Structural factors controlling the rupture process of a megathrust earthquake at the Nankai trough seismogenic zone

Geophys. J. Int. (2002) 149, 815 835 Structural factors controlling the rupture process of a megathrust earthquake at the Nankai trough seismogenic zone S. Kodaira, 1 E. Kurashimo, 2 J.-O. Park, 1 N. Takahashi, 1 A. Nakanishi, 1 S. Miura, 1 T. Iwasaki, 2 N. Hirata, 2 K. Ito 3 and Y. Kaneda 1 1 Japan Marine Science and Technology Center, Natsushima 2-15 Yokosuka 237-0061 Japan. E-mail: kodaira@jamstec.go.jp 2 Earthquake Research Institute, University of Tokyo, Yayoi 1-1-1 Bunkyo-ku 113-0032 Tokyo, Japan 3 Disaster Prevention Research Institute Kyoto University, Uji 611-0011 Kyoto, Japan Accepted 2002 January 7. Received 2002 January 7; in original form 2001 May 11 SUMMARY The Nankai Trough is a vigorous subduction zone where large earthquakes have been recorded with a recurrence time of 100 200 yr. The 1946 Nankaido earthquake is well known as an unusual event among these earthquakes, because the rupture zone estimated from long-period geodetic data is more than twice as large as that derived from seismic wave data. In the summer of 1999, an onshore offshore deep seismic survey was performed along a 355 km long profile in the western Nankai Trough seismogenic zone. Seismic signals both from an airgun array (207 l) and land explosions (maximum of 500 kg) were recorded simultaneously by 98 oceanbottom seismographs and 93 land seismic stations. Conventional 2-D seismic reflection data were also acquired along part of the offshore profile. From the wide-angle seismic data, we found a subducting seamount at the centre of the proposed rupture zone with dimensions of 13 km thick by 50 km wide at 10 km depth. The seismic velocity image also shows that the seamount is now colliding with the Japanese island arc crust. From this significant structure, this paper proposes that the subducted seamount functioned as a barrier at least during the 1946 earthquake, i.e. the rupture of the 1946 earthquake extended over the entire locked zone to the east of the subducted seamount, and then the rupture was deflected around the subducted seamount at the down-dip end of the locked zone between Cape Muroto and Cape Ashizuri. Another significant structure, a highly reflective layer, is obtained beneath Shikoku Island. A very slow P-wave velocity (3 km s 1 ) is necessary in a thin layer at the base of the island arc crust in order to explain the observed high-amplitude reflection phases. An area of low resistivity obtained by a previous magnetotelluric study corresponds to the highly reflective layer. This suggests a possible water layer at the base of the island arc crust. The water may be generated by dehydration of the downgoing probably partially serpentinized mantle, which is implied by a low P-wave velocity (7.5 km s 1 ) beneath the subducted seamount. A locally observed non-slip region during the 1946 earthquake at the eastern part of Shikoku Island is interpreted as a result of weak coupling at the possible water layer. Key words: crustal structure, earthquake, Nankai Trough, seismic refraction, subduction. 1 INTRODUCTION The Nankai Trough, off southwest Japan, is an active plateconvergent margin where the Philippine Sea plate is subducting beneath the Eurasian plate (Fig. 1). The Izu Bonin arc and the Kyusyu Palau ridge separate the Nankai Trough from other subduction zones. The Kinan seamount chain is recognized in the Shikoku Basin to the southeast of the Nankai Trough along a plate-convergent direction (e.g. Kobayashi et al. 1995). The present convergence rate at the Nankai Trough is estimated to be 43 46 mm yr 1 on the basis of detailed earthquake mechanism studies and geological constraints (Seno et al. 1993). A study of very long baseline interferometry (VLBI) also shows a convergence rate of 39 ± 2mmyr 1 (Hashimoto & Jackson 1993). Active subduction at the Nankai Trough has been the cause of repeated great interplate earthquakes, the history of which can be traced back for over a thousand years (Ando 1975). It has been proposed, by several studies (e.g. Sugi & Uyeda 1984; Shiono & Sugi 1985), that the variation in the age of the subducting oceanic crust at the trough might be a factor in controlling the seismic activity along the Nankai seismogenic zone The Shikoku Basin located at the northern part of the Philippine Sea plate is estimated to have C 2002 RAS 815

816 S. Kodaira et al. Figure 1. Map showing the area around Japan. The contour interval of water depth is 2000 m. Thick dashed lines indicate plate boundaries. A framed area represents a survey area of this study. SB, Shikoku Basin; KS, Kinan seamount chain; PB, Parece Vela Basin. opened from 26 to 15 Ma by backarc spreading behind the Izu Bonin arc. Spreading ceased at 15 Ma (Kobayashi & Nakada 1978; Hibbard & Karig 1990; Okino et al. 1994; Kobayashi et al. 1995). Since then, the subduction of the Philippine Sea plate has been to the northwest, which is approximately along the extinct spreading axis. The age of the Philippine sea plate subducting in the Nankai Trough thus varies to both the southwest and the northeast of the extinct axis from 25 to 15 Ma (Kobayashi et al. 1995). The last large thrust events in 1944 (Tonankai earthquake, Ms = 8.2) and 1946 (Nankaido earthquake, Ms = 8.2) were studied by several earthquake seismologists. Two conflicting results concerning the rupture process of the 1946 Nankaido earthquake have, however, remained as an unsolved question. The contradictory results were derived independently from seismic wave data and geodetic data, i.e. the geodetic data show a fault size of 2.5 10 4 km 2, with a slip of 5 18 m (Fitch & Scholz 1971), while a fault model obtained from seismic data shows a size of 1 10 4 km 2 with a slip of 3 m (Kanamori 1972). The area and size of the fault derived from seismic data correspond to the 1 day aftershock area (Mogi 1968), which is located in the eastern half of the fault area determined by the geodetic data. The slip of 3 m represents almost complete release of elastic strain inferred to have accumulated between the 1854 earthquake and the 1946 events (Kanamori 1972). What happened in the seismogenic zone of the 1946 Nankaido earthquake? Structural images of the entire seismogenic zone, including subducting oceanic lithosphere and island arc crust, provide fundamental information to address the above question. Crustal and deep sedimentary structures of the western Nankai Trough seismogenic zone, from the trough axis to the coastline of Shikoku Island, have been obtained by several seismic surveys. Numerous seismic reflection studies have, however, focused mainly on the evolution of the forearc accretionary prism. For example, Moore et al. (1990) obtained high-quality multichannel seismic reflection data around the deformation front in the western Nankai Trough. Their interpretation of these data was that the accretionary prism is created by scraping off as much as a 1 km thick clastic sediment section from the incoming oceanic plate. The internal structure of the prism is characterized by anticlines separated by mainly landwarddipping thrusts with a spacing of a few kilometres (Moore et al. 1990; Taira et al. 1991). Their data also showed that a décollement, just landward of the trough axis, lies several hundred metres above the top of the oceanic crust (Taira et al. 1991). The décollement steps down to near the top of the crust 30 40 km landward of the trough axis (Moore et al. 1990). Recently, Park et al. (1999a, 2002) obtained clear images at the proposed up-dip limit of the seismogenic zone. They imaged a subducted seamount and a splay fault system consisting of several sigmoidal out-of-sequence thrusts (OOST). Most of the OOSTs are apparently developed from the subducting oceanic crust to the seafloor, cutting through underthrust sediments and the overriding accretionary prism. Since these previous seismic reflection images did not reach the down-dip limit

