Depth-dependent stress field in and around the Atotsugawa fault, central Japan, deduced from...

Depth‐dependent stress field in and around the Atotsugawa fault,central Japan, deduced from microearthquake focal mechanisms:Evidence for localized aseismic deformation in the downwardextension of the fault

Kazutoshi Imanishi,1 Yasuto Kuwahara,1 Tetsuya Takeda,2 Takashi Mizuno,1,3 Hisao Ito,4

Kiyoshi Ito,5,6 Hiroo Wada,5 and Yoshikatsu Haryu2,6,7

Received 30 July 2010; revised 2 November 2010; accepted 22 November 2010; published 21 January 2011.

[1] Focal mechanisms have been determined from P wave polarity data as well as bodywave amplitudes for 154 microearthquakes that occurred around the Atotsugawa fault incentral Japan between 2002 and 2004. While we found many microearthquakes with apure strike‐slip mechanism that is similar to the faulting style of the Atotsugawa fault, aconsiderable number of microearthquakes with reverse faulting components are alsooccurring. Most of the P axes are horizontal and oriented in the WNW‐ESE direction,which conforms to the general tectonic trend in this area. In contrast, the T axes have awide range of plunge, suggesting that reverse‐faulting‐type earthquakes as well asstrike‐slip ones are occurring. The most conspicuous feature in the focal mechanismdistribution is the depth dependence, where shallow earthquakes are primarily reversefaulting and strike‐slip earthquakes become predominant with depth. A stress tensorinversion reveals that the shallower part is characterized by a mixture of reverse andstrike‐slip faulting regimes and that a pure strike‐slip faulting regime appears only aroundthe bottom of the seismogenic zone. Together with other geophysical observationalevidence for the fault, we suggest that the existence of a localized aseismic deformationbelow the Atotsugawa fault is the simplest scenario that can explain the observed stressfields. This scenario provides a stress accumulation mechanism of disastrous shallowinland earthquakes, in which the localized aseismic deformation accumulates stress ontothe fault plane in the seismogenic zone during an earthquake cycle and the mainshock would occur when the failure stress is reached on the fault.

Citation: Imanishi, K., Y. Kuwahara, T. Takeda, T. Mizuno, H. Ito, K. Ito, H. Wada, and Y. Haryu (2011), Depth‐dependentstress field in and around the Atotsugawa fault, central Japan, deduced from microearthquake focal mechanisms: Evidence forlocalized aseismic deformation in the downward extension of the fault, J. Geophys. Res., 116, B01305,doi:10.1029/2010JB007900.

1. Introduction

[2] It is widely known that tectonic stress graduallyaccumulates in locked portions of plate boundaries, owingto relative plate motion, and that large interplate earthquakes

occur when the accumulated stress exceeds the strength ofthe locked portions. In particular, recent seismic and geo-detic observations of small repeating earthquakes [e.g.,Nadeau and McEvilly, 1999; Igarashi et al., 2003; Uchidaet al., 2007] and of episodic tremor and slip [e.g., Dragertet al., 2001; Obara, 2002; Shelly, 2010] contribute to adeeper understanding of the generation mechanism of largeinterplate earthquakes and provide a new way to monitorstress buildup at plate boundaries. In contrast, understandingof the generation mechanism of disastrous shallow inlandearthquakes (intraplate earthquakes) has advanced slowly[e.g., Stein and Mazzotti, 2007]. Scholz [2002] classifiedshallow intraplate earthquakes into two types. The first typeoccurs in broad areas near plate boundaries (e.g., inlandearthquakes in Japan and New Zealand) or in diffuse plateboundaries (e.g., the Basin and Range provinces of westernNorth America, the Andes, and the Himalayan‐TibetanPlateau). The second type occurs in stable continental areas

1Geological Survey of Japan, National Institute of Advanced IndustrialScience and Technology, Tsukuba, Japan.

2National Research Institute for Earth Science and Disaster Prevention,Tsukuba, Japan.

3Now at Schlumberger Kabushiki Kaisha Integration Center, Sagamihara,Japan.

4Center for Deep Earth Exploration, Japan Agency for Marine‐EarthScience and Technology, Yokohama, Japan.

5Disaster Prevention Research Institute, Kyoto University, Uji, Japan.6Now at Hanshin Consultants Co., Ltd., Osaka, Japan.7Association for the Development of Earthquake Prediction, Tokyo,

Japan.

Copyright 2011 by the American Geophysical Union.0148‐0227/11/2010JB007900

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 116, B01305, doi:10.1029/2010JB007900, 2011

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http://dx.doi.org/10.1029/2010JB007900

and is considered not to be associated with plate boundaries.The most famous example of this type is the New Madridseismic zone in the central United States. During the last twodecades, the aseismic deformation zone beneath the seis-mogenic layer came to be considered as a key for theoccurrence of both types of large intraplate earthquakes, andsome models incorporating a local stress accumulationmechanism have been proposed [e.g., Liu and Zoback, 1997;Kenner and Segall, 2000; Iio and Kobayashi, 2002a;Hasegawa et al., 2005]. Recent studies on the first type ofintraplate earthquakes in Japan also support the existence ofan aseismic deformation zone beneath the seismogeniclayer, on the basis of high‐quality geophysical data andnumerical simulation [Iio and Kobayashi, 2002b; Mizunoet al., 2005; Kawanishi et al., 2009; Ando and Okuyama,2010]. However, such observational results are still rareand the style of deformation (i.e., localized or widely dis-tributed in volume) still remains open to discussion. Furtherobservational constraints at various active faults are neededto better understand the stress accumulation mechanism ofshallow inland earthquakes.[3] The Atotsugawa fault (Figure 1) in central Japan is

probably the best field to study the stress accumulationmechanism of large inland earthquakes in Japan because ofits high level of seismicity, high deformation rate, simplefault geometry, and the existence of many related studies[e.g., Ito and Wada, 2002; Hirahara et al., 2003; Mizunoet al., 2005; Kato et al., 2006]. Here, the Atotsugawa faultis categorized as being one of the first type of intraplateearthquakes, as defined by Scholz [2002]. Mizuno et al.[2005] suggested a possibility that localized aseismic slipoccurs below the Atotsugawa fault by analyzing shearwave splitting across the fault. They found that the anglebetween the polarization direction of the leading shearwave (LSPD) and the strike of the fault decreases withapproach to the fault. On the basis of the assumption thatLSPD corresponds to the direction of the maximum hori-zontal stress [e.g., Crampin, 1978], they concluded that thespatial variation of LSPD can be explained by incorpo-rating aseismic slip along the deep extension of the faultthrough numerical modeling. However, the assumptionadopted by Mizuno et al. [2005] is not necessarily appli-cable each time, because LSPD may be affected bystructural anisotropies such as fault zone fabric [e.g., Zinke andZoback, 2000;Balfour et al., 2005;Boness and Zoback, 2006].[4] If the deep extension of the fault slips aseismically, it

