Crustal Deformation at the Southernmost Part of the Ryukyu Subduction (East Taiwan) As Revealed by New Marine Seismic Experiments

TECTO-125443; No of Pages 21 Tectonophysics xxx (2012) xxx–xxx

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Crustal deformation at the southernmost part of the Ryukyu subduction (East Taiwan) as revealed by new marine seismic experiments

Thomas Theunissen a,⁎,1, Serge Lallemand a,f, Yvonne Font b, Stéphanie Gautier a,f, Chao-Shing Lee c,f, Wen-Tzong Liang d,f, Francis Wu e, Théo Berthet a a Geosciences Montpellier, University of Montpellier 2, CNRS, France b University of Nice Sophia-Antipolis, Institut de Recherche pour le Développement (UR 082), Observatoire de la Côte d'Azur, Géoazur, Villefranche-Sur-Mer, France c NTOU, Keelung, Taiwan d IES, Academia Sinica, Taipei, Taiwan e Department of Geological Sciences and Environmental Studies, Binghamton University, NY, USA f LIA (Associated International Laboratory) ADEPT, France-Taiwan article info abstract

Article history: The southernmost part of the Ryukyu subduction, where the Philippine Sea Plate is subducting under the Received 10 June 2011 Eurasian Plate, is known to be a very seismically active region of transition from a north-dipping subduction Received in revised form 19 March 2012 along the Ryukyu subduction to an ~SE–NW collision along the Taiwanese orogenic wedge. In this paper, we Accepted 11 April 2012 will focus on the Ryukyu forearc area close to Taiwan where the deformation is paroxysmal. In order to de- Available online xxxx cipher the nature of the seismic deformation in this region, a three month passive experiment, combining 22 Ocean Bottom Seismometers and 51 onland stations, has been led. Starting from an a-priori heterogeneous Keywords: Passive experiment RATS (Ryukyu Arc: model, we have obtained 801 well-located earthquake hypocenters, a precise P-wave tomography model Tectonics and Seismology) and 14 focal mechanisms. The seismicity along the Ryukyu forearc is mainly located not only in the vicinity Collision-Subduction transition east of Taiwan of the Interplate Seismogenic Zone (ISZ) but also within both the subducting PSP and the overriding plate. Ryukyu forearc Seismicity within the upper-plate is essentially localized east of Nanao basin where NW–SE extension occurs, Absolute earthquake location and northwest of the Hoping basin where strike-slip dominates. As revealed by both the P-wave velocity Focal mechanisms structure and the newly derived seismicity, we argue that a sub-vertical step offsetting the subducting PSP 3D approach (a priori 3D P-wave around 10 km may support the presence of a trench-parallel tear. The PSP also undergoes extension in its velocity model) upper part that is probably caused by buckling and slab pull. The P-wave velocity structure reveals three other major features: (1) a continuity between the Central Range and the Ryukyu Arc with a shallower Moho (~30 km depth) between ~122.3°N and ~122.5°N along the Ryukyu Arc, (2) high P-wave velocities along the eastern side of the Central Range and, (3) two bodies with similar high crustal velocities (6.5– 7.0 km/s) at 12–18 km depths, embedded within the Ryukyu arc basement, just north of Hoping Basin and north of the Nanao Basin. © 2012 Elsevier B.V. All rights reserved.

1. Introduction 1993; Peterson et al., 1984; Ruff and Kanamori, 1983). The southernmost part of the Ryukyu Subduction system, in particular between The Ryukyu Subduction zone between Kyushu Island (Japan) and Taiwan and the Gagua Ridge, is a region of transition between an obli- Taiwan is known to have generated only a few large thrust interplate que subduction (Ryukyu) and an active collision (Taiwan orogen) earthquakes during the period of 1900–2010 at its two extremities (Kao et al., 1998b). This region results from the meeting and ~5 My (Heuret et al., 2011; Shiono et al., 1980), i.e., in the northern part off- evolution along the South China passive margin, of two subductions shore SW Japan and close to Taiwan west of 124°E (Fig. 1a and b). No with opposite polarity: the east-dipping Eurasian Plate (EP) of the evidences of major MW >8.0 historical shallow earthquakes have been Manila Subduction and the northwest-dipping Philippine Sea Plate reported (Abe, 1981; Kanamori, 1986) suggesting that the plate inter- (PSP) of the Ryukyu Subduction, respectively to the south and face is seismically weakly coupled (Kanamori, 1971; Pacheco et al., northeast of Taiwan (Chai, 1972; Lallemand et al., 2001; Suppe et al., 1984; Teng, 1990; Tsai et al., 1977; Wu, 1978; Yen, 1973)(Fig. 1). Off- shore, east of Taiwan and in the transitional domain between subduction and collision along the Ryukyu forearc, the high level of seismicity ⁎ Corresponding author. E-mail address: [email protected] (T. Theunissen). characterizes a paroxysmal deformation (Chen et al., 2009; Hsu, 1961; 1 Now at IRAP, OMP, UPS3, Toulouse, France. Kao et al., 1998b; Tsai, 1986; Wang, 1998; Wang and Shin, 1998; Wu,

Fig. 1. Tectonic settings. a: Thrust events along the Ryukyu subduction zone during the period 1977–2009 and major earthquakes (MS >7) during the period 1900–2010. b: Close- up view of the southernmost part of the Ryukyu subduction. Known slip area of Slow Slip Events (SSE) and two larger earthquakes known (1771, 1920) are also added in purple. c: Structural and kinematic context. GPS velocity ﬁeld comes from Hsu et al. (2009). CP: Coastal Plain, DF: Deformation Front, WF: Western Foothill, LFS: Lishan Faults system, CeR: Central Range, LVF: Longitudinal Valley Fault, CoR: Coastal Range, HB: Hoping Basin, NB: Nanao Basin, YF: Yaeyama Fault, EYF: East Yaeyama Fault. S102 is the reference of the GPS station on Lanyu Island supposed to represent the velocity of the non-deformed PSP. (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article.)

1978)(Fig. 2). There, more than 10 major events with magnitude be- sources. This work carries out a 3D approach that uses an a priori 3D tween 7 and 8 occurred since the beginning of the last century but the P-wave velocity model and 3D hypocenter determination as initial source of each of them is not known (Theunissen et al., 2010). Efforts inputs to perform the tomographic inversion and focal mechanism de- to image this area, mainly based on active seismic offshore (reﬂection termination. East of Taiwan, some studies have provided earthquake and refraction) or passive seismic using onland seismic stations, have location and seismic tomography to image and understand the seismic led to divergent interpretations. Previous studies were non-conclusive deformation pattern offshore. Many of them used a combination of regarding the geometry and mechanism of offshore active faults (e.g., seismic stations located on Taiwanese and Japanese islands in order Font and Lallemand, 2009). Responses to key questions are still out- to better highlight the area offshore (Chou et al., 2006, 2009; Font and standing: how is the deformation accommodated offshore northeast Lallemand, 2009; Font et al., 2004; Hsu et al., 2001; Kao and Rau, Taiwan? What are the type and the origin of the seismicity along the 1999; Lin et al., 2004; Wu et al., 2008, 2009b). However, in all these Ryukyu forearc? What is the nature of the forearc domain? Is the sub- studies, no Ocean Bottom Seismometer (OBS) has been used to improve duction interface close to Taiwan likely to generate a major the azimuthal coverage, resulting in poorly resolved crustal structures earthquake? (especially at shallow depth) and large uncertainties on hypocenter po- In this study, before to answer these questions, we aim to image sition (especially the depth). Only Lin et al. (2007) have used a combi- the seismic wave velocity structure to describe tectonic features and nation of OBS deployed during 12 days in the Okinawa basin and to characterize the geometry and deformation type of offshore active permanent stations to study the micro-seismicity in the back-arc faults in order to contribute to the determination of large event basin. To improve azimuthal coverage and P-wave velocity structure

Vertical N-S cross-sections Map : seismicity 1991-2008 (M>~3.5) Distance (km) 0 50 100 150 200 121˚ 122˚ 123˚ 124˚ 0 SCSC 19171917 −20 19511951 19221922 19631963 25˚ −40 19471947 −60 19781978 19221922 −80 Depth (km) −100 19631963 SCSC 19221922 1 19201920 19661966 −120 0 50 100 150 200 20022002 0 HCHC NCNC 24˚ 19201920 −20 20022002 HCHC −40 19511951 −60

Depth (km) −80 −100 2 23˚ −120 19781978 12345 0 50 100 150 200 0 −20 HCHC −40 −60 22˚ Depth (km) −80 −100 0 10 20 30 40 50 60 70 80 90 100 110 120 130 3 Focal depth (km) −120 0 50 100 150 200 0 50 100 150 200

0 0 19171917 −20 19661966 −20 NCNC −40 19221922 −40 −60 −60 Depth (km)

Depth (km) −80 −80 19471947 −100 4 −100 5 −120 −120

Fig. 2. Seismicity recorded east of Taiwan by permanent CWB and JMA seismic networks (1991–2008, ML >~3.5). 3D location procedure from Theunissen et al. (2010). Three seismicity clusters are visible along the Ryukyu forearc: the Suao cluster (SC), the Hoping cluster (HC) and the Nanao cluster (NC). Major earthquakes (MW >7) are represented by stars (Theunissen et al., 2010). PSP slab roof in red on vertical sections is from Font et al. (2003). (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article.)