Factors controlling an earthquake in the Nankai trough 817 of the seismogenic zone, several wide-angle ocean-bottom seismic studies tried to obtain deeper images of the subducting oceanic crust in the 1980s (e.g. Kinoshita & Matsuda 1989; Nishizawa & Suyehiro 1989). They were able to trace the subducting oceanic crust down to 7 10 km depth. This is substantially seaward of the down-dip limit of the seismogenic zone. Kodaira et al. (2000a) obtained a first crustal model reaching to the down-dip limit of the Nankai seismogenic zone by using wide-angle ocean-bottom seismic data. They concluded that the down-dip limit extends to 25 km depth along the subducting oceanic crust, but does not reach the deep end of the oceanic crust-island arc crust contact zone. Nakanishi et al. (2002) also showed that the down-dip limit of the rupture zone of the 1944 and the 1946 earthquakes lies around 23 25 km depth along the entire western Nankai Trough by compiling recently acquired wide-angle ocean-bottom seismic data. As mentioned above, recent progress of seismic imaging provides fine-scale structure at the up-dip limit of the seismogenic zone and the overall structural features of the entire seismogenic zone. There has still been little consideration of the most essential question concerning subduction earthquake rupture processes; i.e. what structural factors control the rupture process of the large subduction earthquakes. One purpose of this study is, therefore, to examine whether significant structural factors exist in the proposed co-seismic slip zone by modelling onshore offshore wide-angle seismic data acquired along a profile extending from the Nankai Trough to the Japanese islands in order to image the entire Nankai subduction system. geophones and a digital recorder (16-bit A/D and 100 Hz sampling). Seismic signals both from the air gun array and the land explosions were recorded by the OBSs and the land stations simultaneously. 3 DESCRIPTION OF SEISMIC DATA AND PHASE IDENTIFICATIONS 3.1 Offshore multichannel seismic data The multichannel seismic reflection data provide a shallow image of the offshore profile (Fig. 3). The section clearly images a sedimentary layer in the trough region and the top of the subducting oceanic crust in the middle of the profile. The reflection from the top of the oceanic crust seems to show 1 2 km of shallowing below the landward slope ( 95 110 km) as plotted by the broken line in Fig. 3(b). A reflector interpreted as being a result of décollement is observed several hundred meters above the top of the oceanic crust in the middle of the section. The décollement seems to step down between CDP 6000 and CDP 7000. Clear reflection events at 11 12 km depth (CDP 3000 5000) are interpreted as the Moho discontinuity of the subducting oceanic crust. A clear continuation of the subducting oceanic crust cannot be traced into the northern part of the section, which enters the proposed co-seismic slip zone (Ando 1975). In this study we mainly model the wide-angle seismic data to image the deeper part of the co-seismic slip zone where the multichannel seismic data could not provide detailed information. 2 DATA ACQUISITION From 1999 May to July the Japan Marine Science and Technology Center (JAMSTEC), in cooperation with the University of Tokyo, Kyoto University, Kyushu University and Kochi University, performed an extensive geophysical experiment at the western Nankai Trough seismogenic zone. Onshore offshore wide-angle seismic data and offshore multichannel seismic data were acquired as a part of this experiment. The location of the onshore offshore seismic profile is shown in Fig. 2. Along the offshore part of the wide-angle seismic profile (185 km long), 98 ocean-bottom seismographs (OBSs) were deployed with 1.6 km spacing except at the southern end of the profile (Fig. 2). Among these OBSs, three (OBS10, OBS22 and OBS33) could not be recovered, and four OBSs (OBS18, OBS19, OBS23 and OBS38) recorded no seismic signal because of mechanical problems of the recorder system. Multichannel seismic data were also acquired between the Nankai Trough and the Tosa-bae, 0 135 km from the southern end of the profile. For the wide-angle and multichannel seismic profile an airgun array with a total volume of 207 l was fired every 200 and 50 m by JAMSTEC s R/V Kaiyo and R/V Kairei, respectively. The OBS used was designed by Shinohara et al. (1993) and is a digital recording version of an OBS originally designed by Kanazawa & Shiobara (1994). The OBS has gimbal-mounted geophones (4.5 Hz) and a hydrophone. The digital recorder has 16-bit analogue-to-digital converter and records data at 100 Hz sampling rate. 93 land stations were deployed with 1 2 km spacing along the onshore parts of the profile (Fig. 2). Three explosive sources (J1 J3) were shot at the middle and the southern end of the onshore profile (Fig. 2). Charge sizes of the shots were 500 kg at J1 and J2, and 100 kg at J3. Recording systems used on the onshore profile (Shinohara et al. 1997) were similar to the OBS system, with 4.5 Hz 3.2 Onshore offshore wide-angle seismic data We made three types of wide-angle seismic record sections depending on the shot receiver pairs. Schematic figures of those are shown in Fig. 4. The first type of record section (Fig. 4a) is a receiver-gather of the air gun shots recorded by the OBSs. This section represents a conventional offshore type of record section. The second type of record section (Fig. 4b) is a receiver-gather of the air gun shots recorded by the land stations. We only used the land stations located on Shikoku Island, since land stations located further north recorded poor quality data from the air gun shots. The third type of record section (Fig. 4c) is a shot-gather of the land explosions. The densely deployed OBSs enabled us to plot the shot-gather record sections such as a conventional land refraction profile even using the onshore offshore data. The second and third types of section provide us with a detailed structure around the coast line, which is usually missing when marine and land seismic data are obtained separately. In the following sections, our phase identifications of the wideangle seismic data are described. Since large structural variations along the entire profile were expected, we referred to the characteristics of wide-angle seismic data and crustal models of previous studies (Nakanishi et al. 1998, 2002; Kodaira et al. 2000a; Takahashi et al. 2002) when we identified phases in our data. 3.2.1 Air gun shots recorded by OBSs Examples of the OBS record sections are shown in Figs 5 and 6. The OBS data observed around the trough show the typical characteristic of oceanic crust, which consists of refraction phases from the subducting oceanic crust (Po in Fig. 5a) and the uppermost mantle (Pn in Fig. 5a), and a reflection phase from the oceanic Moho (PmP in Fig. 5b). The character of record sections varies gradually

