Crustal structure across the coseismic rupture zone of the 1944 Tonankai earthquake, the central Nankai Trough seismogenic zone

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 17, NO. B1, 7, 129/1JB424, 2 Crustal structure across the coseismic rupture zone of the 1944 Tonankai earthquake, the central Nankai Trough seismogenic zone Ayako Nakanishi, 1 Narumi Takahashi, 2 Jin-Oh Park, 1 Seiichi Miura, 1 Shuichi Kodaira, 1 Yoshiyuki Kaneda, 1 Naoshi Hirata, 3 Takaya Iwasaki, 3 and Masao Nakamura 3 Received 5 May ; revised 27 November ; accepted 27 June 1; published 16 January [1] Differences in the coseismic rupture process between the 1944 Tonankai and the 1946 Nankai earthquakes have been studied by many fault models. To understand what factors control coseismic rupture zones, it is important to investigate differences in deep crustal structures of the rupture zones between the 1944 and 1946 earthquakes. The previously published crustal structure of the rupture zone of the 1946 earthquake shows that the coseismic rupture extends to the Neogene-Quaternary accretionary prism. However, little is known about the structure of the rupture zone of the 1944 earthquake. To obtain a complete image of the seismogenic zone of the 1944 earthquake, a wide-angle seismic survey was performed across the presumed coseismic rupture zone of the 1944 earthquake from ocean to land. Our model for the crustal structure is based on ocean bottom seismographic data. The crustal structure appears characteristic for subducting oceanic crust and a Neogene-Quaternary accretionary prism bounded by an island arc crust. The Neogene-Quaternary accretionary prism reaches a maximum thickness of 7 km at 5 km distance landward from the deformation front. The subducting oceanic crust can be traced down to 35 km. The subduction angle becomes steeper landward, reaching up to 11 beneath the island arc crust. The depth of the top of subducting oceanic crust at the downdip limit of the rupture zone is 23 km, while the updip limit is located beneath the island arc upper crust. Similar structures of the updip and downdip limits are also published for several other subduction zones. INDEX TERMS: 325 Marine Geology and Geophysics:Marine seismics (935), 815 Tectonophysics: Plate boundarygeneral (34), 93 Information Related to Geographic Region: Asia; KEYWORDS: crustal structure, ocean bottom seismograph, Nankai Trough, seismogenic zone, great earthquake Introduction 1 Institute for Frontier Research on Earth Evolution, Japan Marine Science and Technology Center, Yokosuka, Japan. 2 Deep Sea Research Department, Japan Marine Science and Technology Center, Yokosuka, Japan. 3 Earthquake Research Institute, University of Tokyo, Tokyo, Japan. Copyright 2 by the American Geophysical Union /2/1JB424$ [2] The Nankai Trough is one of the active convergent boundaries where historic great earthquakes were generated repeatedly. These large thrust earthquakes in the Nankai Trough are attributed to the subduction of the Philippine Sea plate beneath the overlying Eurasian plate (Figure 1). On the basis of previous geodetic and seismological studies, there are remarkable regularities in the space-time distribution of great earthquakes along the Nankai Trough. The recurrence interval between the great earthquakes is estimated at 1 years [e.g., Ando, 1975]. Such large thrust earthquakes recur at a certain portion of a plate contact zone with its updip and downdip limit [e.g., Tichelaar and Ruff, 1993], which is called a seismogenic zone. However, little is known about the nature of earthquake mechanisms in the seismogenic zone. Several previous studies had given important informations to progress our understanding the seismogenic zone so far. On the basis of the geodetic and tsunami data, Ando [1975] estimated the coseismic rupture (slip) distribution of the historic great earthquakes as a simple model consisting of four ruptured fault planes denoted by A, B, C, and D from west to east, respectively (Figure 1). These fault planes are recognized as the coseismic rupture zone [Ando, 1975]. The last great earthquakes occurred in 1944 (Tonankai earthquake, M s = 16) and 1946 (Nankai earthquake, M s = 19); both hypocenters were located off the Kii Peninsula at depth of 3 km, with a large uncertainty from to 4 km depth [Kanamori, 1972]. They are both interpreted as low-angle thrust faults at the plate boundary with oceanic side underthrusting northwestward against southwestern Japan [Kanamori, 1972; Ando, 1975]. It is known that the rupture zone of the 1946 event covers both areas A and B, although the 1944 event only ruptured area C. It is, however, not clear what factors control the size of the coseismic rupture zone. Furthermore, it has been proposed that there is significant difference in rupture processes of the rupture zones [Kanamori, 1972; Ando, 1975]. Two contradicting results for the coseismic rupture zone of the 1946 event are derived independently from seismic and geodetic data. Areas A and B are estimated as a coseismic rupture zone from the geodetic and tsunami data [Ando, 1975], whereas only area B is estimated from the seismic wave data reported in the world, e.g., International Seismological Summary [Kanamori, 1972]. Area B also corresponds to the 1-day aftershock area [Mogi, 1968]. Considering these facts, Ando [1975] proposed that the coseismic rupture zone is divided into two parts (e.g., an eastern brittle block (area B) and a western ductile block (area A)) and suggested that the mode of fracturing occurred first in area B, exciting the seismic and tsunami wave efficiently, and then the fracture propagated slowly westward, causing extensive crustal deformation but EPM 2-1

2 EPM 2-2 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE Figure Map of area around the Nankai Trough. The wide-angle seismic survey was conducted along profile KR98 The air gun was shot along the solid line. Numbered circles indicate OBS positions. Open stars indicate stations on land. MCS data were acquired along the thin white line between OBSs 6 and 1 MO14 and KR981 indicate previous OBS profiles [Kodaira et al., ; Takahashi et al., 1999]. Rectangular areas A to D show coseismic rupture zones of great earthquakes [Ando, 1975]: Areas A and B were ruptured by the 1946 Nankai earthquake, and area C was ruptured by the 1944 Tonankai earthquake. Solid stars indicate epicenters of the 1946 and 1944 events, respectively. Area D has not experienced a major earthquake since the 1854 Tokai earthquake. Onshore geology of the Shikoku Island (Shikoku Is) and Kii Peninsula (Kii Pen) is indicated by Geological Survey of Japan [1992]. MTL, Median Tectonic Line; BTL, Butsuzo Tectonic Line; Sa, Sanbagawa Belt; Chi, Chichibu Belt; NS, Northern Shimant belt; SS, southern Shimanto Belt; PSP, Philippine Sea plate; EUR, Eurasian plate; T, Tosa Basin; M, Muroto Basin; K, Kumano Trough; E, Enshu Trough. weak seismic and tsunami waves. In sharp contrast to the 1946 event the areas of aftershocks, the tsunami generation, and the crustal deformation are all consistent with each other for the 1944 event, which corresponds to area C. This result suggests that the breakage at the 1944 event in area C was predominated by brittle rapid fracturing [Ando, 1975]. [3] Hyndman et al. [1995] proposed a model of a thermally constrained locked zone for the Nankai Trough during the interseismic period (interseismic locked zone), which is almost consistent with the coseismic rupture zone, as mentioned above. The thermal model is calculated by the finite element method [Hyndman and Wang, 1993] based on the extensive heat flow data [Hyndman et al., 1995]. Hyndman et al. [1995] also proposed that a computed elastic dislocation model based on the thermal results agrees with the coseismic rupture zone of Ando [1975]. Their thermal model indicates that the zone of locked stick slip where earthquakes can occur is limited downdip by a temperature of 35 C and by the transition to the stable sliding zone into which coseismic displacement can reach temperatures up to 45 C. The updip limit of the stick-slip zone is estimated at 1 15 C. Moreover the downdip limit of the interseismic locked zone corresponds well to that of the coseismic rupture zone of the 1944 event determined by Ando [1975]. However, the updip limit of the thermally constrained locked zone is 25 km seaward from that of the coseismic rupture zone of the 1944 event. [4] Both the coseismic rupture and interseismic locked zones can be considered as indicators defining the seismogenic zone. To understand what factors control the updip and downdip limit of the seismogenic zone, generation of great earthquakes, and rupture

