wide-angle reflection analysis

Tectonophysics 475 (2009) 438–453

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Tectonophysics j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / t e c t o

Collision and subduction structure of the Izu–Bonin arc, central Japan, revealed by refraction/wide-angle reflection analysis Ryuta Arai a,⁎, Takaya Iwasaki a, Hiroshi Sato a, Susumu Abe b, Naoshi Hirata a a b

Earthquake Research Institute, The University of Tokyo, 1-1-1 Yayoi, Bunkyo-ku, Tokyo 113-0032, Japan JGI Incorporated, 1-5-21 Otsuka, Bunkyo-ku, Tokyo 112-0012, Japan

a r t i c l e

i n f o

Article history: Received 20 October 2008 Received in revised form 1 May 2009 Accepted 24 May 2009 Available online 3 June 2009 Keywords: Refraction/wide-angle reflection Crustal structure Izu–Bonin arc Arc–arc collision Subduction of Philippine Sea plate Delamination

a b s t r a c t Since the middle Miocene, the Izu–Bonin arc (IBA) has been colliding from the south with the Honshu arc in central Japan associated with subduction of the Philippine Sea plate (PSP). This process is responsible for forming a complex crustal structure called the Izu Collision Zone (ICZ). To obtain direct evidence of the deep structure dominated by collision and subduction, an intensive seismic experiment using explosive and vibroseis sources was conducted in the eastern part of the ICZ. CMP reflection and refraction/wide-angle reflection data were acquired on a 130-km-long seismic line crossing the Kanto Mountains (KM), the Tonoki– Aikawa Tectonic Line (TATL) and the Tanzawa Mountains (TM) from north to south. The TATL is considered to be a collision boundary separating the Honshu arc (KM) from the IBA (TM). The structural model constructed by refraction tomography and forward ray tracing shows remarkable lateral velocity variation across the TATL and some clear reflectors in the deep crust. A north dipping reflector beneath the KM was interpreted to be the deeper extension of the TATL. From the geometry of reflectors, we interpret the Tanzawa block is delaminated from the subducting slab due to the collision to form a wedge-like body thrusting between the upper and lower crust of Honshu. The TM itself shows a velocity structure almost consistent with the upper part of the crust of the Izu intraoceanic arc, south of our study area. The relocated hypocenter distribution using our velocity model shows high seismic activity concentrated around the collision boundary, which is in a marked contrast of low seismicity within the Tanzawa block. These features of seismicity are strongly dominated by the collision of the IBA with the Honshu arc. © 2009 Elsevier B.V. All rights reserved.

1. Introduction Collision and subduction processes are key factors which control crustal deformation and evolution at convergent plate margins. Recent geophysical expeditions including seismic surveys have revealed crustal structures dominated by continent–continent collision such as the Alps (TRANSALP Working Group, 2002; Bleibinhaus et al., 2006; Brückl et al., 2007) and subduction of oceanic plates such as the Andes (ANCORP Working Group, 2003). For arc–arc collision zone associated with plate subduction, however, there is little evidence that directly demonstrates complex structure of collision style to a deeper crustal level. The Japanese Islands, which consist of several island arcs located along the eastern margin of the Asian Continent, provide good research fields for studies of crustal evolution and deformation processes of island arcs ongoing in subduction zones (Fig. 1(a)). They exist in a complex tectonic environment involving subduction, accretion, back-arc spreading and arc–arc collision. The main part of the Japanese Islands is geologically divided into two arcs of NE and SW Japan. The NE Japan arc consists of western Hokkaido and northern Honshu, beneath which the ⁎ Corresponding author. E-mail address: [email protected] (R. Arai). 0040-1951/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2009.05.023

Pacific plate is subducted. The SW Japan arc characterized by subduction of the Philippine Sea plate (PSP) is composed of western Honshu, Shikoku and Kyushu. These arcs were located at the eastern edge of the Asian Continent during the Mesozoic to early Miocene time, and rotated to their present locations due to the Miocene back-arc opening of the Sea of Japan (Otofuji et al., 1985). There are two arc–arc collision zones in the Japanese Islands. In Hidaka collision zone, Hokkaido, the Kuril forearc has been collided against the western part of the Hokkaido (the NE Japan arc) since the Miocene time (Fig. 1(a)). This area was well investigated by the reflection and refraction/wide-angle reflection method (e.g. Ito, 2000, 2002; Iwasaki et al., 2004). These studies show the delamination structure of the Kuril forearc, where the upper 21 km crust of the Kuril forearc is obducted to form a metamorphic belt, while its lower crust is subducted under the NE Japan arc. Another arc–arc collision is ongoing in the northernmost part of the Izu–Bonin arc (IBA), located south of Honshu (Fig. 1(a)). This arc itself is known as a good example of an island arc system in an intraoceanic convergent margin and well investigated by seismic surveys (Richards and Lithgow-Bertellone, 1996; Stern et al., 2003; Kodaira et al., 2007). Because of its northeastward motion associated with the subduction of the PSP and the back-arc opening of the Sea of Japan mentioned above, the IBA