818 S. Kodaira et al. Figure 2. Survey area with the onshore offshore wide-angle seismic profile. Contour interval of water depth is 1000 m. Circles indicate OBS positions numbered from 1 to 98 toward the north. Squares indicate positions of land stations numbered from 1 to 93 toward the south. J1 J3 indicate positions of land explosions. Rectangular areas labelled A and B show co-seismic slip zones of the 1946 Nankaido earthquake obtained by Ando (1975). Thick dashed area and star show the 1 day aftershock area (Mogi 1968) and an epicentre of the 1946 earthquake (Kanamori 1972). Geological units are simplified after Taira et al. (1991). Topography data used were compiled by the Hydrographic Department, Japan Maritime Safety Agency. CS, cape Shiono; CM, cape Muroto; CA, cape Ashizuri. from the typical oceanic character at the foot of the landward flank of the trough, 90 km from the southern end of the profile. Refraction arrivals interpreted to be from the sedimentary wedge (Psed in Figs 5c and d) are observed at near offset on OBSs at the landward flank to the Tosa-bae. For example, the refraction arrivals from the sedimentary wedge are recognized at 5 20 km offset on OBS58 (Fig. 5c) and at 5 25 km offset on OBS70 (Fig. 5d). More landward towards the Tosa-bae, Psed is not apparent, but refraction arrivals from the island arc upper crust are observed at near offsets (e.g. Pu in Figs 6a and b). This implies that the sedimentary wedge becomes thinner and that the island arc upper crust becomes shallow toward the land at Tosa-bae. The Po phase following the Pu phase is observed on the OBSs landward of the Tosa-bae (e.g. Fig. 6); however, we could identify an offset in traveltimes between Po and Pu. This is clearly identified, for example, at 40 km offset on OBS83 (Fig. 6b) and 40 km offset on OBS98 (Fig. 6d). The existence of this offset in the traveltimes implies a possible velocity inversion with depth (a low-velocity layer) below the island arc upper crust. Refraction arrivals from the uppermost mantle (Pn in Figs 5 and 6) are observed following the Po phase

Factors controlling an earthquake in the Nankai trough 819 Figure 3. (a) Poststack depth migrated section of MCS data acquired as a part of the offshore profile, from the trough to the Tosa-bae. (b) Interpretation of the MCS section (after Park et al. 1999b). at far offset. Intercept times of the Pn phases become larger toward the land; e.g. 6sonOBS34 (Fig. 5b) and 8 s on OBS83 (Fig. 6b). This implies landward deepening of the uppermost mantle. We also identified a wide-angle reflection phase from the top of the mantle (PmP in Figs 5 and 6) as a major later phase, which approaches the Pn phase. 3.2.2 Air gun shots recorded by land stations The air gun shots are clearly observed by the land stations deployed on Shikoku Island (Fig. 7). This indicates that onshore recording of the offshore shots is a useful method of revealing the structure of the ocean land transition area (cf. Trehu et al. 1994; Gerdom et al. 2000). First arrivals are characterized by two major refraction phases; refraction phases from the island arc crust (Pu) and the uppermost mantle (Pn). The refraction phase from the subducting oceanic crust (Po) is also observed on the stations near the coastline, in which an offset of traveltimes between Pu and Po phases is recognized (Fig. 7a). The most prominent character of the record section of the air-gun shots observed by land stations is a large-amplitude reflection phase (PoP) that approaches the Pu phase (Fig. 8). We interpreted this phase as a reflection from the bottom of the island arc crust, and the observed remarkably high amplitude seems to indicate a large impedance contrast. Similar deep reflection phases were observed at the Cascadia margin (e.g. Hyndman 1988; Gerdom et al. 2000). Figure 4. Shot-receiver configurations of the onshore offshore profile. (a) Airgun shots recorded by the OBSs. (b) Airgun shots recorded by the land stations located on Shikoku Island (Land 29 93). (c) Land explosions recorded by the OBSs and the land stations. Solid triangles indicate the positions of the land explosions. 3.2.3 Land explosions recorded by OBSs and land stations As mentioned above, the densely deployed OBSs provide a conventional land survey type of record section using the onshore offshore seismic data (Fig. 9). All OBSs and land stations clearly recorded the explosion at the southern part of the onshore profile,

820 S. Kodaira et al. Figure 5. Observed airgun shots recorded by OBSs (top panel of each figure) and ray theoretical synthetic seismograms calculated from a final model (lower panel of each figure). A bandpass filter (5 20 Hz) was applied to the observed data. Amplitudes are scaled proportionally to the square root of offset. The reduction velocity is 8 km s 1. The horizontal axes represent the distance from the southern end of the profile and the source receiver offset at the bottom and the top, respectively. (a) OBS05. Pn, refraction phase from mantle; Po, refraction phase from the oceanic crust; So, converted S-wave phase. (b) OBS34. Psed, refraction phase from the sedimentary wedge; PmP, reflection from Moho. (c) OBS58. Pu, refraction phase from the island arc upper crust. (d) OBS70.

Factors controlling an earthquake in the Nankai trough 821 Figure 6. Observed airgun shots recorded by OBSs (top panel of each figure) and ray theoretical synthetic seismograms calculated from a final model (lower panel of each figure). A bandpass filter (5 20 Hz) was applied to the observed data. Amplitudes are scaled proportionally to the square root of offset. The reduction velocity is 8 km s 1. The horizontal axes represent the distance from the southern end of the profile and the source receiver offset at the bottom and the top, respectively. (a) OBS76. (b) OBS83. (c) OBS92. (d) OBS98.

822 S. Kodaira et al. Figure 7. Observed airgun shots recorded by land stations (top panel of each figure) and ray theoretical synthetic seismograms calculated from a final model (lower panel of each figure). A bandpass filter (5 20 Hz) was applied to the observed data. Amplitudes are scaled proportionally to the square root of offset. The reduction velocity is 8 km s 1. The horizontal axes represent the distance from the southern end of the profile and the source receiver offset at the bottom and the top, respectively. (a) Land92. PoP,reflection phase from the top of the oceanic crust. (b) Land73.

