Earth Planets Space, 57, 825–838, 2005

High resolution 3-D velocity structure in the source region of
the 2000 Western Tottori Earthquake in southwestern Honshu, Japan
using very dense aftershock observations
Takuo Shibutani1 , Hiroshi Katao1 , andGroup for the dense aftershock observations of the 2000 Western Tottori Earthquake2
1 Disaster Prevention Research Institute, Kyoto University, Kyoto 611-0011, Japan
2 The members of the group are listed at the end of this paper.

(Received June 24, 2004; Revised May 17, 2005; Accepted July 6, 2005)

We carried out high density aftershock observations a week after the 2000 Western Tottori Earthquake for 40
days. We deployed 72 seismic stations in and around the aftershock area. The average spacing of the stationsin the aftershock area was 4–5 km. We determined accurate hypocenters and focal mechanisms for ∼1,000
aftershocks and obtained a high resolution 3-D velocity structure in the source region. High P and S wave
velocity anomalies (> 4%) near the southeasternmost aftershock area at 2 km depth correlated with Jurassic
to Late Cretaceous plutonic and high pressure metamorphic rocks. The depth distribution of the P and S wave
velocities along the mainshock fault showed that high velocity anomalies were located at the shallow southeasternedge and the deeper central part of the aftershock area. The ratio between P and S wave velocities in the highvelocity anomalies was a little higher (∼1.75) than the average value (∼1.70) in the upper crust. These results
indicate that the high velocity anomalies could correspond to the plutonic or metamorphic rocks. The distributionsof the high velocity anomalies and large slips of the mainshock were complementary. These suggest that thehigh velocity anomalies could be stronger than the surrounding materials and might behave as barriers to themainshock rupture.
Key words: The 2000 Western Tottori Earthquake, aftershock observation, 3-D velocity structure, aftershock
distribution, focal mechanisms.

1.

Introduction

A large earthquake occurred on 6 October 2000 in thewestern part of Tottori Prefecture, southwestern Honshu,
Japan. The Japan Meteorological Agency (JMA) magnitude (Mj) of the mainshock was 7.3, while the moment
magnitude was estimated as 6.6 (Fukuyama et al., 2001).
It injured nearly 100 people and damaged more than 1,500
houses in and around the source area.
A quick report on this earthquake was compiled by
Umeda et al. (2001). Ohmi et al. (2002) provided detailed
descriptions on the focal mechanism of the mainshock andthe aftershock distribution of this event. The rupture process during the mainshock was studied by Sekiguchi and
Iwata (2001) and Iwata and Sekiguchi (2002).
This earthquake was preceded by swarm-like seismic activity in 1989, 1990 and 1997 during which there were six
Mj ≈ 5 earthquakes. The relation between these preceding swarms and the occurrence of the 2000 Western Tottori Earthquake was investigated by Shibutani et al. (2002).
They found that the preceding swarms occurred on the same
fault plane as that of the mainshock in 2000 and that thearea of the preceding activity did not overlap with the areaof large slip of the mainshock.
c The Society of Geomagnetism and Earth, Planetary and Space SciCopyright ences (SGEPSS); The Seismological Society of Japan; The Volcanological Societyof Japan; The Geodetic Society of Japan; The Japanese Society for Planetary Sciences; TERRAPUB.

Following the mainshock we carried out a dense aftershock observation with 72 stations from 13 October that
continued for one and a half months. The purpose of thisobservation was to determine precisely the locations andthe focal mechanism solutions of the aftershocks, to obtain a detailed 3-D velocity structure in the source region
by means of tomographic inversions of P and S travel timedata, and to image 3-D distributions of reflectors and scatterers by utilizing array analyses of reflected and scatteredwaves.
In this paper we describe the field observations, and show
the results of the aftershock distribution, the focal mechanisms of the aftershocks and the 3-D velocity structure in
the source region.
Hirahara et al. (1992) studied the relationship between
the slip distribution of the 1984 Western Nagano Prefecture
Earthquake (Mj = 6.9) and the 3-D velocity structure in
the source region. They found that the area with large
dislocations corresponded to low velocity anomalies on thefault plane. Zhao et al. (1996) studied the 1995 Hyogoken Nanbu Earthquake (Kobe Earthquake, Mj = 7.3) andfound low velocity anomalies with high V p/V s values near
the hypocenter (V p and V s indicate the P and S wave
velocities, respectively), and suggested that fluids mightplay an important role in the initiation of the mainshock.
We checked whether such features could be recognized forthe 2000 Western Tottori Earthquake. We also investigated825

826 T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE

130°

140°
20 km

35.5°

40°

01

Japan Sea
u

IKUM

02

nsh

Ho

Tottori Pref.
12

05
04

NITH

91

Shimane Pref.

03

11

42
KYT

Mt. Daisen

16

15
MZKH
20
13
06
43
51
21 18
22
57
HKTH SMNK
25
52
23
92
56 26
28
45
53 2927
32
24
46
30
55
54
31
34
47
MTC
35
36
37 38 40
48
10
49
39 33
TRT
14

SNT
JKR

44

QMT

50
SGOH

35°

41

SAIJ

Hiroshima Pref.

NKR1

133°

Okayama Pref.

133.5°

134°

Fig. 1. Map showing the location of the stations (cross: temporary onsite recording, cross with circle: temporary telemetered, cross with square:permanent telemetered) used in the aftershock observation of the 2000 Western Tottori Earthquake. The part of the codes “wt” common to thetemporary onsite recording stations is omitted in this figure. Dots denote the aftershocks used in this study. The star shows the starting point of themainshock rupture. The triangle denotes Mt. Daisen, a Quaternary volcano which has been inactive for 20,000 years. The dashed lines indicate theborders of prefectures. Active faults are shown by solid lines.

the relationship between the 3-D velocity structure and the the observation period, and extracted corresponding wavegeology in the source region. Finally, we want to clarify forms from the continuously recorded data.
the relationship between the source process of the 2000
We manually picked the arrival times of the P and S
Western Tottori Earthquake and the 3-D structure.
waves and the polarity of the P wave for 1,087 local events
which occurred from 15 October to 25 October. The on2. Aftershock Observations
sets of the P waves were sharp for most of the waveforms,
We deployed 57 stations with onsite recording and 2 so that we could read the arrival times with high accuracy
telemetered stations in and around the aftershock area from (typically ∼0.01 s). The P arrival times were read at more
13 October to early in December. We also used the data than 50 stations for 94% of the analyzed aftershocks. S
from 13 permanent stations together with the data from arrival times were picked on one of the horizontal compothe temporary stations. As shown in Fig. 1, the stations nents. Most of the S wave onsets were also sharp and were
were placed to densely cover the aftershock area along the read at more than 45 stations for 90% of the analyzed aftermain fault. The average station spacing was 4–5 km near shocks.
the center and 10–20 km near the periphery. The area of
The hypocenters of the aftershocks were initially deterthe stations extended over 65 km in the NW–SE direction mined using HYPOMH (Hirata and Matsu’ura, 1987) forand 100 km in the NE–SW direction. Details of the station a 1-D velocity model shown by the dotted lines in Fig. 2,
information are listed in Table 1.
which is used routinely for this region (Oike, 1975; Ohmi
At most of the stations, we installed 3 component short- et al., 2001).
period velocity seismometers (Lennartz LE3D [1 Hz]; Sercel (Mark Products) L22D [2 Hz]; L28B [4.5 Hz]; L15 [4.5 4. 1-D Velocity Models and Station Corrections
With the manually picked arrival times of the P andHz]), and recorded continuously with a sampling rate of 100
Hz (Clovertec DAT-2GC [16-bit A/D]). The GPS time in- S waves and the initially determined origin times andformation was also recorded every 6 hours for time correc- hypocenters, we determined simultaneously the hypocentral parameters for the 1,087 events, 1-D velocity models oftions.
More than 70 people took part in this aftershock deploy- the P and S waves and the travel time corrections for 68
ment from Kyoto University, University of Tokyo, Tohoku stations (Table 1) using a Joint Hypocenter Determination
University, Hokkaido University, Ehime University, Yama- (JHD) method (Kissling et al., 1994). Five stations were
gata University, Tottori University and Kyushu University. not used in the JHD relocation and the tomographic inversion (shown in the next section) because of insufficient time
All the participants are listed at the end of this paper.
accuracy. In this inversion the weight of the S waves was
3. Initial Hypocenter Determinations
50% of that of the P waves, and the damping factors were
We selected local events (Mj ≥ 1.7) from the JMA cata- set to 0.01 for the hypocentral parameters, 1.0 for the veloclog which occurred in and around the aftershock area during ity parameters and 0.1 for the station correction parameters.

