Stress field in the 2008 Iwate Miyagi earthquake (M7.2) area

Transcription

1 Article Volume 12, Number 6 18 June 2011 Q06006, doi: /2011gc ISSN: Stress field in the 2008 Iwate Miyagi earthquake (M7.2) area Zhouchuan Huang Department of, Tohoku University, Sendai , Japan (zhouchuan.huang@gmail.com) School of Earth Sciences and Engineering, Nanjing University, Nanjing , China Dapeng Zhao Department of, Tohoku University, Sendai , Japan (zhao@aob.gp.tohoku.ac.jp) Liangshu Wang School of Earth Sciences and Engineering, Nanjing University, Nanjing , China (lswang@nju.edu.cn) [1] We determined the focal mechanism solutions (FMS) of 191 crustal earthquakes as well as the stress tensor in the source area of the 2008 Iwate Miyagi earthquake (2008 IMEQ, M7.2) that occurred in the central portion of northeast (NE) Japan. The FMS and the stress tensors were determined by using both 1 D and 3 D velocity models, which exhibit almost the same results. The differences caused by the use of 1 D and 3 D models can be neglected when compared with the differences due to the different methods, which indicates that the FMS and the stress tensor determined with a 1 D model are accurate enough to study the crustal stress field in the study region. The obtained P axis (s 1 ) trends WNW ESE subhorizontally, and the T axis (s 3 ) is oriented subvertically in a NNE SSW belt perpendicular to s 1. The s 1 orientation is consistent with the motion of the Pacific plate relative to NE Japan, which indicates that the plate boundary forces dominate the intraplate stress regime. Both temporal and spatial variations of the stress field in the IMEQ source area are detected, which may be induced by the stress rotation accompanying the main shock and its aftershocks. The seismogenic faults in the study area are estimated to be very weak, which argues against the concept of strong crust. The faults may be weakened by the high temperature magma and the fluids in the lower crust and uppermost mantle that intrude upward into the shallower crust. Components: 6000 words, 11 figures. Keywords: 3 D velocity model; arc magma and fluids; focal mechanism solution; plate boundary forces; stress rotation; stress tensor inversion. Index Terms: 7205 Seismology: Continental crust (1219); 8124 Tectonophysics: Earth s interior: composition and state (1212, 7207, 7208, 8105); 8164 Tectonophysics: Stresses: crust and lithosphere. Received 15 March 2011; Accepted 22 April 2011; Published 18 June Huang, Z., D. Zhao, and L. Wang (2011), Stress field in the 2008 Iwate Miyagi earthquake (M7.2) area, Geochem. Geophys. Geosyst., 12, Q06006, doi: /2011gc Introduction [2] Determining the focal mechanism solutions (FMS) of earthquakes has become a fundamental but routine procedure in seismology. The FMS of M > 5.5 earthquakes are usually determined in hours after the earthquakes by the Global Centroid Moment Tensor (CMT) project [e.g., Dziewonski Copyright 2011 by the American Geophysical Union 1 of 13

2 Figure 1. Seismicity (gray crosses) in the source area of the 2008 Iwate Miyagi earthquake (M 7.2). White circles denote the 191 earthquakes whose focal mechanism solutions are determined in this study. The dashed line shows the boundary between the northern and southern parts of the study area as mentioned in the text. The red stars show the main shock and three major aftershocks. Solid triangles denote the active volcanoes. The lower right and upper left insets show the present study area and the main shock fault type, respectively. and Woodhouse, 1983]. These FMS provide important information on the stress field in various regions worldwide, such as the mid ocean ridges and subduction zones. The FMS for the smaller earthquakes are also determined in regions where dense seismic networks are operated. For example, in the Japan Islands the FMS of M > 3.5 earthquakes are routinely determined by the data centers of the Broadband Seismograph Network (F net) and the High Sensitivity Seismograph Network (Hi net) deployed on the Japan Islands. [3] The FMS are usually determined by using a 1 D seismic velocity model. Now that 3 D velocity models have been determined in more and more regions, there is a question whether the use of a 3 D velocity model can improve the accuracy of the FMS and the stress regime in a region. In Northeast (NE) Japan, the Pacific plate is subducting beneath the Okhotsk (or North American) plate at a rate of 9 cm/year from the Japan Trench. Many seismic tomography studies have been made in NE Japan, which revealed