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Fig. 1. (a) Location map of the Japanese Islands and the surrounding area. A red rectangle shows the area of the Izu Collision Zone (ICZ, see also Table 1). (b) Survey map of the study area. Blue, pink, black, and light blue lines denote the refraction/wide-angle reflection line for large energy shots. The blue, pink and light blue lines also represent the dense reflection lines (the Kiryu Line, the Sagamiko Line and the Matsuda Line, respectively). Dynamite shots (SP1, 6–13) are indicated by red stars, and vibroseis multiple shots in the northern part by yellow stars (SP2–5) and in the southern part by blue stars (SP14, 15). An active fault and major tectonic boundaries are indicated by red lines (Fukaya Fault, Tonoki–Aikawa Tectonic Line (TATL) and Kozu–Matsuda Fault (KMF). The Kanto Mountains (KM), the Tanzawa Mountains (TM), the Misaka Mountains and the Koma Mountains are represented by shaded zones. The KM are composed of Paleozoic to Mesozoic accretionary prisms such as the Sambagawa Metamorphic Belt (SMB), the Chichibu Belt (CB), and the Shimanto Belt (SHB). The Tanzawa group (TZ) consists of Miocene volcanic rocks. The Kanto Plain (KP) is covered with the Holocene sediments. Abbreviations of geological units and their meanings are listed in Table 1.

started to collide with the Honshu Island in the middle Miocene. This collision process caused wide intra-arc deformations resulting in the formation of the Izu Collision Zone (ICZ) (Kano et al., 1990; Yamakita and Otoh, 2002). Some fragments of the IBA such as the Tanzawa Mountains (TM) and the Izu peninsula (IP) were accreted onto the Honshu crust, and several active faults represented by the Tonoki– Aikawa Tectonic Line (TATL) and the Kozu–Matsuda Fault (KMF), were formed in the ICZ. In Table 1, abbreviations of geological units and their meanings are listed. Due to the complex tectonic processes, the deeper structure of the ICZ was left unclarified. In 2003, an integrated seismic experiment was carried out across the eastern flank of the ICZ (Sato et al., 2005; Fig.1(b)). Sato et al. (2005), based on the CMP reflection processing from this experiment, firstly succeeded in mapping the north dipping PSP, and the delamination between the upper Tanzawa block and the PSP. Besides very dense CMP data, this experiment yielded a large amount of travel time data of first arrivals and later phases (wide-angle reflections), which are not fully used in the usual CMP reflection processing. In this paper, an integrated analysis was applied to those data sets to provide direct evidence of collision and subduction of the IBA in terms of seismic

wave velocity structure. Our analysis consists of seismic tomography, forward modeling by ray-tracing and amplitude analysis of later phases. The obtained velocity structure model by our analysis is complementary to the crustal image from the CMP reflection method because our data sets provide more or less independent information on near-vertical

Table 1 List of abbreviations used in the text. ICZ IBA PSP TM IP KP TATL KMF TZ SMB CB SHB

Izu Collision Zone Izu–Bonin arc Philippine Sea plate Tanzawa Mountains Izu Peninsula Kanto Plain Tonoki–Aikawa Tectonic Line Kozu–Matsuda Fault Tanzawa group Sambagawa Metamorphic Belt Chichibu Belt Shimanto Belt

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Table 2 Data acquisition parameters. Seismic line

Refraction

Dense reflection line Kiryu

Length (km) Source type Charge (kg)/sweep No. of sweeps No. of shots Shot interval No. of shots in tomography Receiver No. of receivers

138 Dynamite 300 (SP1,13) 200 (SP6–12) 9

Sagamiko

Matsuda

53 Four vibrators Four vibrators 6–30 Hz (or 35 Hz), 24-s sweep length

55 Four vibrators

16 Four vibrators

75–715 6

20 347 100 m 42

50 77 500 m 36

200 94 100 m 17

480

965

221

9 5 10-Hz geophones at 50-m receiver interval 2518 1278 (SP2–5) 621 (SP14,15)

Sampling rate

reflection data. Combining the result of this study with the image of the reflection method in Sato et al. (2005) ensures more reliable geophysical/geological interpretations on the obtained structure model. Seismicity also provides useful insights into crustal structure and collision process. In the study area, a number of permanent seismic stations have been deployed to study local seismic activity by several organizations (Earthquake Research Institute (ERI), National Research Institute for Earth Science and Disaster Prevention (NIED), Hot Springs Research Institute of Kanagawa Prefecture (HSRI) and Japan Meteorological Agency (JMA)).

4 ms

JMA is providing arrival times of local events observed at these stations. From these arrival time data, we also carried out hypocenter relocation in order to discuss the relationship between seismic activity and crustal heterogeneity arising from collision and subduction. 2. Tectonic setting The ICZ was formed through the collision process between the IBA and the central part of Honshu (Fig. 1(b) and Table 1). This area has

Fig. 2. Examples of seismic records from dynamite shots with calculated travel times. A 3- to 15-Hz band-pass filter is applied. Each trace is normalized by its maximum amplitude. Clear later phases are indicated by black arrows. (a) Seismic record of SP1. 2-s automatic gain control (AGC) is applied. (b) Seismic record of SP12. No AGC is applied.