offshore along the Ryukyu forearc, a passive and an active experiment 2. Geodynamic and tectonic context east of Taiwan have been performed east of Taiwan in 2008 and 2009 called RATS1 and RATS2 respectively (RATS for Ryukyu Arc Tectonics and 2.1. Convergence accommodation Seismology). During the three months of RATS1, 15 OBS were deployed above the Ryukyu forearc and 24 OBS were deployed during RATS2 In the vicinity of Taiwan, the convergence rate is about 8 cm/yr be- along a line crossing through the Ryukyu margin (Fig. 3)(see tween the PSP and the South China Block (SCB) (~Eurasian Plate, EP). Klingelhoefer et al., this issue). The passive seismic network has been West of the Gagua Ridge, the convergence between the PSP and the combined with 7 OBSs from the TAIGER experiment (TAiwan Integrated Ryukyu Arc shows an important obliquity between 40° and 60°. Be- Geodynamics Research) and 51 inland seismic stations. Results from the cause the Ryukyu Arc moves southward in response to the opening RATS2 active experiment have been used to update the 3D a-priori ve- of the back-arc Okinawa basin, the convergence rate between the Ryu- locity model from Font et al. (2003). This 3D a-priori velocity model and kyu Arc and the PSP has been estimated to 107 mm/yr (Lallemand and results from the passive experiment including earthquake locations, Liu, 1998). According to new GPS data (Nakamura, 2004; Nishimura local earthquake tomography and focal mechanisms determinations et al., 2004), we re-evaluate the convergence rate to ~140 mm/yr at will be presented in this paper. This study highlights the deformation 123°E along a direction ~337°N. The obliquity is thus reappraised be- along the Ryukyu margin east of Taiwan and improves our understand- tween 30° and 50°. Despite of this important obliquity, the mean ing of the geodynamics in this region. direction of slip vectors of thrust events, 345±12° (Kao et al.,

121˚ 122˚ 123˚ 124˚ 121˚36' 122˚00' 122˚24' 122˚48' 123˚12'

EXTENDED NETWORK ILAILA EGSEGS NEAR FIELD NETWORK 24˚48' TWY TWETWE ANP a b TWS1 NWF ENTENT TAP TWCTWC TAP1 TWB1 25˚ NCU HSN TWA EGS ILA YONAYONA SBCB TWE YHNB ENAENA YOJYOJ NSK ENT TWC NST YNGFYNGF 24˚24' TO01TO01 YNGF HATO EHPEHP P001P001 NNSB ENA YOJ IGKF NNS P002P002 YONA TDCB EHP IRIF NACBNACB P013P013 TO02TO02 P004P004 TWT ISH2 WHF NACB KURO P003P003 CHGB HATS TWDTWD P005P005 TWD P014P014 P007P007 24˚ HWA2 P006P006 ESF HWA2HWA2 P010P010 24˚00' WDT ESL P008P008 P011P011 ESFESF P009P009 EGC P015P015 YUS EHY TO04TO04 P012P012 YULB ESLESL TWF1 TO09TO09 ELD EEGCGC TO10TO10 CHK 23˚36' 23˚ TO11TO11 TWGB TWG TO12TO12 LDU ECL TTN TWH

Stations used in this study 07/19/2008 - 10/22/2008

LAY OBS INSU/CNRS RATS1 BATS 22˚ OBS NTOU F-NET OBS of the RATS2 active seismic experiment OBS TAIGER CWB 05/03/2009 - 05/20/2009 JMA

Fig. 3. Seismic networks used in this study. The extended network (a), composed of 73 stations, is used for travel-time tomography and 3D focal mechanisms determination. The near ﬁeld network (b) includes 38 stations used for the initial location procedure.

1998b), is close to the convergence direction. This suggests that the the collision between the subducted slab of PSP and the root of the Tai- convergence component parallel to the trench is mainly accommodat- wan orogen (Chou et al., 2006; Font et al., 1999; Kao and Jian, 2001; ed along the ISZ and that there is almost no need of strain partitioning. Wang, 2005; Wang et al., 2004). This deep collision may cause impor- This suggestion is in agreement with the lack of evidence of a forearc tant seismic deformation. Physical models proposed by Chemenda et basement sliver (Chemenda et al., 2000). If strain partitioning does al. (1997, 2001) have suggested that an incipient westward dipping exist, the remaining strain is rather accommodated by a reconﬁgura- subduction (trending ~ N–S) could develop offshore east of Hualien. tion of the local plate kinematics in relation with considerable intra- Later, Malavieille et al. (2002), Font (2002) and Bos et al. (2003) plate deformation in this region (Chiao et al., 2001; Kao et al., 1998b; have found some evidences of such intra-PSP reverse fault, trending Lallemand et al., 1997). The only evidence of partitioning within the N–S, east of the Luzon arc and beneath the Ryukyu forearc. Neverthe- upper-plate occurs along the Yaeyama Fault near Taiwan (Fig. 1c), less, no clear manifestations of westward PSP subduction beneath Tai- south of the thrust events, and within the sedimentary accretionary wan were found. To accommodate the increasing stress in the wedge (Dominguez et al., 1998; Lallemand et al., 1999). northernmost domain of the collision, i.e. Hualien region, a NW–SE In addition to the high subduction obliquity, the subduction seg- tear fault within the PSP oceanic lithosphere has also been proposed ment between the Gagua Ridge and the Taiwan Island is also inﬂu- (Lallemand et al., 1997; Malavieille et al., 2002). However, such struc- enced by the collision between the Luzon volcanic Arc (carried by ture has never been clearly imaged in this region. the PSP and originated from the Manila Subduction zone) and the Eurasian passive margin that causes the Taiwanese orogenic wedge 2.2. A privileged geodynamical model (Biq, 1972; Chai, 1972; Ho, 1986; Malavieille et al., 2002). Along the major collision suture (eastward dipping thrust of the Longitudinal From kinematic plate reconstruction (Lallemand et al., 2001), it ap- Valley Fault system, LVF) (Biq, 1965; Hsu, 1976), the shallow northern pears that the 400 km-long southernmost segment of the Ryukyu sub- part of the Coastal Range is characterized by a much smaller conver- duction (from Taiwan to Miyako island) is most probably neo-formed, gence rate north of 23.5°N than in the south (Fig. 1c), and a larger associated with the westward propagation of PSP slab. Recent geolog- left-lateral component. The east-dipping deeper part of the LVF ex- ical history such as the subduction of the Luzon Arc beneath the south- tends offshore (Chung et al., 2008; Kuochen et al., 2004; Shyu et al., ernmost Ryukyu Arc should be taken into account to understand the 2005) with a possible inversion to the north of the LVF (north of mechanical and seismological behavior of this area (Lallemand et al., 23.5°N) into a west-dipping shallow thrust (e.g. Kim et al., 2006; 2001; Teng, 1990). Such kinematic evolution has large implications Rau et al., 2007). Increasing of lateral E–W compression within the for the forearc domain deformation and its current seismicity. subducting PSP is interpreted as a result of the transmitted strain originated from the collision (e.g., Kao and Chen, 1991; Kao et al., 1998b). 2.3. Geomorphological and structural context In response to this lateral compression, buckling of the subducting PSP and possible west-dipping slivering occurred in this area (Bos et al., From the Ryukyu trench toward the north, the subduction domain 2003; Chou et al., 2006; Font et al., 1999; Malavieille et al., 2002; is organized with a large accretionary prism, a series of three sedimen- Wang et al., 2004). Some authors suggested that the PSP was also tary forearc basins lying above the Ryukyu Arc basement, the Ryukyu buckling at depth in response to the E–W compression generated by Arc itself and the back-arc Okinawa Trough. The whole system trends

NW–SE, contrasting with the NE–SW main direction of the Ryukyu The earliest one (MW of 8) occurred in 1771 east of Gagua Ridge at Subduction between Miyako Island and Japan. Intense crustal defor- shallow depth (0–20 km) and has been reported as an earthquake mation affects the Ryukyu margin. To the north, E–W trending normal responsible for one of the most devastating Japanese tsunami faults in relation with the opening of Okinawa Trough are located (Nakamura, 2009b; Nakata and Kawana, 1995). There, the seismicity south of the Ilan Plain and offshore, in the continuation of the Lishan rate is relatively low and slow slip events (observed since 1997) Fault (Fig. 1c) (Hsu et al., 1996; Lai et al., 2009; Lin et al., 2009; Rau occur biannually on the interplate seismogenic zone (ISZ), at depths et al., 2008; Sibuet et al., 1987, 1998). South of the Okinawa Trough, between 20 and 40 km and over a length of 180 km. Average equiva- near 122.25°E, a N–S strike-slip fault system is proposed to accommo- lent Mw are estimated at 6.6 each time (Heki and Kataoka, 2008) date the southward displacement of the Ryukyu Arc (Okinawa Trough (Fig. 1b). In 1920, a MW7.7 event (Theunissen et al., 2010) (estimated opening and trench retreat) relatively to Taiwan (Fig. 1c) (Lallemand at MW7.8 by Pacheco and Sykes, 1992 and MS8.1 by Wang and Kuo, and Liu, 1998; Wu, 1978) even though fault traces on the sea-floor 1995) occurred closer to Taiwan, at shallow depth (b20 km). The or seismicity distribution do not clearly demonstrate this fault system. revisited hypocenter is in agreement either with the downdip limit This fault system would separate the Hoping (and Suao) sedimentary of an E–W trending splay-fault or with the ISZ (Theunissen et al., basins from the Hoping basement high (Fig. 1c, Font et al., 2001). The 2010), but doubts still exist on which fault triggered this large event. Hoping Basin contains more than 9 km of sedimentary thickness and Neighboring the 1920 event (at epicentral distanceb50 km) three the nature of the basin floor is still enigmatic. The Hoping Basement other significant earthquakes occurred in 1922 (Sept. 1st, Mw7.4), High is proposed to be made of Ryukyu Arc basement uplifted by the 1963 (Mw7.2) and 2002 (Mw7.1). The available focal mechanisms subduction (or the underplating) of some local asperity. East of the (for two earthquakes only) indicate typical thrust events. The latest Hoping Basement High, the Nanao sedimentary basin develops over one (on March 31, 2002) has been associated with an important a rough Ryukyu Arc basement that suggests crustal deformation. The after slip that lasted 5 years (Fig. 1b), located between 30 and 60 km

