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Tectonophysics 388 (2004) 7 – 20
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Configuration of subducting Philippine Sea plate and
crustal structure in the central Japan region
Takashi Iidakaa,*, Tetsuya Takedaa, Eiji Kurashimoa, Tomonori Kawamuraa,
Yoshiyuki Kanedab, Takaya Iwasakia
a
Earthquake Research Institute, University of Tokyo, Yayoi 1-1-1-, Bunkyo, Tokyo 113-0032, Japan
b
Japan Marine Science and Technology Center, Natsushima 2-15, Yokosuka, Japan
Received 30 September 2003; received in revised form 3 February 2004; accepted 13 June 2004
Available online 2 September 2004
Abstract
A seismic experiment with six explosive sources and 391 seismic stations was conducted in August 2001 in the central Japan
region. The crustal velocity structure for the central part of Japan and configuration of the subducting Philippine Sea plate were
revealed. A large lateral variation of the thickness of the sedimentary layer was observed, and the P-wave velocity values below
the sedimentary layer obtained were 5.3–5.8 km/s. P-wave velocity values for the lower part of upper crust and lower crust were
estimated to be 6.0–6.4 and 6.6–6.8 km/s, respectively. The reflected wave from the upper boundary of the subducting
Philippine Sea plate was observed on the record sections of several shots. The configuration of the subducting Philippine Sea
slab was revealed for depths of 20–35 km. The dip angle of the Philippine Sea plate was estimated to be 268 for a depth range of
about 20–26 km. Below this depth, the upper boundary of the subducting Philippine Sea plate is distorted over a depth range of
26–33 km. A large variation of the reflected-wave amplitude with depth along the subducting plate was observed. At a depth of
about 20–26 km, the amplitude of the reflected wave is not large, and is explained by the reflected wave at the upper boundary
of the subducting oceanic crust. However, the reflected wave from reflection points deeper than 26 km showed a large amplitude
that cannot be explained by several reliable velocity models. Some unique seismic structures have to be considered to explain
the observed data. Such unique structures will provide important information to know the mechanism of inter-plate earthquakes.
D 2004 Elsevier B.V. All rights reserved.
Keywords: Philippine Sea plate; Reflected waves; Tokai region; Plate boundary; Crustal structure
1. Introduction
* Corresponding author. Tel.: +81 358415804; fax: +81
356897234.
E-mail address: [email protected] (T. Iidaka).
0040-1951/$ - see front matter D 2004 Elsevier B.V. All rights reserved.
doi:10.1016/j.tecto.2004.07.002
In the Tokai region, central Japan, the Philippine Sea
plate is descending beneath a continental plate at a
velocity of several cm/year. In this area, large inter-plate
earthquakes have occurred repeatedly. A knowledge of
the physical properties at the plate boundaries will help
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T. Iidaka et al. / Tectonophysics 388 (2004) 7–20
to understand the mechanism of large inter-plate earthquakes. The Tokai region is very important for understanding the mechanism of large inter-plate earthquakes.
Many numerical simulations of large earthquakes have
focused on this area (e.g., Kato and Hirasawa, 1999). In
numerical simulation studies the configuration of the
upper boundary of the subducting plate and the physical
property at the plate boundary are important. The
geometry of the subducting plate is one of the very
important parameters in the numerical simulation, but it
has not yet been well determined. Little is known of
physical properties at the plate boundary.
The detailed configuration of the subducting
Philippine Sea plate was not clear because it had been
Fig. 1. Location map of the profile line in 2001 experiment (A–AV), which crosses central Japan from north to south. The profile line traverses
the island arc from the Philippine Sea to the Japan Sea. 391 seismic stations are located on the line. The profile line of the experiment in 1985
(B–BV) with an east–west direction is also shown by small symbols (Matsu’ura et al., 1991). Shot points are shown by stars. The research area is
shown in the inset. The thick purple line in the Philippine Sea is the profile line of the experiment of JAMSTEC. The seismic stations of the
experiment of Aoki et al. (1972) are shown by triangles. The orange lines crossing central Japan are profile lines analyzed by Takeda (1997).
