JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 117, B05402, doi:10.1029/2011JB008992, 2012

Upper-mantle earthquakes beneath the Arafura and south Aru Trough: Implications for continental rheology R. A. Sloan1 and J. A. Jackson1 Received 4 November 2011; revised 5 March 2012; accepted 8 March 2012; published 1 May 2012.

[1] The extent and controls of long-term elastic strength and seismicity in the upper continental lithospheric mantle (UCLM) are controversial topics in continental tectonics. One key issue is the scarcity of UCLM earthquakes, even where the UCLM is likely to be colder than 600C. The rarity of these earthquakes could be because the UCLM generally relatively hydrous causing it to deform aseismically even when colder than 600C, unless it is deforming at exceptionally high strain rates. Alternatively, the UCLM could be relatively anhydrous, and potentially seismogenic at temperatures below 600C; in which case the rarity of UCLM earthquakes may be because areas where the UCLM is colder than 600Chave such a thick seismogenic layer, and such a cool mantle root, that they deform exceptionally slowly. The identification and study of UCLM earthquakes allows us to distinguish between these possibilities. Here we show that two earthquakes occurred in the UCLM beneath the epicontinental Arafura Sea. Both earthquakes occurred where the UCLM is probably cooler than 600C and one of these earthquakes lies 25 km below the Moho in an where there is no evidence of unusually high strain rates. There at least, it is probable that the UCLM is relatively anhydrous, and seismogenic at temperatures below 600C. We also find evidence that where the UCLM is colder than 600C also have a seismogenic lower crust. This results in a single, extremely strong layer comprising the entire crust and the UCLM down to the 600Cisotherm. Citation: Sloan, R. A., and J. A. Jackson (2012), Upper-mantle earthquakes beneath the Arafura Sea and south Aru Trough: Implications for continental rheology, J. Geophys. Res., 117, B05402, doi:10.1029/2011JB008992.

1. Introduction [3] The challenge we face is to make observations that constrain the actual rheological behavior of upper-mantle [2] One of the largest gaps in our knowledge of the factors material beneath the . One method is to accurately governing continental tectonics is the rheology of the upper determine the centroid depths of earthquakes. This reveals continental lithospheric mantle (UCLM). Views on the long- the depth to the seismic-aseismic transition and places an term strength and the deformation mechanism of the UCLM important bound on the UCLM conditions that control vary widely. Experimental research on the rheological seismogenic behavior. properties of mantle minerals has shown that the flow laws [4] In oceanic lithosphere the seismogenic and elastic which control deformation in mantle materials are extremely thicknesses increase with plate age, suggesting both are sensitive to minor compositional changes, especially the controlled by temperature [Watts et al., 1980; Chen and presence or absence of small amounts of hydrogen dissolved Molnar, 1983; Wiens and Stein, 1983]. When oceanic geo- in nominally anhydrous minerals such as olivine and therms are calculated, taking into account the temperature pyroxene [Mackwell et al., 1998; Hirth and Kohlstedt, 2003; dependence of thermal conductivity, the seismic-aseismic Boettcher et al., 2007]. If the UCLM is relatively hydrous transition follows the 600C isotherm [Denlinger, 1992; these minerals contain a significant amount of hydrogen and McKenzie et al., 2005]. However, in oceanic lithosphere it is the UCLM will undergo aseismic deformation at tempera-  likely that the lithospheric mantle has been dehydrated by tures well below 600 C and will lack significant long-term melting beneath the mid- ridge. It is far from certain strength [Maggi et al., 2000]. Alternatively, relatively that the UCLM should follow the same pattern, especially anhydrous UCLM should remain seismogenic to tempera- beneath ancient shields where the UCLM may have been tures of up to 600C[Boettcher et al., 2007], and could gradually hydrated by metasomatic melts from the astheno- contribute significantly to the long-term strength of the sphere over a long period of time [Harte et al., 1993; Maggi continents in some areas [McKenzie et al., 2005]. et al., 2000]. [5] In the continents, UCLM earthquakes were once

1 thought to be present in many places, leading to the con- Bullard Laboratories, Department of Sciences, University of Cambridge, Cambridge, UK. clusion that the UCLM was generally seismogenic, and contributed greatly to the long-term strength of the con- Copyright 2012 by the American Geophysical Union. tinents [Chen and Molnar, 1983]. Indeed, the UCLM was 0148-0227/12/2011JB008992