strong structural heterogeneities in the crust and upper mantle, such as the high velocity subducting Pacific slab and the arc magma related low velocity anomalies in the mantle wedge (see Zhao et al. [2011] for a recent review). These structural heterogeneities may affect the FMS estimation with the P wave first motion (polarity) data or through the waveform inversion. In addition, the earthquakes occurred on a preexisting weak zone or fault and the stress directions inferred from individual FMS may show considerable lateral variations [McKenzie, 1969; Scholz, 2002]. The regional stress field can only be estimated from a group of events in a limited area by using either the FMS [e.g., Gephart and Forsyth, 1984; Michael, 1987] or the P wave polarity data directly [e.g., Horiuchi et al., 1995; Zhao et al., 1997]. [4] The 2008 Iwate Miyagi earthquake (IMEQ) with the JMA (Japan Meteorological Agency) magnitude 7.2 occurred in the central portion of NE Japan on June 14, 2008 (Figure 1). Its main shock occurred on a west dipping reverse fault, where slips up to 6 8m on the fault plane were derived from the waveform inversion [Fukahata et al., 2008; Toda and Maruyama, 2009; Suzuki et al., 2010]. Numerous aftershocks occurred in a 40 km long, NNE SSW trending belt along the fault, which were recorded by many permanent and portable seismograph stations installed by the JMA, Hi net and Japanese universities. Several researchers have used the arrival time data from the aftershocks to study the 3 D crustal structure in the Iwate Miyagi source area [e.g., Wang et al., 2008; Okada et al., 2010; Cheng et al., 2011]. [5] In this work, we first determined the FMS of 191 Iwate Miyagi aftershocks using P wave polarity data in both 1 D and 3 D velocity models to investigate the potential effect of the structural heterogeneity on the FMS determination. Then we used a large number of P wave polarity data to conduct the stress tensor inversions to investigate the spatial and temporal variations in the stress field in the 2008 IMEQ area. 2. Data and Method [6] The 191 events (M > 3.0) used in this study (Figure 1 and Data Set S1 in the auxiliary material) 2of13

3 Figure 2. The differences in the takeoff angle (blue) and azimuth (red) of the 6693 P wave polarity data calculated for the 1 D and 3 D velocity models [Huang et al., 2011], which are plotted against the epicentral distance. The corresponding histograms are shown in the insets. were recorded by the JMA and Hi net seismograph stations installed on the Japan Islands. 1 Most of the events are the aftershocks of the 2008 IMEQ (Data Set S1) while only 9 of them occurred before the main shock. The 191 events generated a total of 6693 P wave polarity data, and there are at least 25 polarity data for each event. [7] For the P wave polarity data, we first calculated their raypaths using the 3 D ray tracing technique of Zhao et al. [1992] for both a 1 D velocity model and a 3 D velocity model of the crust and upper mantle under a wide region from the Japan Trench to the Japan Sea [Huang et al., 2011]. The azimuths and takeoff angles were then calculated and compared (Figure 2). The double couple FMS of the 191 events were determined by minimizing the number of inconsistent polarity data applying a grid search algorithm [Horiuchi et al., 1995] (Figure S1 and Data Set S1). The step in the grid search is 3 for both the strike and dip of the nodal plane. The inconsistent first motions for most of the events are less than 10% of the total polarity data and they reach 15% for only 9 events. [8] We used the stress tensor inversion method of Horiuchi et al. [1995] to invert for the stress field in the source area of the 2008 IMEQ. Two assumptions are made following the previous studies: (1) the stress field is uniform in the stress inversion area, and (2) the direction of the faulting is parallel to the direction where the shear stress becomes maximum [Gephart and Forsyth, 1984; Horiuchi 1 Auxiliary material data sets are available at ftp://ftp.agu.org/ apend/gc/2011gc Other auxiliary material files are in the HTML. et al., 1995]. The method of Horiuchi et al. [1995] adopts a grid search approach (the step is 6 ) to invert the P wave polarity data to estimate four of the six elements of the stress tensor. The absolute magnitude of the stress cannot be resolved, nor can the isotropic component of the stress tensor without assuming a particular failure criterion. The four resolvable elements are the orientations of the three principal stresses and the stress ratio. The stress ratio is a measure of the relative amplitudes of the stress components, and it is defined as R = (s 1 s 2 )/ (s 1 s 3 ), where s 1, s 2, and s 3 are the maximum, intermediate, and minimum stresses, respectively. The stress tensor is inverted by using the P wave polarity data whose azimuths and takeoff angles Figure 3. Distributions of the P and T axes of the 191 focal mechanism solutions determined with the (a) 1 D and (b) 3 D velocity models. The differences in the P and T axes between the (c) 1 D and(d)3 D velocity models. 3of13

