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Suboceanic earthquake location and seismic structure in the Kanto district, central Japan
Earth and Planetary Science Letters 241 (2006) 789 – 803 www.elsevier.com/locate/epsl
Suboceanic earthquake location and seismic structure in the Kanto district, central Japan Zhi Wang*, Dapeng Zhao Geodynamics Research Center, Ehime University, Matsuyama 790-8577, Japan Received 1 August 2005; received in revised form 6 November 2005; accepted 7 November 2005 Available online 15 December 2005 Editor: S. King
Abstract We present a combined method, using sP depth-phase data and double-difference arrival times, to determine the precise hypocenter locations of earthquakes that occur under the Pacific Ocean outside of the area covered by the land-based seismic network. We assess the effectiveness of the combined method using a data set of P- and S-wave arrival times and sP depth phase from suboceanic earthquakes recorded by both land-based seismic stations and offshore seismic stations (OFS). The hypocenters of the offshore earthquakes relocated using the combined method are consistent with those determined using the standard location method and OFS data. The differences in the hypocenters relocated by the two methods are less than 4 km. We applied the method to the subduction region that underlies the Kanto district, central Japan, and located a large number of earthquakes that occurred beneath the Pacific Ocean. We then determined the detailed 3D seismic velocity structure by inverting a large number of arrival times of P- and S-waves and sP depth phase from the relocated earthquakes in the study region. High-velocity anomalies related to the cold subducting Pacific slab and low-velocity anomalies related to the hot mantle wedge are clearly imaged. Beneath active volcanoes, low-velocity zones are visible from the surface to a depth of 100 km, reflecting fluids released by dehydration of the subducting Pacific slab. Strong lateral heterogeneities are revealed on the upper boundary of the Pacific slab beneath the forearc region. The low-velocity areas under the offshore region are associated with low seismicity and weak interplate coupling. A lowvelocity layer is imaged along the upper boundary of the Philippine Sea slab in the northern part of Kanto district, which may reflect dehydration of the slab. Our tomographic images indicate that the overlaying Philippine Sea plate has effects on the spatial distribution of active volcanoes related to the subducting Pacific slab in the study region. D 2005 Elsevier B.V. All rights reserved. Keywords: hypocenter location; double-difference travel times; sP depth phase; seismic tomography
1. Introduction Seismic tomography has played an important role in understanding the crust and upper mantle structure and * Corresponding author. Tel.: +81 89 927 8258; fax: +81 89 927 8167. E-mail addresses: [email protected] (Z. Wang), [email protected] (D. Zhao). 0012-821X/$ - see front matter D 2005 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2005.11.015
seismotectonics of subduction zones. The conventional tomographic method, however, does not take advantage of recent developments in earthquake location techniques such as the double-difference method [1,2] and the use of sP depth phase [3–7]. The accuracy of earthquake hypocenter estimates is affected by several factors, including the available phases, arrival time accuracy, network geometry [8], and the hypocenter location method. The poor hypocenter locations for
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offshore earthquakes determined using only the first arrival times of P- and S-waves recorded by the landbased seismic stations impose a limitation on determining the velocity structure of suboceanic regions. Many researchers have attempted to locate earthquakes that occurred beneath the suboceanic regions of subduction zones where sparse seismic stations exist. Nishizawa et al. [9] used arrival times recorded by both land-based and ocean-bottom seismometers (OBS) to locate shallow earthquakes near the Japan Trench. Short-period and long-period teleseismic depth phases [10–13] have been used to locate earthquakes within the suboceanic region, however, teleseismic depth phases (pP, sP) are normally usable only for relatively large earthquakes that have enough energy to generate teleseismic waves. An sP depth phase was detected from seismograms of the suboceanic earthquakes recorded by land-based stations in Northeast (NE) Japan [3]. Their study showed that the sP depth phase could be observed from suboceanic events with epicentral distances of ~150 km. Theoretical sP–P times were calculated by using a 2D ray-tracing method to determine the focal depths of many small offshore seismic events beneath the forearc region of the NE Japan subduction zone [3–5]. Wang and Zhao [6] presented a method that enables the combined use of Pwave, S-wave and sP depth-phase arrival times to relocate the hypocenters of offshore earthquakes
under the Pacific Ocean. This method is considered to be an effective way to accurately relocate hypocenters for earthquakes that occurred outside the area of the seismic network. On the other hand, the doubledifference technique [1] is considered to be a useful method to locate earthquakes accurately under a seismic network. Recent studies have shown substantial improvements in the hypocentral locations determined with this method [1,2]. The double-difference location technique can be extended to determine the relative locations of earthquakes by using well relocated events known as master events (MEs) [1]. In this work, based on the tomographic method of Zhao et al. [14] and taking advantage of the new hypocenter location techniques [1,2], we present an efficient method of relocating earthquakes that occurred under the Pacific Ocean, outside of the area covered by the seismic network. This method makes use of the sP depth phase [6] together with the doubledifference method [1,2,7,15]. To demonstrate the effectiveness of the method, we applied it to a real data set including the sP depth phase and P- and S-wave arrival times from offshore seismic events recorded by both OFS under the Pacific Ocean and land-based stations in the Kanto district, Japan (Fig. 1). The earthquake hypocenters relocated by this study are consistent with those determined using the standard location method [14] and OFS data.
Fig. 1. Depth distribution of the descending Philippine Sea (PHS) plate (thick contour lines) and the Pacific plate (thin contour lines) beneath the Kanto district in central Japan (modified from [18]). The PHS plate descends from the Sagami trough to the northwest beneath the Kanto district [55]. The Pacific plate subducts to the WNW under the PHS plate [56]. Two arrows indicate the directions of plate motions. Open circles denote the large historic earthquakes (M N 6.0) which occurred during the period from 1923 to 1998 [49]. Black solid triangles represent the active volcanoes. The bold dashed line indicates the Sagami trough and Suruga trough. The shaded rectangle in the insert map indicates the present study region.
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2. Tectonic background of the study region The Kanto district is a distinct region characterized by the presence of three tectonic plates: the Pacific plate, the Philippine Sea (PHS) plate, and the Eurasian plate (Fig. 1). The PHS plate descends beneath the Eurasian plate along the Sagami and Suruga Troughs, while the Pacific plate is subducting to the WNW under the PHS and Eurasian plates along the Japan Trench (Fig. 1). During the last two decades many researchers have investigated the structure and tectonics of this region using various approaches [e.g., [16–24]]. Geomorphological studies suggested that the large interplate earthquakes have recurred along the Eurasian–PHS plate boundary at least in the last 1600 years [25]. Plate subduction and plate interaction are responsible for the large damaging earthquakes that occur in this region, such as the 1923 Kanto earthquake (M7.9) that caused significant damage and loss of human lives; such destructive earthquakes occur at shallow levels close to densely populated areas. The large earthquakes that occur under this region might be related to the motion of the subducting slab, however, few studies of the seismic velocity structure of suboceanic regions have been made due to difficulties in carrying out OBS observations in oceanic regions. In this work, we used a large number of arrival times of sP depth phase, and Pand S-waves generated from 6353 land-based and offshore earthquakes. The data were relocated by a combined method to determine a detailed 3D seismic structure under the Kanto district. The tomographic image confirms the results of previous studies as well as revealing new features that underlie the Kanto region.
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earthquake, then reflected at the ocean floor and immediately converted to P-wave, and finally reaching a land-based station (Fig. 2a). Its travel time is very sensitive to the focal depth of the suboceanic earthquake [3,4,6]. Thus, it can be used for relocating the hypocenters of suboceanic events. We found that the sP depth phase can be handpicked from about 13% of the offshore earthquakes (M N 3.0) recorded by the dense Hi-net seismic network, and that the epicentral distances of the observed sP depth phases range from 80 to 360 km in the study region. When calculating the ray path and travel time of the sP depth phase, a threedimensional perturbation algorithm is used to determine the reflected point (also known as the bounce point) in the suboceanic region [6]; the 3D ray-tracing algorithm of Zhao et al. [14] is then used to trace the ray segments from the hypocenter to the bounce point and to the station. We then use the sP depth phase together with the P- and S-wave arrival times to accurately relocate the suboceanic events (master events). We found that the depth of the ocean floor within the Pacific Ocean, which ranges from tens to thousands of
3. Methods In this study, we present a combined method of relocating offshore earthquakes that occur outside of the landbased seismic network. We first use sP depth-phase data to relocate the hypocenters of the selected master events [6], and then employ the double-difference method [1,15] to relocate new events that are very close to the master events; we selected those new events that were within 6 km of master events. We then applied the 3D tomographic method of Zhao et al. [14,26] to invert for the seismic velocity structures using sP depth phases, P- and S-wave arrival times from the relocated earthquakes. 3.1. Hypocenter location of the master events The sP depth phase is a remarkable converted later phase, being radiated as an S-wave from a suboceanic
Fig. 2. (a) Schematic diagram showing the ray path of the sP depth phase from the hypocenter (cross) to the station (triangle). (b) Example of 3-component seismograms showing the distinct sP depth phases between P- and S-phases observed at a Hi-net station (NGRH) from a suboceanic event. The hypocentral parameters are listed above the seismogram. The epicentral distance of the corresponding station (NGRH) is shown on the left of the seismogram. From left to right in the seismogram, three vertical lines show P-wave, sP depth phase and S-wave arrivals. Focal depth (in km) is determined by Hi-net with the routine location method.
