Seismic activity beneath the Nankai trough revealed by DONET ocean-bottom observations
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We conducted a detailed investigation of seismic activity from January 2011 to February 2013 along the Nankai trough off the Kii Peninsula, central Japan, by using data obtained from the DONET ocean-bottom observation network. The hypocenters are mostly within the subducting Philippine Sea (PHS) plate, although a few are along the plate boundary or in the sedimentary wedge below the Kumano forearc basin. The seismic activity can be separated into events above and below 20 km depth, which corresponds approximately to the Moho. The hypocenter distributions are distinctly different for these groups. The seismic activity in the oceanic crust can be further separated into three clusters. Most of the seismic activity recorded in our data represents aftershocks of the 2004 off the Kii Peninsula earthquakes (M JMA = 7.1, 7.4, and 6.5), which occurred in the PHS plate. The hypocenter distribution in the oceanic crust correlates well with the location of the Paleo-Zenisu ridge, which is formed by a chain of seamounts that is subducting beneath the forearc basin. The hypocenters in the uppermost mantle are aligned on a plane dipping to the southeast, consistent with the existence of a thrust fault cutting through the lithosphere of the oceanic plate. The focal mechanisms of the earthquakes show that the axis of compressive stress in the PHS plate is oriented N–S, almost perpendicular to the direction of plate convergence, indicating a complex tectonic regime in this region. These results suggest that intraplate shortening may be occurring in the subducting oceanic plate.
KeywordsHypocenter distribution Paleo-Zenisu ridge Ocean-bottom seismometers 2004 off the Kii Peninsula earthquakes
The most recent significant seismic activity in this region was the sequence of 2004 off the Kii Peninsula earthquakes (M JMA = 7.1, 7.4, and 6.5). These earthquakes occurred below the trough axis and were intraplate earthquakes within the oceanic crust of the subducting PHS plate (Fig. 1). Although the centroid moment tensor (CMT) solutions for these earthquakes obtained by F-net operated by the National Research Institute for Earthquake Prediction and Disaster Prevention (NIED) show a focal mechanism with P axis trending N–S (see Fig. 1), details of the rupture process of the mainshock are controversial. Some researchers consider that this event was a composite of movements on both a dip-slip fault below the trough axis and a strike-slip fault below the Kumano forearc basin (e.g., Yagi 2004; Satake et al. 2005; Hara 2005). Other researchers have proposed a model in which the mainshock ruptured a reverse fault below the forearc basin (e.g., Yamanaka 2004; Baba et al. 2005; Matsumoto and Mikada. 2005; Saito et al. 2010). After the 2004 earthquakes, very-low-frequency (VLF) earthquake activity was triggered above the source region (Obara and Ito 2005). Recent studies show that the VLF earthquakes are in the sedimentary wedge or at the plate boundary (Ito and Obara 2006; Sugioka et al. 2012). Kitajima and Saffer (2012) considered that elevated pore pressures in the sediments play an important role in the generation of VLF earthquakes.
The Zenisu ridge, formed by a chain of seamounts rising 2,000 m from the bottom of the Nankai trough, lies southeast of the DONET stations (Fig. 1). The ridge runs almost parallel to the trough axis and a thrust cutting through the lithosphere of the PHS plate is found below this ridge (e.g., Lallemant et al. 1989; Nakanishi et al. 1998, 2002b; Mazzotti et al. 2002). The slip on the thrust below this ridge may indicate shortening of the oceanic plate in this region (e.g., Mazzotti et al. 2002). Northwest of this ridge, seafloor morphologic data (Lallemand et al. 1992; Okino and Kato 1995) and geomagnetic data (Le Pichon et al. 1996) show a structure below the Kumano forearc basin that represents a subducting seamount. This seamount is part of the Paleo-Zenisu ridge, which was also formed by a chain of seamounts that is almost parallel to the trough axis, 20–30 km wide, about 200 km long, and with a maximum height of 2.5 km. Park et al. (2003) made a detailed estimate of the extent of the Paleo-Zenisu ridge from seismic reflection and refraction data and showed that it correlates well with the seaward edge of the coseismic rupture zone of the 1944 Tonankai earthquake (Mw = 8.1). They considered that the ridge affects mechanical coupling at the plate boundary.
In a previous study (Nakano et al. 2013), based on data from January to August 2011, we showed that DONET ocean-floor observation data provide a better detection capability and higher precision of earthquake source locations below the Nankai trough than analyses based only on onshore observation data. However, DONET was still under construction during most of the analysis period; construction of the final planned station was finished in the end of June 2011. Accordingly, the detection capability and resolution of earthquake source location was still limited, especially south of the observation network. In this study, we extended our analysis period to cover the period from January 2011 to February 2013. Adding 1.5 years data from full-operation network, we have developed more comprehensive earthquake catalogue off the Kii peninsula. The hypocenters correlate well with the aftershock distribution of the 2004 earthquakes. We also used the new data to investigate earthquake focal mechanisms. This new and larger dataset forms the basis of our discussion of the tectonic implications of seismic activity along the Nankai trough.