T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE 827
Table 1. List of the stations deployed in the aftershock observation and the station corrections obtained by the JHD method.
Code
wt01
wt02
wt03
wt04
wt05
wt06
wt07
wt08
wt91
wt92
wt10
wt11
wt12
wt13
wt14
wt15
wt16
wt17
wt18
wt19
wt20
wt21
wt22
wt23
wt24
wt25
wt26
wt27
wt28
wt29
wt30
wt31
wt32
wt33
wt34
wt35
wt36
wt37
wt38
wt39
wt40
wt41
wt42
wt43
wt44
wt45
wt46
wt47
wt48
wt49
wt50
wt51
wt52
wt53
wt54
wt55
wt56
wt57
SMNK
MTC
KYT
SNT
TRT
QMT
NKR1
IKUM
JKR
SAIJ
HINH
MZKH
HKTH
NITH
SGOH

Longitude133.1755
133.1326
133.2236
133.1984
133.1366
133.2325
133.2123
133.3127133.1121133.1836
133.2238
133.2591
133.3320133.2502
133.3076
133.3430
133.4555
133.4079
133.4668
133.3649
133.3214133.4145
133.3634
133.2764
133.4002
133.4695
133.4233
133.3708
133.3142133.3539
133.5426
133.4969
133.4582
133.4739
133.3522
133.2856
133.3759
133.4302
133.4813
133.4370
133.5130133.4303133.5761
133.5695
133.6839
133.5287
133.6311
133.5844
133.5956
133.5353
133.5481
133.3242
133.3455
133.3389
133.3818
133.4439
133.3980
133.4364
133.3136
133.4724
133.8338
134.0203133.2016
133.8491
132.8228
133.0300
133.8200133.1150
133.3936
133.4942
133.2628
133.0883
133.3196

Latitude35.5572
35.491235.379135.335335.384435.300735.189335.3179
35.1995
35.270135.141535.3634
35.370235.334235.3500
35.352735.392135.337135.3279
35.3224
35.338935.2940
35.272335.2638
35.2274
35.2600
35.2275
35.233735.2557
35.2474
35.2269
35.2020
35.2150
35.1256
35.192335.178735.167635.155435.1406
35.1233
35.155435.027135.473635.323435.301535.2543
35.2319
35.190135.147635.1359
35.0643
35.2966
35.2826
35.2366
35.211635.214435.259035.270935.2924
35.189035.4392
35.410535.102235.0885
34.941035.4967
35.378334.998335.2278
35.352235.293135.220335.0461Height, m
10
20
60
660
400
190
680
120
460
340
545
40
180
140
70
80
185
140
300
110
70
140
330
360
410
190
230
550
460
405
760
420
260
660
260
380
445
610
770
520
620
360
115
725
560
335
600
550
525
520
420
170
190
430
255
260
240
360
210
430
100
200
480
330
330
20
180
480
308
222
54
239
461

 P, s
0.216
0.465
0.071
−0.015
0.114
0.036
—
—
−0.024
0.003
0.010
0.092
0.133
0.041
0.073
0.066
0.076
—
0.044
—
0.079
0.007
0.005
−0.0060.033−0.014
0.051
0.000
0.095
0.005
−0.116−0.107−0.053−0.0980.019
0.022−0.018−0.030−0.122−0.090−0.151−0.043
0.191
0.038
0.071
−0.106−0.137−0.122−0.071−0.079−0.102
0.015
0.017
0.044
0.013
−0.040
0.003
0.011
0.001
−0.114
0.119
0.175
0.002
−0.071−0.0040.411
0.083−0.025—−0.0160.033
0.041−0.014

 S, s
0.348
0.923
0.121
−0.085
0.178
0.059
—
—
−0.077−0.041−0.014
0.168
0.218
0.054
0.114
0.171
0.213
—
0.144
—
0.144
−0.031−0.043−0.0900.129−0.061
0.146−0.0400.206−0.012−0.269−0.215−0.142−0.135−0.034−0.023−0.117−0.015−0.206−0.142−0.258−0.161
0.415
0.143
0.401
−0.254−0.255−0.274−0.094−0.112−0.189−0.030−0.021
0.091
0.017−0.117−0.0240.066−0.066−0.2310.178
0.375−0.048−0.121−0.1340.804
0.132−0.043—−0.096−0.026−0.016−0.135

Remarks
TF
TF
TF
TF
TF
TF
TF
TF
TF, – 10/22 11:30
TF, 10/22 13:16 –
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF
TF, – 11/07 16:50
TF
TF
TN
TN
P, KU
P, KU
P, KU
P, KU
P, UT
P, JMAP, JMAP, JMA
P, NIED
P, NIED
P, NIED
P, NIED
P, NIED P and  S denote the station corrections (delays) for the P and S times, respectively. The values are relative to the P timeof station wt27. TF: temporary station with onsite recording; TN: temporary telemetered station; P: permanent telemeteredstation. KU: RCEP, DPRI, Kyoto University; UT: ERI, University of Tokyo; JMA: Japan Meteorological Agency; NIED:
National Research Institute for Earth Science and Disaster Prevention.

828 T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE

23

VELOCITY, km/s
4
5

6

7

1.6

1.7

Vp/Vs
1.8
1.9

2.0
-2

0

0

2

2

4

4

6

6

8

8
Vp

Vs

10

10

12

12

14

14

16

16

18

DEPTH, km

DEPTH, km

-2

18
2

3

4
5
VELOCITY, km/s

6

7

1.6

1.7

1.8
1.9
Vp/Vs

2.0

Fig. 2. 1-D models of the P and S wave velocities from −2 to 18 km depth. The V p/V s is also shown on the right side. The broken lines indicate
the velocity models which were used for the initial determination of the aftershocks. The solid lines denote the velocity models which were obtained
as the result of the joint hypocenter determination (JHD) method (see the text for details). The open circles indicate the depth points of the initialmodels for the tomographic inversion.

Number of events

500

(a)

(b)

(c)

500

400

400

300

300

200

200

100

100

0

0
0.0

0.1
0.2 0.0
Error_EW, km

0.1
0.2 0.0
Error_NS, km

0.1
Error_Z, km

0.2

Fig. 3. Histograms of the errors in the hypocenters of the aftershocks relocated with the JHD technique. (a) EW direction, (b) NS direction and (c) Z
direction.