4 Figure 4. Focal mechanism solutions (FMS) of the 191 earthquakes in the 2008 Iwate Miyagi earthquake area determined by using the 3 D velocity model [Huang et al., 2011]. The magnitude scale of the earthquakes is shown at the bottom. The FMS of the main shock and four major aftershocks determined by the Hi net, F net are also shown. Solid triangles denote the active volcanoes. were determined by using the 1 D and 3 D velocity models. Besides the 191 events as mentioned above, P wave polarity data from additional 54 events with 10 recordings were included in the inversion. The uncertainties of the inverted stress tensor were estimated by using a statistical method with 2000 times of nonparametric bootstrap resamplings [Michael, 1987]. 3. Results [9] Figure 3 summarizes the P and T axes determined by the 191 FMS using the 1 D and 3 D velocity models (Data Set S1). Most of the solutions are robust, though some are rather nonunique (Figure S1). The FMS resulting from both velocity models show dominant WNW ESE trending, subhorizontal P axis and NNE SSW trending, subvertical T axis (Figures 3a, 3b and S2). The principal stresses are consistent with those of the main shock and several large aftershocks (Figure 4), which denote the typical reverse faulting. Figure 5 shows the spatial distributions of the P and T axes in the 2008 IMEQ area, suggesting insignificant lateral variations in the stress field. 4of13

5 Figure 5. Distributions of the (a) P axis and (b) T axis of the 191 FMS (Figure 4) projected to the surface and to the eastwest and north south vertical cross sections. The lengths of short bars denote the angle between the axes and the corresponding planes with the scale shown in the bottom right. Red stars show the main shock and three large aftershocks. Solid triangles denote the active volcanoes. 5of13

6 Figure 6. (left) Results of the stress tensor inversion using the 6693 polarity data (Figure 2) determined with (a) the 1 D velocity model, (b) the 3 D velocity model of Huang et al. [2011], and (c) the 3 D velocity model of Cheng et al. [2011]. The maximum, intermediate, and minimum principle stresses are shown as circles, triangles, and squares, respectively. Large symbols denote the optimal orientations by 2000 times nonparametric bootstrap resamplings, while small symbols denote the 95% confidence areas. (right) Histograms of the optimal values of the stress ratio R (red arrows) and the 95% confidence areas (blue columns). [10] Figures 6a and 6b show the results of the stress tensor inversion using the 1 D and 3 D velocity models [Huang et al., 2011], respectively. Similar to the stress field derived from the FMS, the stress tensors for both velocity models indicate WNW ESE trending, subhorizontal P axis (s 1 ) and subvertical T axis (s 3 ). The optimal stress ratio R is 0.50 with a 95% confidence range from 0.4 to 0.7. The direction of the maximum stress s 1 is consistent with both the motion of the Pacific plate 6of13

7 Figure 7. Spatial distribution of the estimated P axes in the 2008 IMEQ source area. The orientation and length of the arrows denote the azimuths and plunges, respectively. The directions of the arrows denote the dip of the axes. Colored circles denote the R values with the scale shown in the inset. The gray parts denote the 95% confidence ranges of the azimuths. Major earthquakes and active volcanoes are shown as stars and triangles, respectively. AA and BB are the two profiles shown in Figure 8. relative to NE Japan and the horizontal principal strain in this area derived from GPS observations [Seno, 1999; Miura et al., 2004]. 