Data and analysis method
DONET is a cabled ocean-bottom observation network system consisting of 20 permanent stations installed on the ocean floor of the Kumano forearc basin (Fig. 1). Seismic signals are recorded by a broadband seismometer (Guralp CMG-3T) and a strong-motion seismometer (Metrozet TSA-100S) at each station. The seismometers are buried 1 m below the ocean floor to minimize environmental noise. The signals from the seismometers are sampled at 200 Hz per channel by a 24-bit analog to digital converter on site and then transferred in real time by fiber optic cable to our laboratory.
For hypocenter determinations we used data obtained from DONET stations, from land stations on the Kii Peninsula operated by the Japan Meteorological Agency (JMA), and from Hi-net and F-net stations operated by NIED. We followed the analysis method of Nakano et al. (2013), which is summarized below.
First, P- and S-wave first arrivals at each station were picked manually. Initial hypocenters were determined by the method of Hirata and Matsu’ura (1987). We assumed the layered P-wave velocity structure compiled for this region by Nakamura et al. (2011). The S-wave velocity (Vs) was calculated assuming a Vp/Vs ratio of 1.8, where Vp is P-wave velocity. We determined the hypocenters for earthquakes for which we had at least six P and/or S readings. The local magnitude (ML) was determined from the maximum amplitude of the vertical velocity seismograms (Watanabe 1971). We incorporated station corrections for P- and S-wave arrival times, which were obtained from the averaged differences between observed and calculated travel times. The error of the initial hypocenter location depends on where the earthquake occurs. The averaged error in each earthquake cluster (see Fig. 4) is as follows: In cluster A, the error is less than 1 km in both horizontal and vertical directions. In cluster B, the error is 1–2 and 6 km in horizontal and vertical directions, respectively. In cluster C and events deeper than 20 km, the error is 1 and 2 km in horizontal and vertical directions, respectively.
After initial determination of hypocenters, we obtained a detailed hypocenter distribution by using the double-difference (DD) method (Waldhauser and Ellsworth 2000). The DD method minimizes errors due to structures not resolved by velocity modeling and improves the precision of relative source locations in a seismic activity swarm. Because the DD method assumes that the recording stations are at sea level, we reduced the source depths of the initial hypocenters by 2 km, which corresponds to the average seafloor depth of the Kumano forearc basin; most of the DONET stations are at about this depth (Fig. 1). The source depths were increased by 2 km after application of the DD method. We used arrival time differences for 165,841 and 233,255 pairs of P and S waves, respectively. Selecting neighboring events within 5 km, we relocated 5,041 events. After relocation, the RMS of the DD time residual was reduced from 423 to 109 ms. We removed some scattered earthquakes that were not well grouped with other events in the obtained hypocenter distribution. We note that for the DD method, the absolute location of earthquake clusters depends on the initial hypocenter distribution. Because the initial hypocenter locations depend on the assumed velocity model, we repeated our analysis using different Vp/Vs ratios. Although the obtained source depths were dependent on the assumed Vp/Vs ratio, it had little effect on the epicenter distribution.
We also determined earthquake focal mechanisms by the method of Imanishi et al. (2006a, b, c). In this method, the ω2-model (Boatwright 1978) is fitted to the amplitude spectrum of the direct P or S waves at each station, and the level extrapolated to DC is used as its amplitude. Assuming a double-couple source, the obtained amplitude distribution is inverted to obtain the focal mechanism. In this study we used P-wave amplitude and polarity. If the P-wave polarity was not clear, the absolute value of amplitude was used. S-waves were not used because their amplitudes can be strongly affected by amplification in the sedimentary layers below the station. Using the P-wave amplitude and polarity distribution, the obtained mechanism is less ambiguous than that traditionally obtained by the P-wave polarity distribution, even when station coverage of the source is not good. We applied the inversion to events with ML > 3. After checking the stability of the results, we obtained focal mechanisms for 24 events. Magnitudes recalculated from the seismic moment obtained by the inversion were between 2.8 and 3.7. The stability of the focal mechanism determination is shown in Appendix.