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meters in the study area, has a great effect on both the travel time and the ray path of the sP depth phase. Therefore, we take into account the effects of seafloor topography when we determine the bounce points of the sP depth phases and the travel times and ray paths of P- and S-waves from the suboceanic earthquakes. 3.2. Extension of the double-difference technique According to Waldhauser and Ellsworth [1], the relative travel time (also called double-difference) is defined as obs i cal drijk ¼ tki tkj tk tkj ð1Þ where drkij is the difference between the observed and calculated differential travel times between two events i and j at the same station k. A general equation for the change in the hypocentral locations between two events i and j can be written as Btkj
Btki Dmi Dmj ¼ drijk Bm Bm
ð2Þ
where Dm i and Dm j are perturbation vectors to the hypocentral parameters for events i and j. Eqs. (1) and (2) are known as the double-difference earthquake location algorithm. Earthquakes located close to the master events (MEs) can then be located using the master events. This is done by splitting the relocation vector for the relocated MEs in Eq. (2) into two components: one, m ¯ , which represents a shift common to all relocated MEs (i.e., shift of the entire cluster relative to the new events), while the other, mV, represents individual shifts between the events. For this purpose, to link a new event, i, to a relocated ME, j, Eq. (2) can be rewritten as below when considering a ME. Bt j Btki ¯ ME Þj ¼ drijk Dmi k DðmVþ m Bm Bm
ð3Þ
or expressed as follows: Btki Bt i Bt i Dxi þ k Dyi þ k Dzi Bx By Bz Btkj Bt j Dð xVþ x¯ Þj þ k D yVþ y¯ j Bx By ! Bt j þ k Dð zVþ z¯ Þj þ Dsi DðsVþ s¯ Þj ¼ drijk Bz
ð4Þ
where x, y, z, s are the four hypocentral parameters of a ¯ new event; xV, yV, zV, sV and x¯, y¯, z¯, s are the hypocentral
parameters of a ME. To relocate a set of new events close to a cluster of MEs, Eq. (4) can be rewritten as 2
d1 6 6 0 6 1 6d 6 6 1 6d 6 6: : : 6 6 4 0 0 2
d2 : : : ::: d2 ::: 0 ::: 0 ::: ::: ::: 0 0 1
::: 3 2
0 d ðnþ1Þ d ðnþ1Þ 0 ::: 1 0 3
0 0 0 ðnþ2Þ
d ::: 0 1
::: ::: ::: ::: ::: ::: :::
3 0 7 d ðnþ1Þ 7 7 d ðnþ1Þ 7 7 7 d ðnþ2Þ 7 7 ::: 7 7 7 5 0 0
dr1 Dm 7 7 6 6 6 Dm2 7 6 dr2 7 7 7 6 6 6 : : : 7 6 dr3 7 7 7 6 6 7 7 6 6 d 6 DmVðnþ1Þ 7 ¼ 6 dr4 7 7 7 6 6 ð nþ2 Þ : : : 7 7 6 6 DmV 7 7 6 6 7 6 ::: 7 6 5 4 0 5 4 0 Dm ¯
ð5Þ
where 1, 2, : : : , n are for the new events, n + 1, n + 2, : : : , are for the MEs, each d represents the four partial derivatives for one event, Dm is the perturbation of the 4 hypocentral parameters of the new events, and DmV, Dm ¯ is the perturbation of the 4 hypocentral parameters for the individual shifts and common shifts of the MEs, respectively. dr is the residual vector [1]. If a master event can be accurately relocated using the sP depth phase, many other offshore earthquakes that are close to the master event can be relocated using Eq. (5). Hereafter we term the combined method the master event location (MEL) method. 4. Locating earthquakes off the Kanto district, central Japan To assess the effectiveness of the MEL method, we applied it to a data set including 6387 P-wave and 4132 S-wave arrival times, and 387 sP depth-phase data collected from 349 offshore earthquakes that occurred within the region 34.0–34.8 N and 139.0– 141.0 E (the box in Fig. 4), where 43 OFS are located in the Pacific Ocean. Fig. 3a–d shows the distribution of the 349 offshore earthquakes and the 43 OFS used in this analysis. The 349 offshore earthquakes, recorded by both land-based seismic stations and the 43 OFS, consist of 76 master events and 273 other events that are located within 6 km of master events. Thus, we can use this data set to demonstrate the accuracy of the hypocenters determined using the MEL method.
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Fig. 3. (a–b) 3D hypocentral distribution of the 76 master events (MEs). (a) Hypocenters of the MEs relocated by Hi-net (red dots) and by sP depth phase together with the P- and S-wave arrival times recorded by the land-based stations (blue stars). (b) Hypocenters of the MEs relocated using the standard location method [14] with P- and S-wave arrival times recorded by the offshore seismic stations (yellow triangles) compared with those determined by sP depth phase (blue stars). (c–d) 3D hypocentral distribution of the 273 new earthquakes that are located within 6 km of MEs. (c) Hypocenters relocated by Hi-net (red dots) and by the MEL method of the present study (blue stars). (d) Hypocenters relocated using the standard location method [14] with P- and S-wave arrival times recorded by the offshore seismic stations (yellow triangles) compared with those determined by the MEL method (blue stars). Solid squares denote the 43 offshore seismic stations used to record P- and S-wave arrival times.
We first relocated the 76 master events using the 387 sP depth phases together with 1278 P-wave and 916 Swave arrival times recorded by the land-based stations (Fig. 3a). For comparison, we relocated the 76 master events by applying the standard location method [14] to 1058 P-wave and 747 S-wave arrival times recorded by the 43 OFS (Fig. 3b). Then, the 276 earthquakes close to the 76 master events were relocated using the MEL method (Fig. 3c). For comparison, we also used the standard location method to relocate the 276 new events using 4051 P-wave and 2469 S-wave arrival times recorded by the 43 OFS (Fig. 3d). The relocated hypocenters are more tightly concentrated than those determined by Hi-net (Fig. 3). Table 1 shows examples of the hypocenters of the master events determined via the standard location method [14] from OFS data, and determined using sP depth-phase data [6]. The differences between the depths determined
using the above two methods is less than 4 km. The average difference in the depths is approximately 2 km. The average epicenter difference that is not shown in Table 1 is less than the depth differences. From Table 1, it is evident that the events relocated using the sP depth phases become shallower (by 3–10 km) than the hypocenters located by the Hi-net seismic network. Fig. 3a– b shows the distribution of hypocenters relocated by the two methods, which indicates that the master events can be relocated accurately with the sP depth phase (Fig. 3b). Table 2 lists the hypocenters located using the standard location method [14] with OFS data and by those located using the MEL method. The average difference of the focal depths between the two methods is about 2 km. This indicates that the hypocenters can be relocated accurately using the MEL method as well as using the standard method with OFS data (Fig. 3d). The distri-
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Table 1 Comparison of focal depths of 10 master events with different location techniques No. Events
Depth2 Depth3 Rdf1 Rdf2 Rdf3
Latitude Longitude Depth1 (Km) 1 34.47 140.26 2 34.54 140.23 3 34.53 139.80 4 34.50 139.88 5 34.63 140.16 6 34.60 139.99 7 34.68 140.24 8 34.34 140.01 9 34.47 139.83 10 34.71 140.27 Average value
69.86 81.25 94.23 89.54 60.51 98.85 58.81 87.00 11.69 62.85
61.57 72.32 89.50 83.21 53.33 91.45 51.83 78.61 11.06 59.93
(Km)
(Km) (Km) (Km)
59.53 71.71 86.18 86.62 54.61 91.46 53.18 82.39 11.16 59.55
8.29 10.33 2.03 8.93 9.54 0.60 4.73 8.05 3.32 6.33 2.92 3.41 7.18 5.90 1.29 7.40 7.39 0.01 6.98 5.63 1.35 8.39 4.61 3.79 0.63 0.53 0.10 2.92 3.30 0.38 6.18 5.82 1.63
Depth1: determined by Hi-net with a routine location procedure. Depth2: determined by using the standard location method [14] with P- and S- wave arrival time data recorded by OFS. Depth3: determined by using sP depth phase together with the first arrival times recorded by the land-based stations. Rdf1: Rdf1 = depth1 depth2. Rdf2: Rdf2 = depth1 depth3. Rdf3: Rdf3 = depth2 depth3.