Factors controlling an earthquake in the Nankai trough 823 Figure 8. Airgun shots recorded on land. A bandpass filter (5 20 Hz) was applied to the observed data. Amplitudes are scaled proportionally to the square root of offset. The reduction velocity is 8 km s 1. The horizontal axes represent the distance from the southern end of the profile and the source receiver offsets at the bottom and the top, respectively. (a) Land57 from the middle of Shikoku Island. A high-amplitude reflection phase from the top of the oceanic crust (PoP) is clearly recognized. (b) Land35 on the north side of Shikoku Island. J2 (Fig. 9b). The seismic signal from J1 and J3 could not be observed by the OBSs located in the southern part of the offshore profile; however, the signal from J1 can be traced up to 110 km offset (Fig. 9a). Two major first arrivals are observed such as the air gun signal recorded on land. The two phases are interpreted as being refraction phases from the island arc crust (Pu) and from the uppermost mantle (Pn). The most remarkable feature of the sections (Fig. 9) is a high-amplitude reflection phase, which is interpreted as being the same reflection phase mentioned above (PoP). We can identify this phase from almost 0 100 km offsets on the J2 section (Fig. 9b). This means that the possible high-impedance contrast extends from the southern coastline to the northern part of Shikoku Island. The same reflection phase is also observed on the J1 section at 30 130 km offsets (Fig. 9a). Another significant phase is found at the southern part of the J2 section. A high-amplitude reflection phase is recognized after the Pn phase; i.e. 0.5 s after the Pn arrival at 120 km offset. This phase is interpreted as an upper-mantle reflection from the subducting oceanic lithosphere. The onshore profile was designed to be slightly oblique to the offshore profile, owing to the restrictions of roads to deploy the stations. The explosion points also had to keep away from highly populated areas. The three types of data mentioned above are considered as a combination of approximately linear profiles (Fig. 4). For example, we did not observe the most offline ray path, which consists of the air gun shot at the southern end of the profile and the land station located at the northern end of the profile. The largest offline ray paths we used are the air gun shots at the middle of the offshore profile recorded at the northern part of Shikoku Island. We therefore assumed a 2-D structure along the profile during modelling of the wide-angle seismic data. 4 MODELLING PROCEDURE An initial model was made by referring to previously obtained crustal images along part of the onshore offshore profile. The migrated seismic reflection section provided images down to the top of oceanic crust under the foot of the landward flank (Fig. 3). Kodaira et al. (2000b) present the result of a first-arrival refraction tomography that used data from the offshore part of the profile. They focused on the structure around the Tosa-bae and showed a fine tomographic image down to 25 km depth. Their tomographic image successfully showed a thick body with P-wave velocity (V p )of5 7 kms 1 below the sedimentary wedge. This thick body is proposed to be a subducted seamount (Kodaira et al. 2000b). Their study, however, did not include later reflection phases and amplitude modelling, because of a restriction of their modelling approach. Kurashimo et al. (in press) modelled traveltimes of the onshore part of the data used in this study. They showed the existence of subducting oceanic crust underlying the island arc crust below Shikoku Island. They also did not include amplitude information in their analysis. Furthermore, since those previous studies separated parts of the onshore offshore profile, the ends of the models (i.e. land sea boundary area) were not resolved. In this study we aimed to obtain a whole crustal transect that explains all observed traveltime data and the overall character of amplitude variations with offset by means of forward modelling (Zelt & Ellis 1988) and an inversion of traveltimes (Zelt & Smith 1992). The forward modelling and the inversion were performed in a layer-stripping way, in which the parameters of successively deeper layers are determined, while the parameters defining the shallower layers remain fixed. The first arrivals and major later reflection phases were digitized with uncertainties (0.05 0.3 s) depending on signal-to-noise ratios. The observed traveltime data were classified into seven groups when the layer-stripping procedure was carried out. The classified groups are: (a) a refraction phase from the sedimentary wedge; (b) a refraction phase from the oceanic crust; (c) a refraction phase from the island arc upper crust; (d) a reflection phase from the top of the oceanic crust; (e) a reflection phase from the base of the oceanic crust; (f) a refraction phase from the uppermost mantle; (g) refraction reflection phases from the island arc lower crust; and (h) a reflection phase from the uppermost mantle. These input data were modelled in the above order during the layer-stripping procedure.

824 S. Kodaira et al. Figure 9. Observed land explosion shots recorded by the OBSs and the land stations (top panel of each figure) and ray theoretical synthetic seismograms calculated from a final model (lower panel of each figure). A bandpass filter (5 20 Hz) was applied to the observed data. The observed data are scaled by the root mean square of background noise in order to reduce site-dependent noise level. The reduction velocity is 8 km s 1. The horizontal axes represent the distance from the southern end of the profile and the source receiver offset at the bottom and top, respectively. (a) Shot J1, (b) Shot J2. An advantage of the inversion is that it allows velocity and interface nodes to be specified independently, allowing the resolution of the final model to be estimated. The resolution value represents the diagonal elements of the resolution matrix and indicates relative ray coverage at the nodes. The node points with a resolution value >0.5 are considered to be well resolved (Zelt & Smith 1992). Even though resolution values are not directly related to absolute parameter uncertainties (Zelt & Smith 1992), we note that high resolution (i.e. higher ray coverage) is usually associated with lower uncertainty. On the basis of similar wide-angle seismic studies (e.g. Kodaira et al. 1998, 2000a), for example, we could consider empirically that velocity parameters with resolutions of greater than 0.5 cannot be varied by more than 0.1 km s 1. The final model was constrained by amplitude modelling, which is calculated by using zero-order asymptotic ray theory (Cerveny et al. 1977; Zelt & Ellis 1988). In this study, we only considered the overall character of the amplitude variations with offset by visual comparison, since there are large uncertainties in a quantitative analysis of the amplitude data (e.g. Mjelde et al. 1997). However, we paid special attention to the following two features, which were not considered during the traveltime modelling; the range of distances over which each refraction phase was observed and the location of

Factors controlling an earthquake in the Nankai trough 825 the critical point of each reflection phase. These features control the velocity gradient in each layer and the velocity contrast between layers, respectively. 5 SEISMOGENIC ZONE STRUCTURE A P-wave velocity model along the onshore offshore wide-angle profile is shown in Fig. 10(a). Fig. 11 presents diagrams of traveltime residuals calculated from this model. In this figure, we plot mean traveltime residuals of each phase group calculated for an OBS at its position, since it is not practical to show all traveltime curves for all OBSs and land stations used. In the following we describe the model in four parts; i.e. the sedimentary wedge, the island arc crust, the subducting oceanic crust and the uppermost mantle. 5.1 Sedimentary wedge The crustal model (Fig. 10a) indicates a thick sedimentary wedge (V p = 2.0 4.2 km s 1 ) beneath the Tosa-bae with a maximum thickness of 9 km. Recent wide-angle seismic surveys also found similar sedimentary wedges in the Nankai Trough; e.g. 8 km thick at 40 km landward of the trough axis at the eastern Nankai Trough (Nakanishi et al. 1994, 1998) and 7 km thick at 70 km landward of the trough axis in the western Nankai Trough to the southwest of Cape Muroto (Kodaira et al. 2000a). These sedimentary wedges are interpreted as a young (Miocene Pleistocene) accretionary prism on the basis of previous geological and seismic studies (e.g. Taira & Tashiro 1987; Taira et al. 1989, 1991). The sedimentary wedge gradually thins toward the Nankai Trough, and continues as a sedimentary layer on the oceanic crust before subduction. Landward, the sedimentary wedge is abruptly thinned 160 km from the southern end of the profile. Figure 10. (a) P-wave velocity model. The horizontal axis represents the distance from the southern end of the profile (Fig. 2). The thick dashed line at the top of the oceanic crust indicates a thin low-velocity layer (300 m thick, 3 km s 1 ). The velocities of the lower part of the lower crust and the upper most mantle beneath Shikoku Island are assumed from a previous land refraction study (Ikami et al. 1982). MTL, Median Tectonic Line. (b) Resolution values calculated from the traveltime inversion. Squares indicate resolution values for interface nodes. Grey-scale contours show a smoothed version of resolution values for velocity nodes.