Furthermore, the damping factors of the velocity parameters were multiplied by 5 for the eighth layer (13–16 km)
and 10 for the ninth layer (≥ 16 km). The velocity parameters of these layers were not well constrained because onlya small number of rays passed through the layers. Low velocity layers were not allowed.
The root mean square (RMS) of the travel time residuals
decreased from 0.177 s to 0.089 s and the variance reductionin the 1-D inversion was 75% (= (1 − (0.0892 /0.1772 ) ×
100).
Figure 3 shows histograms of the errors of the relocatedhypocenters. The errors were smaller than 0.10 km (EW),
0.09 km (NS) and 0.21 km (Z) for more than 90% of the analyzed aftershocks. The errors were significantly improved
as well as the hypocenters after the JHD relocation. The relocated hypocenters relative to the initial locations were at
most 0.4 km in a N30◦ W direction and 0.7 km shallower,
for more than 60% of the analyzed aftershocks. On theother hand, the relocated hypocenters relative to the JMA

hypocenters were at least 1.6 km in the same direction and1.9 km shallower, for more than 60% of the analyzed aftershocks. The differences between the relocated hypocenters and the JMA hypocenters were much larger than thosebetween the relocated hypocenters and the initial hypocenters. This indicates that the initial hypocenters were determined quite accurately because the distribution of the stations was dense and the travel time residuals, which should
correspond to the station corrections, were not mismapped
to biases in the hypocenters.
The resulting velocity models are shown by the solidlines in Fig. 2 and listed in Table 2. The P wave velocity
increases gradually from 5.96 km/s at 3 km to 6.24 km/s at
10 km. The S wave velocity also increases gradually from3.50 km/s at 3 km to 3.64 km/s at 10 km. The ratio of V p to
V s is nearly constant, 1.70–1.71, in this depth range. The
resolutions are greater than 0.9 and the standard errors are
smaller than 0.03 km/s.
The resolution of the first layer (−2–0 km) is small for

T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE 829

133°

133.5°

134°

(a)
20 km

35.5°

Shimane Penn.

Lake Naka-Umi
Mt. Daisen

Lake Shinji

O - C (P)- 0.1 s

35°

0.1 s

(b)
20 km

35.5°

Shimane Penn.

Lake Naka-Umi
Mt. Daisen

Lake Shinji

O - C (S)- 0.2 s

35°

0.2 s

133°

133.5°

134°

Fig. 4. Travel time corrections for the P and S waves derived from the JHD method. (a) P wave and (b) S wave. The values are relative to the P time ofstation wt27 which is indicated by being surrounded by the square in (a). Circles indicate positive residuals, while crosses indicate negative residuals.
Reference sizes are shown in the insets. The star indicates the starting point of the mainshock rupture. The triangle denotes the peak of Mt. Daisen.
Table 2. 1-D velocity structure obtained by the JHD method.

Layer

D, km

1
2
3
4
5
6
7
8
9

−2.0
0.0
1.0
3.0
5.0
7.0
10.0
13.0
16.0

V, km/s
4.57
4.71
5.70
5.96
6.15
6.23
6.24
6.24
6.60

P waveSE, km/s
0.024
0.038
0.008
0.004
0.004
0.005
0.026
0.003
0.000

R
0.102
0.411
0.992
0.998
0.998
0.996
0.906
0.006
0.000

V, km/s
2.31
2.55
3.24
3.50
3.61
3.64
3.64
3.66
3.81

S waveSE, km/s
0.031
0.036
0.006
0.002
0.002
0.003
0.009
0.011
0.000

R
0.181
0.602
0.996
0.999
0.999
0.999
0.989
0.080
0.000

D denotes the depth to the top of the layer. V indicates the wave velocity in the layer. SE and R are the standard error and theresolution of the velocity, respectively.

both the V p and V s, which means that the velocities in thelayer are not well constrained. This is partly because therays in the first layer are nearly vertical and their proportion
to the total length is small, and partly because of the trade
off between the velocities in the first layer and the station
corrections. However, 4.57 km/s seems to be acceptable for
the P wave velocity in the surface layer in this area. Resultsfrom a refraction survey (Tottori Prefectural Government,
2004) showed that V p in the surface layer ranged from 4.0
to 5.0 km/s.
The obtained station corrections for the P and S travel

times are illustrated in Fig. 4 and listed in Table 1. The station corrections are relative to the P time of station wt27,
which is located in the area of the Neu Granitic Pluton(mentioned in section 8), and at which the travel time residual averaged over all the analyzed aftershocks is very small(0.01 s for P and 0.08 s for S). Large positive residuals
in both P and S travel times are found at stations in the
Shimane Peninsula. They are due to thick sediments in the
Shinji Graben in which Lake Naka-umi and Lake Shinji arelocated (Sawada et al., 2001). Other positive residuals arefound around Mt. Daisen which is a Quaternary volcano

830 T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE

50

40

30

A

20

a

b

c

d

e

f

g

h

Y, km

10

0

-10

B

-20

-30

-40
-50

-40

-30

-20

-10

0
X, km

10

20

30

40

50

Fig. 5. Horizontal grid (inverted triangles) on which the P and S velocities are obtained in the travel-time tomography. The y-axis is rotatedanticlockwise from North by 30◦ , so it is parallel to the distribution of the aftershocks. Cross sections of the resulting velocity structure along
the thick lines AB, ab, cd, ef and gh are shown in Figs. 7 and 8. Open circles and crosses indicate the aftershocks and the stations, respectively.

and whose last activity was over 20,000 years ago (Tsukui, the station corrections should be subtracted from the P and1985).
S times before the tomographic inversion to obtain a better
velocity structure of the surrounding region.
Eberhart-Phillips (1990) and Thurber (1993) pointed out
5. Seismic Tomography and 3-D Velocity Structhat it is better to invert P and S times for V p and V p/V sture
We carried out a tomographic inversion of the travel rather than for V p and V s because the S times tend to betimes of the P and S waves. The data set was the same as fewer in number and lesser in quality than the P times, so
we used in the 1-D inversion. 65,126 P times and 58,130 that V s model solutions have lower resolution and greater
S times from the 1,087 events to the 68 stations were in- uncertainty than V p. However, the S times are comparable
verted for the 3-D heterogeneous structure of the P and S to the P times in both number and quality in our data set
velocities. We used a program developed from SIMUL3M as mentioned above. We successfully tuned the damping(Thurber, 1983). We extended it to use the S times and ob- factors for V p and V s so that the resolutions of V p andV s became nearly equal (see Fig. 7(g) and (h)). Therefore,
tain the S velocity structure, similar to the P√wave analysis.
The weights of the S times were 0.76 (≈ 1/1.73) of the our choice to invert the P and S times for V p and V s is
P times. The distribution of the grid is shown in Fig. 5 justified.
The RMS of the travel time residuals decreased fromwith the events and the stations. The grid intervals are 2
0.089
s for the 1-D inversion to 0.042 s. The variance re× 2 × 2 km in the vicinity of the aftershock area, so that
ductionin the tomographic inversion was 78% (= (1 −small scale heterogeneities in the velocity structure can be
2
2
0.042
/0.089) × 100). The total variance reduction in bothresolved. The initial velocity model for the tomographic
inversionswas
94% (= (1 − 0.0422 /0.1772 ) × 100). The
inversion was obtained by resampling the 1-D model obresulting
velocity
structure is shown in Figs. 6, 7, 8 and 11.
tained in the previous section, as shown in Fig. 2. The
Indrawingthesefigures,
we multiplied the perturbations ofstation corrections obtained in the previous section werethevelocitiesbythecorresponding
diagonal elements of thesubtracted from the observed P and S times. We carried
resolution
matrix,consideringtheresolutions
as the weights
out the first few runs of the tomographic inversions withoutforthevelocity
perturbations.
The
diagonal
elementsof thethe station corrections. The results showed, for example,
resolution
matrixhave
valuesbetween0
and1.
Valuesofthat a large low velocity anomaly in the Shimane Penin0
indicate
that
the
model
parametersarecompletely
unresula was smeared southeastwards and downwards to 6 km
depth. The depth distribution of velocity anomalies along solved and the initial values should be taken. Values of 1
the mainshock fault did not change, for runs with and with- indicate that the corresponding model parameters are perout the station corrections. Therefore, we concluded that fectly resolved and the final solutions can be calculated by