The seismic activity in the oceanic crust (Fig. 4a) can be separated into three clusters. The activity in the north (cluster A in Fig. 4a) corresponds well with the aftershock distribution of the 2004 events (Sakai et al. 2005), which showed a NW–SE trend. The distribution of cluster A hypocenters in our new data is narrower than the active region immediately after the 2004 mainshock. In addition to the overall NW–SE trend within cluster A, there is an ENE–WSW trend, almost parallel to the trough axis. This trend was not so clear in our previous study. Cross sections along these trends show that the hypocenters are at depths between 5 and 20 km (Fig. 3a, c).
Another cluster of seismic activity to the southeast of cluster A (cluster B in Fig. 4a) shows a clear NE–SW trend, which activity was not clearly recognized in our previous study. This may be due to weak detection capability in this region. The F-net CMT solution of the mainshock of the 2004 events is in this region (Fig. 1). Aftershocks immediately after the 2004 mainshock were widely distributed in this region (Sakai et al. 2005), whereas our data show activity to be limited to a narrower region.
In cluster C (west of cluster B, Fig. 4a), the hypocenters are distributed in a narrow region with a N–S trend in the northern part of the cluster, but we found a more scattered E–W trend south of the trough axis. Some hypocenters are within the sedimentary wedge at the northern end of cluster C (Fig. 3b, d). This cluster lies outside the 2004 aftershock region; Nakano et al. (2013) attributed this cluster to activity along pre-existing weak structures in the incoming PHS plate. We note that there are some ambiguities in the source depths of earthquakes in clusters B and C, which may reflect the assumed velocity structure.
We found only limited seismic activity along the plate boundary, and very little in the region corresponding to the large slip of the 1944 Tonankai earthquake determined by Kikuchi et al. (2003) (Fig. 3a). We also found only a few earthquakes in the sedimentary wedge. Although very-low-frequency (VLF) earthquakes are known to occur in the sedimentary wedge and at the plate boundary in this region (Obara and Ito 2005; Ito and Obara 2006; Sugioka et al. 2012), their distribution has little overlap with that of the ordinary earthquakes (Fig. 2). The lack of ordinary earthquakes may reflect a sedimentary wedge consisting of soft sediments in which insufficient stress can accumulate to nucleate earthquakes that radiate high-frequency seismic waves. However, the earthquakes we detected in the sedimentary wedge do not show characteristics of VLF earthquakes.
Most of the seismic activity in our data from around the Nankai trough corresponds well with the aftershock distribution of the 2004 off the Kii Peninsula earthquakes, indicating that they represent continuing aftershock activity. The CMT solutions of the 2004 events show that the sources of the foreshock, mainshock, and largest aftershock were below the trough axis (Fig. 1) at depths between 5 and 15 km. The mainshock caused tsunamis of up to 0.9 m height along the coast of the Kii Peninsula (e.g., Satake et al. 2005; Baba et al. 2005; Matsumoto and Mikada 2005), which suggests that the 2004 events occurred in oceanic crust, although aftershock activity has extended to the uppermost mantle (Sakai et al. 2005). The CMT solutions show significant non-double-couple components (about 15–20 %), implying complex rupture processes.
Many source models have been proposed for the mainshock of the 2004 events, and the subject remains controversial. One group proposes that the mainshock is the composite result of dip-slip on a fault below the trough axis, consistent with the CMT solution, and a strike-slip fault of NW–SE strike below the Kumano forearc basin, corresponding to the activity of cluster A (e.g., Yagi 2004; Satake et al. 2005; Hara 2005). Another group proposes that the mainshock occurred on a NW–SE striking reverse fault extending between the trough axis and the forearc basin and dipping to the southwest. This model is based on the aftershock distribution corresponding to cluster A and the initial seawater uplift estimated from tsunami observations (e.g., Yamanaka 2004; Baba et al. 2005; Matsumoto and Mikada 2005; Saito et al. 2010), although the reverse fault of this model is in conflict with the CMT solution.
The earthquakes in cluster A are distributed in an area extending over several tens of kilometers. This cluster corresponds well with the aftershock distribution, but we could not recognize a planar structure within the hypocenter distribution. Thus, it is difficult to identify a fault model for the 2004 mainshock. The distribution of the earthquakes below the trough axis (cluster B) has a NE–SW orientation, which is consistent with the strike of one of the nodal planes of the CMT solution, although no planar structures were evident in this distribution.
The Zenisu ridge lies southeast of the study area (Fig. 1). Below this bathymetric high a thrust is found in the lithosphere of the PHS plate (e.g., Lallemant et al. 1989; Nakanishi et al. 1998, 2002b; Mazzotti et al. 2002). Northwest of this ridge, the Paleo-Zenisu ridge is a similar bathymetric high formed on the PHS plate, which is also subducting beneath the Kumano forearc basin (Fig. 2). The formation of these ridges was a result of intraplate shortening in the subducting plate (Mazzotti et al. 2002). Park et al. (2003) showed that the southwestern edge of the Paleo-Zenisu ridge corresponds well with the area of dense earthquake activity in cluster A (Fig. 2).