T. Iidaka et al. / Tectonophysics 388 (2004) 7–20
estimated from the distribution of micro-earthquakes
(e.g., Ishida, 1992; Yamazaki and Ooida, 1985;
Harada et al., 1998). Seismic reflection and refraction
studies are required to know the detailed configuration
of the subducting Philippine Sea plate. Several seismic
experiments with explosive sources have been done in
this area (Aoki et al., 1972; Matsu’ura et al., 1991;
Iidaka et al., 2003) (Fig. 1). In the study of Matsu’ura
et al. (1991), the seismic survey line was located
parallel to the Nankai trough with a length of 53 km
(B–BV in Fig. 1). Two later arrivals were observed and
interpreted as reflected waves at the boundaries with
depths of about 12–16 and 23–28 km. The boundary
with a depth of 23–28 km was considered to be the
upper boundary of the subducting Philippine Sea
plate. However, it is difficult to know the configuration of the subducting plate because the survey line
was located parallel to the Nankai trough. A seismic
experiment with the profile line along the dip direction
of the descending plate is required to reveal the
configuration of the subducting plate. Iidaka et al.
(2003) researched the crustal structure and the configuration of the subducting Philippine Sea plate by
forward-modeling using data from an active source
seismic experiment. Clear reflected waves from the
upper boundary of the subducting Philippine Sea plate
were observed. However, the configuration of the
subducting plate was not estimated well because of the
trade-off between the dip-angle of the boundary and
velocity structure.
In the Nankai region, which is located west of the
Tokai region, huge earthquakes have occurred cyclically. A low-velocity layer at the top of the subducting
Philippine Sea plate had been reported from active
source experiments (Kodaira et al., 2000, 2002;
Kurashimo et al., 2003) and ScSp phase studies
(e.g., Nakanishi, 1980; Nakanishi et al., 1981) in the
Nankai region.
A detailed analysis of the reflected waves is
required to reveal the physical properties at the upper
boundary of the subducting plate. A joint seismic
experiment was conducted in the Tokai and Chubu
areas, central Japan, in August, 2001 (Fig. 1). The
seismic experiment consisted of three parts: (1)
refraction study with explosive sources by the
Research Group for Seismic Expedition in Central
Japan, which is organized by universities, JAMSTEC
(Japan Marine Science and Technology Center), and
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other government organizations (Iidaka et al., 2003),
(2) refraction and reflection surveys in the oceanic area
by JAMSTEC (Kodaira et al., 2003), (3) reflection
study with CDP cable by several universities (Sato et
al., 2001). Here, we show the results of the refraction
study in central Japan. The objectives of the experiment are to know the large-scale structural variation of
the island-arc crust across central Japan, and to know
the configuration of the subducting Philippine Sea
plate. In the study, we research the lateral heterogeneity of the subducting Philippine Sea plate and
crustal structure using data from seismic experiments.
2. Data
The seismic experiment was conducted on 25 and
26 August, 2001. A 261.6-km-long profile extended
in the N–S direction to traverse the island-arc of Japan
from the south coast to the north coast (Fig. 1). We put
391 seismic stations along the survey line. Six
explosive sources were shot on the seismic survey
line. The charge sizes of the shots were 500 kg for J1,
J2, J3, J4, and J5, and 100 kg for T6. Verticalcomponent sensors with a natural period of 2.0 Hz
were used at 328 seismic stations. Three-component
geophones (4.5 Hz) were used at 63 seismic stations.
The average spacing of the seismic stations was 669
m. Digital recorders were used at the seismic stations
with a sampling frequency of 100 Hz.
3. Analysis
3.1. Crustal structure in the Tokai–Chubu region
A velocity model was constructed from observed
data by the following process. (1) First arrivals and
later phases at each seismic record were picked. (2)
The observed arrival time data located very close to
the shot points were used to estimate the velocity
structure of the shallowest layer. (3) The seismic
structure was estimated by forward modeling using a
ray-tracing method (Zelt and Smith, 1992).
The record sections are shown with a reduction
velocity of 6 km/s (Fig. 2). The amplitude of the
waveforms is normalized by the maximum amplitude
value of each trace. In the record section of shot J1 (Fig.