B05402 1of13 B05402 SLOAN AND JACKSON: UCLM EARTHQUAKES BENEATH THE ARAFURA SEA B05402 thought to contribute more to the long-term strength of the estimates from upper mantle xenoliths form southern India continents than the lower crust, which is aseismic in many also suggest the Indian Shield has Moho temperatures of deforming regions. This view contributed to the widespread 500C[Priestley et al., 2008]. This observation supports acceptance of a laminated or “jelly sandwich” model of the view that the UCLM is relatively anhydrous and is continental rheology [Chen and Molnar, 1983]. Reassess- capable of seismic deformation at temperatures up to ments of earthquake and Moho depths in active regions have 600C. However, some caution is required due to the tec- since revealed that UCLM earthquakes are in fact rare. It tonic setting of these microearthquakes. India is being was therefore suggested that, in most areas, the UCLM is thrust under Tibet at 20 mm yrÀ1 [Larson et al., 1999] aseismic and contributes little to long-term elastic strength and these UCLM events occur where the Indian plate first [Maggi et al., 2000; Jackson et al., 2008]. Geotherm mod- bends and then unbends in a ramp-and-flat geometry eling, taking into account modern estimates of crustal beneath the Himalaya. The UCLM in this region is thickness, radiogenic heat production and temperature deforming at relatively high strain rates, which may allow it dependent conductivity, revealed that Moho temperatures in to undergo seismic deformation at higher temperatures than actively deforming areas are generally greater than 600C it would under normal conditions [Priestley et al., 2008]. [McKenzie et al., 2005]. The UCLM in those areas would be [9] One, possibly two, UCLM earthquakes have also been expected to deform aseismically, and be relatively weak, reported in the Andean UCLM in an area where a shallow regardless of its composition. and subhorizontally subducting oceanic slab underplates the [6] There are, however, many continental regions, par- Andean continental lithosphere [Emmerson, 2007]. These ticularly in ancient shields, which are underlain by rela- earthquakes occur in the UCLM between the cold subduct- tively thick lithosphere (>150 km), have low-to-moderate ing slab and the surface. This region is therefore being crustal thicknesses (<40–45 km), and lack large amounts of cooled from both above and below, resulting in temperatures radiogenic crustal heat production. In these areas (e.g. the colder than 600C[Emmerson, 2007]. In this unusual setting Canadian and Siberian shields) the Moho temperature is both coupling between the subducting oceanic slab and the significantly cooler than 600C. If the UCLM is relatively UCLM and the release and upward migration of fluid from anhydrous, and UCLM seismicity is controlled by the the subducting slab could result in unusually high strain rates 600C isotherm, then the UCLM in such places should be in the area where the earthquakes occurred. capable of seismic deformation, and will contribute greatly [10] It is therefore important to test the hypothesized to long-term elastic strength. Alternatively, if the UCLM in 600C seismic-aseismic transition temperature in areas such places is relatively hydrous, perhaps due to pervasive where the UCLM is cooler than 600C and is not deforming metasomatism, the UCLM is likely to deform aseismically, at unusually high strain rates. However, most ancient shields and will contribute little to long-term elastic strength. are deforming so slowly that earthquakes occur only very [7] One way of resolving the ambiguity in UCLM rhe- occasionally. The rare earthquakes that do occur in these ology would be to consider the effective elastic thickness of settings are therefore potentially very informative, especially regions where the 600C isotherm is expected to lie well if they occur in the UCLM. Such earthquakes provide the below the Moho. Unfortunately, whilst in principle it is opportunity to test whether the rarity of UCLM seismicity possible to use the relationship between gravity and topog- beneath ancient shields is because the UCLM there is rela- raphy to estimate the effective elastic thickness [Forsyth, tively hydrous and so deforms aseismically, or because the 1985; McKenzie and Fairhead, 1997], in practice differ- UCLM is anhydrous and deforms seismically, but the great ent methods produce estimates from ancient shields which seismogenic thickness and thick cold mantle root result in vary by an order of magnitude (from 20 to 200 km) such slow deformation that this has not yet been recognized. [Pérez-Gussinyé and Watts, 2005; Kirby and Swain, 2009; Pérez-Gussinyé et al., 2009; McKenzie, 2010]. This is 2. Geological Setting because these methods are only reliable in areas with sig- nificant topography, and high coherence between free-air 2.1. Arafura Sea gravity anomalies and surface loads. Shields typically have [11] One area with sufficiently thick lithosphere and thin very subdued topography and low coherence due to the crusttohaveUCLMcolderthan600C is the Arafura Sea, a action of subaerial erosion over millions of years, preventing shallow (mainly <200 m deep) epicontinental sea separating the accurate determination of their effective elastic thickness from (Figure 1a). It is part of the [McKenzie, 2010]. Even when well-constrained elastic Australian plate and is bordered by the Banda arc subduction thickness estimates can be obtained they cannot constrain zone to the west and the New Guinea highlands to the north. which part or depth level of the lithosphere provides long- Whilst these boundaries are very seismically active, the inte- term elastic strength. rior of the sea is virtually aseismic (Figure 1a). Two seismo- [8] Determining the depth to the seismic-aseismic transi- genic slabs of subducted oceanic lithosphere exist in the area, tion in ancient shields therefore provides our best chance of beneath the in the west and New Guinea in the constraining UCLM rheology. Monsalve et al. [2006] used a north, but they do not extend beneath the central Arafura Sea. local network in the eastern Himalaya to show that micro- [12] Onshore receiver functions in NW Australia show that seismicity occurs within the uppermost 5–10 km of the the crustal thickness varies from 38–46 km [Clitheroe et al., Indian UCLM (as well as in the lower crust) where it is forced 2000] and in the Arafura Sea estimates from seismic refrac- below Tibet. North India has relatively thick lithosphere tion profiles (yellow triangles in Figure 1a) vary from 27– and relatively thin crystalline crust, and so the Moho tem- 40 km [Jacobson et al., 1979; Rynn and Reid, 1983]. The perature where these microearthquakes occur is likely to be upper end of this range is found where the Australian plate below 600C[Priestley et al., 2008]. Pressure-temperature bends down beneath the Banda arc, whereas the only

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Figure 1. Topography, seismicity and lithospheric thickness in the Arafura Sea and surrounding areas. (a) Earthquakes, taken from the updated EHB catalogue [Engdahl et al., 1998], are shown as black (cen- troid depth <100 km) or red (centroid depth >100 km) dots. Yellow triangles show the locations of seismic refraction profiles and the calculated Moho depth is shown beside receiver site [Jacobson et al., 1979; Rynn and Reid, 1983]. The red star indicates the location of the near-Moho 1992 earthquake discussed in the text. The 2000 Arafura Sea earthquake is marked by the strike-slip focal mechanism (compressional quadrants shaded). GC, BS indicate the and the Banda Sea, and T, NG and AI indicate Timor, New Guinea and the Aru islands respectively. C and J indicate the locations of two short seismic refraction profiles referred to in section 4.1. The area surrounded by a continuous line is shown in more detail in Figure 2. (b) Lithosphere thickness map (section 4.2) [McKenzie and Priestley, 2008] of the area enclosed by dashed lines in Figure 1a. Points at which lithospheric thickness has been calculated are shown by large black dots, and the four points surrounding the 2000 Arafura Sea earthquake are ringed in yellow. The geotherms at these points are examined in Figure 5. measurement of less than 30 km (between the Aru islands [13] Figure 1b shows a map of lithospheric thickness and New Guinea) is an unreversed seismic refraction line derived from surface-wave tomography [McKenzie and with anomalously low (7.63 km sÀ1) mantle P-wave velocity Priestley, 2008]. The lithospheric thickness has been calcu-   (Vp)[Jacobson et al., 1979]. Close to Australia, away from lated on a 2 Â 2 grid (large black dots in Figure 1b) and has the edge of the plate, seismic refraction studies suggest the a vertical resolution of 30 km. The 50–160 s period range crustal thickness is 31–34 km. that is used limits the horizontal resolution to 300 km. The

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Figure 2. Bathymetry and seismicity of the Aru Trough and surrounding areas. Grey focal mechanisms represent earthquakes with reliable focal mechanisms taken from the CMT catalogue. Black, green and red focal mechanisms represent waveform-modeled earthquakes with reliable centroid depths in the upper crust (<20 km), the lower crust (20–32 km) and possibly the upper mantle (32+ km) respectively. A–B marks the line of the cross section shown in Figure 5. Arrows and yellow squares indicate GPS motions relative to the Australian plate taken from Bock et al. [2003] and the ovals represent the 2D 95% confidence interval. White triangles mark the end points of the reversed seismic refraction profile reported by Jacobson et al. [1979]. Contours of lithosphere thickness (see Figure 1b) are shown as thin lines. Lithosphere thick- ness values are calculated at 2 by 2 intervals shown as large black dots, and values have a 300 km hor- izontal resolution. ST, and TT indicate the Seram and Tanimbar troughs respectively. WD, BH and IJ mark the Weber Deep, the Bird’s Head Peninsula and Irian Jaya. The thick black line labeled TAF represents the Tarera-Aiduna fault, and its possible offshore extension is marked by a dashed line. methods and uncertainties [Priestley and McKenzie, 2006; geological observations and earthquake focal mechanisms McKenzie and Priestley, 2008] are described in section 4.2. all indicate that the Aru Trough is extending [Jacobson et al., Figure 1 shows that there is a steep gradient in lithospheric 1979; Bowin et al., 1980; Charlton et al., 1991; Milsom et al., thickness beneath the Arafura Sea, from ≤125 km on the 1996; Bock et al., 2003]. To the north, it projects to a north, west and east margins to >200 km towards Australia in tectonically complicated region known as the Snellius II the south. This feature is also observed by regional tomo- triple junction, where it appears to meet the NW trending graphic studies [van der Hilst et al., 1998; Debayle and Seram trough and a poorly defined left-lateral strike-slip Kennett, 2000; Fichtner et al., 2009]. Fishwick et al. [2008] zone associated with the EW Tarera-Aiduna fault system studied the velocity structure of Australia and the S Arafura [Jongsma et al., 1989]. Sea using surface wave tomography. They found positive Vs [15] Numerous papers have described this area, including a anomalies consistent with the presence of a thick lithospheric seismic refraction profile [Jacobson et al., 1979], seismic mantle root beneath the S Arafura Sea, however the N Ara- reflection profiles [Bowin et al., 1980; Jongsma et al., 1989; fura Sea and Aru Trough lay outside of the study region. Untung, 1985], gravity and heat flow data [Bowin et al., 1980], geological studies of the Kai islands [Charlton et al., 2.2. Aru Trough 1991; Milsom et al., 1996], GPS data [Bock et al., 2003] [14] The Aru Trough (Figure 2) is a seismically active and seismicity [Cardwell and Isacks, 1978; McCaffrey, NNE-SSW trending 3.5 km deep depression in the W 1988, 1989; McCaffrey et al., 1991]. The tectonic interpre- Arafura Sea flanked on the east by the Aru islands and on the tation of the Aru Trough is controversial (see section 5.2). It west by the Kai islands. It appears to be a bathymetric con- is the site of active normal faulting [Jacobson et al., 1979; tinuation of the shallower Tanimbar and Timor “Troughs” to Bowin et al., 1980; McCaffrey, 1989] and Jongsma et al. the south which are foreland basins associated with thrusts [1989] suggested that extension began 3 Ma on the basis of the Banda -arc collision; but GPS constraints, of seismic sequence stratigraphy.