might be possible to detect a depth dependence in the stressfield caused by the stress concentration at the aseismic slipfront. Among various stress indicators such as shear wavesplitting, strain data, and others, earthquake focal mechan-isms are thought to be the most effective means to detect thedepth dependence, because they constrain the stress field atdepths where earthquakes are actually occurring. Focalmechanisms are ordinarily determined from P wave polaritydata. In the case of microearthquakes, however, it is difficultto obtain a unique focal mechanism solution, because thenumber of stations detecting events decreases and theirazimuthal coverage becomes poor. One approach to over-coming this problem is to use S/P amplitude ratios [e.g.,Kisslinger, 1980; Julian and Foulger, 1996; Hardebeck andShearer, 2003] or absolute P and S amplitudes [e.g., Slunga,1981; Nakamura et al., 1999; Igarashi et al., 2001]. These

approaches have been successfully applied to micro-earthquakes and have been shown to be effective even if thenumber of P wave polarities is insufficient and the obser-vation station coverage is poor. We prefer the latterapproach for microearthquakes, because there are manyseismograms in which only S wave amplitude can be used inthe analysis owning to poor signal‐to‐noise ratio of P waveamplitude.[5] In this study, we investigate the depth dependence of

the stress field along the Atotsugawa fault plane fromnumerous focal mechanism solutions. Because most of theearthquakes occurring in and around the Atotsugawa faultare smaller than 2 in magnitude, we carried out a densetemporary seismic observation around the fault and deter-mined the focal mechanisms using P wave polarity data aswell as absolute body wave amplitudes. On the basis of theestimated stress fields and other observational evidence, wediscuss a possible stress accumulation mechanism of theAtotsugawa fault.

2. Seismological Setting

[6] A number of active Quaternary faults are concentratedin the study region. They are conjugate sets, where oneconsists of right‐lateral strike‐slip faults trending in an ENE‐WSW to NE‐SW direction and the other consists of left‐lateral strike‐slip faults trending in a NNW‐SSW to NW‐SEdirection (Figure 1). The Atotsugawa fault system is locatedbetween Mt. Tateyama and Mt. Hakusan and is composed ofthree near‐vertical right‐lateral strike‐skip faults, the Atot-sugawa, Ushikubi, and Mozumi‐Sukenobu faults [e.g.,Takeuchi et al., 2003]. The Atotsugawa fault is traced geo-morphologically for about 69 km with a strike of N60°E.The Ushikubi fault, with a length of about 54 km, runsalmost parallel to the Atotsugawa fault about 6 km north ofit. The Mozumi‐Sukenobu fault is about 23 km long andbranches from the Atotsugawa fault at its eastern termi-nation. Paleoseismic trenching data suggest that recurrenceintervals of the Atotsugawa, Ushikubi, and Mozumi‐Sukenobu faults are about 2400 ± 100 years (Headquartersfor Earthquake Research Promotion, Long‐term evaluationof the Atotsugawa fault (in Japanese), 2004, available athttp://www.jishin.go.jp/main/chousa/04sep_atotugawa/index.htm), 6050 ± 1050 years (Headquarters for Earth-quake Research Promotion, Long‐term evaluation ofthe Ushikubi fault (in Japanese) 2005, available at http://www.jishin.go.jp/main/chousa/05mar_ushikubi/index.htm),and 13,500 ± 6300 years [Takeuchi et al., 2003], respec-tively. The most recent destructive earthquake in this areais the 9 April 1858 Hietsu earthquake (M7.0), which wasconsidered to be a faulting on the Atotsugawa fault[Matsuda, 1966]. Takeuchi et al. [2003] suggested that theMozumi‐Sukenobu fault was also active in the 1858 event.The average horizontal fault slip rate of the Atotsugawafault inferred from the topographical offsets is 2–3 mm/yr(Headquarters for Earthquake Research Promotion, Long‐term evaluation of the Atotsugawa fault (in Japanese),2004, available at http://www.jishin.go.jp/main/chousa/04sep_atotugawa/index.htm). This slip rate falls within thefastest category in Japan [Research Group for Active Faultsof Japan, 1991]. The average horizontal fault slip rates of

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the Ushikubi and Mozumi‐Sukenobu faults are consideredto be less than that of the Atotsugawa fault.[7] On the basis of his analyses of triangulation data

compiled over the past 100 years,Hashimoto [1990] detecteda zone with high strain rates fromNiigata to the Kinki district,central Japan. This zone, with high strain rates, was alsoclearly identified in the recent observations from dense GPSarrays by the Geographical Survey Institute (Figure 1), and itwas named the Niigata‐Kobe Tectonic Zone (NKTZ) [Sagiya

et al., 2000]. The strain rate in the NKTZ was determined toexceed 10−7/yr, which is tenfold that of the average in Japan[Sagiya et al., 2000; Mazzotti et al., 2001]. There are manyactive Quaternary faults within this zone and large inlandearthquakes have also occurred intensively. The Atotsugawafault system is one of the active faults in the NKTZ.[8] Seismicity in the study region was investigated in

detail byMikumo et al. [1988], Ito and Wada [2002], and Itoet al. [2007]. Figure 1 shows seismicity in the study area

Figure 1. Overview of the Atotsugawa fault system and adjacent regions, central Japan. Active faults arerepresented by red lines and AGF represents the Atotsugawa fault. Two near‐vertical right‐lateral strike‐slip faults named the Ushikubi fault (UKF) and the Mozumi‐Sukenobu fault (MSF) are running near theAtotsugawa fault. Black circles represent earthquake locations shallower than 20 km determined by theJapan Meteorological Agency (JMA) during the period from January 2000 to December 2009. Earthquakelocations in the X‐Y cross section are shown at the bottom. The thick orange line in the top left insetshows the concentrated deformation zone (Niigata‐Kobe Tectonic Zone, NKTZ).