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T. Iidaka et al. / Tectonophysics 388 (2004) 7–20
Fig. 2. Record sections of shots J1 (a), J3 (b) and J5 (c). In the record section of J1, delayed arrivals appear at the Tonami basin (20–40 km in
distance). The remarkable two later-phases (L1, L2) are shown by arrows in the record section of J5 (200 to 20 and 100–0 km).
T. Iidaka et al. / Tectonophysics 388 (2004) 7–20
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Fig. 3. The reflected waves at the upper boundary of the subducting Philippine Sea plate (arrows) for shots of J4 (a) and T6 (b).
2a), the arrival times are delayed at ranges of 20–40 km.
Pn arrivals are first arrivals at ranges greater than 130
km. For shot J3 in the centre of the profile, the apparent
velocity is almost 6 km/s with some perturbation at
ranges greater than 20 km (Fig. 2b). In the record
section of J5 (Fig. 2c), two clear later arrivals are shown
at ranges of 200 to 20 km and 100 to 0 km. The
amplitude of the later arrivals, which are observed at
ranges of 200 to 20 km from shot J5, is very large
compared to the first arrivals (Figs. 2c and 3).
The obtained P-wave seismic structure is shown in
Fig. 4. The observed and modeled travel times are
shown in Fig. 5. The shallower part of the crustal
structure is well constrained by the travel times of the
Fig. 4. P-wave velocity structure model at central Japan. Several reflectors are shown in the crust. L2 in the crust is the boundary of geological
segments. L1 is the upper boundary of the subducting Philippine Sea slab. The reflected waves of the boundaries (L1 and L2) are shown in Fig. 2c.
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Pg phase. A lateral large variation of the sedimentary
layer was found. The Tonami basin is located in the
northern part of the survey line. The delay of arrivals
at profile distances of 20–40 km can be explained by
the structure with the low-velocity Tonami basin
(Figs. 4 and 5a). The thickness of the sedimentary
layer beneath the Tonami basin is estimated to be ~3
km. A layer with a P-wave velocity of 5.3–5.8 km/s
and a thickness of ~5 km is located beneath the
sedimentary layer along the length of the profile. The
P-wave velocity for the lower part of the upper crust is
estimated to be 6.0–6.4 km/s. The lower crust P-wave
velocity is 6.6–6.8 km/s. The thickness of the lower
crust is about 10 km.
The uppermost mantle velocity in the central
Japan region is estimated from the arrival time data
of the Pn phase. In this experiment, arrival-time data
with apparent velocity values of 7.7–8.0 km/s are
observed at ranges of about 140–170 km for shots of
J1 and J2. The velocity of the uppermost mantle is
7.6–7.9 km/s.
3.2. Configuration of the subducting Philippine Sea
plate
Two remarkable clear later arrivals are observed on
the record section of shot J5, which is the southernmost shot point (Fig. 2c). The origins of the later
phases were investigated. The two later arrivals are
explained by the reflected waves at two boundaries
beneath J5. The upper boundary is located at a depth
range of about 10–20 km (L2 in Fig. 4). Seismic
experiments with explosive sources have been done in
this area (Matsu’ura et al., 1991; Sato et al., 2001;
Kawamura et al., 2003). Sato et al. (2001) concluded
the reflector was the lithological boundary between
the Southern Shimanto belt and the Northern Shimanto belt.
The deeper boundary was located at depths of 20–
35 km (Iidaka et al., 2003). The location of the
deeper boundary (Figs. 4 and 5) is consistent with
that of the upper boundary of the subducting
Philippine Sea plate, which was estimated from
iso-depth lines on a seismicity map (Yamazaki and
Ooida, 1985). The later phase was identified as a
reflected wave at the upper boundary of the
subducting Philippine Sea plate. The configuration
of the subducting Philippine Sea plate was estimated
by Iidaka et al. (2003). However, the configuration
of the subducting plate was not estimated well
because the reflected wave was only observed at a
record section of one shot. It was very difficult to
locate accurately the subducting Philippine Sea plate
because of the trade-off between the dip-angle of the
boundary and velocity structure. To estimate the
location of the subducting Philippine Sea plate
precisely, the reflected waves have to be observed
at record sections of several shots at different
locations. In this study, the reflected waves are
detected on the record sections of shots J4 and T6 in
addition to shot J5 record section (Figs. 3 and 5e).