3 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE EPM 2-3 Figure (a) Time-migrated MCS section acquired between OBSs 6 and 14 on profile KR986, (b) its interpretation, and (c) depth-converted MCS section and its interpretation. Positions of OBSs are shown by numbered solid circles. Depth-converted section is plotted at the same scale (both depth and distance) as our crustal morel as shown in Figure

4 EPM 2-4 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE Figure Observed and synthetic seismograms at OBS 13: (a) Observed seismograms and (b) synthetic seismograms. Observed seismograms, which were recorded on the vertical component, have been digitally band-pass filtered (5 1 Hz). Reduction velocity is 8 km/s. Trace amplitudes are scaled by the square root of the shot distance. processes, it is thus important to investigate differences and similarities in deep crustal structures of the coseismic rupture zones between the 1944 and 1946 earthquakes. [5] Kodaira et al. [] present a detailed model of the deep crustal structure across area A, which is the coseismic rupture zone of the 1946 Nankai earthquake [Ando, 1975]. The model is characterized by a gentle slope of the subducting oceanic crust and thick overlying sedimentary wedge. It reaches a maximum thickness of 9 km at 7 km from the deformation front. The subduction angle of the subducting oceanic crust becomes steeper landward with value of 7 beneath the island arc crust. Kodaira et al. [] also show that the downdip limit of the rupture zone does not extend to the deep end of the contact zone between oceanic crust and island arc crust and the updip limit extends to the Neogene-Quaternary accretionary prism. Moreover, variation of the crustal structure along the Nankai Trough was obtained from crustal surveys along the seaward edge of area B [Mochizuki et al., 1998; Sato et al., 1998]. Mochizuki et al. [1998] found an abrupt change in thickness of the subducting plate at the location between areas B and C. Sato et al. [1998] also found depth variation of the subducting plate at the southwestern edge of area B. They suggested that these structural change might affect the process of interplate earthquake occurrence. [6] The geometry of the subducting oceanic plate (slab) is estimated by studies of hypocentral distributions [e.g., Nakamura et al., 1997]. However, Nakamura et al. [1997] determined the slab geometry only beneath southwestern Japan using the land seismic network. It is impossible to clarify the slab geometry off the coast around the Nankai Trough area without a deep crustal survey, since the seismic activity is known to be very low. [7] This paper presents the crustal and uppermost mantle structure across area C, the coseismic rupture zone of the 1944 Tonankai earthquake [Ando, 1975], mainly by use of wide-angle ocean bottom seismographic (OBS) data. The OBS profile location is shown in Figure [8] In the survey area, there is a forearc basin, the Kumano Trough (Figure 1), and an outer ridge, which is a topographic high along the seaward slope of the forearc basin and traps the basin sediments [Mogi, 1975]. Similar forearc basins with outer ridges are developed on the continental slope along the Nankai Trough; they are Tosa Basin, Muroto Basin, Kumano Trough, and Enshu Trough from west to east (Figure 1). Awata and Sugiyama [1989] interpreted these forearc basins as individual geological units, and the area of each unit corresponds with the presumed coseismic rupture zone of great earthquakes, areas A D of Ando [1975]. In particular, a well-developed outer ridge is recognized off the Kumano Trough. The outer ridge off the Kumano Trough consists of sediments of Neogene and Quaternary age uplifted in the Paleogene to Miocene [e.g., Okuda, 1984]. Remarkable coincidences between the geological structure of outer ridges and the structural high of the forearc model [e.g., Dickinson and Seely, 1979] were pointed out by Sakurai and Sato [1983]. They speculated that the outer ridge off the Kumano Trough is the uplift caused by the landward understuffing of the subduction complex.

5 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE EPM 2-5 Figure Observed and synthetic seismograms at OBS 1: (a) observed seismograms and (b) synthetic seismograms. However, it is not clear how the existence of the outer ridge can be related to the crustal structure and the coseismic rupture process of the seismogenic zone. [9] In this paper our crustal model is compared with the crustal model across the rupture zone of the 1946 Nankai earthquake [Kodaira et al., ] to discuss differences in crustal structures. The spatial relationship of the crustal structure and presumed coseismic rupture zone of great earthquakes between areas A and C is also examined. Moreover, we discuss the spatial relationship of the outer ridge off the Kumano Trough and the rupture zone of the 1944 event. Our crustal model is finally compared with other several subduction seismogenic zones to discuss characteristics of the structure of the seismogenic zones. Data Acquisition [1] In June July 1998 the Japan Marine Science and Technology Center (JAMSTEC), in cooperation with the University of Tokyo, performed a seismic survey across the central Nankai Trough and the Kumano Trough using R/V Kairei of JAMSTEC. Figure 1 shows the locations of the multichannel seismic (MCS) reflection and the refraction and wide-angle reflection profiles of this study. Fourteen OBSs were deployed along a -km-long profile with a spacing of 1 15 km. The MCS reflection data were acquired using a 1-channel hydrophone streamer (3 m long) along the OBS profile from OBS 6 to OBS 1 We used four air guns (each 17 L) as controlled sources. They were shot every 5 m one way and m return with a pressure of 14 MPa. With a shot time interval of over 4 s on both ways of the 5 m and m shooting, the noise generated by the direct water wave of previous shots was reduced enough to ensure good quality of the OBS data. To investigate the deep crustal structure beneath the Japanese island arc, air gun shots were also recorded by three temporary stations (KWR, KHR, and AIZ in Figure 1) and existing telemetered stations (INN in Figure 1) along the landward extension of the OBS profile. The total length of the wide-angle profile is 242 km. [11] The OBSs used in this survey were developed by Shinohara et al. [1993]. The OBS represents the digital recording version of an OBS originally designed by Kanazawa [1986] and Kanazawa and Shiobara [1994]. The OBS has three-component gimbal-mounted geophones (5 Hz) and a hydrophone. The digital recorder has a 16-bit A/D converter and stores data on a digital audio tape at 1 Hz and can store 17 days of continuously recorded data. [12] Because of strong sea currents, a free-fall-type OBS drifts away during descent; hence the location of the OBS sites may differ from the ship s position at the time of deployment. To determine the correct OBS position, we measured the travel time between the ship and the OBS from different positions using an acoustic transponder system that is mounted on each OBS mainly for the sinker releasing. Using those data as well as the arrival times of the direct water wave from the air gun shots, we redetermined the seafloor OBS positions by a nonlinear inversion method [Shiobara et al., 1997]. Moreover, in this calculation we considered one-dimensional acoustic velocity structure calculated from conductivity-temperature-depth (CTD) data and expendable