826 S. Kodaira et al. Figure 11. Traveltime residuals of observed phase groups. The mean traveltime residual of each OBS and land station is plotted at its position. The mean residual for a land explosion is plotted at its position. Solid and open circles indicate the mean residuals for the arrivals observed at the north and the south of the stations/shots, respectively. Bars indicate standard deviations. Shaded areas indicate the maximum range of uncertainties of the observed arrival times (0.3 s). The number of the observed arrival times is indicated. T rms, root-mean-square of the traveltime residual for each phase group. (a) Refraction phase from the sedimentary wedge. (b) Refraction phase from the oceanic crust. (c) Refraction phase from the island arc upper crust. (d) Reflection phase from the top of oceanic crust. (e) Reflection phase from the base of oceanic crust. (f) Refraction phase from the uppermost mantle. (g) Refraction reflection phases from the island arc lower crust. The structure of the sedimentary wedge is obtained by modelling the refraction first arrival labelled Psed in Fig. 5. Traveltime fitting of this phase is shown in Fig. 11(a), which demonstrates the calculated traveltimes are fitted to within 0.25 s. The velocities in the sedimentary wedge are well resolved (resolutions > 0.5; e.g. Zelt & Smith 1992). The synthetic seismograms (Fig. 5) are also generally consistent with the offset where Psed phases are observed. This indicates that velocity gradients are well constrained. In the sedimentary wedge, the crustal model indicates two interfaces where velocity gradients change. We place these two interfaces in the sedimentary wedge, since large changes of velocity gradients are recognized in the tomographic result (Kodaira et al. 2000b). 5.2 Island arc crust We found a crustal block with V p = 5.4 6.6 km s 1 landward of the sedimentary wedge. This block thins toward the trough (the maximum thickness is 32 km beneath the Median Tectonic Line), and also seems to gradually thin northward. The model indicates four layers in this crustal block, i.e. V p = 5.4 6.0, 6.0 6.2, 6.4 and 6.6 km s 1. Similar structures are widely observed between the trough and the Japanese islands along the Nankai Trough; e.g. at the eastern Nankai Trough (Nakanishi et al. 1998), the central Nankai Trough (Nakanishi et al. 2002), and the western Nankai Trough (Kodaira et al. 2000a; Takahashi et al. 2002). We therefore call the upper two layers island arc upper crust, and the lower two layers island arc lower crust. The velocities and the geometry were obtained by the refraction arrival labelled Pu in Figs 5 7. Calculated traveltime residuals for Pu phase are within ±0.2 s (Fig. 11c) at the OBSs and the land stations located between 130 and 280 km from the southern end of the profile (Fig. 11c). The root-mean-square of the residuals is estimated to be 0.15 s (Fig. 11c), which is comparable to the observed uncertainty of the traveltimes. The model indicates changes in the velocity gradient in the island arc upper crust. The first layer shows a higher velocity gradient (e.g.

Factors controlling an earthquake in the Nankai trough 827 0.1 s 1 at the coast) than the second layer (e.g. 0.02 s 1 at the coast). The change in the velocity gradient (high gradient at the top and low gradient at the bottom) is necessary to explain the observed amplitude variation with offset of the Pu phase. For example, the Pn phase is recognized with high amplitude between 5 50 km offsets in the shot-gather record section of J2; however, the amplitude becomes weak at offsets greater than 50 km offsets (Fig. 9). This observed amplitude variation is reproduced well in the synthetic seismograms of J2 (Fig. 9b). The geometry of the top of the island arc crust and the velocities are only well resolved between 130 and 280 km in the first layer of the upper crust (Fig. 10b) caused by limitations of the observable range of the Pu phase (Fig. 11c). The resolution diagram implies a detailed discussion of the velocities is not allowed at the northern half of the upper crust. The clear wide-angle reflection phase from the intracrustal interface (e.g. PiP in Fig. 9b) is observed on the shot-gather record section of J2 and on several land stations. The interface between the island arc upper and lower crust is well constrained by the PiP phase between 150 220 km (Fig. 10b). Kodaira et al. (2000a) also observed a clear intracrustal reflection phase landward of the sedimentary wedge along a profile 50 km west of our profile. It is thus likely that a sharp interface exists between the island arc upper and lower crust. The traveltimes of the wide-angle reflections from the top of the subducting oceanic crust (PoP) were used to estimate the velocities in the upper part of the lower crust (V p = 6.4 kms 1 ), and the geometry of the subducting oceanic crust since there are no observed refracted arrivals from the lower crust. For this reason, the velocities of the lower crust are not as well constrained as those of the upper crust. We assumed a velocity of 6.6 km s 1 in the lower part of the lower crust by referring to a previous land refraction study on Shikoku Island (Ikami et al. 1982). In the shot-gather record section of J2, we recognize a wave train of between 325 and 355 km (Fig. 12c), even through the exact arrival time from that phase is not clear. We calculated traveltimes of the wide-angle reflections from a landward-dipping interface at the base Figure 12. Magnification of the J2 section showing deep reflection phases. (a) Observed section of 0 85 km from the southern end of the profile. Two major phases are indicated by arrows. (b) Calculated traveltime curve of a reflection phase from an uppermost mantle reflector (Fig. 12e) is superimposed. Pn, refraction from the uppermost mantle. (c) Observed section of 310 355 km from the southern end of the profile. The wave train can be recognized between the thick arrows. Onsets of this phase are not clearly observed. (d) Calculated traveltime curve of a reflection from the base of the island arc crust (Moho) (Fig. 12e) is superimposed. Weak arrivals are also recognized before the Moho reflection. Those arrivals are interpreted to be a refraction phase from the Moho. PmPi, reflection phase from the base of the island arc crust (Moho); Pni, refraction from the Moho. (e) Ray diagram for calculating the traveltimes of the deep reflections (Figs 12b and d).

