Localized polymorphic phase transformations in high-pressure faults and applications to the physical mechanism of deep earth-quakes, J. Geophys. Res., 92, 13,789–13,800.

Kirby, S. (1995), Interslab earthquakes and phase-changes in subductinglithosphere, Rev. Geophys., 33, 287–297.

Kirby, S. H., W. B. Durham, and L. A. Stern (1991), Mantle phase changesand deep-earthquake faulting in subducting lithosphere, Science, 252,216–225.

Kirby, S. H., E. R. Engdahl, and R. P. Denlinger (1996a), Intermediate-depth intraslab earthquakes and arc volcanism as physical expressions ofcrustal and uppermost mantle metamorphism in subducting slabs, inSubduction: Top to Bottom, Geophys. Monogr. Ser., vol. 96, edited byG. E. Bebout, pp. 195–214, AGU, Washington, D. C.

Kirby, S. H., S. Stein, E. A. Okal, and D. C. Rubie (1996b), Metastablemantle phase transformations and deep earthquakes in subducting ocea-nic lithosphere, Rev. Geophys., 34, 261–306.

Lay, T., L. Astiz, H. Kanamori, and D. H. Christensen (1989), Temporalvariation of large intraplate earthquakes in coupled subduction zones,Phys. Earth Planet. Inter., 54, 258–312.

Marton, F. C., C. R. Bina, S. Stein, and D. C. Rubie (1999), Effects of slabmineralogy on subduction rates, Geophys. Res. Lett., 26, 119–122.

McGeary, S., A. Nur, and Z. Ben-Avraham (1985), Spatial gaps in arcvolcanism: The effect of collision or subduction of oceanic plateaus,Tectonophysics, 119, 195–221.

Meade, C., and R. Jeanloz (1991), Deep-focus earthquakes and recycling ofwater into the Earth’s mantle, Science, 252, 68–72.

Molnar, P., D. Freedman, and J. S. F. Shih (1979), Lengths of intermediateand deep seismic zones and temperatures in downgoing slabs of litho-sphere, Geophys. J. R. Astron. Soc., 56, 41–54.

Okal, E. A. (2001), ‘‘Detached’’ deep earthquakes: Are they really?, Phys.Earth Planet. Inter., 127, 109–143.

Okal, E. A., and S. H. Kirby (1998), Deep earthquakes beneath the Fijibasin, SW Pacific: Earth’s most intense deep seismicity in stagnant slabs,Phys. Earth Planet. Inter., 109, 25–63.

Pilger, R. H., Jr. (1981), Plate reconstructions, aseismic ridges, and low-angle subduction beneath the Andes, Geol. Soc. Am. Bull., 92, 448–456.

Raleigh, C. B., and M. S. Paterson (1965), Experimental deformation ofserpentinite and its tectonic implications, J. Geophys. Res., 70, 3965–3985.

Richter, F. M. (1979), Focal mechanisms and seismic energy release of deepand intermediate depth earthquakes in the Tonga-Kermadec region andtheir bearing on the depth extent of mantle flow, J. Geophys. Res., 84,6783–6795.

Sacks, I. S. (1977), Interrelationships between volcanism, seismicity, andanelasticity in western South America, Tectonophysics, 37, 131–139.

Savage, J. C. (1969), The mechanics of deep-focus faulting, Tectonophys-ics, 8, 115–127.

Savage, M. K. (1999), Seismic anisotropy and mantle deformation: Whathave we learned from shear wave splitting?, Rev. Geophys., 37, 65–106.

Scholz, C. H. (1990), Mechanics of Earthquakes and Faulting, 439 pp.,Cambridge Univ. Press, New York.

Simkin, T., and L. Siebert (1994), Volcanoes of the World, Geosci. Press,Tucson, Ariz.

Skomurski, F. N., M. R. Brudzinski, and W.-P. Chen (1999), Complexrupture associated with a large outboard earthquake, Eos Trans. AGU,80(46), Fall Meet. Suppl., Abstract S12F-05.

Smalley, R. F., and B. L. Isacks (1987), A high-resolution local-networkstudy of the Nazca plate Wadati-Benioff zone under western Argentina,J. Geophys. Res., 92, 13,903–13,912.

Tavera, H., and E. Buforn (2001), Source mechanism of earthquakes inPeru, J. Seismol., 5, 519–539.

Vassiliou, M. S., B. H. Hager, and A. R. Raefsky (1984), The distribution ofearthquakes with depth and stress in subducting slabs, J. Geodyn., 1, 11–28.

Wadati, K. (1927), Existence and study of deep-focus earthquakes (inJapanese), J. Meteorol. Soc. Jpn., Ser. 2, 5, 119–145.

Wang, K. (2002), Unbending combined with dehydration embrittlement asa cause for double and triple seismic zones, Geophys. Res. Lett., 29(18),1889, doi:10.1029/2002GL015441.

Wiens, D. A., and S. Stein (1983), Age dependence of oceanic intraplateseismicity and implications for lithospheric evolution, J. Geophys. Res.,88, 6455–6468.

Wiens, D. A., J. J. McGuire, and P. J. Shore (1993), Evidence for transfor-mational faulting from a deep double seismic zone in Tonga, Nature, 364,790–793.

Zhang, H. J., C. H. Thurber, D. Shelly, S. Ide, G. C. Beroza, andA. Hasegawa (2004), High-resolution subducting-slab structure beneathnorthern Honshu, Japan, revealed by double-difference tomography,Geology, 32, 361–364.

�����������������������M. R. Brudzinski, Department of Geology and Geophysics, University of

Wisconsin-Madison, 1215 W. Dayton St., Madison, WI 53706, USA.(brudzins@geology.wisc.edu)W.-P. Chen, Department of Geology, University of Illinois at Urbana-

Champaign, 1301 W. Green St., 245 NHB, Urbana, IL 61801, USA.(wpchen@uiuc.edu)

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