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Reactivations of boundary faults within a buried ancient rift system by ductile creeping of weak shear zones in the overpressured lower crust: The 2004 mid-Niigata Prefecture Earthquake
Tectonophysics 486 (2010) 101–107
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Reactivations of boundary faults within a buried ancient rift system by ductile creeping of weak shear zones in the overpressured lower crust: The 2004 mid-Niigata Prefecture Earthquake Aitaro Kato ⁎, Takashi Iidaka, Takaya Iwasaki, Naoshi Hirata, Shigeki Nakagawa Earthquake Research Institute, University of Tokyo, Tokyo, Japan
a r t i c l e
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Article history: Received 28 September 2009 Received in revised form 15 February 2010 Accepted 15 February 2010 Available online 20 February 2010 Keywords: Tomography Aftershocks Reflective lower crust Boundary faults within the buried ancient rift system The 2004 mid-Niigata Prefecture Earthquake Overpressured crustal fluids
a b s t r a c t We elucidated fine-scale heterogeneities of the seismogenic structure associated with the 2004 mid-Niigata Prefecture Earthquake (thrust fault), Japan, by deploying a dense portable seismic array in the southwestern edge of the source region to observe the aftershocks. A velocity model inverted from double-difference tomographic analysis with first arrival times shows that most aftershocks were aligned between sedimentary strata in the hanging wall and the basement in the footwall. The basement is characterized by clear step-like and tilted block structures that gradually deepen toward the west. The domino-tilted block structures of the basement reveal evidence of a Miocene rift system buried beneath the thick sedimentary sequence. The aftershocks appear to be aligned roughly along preexisting boundaries of the step-like array of tilted blocks. In addition, we used the Natural Earthquake Reflection Profile (NERP) method to image reflective zones in the lower crust, defining a set of relatively strong reflectors. Low seismic velocity anomalies combined with high electrical conductivity in the lower crust suggest that a plausible explanation for the reflective zones is the presence of reservoirs of overpressured crustal fluids. Given these considerations, we propose that pre-existing boundary faults within the buried ancient rift system were reactivated through ductile creeping of weak shear zones in the overpressured lower crust. © 2010 Elsevier B.V. All rights reserved.
1. Introduction The 2004 mid-Niigata Prefecture Earthquake of Mw 6.6 occurred in the back-arc area of the main Japanese island as a shallow inland earthquake at 17:56 (JST=UT+9 h) on October 23, 2004. The focal area was situated within a zone of high E–W contractional strain rates larger than 10− 7 per year, which were detected by geodetic measurements and geological studies (Sagiya et al., 2000; Okamura et al., 2007). The focal mechanism of this earthquake was of a reverse fault type, with a strike of approximately N35°E (Fig. 1). The orientation of the P axis of the focal mechanism coincided fairly well with the regional stress field (e.g., Townend and Zoback, 2006; Kato et al., 2006). The earthquake was remarkable for the number of comparable large-magnitude events in the sequence (MN 4) [Japan Meteorological Agency (JMA) catalog]. The focal area of the mainshock rupture was located in the eastern margin of a thick (locally N 6 km deep) deformed middle Miocene– Pleistocene sedimentary basin (the Niigata Basin), which is characterized by a NNE–SSW-trending anticline–syncline system that forms the Uonuma and Higashi-Kubiki Hills (Yanagisawa et al., 1986; Okamura et al., 2007). This sedimentary basin was formed as a back-
⁎ Corresponding author. E-mail address: [email protected] (A. Kato). 0040-1951/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2010.02.012
arc basin in a rift structure that developed during the opening of the Japan Sea (25–15 Ma) (e.g., Sato and Amano, 1991). Miocene sediments mainly consist of muddy sediments and turbidites deposited in deep marine environments. From Pliocene to Pleistocene, the basins were filled with fluvial to shallow marine and coarse sediments. The basins were rapidly uplifted by folding and faulting during Pleistocene, and they emerged above sea-level at that time. This basin is bounded to the east by the Muikamachi Fault and its northeastern extension (Fig. 1), where pre-Neogene basement rocks (granitic, sedimentary and volcanic rocks) dating back to more than 30 Ma are widely exposed. Previous studies (e.g., Sato, 1994; Okamura et al., 1995; Kato et al., 2009) have inferred that parts of the normal faults within the rift system have subsequently been reactivated as a reverse fault system since the tectonic stress regime changed from extensional to compressional in the late Pliocene (2–3 Ma), through a process of compressional inversion (Williams et al., 1989). Intensive studies, such as hypocenter relocations, focal mechanism distributions, seismic tomography and imaging of reflection planes in the crust, have been conducted using dataset recorded at dense portable seismic networks deployed after the mainshock (e.g., Sakai et al., 2005; Kato et al., 2005; Okada et al., 2005, Shibutani et al., 2005, Matsumoto et al., 2005). Aftershock distributions estimated from these portable seismic stations showed that the mainshock occurred along a steep westward-dipping fault plane (50–60°), and many
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Fig. 1. Map of the seismic array and aftershocks cited from Kato et al. (2009). Aftershocks are shown as circles scaled to earthquake magnitude and color coded to depth. Moment tensor focal mechanism of the mainshock rupture (NIED) is plotted at the epicenter. The solid squares denote a portable seismic array installed in 2005 at the southwest edge of the aftershock area, and the black triangles are permanent seismic stations. The Muikamachi (MUKF) fault and other major active faults are shown by red lines. The black rectangle indicates the regional location of Fig. 2a. The inset shows the location of the studied area within the Niigata-Kobe Tectonic Zone (NKTZ). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
planar alignments of aftershocks associated with the subsequent comparable large-magnitude aftershocks. The aftershocks directly related to the mainshock rupture were aligned along a boundary between sediments in the hanging wall and the basement in the footwall. These previous studies suggested that the complex structures associated with crustal stretching and folding had significant potential to nucleate the mainshock and trigger a sequence of largemagnitude aftershocks. However, more detailed seismic observations of the crustal heterogeneities associated with the nucleation of the mainshock rupture are essential in order to understand the relationship between the crustal structures and the earthquake generation process. We therefore deployed a linear seismic array consisting of 115 stations with a total length of about 13 km, at the southwestern edge of the source region one year after the mainshock. Both a seismic tomography study and a Natural Earthquake Reflection Profiling (NERP) analysis using aftershock data retrieved from the seismic array provided us valuable opportunities (1) to investigate the fine-scale velocity structure in the seismogenic zone, with a spatial resolution of 2 km; (2) to map crustal reflectors beneath the seismogenic zone; and (3) to demonstrate that the Miocene rifting faults were reactivated through ductile creeping of weak shear zones in the overpressured lower crust.