6 EPM 2-6 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE Figure Observed and synthetic seismograms at OBS 7: (a) observed seismograms and (b) synthetic seismograms. bathythermographic (XBT) data obtained around the survey area. The accuracy of determined OBS positions is enough for present OBS survey, which becomes order of 1 m. Data Multichannel Reflection Data [13] Figure 2 shows the time-migrated and depth-converted MCS reflection sections. The depth-converted section (Figure 2c) is consistent with our crustal model derived from wide-angle OBS data as mentioned in sections 2 and The section clearly images the sedimentary layers around the trough, a decollement zone, the shallow parts of the accretionary sediments, and the top of the subducting oceanic crust. It seems that the Shikoku Basin sediments, which are divided into two layers, start to subduct landward from OBS 1 The deformation front is recognized between OBSs 9 and The decollement and the top of the oceanic crust can be traced landward to the outer ridge at 3 km distance from the deformation front. Although the decollement zone, having twoway travel time of.2 s, can be clearly imaged in the MCS section, we have not included northwest of the deformation front in our model. The subducting sediments beneath the decollement are considered to be a low-velocity layer, and their seismic velocities could not be determined by t-p method as mentioned in section Moreover, the thickness of the subducting sediments can be estimated to be too thin to be resolved by the OBS data. Several reflectors can be recognized in the accretionary sediments between OBSs 8 and We interpret these reflectors as thrust faults. One of those reflectors might be an out-of-sequence thrust (OST). However, none of these thrusts reaches the seafloor. Wide-Angle Data [14] Data quality was very good on the almost all OBSs and land stations. Sample record sections are shown in Figures 3, 4, 5, 6, 7, 8, and The wide-angle data are classified into three patterns. The record sections from the Shikoku Basin to the Nankai Trough (OBSs 14 1) simply show four main phases; they are refracted arrivals from oceanic layer 2 (), layer 3 (), uppermost mantle (), and reflection from the Moho (). As shown in Figures 3 and 4, for example, and are identified at offsets between 6 and 14 km and 14 and 35 km, respectively. is observed as high-amplitude later arrivals at offsets of km. can be recognized at offsets over 14 km. The refracted and reflected arrivals from the sediments (Psed,, and Rsed) are identified as later phases. The record sections of OBSs 9 2 are characterized by first arrivals with a lower apparent seismic velocities of 3 5 km/s ( and shown in Figures 5, 6, and 7). These phases are observed up to km offset and interpreted as refracted arrivals from the sedimentary layers beneath the continental slope. Several later arrivals are also recognized; they are considered to be reflection arrivals from the subducting oceanic crust (P and P) and Moho (). Record sections of OBS 1 and all land stations show first arrivals with apparent velocities of 5 6 km/s ( and Plc), widely observed up to 1 km offset (Figures 8 and 9). These arrivals are interpreted as refractions from

7 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE EPM 2-7 Figure Observed and synthetic seismograms at OBS 4: (a) observed seismograms and (b) synthetic seismograms. island arc crustal blocks. Intracrustal reflections (PcP) and are observed as prominent later phases on the land stations. can be clearly observed along the entire profile. Modeling Procedure [15] The P wave velocities within the shallow sedimentary layers were estimated by applying the t-p method [Diebold and Stoffa, 1981; Shinohara et al., 1994] to the OBS data. The t-p inversion gave us a one-dimensional velocity structure beneath each of the OBSs. Reflectors in the shallow sedimentary layers and the top of the oceanic crust between OBS 7 and OBS 14 were picked from the time-migrated MCS section, as refracted and wideangle reflected arrivals were not always observed by OBSs. We assumed that the lateral velocity variation of the sedimentary layers could be expressed by interpolating these inversion results between the OBSs. The geometry of the sedimentary layers down to the top of the oceanic crust were calculated by using estimated velocities as a result of the t-p inversion. We then put this structure of the sedimentary layers into the initial model of two-dimensional velocity structure for the following forward modeling as a priori information. In particular, the geometry and velocities between OBS 9 and OBS 14 were kept fixed throughout the following modeling procedure because they were derived from only the t-p inversion and the MCS data. [16] The velocity model, from which refracted and wide-angle reflected arrivals were recorded by OBSs and land stations, was determined mainly by a trial-and-error approach. The observed first and clear later arrivals were picked from all wide-angle seismic

8 EPM 2-8 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE Figure Observed and synthetic seismograms at OBS 3: (a) observed seismograms and (b) synthetic seismograms. sections. Uncertainties between.4 and.26 s, depending on signal-to-noise ratios were assigned to the picked travel times by comparing the energy before and after the pick. First, we calculated theoretical travel time curves and synthetic seismograms and modified the model parameters iteratively until all these picked observed data could be explained consistently. We used a twodimensional ray tracing technique [Zelt and Smith, 1992] for travel time calculations. We then applied two-dimensional travel time inversion [Zelt and Smith, 1992] to our crustal model in order to refine and evaluate our model objectively. This inversion is performed in a layer stripping procedure in which the parameters of deeper layers are successively determined while keeping the parameters of the shallower layers fixed. [17] The reliability of the final model could be examined in terms of its resolution and uncertainty. This inversion procedure allows calculation of the resolution of each parameter specifying the final crustal model [Zelt and Smith, 1992]. The resolution values range between zero and one and indicate the relative number of rays that sample each model parameter. Model parameters with a resolution >.5 are generally considered to be well resolved [Zelt and Smith, 1992]. In order to demonstrate the reliability and uniqueness of the crustal model we use two parameters given by Zelt and Smith [1992]; they are root-meansquare of misfit between calculated and observed travel times (T RMS ) and misfit normalized by uncertainty of the observed travel times (c 2 ). [18] The uncertainties of the model parameter (interface and velocity values) can be estimated following Zelt and Smith [1992]. The value of a model parameter is perturbed until the travel time fit of the perturbed model clearly deviates from that of the unperturbed model. The maximum perturbation that allows a comparable fit is a measure of the absolute uncertainty [Zelt and Smith, 1992]. The process is time consuming, and we therefore only tested a few representative model parameters. These tests show that model parameters with such high resolution cannot be perturbed by more than, for example,.1 km/s for the velocity of oceanic layer 3, or.9 km for the Moho depth, in order to fit the calculated travel time within the estimated uncertainty.