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Fig. 3. Initial and final velocity model in refraction tomography. The origin of the horizontal axis is taken at SP1 at the northern end of the seismic line. Locations of the large energy shots are denoted by black stars, and vibroseis shots of dense reflection lines by red circles. (a) Initial model in which the low velocity zone corresponding to the KP is roughly taken into account. (b) Final model after 4 iterations. The surface geological structures are also presented. It is noted that lateral velocity variation corresponds well to the geological units. The Fukaya Fault is interpreted to be dipping southward. The TATL, which is the collision boundary between the Honshu arc and the Izu–Bonin arc (IBA), is interpreted to be dipping northward along the low velocity body of the Tanzawa block. Gray color means no ray coverage.

been intensively studied over the past few decades, especially using geological approaches (Sugimura, 1972; Amano, 1991; Matsuda, 1989; Aoike, 1999; Taira et al., 1998). Petrological and geological data strongly indicate that the Koma Mountains, the Misaka Mountains, the TM and the IP were accreted onto the Honshu crust, forming several tectonic boundaries in and around the ICZ. For example, the TATL, located at the northern flank of the TM, is considered to be a collision boundary between the Honshu arc and the IBA (Amano, 1991; Matsuda, 1989). The TM consist of early to middle Miocene volcanic rocks called the Tanzawa group (TZ) and Neogene intrusive rocks represented by tonalites. The tonalites widely exposed at the land surface around the center of the TM are considered to form the middle crust of the IBA (Kodaira et al., 2007). Some petrological studies suggested that the intrusion of the tonalites occurred just before the

collision of the Tanzawa block (Amano, 1991). Thus these rocks are of great importance in understanding the collision process of the IBA. Sugimura (1972) firstly proposed that the KMF, which connects the Sagami trough and the Suruga trough developed on both sides of the ICZ, was the northern boundary of the Izu–Bonin oceanic crust. Since then, numerous geological studies on the origin of crustal block forming the ICZ have been conducted (Matsuda, 1989; Aoike, 1999). Matsuda (1978) suggested that the present KMF was a former trough (Ashigara trough) filled with non-volcaniclastic rocks of the Ashigara group formed after the collision of the IP. This idea was also supported by Huchon and Kitazato (1984) and Ito (1985). It was also proposed that the TM was another segment of the IBA crust which collided with Honshu before the IP. The clastic rocks in the northern and eastern flank of the TM filled the trough between the Kanto Mountains (KM)

Fig. 4. Results of the checkerboard resolution test. (a) The size of velocity anomaly is 10 km × 2 km. (b) 6 km × 1.2 km.

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Fig. 5. The velocity model in the shallow part by forward ray tracing. Geometries and velocities of the individual layers, and reflective boundaries including shallow reflector R0 (a red line) are shown. Shaded areas indicate where there is no ray coverage. The surface geological structures are also presented.

and the TM (Niitsuma and Matsuda, 1985; Niitsuma and Akiba, 1985). Aoike (1999) indicated that the collision boundary was migrated from the TATL to the KMF during 15–4 Ma from depositional ages of troughfill sediments along the segment boundary. From mass balance calculation using growth rates of the IBA and the NE Japan arc, Taira et al. (1998) presented a cross-section across the ICZ in NNW-SSE direction, showing the wedge-like structures of the IBA crusts. However,

the deep part of these blocks and clear images of tectonic boundaries directly dominated by the collision process are left enigmatic. The KM are mainly composed of Paleozoic to Mesozoic accretionary prisms represented by the Sambagawa Metamorphic Belt (SMB), the Chichibu Belt (CB) and the Shimanto Belt (SHB) from north to south (Taira et al., 1989; Suzuki, 2002; Kano et al., 1990; Fig. 1(b)). Their trends are changed from almost E–W to NNE-SSW direction

Fig. 6. Travel time plots of first arrivals. Observed and calculated travel times are shown by filled and open circles, respectively. Large open circles indicate the locations of shots. (a), (b) and (c) show first arrivals of dynamite shots in the northern end of the profile (SP1), in the KM (SP6) and just north of the TATL (SP10), respectively.

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Fig. 7. Seismic records for vibroseis shots on a reflection line (Kiryu) with calculated travel time curves, and their ray diagrams. A 3–20 Hz band-pass filter and 1-s AGC are applied. Reduction velocity is 6 km/s. Each trace is normalized by its maximum amplitude. (a) Seismic record for SP326. (b) Ray diagram for SP326. (c) Seismic record for SP347. (d) Ray diagram for SP347.

around the ICZ. The SMB consists of the high-pressure metamorphic rocks of Cretaceous age. The CB is mainly composed of Mesozoic strata of sandstone and mudstone, and also includes chert and limestone. The SHB is similar to the CB in its composition, as characterized by Cretaceous to early Miocene strata. On the northern side of the KM, the Kanto Plain (KP) is covered with Holocene coarse fragment and sand. The Fukaya Fault is located at the northwestern part of the KP (Fig. 1(b)).

The PSP is forming subduction zones on both sides of the ICZ (Fig. 1(b)). On the eastern side, it is subducted from the Sagami trough, and large interplate earthquakes such as the 1703 and 1923 great Kanto earthquakes occurred during the last few hundred years (Matsu'ura et al., 1980; Ishibashi, 2004). In spite of the social as well as geological/geophysical importance, the subduction geometry itself was not clarified in this region. Some papers report the image of the PSP using hypocenter distribution and focal

Fig. 8. The final velocity model. Tectonic interpretations are also presented. Seismic velocity is indicated by color contour and numerals. Red solid lines denote clear reflectors in the crust, which are referred to as R0–R6 as described in the text. Shaded areas indicate where there is no ray coverage.