2. Data and methods We installed a linear seismic array consisting of 115 portable stations at the southwestern edge of the aftershock zone; the waveform data were acquired from October 27 to November 4, 2005 (Figs. 1 and 2). Short-period vertical-channel sensors with 4.5 Hz natural frequency were laid out along the array and spaced about 120 m apart. These signals were recorded continuously at a sampling rate of 125 Hz using a small recorder (LS8200SD) (Kurashimo et al., 2006). To increase the accuracy of aftershocks, an additional nine portable stations with 1 Hz natural frequency sensors (three-component) were deployed near the array. A GPS receiver was equipped within each recorder to maintain the accuracy of the internal clocks of the seismic recorders. We merged the data from these portable seismic stations with those from permanent stations distributed around the linear array. A total of 69 aftershocks were detected from continuous waveforms observed by the seismic array (blue circles in Fig. 2). We then manually picked P- and S-wave arrival times from the detected aftershock waveforms. The picking accuracy ranged from 0.02 to 0.05 s for the P-wave, and from 0.04 to 0.1 s for the S-wave. To increase the accuracy of aftershock locations and the spatial resolution of the tomography model at the northwestern area away from the seismic array (Fig. 2), we added an aftershock dataset to the
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Fig. 2. (a) Distributions of portable seismic stations and relocated aftershocks near the seismic array (gray circles observed in 2004, and blue circles in 2005). The solid squares denote short-period vertical-channel sensors, and the open squares correspond to three-component portable stations installed in 2005. Open triangles indicate portable stations deployed immediately after the mainshock in 2004 (Kato and the research team of aftershock observations for the 2004 mid-Niigata Prefecture Earthquake, 2007). Solid triangles are permanent seismic stations. The grid used in the tomographic analysis is plotted with purple crosses, and the horizontal dashed line of Y = 0 km corresponds to the profile in Figs. 4 and 5. Major active faults are shown as red lines. (b) Depth section of the grid node and all of the relocated aftershocks along the profile of Y = 0 km.
dataset shown above (gray circles in Fig. 2), which we retrieved from the dense aftershock observation conducted immediately after the mainshock rupture (Kato and the research team of aftershock observations for the 2004 mid-Niigata Prefecture Earthquake, 2007; 58 events near the array from 2004/10/27 to 11/24). 2.1. Double-difference tomography and aftershock relocations To estimate the detailed velocity structure and accurate hypocenters in the source region, we applied double-difference tomography method (Zhang and Thurber, 2003) to the first arrival times of the P- and S-waves. The initial hypocenter locations for the tomography analysis were determined by applying a maximum likelihood estimation algorithm (Hirata and Matsu'ura, 1987) to the observed arrival times. We employed two different one-dimensional velocity structures (Fig. 3) (Sakai et al., 2005; Kato et al., 2005). At shallow depths of less than 3 km, the northwestern side of the Muikamachi fault has a remarkably P-wave low velocity, whereas the southeastern side has a moderately high velocity
(Fig. 3). The boundary between these velocity structures roughly coincides with that of the Muikamachi fault and its northeastward extension (Fig. 2). Furthermore, we adopted the station corrections, which were evaluated using average travel time residuals at each station, to locate the initial hypocenters before the tomography analysis. The initial velocity structure used for tomography analysis is the same as that employed for the hypocenter determinations (Vp /Vs is set as 1.73 in all the grids). The absolute arrival times for the P- and S-waves used in the tomography were 12,426 and 6259 respectively. The differential arrival times for the manually picked P- and S-waves reached 50,769 and 19,299, respectively. We also used the differential arrival times obtained by the waveform cross-correlation method in the time-domain (Schaff et al., 2004; Kato et al., 2006). We were left with a more accurate dataset of differential arrival times that contained 74,339 P-wave observations and 33,128 S-wave observations; all had a normalized cross-correlation coefficient larger than 0.7. The grids were located at −300, −20, −12, −10, −8, −6, −4, −2, 0, 2, 4, 6, 8, 16, and 300 km on the X (N125°E) axis; −300, −20, −10, 0, 10, 20, and 300 km on the Y (N35°E) axis; and
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data were filtered with a pass-band filter from 4 to 12 Hz, (2) the direct P-wave and S-wave were muted with a window length of 0.6 s and 1.0 s. 3. Tomographic image with relocated aftershocks
Fig. 3. Initial P-wave velocity structures employed for tomography analysis, which consist of two one-dimensional velocity structures at the eastern side of the Muikamachi fault, and within the Niigata Basin area (Sakai et al., 2005).