bution of the hypocenters determined by the two methods is shown in Fig. 3d. The difference in the focal depth between the methods is less than 4 km. Thus, we consider that suboceanic earthquakes that occur outside the area of the land-based seismic network can be accurately relocated using the standard location method with OFS data or by using our MEL method. 5. Tomographic inversion To determine the 3D velocity structure in the study region, we used a large number of P-wave, S-wave and sP depth-phase arrivals from 6353 earthquakes recorded by the land-based seismic stations located in the Kanto district and OFS in the Pacific Ocean. Fig. 4 shows the 3D hypocentral distribution of the 6353 earthquakes, which consist of three data sets. Table 3 shows the details of the three data sets. The first data set (master events) includes P-wave, S-wave and sP depthphase arrivals from the suboceanic events that occurred from April 2004 to May 2005 (blue dots in Fig. 4). A total of 1158 sP depth-phase arrivals were handpicked from 299 suboceanic earthquakes that were identified from 5218 suboceanic earthquakes under the Pacific Ocean. Only very clear sP depth phases were handpicked from the 3-component seismograms (Fig. 2b). The picking accuracy of the sP depth phases, P-wave and S-wave arrivals are estimated to be ~0.05–0.15, ~0.03–0.1 and ~0.05–0.15 s, respectively. The second data set consists of the P-wave and S-wave arrival times
from 1129 suboceanic earthquakes (green dots in Fig. 4) that are located close to the master events in the first data set. The 1129 offshore events were selected from 7603 suboceanic earthquakes under the Pacific Ocean according to their locations relative to the master events; the distance between the earthquake and corresponding master event is required to be less than 6 km. The third data set includes P-wave and S-wave arrival times from 4925 earthquakes (crosses in Fig. 4) that occurred beneath the land area of Kanto district. As a result, a total of 6353 earthquakes were used for the tomographic inversion. The spatial distribution of the three groups of earthquakes is quite uniform in the study area (Fig. 4). The number of P- and S-wave arrival times and sP depth phases is 107 532, 77 409 and 1158, respectively. In this study we used 227 seismic stations that recorded the arrival time data of the 6353 events from June 2002 to May 2005 (Fig. 4). These seismic stations belong to the University of Tokyo, Kyoto University, the Japan Meteorological Agency (JMA), the High-Sensitivity Seismic Network (Hi-net), and the J-Array network. Fig. 5 shows the results of Vp and Vs checkerboard resolution tests (CRT) with a grid separation of 10–15 km in the horizontal direction and 10–20 km in depth. We assigned positive and negative velocity perturbations of F 3% to the grid nodes, and then calculated synthetic travel times for this model. The CRT results for P- and S-wave structures show good resolution beneath the forearc region at depths of 10–50 km (Figs. 5 and 6). At depths of 60–150 km, the resolution is good under the land area of Kanto district (Fig. 6). Table 2 Comparison of focal depths of 10 offshore events determined with different location methods No. Events
Depth2 Depth3 Rdf1 Rdf2 Rdf3
Latitude Longitude Depth1 (Km) 1 34.81 140.11 2 34.45 140.24 3 34.68 140.21 4 34.50 140.03 5 34.74 140.03 6 34.68 140.14 7 34.40 139.83 8 34.40 139.76 9 34.83 139.96 10 34.83 140.13 Average value
62.12 85.37 62.47 84.08 60.14 67.64 89.76 96.82 54.46 57.28
56.18 78.89 60.64 80.02 57.08 59.41 80.32 91.81 47.74 52.56
(Km)
(Km) (Km) (Km)
54.57 78.54 60.26 81.47 55.83 55.63 82.64 89.71 49.72 55.23
5.94 7.55 1.61 6.49 6.83 0.34 1.84 2.21 0.37 4.06 2.61 1.45 3.06 4.31 1.25 8.23 12.01 3.78 9.44 7.12 2.32 5.01 7.11 2.10 6.72 4.74 1.98 4.72 2.05 2.67 5.55 5.65 1.79
Depth1: determined by Hi-net. Depth2: determined by using the standard location method with P- and S-wave arrival time data recorded by OFS. Depth3: determined by using the MEL method. Rdf1, Rdf2, Rdf3: the same as those in Table 1.
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Fig. 4. 3D hypocentral distribution of the 6353 earthquakes analyzed in this study. Crosses represent the hypocenters of 4925 earthquakes that occurred under the land area of the Kanto district. Blue dots denote the 299 earthquakes (M N 3.0) (master events) selected from 5218 offshore events. We handpicked 1158 sP depth phases from the 299 master events. Green dots represent the 1129 events selected from 7603 subocanic earthquakes according to the distance (b6 km) from the corresponding ME. Red solid squares denote the 227 stations that recorded the sP depth phase, P- and S-wave arrival times from June 2002 to May 2005. The large rectangle denotes the region in Fig. 3 selected for testing the MEL method.
The final inverted results were obtained after 8 iterations. The hit count at the grid node is greater than 10 for the area where the velocity image is determined. The root-mean-square (RMS) travel time residual calculated from the hypocenters relocated by this study is reduced drastically from 0.543 to 0.203 s for P-waves and from 0.596 to 0.232 s for S-waves. Fig. 7 shows plan views of the Vp and Vs images at four depths, together with the distribution of background seismicity, large interplate earthquakes, active volcanoes, and active faults. Beneath the active volTable 3 Data sets used for tomographic inversion Content
Events
sP phase
P-wave
S-wave
Recorded stations
M JMA
MEa CMEb LEc Total
299 1129 4925 6353
1158
5061 19109 83362 107532
3887 13756 59766 77409
N35 N25 N25
N3.0 N2.0 N1.5
a
1158
Master events were relocated by using sP depth phase and P- and S-wave arrival times jointly. b Earthquakes that were within 6 km of master events were relocated by the MEL method of this study. c Earthquakes under the land area were relocated with the doubledifference method [1,7].
canoes, low-velocity zones are imaged clearly from the surface to a depth of 100 km. Fig. 8 shows four vertical cross-sections of the Vp and Vs images together with the distribution of active volcanoes and seismicity along the lines shown in the insert map. Low-velocity areas are visible under the offshore region at depths of 10–30 km (Fig. 7). The subducting Pacific slab is imaged clearly as a high Vp and Vs zone (Fig. 8), similar to the previous results [6,14,27– 30]. In the northern section of the study region, highvelocity bodies are imaged at depths of 35–65 km, which represent the subducting Philippine Sea plate (Figs. 7 and 8). 6. Discussion 6.1. Interpretations of the tomographic images Strong low-velocity anomalies are visible in the crust and uppermost mantle under the volcanic front of the Kanto district (Figs. 7 and 8). The CRT results for the Pand S-wave structures (Figs. 5 and 6) show good resolution at depths of 0–100 km; the low-velocity anomalies are therefore considered to be reliable features. Similar structures in the mantle wedge beneath the arc
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Fig. 5. Results of checkerboard resolution test for P-wave (a) and S-wave (b). The input amplitudes of the velocity perturbation are F 3%. The depth of each layer is shown below each map. The velocity perturbation scale is shown at the bottom.
volcanoes have been determined by many researchers in other subduction zones, such as the Aleutians [31], Alaska [32], Mediterranean [33], and Northeast Japan [6, 14], although the resolution scales are different from
one study to another. We consider that at depths of 80– 150 km, aqueous fluids are produced by progressive metamorphic dehydration reactions involving numerous hydrous minerals [34–38]. Such fluids might trigger
Z. Wang, D. Zhao / Earth and Planetary Science Letters 241 (2006) 789–803 Fig. 6. Vertical cross-sections showing results of the checkerboard resolution test of P-wave (left) and S-wave (right) along the lines shown in the insert map. Solid thick lines and gray solid triangles on the top of each section denote the land area of the Kanto district and active volcanoes, respectively. The velocity scale is shown on the right. 797
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Fig. 7. Plan views of P-wave (a) and S-wave (b) velocity images (in percent) at four depths determined by tomographic inversion. The depth of each layer is shown in the lower left corner of each map. Red color denotes low velocity while blue color represents high velocity. White stars indicate locations of the large historic earthquakes which occurred from 1923 to 1998 [49]. Crosses represent the locations of earthquakes along each profile. Active volcanoes are shown by red triangles. Curved solid lines at each depth indicate the locations of the active faults. The velocity perturbation scale and magnitude scale of the great earthquakes are shown at base of the figure.