Nanao Basin ends on a second uplifted basement high (Nanao Base- depth, producing a cumulative MW7.4 (Nakamura, 2009a). The two ment High) interpreted as the result of the Gagua Ridge subduction remaining signiﬁcant events occurred in 1922 (Sept. 14th, Mw7.1) (Dominguez et al., 1998; Konstantinou et al., 2011). Both basement and 1966 (Mw7.5) affecting the overriding Ryukyu margin. The 1966 highs are affected by ~N–S trending normal faults. event was associated with strike-slip deformation (Fig. 2). At the junction between the Taiwanese orogenic wedge, the Ryukyu In the same area, east of Taiwan and west of the Gagua Ridge, the Arc and the PSP, approximately above the Hoping Basin, a triple junc- seismicity rate is very high, especially beneath the arc and forearc tion where evidence of few strike-slip and normal faults, limited in (Fig. 2)(Chen et al., 2009; Hsu, 1961; Kao et al., 1998b; Tsai, 1986; space, accommodates extension and rotation in the crust (Angelier et Wang, 1998; Wang and Shin, 1998; Wu, 1978). Instrumental seismic- al., 2009; Font, 2002; Hou et al., 2009; Wu et al., 2009a). ity is distributed, in part within the subducting Philippine Sea Plate (i.e. Benioff zone), in part along the ISZ, and, in part within several ac- 2.4. Seismicity tive and shallow clusters that affect the overriding margin. To the north (24.5°N to 25°N), shallow deformation is clearly associated As close to Japan, historical instrumental seismicity of the southern with the back-arc activity. Over the forearc basin, the Hoping cluster part of the Ryukyu Subduction zone is characterized by frequent (noted HC in Fig. 2) aligns with the Hoping Canyon (along an ~E–W di-

Mw b7.5 seismic events (Shiono et al., 1980; Theunissen et al., 2010) rection) and earthquakes are mainly distributed within the overriding (Figs. 1 and 2). Prior this period, historical large seismic episodes, margin, from the ISZ to superficial depth (Font et al., 2004). This activ- last one ~2 ky ago certainly associated with a mega-tsunami (Nakata ity has been associated to either a splay-fault or a high-angle back- and Kawana, 1995), have been revealed from paleo-seismological thrust fault (Font and Lallemand, 2009). Repeating earthquakes com- investigations in Japanese islands east of Miyako but nothing close to patible with the interplate mega-thrust and located in the HC have Taiwan (Inagaki et al., 2007; Ota and Omura, 1992; Pirazzoli and been identified (Igarashi, 2010). East of Nanao Basin, the shallow seis- Kawana, 1986; Sugihara et al., 2003). micity, called Nanao cluster (noted NC in Fig. 2), may be explained by On the southern part of the Ryukyu Subduction, near Taiwan, N–S trending normal faulting visible on both seismic lines and ba- seven significant earthquakes have been reported (Figs. 1 and 2). thymetry (Fig. 1c) (Lallemand et al., 1999) in agreement with CMT

Table 1 Instruments and data used in this study. Number of picked phases corresponds to the number used in the 3D tomographic inversion. SP: short-period, BB: broadband. INSU/CNRS: Institut National des Sciences de l'Univers/Centre National de la Recherche Scientiﬁque; NTOU: National Taiwan Ocean University; All OBSs had four components (2 vertical including the hydrophone and two horizontal). JMA and CWB arrival-times and polarities have been manually picked by routine analysis at each seismological center while all other data have been obtained by manual picking.

Network Country No. Instruments Arrival-times Number of polarities used

P S P S Ampl

Earthquake locations and tomography Offshore INSU/CNRS France 12 SP/L-28LB Sercel 6241 6549 164 182 250 801/1035 events NTOU Taiwan 3 SP/Microbs Ifremer 354 82 5 –– TAIGER Taiwan/USA 7 BB/L4, CMG, KWB 807 1252 82 –– Onland BATS Taiwan 1 BB/STS-2 318 324 13 16 23 F-NET Japan 1 BB/STS-2 96 74 4 8 12 CWB Taiwan 13 SP/S13 1522 1787 31 –– JMA Japan 1 SP/E93 51 50 0 –– Total 38 9389 10,118 299 206 285 Focal mechanisms determination Distant stations BATS Taiwan 7 BB/STS-2,CMG 1139 1077 57 47 65 14 events onland CWB Taiwan 22 SP/S13 1113 2586 50 –– JMA Japan 5 SP/E93 179 206 0 –– F-NET Japan 1 BB/STS-2 41 40 3 3 4 Total 35 2472 3909 110 50 69 Total 73 11,861 14,027 409 256 354

focal mechanism of the 2001 event (MW =6.8). However, seismicity linear problem (Eberhart-Phillips and Michael, 1993; Kissling et al., interpretation in terms of tectonic activity is hindered by the uncer- 1995a, 1995b; Thurber, 1992). tainties on hypocenter positions offshore. The a priori 3D P-wave velocity model used in this study is an updated version of the a priori 3D model built by Font et al. (2003). The 2003 model combined a tomographic model onland (Rau and 3. Data Wu, 1995) and an a priori model offshore. The a priori model offshore used all available velocity structure information and interface posi- 3.1. RATS1 experiment and data quality tions from geophysical campaigns. We call “interface” high impedance contrast limits either between sedimentary and crustal structures or The RATS1 passive seismological experiment was deployed during either at the Moho discontinuity. Readers should refer to Font et al. three months from July, 19 to October, 22, 2008 (Fig. 3 and Table 1). (2003) for more details. The RATS network included 12 short-period INSU-CNRS OBSs and 3 Main modifications of the 2003 offshore model concern the imple- short-period NTOU OBSs. For the earthquake location, 5 nearby net- mentation of the topography and the integration of recent active seis- works were combined including: 7 broadband TAIGER OBS from the mic data in the area (RATS2) and slab position. 3D-contours of the top US National OBSIP (e.g. http://taiger.binghamton.edu/)(Wu et al., of the Ryukyu slab result from a combination of data issued from active 2007), 1F-NET broadband seismic station on Yonaguni Island (Okada seismic observations (down to 25 km in depth) (e.g., Klingelhoefer et al., 2004), 1 broadband BATS seismic station (Kao et al., 1998a), et al., this issue) and position of the Wadati–Benioff zone in depth 13 short-period CWB seismic stations (e.g., Shin and Teng, 2001) and from the EHB2 earthquakes location (Engdahl and Villaseñor, 2002; 1 short-period JMA station (Okada et al., 2004). A total of 38 stations Engdahl et al., 1998). The Moho discontinuity is improved for both the surrounding the studied area are included in a unique dataset to define oceanic and the continental lithospheres, i.e. below the Ryukyu arc anearfield network (Fig. 3b). For tomography inversion and focal and the Okinawa back-arc. The top of the oceanic and continental mechanisms determination, we further include 35 regional stations crust (acoustic basement) is well-documented thanks to seismic reflec- from CWB, BATS and JMA seismic networks. The extended regional net- tion lines. These 3D envelopes of main tectonic units (sedimentary work is therefore composed of 73 stations (Fig. 3aandTable 1). layers, crustal basement and upper-mantle) are implemented into a se- OBS position and sensor horizontal component orientation have ries of N–S 2D lines, every 5 km, to interpolate the velocity model, with been precisely determined from active seismic recording. The orienta- a 1×1 km resolution (using RAYINVR program from Zelt and Ellis, 1988; tion of the vertical component and homogeneity of the network in Zelt and Smith, 1992). Then, interpolation along perpendicular lines terms of polarity were also checked. Polarities of broadband stations (each one km) allowed computing the 3D model with an evenly spaced have been verified using one Vanuatu teleseismic event that is ~60° 1×1×1 km grid. The 2003 3D model has been partially modified by away from Taiwan (2008/09/08 18 h52 UTC; M 7.0; depth=135 km). W implementing the result of tomographic inversion of Wu et al. (2009b) For short-period OBSs, polarities have been checked using a regional in the inland Taiwan (Fig. 4). Result of this construction reveals some Sichuan earthquake (2008/10/01 08 h32 UTC; M 5.7; distance=17°). W main differences with the 2003 model in addition to a higher definition: The RATS1 network recorded more than 4000 local events. In this a vertical backstop at the toe of the Ryukyu Arc, upper-mantle velocity study, we measured P and S-wave arrivals for 1300 events. Among below the oceanic crust reduced from 8.0 to 7.8 km/s and a Moho dis- them, 1035 events were located between the east coast of Taiwan continuity beneath Taiwan, defined by a deeper 7.8 km/s isopleth. and the Gagua Ridge, from the Huatung basin to the Ryukyu arc, i.e. The resulting 1×1×1 km3 grid has an origin at 120.9°E–22°N (SW the target area. From this dataset, 801 events were selected for the corner), extends 350 km eastward and 370 km northward, and reaches joint inversion procedure (see Section 4). Coda magnitude (M ) esti- c 200 km depth. In this study, as required by the specificparameteriza- mate of these earthquakes ranges between about 0.5 and 3.9. Local tion of each method (discussed in the next section), two different magnitude (M ) estimated by the CWB gives a maximum of 4.9 within L grids are calibrated for (1) the initial 3D earthquake location process these 3 months. The coda magnitude estimate is often misestimated and (2) the initial 3D velocity model implemented in the tomography according to the M . Because of both an incomplete catalog and a L process. The velocity model used for the initial earthquake location bias in the magnitude estimate, magnitudes will not be further dis- has an origin at 121.24° isopleth 23.3°N, extends 220 km eastward cussed. About 400 events have been detected by our combined seismic and 180 km northward (Fig. 3b), and reaches 102 km depth. It contains network with at least 2 inland stations and 8 OBSs (7 inland stations the near-field network composed of 38 seismic stations. It is defined on and 13 OBSs, in average). The other 635 earthquakes were only an evenly spaced 4×4×1 km3 grid, whereas initial velocity model used recorded by OBSs (at least 4, and 12 in average). 77 earthquakes for the arrival-times tomography has the same horizontal extension were large enough to be located using the combination of CWB and than the initial grid (Fig. 3a) but reaches 132 km depth with an evenly JMA permanent seismic networks. Comparison of hypocenters deter- spaced 10×10×6 km grid. It contains the extended network of 73 sta- mination, with or without OBS, is discussed in Section 5.2. tions. These two models are obtained by averaging the initial grid.