4. Discussion 4.1. Influence of Velocity Models on the Stress Inversion [11] Figure 2 shows the differences in the azimuth and takeoff angle of the 6693 P wave polarities determined with the 1 D and 3 D velocity models. For most of the polarity data, the differences are generally small, being 1 for the azimuth and 2 for the takeoff angle. For some data with a short epicentral distance (e.g., < 100 km), however, the differences in the takeoff angle reach as large as 10. The 191 FMS obtained do not show significant differences between the 1 D and 3 D velocity models. The differences in the P and T axes (DP, DT) for most robust FMS are not greater than 6 (Figure 3 and Data Set S1). The use of the 3 D velocity model does not significantly improve the FMS. We also compare the FMS with those determined by the Hi net and F net for the main shock and four large aftershocks of the 2008 IMEQ (Figure 4). The Hi net solutions are also determined with the P wave polarity data but a different 1 D velocity model is adopted, while the F net solutions are determined by moment tensor inversion of waveforms [Fukuyama et al., 1998]. In spite of the similar pattern, significant discrepancies are visible between these solutions. In fact, the differences in the solutions caused by the 1 D and 3 D models are negligible as compared with the discrepancies between the Hi net and F net solutions. [12] The inversions of the P wave polarity data with the 1 D and 3 D velocity models have yielded almost the same stress tensors, for both the optimal values and the 95% confidence ranges. There is only a difference of 6 in the azimuth of the optimal maximum stress s 1. The stress tensor results also show small differences for the 1 D and 3 D velocity models. We note that the resolution of the adopted 3 D velocity model is km [Huang et al., 2011], which may be too low to improve the stress tensor inversion. So we further determined the FMS of the 191 events and inverted for the stress tensor in 2008 IMEQ focal area with a higher resolution (5 km) 3 D velocity model [Cheng et al., 2011]. Although the stress ratio R seems improved with a more convergent value (R = 0.6), the three stress components are almost the same as those for the 1 D model (Figure 6c). The FMS also show the same feature (Data Set S1). Yukutake et al. [2007] adopted a similar approach to study the stress field in the source area of the 2000 Western Tottori earthquake (M 7.3), and concluded that the best resolved stress parameters are not affected by the use of a 3 D velocity model, similar to the present result. Thus it seems that the 1 D velocity models would be good enough for determining accurate FMS and inverting for the stress tensors Spatial and Temporal Variations of the Stress Field [13] We adopted a moving window method to study the spatial variations of the stress field in the 2008 IMEQ source area. We set up a grid net with a lateral grid interval of 0.1 in the study area (Figure 7). Events within 15 km around each grid 7of13

8 Figure 8. The orientations of the P axes and R values plotted along the profile parallel (AA ) and normal (BB ) to the fault as shown in Figure 7. The optimal values and their uncertainties are shown as solid circles and short bars, respectively. Note that the plunges are positive when the axis dips toward WNW and negative when it dips toward ESE. The shaded portions denote the corresponding probability density functions by bootstrap resamplings. Major earthquakes are shown as stars. node are used to invert for the stress tensor around that grid node (Figure S2). The number of events used for the stress inversion for each node is larger than 20. For the grid nodes near the main shock, more than 130 events were included for each inversion. Figures 7 and 8 show the spatial distribution of the inverted maximum stress s 1 and the stress ratio R. The most notable feature is that the maximum stress s 1 varies laterally in the source area; it generally dips toward ESE in the southern part while toward WNW in the northern part (Figure 8b). The stress ratio R also shows lateral variations. In the southern part where the focal depths are generally shallower, the R values are smaller due to the smaller vertical components (s v or s 3 ) of the stress tensors (Figure 8c). Along the profile (BB ) that is perpendicular to the fault, the R value also decreases toward the updip direction (ESE) of the fault due to the smaller depth and the reduced s 3 (Figure 8f). [14] In order to detect the possible temporal variation of the tectonic stress in the 2008 IMEQ source area, we divided our data set to several subsets in nonoverlapping time intervals (Figures 9 and S3). At least 20 events are included in most of the subsets, except for the interval (N1) before the main shock that contains 15 events. By considering the spatial stress variations as shown above, the northern (200 events, N1 N10) and southern (45 events, S1 S2) parts are analyzed separately (Figure 9). Figure 10 summarizes the orientations of the maximum principle stress s 1 and the stress ratio R versus time. While the stress ratio is rather scattered, s 1 shows notable temporal variations. 8of13