−150, 0, 2, 4, 6, 8, 10, 12, 15, 25, and 300 km on the Z (depth) axis (Fig. 2). The root mean square (RMS) travel time residual was reduced from 0.20 s to 0.05 s after 24 iterations. 2.2. Natural Earthquake Reflection Profiling (NERP) method We applied the NERP method (Nakagawa et al., 2005) to the linear array data (spaced 120 m apart) to image crustal reflections incorporating the Common Reflection Point (CRP) transform of the relocated aftershock waveforms. The CRP transform consisted of the following steps. (1) Calculation of the CRP for a horizontal reflector (deeper than a hypocenter) and travel times between possible combinations of the event and receivers based on a 1D velocity structure constructed from the tomography analysis (Table 1). Here, we assumed the PxP mode of reflections. (2) Mapping the amplitude of the waveform equivalent to the calculated travel time onto the CRP. We repeated (1)–(2) procedures for the next depth reflector by Δd (=20 m). We then obtained the CRP transformed image of the single event. The CRP transform relocated and re-sampled the data to a regular grid to place the reflectors at points with the assumption that the reflectors are horizontally layered. Finally, we stacked the CRP transformed sections of all events that lay in the same bin (spaced 100 m apart) on the profile of Y = 0 km. To exclude noisy events, we visually checked the waveform traces of all recorded events and selected 54 events for the present analysis. The pre-processing before the CRP transform consisted of two steps: (1) automatic gain control (AGC) with a time window of 3 s, which is equivalent to about 9 km in depth, was applied to the data, and the
Table 1 List of the velocity structure used for the NERP method, constructed from tomography analysis. Depth (km)
Vp (km/s)
0.00 2.00 4.00 6.00 8.00 10.00 16.00 16.01 32.01
2.35 3.85 3.90 5.24 5.94 6.00 6.10 6.70 6.71
To evaluate the model resolution of the final velocity structure, we conducted several following tests. We initially estimated the model resolution of the tomography (diagonal elements values of the resolution matrix) applying the solution technique provided by the simul2000 algorithm (Thurber and Eberhart-Phillips, 1999) to only the absolute arrival times. Because we did not include the double-difference data when estimating the resolution, the simil2000 underestimated the actual model resolution (Thurber et al., 2007). Fig. 4a shows that most model areas of Vp beneath the linear array had good resolution, as estimated from resolution diagonal elements value greater than 0.4, except at the model edges. Next, a checkerboard velocity model on the mainshock fault plane was created from the final velocity structure using perturbations of ±5%. Synthetic travel times were then calculated for this checkerboard model, adding random noises uniformly distributed from −0.05 s to 0.05 s for the P-wave, and from −0.1 s to 0.1 s for the S-wave. Subsequently, the synthetic data were inverted using the final velocity structure as a starting model. Fig. 4a shows the result of the checkerboard resolution test of Vp. P-wave velocity anomalies were well recovered within areas showing resolution diagonal elements greater than 0.4 (Fig. 4a). The depth section of Vp beneath the linear array is shown in Fig. 4b (Y = 0 km). Most relocated aftershocks were distributed between a low velocity layer and a high velocity body. The P-wave low velocity layer in the hanging wall is considered to be made up of a sedimentary sequence that accumulated in half-grabens formed by crustal stretching during the opening of the Japan Sea (Kato et al., 2005, 2006, 2009). Conversely, the high velocity body in the footwall is thought to correspond to the pre-Neogene basement assemblage (N30 Ma). According to the definition by Kato et al. (2009), we defined the basement as high velocity bodies in which Vp is greater than 5.7 km/s (bounded below by the white curves shown in Fig. 4b). Note that the top of the basement shows a clear step-like structure that gradually deepens to the west. There are at least three steps along this section, with widths of about 5 km. In addition, the westward-tilted block structures in the basement are outlined by aftershock alignments, and the top surface geometries of the basement (broken white lines in Fig. 4b). Aftershocks concentrated along the top surface of the basement show a cross-cutting geometry consisting of westward- and eastwarddipping planes (shown as arrows in Fig. 4b). These aftershocks appear to be aligned along the top surface geometry of the basement, although there is a slight gap between them due to the limits of the spatial resolution of the velocity model. We here interpret that the eastward-dipping alignment (ED) of the aftershocks coincided with an upper side of the tilted block A, and the westward-dipping alignment (WD) corresponded to a western-flank of the tilted block B. The WD and ED aftershock alignments occurred on planes dipping 50°-NW and 30°-SE, respectively. Because the mainshock rupture occurred on a plane dipping 50–60°-NW (e.g., Sakai et al., 2005), the WD plane corresponded to the southwestern edge of the mainshock fault plane. With deepening of the basement-cover contact to the west, the sedimentary succession in the hanging wall deepens to about 9 km. Indeed, the westward-tilt of the sedimentary strata was delineated by a seismic reflection survey conducted by the Japan National Oil Corporation (JNOC, 1988) in the vicinity of the studied area (Fig. 3 in Sato and Kato, 2005). The velocity model within the sedimentary basin correlates well with the seismic reflection profile. Elevation of topography increases toward the Muikamachi fault, which illustrates that the cover sequence has been deformed by upward movements of the step-like tilted blocks in the basement. There seem to be three
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The step-like tilted block structures within the basement provide clear evidence of the buried Miocene rift system formed during the spreading of the Japan Sea, as suggested by Kato et al. (2009). Most aftershocks in the studied area occurred near the contacts between the sedimentary cover sequence and the basement. These near-orthogonal faults, corresponding roughly to the pre-existing boundaries of the tilted blocks, were both reactivated as reverse faults. However, the steep NW-dipping faults were far from optimally oriented for reactivation in the inferred regional stress field with horizontal maximum compressive stress (Kato et al., 2006; Townend and Zoback, 2006), the implication being that those faults were mechanically weakened by high pore fluid pressures or clay minerals with low frictional coefficients within the fault zones (e.g., Kato et al., 2006; Sibson, 2007). 4. Reflection image
Fig. 4. (a) Results of the checkerboard resolution test of the Vp, and the superimposed resolution diagonal elements value contoured by solid line (0.4). (b) Depth section through the Vp model with superimposed relocated aftershocks distributed within ± 4 km from the section of Y = 0 km (gray circles observed in 2004, and red circles in 2005). Masked areas correspond to low model resolution. White curves denote iso-velocity contours of Vp = 5.7 km/s, and white broken lines show faults suggested from aftershock alignments and the top surface geometries of the basement. Three tilted blocks are labeled as A, B and C. Eastward- and westward-dipping alignments of aftershocks near the top surface of the basement are indicated by black (ED) and white (WD) arrows. The top figure shows a topography variation along the profile. Red lines correspond to the three plateaus in the topography.