To attempt to explain the earthquake distribution around the Zenisu ridge, Mazzotti et al. (2002) computed the changes in Coulomb failure stress (CFS) in the oceanic crust that would be caused by slip on a thrust in the uppermost mantle. They found two areas of CFS increase in the oceanic crust for a thrust (Fig. 7b), one at the upward extension of the thrust and the other on the upthrown side of the thrust, similar to clusters A and B of this study (Fig. 7a). Therefore, if we assume a thrust in the uppermost mantle below the Nankai trough, the spatial distribution of the seismic activity of our data corresponds well with the resultant increase of CFS in the oceanic crust.
Despite the similarity of the spatial patterns of the seismically active regions shown in Fig. 7a, b, the thrust below the Nankai trough assumed in this study dips southeast (Fig. 7a), whereas the thrust below the Zenisu ridge of previous studies (e.g., Nakanishi et al. 1998, 2002b; Mazzotti et al. 2002) dips northwest. In mechanical terms, these thrusts are not contradictory; both reflect compressive stress within the PHS plate. However, the position of the Paleo-Zenisu ridge on the surface of the PHS plate does not correspond to the area that would be uplifted by slip on the assumed thrust below the Nankai trough (Fig. 7a), so thrusting may not be the primary factor in formation of the ridge. The focal mechanisms of the earthquakes in the mantle are reverse faults, but neither of the two nodal planes matches the dip of the assumed thrust (Fig. 5). Therefore, we consider that the seismic activity of our data set reflects compressional deformation in the uppermost mantle, possibly indicating that intraplate shortening of the PHS plate is occurring below the Nankai trough. However, our data set does not allow us to clarify the fault structure.
The focal mechanisms obtained in this study show that the P axis is oriented NNE–SSW. The P axis of the CMT solutions obtained for the 2004 earthquakes (Fig. 1) and aftershocks (Ito et al. 2005) are of similar orientation. Similar focal mechanisms have also been obtained for earthquakes before the 2004 events (e.g., Obana et al. 2004, 2005). These results suggest that the axis of compressive stress in the PHS plate is oriented N–S to NNE–SSW, almost perpendicular to the direction of motion of the PHS plate with respect to the Eurasian plate (see Fig. 1).
West of the Kii peninsula, Mochizuki et al. (2010) showed that the P-axis of intraslab earthquakes is oriented almost normal to the trough axis. This feature may be expected in a region where interplate coupling is strong. Yoshioka and Matsuoka (2013) showed strong intrerplate coupling ratio along the Nankai trough below the Kii peninsula. Accordingly, we would have expected a similar stress field in the PHS plate below southeast of the Kii peninsula, too. Bending of the plate due to subduction may also affect the stress field within the PHS plate (e.g., Seno 2005; Miyoshi and Ishibashi, 2005), in which compressional and tensional stress perpendicular to the trough axis dominate in the lower and upper parts, respectively, within the lithosphere. However, these mechanisms could not explain the observed direction of compressive stress oriented perpendicular to the direction of plate convergence. The focal mechanisms of the VLF earthquakes are mostly thrusts striking parallel to the trough axis (Fig. 2), which is a common phenomenon for earthquakes at plate boundaries. Taking into consideration these observations, the stress field in the PHS plate in our study area is clearly anomalous. One possible explanation is that the stress field has been disturbed by collision at the Izu Peninsula east of our study area (e.g., Obana et al. 2004; Seno 2005; Miyoshi and Ishibashi 2005), but the large difference in orientation of the axes of compressional stress are not well explained by this model. This question requires further investigation.
DONET ocean-bottom seismic observations have revealed detailed earthquake distributions and tectonic features below the Nankai trough off the Kii Peninsula, which could not have been obtained by onshore observations alone. Our study indicates that most of the seismic activity in the study area represents aftershocks of the 2004 off the Kii Peninsula earthquakes. This seismic activity extends into the uppermost mantle. Focal mechanism solutions suggest that the axis of compressive stress in the PHS plate is oriented N–S, almost perpendicular to the direction of plate convergence in this region. Our data indicate that intraplate shortening of PHS plate may be occurring below the Nankai trough.
We used data obtained from the Hi-net and F-net systems operated by NIED and onshore observations obtained by JMA. We thank M. Takaesu and S. Yada for their support with data processing. We also thank Drs. H. Sugioka and Y. Ito for sharing the source parameters of VLF earthquakes. All figures were drawn using Generic Mapping Tools (Wessel and Smith 1998). We greatly appreciate comments from two anonymous reviewers.
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