T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE 831

δVp/Vp

Y, km

0

20

40

Y, km

20

40

0

δ(Vp/Vs)

Vs CBR
20

40

0

20

40

0

20

40
40

20

20

0

0
2 km

2 km

2 km

2 km

1.76

40

40

20

20

0

0
Z=4 km

Y, km

0

40

Z=2 km

4 km

4 km

4 km

4 km

1.70

40

40

20

20

0

0
6 km

Z=6 kmY, km

δVs/VsVp CBR

6 km

6 km

6 km

1.70

40

40

20

20

0

0
8 km

Z=8 km

8 km

0
20
X, km

40

0
20
X, km

0
2
δVp/Vp, %4

0 1 2
Vp CBR, %

8 km

40

8 km

0
20
X, km

40

0
20
X, km

0
2
δVs/Vs, %4

0 1 2
Vs CBR, %

40

1.71
0
20
X, km

40

0.00
δ(Vp/Vs)

0.05

Fig. 6. V p and V s perturbations and their checker board resolutions at 2, 4, 6 and 8 km depths (top to bottom). The horizontal distributions of V p/V s
are also shown in the rightmost figures. The number at the bottom right corner denotes the average value of V p/V s at that depth. The star indicatesthe starting point of the mainshock rupture. The triangle denotes the peak of Mt. Daisen. Thin lines indicate active faults.

adding the perturbations to the initial values. Therefore, forall cases, we can calculate the final solutions by adding theweighted perturbations to the initial values. Since the resolutions of V p and V s are nearly equal as shown in Fig. 7(g)and (h), the differences between the V p resolutions andthe V s resolutions have little effect on the perturbations of
V p/V s. The checker board tests show how well an initialalternating pattern with ±2.5% velocity anomaly can be recovered by the simulated tomography with the same set ofpaths as in the actual case.
Figure 6 shows the lateral distribution of the perturbations of V p and V s at the depths of 2, 4, 6 and 8 km. The
checker board tests and the perturbation of V p/V s are also
shown in Fig. 6. At 2, 4 and 6 km, the checker board tests
show good recovery in relatively wide areas in and aroundthe aftershock distribution for both V p and V s. However, at
8 km, the good recovery is limited to the vicinity of the aftershock area. The high and low velocity anomalies of V phave, on the whole, the same patterns as those of V s. At
2 km quite a large low velocity anomaly is located southeastward of the starting point of the mainshock rupture.

At 4 km another low velocity anomaly exists northwestward of the mainshock starting point. In this low velocity anomaly, |δV p/V p| is larger than |δV s/V s| and consequently V p/V s is low. The low velocity anomaly continues
to 6 km in V p, but not in V s, which results in a low V p/V s
also at 6 km. At 6 and 8 km, the vicinity of the mainshock
starting point shows a high velocity anomaly of 2–4%. In
the area around Daisen Volcano low velocity anomalies arefound at various depths, however, the resolutions are unfortunately not sufficient to discuss in detail.
Figure 7 shows the depth distribution of V p and V s on
the mainshock fault plane (along the line AB in Fig. 5). The
checker board tests of V p (e) and V s (f) are good in almost
the whole area of the aftershock distribution, where most of
the diagonal components of the resolution matrix for bothV p (g) and V s (h) are greater than 0.5. High V p anomalies
with 4% or greater are located in patches at the central partand the southeastern edge of the aftershock distribution.
In the high V p patches, V s also shows a high velocity
anomaly and V p/V s is ∼1.75, which is significantly higher
than the average value (∼1.70) in the upper crust (3–16832 T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE

Y, km
10

A
20

B

Y, km
10

A
20

0

B
0

DEPTH, km

0

0

10

10
(a)

(b)

5.0

5.5
6.0
Vp, km/s

6.5

3.0
3.5
Vs, km/s

DEPTH, km

0

0

10

10
(c)

(d)

0
2
δVp/Vp, %02
δVs/Vs, %4

4

DEPTH, km

0

0

10

10
(e)

(f)

0
1
Vp CBR, %2

01
Vs CBR, %2

DEPTH, km

0

0

10

10
(g)

(h)
20

10

0

0.0 0.2 0.4 0.6 0.8
Vp R
20

20

10

0

0.0 0.2 0.4 0.6 0.8
Vs R

1.0
Y, km
10

0

1.0

-10

0
DEPTH, km

10(i)1.65 1.70 1.75 1.80 1.85 1.90
Vp/Vs
Fig. 7. Cross sections of the resulting velocity structure along the line AB in Fig. 5, which is almost the same as the mainshock fault plane. The depth
distributions of V p (a), V s (b), V p perturbations (c), V s perturbations (d), checker board resolutions of V p (e) and V s (f), diagonal components of
the resolution matrix of V p (g) and V s (h), and V p/V s (i). Open circles indicate the aftershocks which were located in the vicinity (±1 km) of thefault plane. The star denotes the starting point of the mainshock rupture.

T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE 833

Vp
0

0

a

Vp CBR
0

10
b

Vs
10

a

0

Vs CBR
0

10
b

10
0

DEPTH, km

10

10
18 km

Y=18 km
0 c

18 km

18 km
c

d

d

0

DEPTH, km

10

10
12 km

Y=12 km
e
0

12 km

12 km
e

f

f

0

DEPTH, km

10

0

10
6 km

Y=6 km
g

6 km

6 km
g

h

h

0

DEPTH, km

10

10
-2 km

Y=-2 km
0
X, km
5.0

5.5
6.0
Vp, km/s

10

6.5

-2 km

-2 km
0
X, km
0
1
Vp CBR, %

10

2

0
X, km
3.0
3.5
Vs, km/s

10

0
X, km
0
1
Vs CBR, %10

2

Fig. 8. Cross sections of the resulting velocity structure along the lines ab (uppermost figures), cd (figures in the second row), ef (figures in the thirdrow) and gh (the lowermost figures) in Fig. 5. The figures in the columns from the left to the right show the depth distributions of V p, the checker
board resolution of V p, V s and the checker board resolution of V s. Open circles denote the aftershocks which occurred in the vicinity (±1 km) of
each cross section. The star indicates the starting point of the mainshock rupture.

km) in this region. The mainshock starting point is locatedin a low velocity anomaly which is surrounded by highvelocity anomalies in the central part. At shallower depths
to the southeast from the mainshock starting point, a low
velocity anomaly extends horizontally, in which very few
aftershocks occur. This low velocity anomaly corresponds
very well with a large slip area (2–3 m) in the mainshock
rupture process (Iwata and Sekiguchi, 2002).
Figure 8 includes four vertical sections along the linesab, cd, ef and gh in Fig. 5. The depth distribution of V p andV s, and their checker board tests are shown in each verticalsection. The vicinity of the mainshock fault shows a highvelocity anomaly rather than a low velocity anomaly. The
velocity field does not show significant changes between
the northeast (N60◦ E) and southwest (S60◦ W) sides of thefault.