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[16] A reversed seismic refraction profile revealed that the reported 200 from the Arafura Sea earthquake, a reversed southern part of Aru Trough is continental and has a crustal profile to the west [Jacobson et al., 1979] (labeled J in thickness of 32 km compared with 40 km in the Tanimbar Figure 1a) and an unreversed profile to the ESE [Curray trough to the south [Jacobson et al., 1979]. Seismic reflec- et al., 1977] Both these profiles report a thin layer of tion profiles show active normal faults with throws of sev- unconsolidated sediment (<0.5 km), a 3–4 km layer of rela- À1 eral hundred meters in the eastern Aru Trough. Large-scale tively fast sediment (Vp = 5.4–5.9 km s ) interpreted as extensional features are absent in the Tanimbar trough fur- Eocene carbonates[Jacobson et al., 1979], and a crystalline À1 ther south [Jacobson et al., 1979]. The steep gradient in basement (Vp = 6.1–6.4 km s )at4–5.5 km depth. The lithospheric thickness in the north Arafura Sea (Figures 1b velocity structure at greater depths is unconstrained. We have À1 and 2) [McKenzie and Priestley, 2008] if projected to the chosen to use a simple velocity model with Vp = 6.5 km s , À1 À3 west, occurs between the Aru and Tanimbar troughs. How- Vs = 3.7 km s and density r = 2750 kg m for the crust and À1 À1 À3 ever, as the horizontal resolution is only 300 km, varia- Vp = 8.0 km s , Vs 4.5 km s and density r = 3,300 kg m tions correlated with surface features on the scale of for the mantle. The seismic refraction profiles reveal a rela- 100 km, such as the Aru Trough, cannot be precise. The tively fast shallow velocity structure and as velocities are Aru Trough was formerly thought to mark the surface likely to increase with depth the assumed model is unlikely to expression of the Australian-Banda subduction zone [Bowin be significantly faster than the average crustal velocity. et al., 1980; Untung, 1985; McCaffrey, 1988, 1989; Tandon [19] We have better constraints on crustal velocities in the et al., 2000]. However, field work has shown that the W Kai Aru Trough. A reversed seismic refraction line from the consists of high-grade metamorphic rocks thrust over an Tanimbar Trough to the southernmost Aru Trough (white imbricated cover sequence interpreted as an accretionary triangles in Figure 2) [Jacobson et al., 1979] passes <30 km wedge whilst the E Kai islands consist of shallowly west- from the earthquakes in the southern Aru Trough. For dipping Eocene-Pliocene sediments cut by NNE east-dipping earthquakes in this area we have used the averaged crustal normal faults [Charlton et al., 1991; Milsom et al., 1996]. and mantle velocities taken from the N end of this refraction À1 À1 Both these geological observations, and the clear extension profile (in the crust Vp = 5.7 km s , Vs = 3.3 km s and À3 À1 between E Kai and the Aru Islands measured by GPS [Bock r = 2700 kg m and in the mantle Vp = 7.85 km s , À1 À3 et al., 2003], show that active convergence is now situated Vs = 4.5 km s and r = 3,300 kg m ). The velocity west of E Kai as shown in Figure 2. structure is relatively slow due to the presence of 5kmof À1 sediments with Vp < 3.3 km s . 3. Methods 3.2. Lithospheric Thickness and Temperature 3.1. Long-Period Body Wave Inversion for Earthquake Modeling Source Parameters [20] The lithospheric thickness map shown in Figure 1b [17] To obtain earthquake source parameters (Figure 3), [McKenzie and Priestley, 2008] has been produced by we used broadband seismograms from the Global Digital using fundamental and higher-mode surface-wave tomogra- Seismograph Network (GDSN) and deconvolved them to phy to map shear wave velocity (V ) as a function of depth change the response to that of a World-Wide Standardized s (z)[Debayle and Kennett, 2000; Ritsema and van Heijst, Seismograph Network (WWSSN) 15–100 long-period 2000; Priestley and Debayle, 2003]. Vs(z) is then converted instrument. This reduces sensitivity to complexities in local  Â  – to a temperature (T) at 25 km depth intervals on a 2 2 velocity structure and allows events of Mw 5.0 6.5 to be grid (large black dots in Figure 1b) using an empirical rela- modeled as point (centroid) sources. We then use the MT5 tionship [McKenzie et al., 2005; Priestley and McKenzie, version [Zwick et al., 1994] of the SYN4 algorithm 2006]. A continuous geotherm is then fitted to these T(z) [McCaffrey et al., 1991], which inverts P and SH waveform estimates using temperature-dependent thermal parameters data for the source time function, scalar moment, strike, dip, [McKenzie et al., 2005] and a mantle potential temperature rake and centroid depth. The source is constrained to be a  of 1315 C. The V (z) profile has a vertical resolution of pure double-couple. P, pP and sP phases are modeled on s 30 km and a lateral resolution of 200–400 km resulting vertical component seismograms with an epicentral range from the 50–160 s period range used [Priestley and spanning 30–90 and S and sS phases are modeled on McKenzie, 2006; McKenzie and Priestley, 2008]. Previous transverse component seismograms in the range 30–80. studies have demonstrated that the values of Vs inferred for Amplitudes are corrected for the effects of geometrical depths ≤100 km are affected by relatively low crustal spreading and anelastic attenuation using Futterman opera- velocities [Debayle and Kennett, 2000; Pilidou et al., 2004; tors with a t* of 1.0 s and 4.0 s for P and SH waves McKenzie and Priestley, 2008]. This results in anomalously respectively. Stations are weighted according to azimuthal high temperature estimates for these depths which we ignore density and the SH waveforms are weighted 0.5 relative to in this study. Temperatures calculated for depths ≥125 km smaller amplitude P waveforms. When possible the wave- are not significantly affected by the choice of crustal velocity forms are aligned using time picks from the broadband structure [Debayle and Kennett, 2000; Pilidou et al., 2004]. records. This is a routine method widely used to obtain These considerations limit geotherm estimates to areas accurate centroid depths [e.g., Maggi et al., 2000]. Centroid where the lithospheric thickness is greater than 125 km. depths determined in this way have uncertainties of typically The base of the lithosphere is identified by the rapid change Æ4 km, compared to Æ10–15 km for the better located in temperature gradient over the thickness of the thermal events in the EHB catalogue [Engdahl et al., 2006]. boundary layer separating advective and conductive heat [18] The velocity structure beneath the Arafura Sea is not transport [Priestley and McKenzie, 2006]. well constrained. Two short seismic refraction lines are

