Seismic waves guided by untransformed oceanic crust subducting into the mantle: the case of the Kanto district, central Japan

Tectonophysics,

355

176 (1990) 355-376

Elsevier Science Publishers

B.V., Amsterdam

- Printed

in The Netherlands

Seismic waves guided by untransformed oceanic crust subducting into the mantle: the case of the Kanto district, central Japan S. HORI National Research Center for Disaster Prevention, Sciences and Technology Agency, Tennodai 3-1, Tsukuba, Ibaraki 305 (Japan) (Received

May 29,1989;

revised version

accepted

October

3,1989)

Abstract Hori, S., 1990. Seismic waves guided by untransformed district,

central

A model

Japan.

of the subduction

occurring

beneath

observed

at specific

association that

agreement suggests

the Kanto

of oceanic district,

stations

the latter

arrives

arrival

comparable

surface

located

range

of 40-60

analysis

indicate

with

a slightly

shallower

velocities

angle

structure

crust

can be directly

to the former.

related

in to

Close

or mantle

waves, respectively.

rocks

One can

Sea plate is apparently

in

of the X, and the X, phases

are

or the lower continental

is laterally

motions,

P and the X, phases

relative

of the Philippine velocities

to take place

ratio of crustal

and tangential

Apparent

earthquakes

from its particle

the initial

- P) and the V,/V,

island.

mantle

(Xn and X, commonly

is considered

including

where the upper surface

if the velocity

that these apparent

km, and

of wavelets

of the main layer of the oceanic

for an upper mantle earthquake

several earthquakes

the case of the Kanto

of the upper

Sea plate. The X, phase. is found,

of the crust of the Honshu

to the seismic velocities

never expected

in a depth

an analysis

and X, travel along the same path as longitudinal

with the bottom

into the mantle:

pair of later P and S phases

time ratio of (X, - S)/(X,

clearly observe these later phases at stations direct contact

through

A distinct

component

at the Earth’s

the relative

that both X,

Japan.

of the Philippine

seismic wave. A principal

between

crust subducting

crust is established

central

motion

oceanic

176: 355-376.

for these events

with the subducting

be a longitudinal shows

Tectonophysics,

homogeneous.

crust,

Travel-time

to the seismic velocity

which

are

curves for

in the source

region. The hypocentral

distribution

of the events related

form such a thin seismic zone as observed Sea plate. The earthquakes contorted

These characteristics subducted

portion

work into account, to eclogitic

from which the later phases

seismic zone. The deeper portion are best explained

of the gabbroic we conclude

rocks at depths

oceanic

Japan

are regarded

crust remains

60 km along the entire northern

A velocity anomaly near the upper surface of the descending oceanic lithosphere has been widely observed as a conversion interface from S,S to P dipping at least down to a depth of 300 km (Okada, 1971, 1979; Snoke et al., 1977, 1979; 0 1990 Elsevier Science Publishers

B.V.

Sea plate does not

thickness

of the Philippine

portion

of this thick and

such events.

if these later phases oceanic

of the Philippine

over the entire

occur only in a shallower

crust at the top of the Philippine

Introduction

0040-1951/90/$03.50

motion

but extends

are observed

does not generate

that the subducting

down to about

to the subducting

in southwest

as the seismic

Sea plate. Taking in a gabbroic boundary

waves

guided

by the

the result of our earlier

phase without of the Philippine

transformation Sea plate.

Nakanishi, 1980; Nakanishi et al., 1981). The ScSp wave is considered to be generated when an ScS wave passes through a thin low-velocity layer. Okada (1979) concluded that partial melting just above the upper surface of the lithospheric slab is the cause of this low-velocity layer. Nakanishi et al. (1981), on the other hand, pointed out that the

356

S. HORI

subducted

oceanic crust itself is possibly

this conversion.

Matsuzawa

related

the PS converted

wave from deep events

northeast

and

Japan

to

et al. (1986) analysed

concluded

that

sion consistent et al. (1985).

beneath

the

upper

with Fukao

Basalt or gabbro of the

oceanic

et al. (1983) and Hori

which constitutes

crust

(Yoder

seismic zone in a depth range from 60 to 150 km

Christensen

constitutes

eclogite under more extensive

seismic

a thin

velocity

rounding

mantle.

low-velocity

6% lower

than

that

with

velocity.

low-velocity

channel

They also suggested could

ducted oceanic crust. Fukao

the

of the sur-

This value is unexpectedly

for a mantle

et al. (1985)

channel

represent

low

that the the

sub-

et al. (1983) and Hori

on the other hand, proposed

that the

transformation from a low-pressure crustal phase to a high-pressure mantle phase in the oceanic crust

does not take place

subduction the phase

immediately

after

the

into the mantle, and that the delay of change makes the subducted oceanic

crust act as a low-velocity waveguide. They investigated subcrustal earthquakes occurring in a depth range from 30 to 60 km beneath southwest Japan and interpreted a prominent pair of P and S later phases, which are frequently observed for these earthquakes, as seismic waves guided by the subducted portion of the gabbroic oceanic crust. Hurukawa and Imoto (1987) and Hurukawa (1987) studied seismic velocities in a source region of the subcrustal earthquakes and also came to a conclu-