plateaus within the topography to the west of the Muikamachi fault trace (red horizontal lines in the top graph in Fig. 4b). We therefore hypothesize that these plateaus might be created by both upward movements of the three tilted blocks and compressive deformations of the sedimentary strata in the hanging wall through compressional inversion.
Fig. 5a shows the final NERP depth section with the relocated aftershocks. Because most aftershocks occurred at depths less than 9 km, no reflection images at shallow depths were obtained. The observed reflective zones were delineated within the middle to lower crust, with relatively strong reflectors at depths of 10 km, 13–14 km, 17 km, 19–21 km, and 23 km, characterizing it as a pervasively reflective lower crust (e.g., Gough, 1986; Ito, 1999). Although some of these reflectors had been reported previously (Matsumoto et al., 2005), the detailed distribution of multiple reflectors had not previously been imaged. A similar reflective lower crust associated with a large inland earthquake was identified below the source regions of the 2000 western Tottori earthquake (Nakagawa et al., 2005) and the 1896 Riku-u earthquake in Japan (Sato et al., 2002). In addition, clear S-wave reflectors appear to be concentrated below the focal area of the 1962 northern Miyagi earthquake in the northeast Japan (e.g., Nakajima and Hasegawa, 2003). It thus seems likely that these reflectors exist below any regions where large intraplate earthquakes occur. Previous regional tomography studies (Nakajima and Hasegawa, 2008; Kato et al., 2009) have suggested that anomalously slow velocities are localized in the lower crust beneath the mainshock source area of the mid-Niigata Prefecture earthquake. These slow anomalies correlate well with the reflective layers delineated in the present study. Fig. 5b shows a depth section of S-wave perturbations along the same profile as the present study, interpolated from a regional (largescale) tomography model (Nakajima and Hasegawa, 2008). Although the spatial resolution of the regional model is fairly coarse (grid nodes with horizontal spacing of 0.1° and depth spacing of 0, 5, 10, 15, and 25 km), the lowermost crust beneath the seismogenic zone is characterized by anomalously slow velocities (shown as SLV in Fig. 5b). Note that the large-scale tomographic model was not able to delineate the ancient rift structure beneath the thick sediments at shallow depths, as imaged in the present study; it showed only low and high velocity layers at shallow depths. Additionally, a wideband magnetotelluric survey (Uyeshima et al., 2005) found a high electrical conductive body in the lower crust (deeper than 15 km) beneath the source region. The combination of anomalously low seismic velocities with high electrical conductivity in the lower crust is most plausibly explained by accumulations of overpressured hydrothermal fluids (e.g., Sibson, 2009). From geological observations of exhumed vein systems (Foxford et al., 2000; Tunks et al., 2004), it seems likely that fluids overpressured to approximately lithostatic values would occupy arrays of flat-lying extension fractures developed by hydraulic fracturing in a compressional stress regime where fluid pressure exceeded the vertical stress (Sibson, 2009). Because the minimum principal compressive stress is vertical in such a regime, the hydrothermally infilled extension fractures would be subhorizontal and elongated parallel to the profile in Fig. 5. The effect of strong fluid overpressuring to near-lithostatic values is to enhance connectivity between pore space and fractures, thereby increasing electrical conductivity and lowering seismic velocities in the rock mass (e.g., Gough, 1986).
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Fig. 5. (a) Final NERP reflection profile along the Y = 0 km with an assumption of PxP reflection mode. Arrows denote depths of relatively strong reflectors. Circles show the relocated aftershocks distributed within ± 4 km from the cross-section of Y = 0 km (gray circles observed in 2004, and red circles in 2005). The heavy line corresponds to the iso-velocity contours of Vp = 5.7 km/s in Fig. 4b. (b) Depth section of S-wave velocity perturbations along the same profile as in Fig. 5a, interpolated from the regional (large-scale) tomography model constructed by Nakajima and Hasegawa (2008). SLV means slow velocity anomaly in the lowermost crust.