Z. Wang, D. Zhao / Earth and Planetary Science Letters 241 (2006) 789–803 Fig. 8. Vertical cross-sections of P-wave (left) and S-wave (right) velocity images (in percent) along lines shown in the insert map. Red color represents low velocity while blue color denotes high velocity. Gray solid circles indicate background earthquakes along each section. Red triangles and bold horizontal lines on top of each section represent active volcanoes and the land area, respectively. The red star denotes the 1923 Kanto earthquake (M7.9) [57]. The solid line represents the upper boundary of the Pacific plate. The two dashed lines in (e–h) show the upper and lower boundaries of the PHS slab. The velocity perturbation scale and earthquake magnitude scale are shown on the right. 799
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partial melting under the active volcanoes and then reduce the seismic velocity of the surrounding material (Fig. 8). Many previous studies have provided evidence for low-velocity anomalies under the volcanic front in the study region (Fig. 8). High heat flow, high temperature and large geothermal gradient are found in the volcanic areas, which could also act to reduce the seismic velocity [16,39]. 3D seismic attenuation structures within the crust and upper mantle under the Kanto region [21,40] clearly demonstrate the existence of the inclined low-attenuation Pacific slab, and high-attenuation bodies in the mantle wedge under active volcanoes. The spatial distribution of the high-attenuation bodies coincides with that of the low-Vp and low-Vs anomalies identified in the present study beneath the volcanic front (Figs. 7 and 8). The low-velocity zones in the mantle wedge above the upper boundary of the subducting Pacific slab at depths of 30–50 km under the forearc region (Figs. 7 and 8) are interpreted as hydrated mantle peridotite rather than partial melting based on the following reasons: 1) The subducted oceanic lithosphere becomes dehydrated with increasing pressure and temperature, and thus the overlying mantle wedge is serpentinized [34]; 2) Serpentinization dramatically reduces the seismic velocity while increasing the Poisson’s ratio, magnetization and electrical conductivity of the forearc mantle [35]. Low-velocity anomalies in the mantle wedge are visible from our tomographic results (Fig. 8), while high Poisson’s ratio [41] and high-attenuation [40, 42] zones in the upper mantle are revealed in the northern portion of Kanto district; 3) The temperature in the low-velocity zones is suitable for serpentinization and much lower than the melting temperature of peridotite because of the low forearc temperature. Furukawa [43] estimated that the temperature is about 200 8C at a depth of 30 km beneath the Kanto district; serpentinite is stable at temperatures lower than about 600 8C [44]. Therefore, we conclude that the lowvelocity zones under the forearc region represent serpentinization of the mantle wedge (Fig. 8). The cold subducting Pacific slab and the PHS plate are imaged as high-velocity zones. The Pacific slab subducts beneath the PHS plate along the Japan Trench (Fig. 8e–h). Fig. 1 shows the contour lines of the depths of the Pacific slab and PHS plate estimated from studies of seismicity patterns, velocity structure and focal mechanisms [18]; these data are consistent with the spatial distribution of the plates imaged by the present study. Widespread low-velocity anomalies are visible in the suboceanic region (with low-velocity perturbations of Vp and Vs of 2–6% and 3–6%, respectively), along
the upper boundary of the Pacific slab (Fig. 8). Features of the low-velocity anomalies are similar to those described in the previous studies [4–6]. The CRT results for the P- and S-wave structures (Figs. 5 and 6) show good resolution under the offshore region at depths of 10 to 50 km, indicating that the low-velocity anomalies should be considered reliable features. Under the suboceanic region, the subducting sediments and oceanic crust contain free water in pore spaces and bound water within hydrous minerals [6]. At shallow depths, free water is expelled by the compaction of subducted sediments and collapse of porosity in the upper oceanic crust, which acts to decrease the seismic velocity of the crust and the uppermost mantle (Figs. 7 and 8). Previous studies indicated that fluids exist widely in the crust and uppermost mantle in the forearc regions of the subduction zone [4,6,34]. We consider that the fluids represent pore water contained in the subducting oceanic crust and sediments released due to the increasing temperature and pressure, and then reduce the seismic velocity there. A low-velocity layer at depths of 20–35 km is imaged clearly along the upper boundary of the PHS plate in the northern part of the Kanto district, indicating that the subducting PHS plate is located beneath the Eurasian plate (Fig. 8e–h). Similar features of low-velocity anomalies were observed in the previous studies of the present region, although the resolution of the previous results was not high [19,41]. A layer of high Poisson’s ratio at depths of 20–35 km has been previously imaged in the lower crust of the study region [41], and interpreted as serpentinized mantle. We consider that the low-velocity layer may reflect dehydration reactions within the PHS plate. Many studies have revealed anomalously high heat flow and strong seismic-wave attenuation in the study region [16,17,21], which provide additional evidence in support of this interpretation. A moderately high heat flow zone exists at the depth levels of the lower crust, which represents a soft ductile zone sandwiched by two brittle layers in the study region [16,17]. Their results coincide with the seismic velocity structures [19], indicating that two layers of concentrated seismicity are associated with active subduction of the oceanic lithosphere: one is in the upper crust at depths of 10–20 km, while the other is in the uppermost mantle at depths of 35–60 km. These results are in good agreement with our tomographic results (Fig. 8); both studies identify a low-velocity layer at depths of 20–35 km sandwiched by two high-velocity layers. A high-attenuation layer was imaged at depths of 25–40 km in the northern section of the study region [21]. We therefore conclude that the low-velocity layer
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may result from fluids released by dehydration reactions within the descending PHS plate; the fluids act to reduce the seismic velocity of the surrounding material (Fig. 8). 6.2. Volcanoes, great earthquakes, and the subducting slabs High-velocity zones are revealed at depths of 25–60 km in the northern portion of the study region, which might reflect the existence of the subducting PHS plate (Figs. 7 and 8e–h). We find that the subducting Philippine Sea plate has affected the spatial distribution of the active volcanoes related to the subducting Pacific slab in the Northern portion of the study region (Fig. 8e–h). In general, low-velocity zones under active volcanoes are sub-parallel to the descending slab in the associated subduction zone [e.g. 6,14]. However, such a trend is ambiguous in the Kanto district due to the fact that the PHS plate overlies the Pacific slab. Iwamori [45] suggested that the influence of the subducted PHS plate on magmatism is to shift the dehydration reactions to a greater depth along the Pacific slab. Low-velocity anomalies are clearly imaged at depths of 80–110 km between the PHS plate and the Pacific slab (Fig. 8e–h), which perhaps reflect dehydration processes within the subducting Pacific slab. The low-velocity material is not parallel to the Pacific slab but deflect towards the backarc due to the overlying PHS plate (Fig. 8e–h). The low-velocity material at the depths of 80–110 km imaged in the cross-sections of E–F and G–H probably represents the effects of volatiles, but not indicating melts due to the cold PHS slab (Fig. 8e–h). Our tomographic results clearly show that in the northern portion, active volcanoes are absent due to the influence of the descending PHS plate; in the middle region, the volcanic front is deflected towards the backarc because of the partial overlapping of the PHS plate; while south of Izu, the active volcanoes are located 120–140 km above the Pacific slab because the PHS plate has not yet subducted there (Fig. 8). These observations indicate that the fluids supplied by the subducting Pacific slab cannot rise to the crust due to the existence of the PHS plate beneath the northern portion of the Kanto district. Thus, we consider that the PHS plate existing between the Eurasian and Pacific plates has two possible effects on the arc magmatism: one is to shift the dehydration reactions to a greater depth along the Pacific slab [45]; the other is to act as a barrier to the upwelling of fluids released from dehydration reactions within the subducting Pacific slab. Further detailed studies with geological, geophysical and geochemical approaches would
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provide us with a better understanding of the relations between the volcanism, the subducting PHS plate and the Pacific slab in the study region. Many large earthquakes occur along the boundaries between the subducting Pacific plate, the Eurasian plate and the PHS plate in the subduction zone. The high seismic risk of the study region has been pointed out by extensive geological and geophysical studies over a long time [18]. The plate subduction and plate interaction are responsible for the large damaging earthquakes that occurred in this region. The strength of interplate coupling between the subducting PHS plate and the Eurasian plate was investigated and the plate boundary was estimated to be located at depths of 10–25 km under this region [46–48]. We found that seismic activity along the boundary between the PHS and Eurasian plates is low compared with that within the overriding Eurasian plate and in the subducting PHS plate (Fig. 8e–h). The 1923 Kanto earthquake (M7.9) occurred on the Eurasian–PHS plate boundary (Fig. 8e, f), where the interplate seismic coupling of the two plates is expected to be strong. Large earthquakes (M z 6.0) that occurred during the period of 1923 to 1998 [49] are shown in white stars in Fig. 7. We divided these earthquakes into two groups according to their hypocentral distribution. One group (including the 1923 Kanto earthquake) occurred along the boundary between the PHS and Eurasian plates, while the second group occurred along the upper boundary of the subducting Pacific slab (Fig. 7). Most of the great earthquakes, such as the 1909 Bousooki (M7.5) and the 1923 Kanto (M7.9) earthquakes, occurred around the low-velocity zones (Fig. 7); this might indicate the existence of weakly coupled sections on the subducting slab boundary. For the subducting Pacific slab, the down-dip limit of the great thrust earthquakes generally corresponds to the intersection of the thrust zone with the forearc mantle [50,51], which might be explained by the distribution of aseismic hydrous minerals such as serpentinite [34,52,53]. Our results show that large earthquakes that occurred along the PHS–Eurasian and the upper Pacific boundaries are located in areas with seismic velocities higher than that of surrounding areas (Fig. 7), which might reflect strongly coupled asperities. Previous studies indicated that fluids exist widely in the crust and uppermost mantle in the forearc regions of the subduction zone [4,34], which is consistent with our tomographic images. We conclude that the spatial distribution of large earthquakes is possibly controlled by lateral heterogeneities, including strongly coupled sections (asperities) and weakly coupled or decoupled sections along the upper slab boundary.
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7. Conclusions
References
In this study we present an efficient combined method for accurately locating a large number of earthquakes that occurred outside the area covered by the seismic network under the Kanto district, Japan. The method makes use of sP depth-phase data together with the double-difference arrival time data. The hypocenters relocated by this method are consistent with those determined from OFS data; the spatial differences in the hypocenters relocated by the two methods are generally less than 2 km. This new approach to hypocenter location can be applied to other offshore regions of subduction zones where few seismic stations exist to record the numerous suboceanic earthquakes. Strong low-velocity anomalies beneath the active volcanoes are clearly imaged, and might reflect aqueous fluids released from the subducting Pacific slab interacting with hot upwelling flow in the mantle wedge, leading to partial melting. A low-velocity seismic layer is visible along the upper boundary of the PHS plate, which may be associated with dehydration reactions within the subducting slab. The cold subducting Pacific slab shows high-velocity anomalies that are ~2–6% higher than that of the normal mantle. We find that the PHS plate existing between the Eurasian plate and the subducting Pacific slab has influenced the spatial distribution of arc volcanoes in northern part of the study region. Widespread low-velocity anomalies are observed on the upper boundary of the subducting Pacific slab under the forarc region, which might reflect weakly coupled or decoupled sections of the plate interface related to serpentinization of the forearc mantle associated with slab dehydration. On the upper boundary of the Pacific slab, under the offshore region, strong lateral heterogeneities are visible, which show a strong correlation with the spatial distribution of the great earthquakes. These large earthquakes occurred outside the low-velocity zones which may be caused by aqueous fluids released from the subducted oceanic crust.