136'E

Fig. 1. Tectonic

map of central

1WE

Japan.

and

and Green,

1966; Cohen

1962; into

P-T conditions

such

crust (Ringwood

et al., 1967; Green

Ringwood,

1972; Mareschal

formation

of gabbro

take place

as the oceanic

trench to the Earth’s velocity of an eclogitic

Tilley,

1975) transforms

as at the base of the continental

and

et al., 1982). A trans-

to eclogite crust

must

therefore

subducts

from

a

interior. Since the seismic rock is slightly higher than

that of a peridotitic rock (Manghnani et al., 1974). the subducted oceanic crust would no longer be a low-velocity layer if the transformation occurs immediately after the subduction. However, Ito and Kennedy (1971) pointed out that the transformation possibly takes a considerable time at low temperatures, so that during subduction the oceanic crust may remain gabbroic to depths well below the continental Moho. Assuming subduction model, the transformation

a plausible depth was

thermodynamically evaluated by Ahrens and Schubert (1975) Delany and Helgeson (1978) and Anderson et al. (1980). However, their conclusions are strongly

140"E

The volcanic

and Salisbury,

a main layer

front is indicated

model-dependent.

144

142"E

by a dotted

Refraction

curve.

IP-Izu

peninsula.

studies

SEISMIC

WAVES

GUIDED

BY SUBDUCTING

UNTRANSFORMED

OCEANIC

have not yet succeeded in resolving the untransformed oceanic crust subducted into the mantle (e.g. Ludwig et al., 1966; Aoki et al., 1972; Yoshii et al., 1973; Spence et al., 1985) since it is too thin and too deep to be detected. The method of Fukao et al. (1983) and Hori et al. (1985) is a novel approach to overcome this difficulty. In this paper we apply their method to the upper mantle earthquakes occurring beneath the Kanto district, central Japan, where the Philippine Sea plate at its northeastern limb is subducting beneath Honshu island. Observation and analysis

Seismicity around the Kanto district As shown in Fig. 1, the Philippine Sea plate is now subducting from the Nankai-Suruga and Sagami troughs and colliding against Honshu island at the north of the Izu peninsula. The Pacific plate, on the other hand, subducts beneath the northeastern Japan arc from the Japan trench and also underthrusts the Philippine Sea plate along the Izu-Bonin trench. These two subducting motions characterize the tectonics around the Kanto district (McKenzie and Morgan, 1969; Nakamura et al., 1984; Kasahara, 1985; Le Pichon and Huchon, 1987). Figure 2a shows the epicentral distribution of the events in a depth range from 30 to 150 km beneath the Kanto district. Two vertical sections of the hypocentral distribution along the E-W direction are shown in Figs. 2b and c, where the upper boundary of the Pacific plate (hereafter abbreviated to UBP after Matsuzawa et al., 1986) is inferred from the hypocentral distribution as proposed by Ishida (1986). The approximate location of the Moho boundary (Milcumo, 1966; Horie and Shibuya, 1979) is also shown by a broken line. The upper mantle earthquakes are classified into two groups. One comprises the events occurring along or below UBP. These earthquakes are interpreted as an activity associated with the subducting motion of the Pacific plate (Maki, 1984; Ishida, 1986). Several clusters of seismic activities are recognized along UBP (e.g., activities in a deeper portion of region B and in region C). These earthquakes are regarded as interplate events be-

CRUST

351

tween the Pacific and the overlying Philippine Sea slabs (Ohtake and Kasahara, 1983; Kasahara, 1985). The other group comprises the events distributed well above UBP. Some of them are also densely distributed in region A and in a shallower portion of region B. These earthquakes are considered to take place in conjunction with the subducting motion of the Philippine Sea plate. Close agreement between the slip direction of the events in region A and the direction of motion of the Philippine Sea plate relative to the Eurasian plate (Seno, 1977; Minster and Jordan, 1978) suggests that these earthquakes are caused by a seismic coupling between the two plates (Kasahara, 1980; Takemoto and Kawasaki, 1983; Kasahara, 1985; Kawasaki and Matsuda, 1987).

Later phase We analysed waveform data of 127 events of the upper mantle earthquakes occurring beneath the Kanto district. Their hypocentral parameters are given by the National Research Center for Disaster Prevention (Matsumura et al., 1988). Figure 3 shows their epicentral distribution and the vertical section along the E-W direction. The upper mantle earthquakes occurring above UBP that are mainly selected for the present study as those related to the subducting Philippine Sea plate. Figure 4a shows an example of three component seismograms of an event occurring in region A (denoted by El in Fig. 3) recorded at the MKB station (see Fig. 8 for its location). One can clearly observe a pair of P and S later phases (denoted by X, and X, respectively). The amplitudes of the initial P and S phases (denoted by P and S respectively) are, in contrast, so small that the onset of the initial S cannot be detected from the waveform in this case. We estimated the onset with reference to the theoretical arrival time calculated from the velocity structure proposed by Ukawa et al. (1984). The “real” onset of the initial S phase is considered to be close to the theoretical value since the hypocentral parameters are well determined, with a number of other “reliable” travel-time data obtained from densely distributed stations. Figure 4b shows another seismogram re-

S. HORI

358

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50km

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EW III STANCE

(km)

(km)

SEISMIC

WAVES

GUIDED

BY SUBDUCTING

UNTRANSFORMED

OCEANIC

359

CRUST

b w

Depth

(km)

0.