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mechanisms (Ishida, 1992; Noguchi, 1998), tomography analysis (Matsubara et al., 2005), and S to P converted waves (Tsumura et al., 1993). The results of these analyses, however, do not have adequate resolution to allow us to discuss the fine-scale structure in and around the ICZ, particularly the relationship between the IBA crust and the PSP. 3. Data acquisition In 2003, an intensive seismic reflection/refraction experiment was undertaken as a part of Special Project for Earthquake Disaster Mitigation in Urban Areas (Sato et al., 2005). A 130-km profile line was set in the eastern flanks of the KM and the TM, crossing the TATL, and the KMF in almost N–S direction (Fig. 1(b)). On this profile, a refraction/wide-angle reflection experiment data were collected using 15 large energy shots [9 explosive shots (SP1, SP6–13) with 200–300kg charge and 6 vibroseis multiple shots (SP2–5, SP14, SP15)]. A total of 2518 geophones were deployed with an average spacing of 50 m. In addition, reflection experiments with densely spaced vibroseis shots (100–500 m spacing) were conducted to image active faults on the three segments of the profile line (Kiryu, Sagamiko and Matsuda Lines). Hereinafter these reflection lines are referred to as “dense reflection lines”. On the Kiryu Line (a blue line in Fig. 1(b)), located on the northern half of the profile, 347 shots and 480 receivers were deployed to obtain a high-resolution image of the KP. The Sagamiko Line in the middle of the profile (a pink line in Fig. 1(b)) was designed to elucidate lateral variation across the TATL using 77 shots and 965 receivers. In the southern end of the profile, 94 shots and 221 receivers were deployed to image fine structures of the KMF (the Matsuda Line, a light blue line in Fig. 1(b)). In these reflection lines, four vibroseis trucks were used as a seismic source. Data acquisition parameters are shown in Table 2. 4. Characteristics of data and modeling procedure Data examples of seismic record from explosive shots are shown in Fig. 2. Fig. 2(a) shows a record section of explosive shot SP1 at the northern end of the profile line. Clear first arrivals can be traced up to an offset of about 70 km. First arrivals observed near the shot points have an apparent velocity of about 3 km/s, representing the low velocity of shallow sediments in the KP, while arrivals with a velocity of 6 km/s in the middle of the profile correspond to a refracted wave from the crystalline upper crust. First arrivals show amplitude decrease beyond 70-km distance. This phenomenon is also due to a thick sedimentary package in KP. We can also recognize several later phases with high amplitudes 0.2–2 s behind the first arrivals, which are interpreted to be wide-angle reflections from the deeper part of the crust. Fig. 2(b) shows a record section from SP12, 106 km south from SP1. As described in Section 5.2, later phases in this record section are interpreted to be reflections from the upper surface of the PSP and the deeper reflector within it. The crustal model was constructed from the following two steps. First, we conducted the refraction tomography analysis (Zelt and Barton, 1998) using only first arrival data. This analysis focuses on the structure of the shallower part of the crust, particularly lateral structural variations around active faults. In this analysis, first arrival data from 14 large energy shots (except for SP14) and 85 vibroseis shots (42 of the Kiryu Line, 36 of the Sagamiko Line and 17 of the Matsuda Line) were used. The latter data were expected to delineate the finer structure of the shallowest part of the crust with reliable resolution. For picking the first arrivals at far offsets, band-pass filters

of 1–15 and 6–30 Hz were applied for the dynamite and vibroseis shots, respectively. Next, forward modeling by ray tracing was performed (Červený and Pšenčík, 1983; Iwasaki, 1988). In this step, later phases which were not used in the tomography analysis were incorporated. In the starting model of this computation, the result from the refraction tomography was included to model the layer geometry and the velocity changes within the individual layer. After obtaining a model explaining the travel time data, synthetic seismograms were calculated based on the asymptotic ray theory, from which the velocity gradient within the individual layers and velocity contrasts across the individual layer boundary were determined. These travel time and amplitude analyses were repeated until the model satisfactorily explained both of the observed travel times and amplitudes. 5. Crustal model 5.1. Refraction tomography We used seismic refraction tomography for determining the velocity structure of the shallow crust. The software package used is FAST (Zelt and Barton, 1998). In this step, we defined a model space of 130 km × 5 km composed of 0.5 km × 0.5 km cells. An origin of modeling was set on a position of SP1 at the northern edge of the profile. The velocities of the cells were determined from 34,426 first arrival times using a damped least squares method. Fig. 3(a) shows an initial model where a basin structure of the KP, which is filled with a low velocity material is taken into account. Fig. 3(b) shows a final model obtained after 4 iterations. Through this inversion, the χ2 residual of travel times was reduced from 44.32 to 4.82. In order to examine the initial model dependence of final solution, we also carried out the inversion for other flat-layered starting models, and confirmed that general features of the images in Fig. 3(b) remain almost unchanged. Checker board resolution tests demonstrate that the model resolution in the upper 3 km crust is capable to resolve ±5% velocity anomalies sized larger than 10 km × 2 km (Fig. 4(a)). A finer size such as 6 km × 1.2 km was successfully recovered in the upper 1–2 km crust below the dense reflection lines due to their dense ray coverage (Fig. 4(b)). Reconstruction tests were also conducted to analyze the resolving capabilities. In this test, synthetic data were constructed from the final model and inverted by the same procedures as in the real data analysis. By this test, almost the same velocity model as the final model was satisfactorily recovered. Prominent large-scale lateral structural variations in Fig. 3(b) almost correspond to the geological structures along the profile (see also Fig. 1(b)). Beneath the northern part of the profile, a low velocity body (b4.5 km/s), which represents the Holocene sediment of the KP, reaches to a depth of about 3 km. We found the lateral velocity variation across the south dipping boundary in the KP, probably representing the Fukaya Fault. The KM consisting of Paleozoic to Mesozoic accretionary prisms (SMB, CB and SHB, see Fig. 1(b) and Table 1) are characterized by higher seismic velocities (N4.0 km/s). In contrast, the TM including volcanic rocks are relatively low velocity (~3.5 km/s). The surface trace of the TATL is situated on the northern edge of this low velocity area. In our model in Fig. 3(b), the TATL is expressed as a boundary between the high velocity KM and the low velocity TM, extending at least down to a depth of 2 km. In the southernmost part of the profile, we notice the uppermost velocity of the southern side of the KMF, the Ashigara group, is slightly smaller than in the north (see also Fig. 1(b) and Table 1).