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Figure 3. Minimum misfit solution for the 23rd December 2000 Arafura Sea Earthquake. The para- meters used are strike = 60, dip = 82, rake = 6, centroid depth = 61 km, Mw = 6.0 and the source time function is shown between the two focal spheres. The synthetic waveforms are shown as dashed lines and the observed waveforms are shown as continuous lines. Stations are identified by a letter A–Z arranged clockwise by azimuth on the focal sphere. The upper circle represents the P focal sphere, and the lower circle represents the SH focal sphere. The amplitiude scale is given, in microns, to the left of each focal sphere and the horizontal time scale is given below the source time function. The P and T axes are shown as a filled and an open circle respectively. Stations not included in the inversion due to relatively high noise levels or the arrival of an unmodeled phase are marked by an asterisk. The inversion window is marked by vertical bars for each station. This model was calculated with a 35 km crustal layer.

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Figure 4. (a) Geotherms calculated at four points (marked in yellow on Figure 1b) surrounding the 2000 Arafura earthquake, using the method described in section 4.2. A 35 km crustal thickness has been assumed. TL indicates the lithospheric thickness calculated for each location, the continuous line indicates the calculated geotherm and the squares indicate temperature estimated from Vs. Only the shaded squares  in each rectangle are used to fit the geotherm. Below 900 C, Vs does not vary enough with temperature to allow an accurate conversion and above the horizontal line at 125 km temperature estimates obtained from Vs increase due to the influence of poorly known low Vs in the crust [Priestley and McKenzie, 2006]. These geotherms must be viewed with caution as the measured Vs can be affected by variations up to 300 km away. (b) Hottest and coldest geotherms shown in Figure 4a with the possible centroid depth range for the 2000 Arafura earthquake. The possible depth-temperature range at which the earthquake is likely to have occurred is shaded and is ≤600C in all cases.

[21] To estimate the geotherm in the area of the 2000 Ara- errors are described in more detail elsewhere [Priestley and fura earthquake we have studied the four points ringed in McKenzie,2006;McKenzie and Priestley, 2008]. yellow in Figure 1b. We calculated geotherms in the manner described above, except that here the expected variation of 4. Results crustal conductivity as a function of temperature [Whittington et al., 2009] is also considered. The radiogenic heat production 4.1. The Arafura Sea À3 in the crust is assumed to be 2 mWm in the upper 5 km and [22] Two earthquakes in the Arafura Sea have well-deter- À 0.2 mWm 3 below [Rudnick and Fountain, 1995]. Although mined centroid depths away from the Aru Trough (Table 1). the lithosphere thickness map [McKenzie and Priestley,2008] Both are unusually deep. The first occurred on 30 September was obtained assuming a constant conductivity in the crust, the 1992 beneath the south Arafura Sea (red star on Figure 1a). crustal geotherm calculated here (Figure 4) remains approxi- McCue and Michael-Leiba [1993] used depth-phase identi- mately linear implying that temperature-dependent conduc- fication to determine that this mb = 5.4 earthquake occurred tivity in the crust does not significantly affect the temperatures at a depth of 38.8 Æ 2.5 km. It was interpreted to be too close we obtain. It is important to note that, because of gradients in to the Moho to differentiate between a lower-crustal or lithospheric thickness within 300–400 km of the event, the upper-mantle origin as there are no reliable constraints on limited lateral resolution of the tomography may influence the crustal thickness or velocity structure in the vicinity of the calculated temperatures. The methods used and expected earthquake. The second, reported for the first time here, is a very unusual event.

Table 1. Arafura Sea Earthquake Source Parameters From Body Wave Modelinga Date Time Focal Mechanism      Year Month Day Hour MinuteLatitude ( ) Longitude ( ) Depth (km) MW Strike( ) Dip( ) Rake( ) Reference 1992 09 30 11 18 À11.39 134.47 39 5.4 –––McC 2000 12 23 07 13 À7.91 135.76 61 6.0 60 83 5 TS aEpicentres and origin times between 1964–2007 are taken from the updated catalogue of Engdahl et al. [1998]. References from which source parameters have been taken are indicated by McC [McCue and Michael-Leiba, 1993] and TS (this study). The depth of the first event was determined by depth-phase modeling and no reliable mechanism is available. The source parameters of the remaining event was obtained through the inversion of teleseismic body waveforms. The focal mechanisms are plotted in Figures 1 and 2.

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Table 2. Aru Trough Earthquake Source Parameters From Body Wave Modelinga Date Time Focal Mechanism      Year Month Day Hour MinuteLatitude ( ) Longitude ( ) Depth (km) MW Strike( ) Dip( ) Rake( ) Reference 1973 04 17 12 34 À4.35 134.22 10 6.4 243 88 7 Ab 1988 07 25 06 46 À6.07 133.64 25 6.7 53 49 À45 TS 1995 01 27 20 16 À4.45 134.45 05 6.8 233 28 À31 TSb 1996 01 10 22 36 À6.17 133.58 37 5.9 80 43 À40 TS 1999 06 22 00 47 À4.47 133.92 14 5.6 178 64 À140 TS 2000 09 02 00 25 À5.07 133.56 30 5.4 71 70 À14 TS 2004 02 07 21 27 À4.11 133.92 12 5.5 102 88 0 TS 2010 01 12 10 58 À5.36 133.73 34 5.3 290 53 À168 TS 2010 10 12 10 21 À4.89 133.66 13 6.1 288 80 À172 TSb aEpicentres and origin times between 1964–2007 are taken from the updated catalogue of Engdahl et al. [1998]. References from which source parameters have been taken are indicated by Ab [Abers and McCaffrey, 1988] and TS (this study). The source parameters of all these events were obtained through the inversion of long-period teleseismic body waveforms. Focal mechanisms are plotted in Figure 2 with the compressional quadrants shaded black (depth < 20 km), green (depth = 20–32 km) or red (depth > 32 km). bEvents where a sub-event was needed to achieve a satisfactory waveform fit.