EW DISTANCE Fig. 2. (a) Epicentral distribution (January-December

of the upper mantle earthquakes

1988). (b) Vertical section of hypocentral

(km) (30 km d depth d 150 km) occurring beneath the Kanto district

distribution

of the events occurring in the rectangular region a-a’

Fig. 2a along the E-W direction (July 1979 - August 1989). (c) Vertical section of hypocentral the rectangular region b-b’ of Fig. 2a along the E-W broken line. UBP-Upper

distribution

direction (July 1979 - August 1989). The Moho boundary

boundary of the subducting

corded at the same station but for a subcrustal earthquake occurring beneath the Kinki district (see Fig. 1 for its location). Although the waveform seen in Fig. 4a is more complicated, the appearances of the initial and the later phases are quite similar to those in Fig. 4b. Fukao et al. (1983) and Hori et al. (1985) concluded that the later phase as seen in Fig. 4b is a particular seismic wave guided by the gabbroic oceanic crust subducting into the mantle. The following analysis will show that the later phase in Fig. 4a is of similar origin. Figure 5 illustrates the particle motions of the initial P and X, phases for the waveform data shown in Fig. 4a. The width of each time window is taken as 1 s from its onset. For simplicity, no correction is made for an effect of the surface response. This diagram shows that the X, phase, as well as the initial P phase, is dominant on the vertical and radial components and weak on the

of

of the events occurring in is indicated by a

Pacific plate proposed by Ishida (1986).

transverse component, and indicates that the X, phase is a longitudinal wave. The apparent angle of incidence of the X, phase (47 “) is, however, slightly larger than that of the initial P phase (38’ ). We examine the polarization of each phase with a principal component analysis method (Montalbetti and Kanasewich, 1970). The polarization is represented by eigenvectors of the covariante tensor C defined as follows:

(la) 1 cij= N-n+1 (i, j=l,...,3)

(lb)

where { xi( t,); i = 1,. . . ,3} are mutually orthogonal three-component seismic signals. The direction of dominant oscillations agrees with the direction

360

S. HORI

36.5

3b.0

1

139.0 E

139.5 E

I

140.0 E

140.5 E

IA1.0 E

L -___

40

:

b0. 5 E80.y %: 100

0 120.

140. (b)

j -100.

-80.

-60.

-40.

-20.

0.

20.

40.

b0.

80.

100.

EW DISTANCE,km Fig.

3.(a) EpicentraJ distribution

in the rectangular

of earthquakes

analysed

region of Fig. 3a along the EW direction.

in the present

study. (b) Vertical

Seismograms 9,ll

and 14.

section of hypocentral

for the events denoted

by

distribution

plotted

El-ES will appear in Figs. 4, 5,

SEISMIC WAVES GUIDED BY SUBDUCTING UNTRANSFORMED OCEANIC CRUST (a) 1lhOOm Apr.10, 1987

NS -

0

10

20

30

40

so

60

70

60

60s

40

. 60

60

70

60

60s

(b) 10kZlm Dec.18.1983

! I

lKl

f4s

EW

0

i 10

I

,I.1 7.0

a

I 1

Fig. 4. Three-component seismogramsrecorded at the MKB station (see Fig. 8 for its location). (a) Records of the event EI of Fig. 3. (b) Records of a subcrustal earthquake Occluding beneath the Kinki district. P, S, Xp and X, represent the onsets of the initial P, the initial S, the later P and the later S phasea, respectively. The onset of the initial S of Fig. 4a is estimated from the theoretical arrival time given by the standard travel-time table proposed by Ukawa et al. (19&1).

of an eigenveotor corresponding to the maximum eigenvalue (Matsumura, 1981). Each component of Z is directly calculated from the E-W, N-S and vertical seismic data arranged as velocity amplitudes without correction for the surface response. The time window of x,( tk) for the summation of eqn. (lb) is approp~ately determined to include a characteristic waveform. Polarization of P and X, obtained in this way is shown in Fig. 6 as an apparent angle of incidence (an angle between the eigenvector and the normal of the Earth’s surface). A positive but weak correlation is recognized between the apparent angles of incidence of P and X,. A considerable number of the plots, however, significantly deviate upward from