9 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE EPM 2-9 from our model is presented in Figure 11, allowing assessment of the quality of our model. Resolution of the interface and velocity calculated from the travel time inversion (Figure 12) and ray diagrams (Figure 13) for the calculated arrivals are also indicated to discuss the reliability of our crustal model. Figure Observed and synthetic seismograms at station KWR on land: (a) observed seismograms and (b) synthetic seismograms. [19] The crustal model finally was confirmed by using synthetic seismograms [Zelt and Ellis, 1988]. We pay special attention to the following two characteristics which are not included during the travel time analysis: the existing range of each refracted phase and the location of the critical point of the Moho reflected phase. These characteristics control the velocity gradient in each layer and the velocity contrast at the Moho, respectively. However, in this study we only made a visual comparison of the observed and synthetic data; we have not made a quantitative analysis of the amplitude information. Crustal Model [] Figure 1 shows the P wave velocity structure model along the wide-angle OBS profile across the rupture zone of the 1944 Tonankai earthquake. In this section we discuss our model, focusing on the most important structure from the Shikoku Basin toward the Japanese island arc. The crustal model shows subducting oceanic crust and a thick sedimentary wedge with seismic velocities <7 km/s beneath the continental slope. Moreover, the thick sedimentary wedge is bounded landward by a crustal block with the velocities over 5 km/s. Beneath this crustal block, a second crustal block with the velocities over 6 km/s is found. They are called island arc upper and lower crust, respectively, following previous studies [e.g., Ikami et al., 1982; Kodaira et al., ]. A comparison of the observed travel times and calculated arrivals Oceanic Crust From Shikoku Basin to Nankai Trough [21] The crustal structure from the Shikoku Basin to the Nankai Trough ( 7 km composed of four layers (Figure 1)). The two upper layers are sediments, and their velocities are 7 8 and 8 km/s, respectively. Beneath the sedimentary layers, two layers with velocities of 5 6 and 7 9 km/s are obtained except between OBS 1 and OBS These two layers are 5 km and 5 km thick, respectively. On the basis of their velocities and thicknesses these layers are interpreted as oceanic layers 2 and 3 of the Philippine Sea plate, respectively [e.g., Purdy and Ewing, 1986; White et al., 1992; Nakanishi et al., 1998; Kodaira et al., ]. The crustal model shows slightly lower velocities of top of oceanic layer 2 ( km/s) and layer 3 (5 km/s) between OBS 1 and OBS 9 at the deformation front as defined from our MCS data (Figure 2). The velocity of the uppermost mantle is 9 km/s. [22] The estimated velocities and geometry of the sedimentary layers well explain the refracted (Psed) and reflected (Rsed) arrivals observed by several OBSs (Figures 3, Figure 11a, 11b, 11c, 11e, and 11f). Observed first arrivals labeled as and in Figures 3a and 11b are well explained as refracted phases from oceanic layers 2 and 3, respectively. The slightly lower velocities of top of oceanic layers 2 and 3 between OBS 1 and OBS 9 are indicated by the variation of the apparent velocities of and on OBSs 9 1 For example, the apparent velocity of the refraction arrivals from oceanic layer 2 from the northwest recorded on OBS 1 (Figure 4) is certainly slower than that recorded on OBS 13 (Figure 3a). Such a high-velocity contrast at the Moho interface in our model explains the large amplitudes of the reflection from this interface () and also seen in synthetic seismograms (e.g., Figures 3 and 4). The velocity of the uppermost mantle is constrained by clearly identified (e.g., Figures 3a and 4). The increasing amplitudes of observed in the range between 17 and km are well reproduced by the synthetic seismograms in Figure 3b. From this result the top of the uppermost mantle with a velocity of km/s beneath the continental slope and the island arc can also be constrained by phases observed on OBSs 9 1 [23] The resolution of the crustal model between OBS 14 and OBS 9 is high (.5), and the ray coverage is good down to the uppermost mantle except for the southwestern end of the profile. However, poor ray coverage in oceanic layer 2 gave several parts of low resolution (.5) in this layer (Figures 12, 13a, and 13c 13e). Sedimentary Wedge and Island Arc Crust [24] The crustal structure beneath the continental slope (7 24 km) shows a thick sedimentary wedge, island arc crustal blocks, and subducting oceanic crust. The sedimentary wedge consists of three layers with velocities of 6 5, 9, and 1 7 km/s. The wedge reaches a maximum thickness of 7 km at 115 km. The wedge becomes thinner and nearly disappears around the coastline. A crustal block with a velocities of 9 km/s (except around OBS 2, where this crustal block shows a slower velocity down to 8 km/s) is found at 115 km (Figure 1). The velocity of the top of this crustal block increases to 8 km/s at the northwestern end of the profile. Beneath the island arc upper crust, another crustal block with a velocity of 7 km/s is found at 145 km (Figure 1). The velocity of the lower crustal block is comparable to that of the island arc lower crust found by Ikami et al. [1982] and Kodaira et al. [] and may increase toward the Kii Peninsula. [25] The velocities and geometry of the sedimentary wedge are constrained by refraction phases, Psed1,, and,