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during the period from January 2000 to December 2009 asroutinely determined by the Japan Meteorological Agency(JMA). A linear distribution of high seismic activity canbe recognized along the Atotsugawa fault, which was firstreported by Wada and Kishimoto [1974]. Linear trends ofseismicity along the Mozumi‐Sukenobu and Ushikubi faultscan also be identified, although the activity on the faults islower than that on the Atotsugawa fault. A cross sectionalong the Atotsugawa fault system is shown at the bottom ofFigure 1. A spatial nonuniformity of seismicity is clearlyrecognized, in which seismicity rate on the northeast sidenear the central part of the fault is significantly low. Theseismicity cutoff depth decreases gradually from the centralpart of the Atotsugawa fault to both ends. This concaveshape could be explained by the change in thermal structure,because there are active volcanoes at both ends of theAtotsugawa fault [Ito and Wada, 2002]. The cross sectionsperpendicular to each fault indicate that the seismicity alongthe faults occurs in a nearly vertical zone without inter-secting each other [Ito et al., 2007].[9] Figure 2 shows focal mechanisms of earthquakes that

occurred during the period from October 1997 to December2009 at depths shallower than 20 km. These mechanismswere routinely determined by the JMA on the basis of theobservation of P wave first motion polarities, which arelisted in the JMA earthquake catalog since October 1997.JMA generally reports focal mechanisms of earthquakeswith magnitudes greater than 3. Most of the earthquakes

occurring in the area are strike‐slip faulting, reverse faulting,or a mixture of these. The JMA earthquake catalog isinsufficient to discuss the general tectonic stress prevailingover the Atotsugawa fault system, since there are few focalmechanisms near the fault system. Mikumo et al. [1988] andKoizumi et al. [1993] reported about 10 focal mechanisms ofearthquakes with magnitudes greater than 3 that occurredduring the period from 1977 to 1984 and from 1978 to 1990,respectively. Most of these moderate size earthquakes havealmost a strike‐slip mechanism with a NWN‐ESE orienta-tion of P axis, which is similar to the faulting style of theAtotsugawa fault system. There are also a few events with areverse‐faulting or normal‐faulting component. Togetherwith a temporary seismic observation, focal mechanisms ofearthquakes smaller than M3 could be determined by usingthe P wave first motions [Wada et al., 2003; Ito et al., 2007;Katsumata et al., 2010]. The result shows that earthquakeswith reverse‐faulting and normal‐faulting components arecertainly occurring around the Atotsugawa fault system,implying a spatial variation of the local stress field.

3. Data

[10] We conducted temporary observations around theAtotsugawa fault at 8 stations from July to November 2002,23 stations from May to October 2003, and 14 stations fromMay to October 2004 (Figure 3). Seismometers employed inthe temporal observations were three‐component velocitygeophones having a natural frequency of 2 Hz (L22‐E,Mark Product Inc.) and fixed on a hard rock surface withPlaster of Paris. The events were recorded on REF TEK 72A(Refraction Technology, Inc.) and LS7000 data loggers(HAKUSAN Corp.) with GPS clocks at sample rates of250 or 200 Hz in offline continuous mode. In Figure 3, wealso show the distribution of the permanent stations usedin our study; these are operated by the Disaster PreventionResearch Institute (DPRI, Kyoto University), EarthquakeResearch Institute (ERI, University of Tokyo), NagoyaUniversity, JMA, National Research Institute for EarthScience and Disaster Prevention (NIED), and GeologicalSurvey of Japan (GSJ, National Institute of Advanced Indus-trial Science and Technology). Each station is equipped with aset of three‐component velocity geophones having a naturalfrequency of 1 or 2 Hz. Seismometers operated by the NIEDand GSJ are installed at the bottom of a borehole at a depthof several hundred meters. Figure 4 shows an example ofvertical component seismograms recorded by temporal sta-tions (black lines) and surrounding permanent stations (shadedlines).[11] We analyzed 481 earthquakes that occurred in and

around the Atotsugawa fault during the period of the tem-porary observations. The hypocenters were determined byapplying a maximum‐likelihood estimation algorithm[Hirata and Matsu’ura, 1987] to manually picked arrivaltimes. The P wave velocity model used in our investigationis shown in Figure 5, which is the same as that of Ito andWada [2002]. The S wave models are assumed by scalingthe P wave velocities by a factor of 1/

ffiffiffi

3p

. The P and S wavearrival times were identified manually using the WIN sys-tem developed by Urabe and Tsukada [1991]. We firstlocated the hypocenters of all events that occurred in 2003without station corrections. We then relocated these events

Figure 2. Spatial distribution of focal mechanism solutionsshallower than 20 km determined by the JMA during theperiod from October 1997 to December 2009. Focal mechan-isms are projected onto the lower hemisphere using the equal‐area projection. Mj denotes the magnitude reported by theJMA. The Mj 4.1 earthquake of 24 March 2005 is shownby an arrow.


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by introducing the station correction, which was obtainedusing the average of the differences between observed andtheoretical travel times at each station. We repeated theabove procedure three times and obtained final locations aswell as station corrections at individual stations. The root‐mean square (RMS) values of the residuals decreased from0.14 to 0.07 s for the P wave and from 0.20 to 0.11 s for theS wave. The locations of all events that occurred in 2002were determined using the same velocity structures andstation corrections. For the 2004 observations, there were anadditional five stations that were not deployed in 2003. Wefirst determined the locations of all events that occurred in2004 on the basis of the same velocity structures and stationcorrections without these five stations. We then calculatedthe station corrections at those five stations by averaging thedifferences between the observed and theoretical traveltimes. Finally, we relocated the 2004 events incorporating

the newly determined station corrections. Magnitudesranged from −1.2 to 3.0, which were determined from themaximum velocity amplitude data of seismograms by themethod of Watanabe [1971]. The average spatial errorscalculated by the maximum‐likelihood estimation algorithm[Hirata and Matsu’ura, 1987] were 122 m in the horizontaldirection and 235 m in depth, which is sufficient for thepurpose of the present study. There are 259 events amongour data set that were used for a 3‐D seismic travel timetomography study [Takeda et al., 2004]. We compared thelocations determined in the present study with those deter-mined by the 3‐D model and found that the average spatialdifferences are 274 m in the horizontal direction and 376 min depth. This comparison demonstrates that our procedurebased on the 1‐D velocity model with station correctionsgives reasonable estimates of earthquake locations. Thegeneral feature of the seismicity along the Atotsugawa fault

Figure 3. Distribution of temporary and permanent seismic stations used for the relocation and focalmechanism determination. The shaded circles represent earthquake locations determined in this study.Active faults are represented by solid lines.


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(e.g., a concave shape of the cutoff depth of seismicity and avery low seismicity rate on the northeast side) is the same asthose in previous studies [e.g., Mikumo et al., 1988; Ito andWada, 2002].