[23] On 23 December 2000 at 07:13:25 UTC an earth- average temperature range at the centroid depth. Figure 4a   quake of Mw 6.0 occurred at 7.92 S 135.76 E, SW of the shows the calculated geotherms for each point assuming a Aru islands (Figure 1). Its focal mechanism and centroid crustal thickness of 35 km. The coolest geotherm is at the depth are well constrained by teleseismic P and SH wave- SW point, where the lithosphere is thickest, and the tem- form analyses, described in more detail in section 4.1. The peratures calculated from Vs(z) (shown as squares) identify reasonable azimuthal station coverage and the separation of the change from advective to conductive gradient relatively depth phases at many stations (such as YSS and YAK for P well. The lithospheric thickness at the other three points is waves and WMQ and DRV for SH waves; Figure 3) con- close to the minimum value (125 km) that can be con- strain the centroid depth to be 60 km. The good fit of strained using this method [Priestley and McKenzie, 2006], synthetic and observed amplitudes for both P and SH also as at shallower depths the Vs(z) estimates may be contami- rules out the possibility that these depth phases represent a nated by the poorly known crustal velocity structure. At second separate event. these three points the rapid change in gradient at the thermal [24] In Figure S1 of Text S1 of the auxiliary material we boundary zone is less clear, and so these geotherms are less explore the effect of varying the velocity model within the well constrained. The coldest and hottest calculated geo- range of plausible crustal thicknesses reported beneath the therms, corresponding to the SW and NE points, are plotted Arafura sea (30–40 km, Figure 1a).1 Within that range in Figure 4b as continuous lines and the possible depth range the centroid depth of the minimum-misfit model varies by of the 2000 Arafura Sea earthquake is shown by horizontal only 2 km. For the earthquake to have occurred in the lower bounds. The range of depth-temperature values at which the crust would require a crustal thickness of greater than 55 km, earthquake is likely to have occurred (shaded in Figure 4b)  which is implausible for an area below . Figure S2 lies below the 600 C isotherm in all cases. Doubling the À of Text S1 of the auxiliary material shows the results of a radiogenic heat production (to 4 mWm 3 in the upper 5 km  centroid depth resolution test intended to explore the effect and 0.4 mWmÀ3 in the lower crust) still results in a 600 C of possible trade-offs between centroid depth and other isotherm deeper than the earthquake centroid depth of parameters. It shows that there is a clear minimum in misfit 61 km, even for the hottest calculated geotherm. The loca- centered on models with a centroid depth of 61 km, and that tion of this earthquake adjacent to a steep boundary in lith- the waveform fit is significantly degraded if the centroid osphere thickness, the difficulty in constraining lithosphere depth is altered by more than Æ2 km. We therefore conclude thickness in areas where it is relatively thin, and the lack of that the likely range of possible centroid depths is limited to strong constraints on the amount and distribution of crustal 57–65 km. radiogenic heat production in this area mean that we cannot [25] This analysis confirms that the earthquake occurred definitively state that this earthquake occurred shallower  within the UCLM. To explore the possible temperatures at than the 600 C isotherm. However, given that a number of the source of this unusual event, we use Vs(z) profiles to different tomographic models suggest that this area is calculate the geotherm at the four points around the earth- underlain by thick lithosphere [van der Hilst et al., 1998; quake ringed in yellow in Figure 1b. The method [Priestley Debayle and Kennett, 2000; McKenzie and Priestley, 2008; and McKenzie, 2006] is described in section 4.2, but it is Fichtner et al., 2009] and most cratonic crustal rocks are important to recall that for each point the calculated Vs(z)is characterized by relatively low radiogenic crustal heat pro- affected by velocity variations up to 300 km away duction [Rudnick and Fountain, 1995] it is likely that this  [Priestley et al., 2006] and, as the earthquake is close to a earthquake occurred in material colder than 600 C. gradient in lithospheric thickness, the geotherms produced must be regarded with caution. However, we used all four of 4.2. The Aru Trough the points surrounding the epicenter to constrain the likely [26] The results of long-period body waveform modeling for eight earthquakes that occurred in the Aru Trough 1Auxiliary materials are available in the HTML. doi:10.1029/ between 1988 and 2010 are shown in Table 2 and Figure 2. 2011JB008992. No previous study has examined the depth range of

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Table 3. Aru Trough Selected Earthquake Source Parameters From the CMT Cataloguea Date Time Focal Mechanism      Year Month Day Hour MinuteLatitude ( ) Longitude ( ) Depth (km) MW Strike( ) Dip( ) Rake( ) Reference 1983 02 11 04 34 À5.73 133.76 38 5.3 36 27 À68 CMT 1985 02 17 21 36 À4.83 133.95 22 5.5 236 62 À19 CMT 1986 04 29 07 28 À4.55 133.61 19 5.4 334 58 À156 CMT 1987 05 26 01 37 À5.53 133.91 35 5.1 202 29 À74 CMT 1989 12 05 03 58 À5.81 133.70 39 5.1 42 32 À60 CMT 1991 06 11 21 40 À4.86 133.77 18 5.2 327 77 173 CMT 2002 07 02 01 55 À5.41 133.85 32 5.1 21 25 À58 CMT 2003 03 15 14 05 À3.90 134.42 10 5.8 13 50 À153 CMT 2004 10 31 05 14 À4.63 133.71 13 5.3 63 76 12 CMT 2007 05 10 18 37 À5.65 133.86 31 5.1 5 44 À127 CMT 2010 09 26 12 12 À5.31 133.93 09 6.0 24 31 À78 CMT 2010 11 03 11 18 À4.61 134.04 15 6.0 323 80 À180 CMT

a Only events with MW > 5 and a double couple component of greater than 85 are included. Epicenters, origin times and depths between 1964–2007 are taken from the updated catalogue of Engdahl et al. [1998]. Focal mechanisms are plotted in Figure 2 with the compressional quadrants shaded light grey. The centroid depths should not be considered reliable.

Figure 5. Cross section through the Aru Trough. (top) Bathymetry taken from SRTM 30+ which is based on Smith and Sandwell [1997]. (middle) Lithosphere thickness. It is important to note that litho- sphere thickness values have a lateral resolution of 300 km and are calculated for a 2 Â 2 grid (shown in Figure 2). This means that the area surrounding the southern end of the cross section is, on average, underlain by 225 km thick lithosphere, and the N end is surrounded by relatively thin lithosphere (<125 km thick); however the location and steepness of the intervening gradient is poorly constrained. The gradient shown is simply a product of the smoothing function used to calculate the contours. (bottom) Focal mechanisms of earthquakes in the Aru Trough projected onto a vertical plane. Moho depth estimate (white triangle) taken from the reversed seismic refraction profile ending the S Aru Trough [Jacobson et al., 1979].