361

the equal-angle line, indicating the same tendency as recognized in Fig. 5. As can be seen in Fig. 4a, the X, - S time is substantially greater than the X, - P time. We directly read the onset of each phase from waveform data and plot X, - S versus X, - P in Fig. 7. Although the X, - S time possibly includes a considerable error because of the ambiguity of .tbe initial S onset, a positive correlation is clearly The rereco8nized between X, -P and X,-S. gression line gives an averaged value of (X, - S)/ (Xr - P) of 1.8, which agrees with a typical value of VJV, of either the crustal or mantle rock (e.g. Ukawa and Fukao, 1981). This observation suggests that both the X, and X, travel along the same path as longitudinal and tangential waves, respectively. The later phases are not always observed at every station. We classify the observation stations into three groups (groups I, II and III) by the appearances of the initial and later phases. Figure 8 shows the distribution of the stations and several examples of the seismic records in the vicinity of their onsets. The later phases are clearly observed at stations indicated by solid circles (group I). No systematic later phases are found at stations indicated by open circles (group II). At stations indicated by solid asterisks (group III) the later phases are seen but are weak. The stations of group I are located in the north of the Izu peninsula (e.g. ENZ) and in the south of the Tokai district (e.g. MKB). The stations of group II, on the other hand, are located near the source region and in the north of the Kanto-Tokai district (e.g. OTR and KIB). At these stations the initial phase appears with its clear onset and is not accompanied by any distinctive later phases. Figure 8 also shows the contour lines of the upper surface of the subducting slab of the Philippine Sea plate as proposed by Ishida (1989). The subduct~g slab undulates beneath the region of interest so that its surface lies at relatively great depths in the source region (i.e., in the large part of the Kanto district) and is in direct contact with the Moho in the north of the Izu peninsula and in the south of the Tokai district. Most of the stations of group I are distributed in this direct contact zone. The later phases are, however, hardly seen at stations located in the

362

S. HORI

STATION:MKB IIhOOm AR?10 1987 I2

t3

I‘

IS

16

17

is

UD

T

i

20

21

22

23

2‘

25

26

27

28

29

30

3,

\ :.“+A V

V

P

19

R

R

_

\38

Fig. 5. Particle motions of the initial P and the X, phases recorded at MKB for the event E, of Fig. 3. Y, R and T represent the vertical, radial and transverse directions. The vertical seismogram is also shown where the time windows to draw the particle motions are indicated by broken lines.

south of the Izu peninsula (e.g. SMD), which is considered to be a microcontinent on the Philippine Sea plate (Sugimura, 1972; Matsuda, 1978). The seismic traces of an event in region A (denoted by E2 in Fig. 3) recorded at stations of group I are assembled in the order of epicentral distances in Fig. 9. The onsets of the initial and later phases are indicated by solid circles and asterisk respectively. We note the following two characteristics. First, the X, and the X, phases have apparent velocities of 7.1 and 4.0 km s-l, respectively. These values are anomalously low in comparison with an upper mantle earthquake, because apparent velocities of any later phases are expected to be greater than the P and S velocities

in the source region if the velocity structure is laterally homogeneous. The observed velocities are rather more comparable to the seismic velocities of the main layer of the oceanic crust or the lower continental crust. Second, a considerable discontinuity in the initial arrival time is recognized between the third and fourth stations from the source. The offset of arrival times is seen for both the initial P and S phases but not for the later phases. Table 1 summarizes the apparent velocities of the initial and the later phases obtained by a least-squares fit for individual events. The averaged values become 7.7, 4.4, 7.0 and 4.0 km .s-’ for the initial P, the initial S, the X, and the X,

SEISMIC

WAVES

GUIDED

BY SUBDUCTING

UNTMNSFORMED

OCEANIC

363

CRUST

361

351N

xii_,, .., ,./ 1

30.0

20.0

10.0 ADParent

incident

annle

40.0 of

P.

,

0

I8

341N

9

_ 1 137E

6 Q

I

138E

I

139E

I

140E

I

141E

50.0

de&

Fig. 6. ReIation between the apparent incident angles of the initial P and the X, phases obtained by a principal component altalysiS.

SMD phases, respectively. The values for the initial phases are, however, of little significance because of the offsets of their arrivals. Figure 10 shows the composite travel-time plots of the initial and later phases. The inclinations of the regression lines

,

xs-$48 xp-P

w

Fig. 8. Top: distribution of the observation stations. Locations of the stations are indicated by solid circles when the later phases are clearly observed (group I) and by open circles when not identified (group II). The later phases are seen but only maqirutlly at stations indicated by solid asterisks (group III). Contour lines of the upper surface of the Philippine Sea plate proposed by Ishida (1989) are drawn at intervals of 10 km. Vertical sections along the broken arrows of A, B and C will appear in Fig. 13. Bottom: examples of vertical seismograms recorded at the stations denoted by three-character codes. The onsets of the initial P and the X, phases are indicated by solid circies and asterisks, respectively.

,/

qb’ ,’

I’

OoC'

I

1

2

3

4

5

6

xp-P, s

Fig. 7. Relation between the X, -P and the Xr -S times. The scale of the vertical axis is compressed to be a half of the horizontal axis. The onset of each phase is directly read from waveform data.

7

again yield apparent velocities of 7.0 and 4.0 km s-l for the X, and X, phases, respectively. On the other hand, a single line cannot be drawn to fit the initial arrivals over the entire distance range because of the offset of arrival times at epicentral distances of 150-170 km. The offset value amounts

Jan. 2. ‘84

139.898”E

36.113-N

47.9 km M4.1

Fig. 9. Record sections of seismograms for the event E2 of Fig. 3, assembled in order of epicentral distances. Tbe origin time, the location of the hypocenter and the magnitude of the event are indicated at the top (the origin time is modified as GMT). (a) Vertical seismograms for P and X,. (b) Transverse seismograms for S and X,. Reduced velocities are 8.0 km s-’ and 4.5 km s-’ for P and S, respectively. The onsets of the initial and later phases are indicated by solid circles and asterisks.