3.2. Update of an a priori 3D velocity model 4. Methods

A signiﬁcant effort of this study concerns the implementation of an In order to better image the seismicity and the type of deformation, entirely 3D approach. As proposed by few authors (Arroyo et al., 2009; we proceed in 3 stages following a 3D approach: Flanagan et al., 2007; Font et al., 2003; Husen and Smith, 2004; Husen (1) retrieve initial earthquake location within the a priori 3D veloc- et al., 2000, 2003), the use of an a priori 3D velocity model that inte- ity model using the MAXI method (Font et al., 2004) (we call grates knowledge about crust and mantle structure and V properties P the resulting dataset MAXI-3D). The MAXI method provides as initial reference for earthquake location and tomography allows an absolute earthquake location (single event procedure) with- improving travel-times estimates and subsequently hypocenter de- in a heterogeneous velocity model. By using this technique, we termination. Our approach considers the coupled hypocenter velocity are able to constrain the focal depth with only the P-wave ar- problem (Crosson, 1976) without using a 1D minimum model rivals and objectively exclude arrival-time outliers during the (Kissling, 1988; Kissling et al., 1994) but within an a priori 3D georea- listic model that represents the southernmost Ryukyu subduction zone. The purpose of the 3D velocity model is to limit errors, both in earthquake location and velocity model, associated with this non- 2 EHB = Engdahl, van der Hilst and Buland.

W Hoping Basin 12 50km 25˚ S Tomography of 3 24˚ Wu et al. [2009]

E 23˚ 122˚ 123˚

~Vertical backstop

P-wave velocity (km/s) 7.8 km/s under the oceanic crust Subducting 2 oceanic crust 3

Fig. 4. View of the a priori 3D P-wave velocity model serving as initial model for initial earthquake location and tomography. The western N–S cross-section (1) cuts through the Taiwan Island, i.e., within the interpolated velocity model of Wu et al. (2009b). The eastern N–S cross-section (2) cuts through the Ryukyu subduction system, i.e., within the offshore a priori 3D velocity model integrating, among others, the RATS2 active seismic experiment results (Klingelhoefer et al., this issue).

location process. Based on MAXI conﬁdence factors, we select computation (Gautier et al., 2006; Latorre et al., 2004) from 962 (among 1035) reliable hypocenters. the initial a priori 3D model and the initial MAXI-3D earth- (2) perform the 3D tomographic inversion with an improved ap- quake location. Hypocenter parameters as well as the P and proach based on an accurate ﬁnite-difference travel-time S-wave velocity models are simultaneously inverted with a

ο ο ο ο 0 ο ο ο ο -30 -60 -90

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Fig. 5. Projected rays distribution. 801 earthquakes (red dots) and 73 stations (blue triangles) are used in the tomography (see Table 1 and Fig. 3). Initial hypocenter determinations results from earthquake location using MAXI technique within the a priori 3D initial velocity model. Histogram represents the depth distribution of all earthquakes. (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article.)

data selection of 801 earthquakes among the 962 of the MAXI- approach and its linearized iterative scheme is given by Latorre et al. 3D dataset (Fig. 5, Table 1). (2004), Vanorio et al. (2005) and Gautier et al. (2006).Inthisap- (3) derive focal mechanisms using take-off angles and azimuth proach, we compute travel-times by solving the Eikonal equation with provided by the tomography result. Polarities and amplitudes a finite-difference algorithm (Podvin and Lecomte, 1991) and rays are of P- and S-waves are computed using the program FOCMEC obtained using an aposterioriray-tracing method that is based on time (Snoke et al., 1984). gradients. More precise travel-times and partial derivatives, both for slowness fields and for hypocenter parameters, are evaluated along 4.1. MAXI method the ray paths. Finally, the scaled and weighted linear system is solved by means of the LSQR method (Paige and Saunders, 1982) and both MAXI provides an absolute earthquake location defined by the the velocity models and the hypocenter parameters are updated. maximum intersection number of hyperbolic Equal Differential Time As proposed by some authors (Le Meur et al., 1997; Spakman and (EDT) volumes (one EDT being described as all grid nodes satisfying Nolet, 1988), normalization or scaling of the derivative matrix is per- the arrival-time differences between 2 stations, ± a tolerance value formed for better reconstruction of the different parameters. This op- known as TERR) (Font et al., 2004; Zhou, 1994). This 3D technique is eration will remove influences of parameter units and also will take well adapted to a strongly heterogeneous environment, avoids the into account the sensitivity of the data to each class of parameters. depth versus origin-time trade-off and objectively filters out possible The parameters used for the regularization of the partial derivative erroneous arrival times. It is well adapted to estimate hypocenter pa- matrix and the damping parameters used for the inversion are both rameters using only first P-arrivals even if earthquakes are outside the fixed through synthetic tests using the ray-based inversion and the network (Font and Lallemand, 2009; Font et al., 2004; Kao et al., 2000). real event-station geometry. A comprehensive description of these Updates on the MAXI technique used in this study integrate (1) a synthetic tests is presented in Gautier et al. (2006). We estimated multiscale approach of the TERR parameter, i.e., an iterative approach that the optimal set of weightings for this tomographic study is 1 for of the TERR parameter; (2) a search volume for final solution limited P waves, 2 for S waves, 5 for both the location and the origin time of by preliminary solutions (3) cleaning of outlier(s) based on EDT inter- earthquakes and finally 0.75 for the damping parameter. The total section statistics rather on travel-times residues; and (4) a final refining number of iterations for the global tomographic procedure with new search based on EDT- intersections on a resampled grid and using ray tracing has been fixed to 20 iterations. An a posteriori analysis of cleaned arrival-times (Theunissen et al., 2009, 2012). both misfit and model perturbation functions show that the conver- MAXI is based on a grid search algorithm and graph theory gence is reached after 15 iterations. for travel-time calculation. Travel-times are computed within a 3D In order to obtain a more reliable and uniform tomographic dataset, model discretized in constant velocity blocks with velocity nodes dis- we selected first arrival times that have higher quality by following the tributed on each facet according to the Shortest Path method (Moser, pick qualities (weights: Wp ≥2andWs ≥3). Then, we removed events 1991). The parameterization of the velocity grid and the MAXI proce- with greatest angle without P-observation (azimuthal gap) higher dure depends on the size of the initial model (computer memory lim- than 180. Finally only events with more than 4P and 2S picked phases itations) and the relative position of earthquakes according to the among the 38 nearest seismic stations were kept in this study. Note network. In our case, the targeted earthquakes are located below that all earthquakes located by MAXI within the a priori 3D velocity the offshore network, i.e., mostly in the crust where there are low model have a RMS lower than 0.465 s. Among the 962 events located velocities in the sedimentary layers and strong velocity variations. Pa- by MAXI, 801 events are selected (Fig. 5). rameterization therefore requires a small grid to better estimate The a priori 3D P-wave velocity model and MAXI-3D hypocenters travel-times using the Shortest Path method. Based on synthetic in- are used as initial input for the 3D tomographic inversion. As input for vestigations, MAXI parameterization are chosen as follows (1) a grid the initial S-wave velocity model, we applied a constant VP/VS ratio 3 size of 4×4×1 km with a node distribution on each facet every on the P-wave velocity model. The Vp/Vs ratio is retrieved from Wadati 500 m and (2) a TERR variation from 0.2 s to 0.6 s with an increment method (Wadati, 1933) using hypocenter parameters from MAXI and of 0.1 s. The minimum value of TERR is chosen to integrate grid size ef- P and S arrivals recorded on the 73 stations of the extended network. fect and numerical approximations. The maximum value of TERR is An average ratio of 1.740±0.015 has been obtained. large enough to take into account small arrival-time errors and small anomalies within the velocity model, and is small enough to exclude 4.3. Focal mechanism determination arrival-time outliers. The maximum value of TERR is chosen from the uncertainties on arrival-time difference between two stations. If we We selected 14 earthquakes, with local magnitude higher than consider two phases with a weight of 3, then the uncertainty on the 3.5 to compute their focal mechanisms. As developed in Nakamura difference could be estimated at 0.4 s (0.2 s+0.2 s) considering an (2002), we combined both P- and S-wave polarities. To read SV and uncertainty on the reading of 0.2 s. SH polarities, seismograms were rotated into radial and tangential Because no accurate S-wave information is available particularly components. S-wave polarity reading is difficult compare to P-wave offshore and from MAXI technique properties, we chose to determine because the S-wave is perturbed by the P-wave coda or SP arrivals the initial MAXI-3D catalog (longitude, latitude, depth and origin time) which come in just before direct S (Booth and Crampin, 1985). Conse- from the well-resolved a priori 3D P-wave velocity model. At this stage, quently, only clear polarities have been read on seismograms and used the use of a global constant VP/VS ratio instead of a 3D S-wave velocity in this study. For mechanisms showing several families of nodal planes model in a region with certainly strong VP/VS variations may lead to with P polarities, we used S/P ratios and S polarities as a last resort to bias absolute hypocenter locations (Maurer and Kradolfer, 1996). discriminate between the solutions (Hardebeck and Shearer, 2003; Kisslinger, 1980). INSU-CNRS OBS signal shows a very low noise level 4.2. Local earthquake tomography certainly because it has been deployed at depths higher than 4000 m and the good weather during the experiment avoided strong bottom We use a delayed travel-time tomography method to invert simul- current and sea wave fluctuation. Consequently, only the 12 INSU- taneously the velocity distribution and the hypocenter parameters CNRS OBSs offshore, the 8 BATS and the 2F-NET seismic stations onland