The configuration of the subducting Philippine Sea
plate is revealed. The dip angle of the Philippine Sea
plate obtained is 268 at the depth range of 20–26 km
(Fig. 4).
A flat Philippine Sea plate model with a dip angle
of 268 explains most of the observed data (Fig. 5).
However, the theoretical arrival-time curve does not
fit well with the observed values at profile distances
of 170–200 km (Fig. 5f-2). A distorted plate
boundary is required to explain the observed traveltime data. An adequate velocity-model with a
distorted plate boundary is obtained by forward
modeling with the ray-tracing method (Fig. 6). A
structural model, with the upper boundary of the
Philippine Sea plate distorted at the depth range of
26–33 km, was derived.
3.3. Depth variation of the reflected-wave amplitude
A remarkable reflected wave at the upper boundary
of the subducting Philippine Sea plate is observed on
the record section of the shot J5 (Figs. 2 and 3). We
researched the amplitude of the reflected waves.
Usually, site effects on amplitude data are expected
to be large. We have to estimate the crustal response
using seismic data with different sources. However,
the seismic stations were set up to observe explosive
shots, and were operated only for several hours. No
far-field earthquake was observed at the seismic
stations. We therefore investigated the amplitude ratio
of the reflected wave at the upper boundary of the
subducting Philippine Sea plate to the first arrival
wave.
The waveforms of the reflected wave at the upper
boundary of the subducting plate are not impulsive, so
T. Iidaka et al. / Tectonophysics 388 (2004) 7–20
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T. Iidaka et al. / Tectonophysics 388 (2004) 7–20
Fig. 5. Observed travel times and calculated travel times using the P-wave velocity model in Fig. 4 are shown by colored bars and black
symbols, respectively. (a), (b), (c), (d), (e), and (f-2) show the travel time data for the shots J1, J2, J3, J4, T6, and J5, respectively. (f-1) shows the
ray diagram of shot J5.
T. Iidaka et al. / Tectonophysics 388 (2004) 7–20
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Fig. 6. Ray diagram using a model with the distorted upper boundary of Philippine Sea plate. The reflection points of the large-amplitude
reflected wave are located at the depth of 26–35 km (white line). The reflection points, where the amplitude data can be explained by the
reflected waves at the layer with the velocity of 6.5 km/s, are located on the yellow line.
a time window with a length of 1.0 s is used for
selecting the first arrivals and reflected waves.
Maximum amplitude values of the two-phases are
picked for each window and compared. The amplitude
ratio (the maximum amplitude within the timewindow of the reflected wave/the maximum amplitude within the time-window of the first arrival) is
plotted along with distance from the shot in Fig. 7. At
a distance of 0–30 km, the amplitude ratio obtained is
less than 1.0. At a distance of 30–70 km, however, the
amplitude ratio is about 1.0–8.0. The large amplituderatio values are observed with a large perturbation. At
the seismic stations with a distance greater than 70 km
from the shot, the amplitude ratio obtained is about
1.0–4.0.
The amplitude of the observed body wave O(f) is
expressed by the following form:
Oð f Þ ¼ A4B4S ð f Þ4Pð f Þ4C ð f Þ4I ð f Þ
A is radiation pattern, B is geometrical spreading
factor, S( f ) is source spectrum, P( f ) is path effect,
C( f ) is crustal response, and I( f ) is instrumental
response. ( f ) is frequency. We are able to assume that
the S( f ), C( f ), and I( f ) parameters are cancelled
using the amplitude ratio of the reflected wave to the
wave of first arrivals. The value A is not considered
here because an explosive source is used. B and P( f )
were evaluated using a ray-tracing method. A raytracing code of SEIS83 (Cerveny and Psencik, 1983)
was used for the calculation.