Fig. 9 Seismic records for dynamite shots with calculated travel time curves, and their ray diagrams. (a) Seismic record for SP1. A 3- to 15-Hz band-pass filter and 2-s automatic gain control (AGC) are applied. Each trace is normalized by its maximum amplitude. Clear reflections are referred to as R0–R6 (see also Fig. 8). (b) Ray diagram for SP1. Black arrows indicate the range of the reflection point. (c) Seismic record for SP8. A 3- to 15-Hz band-pass filter and 2-s AGC are applied. (d) Ray diagram for SP8. (e) Seismic record for SP9. A 3- to 15-Hz band-pass filter and 2-s AGC are applied. (f) Ray diagram for SP9. (g) Seismic record for SP12. A 3- to 15-Hz band-pass filter is applied. (h) Ray diagram for SP12.

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Fig. 9 (continued).

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Fig. 10. Comparison between observed and synthetic seismograms for SP1. Each trace is normalized by its maximum amplitude. (a) Observed record section for SP1. A 1–15 Hz bandpass filter is applied. No AGC is applied. (b) Synthetic seismograms for SP1 calculated from the final model in Fig. 8. (c) Synthetic seismograms for SP1 calculated from a model in which velocity below R2 is assumed to be 6.7 km/s. It is noted that calculated amplitudes from R2 are much lower than the observed.

5.2. Forward modeling by ray tracing Next, we determined the velocity structure of the entire crust by forward ray tracing. Fig. 5 shows velocity structure down to a depth of 11 km. General velocity variation in the uppermost crust is consistent with Fig. 3(b), particularly bowl-shaped low velocity part corresponding to the KP, which consists of 2.0–3.2 km/s and 3.8–4.9 km/s layers. Layer boundaries are modeled both from the first arrivals (Fig. 6) and later phases observed in the dense reflection lines (Fig. 7). Fig. 6 shows first arrival time plots for some large energy shots (SP1, SP6, and SP10). Most of the first arrivals are satisfactorily explained within an error of 0.1 s. It is noted that the apparent velocities are notably different across SP10, which is located just north of the TATL (Fig. 6(c)). The resultant velocities of the KM in the northern side of the TATL are ranging from 3.6 to 5.8 km/s in the uppermost three

layers, clearly higher than those of the TM (3.3–5.7 km/s) in the southern side (Fig. 5). In this analysis, travel times of reflections observed in the dense reflection lines (Fig. 7) were also used to determine layer geometries. This forward modeling was repeated until a model satisfactorily explaining both of the observed first arrivals and reflection phases was obtained. Calculated travel times (solid curves in Fig. 7) agree well with the observed phases. The deeper structure was constructed from the wide-angle reflections observed in the record sections of large energy shots as shown in Fig. 2. The layer geometry of the deeper crust was determined from the travel times of the later phases. Then, synthetic seismograms were computed to check whether the model explains the observed amplitude characteristics including the relative amplitudes of reflected and refracted waves on the same trace and their amplitude variation with offset.

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Fig. 11. Comparison between observed and synthetic seismograms for SP9. Each trace is normalized by its maximum amplitude. (a) Observed record section for SP9. A 1–15 Hz bandpass filter is applied. No AGC is applied. (b) Synthetic seismograms for SP9 calculated from the final model in Fig. 8. (c) Synthetic seismograms for SP9 calculated from a model in which velocity below R2 is assumed to be 6.7 km/s. It is noted that calculated amplitudes from R2 and R3 are much lower than the observed.

The final model in the whole crust from the forward analysis is shown in Fig. 8. Red lines represent clear reflectors in the deep crust, which are named as R0–R6. For dynamite shots of SP1, SP8, SP9 and SP12, record sections with calculated travel times of deep reflections and their ray diagrams are presented in Fig. 9. For SP1, SP9 and SP12, synthetic seismograms are presented in Figs. 10–12.