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Acknowledgment We thank Profs. S. King and T. Seno and an anonymous reviewer for their thoughtful comments and suggestions. We thank Hi-net, Japan Meteorological Agency and J-Array data centers for providing us P-wave and S-wave arrival time and waveform data via Internet. All figures in this work are made by using GMT [54]. This work was partially supported by research grants (Kiban-B 11440134, and Kiban-A 17204037) from Japan Society for the Promotion of Science to D. Zhao.
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Suboceanic earthquake location and seismic structure in the Kanto district, central Japan Zhi Wang*, Dapeng Zhao Geodynamics Research Center, Ehime University, Matsuyama 790-8577, Japan Received 1 August 2005; received in revised form 6 November 2005; accepted 7 November 2005 Available online 15 December 2005 Editor: S. King
Abstract We present a combined method, using sP depth-phase data and double-difference arrival times, to determine the precise hypocenter locations of earthquakes that occur under the Pacific Ocean outside of the area covered by the land-based seismic network. We assess the effectiveness of the combined method using a data set of P- and S-wave arrival times and sP depth phase from suboceanic earthquakes recorded by both land-based seismic stations and offshore seismic stations (OFS). The hypocenters of the offshore earthquakes relocated using the combined method are consistent with those determined using the standard location method and OFS data. The differences in the hypocenters relocated by the two methods are less than 4 km. We applied the method to the subduction region that underlies the Kanto district, central Japan, and located a large number of earthquakes that occurred beneath the Pacific Ocean. We then determined the detailed 3D seismic velocity structure by inverting a large number of arrival times of P- and S-waves and sP depth phase from the relocated earthquakes in the study region. High-velocity anomalies related to the cold subducting Pacific slab and low-velocity anomalies related to the hot mantle wedge are clearly imaged. Beneath active volcanoes, low-velocity zones are visible from the surface to a depth of 100 km, reflecting fluids released by dehydration of the subducting Pacific slab. Strong lateral heterogeneities are revealed on the upper boundary of the Pacific slab beneath the forearc region. The low-velocity areas under the offshore region are associated with low seismicity and weak interplate coupling. A lowvelocity layer is imaged along the upper boundary of the Philippine Sea slab in the northern part of Kanto district, which may reflect dehydration of the slab. Our tomographic images indicate that the overlaying Philippine Sea plate has effects on the spatial distribution of active volcanoes related to the subducting Pacific slab in the study region. D 2005 Elsevier B.V. All rights reserved. Keywords: hypocenter location; double-difference travel times; sP depth phase; seismic tomography
1. Introduction Seismic tomography has played an important role in understanding the crust and upper mantle structure and * Corresponding author. Tel.: +81 89 927 8258; fax: +81 89 927 8167. E-mail addresses: [email protected] (Z. Wang), [email protected] (D. Zhao). 0012-821X/$ - see front matter D 2005 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2005.11.015
seismotectonics of subduction zones. The conventional tomographic method, however, does not take advantage of recent developments in earthquake location techniques such as the double-difference method [1,2] and the use of sP depth phase [3–7]. The accuracy of earthquake hypocenter estimates is affected by several factors, including the available phases, arrival time accuracy, network geometry [8], and the hypocenter location method. The poor hypocenter locations for
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offshore earthquakes determined using only the first arrival times of P- and S-waves recorded by the landbased seismic stations impose a limitation on determining the velocity structure of suboceanic regions. Many researchers have attempted to locate earthquakes that occurred beneath the suboceanic regions of subduction zones where sparse seismic stations exist. Nishizawa et al. [9] used arrival times recorded by both land-based and ocean-bottom seismometers (OBS) to locate shallow earthquakes near the Japan Trench. Short-period and long-period teleseismic depth phases [10–13] have been used to locate earthquakes within the suboceanic region, however, teleseismic depth phases (pP, sP) are normally usable only for relatively large earthquakes that have enough energy to generate teleseismic waves. An sP depth phase was detected from seismograms of the suboceanic earthquakes recorded by land-based stations in Northeast (NE) Japan [3]. Their study showed that the sP depth phase could be observed from suboceanic events with epicentral distances of ~150 km. Theoretical sP–P times were calculated by using a 2D ray-tracing method to determine the focal depths of many small offshore seismic events beneath the forearc region of the NE Japan subduction zone [3–5]. Wang and Zhao [6] presented a method that enables the combined use of Pwave, S-wave and sP depth-phase arrival times to relocate the hypocenters of offshore earthquakes
under the Pacific Ocean. This method is considered to be an effective way to accurately relocate hypocenters for earthquakes that occurred outside the area of the seismic network. On the other hand, the doubledifference technique [1] is considered to be a useful method to locate earthquakes accurately under a seismic network. Recent studies have shown substantial improvements in the hypocentral locations determined with this method [1,2]. The double-difference location technique can be extended to determine the relative locations of earthquakes by using well relocated events known as master events (MEs) [1]. In this work, based on the tomographic method of Zhao et al. [14] and taking advantage of the new hypocenter location techniques [1,2], we present an efficient method of relocating earthquakes that occurred under the Pacific Ocean, outside of the area covered by the seismic network. This method makes use of the sP depth phase [6] together with the doubledifference method [1,2,7,15]. To demonstrate the effectiveness of the method, we applied it to a real data set including the sP depth phase and P- and S-wave arrival times from offshore seismic events recorded by both OFS under the Pacific Ocean and land-based stations in the Kanto district, Japan (Fig. 1). The earthquake hypocenters relocated by this study are consistent with those determined using the standard location method [14] and OFS data.
Fig. 1. Depth distribution of the descending Philippine Sea (PHS) plate (thick contour lines) and the Pacific plate (thin contour lines) beneath the Kanto district in central Japan (modified from [18]). The PHS plate descends from the Sagami trough to the northwest beneath the Kanto district [55]. The Pacific plate subducts to the WNW under the PHS plate [56]. Two arrows indicate the directions of plate motions. Open circles denote the large historic earthquakes (M N 6.0) which occurred during the period from 1923 to 1998 [49]. Black solid triangles represent the active volcanoes. The bold dashed line indicates the Sagami trough and Suruga trough. The shaded rectangle in the insert map indicates the present study region.
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2. Tectonic background of the study region The Kanto district is a distinct region characterized by the presence of three tectonic plates: the Pacific plate, the Philippine Sea (PHS) plate, and the Eurasian plate (Fig. 1). The PHS plate descends beneath the Eurasian plate along the Sagami and Suruga Troughs, while the Pacific plate is subducting to the WNW under the PHS and Eurasian plates along the Japan Trench (Fig. 1). During the last two decades many researchers have investigated the structure and tectonics of this region using various approaches [e.g., [16–24]]. Geomorphological studies suggested that the large interplate earthquakes have recurred along the Eurasian–PHS plate boundary at least in the last 1600 years [25]. Plate subduction and plate interaction are responsible for the large damaging earthquakes that occur in this region, such as the 1923 Kanto earthquake (M7.9) that caused significant damage and loss of human lives; such destructive earthquakes occur at shallow levels close to densely populated areas. The large earthquakes that occur under this region might be related to the motion of the subducting slab, however, few studies of the seismic velocity structure of suboceanic regions have been made due to difficulties in carrying out OBS observations in oceanic regions. In this work, we used a large number of arrival times of sP depth phase, and Pand S-waves generated from 6353 land-based and offshore earthquakes. The data were relocated by a combined method to determine a detailed 3D seismic structure under the Kanto district. The tomographic image confirms the results of previous studies as well as revealing new features that underlie the Kanto region.
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earthquake, then reflected at the ocean floor and immediately converted to P-wave, and finally reaching a land-based station (Fig. 2a). Its travel time is very sensitive to the focal depth of the suboceanic earthquake [3,4,6]. Thus, it can be used for relocating the hypocenters of suboceanic events. We found that the sP depth phase can be handpicked from about 13% of the offshore earthquakes (M N 3.0) recorded by the dense Hi-net seismic network, and that the epicentral distances of the observed sP depth phases range from 80 to 360 km in the study region. When calculating the ray path and travel time of the sP depth phase, a threedimensional perturbation algorithm is used to determine the reflected point (also known as the bounce point) in the suboceanic region [6]; the 3D ray-tracing algorithm of Zhao et al. [14] is then used to trace the ray segments from the hypocenter to the bounce point and to the station. We then use the sP depth phase together with the P- and S-wave arrival times to accurately relocate the suboceanic events (master events). We found that the depth of the ocean floor within the Pacific Ocean, which ranges from tens to thousands of
3. Methods In this study, we present a combined method of relocating offshore earthquakes that occur outside of the landbased seismic network. We first use sP depth-phase data to relocate the hypocenters of the selected master events [6], and then employ the double-difference method [1,15] to relocate new events that are very close to the master events; we selected those new events that were within 6 km of master events. We then applied the 3D tomographic method of Zhao et al. [14,26] to invert for the seismic velocity structures using sP depth phases, P- and S-wave arrival times from the relocated earthquakes. 3.1. Hypocenter location of the master events The sP depth phase is a remarkable converted later phase, being radiated as an S-wave from a suboceanic
Fig. 2. (a) Schematic diagram showing the ray path of the sP depth phase from the hypocenter (cross) to the station (triangle). (b) Example of 3-component seismograms showing the distinct sP depth phases between P- and S-phases observed at a Hi-net station (NGRH) from a suboceanic event. The hypocentral parameters are listed above the seismogram. The epicentral distance of the corresponding station (NGRH) is shown on the left of the seismogram. From left to right in the seismogram, three vertical lines show P-wave, sP depth phase and S-wave arrivals. Focal depth (in km) is determined by Hi-net with the routine location method.