9 Figure 9. (a) Distribution of the 12 groups (N1 N10, S1 S2) of earthquakes that are used to study the temporal variations of the stress field. The numbers in the brackets show the number of earthquakes for each group. (b) An enlarged version of the yellow area in Figure 9a. In particular, there is a 24 rotation for the plunge of s 1 (N4) in the northern region (note that the resolution is 6 ), though they overlap in the 95% confidence region. Note that the plunges were measured relative to the WNW direction. The stress rotation has been found following large earthquakes in some regions, e.g., in southern California following the 1992 Landers earthquake (M7.3) and the1994northridgeearthquake(m6.7)[zhao et al., 1997; Hardebeck and Hauksson, 2001]. It was proposed that the stress rotation (D) depends on the orientation of the preearthquake s 1 relative to the fault () and the ratio of the earthquake stress drop to the background deviatoric stress level (Dt/t) (Figure 11) [Hardebeck and Hauksson, 2001]. In the case of the 2008 IMEQ, is estimated to be in the range of 30 to 40 [Abe et al., 2008]. As long as the stress drop is positive (Dt/t >0),theD is generally negative (D < 0) (Figure 11), i.e., s 1 rotates closer to the fault and its plunge becomes positive (Figures 10 and 11). There seems also some temporal variations in the azimuth of s 1 following the main shock (Figure 10a). In fact, most earthquakes (main shock and aftershocks) in the 2008 IMEQ area are not pure dip slip events; they have considerable strike slip components (Figure 4). Therefore the azimuths of s 1 may have also rotated temporally. In the southern part, however, s 1 rotates 24 far away from the fault plane (S2 in Figure 10). Takada et al. [2009] constructed a segmented fault model to explain the coseismic displacement of the 2008 IMEQ detected by ALOS/PALSAR. In their model, the plunge of the fault in the southern part is as large as 55. This large plunge will induce the positive stress rotation (D > 0) following the model of Hardebeck and Hauksson [2001] (Figure 11), which explains well the present result. [15] Nevertheless, the stress rotations (D) in hours immediately following the main shock (N2 and N3) in the northern part are positive (the s 1 rotates away from the fault) (Figure 10). The stress ratio R in that period also tends to be larger, suggesting a 9of13

10 and the stress tends to increase (Dt/t < 0) immediately after the earthquake. [16] But the stress rotation due to the negative stress drop is no more than 10 for in the range of (Figure 11), which can account for only part of what we obtained (18 ). Additional factors are needed to explain the immediate stress rotation following the main shock. The 2008 IMEQ is located near the volcanic front that runs through in the middle of NE Japan. Seismic tomography revealed significant low velocity anomalies in the lower crust and uppermost mantle beneath the source area [Wang et al., 2008; Okada et al., 2010; Cheng et al., 2011], which may reflect the arc magma and fluids ascending from the mantle wedge. Sibson [2009] proposed that the fluids in the middle crust below the seismogenic layer in NE Japan are overpressured. The preearthquake fluid pressure approaches the lithostatic pressure, while it falls suddenly to the hydrostatic pressure after the earthquake ruptures [Sibson, 2009]. The fluids may occasionally intrude into the seismogenic layer and are expected to transport through the faults and cracks in the source area. The sudden release of the fluid pressure as well as its concurrent transportation will surely change the stress status in the source area [Hardebeck and Hauksson, 1999; Cappa et al., 2007]. Thus, the crustal fluids are important in understanding the stress field in the 2008 IMEQ area, though the details are still unclear. Figure 10. Temporal variations of the (a) azimuth and (b) plunge of the P axis and (c) the R value for the northern part (blue bars with gray shades) and the southern part (red bars with green shades) of the study area. The results are plotted at the beginning of each time interval marked in Figure 10b. The other labeling is the same as that in Figure 8. larger s 1 than that before the main shock, i.e., negative stress