(Aki and Lee, 1976; Benz et al., 1996; Spakman and Nolet, 1988; are used to read SV/SH polarities and amplitudes. NTOU OBSs have a too Spencer and Gubbins, 1980; Thurber, 1992). The inversion method noisy signal and TAIGER OBSs orientation was unavailable for this study. provides a smooth velocity model estimated on a 3D, regularly spaced, For the focal mechanism construction, we combine P- and S-wave rectangular grid. A comprehensive description of the ray-theoretical polarities together with S/P ratios using FOCMEC program (Snoke et

Please cite this article as: Theunissen, T., et al., Crustal deformation at the southernmost part of the Ryukyu subduction (East Taiwan) as revealed by new marine seismic experiments, Tectonophysics (2012), doi:10.1016/j.tecto.2012.04.011 T. Theunissen et al. / Tectonophysics xxx (2012) xxx–xxx 9 al., 1984). Azimuth and take-off angles used to locate polarities on a the cloud defined by the distribution of each determination for one lower hemisphere representation are directly extracted from the to- given earthquake gives us an idea of the uncertainty and the stability mographic inversion forward modeling. Then, the program performs of our earthquake location. To perturb travel-times, we use three dif- a systematic grid search for acceptable focal mechanism solutions. A ferent initial velocity models, in one hand, and 3 different tomographic search interval of 2° is used in this study. inversion procedures, in the other hand. In particular, uncertainties of S Even if velocity gradients are smoothed during tomography process, phases, for which erroneous arrival times lead to significant incorrect we expect that the resulting 3D ray-path tracing will better constrain depth estimate (Gomberg et al., 1990), are thus approached in different the focal mechanism solutions. It has been demonstrated that the use ways. of a 3D velocity model can reduce the number of inconsistencies in For that purpose, initial velocity models are the a priori 3D velocity the solutions if the 3D model is well resolved (Eberhart-Phillips, 1989; model of Font et al. (2003), the update a priori 3D velocity model (this Michael, 1988; Rau et al., 1996). Moreover, as shown by Béthoux et al. study) and the best minimum 1D velocity model. We used VELEST 3.3 (2007), proximity of strong velocity heterogeneities may generate program (Kissling et al., 1994; Kissling et al., 1995a, 1995b) to invert a large discrepancies between 1D and 3D solutions. In a subduction con- 1D minimum velocity model adapted to our dataset. We followed text, where important crustal and/or lithospheric deformation is super- three inversion procedures: (1) inversion of P phases solely simulta- imposed, such approach seems thus straightforward. neously with hypocenters parameters, (2) using the result of (1) to invert simultaneously P-, S-wave velocities and hypocenters parameters 5. Resolution estimate and uncertainties and (3) joint inversion of P, S and hypocenter parameters. We run these three tests with the three available initial velocity models and 5.1. Velocity models (P and S) their associated initial hypocenter parameters. At the end, we obtain 12 different hypocenter determination datasets assumed as possible Checkerboard anomaly tests were used to assess the resolution of solutions. the tomographic models (Kissling et al., 2001). They provide a global Comparisons of all pair combinations among 12, i.e. 66 at total, insight of the local resolution length by identifying the well-resolved allow an evaluation of the impact on the hypocenter determinations area and defining the minimum anomaly size that is expected to be re- for perturbed travel-times (P and S or P phases only). Results show solved in the study. This is an a posteriori procedure because the final that hypocenter determinations given by our tomography result, i.e. tomographic model is required for performing the checkerboard tests. simultaneous inversion of P-, S-wave velocities and hypocenter pa- These tests consist in the construction of synthetic input velocity rameters from the a priori 3D model built in this study, have a models by adding a velocity perturbation to the final tomographic mean position according to other solutions in average. In any case, models (Vp =800 m/s and Vs =400 m/s). This velocity perturbation our solution is close to the barycenter of each cloud. This suggests is strong compared to the numerical noise level and also small enough that the solution converges toward our determination. On average, to avoid noticeable disturbances in the ray coverage. Synthetic travel- other positions are distant by about 3.0 km±1.1 km (1σ) in horizon- times are computed (Podvin and Lecomte, 1991) in the input velocity tal and about 3.1 km±1.0 km (1σ) in depth to the determination models using the source-receiver distribution of the real dataset. A used in this study. Moreover, the mean total distance between our so- noise term is added to the synthetic data set from a uniform distribu- lution and all others is about 4.6 km±1.5 km (1σ). Upper bound on tion between −0.05 s and 0.05 s. This simulates errors in the arrival absolute location uncertainty provided by our inversion is 4.1 km in times such as for example picking errors. The resulting synthetic data- horizontal and vertical which is the mean dispersion with 1σ around set is then inverted using the same procedure and the same parame- our solution during all 12 runs led in this analysis. terisation that was used for the real dataset. Finally, the recovered velocity is compared to the input model in order to estimate the 5.2.2. Previous catalogs model resolution for some parameters like the amplitude, the location, Hypocenter positions for a 3 month-period using OBS records the size and the shape of the reconstructed anomalies, as well as earth- shall differ from the position of known clusters in the area. The posi- quake parameters. tion discrepancy might be due to variation in fault activity that can be Checkerboard tests revealed that the dataset, both with P- or specific during the considered time period or to the fact that we sam- S-wave, is able to reconstruct in shape 20×20×12 km3 and 30×30× ple smaller magnitudes. To distinguish between those cases and as- 12 km3 pattern velocity anomalies in the target area, below the OBS sure the OBS impact on earthquake location, we have performed network and to the west at the transition with Taiwan, down to a earthquake location of RATS events also recorded by CWB and JMA depth of about 40–50 km (Fig. 6)despitesomedifficulties to retrieve stations, without using OBS records. the exact amplitude of the checkerboard. Seventy-seven (77) earthquakes have been recorded from both the CWB and JMA permanent seismic networks and OBS. We re-located 5.2. Hypocenters those events without OBS phases thanks to the MAXI technique, using the 2003-velocity model (VM-2003, Font et al., 2003) and the 5.2.1. RATS determinations new velocity model resulting from the tomographic inversion (this Uncertainties in absolute event locations result from a combination study, VM-TOMO-2011). Solutions obtained using the MAXI method of the network geometry, arrival time measurement errors and errors in VM-2003 and VM-TOMO-2011, as well as the CWB determinations in travel-times estimates, i.e., errors from the ray tracing method and in a 1D velocity model, are compared with our solutions resulting from from the difference between the real Earth and the velocity model. To a joint inversion using OBS records (Fig. 7). judge the location accuracy, a bootstrap approach can be applied in Comparison with CWB determination (Fig. 7a) indicates an epicen- which random perturbations representing picking errors are added to tral and vertical absolute misfits of 4.9±3.8 km and 7.4±11.6 km re- the travel-times and the event is relocated many separated times to ob- spectively. The relative vertical misfitis−2.4±13.5 km. One may also tain an estimate of the probable scatter in the calculated locations due observe that the misfit dramatically increase at distances exceeding to uncertainties in the picks. This technique offers the advantage of ac- 40 km from the coast. Earthquake positions obtained without OBS counting for the nonlinearities in the problem and the fact that some and within the VM-2003 model (Fig. 7b) show a similar mean absolute stations and some ray paths are much more important than others in misfit (5.0±3.1 km in horizontal and 7.4±5.9 km in vertical) with constraining the location (Billings et al., 1994). hypocenters systematically shallower than expected (relative vertical To estimate earthquake location uncertainties and to check the misfit of 4.9±8.1 km). Misfits are much reduced using the VM-TOMO- stability of the solution, we thus perturb travel-times. Accordingly, 2011 model providing a mean absolute misfit of 2.1±1.5 km in

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Fig. 6. Checkerboard test: (a) N–S vertical cross-sections. On the left: 30×30×12 km3 P-wave checkerboard anomalies used to compute synthetic travel-times. On the right: Re- trieved P-wave velocity model anomalies after tomographic inversion. (b) Horizontal cross-sections. Same description as (a).

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Fig. 6 (continued).

HM VAM VRM HM VAM VRM HM VAM VRM CWB (1D) 20 MAXI VM-2003 20 MAXI VM-TOMO-2011 20 Shallower Shallower Shallower a 10 b 10 c 10 0 0 0 Average (km) Average Average (km) Average Average (km) Average

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Fig. 7. Comparison of different earthquake determinations obtained from permanent networks without using OBS with hypocenter position obtained after tomography inversion using OBS. 77 earthquakes are compared. (a) 1D location provided by CWB, (b) 3D location within the 2003 a-priori velocity model VM-2003 (Font et al., 2003) using MAXI technique and (c) 3D location within the new 3D model obtained from the tomography inversion VM-TOMO-2011 using MAXI technique. The hypocenters are represented in map and E-W cross-section. Squares represent results of earthquake location from permanent networks without using OBS. The stroke marks the difference in position with the results of this study. Upper right in-frame shows, from left to right, statistics on the horizontal misfit HM (epicentral distance), the vertical absolute misfit VAM and the vertical relative misfit VRM (positive misfit means that the hypocenter determination in this study is deeper than those from permanent networks without OBS). The vertical bar represents 1σ deviation. and Table 2), we have checked the validity of our data and our approach and the downgoing plate. Figs. 8 to 12 present results from earth- (Fig. S1). Note that this earthquake is an intermediate-depth earthquakes location, focal mechanism determination and P-wave velocity quake located close to Taiwan west of the target area. BATS focal mech- structures. 801 earthquakes have been located during the joint inver- anisms are obtained by full waveform inversion, from all 3 components, sion and 14 focal mechanisms have been determined (Fig. 8) using at very low frequency (0.03–0.08 Hz) in order to avoid effects of strong the absolute 3D P- and S-wave velocity models obtained during the lateral heterogeneities and possible epicentral mislocation. Few 1D ve- inversion. locity models are used to adjust the depth of the Moho for each station. The description of the inversion algorithm is given by Kao et al. (1998a, 6.1. Microseismicity distribution 1998b) and Kao and Jian (1999). The focal mechanism obtained in this study from P polarities, S/P amplitude ratios and S polarities is similar to In Fig. 8, the epicenter distribution shows a similar pattern than the BATS reference with differences for the two nodal planes of 29° and previous studies that have used permanent seismic networks for 8° for the strike, 14° and 2° for the dip and, 13° and 20° for the rake. This earthquake location process (see Fig. 2 for example). From west to result validates the approach that gives weight to amplitude ratio and east, we recognize a band of earthquakes that parallels the east coast