09h22m27.44s

r $j E

SEISMIC

WAVES

GUIDED

BY SUBDUCTTNG

UNTRANSFORMED

OCEANIC

to more than 1.5 s for both the initial P and S phases at this distance. Straight lines fitted to the travel-time data before and after the offset give

TABLE 1 List of apparent velocities of the later phases NFJ

C, (km s-l)

C &I

N,

G (km s-l)

G (km s-‘)

s-l)

4

7.62

6.64

3

4.37

4.00

8

7.46

6.96

4

4.01

4

8.79

8.47

1

4.53 -

4

7.58

7.01

3

4.45

3.90

5

7.60

6.72

4

4.23

3.89

9

7.78

6.80

7

4.47

4.03 4.09

6

7.95

7.00

2

4.42

8

7.44

6.97

3

4.00

3.99

10

7.69

6.99

3

4.50

4.19

4

8.53

6.47

1

9

7.63

7.30

1

-

11

7.76

7.20

2

4.99

3.72

9

7.68

7.11

5

4.42

4.06

1

3

6.68

6.18

6

7.68

6.99

3

4.65

4.13

10

7.72

7.17

6

4.25

3.93

5

8.92

1.27

1

5

8.51

6.79

1

-

5

8.46

6.68

1

-

11

7.80

7.08

5

4.43

4.14

7

7.60

6.92

4

4.48

3.81

6

7.92

6.56

3

4.23

3.82

11

7.71

6.99

2

4.33

4.04

16

7.75

7.01

3

4.89

4.37

8

7.76

7.05

3

4.29

13

7.51

6.69

1

4.44 -

3

7.67

6.61

1

6

7.45

7.10

1

5

7.31

6.69

1

13

7.51

7.15

4

5

8.08

6.51

1

12

7.63

7.22

1

5

6.97

1.24

1

-

4

8.10

6.55

1

-

6

7.05

7.2s

1

-

7

7.56

7.05

4

4.40

3.93

8

7.38

6.77

5

4.38

4.08

4.07 -

365

CRUST

apparent P and S velocities of 8.6 and 5.0 km s-l and 8.1 and 4.6 km s-l. respectively. These velocities are higher than the sub-Moho velocities of 7.8-7.9 and 4.4-4.5 km s-t in this region (Horie and Shibuya, 1979; Ukawa and Fukao, 1981). Since there is no systematic offset in the later arrivals, the offset in the initial arrivals is considered to have its origin in the propagation path of the initial phases. Figures lla and b show the P wave seismograms for two events, the locations of which are denoted by E3 and E4 in Fig. 3. These two hypocenters are on either side of region A so that their travel-time difference at a common station reflects the P velocity averaged over region A. The systematic later phases, with an apparent velocity of 6.8 km s-1, are clearly observed in both sections. Note that the travel-time curves of these phases have the same intercept time. This observation implies that the seismic wave appearing as the later phase travels through region A at a speed of about 7 km s-l. The agreement of this value with the apparent velocities given in Table 1 indicates that the latter values directly reflect the in-situ P velocity. This seismic velocity in the source region is significantly lower than the velocity of the subMoho mantle and rather more comparable to that of gabbroic or basaltic rocks. Spatial characteristics of the event exhibiting the later phases In southwest Japan a thin slab-like dipping seismic zone in conjunction with the subducting

3.78

Notes to Table 1 Averages:5 = 7.69f 0.33 km s-l C,=

6.96kO.28

c,=4.43*0.19

km s-l km s-1

9

7.77

6.86

7

4.67

4.17

10

7.78

6.80

5

4.42

3.83

13

7.66

6.75

6

4.30

3.81

Np--Number

12

7.51

6.80

1

10

7.53

7.16

5

4.08

7

7.56

7.05

1

4.48 -

parent velocity of the X,

9

7.31

6.57

1

-

the initial S phase; C;*-apparent

16

7.78

7.12

5

4.54

The average is the mean value weighted by the number of

4.05

c,,=

4.00*0.15

km s-1

of observations of the initial P and the X,

phases; C,-apparent

velocity of the initial P phase. C.+--Ap-

phase; lV--number

of the initial S and the X,

observations.

of observations

phases; C,-apparent

velocity of

velocity of the X,

phase.

366

S. HOI21

m

'0

(~~~‘~~~/‘~/~~‘~~~~‘~~~~‘~~~(‘~~~~’(~~~ 50

100

150

200

DISTANCE.