6.

Aftershock Distribution

Figure 9 shows the distribution of aftershocks whichwere simultaneously determined in the tomographic inversion. As shown in the epicentral distribution and the depth
distribution along the Y axis, the earthquakes are locatedin the vicinity of the fault plane of the mainshock and distribute from 2 to 13 km in depth in the southeastern twothirds of the aftershock area (−7 ≤ Y ≤ 15). In contrast, inthe northwestern third of the aftershock area (15 ≤ Y ≤ 20),
the earthquakes spread out spatially at shallower depths (2–7 km). Only small slip (< 1 m) was observed in the shallow
depths (< 10 km) of the northwestern third area and large
slip (≥ 2 m) was detected in the middle of the southeastern
two-third area (Iwata and Sekiguchi, 2002). The aftershock
activity in an hour after the mainshock was low in the northwestern area in comparison to the southeastern area (Ohmiet al., 2002). Therefore, the rupture directly associated tothe mainshock was considered to occur in the southeasternarea. This might be a cause of the differences of the aftershock distributions in the two areas. As shown in the crosssection along DD , the fault plane is almost vertical in thevicinity of the starting point of the mainshock (the star).
However, it dips to the southwest in the farthest southeastern part of the aftershock area as shown in the cross sections
along EE and FF . The aftershocks branch to the southwest
at Y = 11 km as shown in the epicentral distribution and thecross section along CC .

834 T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE

15

10

5

0
25

0

Y, km

10

B

B'

C

C'

D

D'

E

E'

F

F'

(a)

0

DEPTH, km-10
10
5
DEPTH, km

5

C

-5
C'

X, km
0

5

E

E'

5
Magnitude

10

5.0
3.0
1.0

15

B

B'

D

D'

F

F'

0
X, km

5

0

5

5

10

10

15
-5

0

5

10

0

-5

A'

X, km
0

15

5

0

A

-5

DEPTH, km

A'

5

DEPTH, km

15

A

DEPTH, km

20

15

X, km
0

-5

15
-5

(b)

0
X, km

5

-5

0
X, km

5

-5

0
X, km

5

(c)

Fig. 9. (a) The cross section of the depth distribution of the aftershocks along the y-axis. (b) The epicentral distribution of the aftershocks determined
simultaneously in the tomographic inversion. The y-axis is rotated anticlockwise from North by 30◦ . (c) The six cross sections of the depth distributionof the aftershocks along the lines AA to FF in (b). In these cross sections, events which occurred in the vicinity (±1 km) of the plane are plotted.
The star indicates the starting point of the mainshock rupture.

The changes of the hypocenters relocated in the tomographic inversion relative to those relocated by the JHD
method were at the most 0.2 km horizontally and −0.05– 0.75 km downwards, for more than 90% of the analyzed
aftershocks. Since the hypocenters changed little in the tomographic inversion, the decrease of the RMS of the travel
time residuals (0.089 s to 0.042 s) was mostly due to theimprovement of the velocity structure.
Ohmi et al. (2002) relocated the mainshock and aftershocks of the 2000 Western Tottori Earthquake with a
JHD technique using data from telemetered stations (permanent stations and the temporary online stations, SMNK
and MTC). They inverted the P and S travel time datafor hypocenters, 1-D velocity model and station corrections
which were represented in a quadratic form of the hypocenter parameters (x, y, z). Fukuyama et al. (2003) relocatedthe mainshock and aftershocks with a double differencetechnique using data from permanent stations. For the epicentral distribution, our results had similar features to thoseof the other studies. However, the results by Fukuyama etal. (2003) showed that the depths of the aftershocks were
deeper than those in our results. And the results by Ohmi etal. (2002) showed that the upper limit of the aftershock distribution became deeper to the northwest, while our results
did not show such a trend. This indicates that the temporary
stations just above the aftershock area greatly improved theaccuracy of the depths of the aftershocks.

above. Because of the high density of the station distribution, the focal mechanisms of most events were well determined. The ambiguities of the solutions were very small,even though the magnitudes of the events were small (Mj ≈1.7).
The resulting focal mechanisms and the P axis directions
are shown in Fig. 10. Following Yamanaka et al. (2002),
we classified the focal mechanisms into four types: normal
fault type with the plunge of the P axis > 60◦ ; reverse fault
type with the plunge of the T axis > 60◦ ; strike-slip typewith the plunge of the null axis > 45◦ ; and the remaining(intermediate) types. More than 80% of the determinedmechanisms are classified into the strike-slip type. Onlya few reverse fault type events are found.
In general, the P axes are approximately in a NW–SE
direction. However, lateral variations of the dominant directions of the P axes are found along the aftershock area.
In the northern part, the P axes are mostly in the NW–SE
direction. In the area around the epicenter of the mainshockand the southern part, events with P axes in a NW–SE direction and events with P axes between E–W and ESE–WNW directions co-exist. The NW–SE direction of P axes
is consistent with a mechanism with a fault plane parallel tothe trend of the entire aftershock distribution, which is thesame as obtained from a moment tensor inversion with thelong period waveforms of the mainshock (Fukuyama et al.,2001). The E–W direction of P axes agrees with the focal
mechanism of the initial rupture of the mainshock, which
was determined by the P wave first motion polarities (Ohmi
7. Focal Mechanisms of Aftershocks
Focal mechanisms of the aftershocks were analyzed us- et al., 2002). In the southeastern end of the aftershock area,
ing the method of Maeda (1992). The hypocenters, P most events have P axes within the E–W and ENE–WSW
wave polarities, takeoff angles and azimuths of ray paths directions. In the area of off-fault induced seismicity thewere available from the hypocenter determination described southwest from the main fault zone, the dominant direction

T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE 835

Northern
Part2

5 km

aroundmainshockNorthern
Part1

SE endSW
InducedSouthernPartP-axis Direction

Reverse fault type
Strike-slip type
Normal fault typeIntermidiate typeFig. 10. Focal mechanisms and the directions of P axes of the aftershocks. Types of the focal mechanisms are distinguished by the symbol for the
epicenter, strike-slip types are solid circles, reverse fault types are open circles, normal fault types are gray circles, and intermediate types are small
solid circles (see the text for the definition of the mechanism types). Bars through the epicenter circles represent the horizontal projections of the Paxes. Characteristic mechanisms for selected events are shown in upper hemisphere stereo projections. The rose diagrams show the relative numberof the P axes in 10◦ sectors for the six regions surrounded by the ellipses. The max at the bottom of the rose diagrams indicated the maximum numberin the sectors. The star indicates the starting point of the mainshock rupture.