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of the Aru Trough, and because the lithosphere in this actively deforming area may not be in thermal steady state, we have not attempted to accurately constrain the geotherm here. However, it is likely that in the north the seismic- aseismic transition occurs at 350C as is typical for most continental orogenic belts, whereas in the south the earth- quakes occurring at depths of 25–37 km must be occurring at substantially higher temperatures, perhaps up to the 600C limit which is seen in other areas of thick litho- sphere and is probably associated with anhydrous crustal compositions [Jackson et al., 2008]. [29] When events across the entire Aru Trough are con- sidered, there is a noticeable gap in seismicity at 15–25 km depth (Figure 6). However, the apparent bimodal distribu- tion is deceptive (Figure 5, bottom). The shallow events all occur in the north Aru Trough where seismicity seems likely to limited only to the upper crust. It is therefore only in the south Aru Trough where mid-crustal events might be expected; and here only four earthquakes have reliable cen- Figure 6. Histogram of reliable earthquake centroid depths troid depths. Until more events are studied, it is impossible in the Aru Trough and wider Arafura Sea. Dark blue dot to determine if the apparent gap is simply an artifact of the short time period studied. represents the Moho depth in the southernmost Aru Trough – from a reversed seismic refraction profile. The associated [30] It is important to consider whether the events at 25 vertical bar represents the uncertainty in Moho depth relative 37 km beneath the south Aru Trough lie within the lower to the centroid depths of the two nearby earthquakes shown crust or the UCLM. Fortunately, the reversed seismic by dark blue rectangles. Crustal thickness estimates in the refraction profile described in section 4.1 gives us a rea- Arafura Sea are taken from the seismic refraction profiles sonably good control on the velocity structure and crustal shown in Figure 1a. The unreversed seismic refraction pro- thickness in the vicinity of the southernmost earthquakes in file E of the Aru islands has been omitted. this group. The quoted Moho depth is 32 km, 5 km shal- lower than the best fitting centroid depth of the deepest earthquake in Figure 5, which occurred on the 10th January 1996. The hypocenter of this earthquake was less than 30 km seismicity in this area. The only previously reported earth- from the seismic refraction profile. quake with a well determined centroid depth is a single [31] The uncertainty in the centroid depths of the waveform-modeled strike-slip earthquake in the N Aru waveform-modeled earthquake is around Æ4 km (Appendix A) Trough reported by Abers and McCaffrey [1988]. Table 3 [Engdahl et al., 2006], divided approximately equally lists Mw > 5 earthquakes which could not be satisfactorily between trade-offs in the waveform fit and uncertainties in modeled using long-period body waveform inversion, but the velocity structure. There are also uncertainties in the for which CMT solutions with >85% double-couple com- velocity structure and Moho depth determined by the seismic ponent are available. Their depths are taken from the EHB refraction profile. However, the same average crustal veloc- catalog for events prior to 2007. Earthquakes reported in the ity was used to model the earthquake centroid depth and to EHB bulletin commonly have depth errors of up to  calculate the Moho depth from the refraction profile. If the 15 km [Engdahl et al., 2006] and so these values are not true velocity structure is significantly faster than that reported reliable; but the high double-couple component indicates by Jacobson et al. [1979] then the Moho depth could increase that their mechanisms are likely to be correct. These CMT by up to 3 km, but the centroid depth would also increase earthquakes are plotted in grey in Figures 2 and 5. by a similar amount. The vertical resolution of a boundary in a seismic refraction profile is limited by the vertical Fresnel 4.3. Variations in Centroid Depth Distribution zone, which is <1 km. The uncertainty of the centroid depth [27] There is a clear change in seismogenic thickness relative to the Moho depth is therefore likely to be Æ3 km, along the Aru Trough (Figure 5). In the north Aru Trough which means that the earthquake at 37 km almost certainly seismicity is restricted to the upper 15 km, and further to the occurred within the UCLM, and the adjacent earthquake at NE, onshore Irian Jaya, six body wave modeled earthquakes 25 km was almost certainly in the lower crust. The two events (black mechanisms in Figure 2) indicate a 20 km seismo- further to the north with centroid depths of 30 and 34 km are genic thickness [Abers and McCaffrey, 1988], typical of too close to the Moho to attribute them unequivocally to the many actively deforming continental areas [Chen and lower crust or the upper mantle, especially as they are further Molnar, 1983; Jackson et al., 2008]. In contrast, four from the location with crustal thickness constraint. earthquakes occurred between 25–37 km depth in the south Aru Trough. Figure 1 shows that, on the regional scale, there 5. Discussion is an increase in lithospheric thickness from the north (<125 km) to the south (225 km). 5.1. UCLM Rheology of the Arafura Sea [28] The lateral resolution of the lithospheric thickness [32] The observation that two moderate-sized earthquakes, map is insufficient to locate precisely variations on the scale in 2000 in the Arafura Sea and in 1996 in the south Aru

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Trough, occurred in the UCLM has important implications 5.2. Wider Implications for Continental Rheology for continental rheology. This is only the third region (after [36] We have shown that, beneath the Arafura Sea at least, northern India and the ) where UCLM earthquakes the UCLM appears to be relatively strong, and capable of have been confidently observed and so they provide a seismogenic behavior. This is not in conflict with numerous valuable opportunity to revisit the controls on seismicity and other studies in orogenic regions that found no evidence for long-term strength in the UCLM. We have shown that the UCLM seismicity [e.g. Maggi et al., 2000; Jackson et al.,  2000 Arafura Sea earthquake occurred 25 km below the 2008] because the UCLM in most active regions is hotter Moho in material that has a temperature in the region of than 600C, the likely seismic-aseismic transition tempera-   500 C (section 5.1, Figure 4) and it is likely that the 1996 ture for relatively anhydrous mantle material [McKenzie S Aru Trough earthquake also takes place in material cooler  et al., 2005]. than 600 C (section 5.2). The presence of these earthquakes [37] A more difficult question is whether this rheology is is consistent with the suggestion that the UCLM, like the also typical of ancient continental shields where the UCLM oceanic lithospheric mantle, can deform seismically where it   is cooler than 600 C. After all, there may well be areas the is colder than 600 C[McKenzie et al., 2005]. In this area at UCLM has been significantly hydrated, leading to a lower least, there is no reason to suppose the UCLM is largely seismic-aseismic transition temperature. Here we return to aseismic through being weakened by other processes such as the problem that seismicity in these shields (which have infiltration by metasomatic fluids, which would cause the thick lithosphere, moderate-thin crustal thickness and low seismic-aseismic transition to occur at lower temperatures radiogenic heat production) is so rare. [Maggi et al., 2000]. [38] As the UCLM beneath the northern Australian shield [33] We also show that as the Aru Trough approaches the is seismogenic, and in most ancient shields the lower crust sharp gradient in lithospheric thickness in the south the is also seismogenic (this study) [Sloan et al., 2011; Craig seismogenic thickness rapidly increases before rifting dies et al., 2011], it is likely that the extremely slow deformation out. This effect has previously been identified in the Baikal of ancient shields is the direct consequence of the great and Laptev rifts in [Sloan et al., 2011] and in the E strength of a single seismogenic layer comprising of the African Rift [Craig et al., 2011]. At least one earthquake in entire crust and the UCLM down to the 600C isotherm, the S Aru Trough (which occurred on the 25th July 1988, combined with the effect of a thick, cold, viscous mantle with a centroid depth of 25 km) occurred in the lower crust. root. Confirmation of this in individual shields must wait This suggests that here, as in N India (section 2) [Monsalve until the depths of more earthquakes can be accurately et al., 2006], a seismogenic UCLM is paired with a seis- determined — a formidable task given their rarity — but at mogenic lower crust. least the alternative hypothesis (that the UCLM is generally [34] It is unlikely to be a coincidence that both the UCLM relatively weak, perhaps through rehydration) has been earthquakes identified here are situated near the edge of a shown not to be generally true. region with thick lithosphere. As discussed in section 2, the [39] If the UCLM in shields is anhydrous and seismogenic rarity of mantle earthquakes in areas with thick lithosphere at temperatures below 600C, and its associated lower crust may well be due to the extreme strength, and the resulting  were to have a relatively hydrous composition, then these lack of deformation, where the UCLM is colder than 600 C. areas would be characterized by a laminated or “jelly sand- The edges of these regions will have warmer geotherms than wich” rheology, with strength concentrated in the upper their central areas, and it is possible that significant stresses crust and upper mantle, separated by a weak lower crust. are associated with rapid changes in lithosphere thickness. However, as discussed above, in many ancient shields Both these factors could slightly increase the strain rates near (including the Arafura Sea) the lower crust is also seismo- the edges of thick lithosphere compared to the thickest part, genic, probably because of an anhydrous granulitic compo- resulting in more frequent, though still very rare, UCLM sition [Jackson et al., 2008]. This common feature of ancient earthquakes. shields may be related to the method of their creation. [35] The 1996 south Aru Trough earthquake occurred [40] One suggestion is that cratons form in settings similar where an active rift reaches an area underlain by thick lith- to modern day Tibet, when a continental collision involving osphere and dies out. Stress concentrations are likely at the highly depleted, buoyant lithospheric mantle creates a thick tip of an active rift and GPS measurements show that the stable lithospheric root [McKenzie and Priestley, 2008]. This Aru Trough as a whole is deforming with a relatively high setting also initially produces in a thick crustal layer. The  Â À15 À1 strain rate ( 6 10 s ). Thus it is possible that the relatively high radiogenic heat production and relatively low 1996 earthquake is comparable to the previously described thermal conductivity of the crustal rocks then result in a UCLM seismicity, in N India where the Himalayan ramp- thermal blanketing effect, which leads to very high tem- and-flat system bends and unbends (section 2) [Monsalve peratures in the middle-to-lower crust. This part of the crust et al., 2006; Priestley et al., 2008], and beneath the Andes then undergoes granulite-grade metamorphism, or even where it is associated with the subhorizontal subduction of melting and granite formation, resulting in both dehydration an oceanic slab (section 2) [Emmerson, 2007]. In all three of and the extraction and upward migration of incompatible these settings UCLM seismicity appears to be associated radiogenic heat-producing elements, which are subsequently with elevated strain rates. However, no such claim can be removed through erosion. This leaves behind a thick, made about the 2000 Arafura Sea earthquake which is buoyant, depleted mantle lithosphere, and an anhydrous located on the GPS-defined stable Australian plate [Bock crust with low radiogenic heat production [McKenzie and et al., 2003] in an area with almost no recorded seismic- Priestley, 2008]. The combination produces a single poten- ity, and therefore must be deforming very slowly. tially seismogenic layer including both the entire crust and a