drop (Dt/t < 0). If this is true, then the positive stress rotation (D > 0) will occur for the of following the synthetic model (Figure 11). The negative stress drop has been physically simulated for the 2008 IMEQ [Ando and Okuyama, 2010] and is also confirmed for the 2004 mid Niigata earthquake [Miyatake et al., 2008] and the 2005 west off Fukuoka earthquake [Horikawa, 2006]. When an earthquake occurs, a corresponding asperity ruptures with a large coseismic slip and the stress on the asperity drops. But in the areas surrounding the asperity, the coseismic slip is very small [Ando and Okuyama, 2010; Suzuki et al., 2010]. The rupture of these areas will be delayed 4.3. Tectonic Implications [17] Despite the discrepancies of the stress tensor for different velocity models and its temporal and spatial variations, the general orientation of the maximum principal stress s 1 (P axis) in the 2008 IMEQ area is consistent with both the motion of the Pacific plate relative to NE Japan and the horizontal principal strain in this area derived from the GPS observations [Seno, 1999; Miura et al., 2004]. Our results follow the general pattern that the regional horizontal stress orientations within the overriding plate are consistent with either the relative or the absolute plate motions, which indicates that the plate boundary forces dominate the intraplate stress distribution [Zoback, 1992; Heidbach et al., 2010]. [18] The stress rotation can be used to estimate the background deviatoric stress level and further constrain the strength of the fault. The Dt/t for the 2008 IMEQ is 0.8 for a 24 rotation (Figure 11). As the average stress drop (Dt) is generally smaller than 20 MPa [Asano and Iwata, 2011; Hok and Fukuyama, 2011], the background deviatoric stress 10 of 13

11 Figure 11. A cartoon to illustrate the temporal stress rotation in the 2008 IMEQ area. (a) The principle stresses before the main shock together with the orientation of the fault. (b) The stresses after the main shock. (c) The curves denote the theoretic stress rotation as a function of angle of s 1 to fault () and the stress drop (Dt/t). The red patches denote the condition of the 2008 IMEQ and that of the southern segment of the earthquake related fault. The blue patch with question mark denotes the possible condition immediately following the main shock when negative stress drop occurs. level is 25 MPa. Measurements in deep boreholes reveal the deviatoric stress on the order of 100 MPa at seismogenic depths [e.g., Townend and Zoback, 2000]. Thus the friction coefficient m in the study area is only , which is obviously lower than the typical laboratory values of [Townend and Zoback, 2000]. These results suggest that the fault of the 2008 IMEQ is very weak, arguing against the concept of strong crust [Townend and Zoback, 2000; Scholz, 2002]. As mentioned above, the 2008 IMEQ is located near the active volcanoes. High temperature arc magma and fluids exist in the lower crust and uppermost mantle [e.g., Cheng et al., 2011; Zhao et al., 2011], which may intrude upwards and weaken the overlying crust. Thus the friction coefficient on the fault can be reduced. [19] At the grid node (39.0, ) right beside the IMEQ main shock where 152 events were used in the stress tensor inversion, the orientations of the maximum stress s 1 are rather convergent (Figure S2). The minimum stress s 3 aligns subvertically in a NNE SSW belt perpendicular to s 1, as indicated by the FMS. The result indicates that, with a common maximum principal stress, the occurrence of earthquakes is somewhat random [Scholz, 2002]. NE Japan is tectonically one of the most active regions in the world. The frequent occurrence of earthquakes and volcanic eruptions has produced many faults and cracks which become the weak planes in 11 of 13

12 the crust. The stress necessary to reactivate these faults and cracks is much lower compared to the stress necessary to generate a new fault [Scholz, 2002]. Thus, as the crustal stress increases, the ruptures first occur along the preexisting, randomly orientated faults and cracks [Gephart and Forsyth, 1984; Scholz, 2002], and so the minimum stress s 3 tends to align in a belt perpendicular to s 1 as observed. 