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meters in the study area, has a great effect on both the travel time and the ray path of the sP depth phase. Therefore, we take into account the effects of seafloor topography when we determine the bounce points of the sP depth phases and the travel times and ray paths of P- and S-waves from the suboceanic earthquakes. 3.2. Extension of the double-difference technique According to Waldhauser and Ellsworth [1], the relative travel time (also called double-difference) is defined as obs i cal drijk ¼ tki tkj tk tkj ð1Þ where drkij is the difference between the observed and calculated differential travel times between two events i and j at the same station k. A general equation for the change in the hypocentral locations between two events i and j can be written as Btkj
Btki Dmi Dmj ¼ drijk Bm Bm
ð2Þ
where Dm i and Dm j are perturbation vectors to the hypocentral parameters for events i and j. Eqs. (1) and (2) are known as the double-difference earthquake location algorithm. Earthquakes located close to the master events (MEs) can then be located using the master events. This is done by splitting the relocation vector for the relocated MEs in Eq. (2) into two components: one, m ¯ , which represents a shift common to all relocated MEs (i.e., shift of the entire cluster relative to the new events), while the other, mV, represents individual shifts between the events. For this purpose, to link a new event, i, to a relocated ME, j, Eq. (2) can be rewritten as below when considering a ME. Bt j Btki ¯ ME Þj ¼ drijk Dmi k DðmVþ m Bm Bm
ð3Þ
or expressed as follows: Btki Bt i Bt i Dxi þ k Dyi þ k Dzi Bx By Bz Btkj Bt j Dð xVþ x¯ Þj þ k D yVþ y¯ j Bx By ! Bt j þ k Dð zVþ z¯ Þj þ Dsi DðsVþ s¯ Þj ¼ drijk Bz
ð4Þ
where x, y, z, s are the four hypocentral parameters of a ¯ new event; xV, yV, zV, sV and x¯, y¯, z¯, s are the hypocentral
parameters of a ME. To relocate a set of new events close to a cluster of MEs, Eq. (4) can be rewritten as 2
d1 6 6 0 6 1 6d 6 6 1 6d 6 6: : : 6 6 4 0 0 2
d2 : : : ::: d2 ::: 0 ::: 0 ::: ::: ::: 0 0 1
::: 3 2
0 d ðnþ1Þ d ðnþ1Þ 0 ::: 1 0 3
0 0 0 ðnþ2Þ
d ::: 0 1
::: ::: ::: ::: ::: ::: :::
3 0 7 d ðnþ1Þ 7 7 d ðnþ1Þ 7 7 7 d ðnþ2Þ 7 7 ::: 7 7 7 5 0 0
dr1 Dm 7 7 6 6 6 Dm2 7 6 dr2 7 7 7 6 6 6 : : : 7 6 dr3 7 7 7 6 6 7 7 6 6 d 6 DmVðnþ1Þ 7 ¼ 6 dr4 7 7 7 6 6 ð nþ2 Þ : : : 7 7 6 6 DmV 7 7 6 6 7 6 ::: 7 6 5 4 0 5 4 0 Dm ¯
ð5Þ
where 1, 2, : : : , n are for the new events, n + 1, n + 2, : : : , are for the MEs, each d represents the four partial derivatives for one event, Dm is the perturbation of the 4 hypocentral parameters of the new events, and DmV, Dm ¯ is the perturbation of the 4 hypocentral parameters for the individual shifts and common shifts of the MEs, respectively. dr is the residual vector [1]. If a master event can be accurately relocated using the sP depth phase, many other offshore earthquakes that are close to the master event can be relocated using Eq. (5). Hereafter we term the combined method the master event location (MEL) method. 4. Locating earthquakes off the Kanto district, central Japan To assess the effectiveness of the MEL method, we applied it to a data set including 6387 P-wave and 4132 S-wave arrival times, and 387 sP depth-phase data collected from 349 offshore earthquakes that occurred within the region 34.0–34.8 N and 139.0– 141.0 E (the box in Fig. 4), where 43 OFS are located in the Pacific Ocean. Fig. 3a–d shows the distribution of the 349 offshore earthquakes and the 43 OFS used in this analysis. The 349 offshore earthquakes, recorded by both land-based seismic stations and the 43 OFS, consist of 76 master events and 273 other events that are located within 6 km of master events. Thus, we can use this data set to demonstrate the accuracy of the hypocenters determined using the MEL method.
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Fig. 3. (a–b) 3D hypocentral distribution of the 76 master events (MEs). (a) Hypocenters of the MEs relocated by Hi-net (red dots) and by sP depth phase together with the P- and S-wave arrival times recorded by the land-based stations (blue stars). (b) Hypocenters of the MEs relocated using the standard location method [14] with P- and S-wave arrival times recorded by the offshore seismic stations (yellow triangles) compared with those determined by sP depth phase (blue stars). (c–d) 3D hypocentral distribution of the 273 new earthquakes that are located within 6 km of MEs. (c) Hypocenters relocated by Hi-net (red dots) and by the MEL method of the present study (blue stars). (d) Hypocenters relocated using the standard location method [14] with P- and S-wave arrival times recorded by the offshore seismic stations (yellow triangles) compared with those determined by the MEL method (blue stars). Solid squares denote the 43 offshore seismic stations used to record P- and S-wave arrival times.
We first relocated the 76 master events using the 387 sP depth phases together with 1278 P-wave and 916 Swave arrival times recorded by the land-based stations (Fig. 3a). For comparison, we relocated the 76 master events by applying the standard location method [14] to 1058 P-wave and 747 S-wave arrival times recorded by the 43 OFS (Fig. 3b). Then, the 276 earthquakes close to the 76 master events were relocated using the MEL method (Fig. 3c). For comparison, we also used the standard location method to relocate the 276 new events using 4051 P-wave and 2469 S-wave arrival times recorded by the 43 OFS (Fig. 3d). The relocated hypocenters are more tightly concentrated than those determined by Hi-net (Fig. 3). Table 1 shows examples of the hypocenters of the master events determined via the standard location method [14] from OFS data, and determined using sP depth-phase data [6]. The differences between the depths determined
using the above two methods is less than 4 km. The average difference in the depths is approximately 2 km. The average epicenter difference that is not shown in Table 1 is less than the depth differences. From Table 1, it is evident that the events relocated using the sP depth phases become shallower (by 3–10 km) than the hypocenters located by the Hi-net seismic network. Fig. 3a– b shows the distribution of hypocenters relocated by the two methods, which indicates that the master events can be relocated accurately with the sP depth phase (Fig. 3b). Table 2 lists the hypocenters located using the standard location method [14] with OFS data and by those located using the MEL method. The average difference of the focal depths between the two methods is about 2 km. This indicates that the hypocenters can be relocated accurately using the MEL method as well as using the standard method with OFS data (Fig. 3d). The distri-
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Table 1 Comparison of focal depths of 10 master events with different location techniques No. Events
Depth2 Depth3 Rdf1 Rdf2 Rdf3
Latitude Longitude Depth1 (Km) 1 34.47 140.26 2 34.54 140.23 3 34.53 139.80 4 34.50 139.88 5 34.63 140.16 6 34.60 139.99 7 34.68 140.24 8 34.34 140.01 9 34.47 139.83 10 34.71 140.27 Average value
69.86 81.25 94.23 89.54 60.51 98.85 58.81 87.00 11.69 62.85
61.57 72.32 89.50 83.21 53.33 91.45 51.83 78.61 11.06 59.93
(Km)
(Km) (Km) (Km)
59.53 71.71 86.18 86.62 54.61 91.46 53.18 82.39 11.16 59.55
8.29 10.33 2.03 8.93 9.54 0.60 4.73 8.05 3.32 6.33 2.92 3.41 7.18 5.90 1.29 7.40 7.39 0.01 6.98 5.63 1.35 8.39 4.61 3.79 0.63 0.53 0.10 2.92 3.30 0.38 6.18 5.82 1.63
Depth1: determined by Hi-net with a routine location procedure. Depth2: determined by using the standard location method [14] with P- and S- wave arrival time data recorded by OFS. Depth3: determined by using sP depth phase together with the first arrival times recorded by the land-based stations. Rdf1: Rdf1 = depth1 depth2. Rdf2: Rdf2 = depth1 depth3. Rdf3: Rdf3 = depth2 depth3.