As shown in Fig. 9(a), the model well explains the travel times of reflections from R1, R2, R4 and R5 for SP1. Amplitude behaviors of first arrivals and prominent reflection from R5 in the offset of 40–70 km are in good agreement with the observed record section (Fig. 10(b)). It is also noted that the amplitude decrease in first arrivals farther than about 70 km is mainly attributed to the bowl-shaped structure and the

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Fig. 12. Comparison between observed and synthetic seismograms for SP12. Each trace is normalized by its maximum amplitude. (a) Observed record section for SP12. A 1–15 Hz band-pass filter is applied. No AGC is applied. (b). Synthetic seismograms for SP12 calculated from the final model in Fig. 8.

high velocity gradient beneath the KP (see also Section 4 and Fig. 2 (a)). Prominent reflections from R5 require a large velocity contrast of about 0.7 km/s for the corresponding boundary at 26-km depth. If we take the velocity contrast to be a small value of 0.2 km/s or a large value of 1.2 km/s, for example, the observed amplitude behaviors are not explained well. Therefore, the estimation error for the velocity contrast is less than 0.5 km/s. In the middle part of the profile, travel times of reflections from R0, R1, R2 and R3 are also well explained in the record section from SP8 and SP9 (Fig. 9(c) and (e)). The travel time data of R2 in the record sections of SP1, SP8 and SP9 strongly indicate the corresponding boundary dips northward. For reflected waves from R3, also seen in Fig. 9(c) and (e), the corresponding boundary is located at a depth of 16–20 km with a slight southward dip. It is noticed that the interface of R2 is merged to R3 to form a wedge-like structure under the KM. By comparing observed and calculated seismograms for SP9 shown in Fig. 11(a) and (b), R0–R3 are estimated to have velocity contrasts of 0.2–0.4 km/s, that is 6.1–6.2 km/s beneath R0, 6.5–6.6 km/s beneath R1, 6.2 km/s beneath R2 and 6.8 km/s beneath R3, respectively (Fig. 8).

In our model the wedge-like body bounded by R2 and R3 has a low velocity of 6.2–6.4 km/s. In order to investigate the validity of this modeling, we calculated synthetic seismograms for SP9 using a model where the velocity of the wedge-like body is assumed to be 6.7 km/s (Fig. 11(c)). Although reflected waves from R2 and R3 are clearly seen in the observed records (Fig. 11(a)), this model can not reproduce these reflections (Fig. 11(c)). Similarly, the reflections from R2 of SP1 are difficult to be explained by this model (Fig. 10(c)). Namely, the value of 6.7 km/s is too small to explain the reflections from R2. If we take lager value of 6.9–7.0 km/s instead of 6.7 km/s, then the velocity below R3 must be taken 7.3–7.4 km/s to explain the amplitudes of reflections from R3. This value, however, seems to be too high for the middle/lower crust of the Japanese Island (Iwasaki et al., 2001). These results strongly indicate that the low velocity TM forms the wedgelike body beneath the KM. The most prominent feature in the southern part of the profile is the northward dipping structure presenting the subduction of the PSP (Fig. 8). In Fig. 12(b), the record section and calculated travel times are presented for SP12. The northward descending geometry of R4

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Fig. 13. (a) Map of the study area including 1241 epicenters of relocated earthquakes indicated by blue circles. White squares denote seismic stations deployed by JMA, NEID, The University of Tokyo and Hot Spring Research Institute of Kanagawa Prefecture. Relocated epicenters are shown by blue dots. Red circles indicate the grid point used in the hypocenter relocation. The black line indicates the 2003 refraction/wide-angle reflection profile. (b). Relocated hypocenters after 7 iterations are indicated by orange circles. Shaded areas indicate where there is no ray coverage. (c). Hypocenter distribution determined by JMA. 1714 hypocenters from the JMA catalog in January 2002 to April 2008 are indicated by orange circles. Shaded areas indicate where there is no ray coverage.

represents the upper surface of the PSP. From this interpretation, reflector R6 is situated within the PSP. Amplitude calculation in Fig.12(b) indicates a large velocity contrast of about 0.8 km/s at R4. The resultant velocity below this boundary becomes 7.1–7.2 km/s. If we take a lower (6.7 km/s) or a higher (7.7 km/s) velocity beneath R4, calculated synthetic seismograms for SP12 can not explain the observation. Therefore, the velocity error for this part is estimated to be less than 0.5 km/s. R6 is estimated to have a velocity contrast of about 0.4 km/s. 6. Seismicity In order to delineate relationships between the seismic activity and complicated tectonic structure around the ICZ, we carried out hypocenter relocations using the present velocity model. The routine hypocenter determination by JMA (Japan Meteorological Agency) is based on a 1-dimensional (horizontally flat layered) velocity structure model. As shown in the present study, however, the velocity structure in and around the ICZ is quite heterogeneous both in the vertical and horizontal directions, which yields significant errors in the routine hypocenter determination.