5. Conclusions [20] We determined the FMS of 191 earthquakes most of which are aftershocks in the 2008 IMEQ source area and inverted for the stress tensors using a large number of P wave polarity data generated by these earthquakes. The FMS and stress tensors are determined by using both the 1 D and 3 D velocity models, which show dominant WNW dipping horizontal P axis and subvertical T axis. The orientation of the P axis (s 1 ) is consistent with the motion of the Pacific plate relative to NE Japan, which indicates that the plate boundary forces dominate the intraplate stress distribution. The differences between the results obtained by using the 1 D and 3 D velocity models are not obvious, and they can be neglected as compared with the discrepancies caused by the different methods. A 1 D velocity model seems accurate enough for determining the FMS and the stress field at least in the present study region. [21] Significant rotation of the maximum principal stress s 1 occurred in the 2008 IMEQ sequence. The principal stress s 1 first rotated closer to the fault (D < 0) and then rotated back to a horizontal direction after three months of the main shock. The stress rotation was largely affected by the angle between the orientations of the fault and s 1 () before the earthquake and the stress drop (Dt/t) caused by the earthquake. In hours immediately following the main shock, s 1 rotated away from the fault (D > 0). The rotation could be affected by the fluid transportation and the negative stress drop (Dt/t < 0) in the 2008 IMEQ area. The stress rotation showed spatial variations. In the northern part where the main shock was located, s 1 rotated closer to the fault (D < 0), which is consistent with the temporal stress rotation following the main shock. In the southern part, in contrast, s 1 rotated away from the fault (D > 0) due to the large plunge of the fault plane there. [22] The background deviatoric stress level is estimated to be only 25 MPa, which corresponds to a small friction coefficient m of in the study area. The result argues that the faults in NE Japan are weak. The fault weakening may be caused by the high temperature arc magma and fluids in the lower crust and uppermost mantle that may intrude upward into the shallower crust. Acknowledgments [23] We thank the data centers of Hi net and JMA seismic networks for providing the P wave polarity and arrival time data used in this study. This work was supported partially by a grant (Kiban A ) to D. Zhao from Japan Society for the Promotion of Science, a grant ( ) from the National Natural Science Foundation of China, and by the Scientific Research Foundation of Graduate School of Nanjing University. D. Zhao and Z. Huang were also supported by the Global COE program of Tohoku University. Comments from T. Becker (Editor) and an anonymous reviewer have greatly improved the manuscript. Most of figures were made by using GMT [Wessel and Smith, 1998]. References Abe, S., H. Saito, H. Sato, S. Koshiya, N. Kato, K. Shiraishi, and T. Kawanaka (2008), Multidisciplinary seismic survey across the Kitakami Lowland, Northeast Japan, paper presented at Japan Geoscience Union Meeting 2008, Chiba City, Japan. Ando, R., and S. Okuyama (2010), Deep roots of upper plate faults and earthquake generation illuminated by volcanism, Geophys. Res. Lett., 37, L10308, doi: /2010gl Asano, K., and T. Iwata (2011), Characterization of stress drops on asperities estimated from heterogeneous kinematic slip model for strong motion prediction for inland crustal earthquakes in Japan, Pure Appl. Geophys., 168, , doi: /s y. Cappa, F., Y. Guglielmi, and J. Virieux (2007), Stress and fluid transfer in a fault zone due to overpressures in the seismogenic crust, Geophys. Res. Lett., 34, L05301, doi: / 2006GL Cheng, B., D. Zhao, and G. Zhang (2011), Seismic tomography and anisotropy in the source area of the 2008 Iwate Miyagi earthquake (M 7.2), Phys. Earth Planet. Inter., 184, , doi: /j.pepi Dziewonski, A. M., and J. H. Woodhouse (1983), An experiment in systematic study of global Seismicity: Centroid moment tensor solutions for 201 moderate and large earthquakes of 1981, J. Geophys. Res., 88, , doi: / JB088iB04p Fukahata, Y., Y. Fukushima, and M. Arimoto (2008), Slip distribution of the 2008 Iwate Miyagi Nairiku, Japan, earthquake inverted from PALSAR data, Eos Trans. AGU, 89(53), Fall Meet. Suppl., Abstract S23B Fukuyama, E., M. Ishida, D. Dreger, and H. Kawai (1998), Automated seismic moment tensor determination by using on line broadband seismic waveforms (in Japanese with English abstract), J. Seismol. Soc. Jpn., 51, Gephart, J., and D. Forsyth (1984), An improved method for determining the regional stress tensor using earthquake focal mechanism data: Application to the San Fernando 12 of 13