4. FocalMechanism Solution ofMicroearthquakesUsing Body Wave Amplitudes

[12] In the present study, we determined focal mechanismsolutions of microearthquakes using absolute P and SH

wave amplitudes as well as P wave polarity. The sameapproach was used by Imanishi et al. [2006a, 2006b] andImanishi and Kuwahara [2009], and was shown to beeffective for microearthquakes. We analyzed earthquakeshaving at least eight P wave polarities. After correcting theinstrumental response, we determined the spectral levels andcorner frequencies of the spectra by fitting the w2 model[Boatwright, 1978] with an attenuation correction. Thespectral levels for lower frequencies were used as observedamplitudes. Theoretical amplitudes were calculated from the

Figure 4. Example of vertical component seismograms recorded by temporary stations (black lines) andsurrounding permanent stations (shaded lines).


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far‐field solutions for a shear point‐source dislocation in ahomogeneous infinite medium, with corrections of theincident angles at the surface and geometrical spreading.The best fit solution of each event was determined byminimizing the residual between the observed and theoret-ical amplitudes, where a grid search approach was appliedfor strike, dip, and slip angles at 5° intervals. We firstdetermined focal mechanisms and seismic moments of allevents that occurred in 2003. We then redetermined them byintroducing the amplitude station correction, which wasobtained using the logarithmic average of the ratios betweenthe theoretical and observed amplitudes at each station.Focal mechanisms and seismic moments of all events thatoccurred in 2002 were determined by using the sameamplitude station corrections. For the 2004 observations, asmentioned, there were an additional five stations that werenot deployed in 2003. Therefore, we first determined thefocal mechanisms and seismic moments of all events thatoccurred in 2004 on the basis of the same amplitude stationcorrections without these five stations. We then calculatedthe amplitude station corrections at those five stations usingthe logarithmic average of the ratios between the theoreticaland observed amplitudes. Finally, we redetermined thefocal mechanisms and seismic moments of the 2004events incorporating the newly determined amplitude stationcorrections. The stability of the solution was checked byplotting all focal mechanisms whose residual was less than1.1‐fold the minimum residual value. We rejected ambigu-ous solutions where multiple solutions were possible. Intotal, we obtained 154 focal mechanisms whose momentmagnitudes ranged from 0.7 to 3.1. It should be noted thatduring the same period there are no routinely determinedfirst‐motion mechanisms by the JMA or moment tensorsolutions determined by the NIED. We also evaluated focalmechanism uncertainties for each event from the average ofKagan angles [Kagan, 1991] between the best fitting solu-

tion and all the solutions whose residual is within 1.1‐foldthat of the minimum residual. Here, the Kagan anglebecomes zero when the two mechanisms are the same, andbecomes 120° when they differ most. The average uncer-tainty of all the focal mechanisms was 14.7°.[13] Figure 6 shows an example of focal mechanism

determination for an earthquake (Mw0.8) that occurred at1754 LT (Japan standard time) on 5 August 2003. As isevident from Figure 6, P wave first‐motion polarity alonecannot constrain the mechanism of this earthquake. Thefirst‐motion P wave polarities are perfectly explained notonly by a pure reverse‐faulting‐type mechanism but also bya pure strike‐slip one. On the other hand, absolute ampli-tudes and the polarity observations require the strike‐slipfaulting‐type mechanism. All of the single‐event focalmechanism solutions determined in our study are shown inFigure 7 together with the first‐motion polarities. The esti-mated solutions explain the first‐motion polarities well.[14] The focal mechanisms are plotted onmaps in Figure 8a,

where different colors are used according to the faulting types;reverse (green), strike‐slip (red), and normal (blue) faultingmechanisms. Following Frohlich [1992], we classified reverseevents as those having T axis plunges of less than 40°, andstrike‐slip and normal fault earthquakes as those having B andP axis plunges of less than 30°, respectively. All remainingevents were defined as “other.”The numbers of reverse, strike‐slip, normal, and other events are 37, 64, 0, and 53, respec-tively. All events are plotted by open circles in a color trianglediagram (Figure 8a, bottom right), which suggests that mostof the other events are a mixture of reverse and strike‐slipcomponents. As was the case with the results obtained byWada et al. [2003], Ito et al. [2007], and Katsumata et al.[2010], our results reveal that many earthquakes with largereverse components are also occurring. In order to investi-gate the spatial pattern of focal mechanisms, we plotted theirfaulting types by means of colored circles in the cross sec-tion (Figure 8b). Here, the color is based on the color tri-angle diagram in Figure 8a. The most conspicuous feature isthe depth dependence in the western half of the entire fault,where the shallow earthquakes are primarily reverse faultingand the strike‐slip earthquakes dominate at the bottom of theseismogenic zone. We also found that the strike‐slip andreverse‐type earthquakes are mixed in all depth ranges nearthe center of the fault. These features had not been reportedin previous studies. The directions of the P and the T axesare shown in Figure 9, where different colors are used todenote different plunge angles. Most of the P axes are hori-zontal and oriented in the WNW‐ESE direction, whichconforms to the general tectonic trend in this area [e.g.,Tsukahara and Kobayashi, 1991; Townend and Zoback,2006]. The T axes have a wide range of plunge, suggestingthat reverse‐faulting‐type earthquakes as well as strike‐slipones are occurring in the area.[15] Katsumata et al. [2010] found 10 normal‐faulting

microearthquakes that occurred in a small cluster near thecentral part of the Atotsugawa fault and at depths greaterthan 8 km. The majority of those events occurred withinone month after the occurrence of the Mj4.1 earthquake of24 March 2005, where Mj is the magnitude reported by theJMA. As shown in Figure 3, the focal mechanism of theMj4.1 earthquake is a strike‐slip faulting type. There aresome observations that a strike‐slip earthquake produced a

Figure 5. P wave velocity (Vp) structure models used forhypocenter determination. The S wave models are assumedby scaling the P wave velocities by a factor of 1/

ffiffiffi

3p

.


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number of normal‐faulting aftershocks [e.g., Polat et al.,2002]. It seems probable that the normal faulting eventswere caused by a local stress perturbation due to the Mj4.1strike‐slip faulting earthquake and they do not reflect thegeneral stress state around the Atotsugawa fault.