10 EPM 2-1 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE Figure Observed and synthetic seismograms at station INN on land: (a) observed seismograms and (b) synthetic seismograms. observed by OBSs 1 9 (Figures 11f 11n). For example, the synthetic seismograms shown in Figure 5b match the existing range of refraction phases, and from the wedge very well. The wide-ranging refraction phase,, which is observed on OBSs 1 6 and on four land stations, leads us to infer the presence of materials with a small velocity gradient (e.g., Figures 6, 8, 9, and 11j 11r). The phase is interpreted as a refracted phase from the island arc upper crust. The same kind of phases is also observed off Shikoku Island [Kodaira et al., ] and around the eastern Nankai Trough [Nakanishi et al., 1]. The calculated travel time curves for the refracted arrivals from the island arc lower crust fit the Plc phase observed on OBSs 1 and 2 and on four land stations (Figures 8, 9, and 11o 11r). Further evidence for inner crustal boundary separating the island arc lower crust and the island arc upper crust can be derived from a prominent reflected phase (PcP) originating at this boundary (Figures 8, 9, and 11o 11r). [26] As shown in Figure 12, the structure of the sedimentary wedge is well resolved with resolution.5 except for the velocity at the bottom of both ends of the second layer. This is also directly recognized from the ray coverage (Figure 13a). The refracted phases from the second layer show a sufficient coverage for the upper part of the layer. There are, however, no seismic rays refracted into the lower part of the layer. The geometry of the top of the island arc upper crust is obtained with high resolution (.75) except for the onshore area, where no refracted rays in the sediments would be observed due to lacking shot on land (Figure 13a). In the process of modeling, the geometry of the top of the island arc upper crust beneath the land area was estimated by assuming sediments with velocities of 2 km/s overlying the island arc upper crust. As shown in Figure 12a, velocity values in the island arc upper crust were well resolved because of a good ray coverage of refracted waves (Figures 13b). The geometry of the island arc upper crust/lower crust interface was well resolved except for the northwestern edge of the profile. However, the high-resolution values for the velocities (.5) (Figure 12) and good ray coverage are restricted to the seaward edge of the island arc lower crust (Figure 12). Subducting Oceanic Crust [27] The subducting oceanic layers 2 and 3 deepen landward and subsequently underlie the sedimentary wedge and the island arc upper and lower crust. Velocities of oceanic layers 2 and 3 are 5 6 and 7 9 km/s, respectively, except around the deformation front as mentioned above. These velocities of the subducting oceanic layer 2 from 11 to 24 km are assumed to be same as those determined beneath the continental slope (OBS 7) in the process of modeling. The velocities of oceanic layer 3 from 16 to 24 km are also assumed to be continuous from those obtained around 16 km. The thicknesses of the subducting oceanic layers 2 and 3 are 2 and 5 km, respectively. The subducting oceanic crust can be traced down to 3 km. The subduction angle becomes steeper toward the Japanese island arc. The subducting

11 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE EPM 2-11 Depth (km) SSE Oceanic layer 2 Oceanic layer deformation front Uppermost mantle continental slope NNW 15 C 35 C 45 C interseismic locked zone transition zone 1944 Tonankai rupture area Sedimentary wedge 2 coast line Island Arc upper crust Island Arc lower crust 8 8 KWR KHR AIZ INN 35 4 Contour interval :.2 km/s Distance (km) Figure P wave velocity model. Positions of the OBSs are shown by numbered solid circles. Velocities are in kilometers per second. The isovelocity contour interval is.2 km/s. Coseismic rupture zone of the 1944 Tonankai earthquake estimated by tsunami and geodetic data [Ando, 1975] is shown at the top of the figure. Interseismic locked and transition zone estimated by geothermal data and critical boundary temperature (15, 35, and 45 )[Hyndman et al., 1995] are also indicated. The shaded area is sedimentary wedge described in text. Moho can be traced down to 3 km depth at the distance of km. The uppermost mantle velocity beneath the island arc crust ( km/s) is almost same as that beneath the area from the Shikoku Basin to the continental slope. [28] The structure of the subducting oceanic crust was determined from the travel time and amplitude information of refraction arrivals of, P, and phases (Figures 6, 7, and 11) interpreted as refracted arrivals from oceanic layer 3, reflected arrivals from the boundary between oceanic layers 2 and 3, and the reflection from the Moho interface, respectively. The good fit between observed and calculated travel times supports our model. We could not identify refraction arrivals from subducting oceanic layer 2 beneath the island arc upper and lower crust because subducting oceanic layer 2 becomes a low-velocity zone with no refracted rays (Figure 13c). Reflected phases from the top of subducting oceanic layer 2 (P) were recognized on OBSs 3 and 7 (Figures 5, 7, 11h, and 11l). The existence of subducting oceanic layer 2 is determined from P, and below the sedimentary wedge (Figure 13c). Moreover, the intercept time of and travel time jump between and the refracted phase from the layer overlying the subducting oceanic crust (e.g., Figures 6 and 7) indicate the existence of the subducting oceanic layer 2 as a low-velocity zone from 11 to 24 km. As is common for the lowvelocity zones, the velocity and thickness cannot be determined uniquely. The way it is modeled here is a compromise between the most simple models with a laterally constant velocity and thickness which is consistent with those of the subducting oceanic layer 2 from to 11 km. Structural variation of the low-velocity zone is estimated within average velocities from to 5 km/s and thicknesses from 7 to 2 km to explain apparent velocities of P. Fluctuation of subduction angle, as a consequence of this structural variation, is <1. The geometry of the subducting Moho beneath the island arc crust explains the phases well (Figures 7, 8, and 9). As shown in Figures 3 9 and 11, are clearly identified all over the profile. From this phase an uppermost mantle velocity of km/s is obtained without significant lateral variation. [29] The structure of the subducting oceanic crust beneath the island arc crust is not well resolved (resolution values.25, Figure 12). However, the coverage of refracted and reflected rays in subducting oceanic layer 3 between and 17 km is very good (Figures 13d and 13e), and the resolution values of geometry and velocity are high (.75) (Figure 12). The geometry of the subducting Moho has a high-resolution value (.75) as well (Figure 12b). However, the resolution values northwest of 17 km are low in subducting oceanic layer 3 because no refracted rays penetrate such deep parts (Figures 12 and 13d). As shown in Figures 12 and 13, a good coverage of refracted rays resulted in high-resolution values of the top of the velocities in the uppermost mantle. Subduction Angle [3] To discuss the crustal structure around the downdip limit of the coseismic rupture zone of the 1944 Tonankai earthquake (Figure 1) it is important to investigate the precise structure of the subducting oceanic crust beneath the island arc crust. However, details of the structure of the subducting oceanic crust beneath the island arc crust are not resolved well because of the low-velocity nature of oceanic layer 2 as well as the poor ray coverage. We thus try to determine the subduction angle of the subducting oceanic crust beneath the island arc crust. An analyzing process that determined the subduction angle beneath the island arc crust is as follows: [31] We evaluate the reliability of the subduction angle beneath the island arc by using T RMS and c 2 as described in section In this analysis the subduction angle was changed from 4 to 16 (Figure 14a). The velocities in the subducting oceanic crust were not varied. The T RMS and c 2 values are calculated for the refracted arrivals from the island arc lower crust, the subducting oceanic crust and the uppermost mantle, and the reflected arrivals from the top of oceanic layer 2, oceanic layer 3, and the Moho interface, as observed at OBSs 1 4 and all four land stations (Figure 14b). As shown in Figure 14b, both the T RMS and c 2 values are minimum at a subduction angle of 11. We thus conclude that the subduction angle becomes steeper toward the Japanese island arc. The subducting oceanic crust dips gently at an angle of 4 from to 6 km (7 13 km in Figure 1),

12 2-12 EPM NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE SSE SSE k P Psed l OBS 3 P PcP PsP P 17 P OBS 12 Psed 1 1 Psed m OBS 2.. Plc OBS n Rsed OBS Psed OBS o PcP KWR Plc OBS f PcP P Rsed OBS 4 OBS 13 Psed Psed e Psed1 Psed1. 1 d c b OBS 5 OBS 14 Rsed NNW NNW j a