of the P axes is approximately ESE–WNW, although the tions, such as the areas A–F, also show high V s anomaprecision of the mechanism determination is not as good as lies and relatively high V p/V s (∼1.75). In these areas thein the main fault zone.
slip is small (Iwata and Sekiguchi, 2002). The aftershock
distribution is terminated in the farthest southeastern highV p anomaly. Geologically this high V p anomaly corre8. Discussion
In this section, we discuss the relation between the result- sponds to Late Cretaceous plutonic rocks between the Saning 3-D velocity structure and the aftershock distribution, gun metamorphic sequence, which are high pressure typerocks that were metamorphized in the Jurassic (Teraoka et
the resistivity structure, and the geological structure.
Figure 11 shows the aftershock distribution with the per- al., 1996). These older plutonic and metamorphic rocks areturbations of V p, V s and V p/V s at 2, 4, 6 and 8 km depths. distributed along the southern rim of early Paleogene pluThe southeastern two-thirds of the NW–SE main branch of tonic rocks, namely the Neu Granitic Pluton (Iizumi et al.,the aftershock distribution is located mostly in high V p and 1982; Teraoka et al., 1996). The older plutonic and metaV s anomalies at 4, 6 and 8 km depths. However, looking morphic rocks correspond well to high velocity anomalies
closely at the aftershock distribution, few earthquakes oc- at 2 km depth, while the Neu Granitic Pluton can be charcur in the high V p anomalies larger than 4%, such as areas acterized by low velocity anomalies. Therefore, the other
A and B. On the other hand the northwestern third of the deeper high V p anomalies with high V p/V s may be conmain branch is located in low V p anomalies at 4 and 6 km sidered as plutonic or metamorphic rocks. They could be
depths (H and G, respectively) and in low V s anomalies at stronger than the surrounding materials and might behave
4 km depth. The NE–SW branch mentioned in section 6 is like barriers to the mainshock rupture and the aftershock
occurrence.
located at the southeastern edge of the low velocity zone.
The low V p anomalies located at G and H show low
The high V p anomalies with relatively large perturba-

836 T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE

20

Vp/Vp
10
5

15

0

-5

20

Vs/Vs
10
5

15

0

-5

20

15

(Vp/Vs)
10
5

0

-5

X, km

5
0

X, km

-5
5

X, km

Z=2 km

2 km

1.76

2 km

Z=4 km

4 km
B

1.70

4 km

C

G
6 km

1.70

6 km

A

0
-5

-5
5
0

D
Z=6 km

-5
5
0

E
H

0
-5
5

0

F

0
-5
5

X, km

5

-5
5
0

Z=8 km
20

8 km

15

-4

10
5
Y, km

-2

0
2
Vp/Vp, %

0

-5

4

20

15

-4

1.71

8 km
10
5
Y, km

-2

0
2
Vs/Vs, %

0

-5

4

20

15

10
5
Y, km

-0.05

0.00
(Vp/Vs)

0

-5

-5

0.05

Fig. 11. Lateral heterogeneities of the resulting velocity structure at 2 km (the uppermost figures), 4 km (figures in the second row), 6 km (figures in thethird row) and 8 km (the lowermost figures). The figures in the columns from the left to the right show V p perturbations, V s perturbations and V p/V s
perturbations. Open circles denote the epicentral distribution of the aftershocks. Green circles indicate the events which occurred in the vicinity (±1km) of each depth slice. The star indicates the starting point of the mainshock rupture. The thin line denotes an active fault (Kamakurayama-nanpo
Fault). Dashed lines indicate the borders of prefectures. See the text for the area A-H surrounded with the yellow lines.

V p/V s (∼1.65). Geologically these low V p anomalies
correspond to non-alkali volcanic and pyroclastic rocks
in the early to middle Miocene (Teraoka et al., 1996;
Sakamoto et al., 1982). Therefore, the above mentioned
NE–SW branch of the aftershock distribution is located at
the farthest southeastern edge of the zone of non-alkali volcanic and pyroclastic rocks.
As mentioned above, the high V p anomalies in the vicinity of the mainshock fault of the 2000 Western Tottori
Earthquake showed V p ≈ 6.5 km/s which was significantly
higher than V p (∼6.2 km/s) in the surrounding region at
the same depths. It is expected that (repeated) mainshock
faulting increased the degree of fractures in the vicinity of
the fault, and as the result the V p and V s were decreased.
However, it is difficult to explain that the velocities wereincreased by mainshock faulting. Furthermore, Figure 11
indicates that the velocity anomalies shown in Fig. 7 are
distributed not only on the fault but also in volumes off thefault. In the farthest southeastern area of the aftershock distribution, where the high velocity anomalies were locatednear the surface, the anomalies corresponded well to highpressure type metamorphic rocks and Late Cretaceous plutonic rocks. These results strongly suggest that the velocity anomalies, which were found using the aftershock data,
originate from the different geology and existed before theearthquake.
The resistivity structure has also been obtained near thesource area of the 2000 Western Tottori Earthquake (Oshiman et al., 2004). comparing the velocity and the resistivity
on a N–S cross section through the vicinity of the mainshock starting point, we found that the high and low velocity anomalies correspond quite well to the high and low
resistivity anomalies, respectively. In the above paragraphwe pointed out a possibility that the high velocity anomalies
correspond to plutonic or metamorphic rocks, which might
be strong and contain few cracks and less fluids. Consequently the high velocity anomalies show high resistivity.
Hirahara et al. (1992) investigated 3-D velocity structurein the source region of the 1984 Western Nagano Prefecture
Earthquake. They found that the areas with large dislocations corresponded to those with relatively low velocities
on the fault plane. This is the same relation we found forthe 2000 Western Tottori Earthquake. Mikumo et al. (1987)
showed that large dislocations occur in and around low velocity zones by numerical calculations of the fault rupture
processes in inhomogeneous media.
Eberhart-Phillips et al. (1990) and Eberhart-Phillips andMichael (1993) investigated 3-D velocity structure in theLoma Prieta Earthquake region and the Parkfield region, respectively. They pointed out strong agreement between theseismic velocity and the resistivity distributions across the
San Andreas Fault (SAF), as we found for the 2000 Western Tottori Earthquake. They also showed that the velocity
structure changes greatly across the SAF. However, our re-

T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE 837

sult indicates that the V p and V s were rather high in thevicinity of the mainshock fault and the velocity structure
does not change across the fault plane of the 2000 Western
Tottori Earthquake, as shown in Fig. 8. This can be dueto the difference in the tectonic settings, where the SAF isa transform fault between the Pacific Plate and the North
American Plate, while the fault of the 2000 Western Tottori
Earthquake is an active fault which was first recognized by
the occurrence of the earthquake. In the vicinity of a mature fault which has repeatedly caused large earthquakes,
a broad fracture zone has been formed and can be recognized as a low velocity anomaly. Therefore, we think thatthe mainshock fault of the 2000 Western Tottori Earthquake
represents young weak planes in a relatively homogeneous
structure, instead of a boundary between different geological structures.
Thurber et al. (1997) found zones of low V p and highV p/V s (≥ 2.0) beneath the SAF. The high V p/V s zones
extend to a depth of about 6 km and cut across areas ofvery different V p values. Therefore, they concluded thatthe high V p/V s values are due to the presence of fluids.
Zhao et al. (1996) found low V p and V s anomalies withhigh V p/V s values near the hypocenter of the 1995 Kobe
Earthquake. They suggested that the anomalies might bedue to overpressurized, fluid-filled, fractured rock matrices
and contribute to the nucleation of the mainshock. However, we did not find such high V p/V s areas which indicated the presence of fluids on the fault plane of the 2000
Western Tottori Earthquake, as shown in Fig. 7. This disagreement might be due to the difference in the tectonic settings and the degree of maturity of the faults.
As mentioned in the section of the focal mechanisms,
most of the aftershocks have strike-slip type focal mechanisms and P axes approximately in a NW–SE direction
for the 2000 Western Tottori Earthquake. On the contrary,
for the 1995 Kobe Earthquake, aftershocks along the mainshock rupture zone showed various directions of the P axes
and variety in the mechanism types. The reverse fault types
and the strike-slip types co-existed in narrow areas (Katao
et al., 1997; Yamanaka et al., 2002). This difference might
be due to differences in the degree of heterogeneity of theinitial stress and strength distributions in the source regionsof the two earthquakes.