5. Stress Field in and Aroundthe Atotsugawa Fault

[16] Using the focal mechanism solutions determined inthe present study, we calculated the stress field around theAtotsugawa fault by applying the inversion method of

Michael [1984]. The inversion solves the orientation of thethree principal stress axes and the relative magnitude of theprincipal stresses defined by � = (S2 − S3)/(S1 − S3), whereS1, S2, and S3 are the maximum, intermediate, and minimumcompressive principal stresses, respectively. We applied thebootstrap resampling technique to calculate confidenceregions for the stress tensor by assuming that a certainpercentage of the planes are picked incorrectly [Michael,1987]. In the present study, we assumed for the bootstrapthat each nodal plane had the same probability of beingchosen during the resampling. On the basis of thisassumption, we used 2000 bootstrap samples to obtain the

Figure 6. Example of focal mechanism determination for an earthquake (Mw0.8) that occurred at1754 LT on 5 August 2003. (top left) A map of the epicenter (open star) and station distribution. (topright) A unique focal mechanism solution cannot be determined by using P wave polarity data alone,while (bottom) it is possible to determine a unique solution (pure strike‐slip faulting) by using absoluteP and SH amplitudes and P wave polarity. Focal mechanism solutions are projected onto the lower hemi-sphere using equal‐area projection, where circles and triangles represent compressional and dilatationalfirst motions, respectively. The size of each symbol is proportional to the logarithmic amplitude.

Figure 7. Focal mechanism solutions of all events determined in the present study (lower hemisphere of equal‐area pro-jection). The circles and triangles represent compressional and dilatational first motions, respectively. Origin time (LT),event number, and moment magnitude are shown above each beach ball. The letter in the top left corner indicates the sub-area (see Figure 10) in which the earthquake is included. Letter “O” denotes an earthquake that is not included in anysubarea.


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Figure 7


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Figure 7. (continued)


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95% confidence region, which is adequate to produce stableconfidence regions up to the 95% level [Michael, 1987]. Weused earthquakes whose epicenters are within 3 km awayfrom the approximate surface trace of the Atotsugawa fault,

and we divided the data set into eastern and western areas inthe analysis. In order to investigate the depth dependence ofthe stress field, the western area was further divided intothree depth intervals as shown in Figure 10. The subarea W3

Figure 8. (a) Spatial distribution of focal mechanism solutions in seven different depth ranges (lowerhemisphere of equal‐area projection), where different colors are used to differentiate reverse (green),strike‐slip (red), and normal (blue) faulting mechanisms. The numbers adjacent to each beach ball corre-spond to the event number in Figure 7. A triangle diagram [Frohlich, 1992] with color scale is shown onthe right. Each focal mechanism is plotted by open circles. (b) Types of estimated focal mechanismsolutions in the X‐Y cross section. The color is based on the triangle diagram in Figure 8a.


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corresponds to the bottom of the seismogenic zone. For theeastern area, we cannot study the depth dependence of thestress field due to the low seismicity rate in the shallowerpart, so the stress field estimated in this area is considered tobe that around the bottom of the seismogenic zone. Thenumbers of focal mechanisms for areas W1, W2, W3, and Eare 22, 23, 22, and 31, respectively. It is noted that the datanear the center of the fault are included in areas W1–W3,although a different stress field around the center might besuggested by a different pattern of focal mechanism distri-bution there. It is considered that more data volume isnecessary to distinguish the different stress field aroundthe center.[17] The results of the stress tensor inversion are shown

in Figure 11. For every area, the maximum principal stressS1 is clearly differentiated from S2 and S3, trending hori-zontally WNW‐ESE. The directions of S2 and S3 are, on theother hand, nonuniform along the Atotsugawa fault, whichgives rise to a depth dependence in the stress field. The 95%confidence regions of S2 and S3 form a girdle and trendroughly NNE‐SSW for areas W1 and W2. Together with theobservation of a relatively low stress ratio �, 0–0.3, theseconfidence regions suggest that the principal stresses S2 andS3 differ only slightly in their magnitude and that the styleof faulting is a mixture of reverse and strike‐slip faulting.In contrast, the stress fields in areas W3 and E are charac-terized by a pure strike‐slip faulting regime, where theorientations of the principal stress axes are nearly verticalfor S2 and horizontal for S1 and S3. The stress ratio � is near0.5 for both areas, implying that the magnitude of S2 isclose to (S1 + S3)/2. This result of nonuniformity distributionof S2 and S3 is discussed in section 6.[18] Lund and Townend [2007] demonstrated that the

direction of maximum “horizontal” compressive stress(SHmax) is not always equal to the trend of the larger of the

Figure 9. P axis and T axis distributions of focal mecha-nism solutions. Different colors are used to indicate differentplunge angles.

Figure 10. Definition of subareas used in the stress tensor inversion (rectangles). White circles representfocal mechanisms used in the inversion, in which their locations are within about 3 km of the surface traceof the AGF. Earthquakes shown by crosses are not used in the stress tensor inversion, because they are toofar from the fault plane. A tomographic P wave velocity cross section along AGF is also shown. Thevelocity is contoured with intervals of 0.1 km/s.


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Figure 11. Stress tensor inversion results in each area. (left) Principal stress axes with their 95% con-fidence regions plotted on lower hemisphere stereonets. The shaded lines on the stereonets indicate theorientation of the Atotsugawa fault. (middle) Misfit angles for the data with respect to the best stress ten-sor determined by the stress tensor inversion. Here, the misfit angle represents the angle between the tan-gential traction predicted by the best solution and the observed slip direction on each plane determinedfrom the focal mechanism. (right) Frequency of the stress ratio �, which belongs to the 95% confidenceregion.


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two subhorizontal principal stresses. The difference dependsnot only on the plunges of the principal stress axes but alsoon the stress ratio �. Following the method of Lund andTownend [2007], we computed the true axis of SHmax

from the four stress parameters determined by the stresstensor inversion (the directions of the three principal stressesand the stress ratio �). Figure 12 shows the frequency ofSHmax computed from the stress parameters that are withinthe 95% confidence regions. Here, 180° is subtracted fromazimuth, if necessary, to ensure that all orientations can beplotted in the range of 0° to 180°. The range of SHmax withinthe 95% confidence regions is generally consistent with thatof the regional one [e.g., Tsukahara and Kobayashi, 1991;Townend and Zoback, 2006]. It seems that the SHmax

direction in area W3 is slightly different from those in otherareas. However, we cannot statistically exclude a possibilityof uniform SHmax direction across the entire region, becausethe 95% confidence regions of SHmax overlap each other. For

all areas, the orientation of SHmax relative to the fault trend isless than the predicted “lock‐up” angle for standard frictioncoefficients of about 0.6 [Sibson and Xie, 1998], suggestingthat the orientation of SHmax is under the condition forreactivating the Atotsugawa fault.[19] Except for the present research, the study by

Katsumata et al. [2010] is the only one that has carried out astress tensor inversion in and around the Atotsugawa faultsystem. They divided the region containing the Toyamaplain as well as the Atotsugawa fault system into 9 subareas,in which each area had a size of 18 and 23 km in the NS andEW directions, respectively. Although the subareas for thestress tensor computation are different from each other, theS1 directions determined in this study agree well with thosedetermined by Katsumata et al. [2010]. This suggests thatthe S1 direction is stable over the region and its estimate isnot influenced by the definition of subareas. For the stressratio, it is not appropriate to compare both results, becausewe found that it changes with subareas. However, the stressratios near the Atotsugawa fault estimated byKatsumata et al.[2010] vary from 0.23 to 0.57, which are within our estimates.