13 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE EPM 2-13 (g). P OBS 8 (p). Plc PcP KHR (h). P P OBS 7 (q). Plc PcP AIZ (i). P P OBS 6 (r). Plc PcP INN Figure 1 Observed arrivals and calculated travel time curves of all OBSs and the land stations. Curves represent calculated travel times. Vertical bars represent uncertainty of the observed arrivals. Reduction velocity is 8 km/s. Rsed, reflection arrivals from the bottom of sediments; Psed1,, and, refraction arrivals from the sedimentary layer on the oceanic crust or the sedimentary wedge;, refraction arrivals from oceanic layer 2;, refraction arrivals from the oceanic layer 3;, reflection arrivals from the Moho interface;, refraction arrivals from the uppermost mantle; P, reflection arrivals from the top of the oceanic crust; P, reflection arrivals from the boundary between oceanic layer 2 and oceanic layer 3;, refraction arrivals from the island arc upper crust; Plc, refraction arrivals from the island arc lower crust; PcP, reflection arrivals from the boundary between the island arc upper crust and the island arc lower crust. while subduction angles of 11 are reached 6 km northwest of the deformation front at 13 km. Discussions Geological Interpretation of Sedimentary Wedge and Island Arc Crustal Blocks [32] It has been recognized that the outer margin of southwest Japan preserves a long history of subduction of oceanic plate [e.g., Taira, 1985; Okuda, 1984; Taira et al., 1989, 1992]. On land, the pre-neogene tectonic divisions of southwest Japan south of the Median Tectonic Line (MTL) can be classified into two terranes (Figure 1): The first, the Jurassic subduction complex is composed of the Chichibu Belt and their metamorphic counterparts, the Sanbagawa Belt. The second, the Shimanto Belt, which is separated from the Jurassic subduction complex by the Butsuzo Tectonic Line (BTL), is classified into the northern (mainly Cretaceous) and southern (mainly Paleogene) Shimanto Belt [e.g., Taira, 1985; Okuda, 1984]. The offshore geology of southwest Japan has been investigated from previous MCS profiles [e.g., Okuda, 1984; Aoki et al., 1982]. These studies clearly illustrated that examples of ongoing sediment accretion processes. According to Okuda [1984] the offshore geology off the Kii Peninsula is stratigraphically described as follows: The acoustic basement beneath the Kumano Basin is correlated to the Tertiary Shimanto Belt. The lower to middle Miocene Tanabe and Kumano formations lie above this. Moreover, the major filling material of the Kumano Basin overlies those formations. A synthesis of offshore geology based on these MCS studies is presented by Taira et al. [1992]. They show a simplified distribution of geotectonic units in the Japanese Islands and offshore. According to speculation of Taira et al. [1992] the southern margin of the southern Shimanto Belt is probably located around OBS 2 of this study (Figure 1). [33] As shown in Figure 15, we interpret the sedimentary wedge and the island arc upper crust of our crustal model by comparing with the distribution of geotectonic unit suggested by Taira et al. [1989]. Differences in geological structure cannot be directly resolved by seismic velocity structure. [34] Taking a stratigraphic interpretation of the offshore geology [Okuda, 1984] into consideration, we interpret the sedimentary wedge as a Neogene-Quaternary accretionary prism that formed after the opening of the Japan Sea. The crustal block with velocities of 9 km/s beneath the island arc had been called island arc upper crust in previous studies [e.g., Kodaira et al., 1996]. Judging from the onshore geology [e.g., Taira, 1985; Okuda, 1984], the island arc upper crust can be related to the Jurassic accretionary prism and the Cretaceous Shimanto Belt at least to the

14 EPM 2-14 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE Figure 1 Resolution values for the crustal model shown in Figure 2, calculated from the travel time inversion: (a) resolution values of the velocity nodes and (b) resolution values of the interface nodes. coast line, as well as the Paleogene Shimanto Belt beneath the upper continental slope [Taira et al., 1992]. Although the layer that is continuous to the island arc upper crust in our model extends southeastward beneath the sedimentary wedge, velocity of this layer is clearly much higher than that of the overlying sedimentary wedge. The older sedimentary rocks are usually considered to have a higher velocity due to compaction and alteration. The layer continuous to the island arc upper crust beneath the forearc basin can thus be related to the Tertiary Shimanto Belt. Moreover, velocities of the island upper crust become higher northwestward around the northwestern end of the profile. This result might be related to the existence of the MTL. Spatial Relationship Between the Crustal Model and the Coseismic Rupture Zone of the 1944 Tonankai Earthquake [35] In this section we discuss a spatial relationship between our crustal model and the coseismic rupture zone of the 1944 Tonankai earthquake defined by the tsunami and geodetic data [e.g., Aida, 1981; Ando, 1975; Inouchi and Sato, 1975]. Although there are several fault models for the 1944 Tonankai earthquake, most downdip limits of the rupture zone are estimated at almost the same location with small uncertainty. The downdip limit of the coseismic rupture zone of the 1944 Tonankai earthquake is located at km of our profile as shown in Figure Although the forearc Moho was not determined around the Kii Peninsula, it was found at 32 km depth in the Tokai district 5 km inland from the coast [Ikami, 1978]. The forearc Moho at a depth of 32 km is also found 7 km landward from the northern end of the profile MO14 off the Shikoku Island [Ikami et al., 1982]. We therefore assumed that the forearc Moho is at 32 km depth northwest of landward end of the profile of this study (KR986) by interpolating these previous results of crustal studies on land. This result is consistent with the depth distribution of the forearc mantle beneath Japan as shown by Zhao and Hasegawa [1993] from a detailed inversion of hypocenters. As a result, the intersection of the slab with the forearc mantle is estimated at the distance of 25 km of our OBS profile (Figure 15). Thus we conclude that the downdip limit of the coseismic rupture zone of the 1944 large event does not extend to the forearc mantle. Furthermore, the depth of the upper surface of the subducting oceanic crust at the downdip limit of the rupture zone is 23 km. This depth is agree with the estimated depth ( 25 km) of the downdip limit of the interseismic locked zone defined from thermal modeling by Hyndman et al. [1995]. However, on the basis of the Wadati- Benioff seismicity, Nakamura et al. [1997] proposed that the downdip limit of brittle fracture in the slab extends to 27 km of our profile over MTL (Figure 15). Comparing the seismicity of Nakamura et al. [1997] with our crustal model, microearthquakes related to the slab become deeper northwestward from the MTL at 23 km of our profile. This result is consistent with hypothesis of thermal constraints on interseismic locked zone