828 S. Kodaira et al. of the island arc crust (32 27 km depth between 275 300 km in the model reference frame). These traveltimes (PmPi in Fig. 12d) are plotted slightly later than the observed wave train. 5.3 Subducting oceanic crust and uppermost mantle 5.3.1 From the trough to the landward flank (0 80 km) A simple oceanic crust structure is obtained from the trough to the foot of the landward flank (0 80 km in the model). The model consists of three layers, which are interpreted to be a sedimentary layer, and oceanic layers 2 and 3 based on their thicknesses and velocities. The sedimentary layer shows velocities of 1.8 3.2 km s 1, and thickens toward the land, e.g. 1.5 and 2.5 km thick at 50 and 70 km, respectively. The igneous oceanic crust (oceanic layers 2 and 3) shows a slightly different character from a normal igneous oceanic crust. The model indicates the velocities of oceanic layers 2 and 3 are V p = 3.8 6.0 and 6.4 7.0 km s 1, respectively. These values are approximately within the bounds of the velocities of normal oceanic layer 2 (V p = 2.5 6.6 km s 1 ) and layer 3 (V p = 6.6 6.8 km s 1 ) (e.g. White et al. 1992), except for the slightly broader velocity range in layer 3 of our model. The thickness of the igneous crust in this area (5.0 6.3 km thick) is, however, remarkably thinner than the average thickness of the igneous oceanic crust. White et al. (1992) concluded that the igneous section of normal oceanic crust averages 7.1 ± 0.8 km thick on the basis of a compilation of a large number of seismic refraction results. Similar thinner oceanic crust has been reported from a wide area along the central to the western Nankai Trough (Yoshii et al. 1973; Nishizawa & Suyehiro 1989; Kodaira et al. 2000a; Nakanishi et al. 2002), e.g. 5.2 5.7 km thick at central Nankai (Nishisaha et al. 2002), 5.3 6.3 km thick at the western Nankai (Kodaira et al. 2000a). From those studies we could, therefore, note that the thinner oceanic crust might be a significant characteristic of the Nankai Trough. The structure down to the top of the oceanic crust in this area was well controlled by the multichannel reflection study and the first-arrival refraction tomography study (Kodaira et al. 2000b). It was not, therefore, necessary to modify the model significantly from the initial model in this part. The multichannel reflection study and the tomographic study, however, did not image the crust mantle boundary structure clearly (i.e. the shape of the Moho). The wideangle reflection phase (PmP in Figs 5 7) provided a precise depth of the Moho, and consequently the thickness of the oceanic crust could be estimated using these data. As indicated in the resolution diagram, the depth of the Moho is estimated with high resolution (>0.75), since a clear PmP phase is observed by almost all OBSs in this area (Fig. 11e). 5.3.2 Subducted seamount (80 150 km) We found anomalously thickened oceanic crust beneath the Tosabae, which is interpreted to be a subducted seamount. Kodaira et al. (2000b) have already obtained an image of the thickened crustal body beneath the Tosa-bae by applying first-arrival refraction tomography to the offshore part of our data. However, the tomographic result shows that the northern edge and deepest part of their model were not well constrained because of poor coverage of seismic rays. This is because of limitations of the data and the modelling approach they used. This study intended to improve the seismic image of those areas by additionally including data from the land stations and the major later arrivals in the data. The new data set successfully imaged the base of the seamount and its landward continuation (Fig. 10). The model indicates the anomalously thick oceanic crust, which is interpreted as a subducted seamount, between 85 and 145 model km. The maximum thickness is 15 km at 120 km (Fig. 10a). This is more than twice as thick as the oceanic crust before subduction. The model also indicates that the thickest part is located 20 km seaward of the peak of Tosabae. We do not find any significant velocity variation in the crust except for a slightly higher velocity at the top of the crust, i.e. the P-wave velocities are 4.6 6.0 and 6.5 7.0 km s 1 in layers 2 and 3, respectively. As mentioned above, the island arc upper crust (V p = 5.4 6.2 km s 1 ) appears beneath Tosa-bae (Fig. 10), immediately landward of the subducted seamount (130 model km). The structure beneath Tosa-bae seems to show that at present the subducted seamount is colliding with the island arc crust beneath the sedimentary wedge. The resolution diagram (Fig. 10b) and traveltime residuals (Fig. 11) show that the entire area of the subducted seamount is well constrained. All resolution values of the interfaces and velocities are estimated to be more than 0.5 in this area. This structure was obtained from the refracted phases from the oceanic crust (Po in Figs 5 and 6), and the uppermost mantle (Pn), and from the reflection from the base of the oceanic crust. As shown in Figs 11(b), (e) and (f), these phases fit the model well in the region of the subducted seamount. The resolution diagram only shows smaller values from the landward foot of the subducted seamount (135 145 km) as an exception. This poorer constraint comes from the absence of a refracted phase traversing this part of the model, since a velocity inversion with depth exists at the top of the crust. Owing to the large velocity contrast between the sedimentary wedge (V p = 2.0 4.2 km s 1 ) and the island arc upper crust (V p = 5.4, 6.2 km s 1 ), refracted phases from each layer are easily distinguished (e.g. Psed and Pu in Fig. 5c). Therefore, at least the upper part of the island arc crust (V p = 5.4 kms 1 layer) could be well determined by clearly observed Pu phases. This is also recognized in the resolution diagrams, which show higher resolution of geometry at the top of the V p = 5.4 kms 1 layer. Because of this higher resolution we believe the interpretation concerning the collision of the subducted seamount is valid, even though the lower part of the upper crust (V p = 6.2 kms 1 layer) shows lower resolution in this region. Another significant structure is obtained below the subducted seamount. The observed refraction phase from the uppermost mantle provided a significantly low P-wave velocity (7.5 7.8 km s 1 ) in the uppermost mantle beneath the subducted seamount. The refraction arrivals used to obtain the mantle velocity (Pn in Figs 5 7) were observed by both the OBSs and the land stations located along the southern part of the onshore profile (Fig. 11f). This implies that the velocity of the uppermost mantle could be estimated with high resolution for the offshore part of the profile. The resolution diagram clearly demonstrates this, i.e. higher-resolution values (>0.5) exist immediately below the subducting oceanic crust at 40 135 model km. From these diagrams it can therefore be seen that there is a significant velocity decrease (7.8 to 7.6 km s 1 ) below the southern edge of the subducted seamount. However, the resolution diagram shows that the landward extension of the low-velocity material is uncertain. In order to explain the traveltimes of the upper-mantle reflection phase (Fig. 12), two uppermost mantle interfaces are necessary at 28 32 and 40 km depth (Fig. 12). The reflection phase from the uppermost mantle interface is observed only in the J2 section (Figs 9b and 12a). Thus, we did not determine the exact shape and