SH polarities in the focal mechanism determination rather than using of Taiwan from shallow depths down to 80 km, a nest of shallow only the minimum number of P polarities errors as major criterion of events (b30 km) east of Suao city and north of the Hoping basin — selection for the ﬁnal solution (Manchuel et al., 2011). Indeed, inde- later called the Suao cluster (SC), and the shallow seismicity in the pendently of using amplitudes ratio or S polarities, some important P forearc area with the Hoping and Nanao clusters. We can also observe polarity errors are visible in this mechanism as well as in solutions of both very shallow (~10 km) and deeper earthquakes (>40 km) north the other 13 earthquakes (Fig. S2). The origin of such important errors of the forearc, in the Okinawa Trough and the PSP slab respectively. In could signify that local small velocity anomalies close to the sources more detail, the Hoping seismic cluster (HC) is deeper (~20 km) and are not resolved. Regarding the constraints given by the P polarities, slightly shifted eastward compared with previous determinations focal mechanisms nos. 1, 5, 6, 7 and 11 are probably less well con- (e.g., Font et al. (2004), Font and Lallemand (2009)). The microseis- strained than other (Fig. S3 and focal mechanisms in gray in Fig. 8). micity mainly concentrates in a NW–SE direction parallel to the convergence between the PSP and the EP extending along the HC and 6. Results the area of shallow earthquakes offshore Suao (SC, see for example the section at 24 km depth in Fig. 9). The Nanao seismic cluster (NC) The tomography process carried out in this paper results from a that occurs farther east remains shallow (b20 km) as in previous full 3D approach as it uses (1) an initial a priori 3D velocity model determination. that integrates information on the geometry of offshore structures based on marine active geophysical data and (2) earthquake location 6.2. P-wave velocity structure determined within this 3D velocity model. Thanks to the OBS records, the resulting Vp model and hypocenter positions provide robust con- The velocity structure developed in this project is the highest res- straints on shallow crustal structures of the active overriding margin olution 3D model that has been derived in the Southwestern Ryukyu

12 (30)

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Fig. 8. Map of the 801 earthquakes located after inversion and the 14 focal mechanisms determined in this study. Bracketed numbers refer to the depth (km) of the event; the other one is the earthquake number (Table 2). Black contour delimits the seismic clusters: SC = Suao Cluster, HC = Hoping Cluster, NC = Nanao Cluster, HuS: shallow Hualien Seismicity.

Arc-NE Taiwan area. The inversion of VP and VS structure with hypo- described by Font et al. (2001) on the basis of seismic reflection inves- centers parameters allows us to well constrain the earthquake depth tigation, is evidenced by Vp ~5.5 km/s, visible on the 12 km depth during the minimization process. However, we have led a complete section of Fig. 9. Immediately to the SE and NW, lower Vp is found re- 3D approach for the VP structure while we have used a constant initial spectively in the Nanao Basin (~5.0 km/s) and in the Hoping Basin VP/VS ratio to describe the initial VS structure. In that sense, we there- (~4.5 km/s) at the same depth. fore assume that the Vp structure is well resolved and sufficient to de- Another striking feature is the high Vp (>6.5 km) narrow anomaly scribe the crustal structure and that VS is not necessary at this step of that parallels the east coast of Taiwan on the island side up 24°30′N the analysis. at depth larger than 6–12 km (Fig. 9). Lower Vp (b6.5 km/s) is found At first, Vp variations within the shallow structures of the Ryukyu west and north of this narrow anomaly at 30 km depth or less, Arc and forearc are in good agreement with previous knowledge and suggesting that the same rock body composes the Central Range and reflect the “high and low” topographic variations of the Ryukyu acous- the southernmost Ryukyu Arc as previously suggested (e.g., Hagen tic basement below the forearc basins. Low Vp (≤4.5 km/s) is visible et al., 1988). This also suggests that the continental Moho (estimated down to 10 to 15 km depth (e.g. 12 km depth horizontal section in by the ~7.5 km/s isopleth) beneath north-easternmost Taiwan Island

Fig. 9, section 4 in Fig. 11, or section 6 in Fig. 12), surrounded by crustal is the deepest (>40 km) below these low Vp. Thereby, the tomog- velocities of the Ryukyu Arc or the Central Range (~6.0 km/s). Slightly raphy (see for example section 8 in Fig. 12) is consistent with the eastward, following the trend of the forearc sedimentary basins, the estimate of 44 km of the Moho depth constrained from receiver func- rise in the basement of the Ryukyu Arc, called Hoping Rise, previously tions at the BATS seismic station NANB (Wang et al., 2010). Also, low

Table 2 Source parameters of the 14 earthquakes for which focal mechanisms have been determined. Events are numbered in chronological order. Epicenters, depth and origin times are those obtained from the 3D inversion. Depth is given below sea level. Local magnitudes (ML) are those reported by the Central Weather Bureau (CWB). Coda magnitudes (MC) are those calculated in this study. Moment magnitudes (MW) are those reported by BATS center (Broadband Array for Taiwan Seismology).

Event Date Origin times, UT Latitude, °N Longitude, °E Depth, km Strike1, deg. Dip1, deg Rake1, deg. Strike2, deg. Dip2, deg Rake2, deg. ML(MC)(MW) (D/M/Y)

1 04/08/2008 13:15:24.7 24.2055 122.2180 18.9 230.3 19.7 44.1 97.9 76.4 104.4 −(3.0)(−) 2 12/08/2008 08:37:13.3 24.0996 122.2390 30.0 227.2 19.9 52.4 86.6 74.4 102.5 3.7(2.9)(−) 3 17/08/2008 17:37:56.3 24.0029 121.6510 45.7 74.2 53.8 54.4 304.7 49.0 128.5 4.6(2.8)(3.7) 4 21/08/2008 06:53:11.1 23.8129 122.5180 19.2 106.4 46.8 −9.5 202.9 83.1 −136.4 4.9(3.5)(−) 5 21/08/2008 06:54:29.8 23.8509 122.4870 20.2 287.2 18.8 −31.2 47.0 80.4 −106.2 4.4(2.7)(−) 6 21/08/2008 06:56:53.2 23.8779 122.4700 20.7 22.6 14.4 −55.9 167.7 78.1 −98.2 3.5(2.9)(−) 7 25/08/2008 01:37:50.8 24.1334 122.2140 21.8 146.5 80.6 −51.4 248.1 39.6 −165.1 −(3.1)(−) 8 03/09/2008 18:55:33.1 23.9982 122.4370 21.7 82.8 44.7 −57.9 221.3 53.4 −117.8 4.3(3.3)(−) 9 03/09/2008 23:11:44.9 24.1316 122.2680 22.2 294.0 72.4 77.4 150.5 21.5 124.5 4.1(3.6)(−) 10 06/09/2008 23:00:35.9 23.9917 121.7810 51.5 355.8 54.8 46.8 234.2 53.4 134.2 4.6(3.7)(3.7) 11 15/09/2008 15:08:26.5 24.0211 122.3170 19.4 57.6 55.6 −70.5 205.5 39.0 −116.0 3.9(3.0)(−) 12 17/09/2008 10:13:56.2 24.2374 122.2400 29.8 352.4 43.2 63.2 207.1 52.3 113.0 −(3.4)(−) 13 20/09/2008 08:42:29.1 24.2373 122.1480 47.1 124.6 57.8 −53.8 250.7 47.0 −133.2 3.8(3.4)(−) 14 07/10/2008 15:01:28.9 23.8135 122.5640 19.5 104.6 68.0 −36.7 210.2 56.3 −153.3 4.2(3.4)(−)

25˚ Hoping Basin 6 km 24 km

Hoping Rise

Nanao Basin

24˚

23˚ 122˚ 123˚ 122˚ 123˚ 25˚ 12 km 30 km

24˚

23˚ 122˚ 123˚ 122˚ 123˚ 25˚ 18 km 36 km

24˚

Vp (km/s)

123456789 23˚

Fig. 9. Map views of horizontal slices through the absolute P-wave velocity model (km/s). The ray-path cover is highlighted by the black contour and the non-modiﬁed initial 3D velocity model is shown in transparency. Arrows point the main velocity anomalies described in the text. The seismic clusters (HC and NC) are also shown. Receiver function at NANB seismic station exhibits a Moho at 44 km depth in agreement with low velocity (~6.5 km/s) at 36 km depth.

velocities (b7.5 km/s) are visible very locally on the 36 km depth hor- depth). Viewed in horizontal sections, the anomaly north of Hoping izontal section (Fig. 9) north of Hualien (top of the narrow high Vp Basin seems to extend within the Central Range of Taiwan at depth anomaly) and in the rifted southern Okinawa Trough at depths of ~18 km previously described along the east coast of Taiwan (Fig. 9). 30–36 km (Fig. 9) suggesting that the Moho is deeper. We also ob- The trench-parallel sections in Fig. 12 show that the PSP crust serve that the Moho rises at a depth of ~30 km at 122°20′E in the deepens toward Taiwan until it reaches the narrow high Vp anomaly Ryukyu Arc while it deepens on both sides. visible in sections 2 to 7.