5. Reactivations of boundary faults by ductile creeping of weak shear zones in the overpressured lower crust The fine-scale tomography image (Fig. 4b) suggests that the seismogenesis of the mid-Niigata Prefecture earthquakes was due primarily to compressional inversion tectonics involving reactivation of weakened boundary faults related to the Miocene rift system. Furthermore, the reflection study reveals the potential existence of overpressured hydrothermal fluids beneath the ancient rift system. Since the strength of the ductile creeping that is a dominant deformation process in the lower crust is lowered by crustal fluids (e.g., Carter and Tsenn, 1987), weak zones could develop beneath the ancient rift system (Iio et al., 2002; Kato et al., 2009). Thus, ductile creeping of weak shear zones in the lower crust may have loaded faults in the upper crust inherited from the Miocene rifting episode, leading to reactivation of this system in the form of intraplate earthquakes (e.g., Kenner and Segall, 2000; Kato et al., 2009). Because the lower curst was affected by Miocene rifting, the possibility also exists that some of the reflectors in the lower crust arise from mafic intrusions in the form of dyke–sill complexes (Hirata et al., 1989). Although regional tomographic studies (Nakajima and Hasegawa, 2008; Kato et al., 2009) suggest that the lower crust has slow anomalies on average beneath the source region, higher resolution studies are needed to resolve the fine-scale heterogeneity within the lower crust and determine the origin of the reflectors.
Given these considerations, we propose that inherited boundary faults within the buried ancient rift system were reactivated through activation of ductile shear zones in the overpressured lower crust. Hydrothermal overpressuring in the lower crust would likely lead to episodic invasion of overpressured fluids into the lower seismogenic zone through fault-valve action (e.g., Sibson, 2007, 2009). The present study provides additional evidence supporting such behavior. In central Italy in 1997, a shallow normal-faulting sequence occurred with accompanying clear earthquake migration and multiple mainshocks (Mw 5.2–6). It has been proposed that this sequence could be driven by the propagation of overpressured fluids trapped below the seismogenic layer through fault-damaged zones (e.g., Miller et al., 2004; Chiarabba et al., 2009). Overpressured fluids below the seismogenic layer are likely to trigger intraplate earthquakes even in different tectonic environments. It is therefore necessary to detect and monitor overpressured fluids below the seismogenic layer during an interseismic cycle. 6. Conclusions We deployed a dense portable seismic array within the source region of the 2004 mid-Niigata Prefecture Earthquake to understand fine-scale crustal heterogeneity. We conducted double-difference tomographic analysis with the first arrival data, and performed the
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natural earthquake reflection profile analysis with the relocated aftershock waveform data. (1) Relocated aftershocks were clearly aligned on both the steeply westward-dipping and the gently eastward-dipping fault planes that separate the sedimentary cover sequence in the hanging wall from the pre-Neogene basement in the footwall. (2) The basement is characterized by clear step-like tilted block structures that gradually deepen to the west. The block width is about 5 km. The step-like tilted block structures of the basement provide evidence of the Miocene rift system buried beneath the thick sediments. (3) Anomalously low seismic velocities combined with high electrical conductivity suggest that stacked reservoirs of overpressured fluids provide a plausible explanation for an array of relatively strong flat-lying reflectors in the lower crust. We propose that pre-existing boundary faults within the buried ancient rift system were reactivated by ductile creeping of weak shear zones in the overpressured lower crust. Acknowledgments We are thankful to H. Zhang for allowing us to use the TomoDD code. We also thank the research team for aftershock observations of the 2004 mid-Niigata Prefecture Earthquake for providing the arrival times and waveform data of aftershocks. We are grateful to one anonymous reviewer, C. Chiarabba and Editor H. Thybo for their constructive comments, which led to substantial improvement to the original manuscript. J. Nakajima kindly provided us with the regional tomography model. We are grateful to Kyoto University, Kyusyu University, the National Research Institute for Earth Science and Disaster Prevention, and the Japan Meteorological Agency for allowing us to use the waveform data collected at online stations. This research was partially supported by the Ministry of Education, Culture, Sports, Science and Technology (MEXT) of Japan, under its Observation and Research Program for Prediction of Earthquakes and Volcanic Eruptions, and by the cooperative research program of the Earthquake Research Institute. References Carter, N.L., Tsenn, M.C., 1987. Flow properties of the continental lithosphere. Tectonophysics 136, 27–63. Chiarabba, C., De Gori, P., Boschi, E., 2009. Pore pressure migration along a normal fault system resolved by time repeated seismic tomography. Geology 37 (1), 67–70. doi:10.1130/G25220A.1. Foxford, K.A., Nicholson, R., Polya, D.A., Hebblethwaite, R.P.B., 2000. Extensional failure and hydraulic valving at Minas da Panasqueira, Portugal: evidence from vein spatial distributions, displacements and geometries. J. Struct. Geol. 22, 1065–1086. Gough, D.I., 1986. Seismic reflectivity, conductivity, water, and stress in the continental crust. Nature 323, 143–144. Hirata, N., Matsu'ura, M., 1987. Maximum-likelihood estimation of hypocenter with origin time eliminated using nonlinear inversion technique. Phys. Earth Planet. Inter. 47, 50–61. Hirata, N., Tokuyama, H., Chung, T.W., 1989. An anomalously thick layering of the crust of the Yamato Basin, southeastern Sea of Japan: the final stage of back-arc spreading. Tectonophysics 165, 303–314. Iio, Y., Sagiya, T., Kobayashi, Y., Shiozaki, I., 2002. Water-weakened lower crust and its role in the concentrated deformation in the Japanese Islands, Earth Planet. Sci. Lett. 203, 245–253. Ito, K., 1999. Seismogenic layer, reflective lower crust, surface heat flow and large inland earthquakes. Tectonophysics 306, 423–433. JNOC, 1988. Report on basic geophysical exploration in onshore area “KubikiTamugiyama”. Japan National Oil Corporation, Tokyo. 45 pp. (in Japanese). Kato, A., the research team of aftershock observations for the 2004 mid-Niigata Prefecture Earthquake, 2007. High-resolution aftershock observations in the source region of the 2004 mid-Niigata Prefecture Earthquake. Earth Planets Space 59, 923–928. Kato, A., Kurashimo, E., Hirata, N., Iwasaki, T., Kanazawa, T., 2005. Imaging the source region of the 2004 Mid-Niigata prefecture earthquake and the evolution of a seismogenic thrust-related fold. Geophys. Res. Lett. 32, L07307. doi:10.1029/ 2005GL022366. Kato, A., Sakai, S., Hirata, N., Kurashimo, E., Iidaka, T., Iwasaki, T., Kanazawa, T., 2006. Imaging the seismic structure and stress field in the source region of the 2004 mid-Niigata Prefecture Earthquake: structural zones of weakness, seismogenic stress concentration by ductile flow. J. Geophys. Res. 111, B08308. doi:10.1029/2005JB004016.