250

300

350

400

KM

Fig. 10. Composite travel-time plots of the initial phases (crosses) and the later phases (octagons}. (a) P wave. (b) S wave. Reduced velocities are the same as Fig. 9.

motion of the Philippine Sea plate is clearly seen as the subcrustal seismic zone {Shiono, 1977; Mizoue, 1983; Yamazaki and Ooida, 1985). The later phases are consistently observed for the subcrustal earthquakes in a depth range of 30-60 km with few exceptions (Fukao et al., 1983; Hori et al., 1985). In the Kanto district, on the other hand, the hypocenters do not form a thin seismic zone but are dispersed above UBP. The dispersed hypocentral dis~bution has been considered to repre-

sent the subducting slab of the Philippine Sea plate (e.g., Noguchi, 1985; I&da, 1986, 1989). The X, and the X, phases are not observed for every event in this dispersed focal zone. Figure 12 shows the location of the hypocenters of the events analysed in the present study, where the upper surface of the Philippine Sea plate proposed by Ishida (1989) is indicated by a broken curve. The hypocenters are indicated by open circles when the later phases are observed and by crosses when

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367

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S. HORI

36.5

36.0

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(a) 139.0 E Fig. 12. Distribution

139.5 E of the upper mantle earthquakes

rectangular region of n-a’

140.0 E

of Fig. 12a. (c) Vertical section of the rectangular

indicate the events from which the later phase is observed parameters of which are well determined by the National

140.5 E

analysed in this study. (a) Epicentral

distribution.

(b) Vertical section of the

region of b-b’ of Fig. 12a. Open circles and crosses

and is not observed,

respectively.

Research Center for Disaster Prevention,

Other earthquakes,

the hypocentral

are also plotted as small dots. The

upper surface of the subducted Philippine Sea plate proposed by Ishida (1989) is indicated by a broken line.

not detected. Small dots indicate the hypocenters of earthquakes the locations of which are well determined by routine data processing of the National Research Center for Disaster Prevention. The cross-sections are taken along the direction of motion of the Philippine Sea plate relative to the Eurasian plate. As can be seen, the hypocenters are clearly separated into two groups: the earth-

quakes exhibiting the later phases are located only in the shallower part to the dispersed focal zone. Model of oceanic crust subduction We found the following characteristics through analyses of many seismograms for upper mantle earthquake occurring beneath the Kanto district:

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BY SUBDUCTING

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(1) A pro~nent pair of P and S later phases and X,) appears a few seconds after the (XP onsets of the initial P and S phases. Their amplitudes are much larger than those of the initial phases. (2) The X, phase can be regarded as a longitudinal wave from its particle motion. (3) The incident angle of the X, phase is slightly greater than that of the initial P phase. (4) The averaged value of (X, - S)/(X, - P) closely agrees with a typical value of V,/V, of the crustal or mantle rock. This observation suggests that the X, and X, phases travel along the same path as longitudinal and tangential waves, respectively. (5) The later phases are well observed at stations distributed in the south of the Tokai district and in the north of the Izu peninsula. Any significant later phases do not appear at stations in the north of the Kanto-Tokai district or in the south of the Izu peninsula. (6) The apparent velocities of the X, and the X, phases are too low in comparison with those normally expected for upper mantle earthquakes. They are comparable to the P and S velocities of the main layer of the oceanic crust or the lower continental crust. There is evidence that these apparent velocities directly reflect the seismic velocities in the source region. (7) An offset of arrival times is clearly seen at a distance of 150-170 km for each of the initial P and S phases. Both the travel-time plots, before and after the offset, give apparent velocities higher than the sub-Moho P and S velocities. In contrast, no significant offset is observed for the travel-time data of the X, or the X, phase. (8) The events related to the subducting motion of the Philippine Sea plate are not confined to the upper portion of the slab but extend over its entire thickness. However, the later phases are only observed for the events at the upper portion. Most of the characteristics of the later phases can be explained by the model proposed by Fukao et al. (1983) and Hori et al. (1985) for an interpretation of the later phases observed for the subcrustal earthquakes occurring beneath southwest Japan. Figure 13 shows a schematic illustration of the explanation. We assume that the oceanic crust

S. IiORt

Fig. 13. Schematic ilIustration explaining the initial and the later phases. Rays of the initial and the later phases are indicated by solid and broken lines, respectively.

remains gabbroic under the sub-Moho condition and that the relevant events are confined within the subducted portion of this untransformed oceanic crust. Gabbro-eclogite transformation is presumed to take place at greater depths. Typical rays of the initial and the later phases are indicated by solid and broken lines respectively. The initial phase at distant stations is interpreted as a direct wave through the mantle or a head wave propagating along the boundary between the mantle and the crustal portions of the oceanic lithosphere. Therefore, the initial phase has an apparent velocity almost equal to the seismic speed of the lithospheric mantle. The subducted portion of the oceanic crust forms a thin low-velocity layer in the mantle. Seismic waves radiated at shallow angles are trapped within this low-velocity layer and cannot escape from it until they reach the bottom of the continental crust where little velocity contrast exists across the contact. We would observe these waves as a prominent later phase with an apparent velocity comparable to the speed of the gabbroic oceanic crust or the lower continental crust. Hori et al. (1985) simulated two-dimensional ray tracing for the later phases of this kind. From their results, the travel time, T (s), of the most energetic branch of later arrivals is expressed as T=(A-lOO)/C+1, (62100)

(2)