Locally, crustal Vp bodies (~6.5–7.0 km/s) are observed in the The last notable feature revealed by this new local tomography is Ryukyu Arc basement embedded into lower Vp rocks north of Hoping the presence of sharp lateral discontinuities in the vicinity of the plates Basin and also north of Nanao Basin (Fig. 9 at 12 km and 18 km interface visible, for example in sections 6 and 7 (Fig. 11) or in section 8

121˚ 122˚

25˚

24˚

23˚ Focal depth (km)

0 10 20 30 40 50 60 70 80 90 100 110 120

Fig. 10. Location map of the vertical sections through the tomography model with hypocenters represented in Figs. 11 and 12. The toponymy is labeled.

(Fig. 12). Isocontours of Vp anomalies are typically offset vertically by (no. 6) showing an E–W T-axis also occurs close to the ISZ or in the 10 km along these discontinuities. upper part of the PSP.

6.3. Focal mechanisms 7. Discussion and preliminary tectonic interpretation

The 14 focal mechanisms revealed a complex seismic pattern 7.1. From collision along the eastern coast of Taiwan to “subduction” (Figs. 8, 11, 12 and Table 2) showing both reverse and normal faulting with variable P and T axes. Two reverse oblique mechanisms show lat- Our microseismicity survey confirms the high rate of seismicity eral compression (n°10 and 12) with WNW–ESE P-axes, one offshore along the eastern coast of Taiwan at depths ranging from the surface Hualien and the other in the western edge of the HC. Both events occur to nearly 80 km (Figs. 2 and 8), the deepest events being located just at depths of 30 and 51 km within the PSP. Four mechanisms are com- north of Hualien. The seismicity is not observed, during the three patible with N–S compression (nos. 1, 2, 3 and 9), 3 of them being lo- months experiment, within the highest values of Vp measured along cated in the HC and one beneath Hualien. The one occurring beneath the narrow band that parallels the coast (see Section 6.2), but imme- Hualien (no. 10), which is confirmed by the BATS determination diately east of it (Fig. 12 section 5 for example, or Fig. 9 on horizontal using waveform inversions (see Section 5.3) is very deep (46 km) section at 12 km depth for shallow events). The events deeper than into the PSP. All the other events reveal extension along three main 40 km are either near or east of the coastline. The high velocity anom- directions: (1) NNE–SSW T-axes for two of them (nos. 7 and 13) at aly has been interpreted by Lin et al. (1998) and also by McIntosh et al. the western edge of the HC either deep into the PSP (depth 47 km) (2005) as deep material presently under exhumation. However, we or close to the ISZ (depth 22 km); (2) five events (nos. 4, 5, 8, 11 and cannot rule out a mantle origin, also proposed by Kim et al. (2006) 14) all located in the SE part of the HC at depths compatible either and Liang et al. (2007), for deeper parts of the PSP sandwiched during with the ISZ or the upper part of the PSP show NW–SE T-axes close the collision. Instrumental seismicity in this region shows thrust-type to the PSP/EP convergence vector; (3) a last event poorly constrained mechanisms with WNW–ESE P-axes like the focal mechanism of deep

Distance (km) Distance (km) 0 20 40 60 80 100 120 140 160 180 0 20 40 60 80 100 120 140 160 180 0

−20

−40 Depth (km) −60 1 3 2

0 20 40 60 80 100 120 140 160 180 0 20 40 60 80 100 120 140 160 180 0

−20

−40

−60 10 3 4

0 20 40 60 80 100 120 140 160 180 0 20 40 60 80 100 120 140 160 180 0

−20

−40 Depth (km) −60 5 13 6 0 20 40 60 80 100 120 140 160 180 0 20 40 60 80 100 120 140 160 180 0 9 7 12

−20

−40 Depth (km) −60 2 7 11 8 0 20 40 60 80 100 120 140 160 180 0 20 40 60 80 100 120 140 160 180 0

−20

−40 5 6 Depth (km) −60 8 9 10

0 20 40 60 80 100 120 140 160 180 0 20 40 60 80 100 120 140 160 180 0

−20

−40 Depth (km) −60 4 14 11 12

123456789 Vp (km/s)

Fig. 11. Vertical cross-sectional views (10 km of total thickness) of the absolute P-wave velocity model (km/s) oriented along the direction ~perpendicular to the trench and parallel to the coast. The ray-path cover is highlighted by the black contour and the non-modiﬁed initial 3D velocity model is shown in transparency. Arrows point the main velocity anomalies described in the text. The seismic clusters (SC, HC and NC) are also shown. Focal mechanisms obtained in this study are reported in cross-section.

event n°10 (Figs. 8 and 12 in section 4). The apparent westward deep- to north, an east-dipping plane (Fig. 12: between 0 and 10 km depth ening of the PSP visible in Fig. 12 (sections 2 to 7), the intermediate- at X=60 km in section 3) turning into a vertical (Fig. 12: section depth seismicity (between 30 km and 50 km) within the PSP litho- 5) and then a west-dipping plane (Fig. 12: section 6). This inversion sphere and focal mechanism no. 10 are in favor of PSP underthrusting is in agreement with an out of sequence deformation associated beneath Taiwan north of LVF associated with lateral compression with the beginning of underthrusting of the PSP under the Central within the PSP mantle. The Luzon Arc would thus sink north of the Range between 24°N and 24.25°N along the east coast of Taiwan.

LVF (Fig. 12: sections 3, 4 and 5) just east of the high VP anomaly (yel- This scheme would be partially in agreement with scenario proposed low arrow on sections), which seems to undergo an important defor- by Rau et al. (2007) in which the west-dipping plane would start at mation in the front of the Luzon Arc. Shallow micro-seismicity, aligned 23.5°N along the northern part of the LVF. The pure collision stage NNE–SSW north of the LVF (called HuS in Fig. 8), reveals, from south would stop between 24°N and 24.25°N where the underthrusting of

Distance (km) Distance (km) 0 20 40 60 80 100 120 140 160 180 0 20 40 60 80 100 120 140 160 180 0

−20

−40 Depth (km) −60 1 2

0 20 40 60 80 100 120 140 160 180 0 20 40 60 80 100 120 140 160 180 0

−20

−40 Depth (km) −60 3 4 10

0 20 40 60 80 100 120 140 160 180 0 20 40 60 80 100 120 140 160 180 0 14

−20

−40 Depth (km) −60 5 6 5

0 20 40 60 80 100 120 140 160 180 0 20 40 60 80 100 120 140 160 180 0

−20

−40

Depth (km) 6 −60 11 9 7 8 8 2 13 1 7 0 20 40 60 80 100 120 140 160 180 0 20 40 60 80 100 120 140 160 180 0

−20

−40 Depth (km) −60 9 12 10

123456789 Vp (km/s)

Fig. 12. Vertical cross-section views (10 km of total thickness) of the absolute P-wave velocity model (km/s) oriented along the direction of the convergence vector PSP/SCB. The ray-path cover is highlighted by the black contour and the non-modiﬁed initial 3D velocity model is shown in transparency. Arrows point the main velocity anomalies described in the text. The seismic clusters (HC and NC) are also shown. Focal mechanisms obtained in this study are reported in cross-section. (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article.)

the PSP under the Central Range begin. In such case, the shortening is the northern part of the Central Range (Suppe et al., 1984; Tan, 1977). accommodated by the northwestward underthrusting of the PSP According to the crustal velocity (~6–7km/s)(Fig. 9 at 30 km depth), below the Central Range north of 24°N and mainly farther east off- this continuity reaches 122.5–123°E. Looking more carefully, the shore between latitudes 23.5°N and 24°N to explain GPS velocity Moho of the Ryukyu Arc is shallower between 122.3°E and 122.5°E ﬁeld north of the Coastal Range (Fig. 1). but an interpretation of it seems premature. The intense seismicity of SC is located at depths less than 30 km mostly within the upper plate. 7.2. Transition from the Ryukyu Arc to the Central Range Unfortunately, we were not able to obtain focal mechanisms in this region, but several authors have described strike-slip faulting like the

Again, our locations of shallow events in the Suao cluster (SC) con- June 5, 1994 ML6.5 Nanao Earthquake (Huang et al., 2012; Lallemand

ﬁrm previous determinations (see Figs. 2 and 8). Based on Vp velocities, and Liu, 1998; Wu et al., 1997) along the SC and normal faulting like we observe that this region of the upper plate north of the Hoping the April 22, 2002 ML4.9 to the east (http://bats.earth.sinica.edu.tw/). Basin seems to extend the Central Range offshore with an ~90° clock- North of the Hoping Basin, this seismicity is aligned in a NW–SE direc- wise bend (Fig. 9). This is in agreement with the ﬁrst similar observation parallel to the convergence between the PSP and the EP as well as tion of Hagen (1988), based on active seismic observations offshore, the seismicity of the Hoping cluster that is located close to the ISZ or but also with onland structural observations (foliation, lineation) at within the PSP crust. We suppose that the Suao cluster is a response