15 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE EPM 2-15 (a) SSE NNW (d) SSE NNW DEPTH (km) Sedimentary layers DEPTH (km) Oceanic layer 3 (b) (e) DEPTH (km) Island arc upper crust DEPTH (km) Moho interface (c) (f) DEPTH (km) Oceanic layer 2 DEPTH (km) Uppermost mantle Figure 1 Ray diagrams for the calculated arrivals: (a) refraction arrivals from the sedimentary layers on the oceanic crust and the sedimentary wedge (Psed1,, and ), (b) refraction arrivals from the island arc upper crust (), (c) refraction arrivals from oceanic layer 2 (), (d) refraction arrivals from oceanic layer 3 (), (e) reflection arrivals from the Moho interface (), and (f) refraction arrivals from the uppermost mantle (). [Hyndman et al., 1995]. The estimated temperature beneath the MTL reaches 45 C, the highest temperature that coseismic rupture for crustal rocks may extend. Therefore earthquakes may not be generated at the plate boundary (the top of the slab) but instead are generated within the subducting oceanic crust and uppermost mantle. However, more precise determinations of source mechanisms and hypocentral distribution considering two- or three-dimensional crustal structure are needed for further discussion. [36] A geological and seismological boundary is located around OBS In terms of geological and crustal structure, there are three characteristic structures around OBS 6: (1) the outer ridge is located just seaward of OBS 6, (2) the decollement zone which is clearly imaged by the MCS section beneath the continental slope becomes indistinct landward from OBS 6 (Figure 2), and (3) the boundary between the Neogene-Quaternary accretionary prism and the island arc upper crust is located beneath OBS From the seismological point of view both the updip limits of the coseismic rupture zone [Ando, 1975] and the interseismic locked zone [Hyndman et al., 1995] do not extend seaward from OBS 6, which means that both updip limits are not shallower than the boundary between the Neogene-Quaternary accretionary prism and the island arc upper crust, although the updip limit of the rupture zone is only poorly constrained by tsunami data and crustal deformation recorded on land. Therefore the behavior of the sedimentary rocks is considered to change around OBS 6 from stable sliding to stick slip. We can also conjecture that the appearance of the island arc upper crust makes the formation of a decollement zone difficult landward from OBS Although several thrusts including a possible OST fault are recognized in the MCS section (Figure 2), they do not reach to the seafloor. It is therefore reasonable to suppose that the coseismic rupture process of the 1944 large event is not responsible for the formation of these thrusts. Thus we conclude that coseismic rupture of the 1944 large event does not extend seaward of OBS Comparison With Crustal Model Across the Coseismic Rupture Zone of the 1946 Nankai Earthquake [37] Little is known about the factors controlling the seismogenic zone and rupture patterns. It is therefore important to find out differences and similarities in deep crustal structures of the coseismic rupture zones between the 1944 and 1946 earthquakes. The crustal model across the coseismic rupture zone of the 1946 Nankai earthquake is presented by Kodaira et al. [] as described in section 1 (Figure 16b). Both crustal models show the subduction structure with the subducting oceanic crust situated beneath the thick sedimentary wedge and an island arc upper and lower crust. There are two notable differences between two models: the subduction angle beneath the island arc crust and the width of the subducting oceanic crust/neogene-quaternary accretionary prism contact zone. First, the subduction angle of our model across the 1944 rupture zone (11 ) is 4 steeper than that of the crustal model across the 1946 rupture zone (7 ), although the initial angles of subduction in both of crustal models are similar (5 ). These different structures in the subduction angle and the depth of the upper surface of the subducting oceanic crust are consistent with the characteristic geometry of the seismic plane derived by hypocenter distribution beneath Japan [Nakamura et al., 1997]. Second, the width of the

16 EPM 2-16 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE (a) 5 SSE NNW Depth (km) ~ Distance (km) (b) TRMS 2 T RMS (s) χ χ Angle deviation (degree).8 Figure 1 Analysis of the reliability and uniqueness of the subduction angle: (a) subduction angle changed during the analysis and (b) T RMS and c 2 for the refraction arrivals from the island arc lower crust, the subducting oceanic crust, and the uppermost mantle and reflection arrivals from the top of the oceanic crust and the boundary between oceanic layers 2 and 3, and the Moho interface estimated as a function of the subduction angle. Solid lines in the model indicate the final model as in Figure subducting oceanic crust/neogene-quaternary accretionary prism contact zone of our model across the 1944 rupture zone (5 km) is shorter than that of the model across the 1946 rupture zone (7 km). Moreover, Takahashi et al. [1999] show that the width of the subducting oceanic crust/neogene-quaternary accretionary prism contact zone in the presumed unruptured zone along profile KR981, located just west of the 1946 rupture zone, is only km. As shown in Figure 1, a wide-angle profile KR981 was also performed using a similar configuration of OBS and stations on land. It seems that the width of the subducting oceanic crust/neogene-quaternary accretionary prism contact zone is proportional to the size of the coseismic rupture zone. [38] The spatial relationship between the crustal model and the rupture zone reveals a common characteristic of both rupture zones; the depth of upper surface of the subducting oceanic crust at the downdip limit of the coseismic rupture zone is 23 km. This uniform depth limit agrees with the hypothesis that the downdip limit of the Nankai Trough seismogenic zone is constrained by temperature as proposed by Hyndman et al. [1995]. The value of 23 km depth is the first precise result derived from our OBS study for the Nankai Trough. The downdip limit of the Nankai seismogenic zone may be constrained by the depth of the subducting oceanic crust (23 km) as well as the 35 C temperature limit which corresponds to thermally activated stable-sliding behavior for crustal rocks (with a transition to 45 C [Hyndman et al., 1995]). Furthermore, it is supposed that the downdip limit of the Nankai Trough seismogenic zone is not limited by the forearc mantle, as the downdip limits of both coseismic rupture zones are still reached above the forearc Moho. [39] Although the location of the updip limit is uncertain, there is a difference in the width of the rupture zone along the Nankai Trough. In the coseismic rupture zone of the 1946 event the updip limit is located in the Neogene-Quaternary accretionary prism, while the updip limit of the rupture zone of the 1944 event only extends below the island arc upper crust. The different positions of the updip limit between the coseismic rupture zones of the 1944 and 1946 events are also supported by MCS sections [Park et al., ]. There are two major differences in the appearance of the decollement zone and the OST faults: (1) In the rupture zone of the 1946 event the decollement zone is clearly traced 3 km landward from the deformation front [Park et al., ], while in the rupture