11 of 13 B05402 SLOAN AND JACKSON: UCLM EARTHQUAKES BENEATH THE ARAFURA SEA B05402 significant amount of UCLM; but the great strength of such Engdahl, R., J. Jackson, S. Myers, E. Bergman, and K. Priestley (2006), a rheology means that it is unlikely to deform at all, except Relocation and assessment of seismicity in the Iran region, Geophys. J. Int., 167(2), 761–778. where it is weakest near its edges. Fichtner, A., B. Kennett, H. Igel, and H. Bunge (2009), Full seismic wave- form tomography for upper-mantle structure in the Australasian region using adjoint methods, Geophys. J. Int., 179(3), 1703–1725. 6. Conclusion Fishwick, S., M. Heintz, B. Kennett, A. Reading, and K. Yoshizawa (2008), Steps in lithospheric thickness within eastern Australia, evidence from [41] Earthquakes in continental mantle lithosphere are surface wave tomography, Tectonics, 27, TC4009, doi:10.1029/ 2007TC002116. rare. This study shows the importance of studying the few Forsyth, D. (1985), Subsurface loading and estimates of the flexural rigidity that do occur, as they provide special opportunities to of continental lithosphere, J. Geophys. Res., 90(B14), 12,623–12,632. investigate lithosphere rheology. The two studied here were Harte, B., R. Hunter, and P. Kinny (1993), Melt geometry, movement and crystallization, in relation to mantle dykes, veins and metasomatism, Phi- both in thick lithosphere associated with the Australian –  los. Trans. R. Soc. A, 342,1 21. shield, one of which was substantially ( 25 km) below the Hirth, G., and D. Kohlstedt (2003), Rheology of the upper mantle and the Moho in a region of very low background strain rate. Both mantle wedge: A view from the experimentalists, in Inside the Subduction occurred in mantle that is likely to be 600C or colder, a Factory, Geophys. Monogr. Ser., vol. 138, edited by J. Eiler, pp. 83–106, AGU, Washington, D. C., doi:10.1029/138GM06. feature also of seismicity in oceanic mantle lithosphere. Jackson, J., D. McKenzie, K. Priestley, and B. Emmerson (2008), New They suggest that the very low level seismicity, and manifest views on the structure and rheology of the lithosphere, J. Geol. Soc., strength, of the ancient continental shields can be attributed 165(2), 453–465. to a thick potentially seismogenic layer that includes both Jacobson, R., G. Shor, R. Kieckhefer, and G. Purdy (1979), Seismic refrac-  tion and reflection studies in the Timor-Aru Trough system and Austra- the lower crust and the upper mantle colder than 600 C — lian , Mem. Am. Assoc. Pet. Geol., 29, 209–222. though the resulting strength of that layer is so great that Jongsma, D., W. Huson, J. Woodside, S. Suparka, T. Sumantri, and such regions are likely to deform hardly at all, except near A. Barber (1989), Bathymetry and geophysics of the Snellius-II triple junction and tentative seismic stratigraphy and neotectonics of the north- their edges where they are hotter and weaker. ern Aru Trough, Neth. J. Sea Res., 24(2–3), 231–250. Kirby, J., and C. Swain (2009), A reassessment of spectral Te estimation in [42] Acknowledgments. I would like to thank D. McKenzie and continental interiors: The case of , J. Geophys. Res., 114, K. Priestley for providing the lithospheric thickness map, D. McKenzie B08401, doi:10.1029/2009JB006356. for calculating the geotherms shown in Figure 4, and G. Hirth and three Larson, K., R. Bürgmann, R. Bilham, and J. Freymueller (1999), Kinemat- anonymous reviewers for their helpful comments. ics of the India- collision zone from GPS measurements, J. Geo- phys. Res., 104(B1), 1077–1093. Mackwell, S., M. Zimmerman, and D. Kohlstedt (1998), High-temperature deformation of dry diabase with application to tectonics on Venus, References J. Geophys. Res., 103, 975–985. Abers, G., and R. McCaffrey (1988), Active deformation in the New Maggi, A., J. Jackson, D. McKenzie, and K. Priestley (2000), Earthquake Guinea fold-and-thrust belt: Seismological evidence for strike-slip fault- focal depths, effective elastic thickness, and the strength of the continen- ing and basement-involved thrusting, J. Geophys. Res., 93(B11), tal lithosphere, Geology, 28, 495–498. 13332–13354. McCaffrey, R. (1988), Active tectonics of the eastern Sunda and Banda Bock, Y., L. Prawirodirdjo, J. Genrich, C. Stevens, R. McCaffrey, arcs, J. Geophys. Res., 93(B12), 15,163–15,182. C. Subarya, S. Puntodewo, and E. Calais (2003), Crustal motion in Indo- McCaffrey, R. (1989), Seismological constraints and speculations on Banda nesia from global positioning system measurements, J. Geophys. Res., Arc tectonics, Neth. J. Sea Res., 24(2–3), 141–152. 108(B8), 2367, doi:10.1029/2001JB000324. McCaffrey, R., P. Zwick, and G. Abers (1991), SYN4 program, IASPEI Boettcher, M., G. Hirth, and B. Evans (2007), Olivine friction at the base of Software Library, 3,81–166. oceanic seismogenic zones, J. Geophys. Res., 112, B01205, doi:10.1029/ McCue, K., and M. Michael-Leiba (1993), Australia’s deepest known earth- 2006JB004301. quake, Seismol. Res. Lett., 64, 201–205. Bowin, C., G. Purdy, C. Johnston, G. Shor, L. Lawver, H. Hartono, and McKenzie, D. (2010), The influence of dynamically supported topography P. Jezek (1980), Arc-continent collision in Banda Sea region, AAPG on estimates of Te, Earth Planet. Sci. Lett., 295(1–2), 127–138. Bull., 64(6), 868–915. McKenzie, D., and D. Fairhead (1997), Estimates of the effective elastic Cardwell, R., and B. Isacks (1978), Geometry of the subducted lithosphere thickness of the continental lithosphere from Bouguer and free air gravity beneath the Banda Sea in eastern from seismicity and fault anomalies, J. Geophys. Res., 102(12), 27,523–27,552. plane solutions, J. Geophys. Res., 83(B6), 2825–2838. McKenzie, D., and K. Priestley (2008), The influence of lithospheric thick- Charlton, T., et al. (1991), Geology of the Kai Islands: Implications for the ness variations on continental evolution, Lithos, 102,1–11. evolution of the Aru Trough and Weber Basin, Banda Arc, Indonesia, McKenzie, D., J. Jackson, and K. Priestley (2005), Thermal structure of Mar. Pet. Geol., 8(1), 62–69. oceanic and continental lithosphere, Earth Planet. Sci. Lett., 233(3–4), Chen, W.-P., and P. Molnar (1983), Focal depths of intracontinental and 337–349. intraplate earthquakes and their implications for the thermal and mechan- Milsom, J., et al. (1996), Extension, collision and curvature in the eastern ical properties of the lithosphere, J. Geophys. Res., 88, 4183–4214. Banda arc, Geol. Soc. Spec. Publ., 106(1), 85–94. Clitheroe, G., O. Gudmundsson, and B. Kennett (2000), The crustal thick- Monsalve, G., A. Sheehan, V. Schulte-Pelkum, S. Rajaure, M. Pandey, and ness of Australia, J. Geophys. Res., 105(6), 13,697–13,714. F. Wu (2006), Seismicity and one-dimensional velocity strcture of the Craig, T., J. Jackson, K. Priestley, and D. McKenzie (2011), Earthquake Himalayan collision zone: earthquakes in the crust and upper mantle, distribution patterns in : Their relationship to variations in litho- J. Geophys. Res., 111, B10301, doi:10.1029/2005JB004062. spheric and geological structure, and their rheological implications, Geo- Pérez-Gussinyé, M., and A. Watts (2005), The long-term strength of phys. J. Int., 185(1), 403–434. and its implications for plate-forming processes, Nature, 436(7049), Curray, J., G. Shor Jr., R. Raitt, and M. Henry (1977), Seismic refraction 381–384. and reflection studies of crustal structure of the eastern Sunda and western Pérez-Gussinyé, M., M. Metois, M. Fernández, J. Vergés, J. Fullea, and Banda arcs, J. Geophys. Res., 82(17), 2479–2489. A. Lowry (2009), Effective elastic thickness of Africa and its relationship Debayle, E., and B. L. N. Kennett (2000), The Australian continental upper to other proxies for lithospheric structure and surface tectonics, Earth mantle: Structure and deformation inferred from surface waves, J. Geo- Planet. Sci. Lett., 287(1–2), 152–167. phys. Res., 105, 25,423–25,450. Pilidou, S., K. Priestley, O. Gudmundsson, and E. Debayle (2004), Upper Denlinger, R. (1992), A revised estimate for the temperature structure of the mantle S-wave speed heterogeneity and anisotropy beneath the North oceanic lithosphere, J. Geophys. Res., 97(B5), 7219–7222. Atlantic from regional surface wave tomography: The Iceland and Azores Emmerson, B. (2007), The relationship between intraplate earthquakes and plumes, Geophys. J. Int., 159(3), 1057–1076. temperature, PhD thesis, Univ. of Cambridge, Cambridge, U. K. Priestley, K., and E. Debayle (2003), Seismic evidence for a moderately Engdahl, E., R. van der Hilst, and R. Buland (1998), Global teleseismic thick lithosphere beneath the Siberian Platform, Geophys. Res. Lett., earthquake relocation with improved travel times and procedures for 30(3), 1118, doi:10.1029/2002GL015931. depth determination, Bull. Seismol. Soc. Am., 88(3), 722–743.