6. Inferred Model for the Stress Field

[20] It is common to assume that one principal stress isvertical and its magnitude is the overburden pressure (Sv)calculated from the density and depth [e.g., Zoback, 1992].If this is the case, the present study indicates that the min-imum horizontal stress (Shmin) has almost the same magni-tude as Sv at depths shallower than about 8 km in thewestern area. Because the reverse‐faulting‐type earthquakesare dominant at the shallower depth (Figure 8), Shmin shouldbe slightly larger than Sv in that depth range. On the otherhand, Shmin becomes less than Sv around the bottom of theseismogenic zone, so strike‐slip earthquakes are generated.The stress field in area E is almost the same as that in areaW3, suggesting that the bottom of the seismogenic zone ischaracterized by a similar stress field, although the depth

Figure 12. Frequency of the true axis of SHmax in each sub-area computed from the 95% confidence regions of thestress tensor inversion results using the algorithm by Lundand Townend [2007]. The shaded vertical line correspondsto the strike of the Atotsugawa fault.

Figure 13. (a) Aseismic slip model in a vertical cross section. The aseismic slip accumulates stress ontothe fault plane at the seismogenic zone during an earthquake cycle, and the main shock will occur whenthe failure stress is reached on the fault. The present study alone cannot constrain the depth range of theaseismic slip zone. (b) Schematic view of principal stress profile at the Atotsugawa fault.


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dependence of the stress field is not clarified, owning to alow seismicity rate in the shallower part. We infer that thedepth dependence of the stress field is caused by theaseismic deformation with right‐lateral strike‐slip faulting inthe downward extension of the fault as suggested by Mizunoet al. [2005]. In this case, a stress concentration occursabove the aseismic deformation, resulting in the increase inSHmax (maximum horizontal stress) and the decrease in Shmin

at the bottom of the seismogenic zone. As described above,Shmin should be slightly larger than Sv in most of the seis-mogenic zone, so the depth‐dependent change in the stressfield can occur as observed in this study. Figure 13 shows aschematic view of the proposed principal stress profile,where we assume, for the sake of simplicity, that SHmax andShmin are zero at the surface.[21] In order to confirm the validity of the present model,

we numerically compute the stress field in and around theAtotsugawa fault by the superposition of a stress change due

to aseismic slip in the downward extension of the fault andvariously assumed external stress fields around the fault. Inthis section, we define s1

e, s2e, and s3

e as maximum, inter-mediate, and minimum compressive principal stresses forthe external stress fields, respectively, to avoid confusionwith symbols used in the stress tensor inversion as well asthat described above.[22] Following Mizuno et al. [2005], we assume that

(1) the seismogenic zone of the Atotsugawa fault iscompletely locked and its deep extension is stably sliding,(2) the deep extension of the fault extends from the bottomof the seismogenic zone at a depth of 12 km to the bottom ofthe lower crust at a depth of 28 km, and (3) the relativedisplacement rate of the aseismic slip is 3 mm/yr, which iscomparable to the average fault slip rate along the Atotsu-gawa fault (Headquarters for Earthquake Research Promo-tion, Long‐term evaluation of the Atotsugawa fault (inJapanese), 2004, available at http://www.jishin.go.jp/main/

Figure 14. Perspective view of the model used in the numerical simulation. The fault plane is com-pletely locked and its deep extension is stably sliding. The X and Y axes correspond to East and Northdirections, respectively.


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chousa/04sep_atotugawa/index.htm) (Figure 14). Becausethe last large earthquake along the Atotsugawa fault, theHietsu earthquake, occurred about 150 years ago in 1858[Matsuda, 1966], the total aseismic slip becomes 0.45 m.We computed the stress change associated with the aseismicslip in an elastic half‐space following Okada [1992] byassuming a shear modulus of 32 GPa and a Poisson’s ratioof 0.25. The slip is linearly tapered as shown in Figure 14to avoid unrealistic stress concentrations at the edges ofthe fault.[23] For external stress fields, we assumed that s1

e ishorizontal and oriented to N105°E. This assumption isconsistent with the present study as well as other seismicobservations [e.g., Townend and Zoback, 2006; Katsumataet al., 2010]. It is also reasonable to assume that the ori-entation of s3

e is vertical, because except for the bottom ofthe seismogenic zone, the stress field around the Atotsuga-

wa fault is generally characterized by a reverse faultingregime. The magnitude of s3

e corresponds to the overburdenpressure (Sv). Under the assumption of the s1

e and s3e

orientations, s2e is horizontal and oriented to N15°E. We

also assumed a uniform differential stress (s1e − s3

e) of150 MPa in the seismogenic depth range. This valueapproximately corresponds to a depth‐averaged stresslevel of the upper crust in a reverse faulting stress regimepredicted for optimally oriented, cohesionless faults underconditions of hydrostatic pore pressure and a friction coef-ficient of 0.6 [e.g., Zoback and Townend, 2001]. Sincethere is no information about the magnitude of s2

e (or s2e −

s3e), we tried to constrain it by comparing the stress fields

deduced from the stress tensor inversion with theoreticallycalculated ones.[24] Through the numerical simulation, we found that the

spatial distribution in the faulting style (strike‐slip or reverse

Figure 15. Computed stress fields (superposition of the stress change due to the aseismic slip andregional stress field) at a depth of 8–11 km. SHmax directions calculated by the algorithm of Lund andTownend [2007] are plotted with 5 km spacing. The same color triangle diagram as that used inFigure 8a is employed to differentiate stress fields. See text for the assumption regarding the regionalstress field.