17 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE EPM 2-17 Quaternary Tertiary Cretaceous Jurrassic Neogene Paleogene continental slope BTL MTL deformation front Neogene-Quaternary Accretionary prism?? coast line Outer Ridge SS NS Chi Sa Depth (km) SSE Oceanic layer 2 Oceanic layer Contour interval :.2 km/s Uppermost mantle 15 C 35 C 45 C interseismic locked zone transition zone 1944 Tonankai rupture area Sedimentary wedge Distance (km) 2 KWR KHR AIZ INN 1 3 Island Arc upper crust Island Arc lower crust NNW Forearc Moho 5-8 Forearc mantle Figure 1 Crustal section consisting of the model in Figure 1, hypocenter distributions beneath the Kii Peninsula [Nakamura et al., 1997], and an estimated result from land refraction surveys [Ikami, 1978; Ikami et al., 1982]. Coseismic rupture zone of the 1944 Tonankai earthquake [Ando, 1975] and interseismic locked and transition zone [Hyndman et al., 1995] are shown at the top of the figure. The model is geologically interpreted on the basis of the onshore geology as in Figure 1 and offshore geology shown by Okuda [1977] and Taira et al. [1992]. zone of the 1944 event it is difficult to follow the decollement zone longer than 15 km landward from the deformation front, and (2) a splay fault system consisting of some sigmoid forearc OST faults, which cut the ocean floor completely, is related to the updip limit of the coseismic rupture zone of the 1946 Nankai earthquake representing possible thrusts [Park et al., ], while only smallscale buried thrusts faults are found along the profile across the rupture zone of the 1944 event. These differences recognized in the MCS sections also mean that the rupture process did not extend to the Neogene-Quaternary accretionary prism during the 1944 earthquake in contrast to the 1946 event. These differences in crustal structure between the coseismic rupture zones of the 1944 and 1946 events may reflect different rupture processes in the ductile A and the brittle C blocks, as suggested by Kanamori [1972] and Ando [1975]. Good correspondence between locations of the forearc basins and the presumed coseismic rupture zones [Ando, 1975] is proposed by Awata and Sugiyama [1989]. Moreover, there are submarine canyons at boundaries between areas B and C and areas C and D. In particular, at the landward extension of this boundary the Jurassic and Cretaceous accretionary belt, which is continuously well ordered from the Shikoku Island, is complicatedly disturbed. Furthermore, there is a Kinan Seamount chain along the seaward extension of boundary between areas A and B. Structural variation of the slab beneath the boundaries between areas C and D, areas B and C, and areas A and B were pointed out by Nakanishi et al. [1], Mochizuki et al. [1998], and Sato et al. [1998], respectively. These structural variations can be considered to be correlated with the existence of submarine canyons or seamount chain. A vertical offset between areas B and C and a remarkable difference in dip angle between areas C and D are shown by Yamazaki et al. [1989] on the basis of observation of microearthquakes using the land seismic network. These structural variations or the existence of topographic anomalies beneath the boundaries of areas A D might be one of causes to stop and change the coseismic rupture process when it propagates to adjacent areas. Comparison of Subduction Seismogenic Zones With Great Interplate Earthquakes [4] In this section we compare our crustal model of the Nankai Trough with other representative subduction zones where coseismic rupture zones of historic large thrust earthquakes had been determined. In particular, following discussions focus on the characteristics of the crustal structure of the updip and downdip limit of coseismic rupture zone. As well as along the Nankai Trough, it is also well known that there are remarkable regularities in the space-time distribution of great earthquakes occurred along the Kurile Trench, beneath which the Pacific plate is subducting beneath the Hokkaido Island of Japan [e.g., Ustu, 1972]. The crustal model across the coseismic rupture zone of the 1952 Tokachi-oki earthquake (M s =3)[Japan Meteorogical Agency,

18 EPM 2-18 NAKANISHI ET AL.: STRUCTURE OF THE NANKAI SEISMOGENIC ZONE a Tonankai block This study SSE deformation front oceanic crust / accretionary prism contact zone coseismic rupture zone NNW sea water Moho Uppermost Mantle Neogene-Quaternary accretionary prism Island arc crust subducting oceanic crust (11 ) ~23km Forearc Moho > 3 km Forearc Mantle b Nankai block after Kodaira et al. [] coseismic rupture zone sea water Moho Uppermost Mantle Neogene-Quaternary accretionary prism Island arc crust subducting oceanic crust (7 ) ~23km Forearc Moho > 3 km Figure 1 Schematic crustal model indicating differences between the different rupture zone: (a) area C ruptured by the 1944 Tonankai earthquake and (b) area A ruptured by the 1946 Nankai earthquake [after Kodaira et al., ]. Dashed lines indicate the width of the oceanic crust and Neogene-Quaternary accretionary prism contact zone. Shaded dashed lines indicate the width of the coseismic rupture zone of the earthquake. 1974] is deduced from wide-angle seismic survey using OBSs [Iwasaki et al., 1989]. This crustal model is characterized by subducting oceanic crust beneath the crustal block corresponding to the island arc upper crust. However, there is no well-developed sedimentary wedge like the accretionary prism of the Nankai Trough. Iwasaki et al. [1989] considered such a difference beneath the continental slope as a reflection of different subduction mechanisms at the margins. Comparing this crustal model with coseismic rupture zone of 1952 earthquake estimated from tsunami data [Aida, 1978], the updip limit is not shallower than the island arc upper crust. This agrees with the crustal structure beneath the updip limit of the 1944 Tonankai earthquake derived from this study. The existence of the forearc Moho was not clarified because the seismic profile of Iwasaki et al. [1989] did not run across the entire coseismic rupture zone. However, assuming a constant subduction angle, the landward extension of top of the subducting oceanic crust can be estimated to reach 25 km depth at the downdip limit, which is almost consistent with the uniform downdip limit depth of the Nankai Trough obtained from our study. Two crustal models deduced by wide-angle seismic profiles off Valparaiso, Chile, are presented by Fleuh et al. [1998]; one is selected to cross the coseismic rupture zone of the 1985 earthquake (M w = ) [Comte et al., 1986], and another is located just north of the rupture zone. These crustal models show the subducting Nazca plate beneath South America. Crustal models of this area have two similarities with those of the Nankai Trough. First, the crustal model has a thick sedimentary wedge beneath the continental, slope. Second, there are differences between crustal models of outside and inside of the coseismic rupture zone, e.g., the different subduction angle and the existence of the low-velocity zone in the sedimentary wedge. Fleuh et al. [1998] interpreted these differences to be responsible for the ridge subduction which is recognized between these two profiles. The coseismic rupture zones of the 1985 event estimated from aftershock distribution [Choy and Dewey, 1988], tsunami modeling [Nakamura, 1992], and dislocation modeling of coseismic deformation [Barrientos, 1995] are all consistent with each other. The updip limit of this rupture zone is not shallower than the area of high-velocity rock (<5 km/s) in the upper plate that is interpreted as part of the continental framework [Fleuh et al., 1998]. Although Oleskevich et al. [1999] proposed that thermal constraint to the downdip limit is provided by the intersection of the subducting oceanic plate with the continental forearc Moho at 4 km depth determined from hypocenter distribution, we believe that the crustal model of Fleuh et al.