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Tectonophysics j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / t e c t o
Reactivations of boundary faults within a buried ancient rift system by ductile creeping of weak shear zones in the overpressured lower crust: The 2004 mid-Niigata Prefecture Earthquake Aitaro Kato ⁎, Takashi Iidaka, Takaya Iwasaki, Naoshi Hirata, Shigeki Nakagawa Earthquake Research Institute, University of Tokyo, Tokyo, Japan
a r t i c l e
i n f o
Article history: Received 28 September 2009 Received in revised form 15 February 2010 Accepted 15 February 2010 Available online 20 February 2010 Keywords: Tomography Aftershocks Reflective lower crust Boundary faults within the buried ancient rift system The 2004 mid-Niigata Prefecture Earthquake Overpressured crustal fluids
a b s t r a c t We elucidated fine-scale heterogeneities of the seismogenic structure associated with the 2004 mid-Niigata Prefecture Earthquake (thrust fault), Japan, by deploying a dense portable seismic array in the southwestern edge of the source region to observe the aftershocks. A velocity model inverted from double-difference tomographic analysis with first arrival times shows that most aftershocks were aligned between sedimentary strata in the hanging wall and the basement in the footwall. The basement is characterized by clear step-like and tilted block structures that gradually deepen toward the west. The domino-tilted block structures of the basement reveal evidence of a Miocene rift system buried beneath the thick sedimentary sequence. The aftershocks appear to be aligned roughly along preexisting boundaries of the step-like array of tilted blocks. In addition, we used the Natural Earthquake Reflection Profile (NERP) method to image reflective zones in the lower crust, defining a set of relatively strong reflectors. Low seismic velocity anomalies combined with high electrical conductivity in the lower crust suggest that a plausible explanation for the reflective zones is the presence of reservoirs of overpressured crustal fluids. Given these considerations, we propose that pre-existing boundary faults within the buried ancient rift system were reactivated through ductile creeping of weak shear zones in the overpressured lower crust. © 2010 Elsevier B.V. All rights reserved.
1. Introduction The 2004 mid-Niigata Prefecture Earthquake of Mw 6.6 occurred in the back-arc area of the main Japanese island as a shallow inland earthquake at 17:56 (JST=UT+9 h) on October 23, 2004. The focal area was situated within a zone of high E–W contractional strain rates larger than 10− 7 per year, which were detected by geodetic measurements and geological studies (Sagiya et al., 2000; Okamura et al., 2007). The focal mechanism of this earthquake was of a reverse fault type, with a strike of approximately N35°E (Fig. 1). The orientation of the P axis of the focal mechanism coincided fairly well with the regional stress field (e.g., Townend and Zoback, 2006; Kato et al., 2006). The earthquake was remarkable for the number of comparable large-magnitude events in the sequence (MN 4) [Japan Meteorological Agency (JMA) catalog]. The focal area of the mainshock rupture was located in the eastern margin of a thick (locally N 6 km deep) deformed middle Miocene– Pleistocene sedimentary basin (the Niigata Basin), which is characterized by a NNE–SSW-trending anticline–syncline system that forms the Uonuma and Higashi-Kubiki Hills (Yanagisawa et al., 1986; Okamura et al., 2007). This sedimentary basin was formed as a back-
⁎ Corresponding author. E-mail address: [email protected] (A. Kato). 0040-1951/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2010.02.012
arc basin in a rift structure that developed during the opening of the Japan Sea (25–15 Ma) (e.g., Sato and Amano, 1991). Miocene sediments mainly consist of muddy sediments and turbidites deposited in deep marine environments. From Pliocene to Pleistocene, the basins were filled with fluvial to shallow marine and coarse sediments. The basins were rapidly uplifted by folding and faulting during Pleistocene, and they emerged above sea-level at that time. This basin is bounded to the east by the Muikamachi Fault and its northeastern extension (Fig. 1), where pre-Neogene basement rocks (granitic, sedimentary and volcanic rocks) dating back to more than 30 Ma are widely exposed. Previous studies (e.g., Sato, 1994; Okamura et al., 1995; Kato et al., 2009) have inferred that parts of the normal faults within the rift system have subsequently been reactivated as a reverse fault system since the tectonic stress regime changed from extensional to compressional in the late Pliocene (2–3 Ma), through a process of compressional inversion (Williams et al., 1989). Intensive studies, such as hypocenter relocations, focal mechanism distributions, seismic tomography and imaging of reflection planes in the crust, have been conducted using dataset recorded at dense portable seismic networks deployed after the mainshock (e.g., Sakai et al., 2005; Kato et al., 2005; Okada et al., 2005, Shibutani et al., 2005, Matsumoto et al., 2005). Aftershock distributions estimated from these portable seismic stations showed that the mainshock occurred along a steep westward-dipping fault plane (50–60°), and many
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Fig. 1. Map of the seismic array and aftershocks cited from Kato et al. (2009). Aftershocks are shown as circles scaled to earthquake magnitude and color coded to depth. Moment tensor focal mechanism of the mainshock rupture (NIED) is plotted at the epicenter. The solid squares denote a portable seismic array installed in 2005 at the southwest edge of the aftershock area, and the black triangles are permanent seismic stations. The Muikamachi (MUKF) fault and other major active faults are shown by red lines. The black rectangle indicates the regional location of Fig. 2a. The inset shows the location of the studied area within the Niigata-Kobe Tectonic Zone (NKTZ). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
planar alignments of aftershocks associated with the subsequent comparable large-magnitude aftershocks. The aftershocks directly related to the mainshock rupture were aligned along a boundary between sediments in the hanging wall and the basement in the footwall. These previous studies suggested that the complex structures associated with crustal stretching and folding had significant potential to nucleate the mainshock and trigger a sequence of largemagnitude aftershocks. However, more detailed seismic observations of the crustal heterogeneities associated with the nucleation of the mainshock rupture are essential in order to understand the relationship between the crustal structures and the earthquake generation process. We therefore deployed a linear seismic array consisting of 115 stations with a total length of about 13 km, at the southwestern edge of the source region one year after the mainshock. Both a seismic tomography study and a Natural Earthquake Reflection Profiling (NERP) analysis using aftershock data retrieved from the seismic array provided us valuable opportunities (1) to investigate the fine-scale velocity structure in the seismogenic zone, with a spatial resolution of 2 km; (2) to map crustal reflectors beneath the seismogenic zone; and (3) to demonstrate that the Miocene rifting faults were reactivated through ductile creeping of weak shear zones in the overpressured lower crust.