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faulting) strongly depends on the magnitude of (s2e − s3

e).Figure 15 shows computed stress fields (superposition of thestress change due to the aseismic slip and assumed externalstress fields) at the depth of 8–11 km for four examples of(s2

e − s3e), where SHmax directions calculated by the algo-

rithm of Lund and Townend [2007] are plotted with 5 kmspacing. The same color triangle diagram as that used inFigure 8 is employed to differentiate stress fields. Figure 15suggests that the difference in the magnitude of s2

e and s3e

can be constrained within some range to produce a strike‐slip faulting regime only at the bottom of the seismogeniczone. On the basis of numerous forward modeling, weconclude that the magnitude of s2

e − s3e should be in the

range of about 0.5 to 1 MPa. We reach the same conclusionas long as the magnitude of s1

e is at least 1 MPa larger thanthat of s2

e as well as s3e, through calculations of various

magnitudes of the external differential stress (s1e − s3

e). Ifthe magnitude of s1

e does not meet the condition, theorientations of computed maximum principal stress scattersignificantly and become inconsistent with the result ofthe stress tensor inversion.[25] We also compared the computed stress ratios with the

results of the stress tensor inversion. The computed stressratios have a value of less than 0.01 in the shallower partand about 0.5 at the bottom of the locked region for thepreferred stress state in Figure 15. This is owing to the factthat the stress change caused by the right‐lateral aseismicslip becomes predominant in the deeper part, while thestress change by the aseismic slip does not reach the shal-lower part and the external stress field becomes predominantthere. It seems probable that the computed stress ratio in theshallower part is somewhat lower than the result of the stresstensor inversion, though it is within the 95% confidenceregion (Figure 11). The external differential stress (s1

e − s3e)

of 5 MPa, for instance, gives a computed stress ratio in theshallower part of about 0.1–0.2, which seems to explain theresults of the stress tensor inversion better. It should benoted that such a small external differential stress requires asignificantly small friction coefficient of less than 0.1 tocause earthquakes in the seismogenic depth zone. It is stillcontroversial whether the friction coefficient in and aroundfault zones is governed by Byerlee’s [1978] law with afriction coefficient of about 0.6 (e.g., see review byTownend [2006]). For example, there is extensive evidenceshowing a lower friction coefficient at the San Andreas faultand at several other large strike‐slip faults and subductionzones [e.g., McGinty et al., 2000; d’Alessio et al., 2003;Hardebeck and Michael, 2004]. Scholz [2006] indicates thatexcept for the creeping section, the San Andreas fault has astrength consistent with a friction coefficient of 0.6 com-parable to that of the adjacent crust. For an inland activefault that caused the 1995 Hyogo‐ken‐Nanbu (Kobe)earthquake, which is categorized as the same type of intra-plate earthquake as that of the Atotsugawa fault, Yamashitaet al. [2004] estimated a friction coefficient of 0.6 bycombining in situ stress measurements with the static stresschange caused by the main shock. In contrast, Spudich et al.[1998] concluded the complete opposite on the basis of sliprake rotations with time. The present study alone cannotconstrain which value of the friction coefficient or externaldifferential stress explains the observations better, although

it provides a clue toward better understanding of a largeearthquake occurrence.

7. Discussion

[26] We consider that the depth dependence of the stressfield is related to the aseismic deformation with right‐lateralstrike‐slip faulting in the downward extension of the fault.In addition to observations of shear wave splitting and thepresent study, there is other observational evidence thatsupports the existence of the aseismic deformation belowthe Atotsugawa fault. We show, in Figure 10, a tomographicP wave velocity along the fault that was obtained byapplying a double‐difference tomography method [Zhangand Thurber, 2003] to arrival time data during the periodfrom April 1995 to December 2004 [Takeda et al., 2004].The P wave velocity is found to decrease with increasingdepth around the bottom of the seismogenic zone. The P andS wave velocity structures below the bottom of the Atot-sugawa fault were imaged by Matsubara et al. [2008], whodetermined 3‐D wave velocity structures beneath the wholeof the Japan Islands. Figure 16 shows P wave velocityperturbation and Vp/Vs at a depth of 15 km constructed fromtheir numerical data, suggesting that the low velocity as wellas a low Vp/Vs anomaly also exist below the Atotsugawafault. This pattern seems to continue down to a depth ofat least 25 km. According to Takei [2002], low Vp and lowVp/Vs reflect the presence of water‐filled cracks with a largeaspect ratio. Laboratory experiments of rock friction suggestthat high pore water pressure reduces the strength and tendsto promote an anelastic deformation [e.g., Scholz, 1998],supporting the existence of aseismic deformation below theAtotsugawa fault.[27] Hirahara et al. [2003, 2007] obtained the crustal

deformation around the Atotsugawa fault on the basis of denseGPS array observations. They proposed a model withoutaseismic deformation below the fault, in which two elasticblocks with a thickness of 15 km are obliquely collidingwith each other along the Atotsugawa fault in an E–W direc-tion. On the other hand, Sagiya et al. [2008] redetermined acrustal deformation pattern around the fault by incorporatingnewly deployed GPS stations and suggested the possibility ofexistence of a concentrated deformation at the deeper exten-sion of the fault. Because the average fault slip rate of theAtotsugawa fault is estimated to be 2–3 mm/yr (Headquartersfor Earthquake Research Promotion, Long‐term evaluation ofthe Atotsugawa fault (in Japanese), 2004, available at http://www.jishin.go.jp/main/chousa/04sep_atotugawa/index.htm),it is reasonable to assume that the relative displacement rateof the aseismic slip is of the same order. In this case, thedisplacement rate observed at the surface is less than 1 mm/yr,so it may be difficult to distinguish which model explainsthe observations better on the basis of GPS observationsover a decade. Our model predicts that the longer‐term GPSobservations will observe a signal caused by the aseismicdeformation in the downward extension of the fault.[28] It is now widely considered that aseismic deformation

below the seismogenic zone plays an important role for aninland large earthquake occurrence [e.g., Iio and Kobayashi,2002a]. For an improved understanding of the stress accu-mulation mechanism, it is important to know the style of


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deformation, whether it is localized in a narrow zone [e.g.,Stuart et al., 1997; Iio and Kobayashi, 2002a] or widelydistributed in volume [e.g., Kenner and Segall, 2000]. Ingeologic studies of exhumed fault zones, there is muchevidence for localized ductile shear zones in the lower crust[e.g., Shigematsu et al., 2004; Cole et al., 2007]. In addition,geophysical evidence for a localized slip below the seis-mogenic zone has recently been found at other faults. On thebasis of deep borehole heat flow measurements around theNojima fault that caused the 1995 Kobe earthquake, Iio andKobayashi [2002b] obtained the temperature profile down toa depth of 30 km. They compared the seismicity cutoffdepth in and around the Nojima fault with the brittle‐ductiletransition depth estimated from the temperature profile andfound that the cutoff depth can be explained by theassumption of a localized aseismic deformation with a widthof 100 m or less. Kawanishi et al. [2009] found a spatialchange in the stress field in and around the eastern part ofthe seismic belt along the Japan Sea coast and showed that itcan be explained by a localized aseismic deformation with awidth of 5 km or less. For the Atotsugawa fault, the stressratio � determined by the stress tensor inversion largelychanges at the bottom of the seismogenic zone (Figure 11).In addition, the orientation of SHmax derived from the shearwave splitting suddenly decreases within 1 km of the fault[Mizuno et al., 2005]. These observations seem to support alocalized slip model in the case of the Atotsugawa fault andthe width of the localized slip should be less than 1 km.[29] It should be noted that except at the bottom of the