NW–SE direction. ISZ. In the neighborhood, the MW7.1 2002/03/31 interplate thrust We have noticed two high-velocity bodies (Vp ~6.5 km/s) about 10 event also occurred at the western edge of the HC around 23.5 km in to 20 km wide and 5 km thick (Fig. 9, section 4 in Fig. 11 and section 9 depth near the “step” in the PSP (Fig. 2). It was followed by 5 years in Fig. 12) that seems embedded into the upper plate. The nature of of after-slip between 30 km and 60 km in depth east of 122.5°E these bodies is still enigmatic. We see on the horizontal sections at (Nakamura, 2009a). The MW7.2 1963/02/13 thrust event (Fig. 2)is 12 and 18 km depth (Fig. 9) that the same velocities are found at sim- also located at the same place (even if this earthquake have not been ilar depths along the eastern side of the Central Range. Their nature relocated). In the end, we suppose that the PSP tear is the locus of a could be metamorphic rocks (Lin et al., 1998; McIntosh et al., 2005) seismic asperity responsible for an increase of the ISZ seismic cou- but as mentioned previously we cannot rule out a mantle origin for pling, for repeating earthquake occurrence in this area (Igarashi, deeper parts of the PSP sandwiched during the collision. At last, as 2010) and for recurrent major earthquakes with magnitude higher there is not a perfect continuity of these bodies below the south than 7. Further investigations are needed to validate this scenario. Ryukyu Arc slope with the eastern side of the Central Range, that is The top of the PSP (~6.0 to ~6.5 km/s) shows a few highs with thus possible to also consider it as pieces of Luzon Arc torn during variable wavelengths as north of the Hoping Rise in section 8 of the the propagation of the PSP toward the west last 1 or 2 My. Fig. 12 (X=95 km, Z=20 km). Font et al. (2001) have suggested that subducting reliefs, like a seamount or an offscraped part of the Luzon 7.3. The Hoping Cluster and the interplate seismogenic zone volcanic arc, were responsible for the uplift of the Hoping Rise. Later, Wang et al. (2004), based on its interpretation of the active seismic One important result concerns the relocation of the numerous line EW-16, proposed that it might result from a buckling and even events that occur within the HC, ﬁrst because a large part of the seismic more slicing of the PSP due to lateral compression exerted by the colli- deformation in the southern Ryukyu is concentrated there and second sion further west. The focal mechanism no.12, juxtaposed close to this because previous (very) shallow determinations were puzzling. The “high” (Fig. 12: section 9), is a thrust event with an E–WP-axisassoci- new determinations are well constrained especially compared with ated with a seismicity showing a west dipping plane in agreement previous in depth. As initially suspected, the lack of OBSs together with this scenario. with the use of inaccurate velocity model has produced mostly shallow determinations as shown in Fig. 7. In this study, we have improved both 7.4. Normal faulting in the upper plate: the Nanao Cluster the azimuthal gap by deploying OBS above the cluster and the 3-D initial velocity model by updating a 3D a priori velocity model. We can Finally the distribution of the microseismicity in the NC conﬁrms now say that most of the microseismicity occurs in the vicinity of the previous shallow locations within the upper plate. Unfortunately, ISZ (see horizontal sections in Fig. 9). However, looking more carefully, we did not obtained any new focal mechanisms in this region but earthquakes are distributed over a fringe of about 10 km in depth (see this area has been investigated with swath bathymetry and multi- Figs. 11 and 12)orevenmore(seesection8inFig. 12) meaning that channel seismics and it was characterized by eastward-facing ~N–S not all of them are generated along the ISZ. normal faults probably caused by trench-parallel stretching of the We suspect that most of these events occur within the subducting margin in response to oblique subduction (Lallemand et al., 1999) plate. Furthermore, 5 focal mechanisms among 12 in the HC show or caused by the subduction of the Gagua ridge (Dominguez et al.,

NW–SE extension (4, 5, 8, 11 and 14 in Fig. 8), and two others (6, 7 1998; Konstantinou et al., 2011). The MW6.8 December 18, 2001 and 13) extension in other directions. Only 3 focal mechanisms (1, earthquake occurred in this cluster and revealed N–S trending normal 2 and 9) are compatible with the ISZ, the last one (12) showing lateral faults. compression within the PSP. The supposed intra-PSP extensional events should reveal either downdip extension caused by the slab 8. Conclusion pull and/or bending caused by the lateral compression at the termination of the subduction zone (Chou et al., 2006; Font et al., 1999; Kao (1) One of the main results of this study is the improvement of the and Rau, 1999; Wang et al., 2004). hypocenter depth determination accuracy offshore. This has Another interesting feature is the sharp velocity gradient associat- been possible thanks to the active and passive experiments ed with an ~10 km vertical offset in the 6.5, 7.0 and 7.5 km/s velocity that allowed to obtain a reﬁne 3D velocity model for the re- contours below the Ryukyu arc slope (see for example sections 5, 6, gion. We have shown that this model greatly improves the hy- 7 and 8 in Fig. 11). The offset suggests a kind of trench-parallel step af- pocenter determination, especially at depth, using CWB and fecting the crust of the subducting PSP or a rough topography on top of JMA stations around even without using OBS. the PSP (as a piece of Luzon Arc). This step could represent the “sur- (2) The mean depth of the Hoping cluster (HC) has been revised face” expression of the PSP tear ﬁrst proposed by Lallemand et al. from a shallow level to a 15–30 km depth range in the vicinity (1997) and further mentioned by Font et al. (2001). Such a tear within of the ISZ. Earthquakes often concentrate nearby a trench- the PSP was supposed to allow the PSP to overthrust the EP along the parallel “step” that offset the crust of the PSP and could repre- LVF and to subduct northwestward beneath NE Taiwan. According to sent the upper part of a lithospheric tear earlier proposed by its position downstream below and north of Ryukyu forearc basins, Lallemand et al. (1997) and Font et al. (2001). This tear would such process is probably not the main reason for plate tearing. This be responsible of a seismic asperity that generates repeating tear could be reactivated by the combination of the PSP collision and earthquakes along the ISZ as mentioned by Igarashi (2010) the interaction with deep crustal root of the orogen extending within and recurrent earthquakes with moment magnitude higher the Ryukyu Arc. One may observe that the seismicity concentrates than 7. M7.7 1920, M7.2 1963 and M7.1 2002 thrust events, lo- near the “step” offsetting the top of the PSP. One would expect the cated in the HC, originated from this asperity at about 122.1– presence of a ramp into the ISZ which would be compatible with the 122.2°E of longitude below the west part of the Hoping rise. high level of seismicity and possibly also with the formation of a (3) The PSP undergoes a severe deformation (1) east of Taiwan splay-fault as suggested by Font and Lallemand (2009), but not yet ob- within the mantle (30–50 km depth) (2) below the Ryukyu served. Back to the MW7.7 1920 largest instrumentally recorded earth- forearc and (3) north of 24°N along the east coast of Taiwan quake that occured in the western part of the HC, the relocation based where the PSP probably underthrusts the northern part of the on an “analog-quake” with the velocity model obtained in this study Central Range. Below the Ryukyu forearc, in addition to an

eventual tearing close to Taiwan, the deformation mainly con- Aki, K., Lee, W.H.K., 1976. Determination of three-dimensional velocity anomalies under a seismic array using first P arrival times from local earthquakes; 1, a homo- sists of extension in its upper part as a result of both bending/ geneous initial model. Journal of Geophysical Research 81 (23), 4381–4399. buckling caused by lateral compression and downdip extension Angelier, J., Chang, T.-Y., Hu, J.-C., Chang, C.-P., Siame, L., Lee, J.-C., Deffontaines, B., Chu, caused by slab pull. This strong bending could accommodate its H.-T., Lu, C.-Y., 2009. Does extrusion occur at both tips of the Taiwan collision belt? Insights from active deformation studies in the Ilan Plain and Pingtung Plain re- steep subduction beneath the thick crust of the Central Range. gions. Tectonophysics 466 (3–4), 356–376. 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Evolutionary model for model built in this study and of the relocated major historical instru- the Taiwan collision based on physical modelling. Tectonophysics 274 (1–3), mental earthquakes (1900–1990) in combination with the velocity 253–274. structure interpretation has to be led to propose a coherent scenario Chemenda, A., Lallemand, S., Bokun, A., 2000. Strain partitioning and interplate friction in oblique subduction zones; constraints provided by experimental modeling. for this region that could explain convergence accommodation pro- Journal of Geophysical Research 105 (B3), 5567–5581. cesses and major earthquakes occurrence. Chemenda, A.I., Yang, R.K., Stephan, J.F., Konstantinovskaya, E.A., Ivanov, G.M., 2001. New results from physical modelling of arc–continent collision in Taiwan; evolutionary model. Tectonophysics 333 (1–2), 159–178. Acknowledgements Chen, R.-Y., Kao, H., Liang, W.-T., Shin, T.-C., Tsai, Y.-B., Huang, B.-S., 2009. Three-dimensional patterns of seismic deformation in the Taiwan region with special implication from the – – Authors wish to thank Shih Min-Hung for his implication and his 1999 Chi-chi earthquake sequence. Tectonophysics 466 (3 4), 140 151. Chiao, L.-Y., Kao, H., Lallemand, S., Liu, C.-S., 2001. An alternative interpretation for slip advices in extraction and preparation data. Also, thanks to Kevin vector residuals of subduction interface earthquakes; a case study in the western- Manchuel for his advices about SEISAN and FOCMEC softwares and most Ryukyu slab. Tectonophysics 333 (1–2), 123–134. picking earthquakes. Furthermore, thanks to Marie Picot, Audrey Cala- Chou, H.-C., Kuo, B.-Y., Hung, S.-H., Chiao, L.-Y., Zhao, D., Wu, Y.-M., 2006. The Taiwan- Ryuku subduction–collision complex; folding of viscoelastic slab and the double buig, Agastin Ludovic and Michaela Chronopoulou for their important seismic zone. Journal of Geophysical Research 111 (B4), 14. picking work. Huang Bor-Shou from Academia Sinica, CWB and JMA Chou, H.-C., Kuo, B.-Y., Chiao, L.-Y., Zhao, D., Hung, S.-H., 2009. Tomography of the are thanked for providing us arrival-time datasets. Thanks to the westernmost Ryukyu subduction zone and the serpentinization of the fore-arc mantle. Journal of Geophysical Research 114, B12301. RATS technical staff or crew from INSU-CNRS or IONTU onboard R/V Chung, L.-H., Chen, Y.-G., Wu, Y.-M., Shyu, J.B.H., Kuo, Y.-T., Lin, Y.-N.N., 2008. Seismogenic OR1. Josiane Tack and Fabrice Grosbeau are acknowledged for the es- faults along the major suture of the plate boundary deduced by dislocation modeling tablishment and improvement of the computing cluster of the labora- of coseismic displacements of the 1951 M7.3 Hualien–Taitung earthquake sequence – – tory, and their advices. Anne Delplanque (GM) is thanked for her help in eastern Taiwan. Earth and Planetary Science Letters 269 (3 4), 415 425. Crosson, R.S., 1976. 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