13 earthquake sequence, J. Geophys. Res., 89, , doi: /jb089ib11p Hardebeck, J., and E. Hauksson (1999), Role of fluids in faulting inferred from stress field signatures, Science, 285, , doi: /science Hardebeck, J., and E. Hauksson (2001), Crustal stress field in southern California and its implications for fault mechanics, J. Geophys. Res., 106, 21,859 21,882, doi: / 2001JB Heidbach, Q., M. Tingay, A. Barth, J. Reinecher, D. Kurfeb, and B. Muller (2010), Global crustal stress pattern based on the World Stress Map database release 2008, Tectonophysics, 482, 3 15, doi: /j.tecto Hok, S., and E. Fukuyama (2011), A new BIEM for rupture dynamics in half space and its application to the 2008 Iwate Miyagi Nairiku earthquake, Geophys. J. Int., 184, , doi: /j x x. Horikawa, H. (2006), Rupture process of the 2005 west off Fukuoka prefecture, Japan, earthquake, Earth Planets Space, 58, Horiuchi, S., G. Rocco, and A. Hasegawa (1995), Discrimination of fault planes from auxiliary planes based on simultaneous determination of stress tensor and a large number of fault plane solutions, J. Geophys. Res., 100, , doi: /94jb Huang, Z., D. Zhao, and L. Wang (2011), Seismic heterogeneity and anisotropy of the Honshu arc from the Japan trench to the Japan Sea, Geophys. J. Int., 184, , doi: / j x x. McKenzie, D. (1969), The relation between fault plane solutions for earthquakes and the directions of the principal stresses, Bull. Seismol. Soc. Am., 59, Michael, A. (1987), Stress rotation during the Coalinga aftershock sequence, J. Geophys. Res., 92, , doi: /jb092ib08p Miura, S., T. Sato, A. Hasegawa, Y. Suwa, K. Tachibana, and S. Yui (2004), Strain concentration zone along the volcanic front derived by GPS observations in NE Japan arc, Earth Planets Space, 56, Miyatake, T., A. Kato, K. Hikima, and T. Kimura (2008), The effects of negative stress drop on the fault rupture the 2004 mid Niigata (Chuetsu), Japan, earthquake, Bull. Earthquake Res. Inst. Univ. Tokyo, 83, Okada, T., N. Umino, and A. Hasegawa (2010), Deep structure of the Ou mountain range strain concentration zone and the focal area of the 2008 Iwate Miyagi Nairiku earthquake, NE Japan Seismogenesis related with magma and crustal fluid, Earth Planets Space, 62, , doi: / eps Scholz, C. H. (2002), The Mechanics of Earthquakes and Faulting, 496 pp., Cambridge Univ. Press, New York. Seno, T. (1999), Syntheses of the regional stress fields of the Japanese islands, Isl. Arc, 8, 66 79, doi: /j x. Sibson, R. H. (2009), Rupturing in overpressured crust during compressional inversion the case from NE Honshu, Japan, Tectonophysics, 473, , doi: /j.tecto Suzuki, W., S. Aoi, and H. Sekiguchi (2010), Rupture process of the 2008 Iwate Miyagi Nairiku, Japan, earthquake derived from near source strong motion records, Bull. Seismol. Soc. Am., 100, , doi: / Takada, Y., T. Kobayashi, M. Furuya, and M. Murakami (2009), Coseismic displacement due to the 2008 Iwate Miyagi Nairiku earthquake detected by ALOS/PALSAR: Preliminary results, Earth Planets Space, 61, e9 e12. Toda, S., and T. Maruyama (2009), Complexity and unpredictability of surface ruptures associated with reverse faulting earthquake in volcanic region: An example from the 2008 Mw 6.9 Iwate Miyagi Japan earthquake, Eos Trans. AGU, 90(52), Fall Meet. Suppl., Abstract T24B 02. Townend, J., and M. D. Zoback (2000), How faulting keeps the crust strong, Geology, 28, , doi: / (2000)28<399:hfktcs>2.0.co;2. Wang, Z., Y. Fukao, S. Kodaira, and R. Huang (2008), Role of fluids in the initiation of the 2008 Iwate earthquake (M7.2) in northeast Japan, Geophys. Res. Lett., 35, L24303, doi: /2008gl Wessel, P., and W. Smith (1998), New improved version of the Generic Mapping Tools released, Eos Trans. AGU, 79, 579, doi: /98eo Yukutake, Y., Y. Iio, H. Katao, and T. Shibutani (2007), Estimation of the stress field in the region of the 2000 Western Tottori earthquake: Using numerous aftershock focal mechanisms, J. Geophys. Res., 112, B09306, doi: / 2005JB Zhao, D., A. Hasegawa, and S. Horiuchi (1992), Tomographic imaging of P and S wave velocity structure beneath northeastern Japan, J. Geophys. Res., 97, 19,909 19,928, doi: /92jb Zhao, D., H. Kanamori, and D. Wiens (1997), State of stress before and after the 1994 Northridge earthquake, Geophys. Res. Lett., 24, , doi: /97gl Zhao, D., S. Yu, and E. Ohtani (2011), East Asia: Seismotectonics, magmatism and mantle dynamics, J. Asian Earth Sci., 40, , doi: /j.jseaes Zoback, M. (1992), First and second order patterns of stress in the lithosphere: The world stress map project, J. Geophys. Res., 97, 11,703 11,728, doi: /92jb of 13