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Priestley, K., and D. McKenzie (2006), The thermal structure of the litho- Untung, M. (1985), Subsidence of the aru trough and the Aru Island, Irian sphere from shear wave velocities, Earth Planet. Sci. Lett., 244, 285–301. Java, Indonesia, Tectonophysics, 112(1–4), 411–422. Priestley, K., E. Debayle, D. McKenzie, and S. Pilidou (2006), Upper man- van der Hilst, R., B. Kennett, and T. Shibutani (1998), Upper mantle struc- tle structure of eastern Asia from multimode surface waveform tomogra- ture beneath Australia from portable array deployments, in Structure and phy, J. Geophys. Res., 111, B10304, doi:10.1029/2005JB004082. Evolution of the Australian Continent, Geodyn. Ser., vol. 26, edited by Priestley, K., J. Jackson, and D. McKenzie (2008), Lithospheric structure J. Braun et al., pp. 39–57, AGU, Washington, D. C., doi:10.1029/ and deep earthquakes beneath India, the Himalaya and southern Tibet, GD026p0039. Geophys. J. Int., 172(1), 345–362. Watts, A., J. Bodine, and M. Steckler (1980), Observations of flexure and Ritsema, J., and H. van Heijst (2000), New seismic model of the upper man- the state of stress in the oceanic lithosphere, J. Geophys. Res., 85(B11), tle beneath Africa, Geology, 28,63–66. 6369–6376. Rudnick, R. L., and D. M. Fountain (1995), Nature and composition of Whittington, A., A. Hofmeister, and P. Nabelek (2009), Temperature- the continental crust: A lower crustal perspective, Rev. Geophys., 33, dependent thermal diffusivity of the EarthÕs crust and implications for 267–309. magmatism, Nature, 458(7236), 319–321. Rynn, J., and I. Reid (1983), Crustal structure of the western Arafura Sea Wiens, D., and S. Stein (1983), Age dependence of oceanic intraplate from ocean bottom seismograph data, Aust. J. Earth Sci., 30(1), 59–74. seismicity and implications for lithospheric evolution, J. Geophys. Sloan, R., J. Jackson, D. McKenzie, and K. Priestley (2011), Earthquake Res., 88(8), 6455–6468. depth distributions in central Asia, and their relations with lithosphere Zwick, P., R. McCaffrey, and G. Abers (1994), MT5 program, IASPEI thickness, shortening and extension, Geophys. J. Int., 185(1), 1–29. Software Libr., vol. 4, Int. Assoc. of Seismol. and Phys. of the Earth’s Smith, W., and D. Sandwell (1997), Global sea floor topography from Inter., Trieste, Italy. satellite altimetry and ship depth soundings, Science, 277(5334), 1956–1962. ’ J. A. Jackson and R. A. Sloan, Bullard Laboratories, Department of Tandon, K., J. Lorenzo, and G. O Brien (2000), Effective elastic thickness Earth Sciences, University of Cambridge, Cambridge CB3 0EZ, UK. of the northern Australian continental lithosphere subducting beneath the ([email protected]) Banda orogen (Indonesia): Inelastic failure at the start of continental sub- duction, Tectonophysics, 329(1–4), 39–60.

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