9.

Conclusions

High density aftershock observation of the 2000 Western Tottori Earthquake successfully reveal high resolution
images of the 3-D V p and V s structure in and around thesource region, as well as precise locations and focal mechanisms of the aftershocks.
The lateral distributions of V p and V s at 2 km depth havegood correlations with geological structure. The depth distributions of V p and V s along the fault plane show thatareas with high V p and V s anomalies (> 4%) and highV p/V s (∼1.75) are located in patches of the central area
and the southeastern edge of the aftershock distribution.
These high velocity anomalies might correspond to plutonic
and metamorphic rocks of the Jurassic to the Late Cretaceous. The cross sections also show that there exist low
velocity anomalies at shallow depths (2–4 km) in the south-

eastern area and at depths (0–6 km) in the northwesternarea. The former might be the core of the Neu Granitic
Pluton of the early Paleogene, while the latter might correspond to non-alkali volcanic and pyroclastic rocks of theearly to middle Miocene.
The total slip during the mainshock is larger in the southeastern region than in the northwestern region of the aftershock distribution. In the southeast area the aftershocks arelocated in the vicinity of the fault plane, while in the northwest they spread out spatially at shallow depths. Furthermore, the slip is small on the high velocity patches and large
in the shallow low velocity anomaly of the southeastern region.
The southwest branch of the aftershocks is located at
the farthest southeastern edge of the low velocity anomaly,
which might correspond to the non-alkali volcanic and pyroclastic rocks.
The high and low velocity anomalies agree quite wellwith the high and low resistivity anomalies, respectively.
We suggest that the high velocity patches on the mainshock fault plane, which might correspond to older plutonic
or metamorphic rocks, contain few cracks and little fluid,
might be strong, and might like barriers to the mainshock
rupture.
Finally, we conclude that pre-existing structural heterogeneities could control the rupture process of the mainshockand the occurrence of the aftershocks in the case of the 2000
Western Tottori Earthquake.
Acknowledgments. We are very grateful to Western General Office of Tottori Prefecture who allowed us to use a room as a baseof the aftershock observation. Our thanks are also due to people inWestern Tottori, Eastern Shimane and Northern Okayama Prefectures who allowed us to put seismometers and data loggers on theirland. We thank Makoto Matsubara for his advices on retrieving
data from DAT tapes. We benefited from the thoughtful reviewsof Edi Kissling and Hiroaki Negishi and the suggestions of Eiichi
Fukuyama, the editor. We used waveform data from the following permanent stations; IKUM, JKR, SAIJ (JMA); HINH, HKTH,MZKH, NITH, SGOH (NIED); NKR1 (ERI, Tokyo Univ.); KYT,
QMT, SNT, TRT (RCEP, Kyoto Univ.). Generic Mapping Tool
(Wessel and Smith, 1991) was used for drawing figures. This
work was partially supported by Special Coordination Funds for
Promoting Science and Technology, from Ministry of Education,
Sport, Culture, Science and Technology (12800013).
The participants of this observation are listed below: S. Fujihara, Y. Fujisawa, Y Fujita, N. Hirano, M. Hirata, S. Itaba, K.
Ito, T. Ito, Y. Kano, Y. Kase, H. Katao, M. Koizumi, Y. Maeda,
T. Mizuno, J. Mori, T. Nakanishi, S. Nakao, K. Nishigami, S.
Ohmi, K. Onoue, T. Shibutani, A. Shito, F. Takeuchi, T. Ueno,
Y. Umeda, F. Yamashita, K. Yoshii, K. Yoshikawa, Y. Yukutake(Kyoto Univ.); N. Hirata, S. Ide, Y. Ichinose, Y. Iio, Y. Inoue, E.
Kurashimo, S. Nakagawa, I. Ogino, K. Sakai, C. Wu (Univ. ofTokyo); Y. Asano, F. Hirose, S. Hori, T. Kono, T. Maeda, Y. Matsuda, J. Nakajima, A. Nakamura, K. Nida, T. Okada, T. Saito,
Y. Suwa, T. Takahashi, H. Ueda, N. Ujigawa, N. Umino (Tohoku
Univ.); K. Katsumata, N. Wada (Hokkaido Univ.); K. Asamori, A.
Kuroki, R. Mita, O. P. Mishra, M. Nakamura, T. Nishino, F. Ochi,
T. Ono, M. Salah, I. Serrano, D. Zhao (Ehime Univ.); A. Hasemi,
T. Shibuya, K. Takizawa (Yamagata Univ.); R. Nishida (Tottori
Univ.); T. Matsushima, K. Uehira (Kyushu Univ.). The affiliationsof the participants at the time of the observation are shown in theparentheses.
Processings such as retrieving waveform data from DAT tapes,
merging them into files, etc. were mainly done by T. Shibutani, N.
Hirano, H. Katao, S. Ohmi, E. Kurashimo, K. Katsumata and T.

838 T. SHIBUTANI et al.: 3-D VELOCITY STRUCTURE IN THE SOURCE REGION OF THE 2000 WESTERN TOTTORI EARTHQUAKE

Okada. Picking up the arrival times was done by T. Shibutani, H.
Katao, S. Nakao, N. Hirano, F. Takeuchi, T. Ueno, A. Shito, Y.
Kano, T. Mizuno, K. Yoshikawa, Y. Maeda, M. Hirata, J. Mori, K.
Nishigami, S. Fujihara. The 1-D and 3-D inversions of the travel
time data were mainly done by T. Shibutani. The focal mechanism
analysis was mainly done by H. Katao.