bution of the hypocenters determined by the two methods is shown in Fig. 3d. The difference in the focal depth between the methods is less than 4 km. Thus, we consider that suboceanic earthquakes that occur outside the area of the land-based seismic network can be accurately relocated using the standard location method with OFS data or by using our MEL method. 5. Tomographic inversion To determine the 3D velocity structure in the study region, we used a large number of P-wave, S-wave and sP depth-phase arrivals from 6353 earthquakes recorded by the land-based seismic stations located in the Kanto district and OFS in the Pacific Ocean. Fig. 4 shows the 3D hypocentral distribution of the 6353 earthquakes, which consist of three data sets. Table 3 shows the details of the three data sets. The first data set (master events) includes P-wave, S-wave and sP depthphase arrivals from the suboceanic events that occurred from April 2004 to May 2005 (blue dots in Fig. 4). A total of 1158 sP depth-phase arrivals were handpicked from 299 suboceanic earthquakes that were identified from 5218 suboceanic earthquakes under the Pacific Ocean. Only very clear sP depth phases were handpicked from the 3-component seismograms (Fig. 2b). The picking accuracy of the sP depth phases, P-wave and S-wave arrivals are estimated to be ~0.05–0.15, ~0.03–0.1 and ~0.05–0.15 s, respectively. The second data set consists of the P-wave and S-wave arrival times
from 1129 suboceanic earthquakes (green dots in Fig. 4) that are located close to the master events in the first data set. The 1129 offshore events were selected from 7603 suboceanic earthquakes under the Pacific Ocean according to their locations relative to the master events; the distance between the earthquake and corresponding master event is required to be less than 6 km. The third data set includes P-wave and S-wave arrival times from 4925 earthquakes (crosses in Fig. 4) that occurred beneath the land area of Kanto district. As a result, a total of 6353 earthquakes were used for the tomographic inversion. The spatial distribution of the three groups of earthquakes is quite uniform in the study area (Fig. 4). The number of P- and S-wave arrival times and sP depth phases is 107 532, 77 409 and 1158, respectively. In this study we used 227 seismic stations that recorded the arrival time data of the 6353 events from June 2002 to May 2005 (Fig. 4). These seismic stations belong to the University of Tokyo, Kyoto University, the Japan Meteorological Agency (JMA), the High-Sensitivity Seismic Network (Hi-net), and the J-Array network. Fig. 5 shows the results of Vp and Vs checkerboard resolution tests (CRT) with a grid separation of 10–15 km in the horizontal direction and 10–20 km in depth. We assigned positive and negative velocity perturbations of F 3% to the grid nodes, and then calculated synthetic travel times for this model. The CRT results for P- and S-wave structures show good resolution beneath the forearc region at depths of 10–50 km (Figs. 5 and 6). At depths of 60–150 km, the resolution is good under the land area of Kanto district (Fig. 6). Table 2 Comparison of focal depths of 10 offshore events determined with different location methods No. Events
Depth2 Depth3 Rdf1 Rdf2 Rdf3
Latitude Longitude Depth1 (Km) 1 34.81 140.11 2 34.45 140.24 3 34.68 140.21 4 34.50 140.03 5 34.74 140.03 6 34.68 140.14 7 34.40 139.83 8 34.40 139.76 9 34.83 139.96 10 34.83 140.13 Average value
62.12 85.37 62.47 84.08 60.14 67.64 89.76 96.82 54.46 57.28
56.18 78.89 60.64 80.02 57.08 59.41 80.32 91.81 47.74 52.56
(Km)
(Km) (Km) (Km)
54.57 78.54 60.26 81.47 55.83 55.63 82.64 89.71 49.72 55.23
5.94 7.55 1.61 6.49 6.83 0.34 1.84 2.21 0.37 4.06 2.61 1.45 3.06 4.31 1.25 8.23 12.01 3.78 9.44 7.12 2.32 5.01 7.11 2.10 6.72 4.74 1.98 4.72 2.05 2.67 5.55 5.65 1.79
Depth1: determined by Hi-net. Depth2: determined by using the standard location method with P- and S-wave arrival time data recorded by OFS. Depth3: determined by using the MEL method. Rdf1, Rdf2, Rdf3: the same as those in Table 1.
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Fig. 4. 3D hypocentral distribution of the 6353 earthquakes analyzed in this study. Crosses represent the hypocenters of 4925 earthquakes that occurred under the land area of the Kanto district. Blue dots denote the 299 earthquakes (M N 3.0) (master events) selected from 5218 offshore events. We handpicked 1158 sP depth phases from the 299 master events. Green dots represent the 1129 events selected from 7603 subocanic earthquakes according to the distance (b6 km) from the corresponding ME. Red solid squares denote the 227 stations that recorded the sP depth phase, P- and S-wave arrival times from June 2002 to May 2005. The large rectangle denotes the region in Fig. 3 selected for testing the MEL method.
The final inverted results were obtained after 8 iterations. The hit count at the grid node is greater than 10 for the area where the velocity image is determined. The root-mean-square (RMS) travel time residual calculated from the hypocenters relocated by this study is reduced drastically from 0.543 to 0.203 s for P-waves and from 0.596 to 0.232 s for S-waves. Fig. 7 shows plan views of the Vp and Vs images at four depths, together with the distribution of background seismicity, large interplate earthquakes, active volcanoes, and active faults. Beneath the active volTable 3 Data sets used for tomographic inversion Content
Events
sP phase
P-wave
S-wave
Recorded stations
M JMA
MEa CMEb LEc Total
299 1129 4925 6353
1158
5061 19109 83362 107532
3887 13756 59766 77409
N35 N25 N25
N3.0 N2.0 N1.5
a
1158
Master events were relocated by using sP depth phase and P- and S-wave arrival times jointly. b Earthquakes that were within 6 km of master events were relocated by the MEL method of this study. c Earthquakes under the land area were relocated with the doubledifference method [1,7].
canoes, low-velocity zones are imaged clearly from the surface to a depth of 100 km. Fig. 8 shows four vertical cross-sections of the Vp and Vs images together with the distribution of active volcanoes and seismicity along the lines shown in the insert map. Low-velocity areas are visible under the offshore region at depths of 10–30 km (Fig. 7). The subducting Pacific slab is imaged clearly as a high Vp and Vs zone (Fig. 8), similar to the previous results [6,14,27– 30]. In the northern section of the study region, highvelocity bodies are imaged at depths of 35–65 km, which represent the subducting Philippine Sea plate (Figs. 7 and 8). 6. Discussion 6.1. Interpretations of the tomographic images Strong low-velocity anomalies are visible in the crust and uppermost mantle under the volcanic front of the Kanto district (Figs. 7 and 8). The CRT results for the Pand S-wave structures (Figs. 5 and 6) show good resolution at depths of 0–100 km; the low-velocity anomalies are therefore considered to be reliable features. Similar structures in the mantle wedge beneath the arc
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Fig. 5. Results of checkerboard resolution test for P-wave (a) and S-wave (b). The input amplitudes of the velocity perturbation are F 3%. The depth of each layer is shown below each map. The velocity perturbation scale is shown at the bottom.
volcanoes have been determined by many researchers in other subduction zones, such as the Aleutians [31], Alaska [32], Mediterranean [33], and Northeast Japan [6, 14], although the resolution scales are different from
one study to another. We consider that at depths of 80– 150 km, aqueous fluids are produced by progressive metamorphic dehydration reactions involving numerous hydrous minerals [34–38]. Such fluids might trigger
Z. Wang, D. Zhao / Earth and Planetary Science Letters 241 (2006) 789–803 Fig. 6. Vertical cross-sections showing results of the checkerboard resolution test of P-wave (left) and S-wave (right) along the lines shown in the insert map. Solid thick lines and gray solid triangles on the top of each section denote the land area of the Kanto district and active volcanoes, respectively. The velocity scale is shown on the right. 797
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Fig. 7. Plan views of P-wave (a) and S-wave (b) velocity images (in percent) at four depths determined by tomographic inversion. The depth of each layer is shown in the lower left corner of each map. Red color denotes low velocity while blue color represents high velocity. White stars indicate locations of the large historic earthquakes which occurred from 1923 to 1998 [49]. Crosses represent the locations of earthquakes along each profile. Active volcanoes are shown by red triangles. Curved solid lines at each depth indicate the locations of the active faults. The velocity perturbation scale and magnitude scale of the great earthquakes are shown at base of the figure.