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The attenuation effect along the ray path has to be
considered. Iwasaki et al. (1994) obtained a detailed
crustal structure in the Tohoku region, northern Japan,
using data from explosive sources. The Q structure
was estimated by analyzing amplitude data of the
explosive sources. The following Q structure model is
used for the calculation based on the results of Iwasaki
et al. (1994). The Q value of 100 is assumed for the
sedimentary layer. The Q value of 150 is given to the
layer beneath the sediment. The lower part of uppercrust and the lower-crust are assumed to have Q
values of 300 and 400, respectively. The Q value at
the uppermost mantle is assumed to be 600. The Q
value of the subducting Philippine Sea slab is given to
be 2000.
The geometry of the subducting Philippine Sea
plate with the flat plate boundary model (Fig. 4) was
used for the calculation. The distorted plate boundary
model (Fig. 6) is much better than flat plate boundary
model for explaining the observed arrival-time data.
The distorted plate boundary model has to be used for
the calculation to evaluate amplitude accurately.
However, the ray-tracing method is used in this
calculation. It is difficult to evaluate the focusingand defocusing-effects on amplitude data using ray
theory. The simple assumption (i.e., flat plate boundary) was therefore adopted for the calculation. In this
study, a thin layer with a thickness of 1 km was
assumed to be located at the top of the subducting
plate based on the results of previous studies
(Kurashimo et al., 2002; Kodaira et al., 2002). Several
velocity models with different P-wave velocities at the
thin layer were examined to explain the observed
amplitude data. The amplitude ratio is calculated for
the models with P-wave velocity values of 8.0, 6.5,
4.0, 3.0, and 2.0 km/s at the thin layer (Fig. 7).
In the distance range of 0–30 km, the observed
amplitude ratio of the reflected wave to the first arrival
can be explained by a model with a P-wave velocity
Fig. 7. Amplitude ratio data of reflected waves at the upper
boundary of the subducting Philippine Sea plate to first arrival
waves. (a) for ranges 0–30 km, (b) for ranges 30–70 km, and (c) for
ranges 70–160 km. The observed data (red lines and symbols) are
compared to the calculated data. The velocity model for the
calculated data has a thin layer with different velocity values at
the top of the slab. Velocity values are 8.0 km/s (Model-a), 6.5 km/s
(Model-b), 4.0 km/s (Model-c), 3.0 km/s (Model-d), and 2.0 km/s
(Model-e).
T. Iidaka et al. / Tectonophysics 388 (2004) 7–20
of 6.5 km/s at the upper boundary of the subducting
slab. The velocity is consistent with that of the
subducting oceanic crust at the Philippine Sea plate
(e.g., Kurashimo et al., 2002).
At distances of 30–70 km, large amplitude ratios
(1.0–8.0) are observed. In the calculation, the amplitude of the reflected wave increases, as the velocity of
the thin layer decreases. However, the observed
amplitude ratio value is much larger than that of the
low-velocity layer with a P-wave velocity of 2.0 km/s.
In the distance range of 70–160 km, the observed
amplitude data have ratios within 1.0–4.0, but the
theoretical amplitude ratios are less than 1.0, even
though an extremely low velocity model (V p=2.0 km/s)
is used.
The rays which arrive at distances of 0–30 km are
reflected at depths of about 20–26 km, and have
amplitudes consistent with those of reflected waves at
the boundary between the mantle wedge and the
subducting oceanic crust. The rays arriving at the
seismic stations at distances of 70 km and 160 km are
reflected at depths of about 30 km and 35 km,
respectively. At reflection points with depths of 26–30
km, extremely large amplitude ratios of 1.0–8.0 are
observed. At the depth range of 30–35 km, the
reflected wave amplitude is several times (1.0–4.0)
larger than that of first arrivals. At reflection points
deeper than 26 km, the amplitude ratio data suggest
that the amplitude of the reflected wave is much larger
than that of the model with an extremely low-velocity
layer (V p=2.0 km/s) located at the top of the slab.
4. Discussion
The crustal structure in this area was researched
by Aoki et al. (1972). The major difference between
the model of Aoki et al. (1972) and our model is
the thickness of the lower crust. A crustal model in
the Chubu region obtained by Aoki et al. (1972)
showed a lower crust with a thickness of 6 km,
which was estimated from data from seismic
stations with spacings of about 10 km. The spatial
density of the seismic stations was insufficient to
detect refracted waves from the lower crust. In this
study, the average spacing of the seismic stations is
669 m, which is sufficient to detect a refracted
wave at the lower crust, and a thickness of about 12
17
km was derived. The seismic velocity model
obtained here explains well the first arrival time
data at a distance range of 70–110 km for the shots
of J5 and T6 (Fig. 5).