2. Data and methods We installed a linear seismic array consisting of 115 portable stations at the southwestern edge of the aftershock zone; the waveform data were acquired from October 27 to November 4, 2005 (Figs. 1 and 2). Short-period vertical-channel sensors with 4.5 Hz natural frequency were laid out along the array and spaced about 120 m apart. These signals were recorded continuously at a sampling rate of 125 Hz using a small recorder (LS8200SD) (Kurashimo et al., 2006). To increase the accuracy of aftershocks, an additional nine portable stations with 1 Hz natural frequency sensors (three-component) were deployed near the array. A GPS receiver was equipped within each recorder to maintain the accuracy of the internal clocks of the seismic recorders. We merged the data from these portable seismic stations with those from permanent stations distributed around the linear array. A total of 69 aftershocks were detected from continuous waveforms observed by the seismic array (blue circles in Fig. 2). We then manually picked P- and S-wave arrival times from the detected aftershock waveforms. The picking accuracy ranged from 0.02 to 0.05 s for the P-wave, and from 0.04 to 0.1 s for the S-wave. To increase the accuracy of aftershock locations and the spatial resolution of the tomography model at the northwestern area away from the seismic array (Fig. 2), we added an aftershock dataset to the
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Fig. 2. (a) Distributions of portable seismic stations and relocated aftershocks near the seismic array (gray circles observed in 2004, and blue circles in 2005). The solid squares denote short-period vertical-channel sensors, and the open squares correspond to three-component portable stations installed in 2005. Open triangles indicate portable stations deployed immediately after the mainshock in 2004 (Kato and the research team of aftershock observations for the 2004 mid-Niigata Prefecture Earthquake, 2007). Solid triangles are permanent seismic stations. The grid used in the tomographic analysis is plotted with purple crosses, and the horizontal dashed line of Y = 0 km corresponds to the profile in Figs. 4 and 5. Major active faults are shown as red lines. (b) Depth section of the grid node and all of the relocated aftershocks along the profile of Y = 0 km.
dataset shown above (gray circles in Fig. 2), which we retrieved from the dense aftershock observation conducted immediately after the mainshock rupture (Kato and the research team of aftershock observations for the 2004 mid-Niigata Prefecture Earthquake, 2007; 58 events near the array from 2004/10/27 to 11/24). 2.1. Double-difference tomography and aftershock relocations To estimate the detailed velocity structure and accurate hypocenters in the source region, we applied double-difference tomography method (Zhang and Thurber, 2003) to the first arrival times of the P- and S-waves. The initial hypocenter locations for the tomography analysis were determined by applying a maximum likelihood estimation algorithm (Hirata and Matsu'ura, 1987) to the observed arrival times. We employed two different one-dimensional velocity structures (Fig. 3) (Sakai et al., 2005; Kato et al., 2005). At shallow depths of less than 3 km, the northwestern side of the Muikamachi fault has a remarkably P-wave low velocity, whereas the southeastern side has a moderately high velocity
(Fig. 3). The boundary between these velocity structures roughly coincides with that of the Muikamachi fault and its northeastward extension (Fig. 2). Furthermore, we adopted the station corrections, which were evaluated using average travel time residuals at each station, to locate the initial hypocenters before the tomography analysis. The initial velocity structure used for tomography analysis is the same as that employed for the hypocenter determinations (Vp /Vs is set as 1.73 in all the grids). The absolute arrival times for the P- and S-waves used in the tomography were 12,426 and 6259 respectively. The differential arrival times for the manually picked P- and S-waves reached 50,769 and 19,299, respectively. We also used the differential arrival times obtained by the waveform cross-correlation method in the time-domain (Schaff et al., 2004; Kato et al., 2006). We were left with a more accurate dataset of differential arrival times that contained 74,339 P-wave observations and 33,128 S-wave observations; all had a normalized cross-correlation coefficient larger than 0.7. The grids were located at −300, −20, −12, −10, −8, −6, −4, −2, 0, 2, 4, 6, 8, 16, and 300 km on the X (N125°E) axis; −300, −20, −10, 0, 10, 20, and 300 km on the Y (N35°E) axis; and
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data were filtered with a pass-band filter from 4 to 12 Hz, (2) the direct P-wave and S-wave were muted with a window length of 0.6 s and 1.0 s. 3. Tomographic image with relocated aftershocks
Fig. 3. Initial P-wave velocity structures employed for tomography analysis, which consist of two one-dimensional velocity structures at the eastern side of the Muikamachi fault, and within the Niigata Basin area (Sakai et al., 2005).