Factors controlling an earthquake in the Nankai trough 829 depth of the interface with sufficiently high resolution. It could be mentioned, however, that an upper-mantle interface exists below the subducting sea mount, and that the calculated traveltimes fit the observed data well when the interface is placed 15 km below the Moho (Figs 12b and e). 5.3.3 Subducting oceanic crust beneath Shikoku Island (150 355 km) The subducting oceanic crust is traced down to 33 km depth beneath Shikoku Island by using the onshore offshore seismic data (Fig. 10). The structure of subducting oceanic crust in this area is determined by the refraction from the oceanic crust (Po), and the reflections from the top (PoP) and the bottom (PmP) of the oceanic crust. The model indicates smoothly dipping oceanic crust with a subduction angle of 7 at 15 20 km depth, while in the deeper part, the oceanic crust is imaged with a steeper angle, 11 at 20 33 km depth. We do not find any velocity variation with the subduction of oceanic layer 3 (V p = 6.5 6.9 km s 1 ). These velocities in the subducting oceanic crust are mainly controlled by the refraction phase Po, which is widely observed by the OBSs and the land stations (Fig. 11b). The resolution diagram also shows that the velocities in oceanic layer 3 are estimated with resolution values of more than 0.5 down to 30 km depth. The velocities of subducting layer 2 are estimated with poorer resolution than those in oceanic layer 3, since there are no refracted waves turning within this layer because of a velocity inversion with depth (low-velocity layer). The velocity and the thickness of oceanic layer2(v p = 4.6 6.0 km s 1 ) are, however, constrained by Po, PmP and Pn phases. The geometry of the subducting oceanic crust is constrained by the clearly observed reflection phases (PoP and PmP) with resolutions of more than 0.5 (Fig. 10). The reflection phase interpreted to be from the top of the oceanic crust is observed with extremely high amplitude from almost zero offset to large offsets (e.g. Fig. 9). The observed amplitude variation with offset could not be explained by a crustal model that has no additional material between the subducting oceanic crust and the island arc crust (Fig. 13b). There are many structural factors affecting the amplitude of the reflection phase; e.g. the attenuation factor (Q value), anisotropy, 3-D inhomogeneity, velocity contrast and density contrast. A full quantitative analysis of all of these factors has not been performed, because of the limitations of the zeroth-order asymptotic ray theory we used in this study and the limitation imposed by the quality of the data. Therefore, we tested the effect of only one factor, the P-wave velocity contrast, on the amplitude variation as the simplest case. In order to generate a high-amplitude reflection, it was necessary to place a large velocity contrast between the island arc crust and the subducted oceanic crust. A thick additional layer prevented the traveltimes from fitting. Therefore, we added a thin layer (300 m), and then varied the velocity in this layerfrom6to3kms 1. The variation of velocity, 3 6 kms 1,is comparable with approximately 0.1 s traveltime difference, which is within the observed uncertainty. The results of this test are shown in Fig. 13. The model with 6 km s 1 in the thin layer provides no visible reflection arrivals, especially near offsets (Fig. 13f). Figs 13(c) (f) clearly demonstrate that the amplitudes of the reflection arrivals become higher as the velocity decreases in the thin layer. The synthetic seismograms from the model having a 3 km s 1 thin layer shows the highest-amplitude reflection arrivals, which is comparable with the amplitude of the first arrival, even near offset. This test does not provide an exact velocity and thickness of the additional layer between the plates. However, it does underline that a model with a thin low-velocity layer between the plates is one of the most likely structures to explain the observed high-amplitude reflection arrivals (PoP). The highamplitude reflection arrivals are only observed by the shot-gather of the land explosions and the land stations at 180 265 model km. From a ray diagram we estimated that reflection points of this phase only extend over 150 240 model km (Fig. 14). 6 DISCUSSION 6.1 Subducted seamount and its role in the rupture process Since the late 1960s the rupture process of the 1946 Nankaido earthquake has been discussed by many seismologists. The main issues are how large was the co-seismic slip zone and how the co-seismic slip extended. Mogi (1968) showed a 1 day aftershock area extending only between Cape Shiono and Cape Muroto (Fig. 15). Kanamori (1972) also estimated the size of the rupture zone to be 80 120 km 2 by using seismic wave data. The size of the rupture zone that he estimated is comparable with the 1 day aftershock zone. On the other hand, Ando (1975) proposed a co-seismic slip zone extending from Cape Shiono to Cape Ashizuri using tsunami wave data. Ando (1975) also mentioned that the slip behaviour might be different between the eastern half (B in Fig. 2) and the western half (A in Fig. 2) of the rupture zone: i.e. brittle rupture in the eastern half and slow slip in the western half. His later study (Ando 1982), however, mentioned that a shorter rise time (less than 2 min) in the western half could explain the tsunami, geodetic and seismic intensity data. Recently, a precise rise time distribution was estimated for 60 60 km 2 subfaults at the rupture zone (Kato & Ando 1997). These authors obtained a rise time of 3 4 min in the western half and 0 min in the eastern half of the rupture zone, except for the most eastern subfault where the rise time was estimated to be 5 min. On the basis of strong ground motion data, Hashimoto & Kikuchi (1999) examined three subevents; the first was off the Kii peninsula (Fig. 2) with M = 6, the second was slightly north of the first subevent with M = 8 and the third occurred 53 s after the second subevent between Cape Muroto and Cape Ashizuri with M = 8 (Figs 2 and 15). They found slow ground motions ( 1 min) that corresponded to the third subevent in the strong ground motion seismographs. Recent progress in developing an inversion technique provided precise distributions of the co-seismic slip (Fig. 15) based on tsunami data (Tanioka & Satake 2001) and geodetic data combined with the tsunami data (Sagiya & Thatcher 1999). These tsunami and geodetic studies also show remarkable difference of the slips in the eastern and western half of the rupture zone. Tanioka & Satake (2001) estimated slips of 3 m extending over the entire rupture zone, from down-dip to up-dip, in the eastern half, and a large slip of 6 m that occurred only at the down-dip end of the rupture zone in the western half (Fig. 15). From these studies, the rupture process of the 1946 Nankaido earthquake may be summarized as follows. The initial break started off Cape Shiono and extended over the entire locked zone between Cape Shiono and Cape Muroto with a short rise time, then the rupture jumped around the down-dip end of the locked zone between Cape Muroto and Cape Ashizuri with a significantly large rise time. A remaining question is what structural factor may control such a complicated dynamic rupture process. We believe the structure we obtained could provide one of the possible answers to that question. A location map of the subducted seamount (Fig. 15) shows that the subducted seamount is just outside the 1 day aftershock zone,

830 S. Kodaira et al. Figure 13. (a) Magnification of the J2 section, 210 270 km from the southern end of the profile, showing a prominent reflection phase. (b) Synthetic seismograms calculated from a model without any additional layer between the island arc crust and the oceanic crust. No visible reflection phase is recognized. (c) (f) Synthetic seismograms calculated from models having a thin layer between the plate at 160 240 km, where velocities in the thin layer were varied from 3 (Fig. 13c) to 6 km s 1 (Fig. 13f). The model having a velocity of 3 km s 1 in the thin layer (Fig. 13c) shows the highest amplitude, from almost zero offset among the models. and is located between the eastern and western half of the rupture zone (Figs 2 and 15). By comparing the location of the seamount and the slip distribution determined by Tanioka & Satake (2001) we recognized that the subfaults with slips of 3 4 m (Figs 7A C and 6A C in Tanioka & Satake 2001) are bounded on the west by the seamount, and west of these subfaults the larger slips of 3 6 m seems to extend around the landward foot of the seamount. It could also be mentioned that the seamount is located between the shorter