SEISMIC

WAVES

GUIDED

BY SUBDUCTlNG

UNTRANSFORMED

OCEANIC

where A, C and t, are the epicentral distance (km), the apparent velocity (7 or 4 km s-l) and the intercept time (s) at a distance of 100 km. Hori et al. (1985) found the values of t, for the P and S later phases to be 16 and 29, respectively, for the subcrustal earthquakes beneath the Tokai and Kinki districts. The regression lines in Fig. 10 represented by eqn (2) show very similar values of t, = 17 for X, and t, = 30 for X,. This close agreement indicates that the later phases observed beneath the Tokai and Kinki districts and those beneath the Kant0 district are of the same origin and that they are quantitatively explained by the model illustrate in Fig. 13. The difference in the incident angles between the initial P and the X, phases are also consistent with the model. As schematically shown in Fig. 13, the seismic wave of the later arrival is radiated from the source at a shallower angle and is incident toward the Earth’s surface also at a shallower angle than that of the initial arrival. The above seismic wave guided by the lowvelocity layer can be observed in a limited region. The ray of the seismic wave should be able to reach the region through a direct contact of the subducted oceanic crust with the bottom of the continental crust. Figure 14 shows the vertical sections along the directions of arrows A, B and C of Fig. 8, where the illustrating of UBP and the subducted slab of the Philippine Sea plate are based on Ishida (1986) and Ishida (1989). The hatched area represents the mantle part of the Philippine Sea plate. In these sections, solid and open triangles represent the stations of groups I and II, respectively. As can be seen, the seismic wave trapped within the oceanic crust comes up to the Earth’s surface along the ray paths of A and B to be observed there, whereas it is never observed along the ray path of C. Hori et al. (1985) showed from numerical experiments that the guided wave consists of multibranched later phases. The most energetic branch is a phase generated by a ray smoothly traveling in the waveguide. Since the subducted portion of the Philippine Sea plate is inclined more steeply along the ray path of A than along that of B, the energetic phase would not arrive at distant stations in this case. This would give a qualitative reason why the later phase is not

371

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Fig. 14. Vertical section of the subducted Philippine Sea plate along the directions of the arrows of A, B and C of Fig. 8, where the hypooenters and the mantle part of the plate are indicated by dots and hatching, respectively. UBP and the subducted Pbilippine Sea plate are based on ishida (1986) and Ishida (1989). Solid lines with an arrow represent the ray of the seismic wave guided by the subducted portion of the oceanic crust. Solid and open triangles represent the stations of group I and group II, respectively.

observed in the southern part of the Izu peninsula. According to the counter map of the subducting slab of the Philippine Sea plate proposed by Ishida (1989) (Fig. 8), the later phase can be observed in the south of the Kant0 district. However, no stations of group I are located in this region (see Fig. 8). This observation may imply the discontinuity of the subducting slab of the Philippine Sea plate beneath the southern Kanto district as proposed by Ishida (1986). This problem should be investigated in a future study including a quantitative

372

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noted by Ef in Fig. 3), which is to be compared to that of E# (Fig. llb) occurring in a shallower part of region B. Although the source-receiver geometry of ES is almost the same as that of E4, the waveform is quite different: the initial phases with large amplitudes are not accompanied by systematic later phases. This difference in waveform is easily understood if the difference in precise location between E4 and E.5 is taken into account. Discussion

Fig. 15. P wave seismograms of the events E5 of Fig. 3. The onsets of the initial phases are indicated by solid circles.

examination such as a three-dimensional ray tracing simulation. In the model shown in Fig. 13, the hypocenter is presumed to be confined within the crustal layer of the subducting oceanic ~thosphere. This would be the case for some of the earthquakes exhibiting the later phases analysed in the present study. However, the hypocenter does not need to be located “within” the crustal layer but can be at its boundary with the overlying mantle. We consider that the earthquakes in region A are interplate earthquakes such as proposed by Kasahara (1980). As shown in Fig. 2b, a considerable activity along UBP is observed in the deeper part of region B. These earthquakes have been considered to be caused by a relative slip between the bottom of the Philippine Sea plate and the upper surface of the underthrusting Pacific plate (e.g. Ohtake and Kasahara, 1983; Kasahara, 19850. Figure If shows the record section of such an event (de-

and conclusions

The later phases that are similar to the X, and the X, phases are also observed in other regions along the northern boundary of the Philippine Sea plate. Figure 16 shows the regions where the events generating the later phases are reported to take place. Tanaka and Oda (1988) investigated the upper mantle earthquakes occurring around the Shikoku district (regions I and II) and found a distinct pair of P and S later phases. They obtained the seismic velocities in the source region as those comparable to the seismic velocities of a gabbroic rock, and concluded that the later phases are the seismic waves guided by the untransformed oceanic crust subducting into the mantle. Okura and Takeuchi (1989) simulated the later phases observed from the events occurring in region I by means of a 3-D Gaussian beam method and confirmed that the model of the oceanic crust subduction is quantitatively valid. It is thus established that the oceanic crust subducting into the mantle remains gabbroic along the entire subduction zone at the northern boundary of the Philippine Sea plate. As suggested by Hori et al. (1985), the light untransformed oceanic crust just below the continental crust is likely to be a source of gravity anomaly: a prominent negative Bouguer anomaly observed in region I (e.g. Hagiwara, 1967) is considered to be related to an overlap of the oceanic crust and the continental crust. The seismic coupling between the two plates along the Nankai-Suruga trough should also be discussed, taking into account positive buoyancy due to the light untransformed oceanic crust. We suggest that this is a feature common to any subduction zone, which has important implications for developing the dynamic theory of plate tectonics.