seismogenic zone, the stress field is different from the faultingstyle of the Atotsugawa fault. This is considered to be for thefollowing reasons. Because the recurrence interval of largeearthquakes on the fault is estimated to be 2500 years[Research Group for Active Faults of Japan, 1991] and the1858 M7.0 Hietsu earthquake is considered to be due to a

faulting along the fault [Matsuda, 1966], the Atotsugawafault is now in an early stage in the earthquake cycle. Wesuggest that the influence of the stress concentration causedby the aseismic deformation will gradually spread upwardwith time and that the shallower part will also become astrike‐slip faulting regime in the future.[30] Finally, we discuss lateral variations of stress fields

along the fault. As mentioned in section 2, the seismicity rateon the northeast side near the central part of the Atotsugawafault is extremely low (Figure 1). In this area, theGeographicalSurvey Institute of Japan [2000] observed a right‐lateral creep‐like movement at a rate of about 1.5 mm/yr on the basis ofrepeated electronic distance measurement (EDM). Figure 10suggests that the area of Region_C characterized by creep‐like movement and low seismicity rate is roughly consistentwith a low‐Vp zone. Kato et al. [2007] revealed that the area isalso characterized by low‐Vp/Vs values and inferred that thelow seismicity rate and creep‐like movement are induced byhigh pore water pressures. If this is the case, a stress concen-tration will occur adjacent to the creep region and strike‐slipearthquakes will dominate there. In fact, there is a tendency forearthquakes with a large strike‐slip component to be dominantat the eastern edge of areas W1 and W2 (see 19, 30, 32, 123,143, 146, and 153 in Figure 8a) as described in section 4.However, the present data set is not necessarily sufficient toconclude that the creep‐like movement actually occurs alongthe fault. Further studies are needed to clarify the existence ofthe creep‐like movement on the basis of long‐term denseseismic and geodetic observations.

8. Conclusion

[31] Stress fields in and around the Atotsugawa fault incentral Japan have been investigated on the basis of numerousfocal mechanism solutions of microearthquakes. We con-

Figure 16. Tomographic imaging below the seismogenic zone [Matsubara et al., 2008]. (left) P wavevelocity perturbation and (right) Vp/Vs ratio at a depth of 15 km are shown in these examples.


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ducted temporary observations in the vicinity of the Atotsu-gawa fault from 2002 to 2004 and attempted to determinefocal mechanisms using P wave polarity data as well as bodywave amplitudes. We successfully obtained focal mechanismsolutions of 154 microearthquakes that occurred during theperiod of temporary observations, while there are no routinelydetermined ones in the same period. In addition to strike‐slipmicroearthquakes with a similar faulting style to the Atot-sugawa fault, we found that many microearthquakes withdominant reverse components are also occurring. A stresstensor inversion reveals the depth dependence in the stressfield, where the shallower part is characterized by amixture ofa reverse and strike‐slip faulting regime, and a pure strike‐slipfaulting regime appears only around the bottom of the seis-mogenic zone. Together with the spatial variation of shearwave splitting across the fault [Mizuno et al., 2005] and thetomographic imaging along the fault [Takeda et al., 2004;Matsubara et al., 2008], we suggest that the existence of alocalized aseismic deformation below the Atotsugawa fault isthe simplest scenario that can explain the observed depthdependence of the stress field. It should be noted that thescenario provides a stress accumulation mechanism of inlandearthquake faults, in which the localized aseismic deforma-tion accumulates stress onto the fault plane at the seismogeniczone during an earthquake cycle and the main shock willoccur when the failure stress is reached on the fault.

[32] Acknowledgments. Seismograph stations used in this studyinclude permanent stations operated by NIED (Hi‐net), JMA, ERI, KyotoUniversity, Nagoya University, and GSJ. We thank Reiken Matsushita forassistance in data processing and Yoshimitsu Okada for the use of his codein the stress calculations. We used “Slick Package” (http://earthquake.usgs.gov/research/software/) coded by Andrew Michael in the stress tensor inver-sion.Wemodified a program coded by Satoshi Ide for estimating focal mech-anism solutions. Comments by Robert Nowack and two anonymousreviewers were helpful in improving themanuscript.We downloaded numer-ical data from http://www.hinet.bosai.go.jp/topics/sokudo_kozo/alljpn_download.php to draw Figure 16. Most figures were generated usingGeneric Mapping Tools [Wessel and Smith, 1998]. This study was partiallysupported by one of the programs of the Earthquake Frontier ResearchProject by the Japan Nuclear Cycle Development Institute (JNC). K. I. wassupported by a Grant‐in‐Aid for Young Scientists (B) of the Ministry ofEducation, Culture, Sports, Science and Technology (17740294).

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Zoback, M. D., and J. Townend (2001), Implications of hydrostatic pore pres-sures and high crustal strength for the deformation of intraplate lithosphere,Tectonophysics, 336, 19–30, doi:10.1016/S0040-1951(01)00091-9.

Y. Haryu and T. Takeda, National Research Institute for Earth Scienceand Disaster Prevention, 3‐1 Tennodai, Tsukuba, Ibaraki, 305‐0006, Japan.K. Imanishi and Y. Kuwahara, Geological Survey of Japan, National

Institute of Advanced Industrial Science and Technology, TsukubaCentral 7, 1‐1‐1 Higashi, Tsukuba, Ibaraki 305‐8567, Japan. ([email protected])

H. Ito, Center for Deep Earth Exploration (CDEX), Japan Agency forMarine‐Earth Science and Technology, 3173‐25 Showa‐machi, Kanazawa,Yokohama, Kanagawa, 236‐0001, Japan.K. Ito, Hanshin Consultants Co., Ltd., 1‐6‐19 Imbashi, Chuo‐ku, Osaka

541‐0042, Japan.T. Mizuno, Schlumberger Kabushiki Kaisha Integration Center, 2‐2‐1

Fuchinobe, Sagamihara, Kanagawa, 229‐0006, Japan.H. Wada, Disaster Prevention Research Institute, Kyoto University,

Gokasho, Uji, Kyoto, 611‐0011, Japan.


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