Factors controlling an earthquake in the Nankai trough 831 Figure 14. Ray diagram corresponding to the observed high-amplitude phase from between the plates. This diagram shows that the highly reflective layer is only located beneath Shikoku Island, 150 240 km from the southern end of the profile. Every 50 rays are plotted for airgun shot. Figure 15. Comparison between the location of the subducted seamount, the 1 day aftershock area (Mogi 1968), co-seismic slips obtained from tsunami data (Tanioka & Satake 2001) and a sequence of subevents obtained from strong ground motion seismograms (Hashimoto & Kikuchi 1999) of the 1946 Nankaido earthquake. The grey-scale shows the co-seismic slips (Tanioka & Satake 2001). Italic numbers and dashed arrows indicate the sequence of the subevents (Hashimoto & Kikuchi 1999). The star shows the epicentre of the 1946 Nankaido earthquake (Kanamori 1972). The off-profile dimensions of the seamount are estimated by reference to the size of the Kosyu seamount that is the northern most seamount of the Kinan seamount chain.

832 S. Kodaira et al. rise time subfaults and the longer rise time subfaults estimated by Kato & Ando (1997). The subevent sequence obtained by the strong ground motion data (Hashimoto & Kikuchi 1999) also seems to extend around the landward edge of the subducted seamount. From these comparisons, we propose a possible scenario of the rupture process of the 1946 Nankaido earthquake: the subducted seamount functioned as a barrier to the shorter rise time slip starting off Cape Shiono, and the longer rise time slip might propagate around the landward edge of the subducted seamount. Several researchers have studied the effects of the subducting seamount on earthquake processes (e.g. Kelleher & McCann 1976; Cloos & Shreve 1996; Scholz & Small 1997). Scholz & Small (1997) summarized those studies and concluded that the coupling between the subducting plate and the overriding plate becomes locally stronger caused by the subduction of a seamount with dimensions of several tens of kilometres. Our study strongly supports their conclusions. 6.2 A possible trapped water layer and its origin A thin, highly reflective, and possibly low-velocity, layer is inferred from the observed high-amplitude reflection phases (Figs 9 and 14). The model shows that the thin low-velocity layer lies between the subducting plate and the overriding plate within an area that extends from 30 km seaward of the coastline to beneath the Median tectonic line (Fig. 16). We propose trapped water between the plates as a possible interpretation of the highly reflective/low-velocity layer by referring to a previous magnetotelluric (MT) study. Yamaguchi et al. (1999) obtained a 2-D resistivity structure from 80 km seaward of the coast to the north of Shikoku Island, which was almost along our seismic profile. Their model indicates a high-resistivity subducting lithosphere (5000 15 000 m) and a low-resistivity layer (75 m) at the top of the lithosphere. The low-resistivity layer extends from offshore Shikoku to beneath the Median Tectonic Line. Yamaguchi et al. (1999) concluded, from tests of several models, that the lowresistivity layer is restricted only to the south of the Median Tectonic Line no extension exists to the north (Fig. 16). The extent of their low-resistivity layer is plotted in Fig. 16. It is considered, from this figure, that the highly reflective/low-velocity layer corresponds to the low-resistivity layer. Similar low-resistivity layers have also been found at several subduction zones; for example, at Vancouver Island (Kurtz et al. 1986), the Coast Range and the westernmost Williamelle Basin (Wannanmaker et al. 1989), and in northern Japan (Utada et al. 1996). Hyndman (1988) concluded that there is trapped water in the low-resistivity layer corresponding to a highly reflective layer at Vancouver Island. We believe that we could apply the same interpretation to our obtained highly reflective/low-velocity layer with the low resistivity; i.e. trapped water in the highly reflective/lowvelocity layer with the low resistively beneath Shikoku Island. Dehydration from subducted uppermost mantle beneath Shikoku Island would be one of the possible origins of the trapped water. Seno & Yamanaka (1996) investigated the effect of a hydrated mantle on intraslab seismicity. According to their study, mantle material affected by a plume would be partially serpentinized, arising from the accumulation of dissolved water through a partial melting zone in the plume, and subduction of the serpentinized mantle would generate an increased brittle response caused by dehydration, which causes the intraslab seismicity. Seno et al. (2001) applied these results to an interpretation of mantle earthquakes beneath southwestern Japan; i.e. the mantle earthquakes beneath a part of western Japan would be caused by the dehydration-increased brittle behaviour in the serpentinized mantle associated with the past backarc igneous activity in the Izu Bonin arc. In the following section, we apply Seno and Yamanaka s mechanism of dehydration-increased brittle behaviour in order to consider an origin for the possible trapped water. Sakamoto & Kim (1999) analysed rock samples collected from Kosyu seamount, which is the northernmost seamount of the Kinan seamount chain (Figs 2 and 15). They concluded, from their chondrite normalized rare-earth element diagram, that the rock sample from Kosyu seamount originated from plume-type magma (P-MORB). Magnetic studies also show that the igneous activity, that generated the Kinan seamount chain, occurred 4 7 Ma as off-ridge igneous activity after backarc spreading of the Shikoku Basin had ceased (Kobayashi et al. 1995; Okino et al. 1999). If we follow Seno and Yamanaka s study, serpentinized mantle could be found beneath the Kinan seamount chain because of an effect of the plume. A P-wave velocity reduction in the mantle would be one piece of evidence for serpentinization (Hess 1962; Christensen 1966). Seno et al. (2001) referred to the uppermost mantle velocity of V p = 7.69 km s 1 as the serpentinized part of the slab mantle beneath the Kanto (Hori et al. 1985). This velocity is comparable with the mantle velocity we obtained (V p = 7.5kms 1 ) beneath the subducted seamount. We, therefore, believe that serpentinization associated with the past igneous activity could be one of the explanations for the low mantle velocity beneath the Figure 16. Interpretation of the model in Fig. 10(a). Seismic activity (after Nakamura et al. 1997) along the profile is plotted on the model. The arrows along the top of the model represent the extents of the low-resistivity layer (Yamaguchi et al. 1999) and the local absence of co-seismic slip during the 1944 Tonankai and the 1946 Nankaido earthquake obtained by geodetic data (Sagiya & Thatcher 1999). Both the no co-seismic slip area and the low-resistivity area show good correspondence with the location of our observed highly reflective layer. A layer of trapped water (thick line) between the plates is considered to be a possible interpretation of the highly reflective/low-resistivity layer. Ages of accretionary prisms are simplified after Taira et al. (1996). MTL, median tectonic line.