SEISMIC

WAVES

GUlDED

BY SUBDUCTING

UNTRANSFORMED

OCEANIC

373

CRUST

Pacific

Ocean

32.0

132.0

t

134.0 E

136.0

E

138.0

E

lQO.0 E

102.0-t

Fig. 16. Location of the untransformed oceanic crust inferred from the later phases. The hatched regions are where the seismic waves guided by the oceanic crust are observed: Z-Tanaka and Oda (1988) and Okura and Takeuchi (1989); ZZ-Tanaka and Oda (1988); ZZZ-Fukao et al. (1983) and Hori et al. (1985); IV--Hot+ et al. (1985); V-the present study. Contours of the subducted slab of the Philippine Sea plate are drawn at intervals of 10 km, based on Shiono (1977), Mizoue et al. (1983), Yamazaki and Ooida (1985) and Ishida (1989).

Fukao et al. (1983) and Hori et al. (1985) showed that the later phases cannot be observed for events deeper than 60 km. This feature is also reported by Tanaka and Oda (1988) and Okura and Takeuchi (1989). Two explanations for this observation are proposed by Hori et al. (1985): (1) gabbro-eclogite conversion starts from a depth of 60 km; and (2) earthquakes at greater depths occur outside the oceanic crust. According to the first explanation the S$,, conversation observed in a depth range of 60-300 km is not directly related to the gabbroic oceanic crust subduction. Helffrich et al. (1989) proposed a thermal and mineralogical model of the slab/mantle interface in aelation to seismic wave conversions. They interpreted the origin of the shallowest S& conversions and the low-velocity waveguide as the subducted portion of the oceanic crust in a metastable gabbroic phase. They suggested that the SCS,conversions at

depths of SO-150 km are caused by another lowvelocity layer in association with melt presence. The second explanation is also intriguing. In the present study we did not observe the later phases from the deeper events (see Fig. 12). In the Kanto district the earthquakes generating no later phases can be regarded as the events occurring in the mantle part of the Philippine Sea plate on the basis of their hypocentral distribution. Whether is as yet the untransformed oceanic crust extends down to depths greater than 60 km unresolved. The hypocentral distribution is quite different between the subcrustal earthquakes beneath southwest Japan and the upper mantle earthquakes beneath the Kanto district: the hypocenters of the former are well confined within a thin slab-like seismic plane but those of the latter are widely dispersed over the entire thickness of the subducting slab. The most significant factor

S. HORI

374

which may be related to this difference is an interaction between the P~~ppine Sea and the under&rusting Pacific plates beneath the Kanto district. Strong seismic coupling between the two plates has been inferred from high inter-plate seismic activities (e.g., region C of Fig. 2). This coupling may cause earthqu~e-generating stress over the entire thickness of the subducting slabs. The subduction of the cold Pacific plate is also likely to cool the overriding Philippine Sea plate, causing the latter to remain in a brittle regime. The dear offset of the initial P and S arrival times seen in Figs. 9 and 10 indicates that the velocity structure is laterally inhomogeneous in their propagation paths. This inhomogeneity does not extend to the propagation paths of the later phases, which do not show such offset. One may suppose that an offset-like rapid change of the Moho depth would cause the offset of the arrival times. However, this explanation is not consistent with the previous studies concerning the crustal structure of the Kanto-Tokai district (e.g. Ikaxni, 19’78; Ashiya et al., 1988). We consider that the inhomogeneity is possibly related to the anomalous mantle in association with the volcanic activity, since the offset apparently occurs across the volcanic front. The low-velocity zone of this kind is also inferred by a recent three dimensional inversion of the velocity structure of the KantoTokai district (I&da and Hasemi, 1988). A quantitative examination of this problem will be discussed in a separate paper (Hori et al., in prep.). We have examined the later phases observed from the upper mantle earthquakes beneath the Kanto district. Together with the earlier observations for the similar phases we have firmly established a model of the oceanic crust subduction. The following results should be noted: (1) The distinct P and S later phase observed from the upper mantle earthquakes in a depth range of 30-60 km are the seismic wave guided by the subducted portion of the oceanic crust. (2) Gabbro-eclogite tr~sformation of the oceanic crust does not take place, at least down to a depth of 60 km, along the entire subduction zone at the northern boundary of the Philippine Sea plate.

Acknowledgements

The author is very grateful to Y. Fukao for his valuable suggestions and critical review. Hypocentral parameters used in this study are based on the database compiled by the Analysing System for Precursorsof Ea~t~q~ake~ of the National Research Center of Disaster Prevention.

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UNTRANSFORMED

OCEANIC

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