6.1. Data and method Fig. 13(a) shows a map of the 54 seismic stations used for the present study. These stations are maintained by the several organizations including ERI, NIED, HSRI and JMA (see also Section 1). This figure also shows the coordinate axes for the hypocenter relocation. We selected 1714 earthquakes in the JMA catalog, occurring in the vicinity of our profile line (a region of 35°19′N to 36°10′N, 139°05′E to 139°20′E and down to a depth of 30 km) from January 2002 to April 2008. Their vertical distribution beneath the profile is presented in Fig. 13(c). From these earthquakes, we used travel times of 1241 events, of which P wave arrival times were observed at more than 10 stations and S wave arrival times at more than 5 stations. In the present study, numbers of absolute arrival times for the P and S wave are 17,775 and 17,858, while, those of double-difference of P and S wave data reach to 85,755 and 85,820 respectively. In the computation, the double-difference tomography code (Zhang and Thurber, 2003) was used, in which only hypocentral parameters were determined with structural parameters fixed. Furthermore, the velocity structure was assumed to be 2.5

Fig. 14. Migrated depth section by reflection method (after Sato et al., 2005).

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Fig. 15. Schematic illustration of collision of the Tanzawa block. Prior to the middle Miocene, the Tanzawa block was separated from the Honshu arc (Stage 1). In stage 2, the Tanzawa block collided with the Honshu crust and thrust into a weak zone characterized by the present active seismicity. The TATL was formed in the boundary. The Tanzawa block could not subduct beneath the Honshu crust and started to be delaminated from the IBA lower crust in Stage 3. In Stage 4, the Izu block collided with the Tanzawa block, forming the KMF in the boundary.

dimensional. Namely, the structural variation along Y-axis (N–S direction) and Z-axis (depth direction) is based on the refraction model (Figs. 8 and 13(a)), but fixed in X axis direction (E–W direction). The Vp/Vs value is set as 1.73 for the entire part of the model. The grid size is fixed at 50 and 2 km for X (E–W) and Y (N–S) direction, respectively. The grid point in Z (depth) direction is located at 0, 1, 2, 3, 4, 5, 6, 7, 8, 9, 10, 12, 14, 16, 18, 20, 23, 26, 30, 35 and 40 km. 6.2. Comparison with the refraction model In the process of hypocenter determination, the root mean square (RMS) of travel time residuals was reduced from 0.200 s to 0.144 s after 7 iterations. The earthquake distribution obtained shows some interesting features (Fig. 13(a) and (b)). First, we see significant difference between the northern and southern side of the TATL. Within the Honshu crust, most hypocenters are distributed within the upper 5–20 km crust with a velocity of 6.0–6.8 km/s. Compared with the routine relocation by JMA (Fig. 13(c)), relocated hypocenters are more or less concentrated along the reflector of R1 at a depth of about 12 km. Beneath the southern part of our profile line, on the other hand, a number of earthquakes occurred in a depth range of 10–25 km, which presumably corresponds to the lower crust of the IBA. It is also noticed that there are almost no earthquakes within the Tanzawa block, but some seismic activities are found in the peripherals of the block, particularly, along the deeper extension of the TATL (R2) and around the reflector R3. It is likely that they are associated with the collision of the IBA and the Honshu arc. 7. Discussion 7.1. Tectonic interpretations based on the velocity model and the reflection image

consistent with each other, particularly for the basin structure of the KP. In the reflection image, the uppermost part of the TATL forms a steeply dipping event. In the velocity structure model, on the other hand, the TATL is recognized as a boundary between the high velocity KM and the low velocity TM (see also Fig. 3(b)). This velocity difference becomes much clearer in the deeper part due to the existence of the prominent reflector R2, which is interpreted to be a deeper extension of the TATL. We see a wedge-like body bounded by R2 (TATL) and the other clear reflector R3 (Fig. 8). According to the amplitude analysis of the present study, the velocity of this body is 6.2–6.4 km/s, 0.2–0.4 km/s lower than the surrounding Honshu arc crust. Based on these features, the wedge-like body is interpreted to be a fragment of the IBA separated from the PSP in the course of the collision (Fig. 8). It is also noticed that the wide-angle reflection data confirmed the northward extension of the upper boundary of the PSP (Fig. 9(a) and (b)). Furthermore, our model shows two other clear reflectors (Fig. 8). One is at a depth of 26 km beneath the KM, and the other within the PSP, almost parallel to R4. These are our important results which are not unclear in the reflection image (Fig. 14). The reflector R5 has a large velocity contrast (0.7 km/s) at a depth of 26 km, and the resulting velocity beneath R5 is estimated to be 7.5 km/s (Fig. 8). Although this value may contain an error (b0.5 km/s), it is reasonable to consider that it represents the velocity of the uppermost mantle beneath the Honshu arc. Actually, Matsubara et al. (2005) showed the three-dimensional P and S wave velocity structure beneath the Kanto district by seismic tomography. In their result, a low velocity zone with high Vp/Vs ratios near the upper boundary of the subducting PSP at depths of about 30–40 km is clearly imaged. They interpret this low velocity zone as the mantle wedge consisting of 20% partially serpentinized peridotites, which have P wave velocity of 7.3 km/s and S wave velocity of 3.9 km/s at 1 GPa (Christensen, 1972). 7.2. Comparison with the crustal structure of the IBA from other studies

Seismic image from the reflection data by Sato et al. (2005) is given in Fig. 14, which should be compared with the refraction/wide-angle reflection model in Fig. 8. The uppermost parts of these results are