References
Eberhart-Phillips, D., Three-dimensional P and S velocity structure in the
Coalinga Region, California, J. Geophys. Res., 95, 15343–15363, 1990.
Eberhart-Phillips, D. and A. J. Michael, Three-dimensional velocity structure, seismicity, and fault structure in the Parkfield region, central California, J. Geophys. Res., 98, 15737–15758, 1993.
Eberhart-Phillips, D., V. F. Labson, W. D. Stanley, A. J. Michael, and B.
D. Rodriguez, Preliminary velocity and resistivity models of the Loma
Prieta earthquake region, Geophys. Res. Lett., 17, 1235–1238, 1990.
Fukuyama, E., M. Ishida, S. Horiuchi, H. Inoue, S. Hori, S. Sekiguchi, T.
Eguchi, A. Kubo, H. Kawai, H. Murakami, S. Yamamoto, and K. Nonomura, NIED Seismic Moment Tensor Catalogue January–December,
2000, Technical Note of the National Research Institute for Earth Science and Disaster Prevention, 217, 1–131, 2001.
Fukuyama, E., W. L. Ellsworth, F. Waldhauser, and A. Kubo, Detailed fault
structure of the 2000 Western Tottori, Japan, Earthquake sequence, Bull.
Seism. Soc. Am., 93, 1468–1478, 2003.
Hirahara, K., N. Hirata, A. Ikami, H. Miyamachi, T. Yabuki, H. Aoki,
I. Fujii, T. Haneda, A. Hasegawa, S. Hashimoto, N. Hirano, S. Horiuchi, Y. Iio, Y. Ishiketa, A. Ito, K. Ito, T. Kanazawa, S. Kaneshima, I.
Karakama, M. Kobayashi, M. Koizumi, T. Kono, M. Kosuga, Y. Kurata,
S. Kuriyama, A. Kuroiso, T. Matsuzawa, T. Mikumo, T. Mitsunami,
K. Miura, Kazuaki Miura, R. Miyajima, M. Mizoue, T. Moriya, A.
Nakajima, I. Nakamura, M. Nakamura, T. Nakamura, K. Nishigami, K.
Oike, T. Okamoto, T. Okura, T. Ooida, T. Ouchi, T. Saeki, K. Sakai, T.
Shibutani, M. Suzuki, S. Suzuki, M. Takahashi, A. Takagi, F. Takeuchi,
T. Tanada, S. Tomita, T. Tsukuda, Y. Umeda, H. Wada, M. Yamada,
A. Yamamoto, K. Yamashina, F. Yamazaki, and M. Yokohama, Threedimensional P and S wave velocity structure in the focal region of the1984 Western Nagano Prefecture Earthquake, J. Phys. Earth, 40, 343–
360, 1992.
Hirata, N. and M. Matsu’ura, Maximum-likelihood estimation of hypocenter with origin time eliminated using nonlinear inversion technique,Phys. Earth Planet. Inter., 47, 50–61, 1987.
Iizumi, S. and Research Group for the Batholith in the San’in Zone,
The Neu Granitic Pluton—Petrographical study on the batholith in the
San’in Zone, Southwest Japan (Part 1)—, J. Geol. Soc. Jpn., 88, 299–
310, 1982 (in Japanese with English abstract).
Iwata, T. and H. Sekiguchi, Source inversion of recent earthquakes using
strong motion records, in a special issue on Strong Motion Prediction,
Chikyu Monthly, 37, 47–55, 2002 (in Japanese).
Katao, H., N. Maeda, Y. Hiramatsu, Y. Iio, and S. Nakao, Detailed Mapping of Focal Mechanisms in/around the 1995 Hyogo-ken Nanbu Earthquake Rupture Zone, J. Phys. Earth, 45, 105–119, 1997.
Kissling, E., W. L. Ellsworth, D. Eberhart-Phillips, and U. Kradolfer, Initial reference models in local earthquake tomography, J. Geophys. Res.,99, 19635–19646, 1994.
Maeda, N., A method of determining focal mechanisms and quantifying the uncertainty of the determined focal mechanisms for microearthquakes, Bull. Seism. Soc. Am., 82, 2410–2429, 1992.
Mikumo, T., K. Hirahara, and T. Miyatake, Dynamical fault rupture processes in heterogeneous media, Techtonophysics, 144, 19–36, 1987.
Ohmi, S., K. Watanabe, T. Shibutani, S. Nakao, N. Hirano, A. Nakao, H.
Takeuchi, F. Takeuchi, S. Miwa, A. Shito, A. Kim, and B. Enescu, The
2000 Tottori-ken Seibu Earthquake—Seismic activity derived from thecompilation of DPRI, JMA and Hi-net data—, Annu. Disas. Prev. Res.
Inst., Kyoto Univ., 44B-1, 273–282, 2001 (in Japanese with English
abstract).
Ohmi, S., K. Watanabe, T. Shibutani, N. Hirano and S. Nakao, The 2000
Western Tottori Earthquake—Seismic activity revealed by the regional
seismic networks—, Earth Planets Space, 54, 819–830, 2002.
Oike, K., On a list of hypocenters compiled by the Tottori Microearthquake
Observatory, Zisin (J. Seismol. Soc. Jpn.), 28, 331–346, 1975 (inJapanese with English abstract).
Oshiman, N. and Research Group for Crustal Resistivity Structure 2001,
Deep crustal resistivity structure in the focal region of the 2000 Western
Tottori Earthquake, Abst. Annu. Conf. Disas. Prev. Res. Inst. 2003, P56,
2004 (in Japanese).
Sakamoto, T., N. Yamada, and Y. Suda, Geological map of Japan
1:200,000 Matsue and Taisha, Geological Survey of Japan, 1982.
Sawada, Y., T. Tokuoka, S. Yamauchi, Y. Sampei, and K. Nishimura, A
newly discovered Pleistocene volcano in Miho Bay, along the axial zoneof the Shinji Rift System, Southwest Japan, J. Geol. Soc. Jpn., 107, 392–
405, 2001 (in Japanese with English abstract).
Sekiguchi, H. and T. Iwata, Near-source ground motions controlled bysource process in the case of the 2000 Tottoriken-Seibu earthquake,
Program. Abst., Seism. Soc. Jpn., A73, 2001 (in Japanese).
Shibutani, T., S. Nakao, R. Nishida, F. Takeuchi, K. Watanabe, and Y.
Umeda, Swarm-like seismic activity in 1989, 1990 and 1997 preceding
the 2000 Western Tottori Earthquake, Earth Planets Space, 54, 831–
845, 2002.
Teraoka, Y., H. Matsuura, H. Makimoto, F. Yoshida, M. Kamitani, T.
Hiroshima, M. Komazawa, and R. Shichi, Geological map of Japan
1:200,000 Takahashi, Geological Survey of Japan, 1996.
Thurber, C. H., Earthquake locations and three-dimensional crustal structure in the Coyote Lake area, central California, J. Geophys. Res., 88,8226–8236, 1983.
Thurber, C. H., Local earthquake tomography: velocities and Vp/Vs—
theory, in Seismic Tomography, edited by H. M. Iyer and K. Hirahara,
842 pp, Chapman and Hall, London, 1993.
Thurber, C., S. Roecker, W. Ellsworth, Y. Chen, W. Lutter, and R. Sessions, Two-dimensional seismic image of the San Andreas Fault in the
Northern Gabilan Range, central California: Evidence for fluids in thefault zone, Geophys. Res. Lett., 24, 1591–1594, 1997.
Tottori Prefectural Government, Investigation of the underground structurein the area related to the 2000 Tottori Seibu Earthquake, 18 pp., Tottori
Prefecture, Tottori, 2004 (in Japanese).
Tsukui, M., Temporal variation in chemical composition of phenocrysts
and magmatic temperature at Daisen volcano, southwest Japan, J. Volcanol. Geotherm. Res., 26, 317–336, 1985.
Umeda, Y., K. Matsumura, T. Shibutani, S. Ohmi, and H. Katao, The
2000 Western Tottori Earthquake—Precursory swarm earthquake, main
shock and aftershock—, J. JSNDS, 19, 501–512, 2001 (in Japanese withEnglish abstract).
Wessel, P. and W. H. F. Smith, Free software helps map and display data,
EOS Trans., Am. Geophys. Union, 72, 441, 1991.
Yamanaka, H., Y. Hiramasu and H. Katao, Spatial distribution of atypicalaftershocks of the 1995 Hyogo-ken Nanbu earthquake, Earth Planets
Space, 54, 933–945, 2002.
Zhao, D., H. Kanamori, H. Negishi, and D. Wiens, Tomography of thesource area of the 1995 Kobe Earthquake: Evidence for fluids at the
Hypocenter?, Science, 274, 1891–1894, 1996.

T. Shibutani (e-mail: shibutan@rcep.dpri.kyoto-u.ac.jp), H. Katao, andGroup for the dense aftershock observations of the 2000 Western Tottori
Earthquake