Z. Wang, D. Zhao / Earth and Planetary Science Letters 241 (2006) 789–803 Fig. 8. Vertical cross-sections of P-wave (left) and S-wave (right) velocity images (in percent) along lines shown in the insert map. Red color represents low velocity while blue color denotes high velocity. Gray solid circles indicate background earthquakes along each section. Red triangles and bold horizontal lines on top of each section represent active volcanoes and the land area, respectively. The red star denotes the 1923 Kanto earthquake (M7.9) [57]. The solid line represents the upper boundary of the Pacific plate. The two dashed lines in (e–h) show the upper and lower boundaries of the PHS slab. The velocity perturbation scale and earthquake magnitude scale are shown on the right. 799
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partial melting under the active volcanoes and then reduce the seismic velocity of the surrounding material (Fig. 8). Many previous studies have provided evidence for low-velocity anomalies under the volcanic front in the study region (Fig. 8). High heat flow, high temperature and large geothermal gradient are found in the volcanic areas, which could also act to reduce the seismic velocity [16,39]. 3D seismic attenuation structures within the crust and upper mantle under the Kanto region [21,40] clearly demonstrate the existence of the inclined low-attenuation Pacific slab, and high-attenuation bodies in the mantle wedge under active volcanoes. The spatial distribution of the high-attenuation bodies coincides with that of the low-Vp and low-Vs anomalies identified in the present study beneath the volcanic front (Figs. 7 and 8). The low-velocity zones in the mantle wedge above the upper boundary of the subducting Pacific slab at depths of 30–50 km under the forearc region (Figs. 7 and 8) are interpreted as hydrated mantle peridotite rather than partial melting based on the following reasons: 1) The subducted oceanic lithosphere becomes dehydrated with increasing pressure and temperature, and thus the overlying mantle wedge is serpentinized [34]; 2) Serpentinization dramatically reduces the seismic velocity while increasing the Poisson’s ratio, magnetization and electrical conductivity of the forearc mantle [35]. Low-velocity anomalies in the mantle wedge are visible from our tomographic results (Fig. 8), while high Poisson’s ratio [41] and high-attenuation [40, 42] zones in the upper mantle are revealed in the northern portion of Kanto district; 3) The temperature in the low-velocity zones is suitable for serpentinization and much lower than the melting temperature of peridotite because of the low forearc temperature. Furukawa [43] estimated that the temperature is about 200 8C at a depth of 30 km beneath the Kanto district; serpentinite is stable at temperatures lower than about 600 8C [44]. Therefore, we conclude that the lowvelocity zones under the forearc region represent serpentinization of the mantle wedge (Fig. 8). The cold subducting Pacific slab and the PHS plate are imaged as high-velocity zones. The Pacific slab subducts beneath the PHS plate along the Japan Trench (Fig. 8e–h). Fig. 1 shows the contour lines of the depths of the Pacific slab and PHS plate estimated from studies of seismicity patterns, velocity structure and focal mechanisms [18]; these data are consistent with the spatial distribution of the plates imaged by the present study. Widespread low-velocity anomalies are visible in the suboceanic region (with low-velocity perturbations of Vp and Vs of 2–6% and 3–6%, respectively), along
the upper boundary of the Pacific slab (Fig. 8). Features of the low-velocity anomalies are similar to those described in the previous studies [4–6]. The CRT results for the P- and S-wave structures (Figs. 5 and 6) show good resolution under the offshore region at depths of 10 to 50 km, indicating that the low-velocity anomalies should be considered reliable features. Under the suboceanic region, the subducting sediments and oceanic crust contain free water in pore spaces and bound water within hydrous minerals [6]. At shallow depths, free water is expelled by the compaction of subducted sediments and collapse of porosity in the upper oceanic crust, which acts to decrease the seismic velocity of the crust and the uppermost mantle (Figs. 7 and 8). Previous studies indicated that fluids exist widely in the crust and uppermost mantle in the forearc regions of the subduction zone [4,6,34]. We consider that the fluids represent pore water contained in the subducting oceanic crust and sediments released due to the increasing temperature and pressure, and then reduce the seismic velocity there. A low-velocity layer at depths of 20–35 km is imaged clearly along the upper boundary of the PHS plate in the northern part of the Kanto district, indicating that the subducting PHS plate is located beneath the Eurasian plate (Fig. 8e–h). Similar features of low-velocity anomalies were observed in the previous studies of the present region, although the resolution of the previous results was not high [19,41]. A layer of high Poisson’s ratio at depths of 20–35 km has been previously imaged in the lower crust of the study region [41], and interpreted as serpentinized mantle. We consider that the low-velocity layer may reflect dehydration reactions within the PHS plate. Many studies have revealed anomalously high heat flow and strong seismic-wave attenuation in the study region [16,17,21], which provide additional evidence in support of this interpretation. A moderately high heat flow zone exists at the depth levels of the lower crust, which represents a soft ductile zone sandwiched by two brittle layers in the study region [16,17]. Their results coincide with the seismic velocity structures [19], indicating that two layers of concentrated seismicity are associated with active subduction of the oceanic lithosphere: one is in the upper crust at depths of 10–20 km, while the other is in the uppermost mantle at depths of 35–60 km. These results are in good agreement with our tomographic results (Fig. 8); both studies identify a low-velocity layer at depths of 20–35 km sandwiched by two high-velocity layers. A high-attenuation layer was imaged at depths of 25–40 km in the northern section of the study region [21]. We therefore conclude that the low-velocity layer
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may result from fluids released by dehydration reactions within the descending PHS plate; the fluids act to reduce the seismic velocity of the surrounding material (Fig. 8). 6.2. Volcanoes, great earthquakes, and the subducting slabs High-velocity zones are revealed at depths of 25–60 km in the northern portion of the study region, which might reflect the existence of the subducting PHS plate (Figs. 7 and 8e–h). We find that the subducting Philippine Sea plate has affected the spatial distribution of the active volcanoes related to the subducting Pacific slab in the Northern portion of the study region (Fig. 8e–h). In general, low-velocity zones under active volcanoes are sub-parallel to the descending slab in the associated subduction zone [e.g. 6,14]. However, such a trend is ambiguous in the Kanto district due to the fact that the PHS plate overlies the Pacific slab. Iwamori [45] suggested that the influence of the subducted PHS plate on magmatism is to shift the dehydration reactions to a greater depth along the Pacific slab. Low-velocity anomalies are clearly imaged at depths of 80–110 km between the PHS plate and the Pacific slab (Fig. 8e–h), which perhaps reflect dehydration processes within the subducting Pacific slab. The low-velocity material is not parallel to the Pacific slab but deflect towards the backarc due to the overlying PHS plate (Fig. 8e–h). The low-velocity material at the depths of 80–110 km imaged in the cross-sections of E–F and G–H probably represents the effects of volatiles, but not indicating melts due to the cold PHS slab (Fig. 8e–h). Our tomographic results clearly show that in the northern portion, active volcanoes are absent due to the influence of the descending PHS plate; in the middle region, the volcanic front is deflected towards the backarc because of the partial overlapping of the PHS plate; while south of Izu, the active volcanoes are located 120–140 km above the Pacific slab because the PHS plate has not yet subducted there (Fig. 8). These observations indicate that the fluids supplied by the subducting Pacific slab cannot rise to the crust due to the existence of the PHS plate beneath the northern portion of the Kanto district. Thus, we consider that the PHS plate existing between the Eurasian and Pacific plates has two possible effects on the arc magmatism: one is to shift the dehydration reactions to a greater depth along the Pacific slab [45]; the other is to act as a barrier to the upwelling of fluids released from dehydration reactions within the subducting Pacific slab. Further detailed studies with geological, geophysical and geochemical approaches would
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provide us with a better understanding of the relations between the volcanism, the subducting PHS plate and the Pacific slab in the study region. Many large earthquakes occur along the boundaries between the subducting Pacific plate, the Eurasian plate and the PHS plate in the subduction zone. The high seismic risk of the study region has been pointed out by extensive geological and geophysical studies over a long time [18]. The plate subduction and plate interaction are responsible for the large damaging earthquakes that occurred in this region. The strength of interplate coupling between the subducting PHS plate and the Eurasian plate was investigated and the plate boundary was estimated to be located at depths of 10–25 km under this region [46–48]. We found that seismic activity along the boundary between the PHS and Eurasian plates is low compared with that within the overriding Eurasian plate and in the subducting PHS plate (Fig. 8e–h). The 1923 Kanto earthquake (M7.9) occurred on the Eurasian–PHS plate boundary (Fig. 8e, f), where the interplate seismic coupling of the two plates is expected to be strong. Large earthquakes (M z 6.0) that occurred during the period of 1923 to 1998 [49] are shown in white stars in Fig. 7. We divided these earthquakes into two groups according to their hypocentral distribution. One group (including the 1923 Kanto earthquake) occurred along the boundary between the PHS and Eurasian plates, while the second group occurred along the upper boundary of the subducting Pacific slab (Fig. 7). Most of the great earthquakes, such as the 1909 Bousooki (M7.5) and the 1923 Kanto (M7.9) earthquakes, occurred around the low-velocity zones (Fig. 7); this might indicate the existence of weakly coupled sections on the subducting slab boundary. For the subducting Pacific slab, the down-dip limit of the great thrust earthquakes generally corresponds to the intersection of the thrust zone with the forearc mantle [50,51], which might be explained by the distribution of aseismic hydrous minerals such as serpentinite [34,52,53]. Our results show that large earthquakes that occurred along the PHS–Eurasian and the upper Pacific boundaries are located in areas with seismic velocities higher than that of surrounding areas (Fig. 7), which might reflect strongly coupled asperities. Previous studies indicated that fluids exist widely in the crust and uppermost mantle in the forearc regions of the subduction zone [4,34], which is consistent with our tomographic images. We conclude that the spatial distribution of large earthquakes is possibly controlled by lateral heterogeneities, including strongly coupled sections (asperities) and weakly coupled or decoupled sections along the upper slab boundary.
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7. Conclusions
References
In this study we present an efficient combined method for accurately locating a large number of earthquakes that occurred outside the area covered by the seismic network under the Kanto district, Japan. The method makes use of sP depth-phase data together with the double-difference arrival time data. The hypocenters relocated by this method are consistent with those determined from OFS data; the spatial differences in the hypocenters relocated by the two methods are generally less than 2 km. This new approach to hypocenter location can be applied to other offshore regions of subduction zones where few seismic stations exist to record the numerous suboceanic earthquakes. Strong low-velocity anomalies beneath the active volcanoes are clearly imaged, and might reflect aqueous fluids released from the subducting Pacific slab interacting with hot upwelling flow in the mantle wedge, leading to partial melting. A low-velocity seismic layer is visible along the upper boundary of the PHS plate, which may be associated with dehydration reactions within the subducting slab. The cold subducting Pacific slab shows high-velocity anomalies that are ~2–6% higher than that of the normal mantle. We find that the PHS plate existing between the Eurasian plate and the subducting Pacific slab has influenced the spatial distribution of arc volcanoes in northern part of the study region. Widespread low-velocity anomalies are observed on the upper boundary of the subducting Pacific slab under the forarc region, which might reflect weakly coupled or decoupled sections of the plate interface related to serpentinization of the forearc mantle associated with slab dehydration. On the upper boundary of the Pacific slab, under the offshore region, strong lateral heterogeneities are visible, which show a strong correlation with the spatial distribution of the great earthquakes. These large earthquakes occurred outside the low-velocity zones which may be caused by aqueous fluids released from the subducted oceanic crust.
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