Detailed velocity model of the IBA south of our study area, namely representing the structure before the collision, was presented by

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Kodaira et al. (2007) from refraction/wide-angle reflection and multichannel seismic reflection data acquired on a 550-km-long marine profile. They showed that the IBA crust mainly consists of five layers, corresponding to the upper crust with P wave velocities of 1.8– 5.8 km/s, the upper and lower part of the middle crust with 6.0– 6.5 km/s and 6.5–6.8 km/s, and the upper and lower part of the lower crust with 6.8–7.2 km/s and 7.2–7.6 km/s. By comparing their result with our model, it is found that the Tanzawa block including the wedge-like body with a velocity of 2.5–6.4 km/s corresponds to the upper crust and upper part of the middle crust. On the other hand, the velocity of 7.1–7.2 km/s between R4 and R6, which clearly forms a part of the subducted PSP, corresponds well to the lower crust of the IBA (Fig. 8). The deeper part of the PSP beneath R6 was determined to be 7.6 km/s with an error of 0.2–0.3 km/s. Kodaira et al. (2007) proposed 7.8–8.0 km/s as the velocity of the uppermost mantle. Therefore, there are two possibilities for the interpretation for R6, namely a boundary between the upper and lower parts of the lower crust or the Moho boundary. From relocated hypocenter distributions in Fig. 13(b), it is apparent that most earthquakes are concentrated between R4 and R6. 7.3. Collision style of the IBA Taira et al. (1998) proposed the cross-section of the ICZ from the mass balance calculation using the growth rates (magmatic addition rates) of the northern IBA since 45 Ma and the Northeast Japan arc since 15 Ma. This section shows the imbrication structure of the accreted IBA crusts and the subduction of the PSP, and there are some similarities to our model such as the north dipping configuration of the accreted IBA crusts. However, the geometry of the lower collision boundary may be different. A wedge-like structure of the reflectors of R2 and R3 in our model suggests that the middle crust of the IBA (the Tanzawa block) is delaminated from the lower crust (PSP), and thrusts between the upper and lower crust of the Honshu arc. It is noted that similar detachments occurring on the boundaries between the upper and lower crusts are also found in continent–continent collisions such as the Alps (Pfiffner, 1992; Schmid et al., 1996) and the Pyrenees (Teixell, 1998). Based on intensive seismic profilings, these papers depicted the complex crustal structures including thrust wedges and detachments at mid-crustal level dominated by continent–continent collision. Their results strongly indicated weak interfaces between the upper and lower crusts. Thus, these phenomena seem to be common in the course of the collision, regardless of its collision type. The Hidaka collision zone in Hokkaido, Northeast Japan (Fig. 1), also has a complex crustal structure associated with an arc–arc collision of the Kuril arc and the Northeast Japan arc. Recent seismic experiments showed a delamination in the Kuril arc crust in the collision area, analogous to the ICZ (Iwasaki et al., 2004). However, there is a significant difference in terms of forms of crustal structure. In the ICZ, the upper part of the IBA crust is accreted onto the Honshu crust. On the other hand, an obduction of the middle/lower crusts of the Kuril arc occurs in the Hidaka collision zone. 8. Conclusions We modeled the crustal structure in the eastern part of the ICZ using refraction/wide-angle reflection data. The obtained model indicates the complex crustal structure including wedge-like body of the Tanzawa block thrusting into the Honshu arc and the subducting PSP corresponding to the lower crust of the IBA. The relocated hypocenter distribution suggests that the seismic activities in the ICZ are also dominated by the collision and subduction processes of the IBA. We propose a new model of the collision process between the Honshu crust and the IBA crust, as shown in Fig. 15. This model is characterized by thrusting and delamination of the Tanzawa block. Stage 1 depicts the presumable structure before the collision. Since there is a gap between the collision age of the Tanzawa block and the

Izu block (Amano, 1991), it is reasonable that these two blocks were originally separated. In stage 2, the Tanzawa block started to collide and thrust into the Honshu arc crust. The structural model obtained by this study shows that this wedge body thrusts into the weak zone between the upper and lower crust in the Honshu (R3 in Fig. 8). There exist seismic activities around this zone (Fig. 13), which suggest that these activities may be induced from the collision of the wedge body. The depositional age of the trough-fill sediments indicates that the collision of the Tanzawa block took place about 6–5 Ma (Amano, 1991). In stage 3, the upper crust and upper part of the middle crust of the IBA (Tanzawa block) were delaminated from the subducting lower crust (PSP). In stage 4, the Izu block collided with the Tanzawa block about 1 Ma (Amano, 1991), forming the KMF in the boundary. Acknowledgements Instruments used in the seismic experiment were provided by the seismic survey company JGI, Incorporated and ERI, The University of Tokyo. We wish to express our sincere thanks to the JGI seismic crew for data acquisition. We are grateful to Japan Metrological Agency for providing the hypocenter and travel time data. We are also grateful to Dr. Heidrun Kopp and two anonymous reviewers for their useful comments and corrections for our manuscript. We used the GMT software (Wessel and Smith, 1998) to draw the figures. This study was supported by the Special Project for Earthquake Disaster Mitigation in Urban Areas from the Ministry of Education, Culture, Sports, Science, and Technology of Japan.

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