The uppermost mantle velocity in the central Japan
region is estimated from the arrival time data of the Pn
phase. In this experiment, arrival-time data with
apparent velocity values of 7.7–8.0 km/s are observed
at the epicentral distance of about 140–170 km for
shots of J1 and J2. The velocity of the uppermost
mantle obtained is 7.6–7.9 km/s. The values are
consistent with those of the previous studies (Aoki et
al., 1972; Takeda, 1997).
The boundary L2 is identified as the lithological
boundary between the Southern Shimanto belt and
the Northern Shimanto belt. The location of the L2
boundary and the velocity value at the shallower part
of the boundary were well defined because many
rays traveled through the region. However, the
velocity value at the lower part of the L2 boundary
was not clear because few rays traveled through the
area. The location of the boundary was only derived
and the velocity change at the boundary was not
estimated.
The location of the upper boundary of the
Philippine Sea slab was estimated by Yamazaki and
Ooida (1985) using a seismicity map with the
assumption that the upper boundary of the slab is
located just above the locations of micro-earthquakes.
The upper boundary of the Philippine Sea slab
obtained here is 2–5 km shallower than that of
Yamazaki and Ooida (1985). In the Shikoku region,
central part of the Nankai trough, Kurashimo et al.
(2002) suggested that the upper boundary of the
Philippine Sea slab was located at 5–10 km above the
seismic region of micro-earthquakes from an analysis
of seismic refraction data. The upper boundary of the
subducting slab seems to be located a few kilometers
above the seismic zone.
The dip angle of the subducting Philippine Sea
plate of 268 is estimated at the depth range of 20–26
km. In this depth range, the reflected waves at the
upper boundary of the subducting Philippine Sea plate
are observed on the record sections of shots J4, J5,
and T6. The reflected waves, for which the reflection
points are deeper than 26 km, are only observed at the
record section of shot J5. It was difficult to estimate
the dip angle of the slab that is deeper than 26 km.
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The record section data of shot J5 (Fig. 6) suggests
that the subducting Philippine Sea plate deeper than
26 km is distorted. The shape of the distortion is not
well defined because of the trade-offs of several
parameters (i.e., velocity and shape, etc.).
In the Nankai region, located west of this research
area, large inter-plate earthquakes have occurred
cyclically. Seismic experiments with explosive sources have been done in this area (e.g., Kodaira et al.,
2000, 2002; Kurashimo et al., 2002, 2003), and a
seismic velocity structure model of the subducting
Philippine Sea plate obtained. The reflected wave
observed at the upper boundary of the subducting
plate has an amplitude that is much larger than that of
first arrival. Kurashimo et al. (2003) tried to estimate
the reflection coefficient at the boundary using the
Amplitude Variation with Offset (AVO) method. A
thin layer with a thickness of 200 m and velocity of
4.0 km/s is required at the top of the subducting
Philippine Sea plate. In the study of Kodaira et al.
(2002), the amplitude analysis of the reflected waves
at the subducting slab suggested that a thin lowvelocity layer with a velocity of 3 km/s was located at
the top of the slab with a depth range of about 15–30
km. Therefore, previous evidence suggests the existence of an extremely low velocity zone at the top of
the subducting Philippine Sea plate.
In this study, the amplitude data analyzed by the
velocity model is consistent with oceanic crust at the
top of the subducting Philippine Sea plate within the
depth range of 20–26 km. At a depth deeper than 26
km, however, the amplitude ratio is extremely large
compared to those of the theoretical models with
large variations of P-wave velocity (2.0–8.0 km/s).
To explain the observed large amplitude data, the
following models were considered: (1) layered
structure at the top of the subducting slab, (2)
focusing effect on the reflected waves. In this
analysis, the maximum amplitude values are
obtained for a time window with a length of 1 s,
because the reflected waves are not an impulsive
phase. If the uppermost part of the subducting
Philippine Sea plate has a layered structure, and
consists of many thin layers, the amplitude of the
reflected wave will be increased by the interference
of the reflected waves at the many layers. If the large
amplitude reflected wave is caused by a layering
structure at the top of the plate boundary, a spectrum
analysis of the large amplitude waveform should
suggest a dominant frequency corresponding to the
layering structure. However, the FFT analysis of the
reflected wave did not show any significant peak in
spectrum data.
If the upper boundary of the Philippine Sea plate
is distorted, some perturbation should be observed in
the arrival-time data of the reflected waves, and a
large amplitude reflected-wave might be observed by
the focusing effect at the distorted interface. A
distorted upper boundary model is required to
explain the observed travel-time data (Fig. 6). The
ray path of the reflected wave, which arrives at
profile distances of 170–200 km, reflects at the upper
boundary of the slab with the depths of 26–30 km.
The depth range is the large amplitude ratio zone
(Figs. 6 and 7b). The large-amplitude reflected
waves might be caused by the focusing effect. In
the Tokai region, a seismic survey with refraction
tomography and pre-stack depth migration of wideangle seismic data imaged a trough parallel cyclic
ridge subduction in the oceanic area (Kodaira et al.,
2003). If the subducted ridge is located at a depth of
26–30 km, the upper boundary should be distorted.
The distorted boundary might cause a large amplitude reflected-wave due to the focusing effect.
However, it has not been modeled well because the
configuration of the distorted boundary was not
obtained accurately, and the ray-tracing method is
not suitable for evaluating the focusing effect on
amplitude data.
5. Conclusions
A seismic experiment with six explosive sources
and 391 seismic stations was conducted in central
Japan. The crustal structure of the area was
obtained. A large lateral variation in the thickness
of the sedimentary layer was found. Beneath the
Tonami basin, the thickness of the sediment is
estimated to be ~3 km. The layer located beneath
the sedimentary layer has a P-wave velocity of 5.3–
5.8 km/s. The P-wave velocity of the lower part of
the upper crust is estimated to be 6.0–6.4 km/s.
The lower crust P-wave velocity is 6.6–6.8 km/s.
The P-wave velocity at the uppermost mantle is
7.6–7.9 km/s.
T. Iidaka et al. / Tectonophysics 388 (2004) 7–20
Large-amplitude reflected waves at the two
boundaries are observed. One of the boundaries is
located at depths of 10–20 km, which is the
lithological boundary between the Southern Shimanto belt and the Northern Shimanto belt. The
other boundary at depths of about 20–35 km is
interpreted to be the upper boundary of the
subducting Philippine Sea plate. The dip angle of
the plate is estimated to be 268 at the depth range of
20–26 km. The upper boundary of the subducting
Philippine Sea plate could be distorted at locations
deeper than 26 km.
To know the physical properties of the upper
boundary of the Philippine Sea plate, the amplitude
ratio of the reflected wave to first arrival was
researched. At depths of 20–26 km, the amplitude
ratio data is consistent with the wave being reflected
at the interface of the oceanic crust of the slab with a
P-wave velocity of 6.5 km/s. However, the amplitude
ratio is much larger than that of acceptable velocity
models at depths deeper than 26 km, which is the
depth range at which the upper boundary of the slab
is distorted. The large-amplitude reflected waves
might be caused by the distorted upper boundary
of the subducting Philippine Sea plate with a
focusing effect. However, the amplitude data has
not been modeled well because the configuration of
the distorted boundary was not obtained accurately,
and the ray-tracing method is not suitable for
evaluating the focusing effect on amplitude data.
Some unique seismic structure has to be considered
to explain the observed data. A unique structure will
provide important information for understanding the
mechanism of inter-plate earthquakes. More detailed
analyses of the waveform are required to reveal the
physical properties at the boundary, and provide
good information for understanding the mechanism
of inter-plate earthquakes.
Acknowledgments
This experiment was supported by a grant from
JAMSTEC. We thank all members of the Research
Group for Seismic Expedition in Central Japan for the
data acquisition employed in this experiment.
RAYINV and SEIS83 computer program codes were
used in the calculation.
19
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