−150, 0, 2, 4, 6, 8, 10, 12, 15, 25, and 300 km on the Z (depth) axis (Fig. 2). The root mean square (RMS) travel time residual was reduced from 0.20 s to 0.05 s after 24 iterations. 2.2. Natural Earthquake Reflection Profiling (NERP) method We applied the NERP method (Nakagawa et al., 2005) to the linear array data (spaced 120 m apart) to image crustal reflections incorporating the Common Reflection Point (CRP) transform of the relocated aftershock waveforms. The CRP transform consisted of the following steps. (1) Calculation of the CRP for a horizontal reflector (deeper than a hypocenter) and travel times between possible combinations of the event and receivers based on a 1D velocity structure constructed from the tomography analysis (Table 1). Here, we assumed the PxP mode of reflections. (2) Mapping the amplitude of the waveform equivalent to the calculated travel time onto the CRP. We repeated (1)–(2) procedures for the next depth reflector by Δd (=20 m). We then obtained the CRP transformed image of the single event. The CRP transform relocated and re-sampled the data to a regular grid to place the reflectors at points with the assumption that the reflectors are horizontally layered. Finally, we stacked the CRP transformed sections of all events that lay in the same bin (spaced 100 m apart) on the profile of Y = 0 km. To exclude noisy events, we visually checked the waveform traces of all recorded events and selected 54 events for the present analysis. The pre-processing before the CRP transform consisted of two steps: (1) automatic gain control (AGC) with a time window of 3 s, which is equivalent to about 9 km in depth, was applied to the data, and the
Table 1 List of the velocity structure used for the NERP method, constructed from tomography analysis. Depth (km)
Vp (km/s)
0.00 2.00 4.00 6.00 8.00 10.00 16.00 16.01 32.01
2.35 3.85 3.90 5.24 5.94 6.00 6.10 6.70 6.71
To evaluate the model resolution of the final velocity structure, we conducted several following tests. We initially estimated the model resolution of the tomography (diagonal elements values of the resolution matrix) applying the solution technique provided by the simul2000 algorithm (Thurber and Eberhart-Phillips, 1999) to only the absolute arrival times. Because we did not include the double-difference data when estimating the resolution, the simil2000 underestimated the actual model resolution (Thurber et al., 2007). Fig. 4a shows that most model areas of Vp beneath the linear array had good resolution, as estimated from resolution diagonal elements value greater than 0.4, except at the model edges. Next, a checkerboard velocity model on the mainshock fault plane was created from the final velocity structure using perturbations of ±5%. Synthetic travel times were then calculated for this checkerboard model, adding random noises uniformly distributed from −0.05 s to 0.05 s for the P-wave, and from −0.1 s to 0.1 s for the S-wave. Subsequently, the synthetic data were inverted using the final velocity structure as a starting model. Fig. 4a shows the result of the checkerboard resolution test of Vp. P-wave velocity anomalies were well recovered within areas showing resolution diagonal elements greater than 0.4 (Fig. 4a). The depth section of Vp beneath the linear array is shown in Fig. 4b (Y = 0 km). Most relocated aftershocks were distributed between a low velocity layer and a high velocity body. The P-wave low velocity layer in the hanging wall is considered to be made up of a sedimentary sequence that accumulated in half-grabens formed by crustal stretching during the opening of the Japan Sea (Kato et al., 2005, 2006, 2009). Conversely, the high velocity body in the footwall is thought to correspond to the pre-Neogene basement assemblage (N30 Ma). According to the definition by Kato et al. (2009), we defined the basement as high velocity bodies in which Vp is greater than 5.7 km/s (bounded below by the white curves shown in Fig. 4b). Note that the top of the basement shows a clear step-like structure that gradually deepens to the west. There are at least three steps along this section, with widths of about 5 km. In addition, the westward-tilted block structures in the basement are outlined by aftershock alignments, and the top surface geometries of the basement (broken white lines in Fig. 4b). Aftershocks concentrated along the top surface of the basement show a cross-cutting geometry consisting of westward- and eastwarddipping planes (shown as arrows in Fig. 4b). These aftershocks appear to be aligned along the top surface geometry of the basement, although there is a slight gap between them due to the limits of the spatial resolution of the velocity model. We here interpret that the eastward-dipping alignment (ED) of the aftershocks coincided with an upper side of the tilted block A, and the westward-dipping alignment (WD) corresponded to a western-flank of the tilted block B. The WD and ED aftershock alignments occurred on planes dipping 50°-NW and 30°-SE, respectively. Because the mainshock rupture occurred on a plane dipping 50–60°-NW (e.g., Sakai et al., 2005), the WD plane corresponded to the southwestern edge of the mainshock fault plane. With deepening of the basement-cover contact to the west, the sedimentary succession in the hanging wall deepens to about 9 km. Indeed, the westward-tilt of the sedimentary strata was delineated by a seismic reflection survey conducted by the Japan National Oil Corporation (JNOC, 1988) in the vicinity of the studied area (Fig. 3 in Sato and Kato, 2005). The velocity model within the sedimentary basin correlates well with the seismic reflection profile. Elevation of topography increases toward the Muikamachi fault, which illustrates that the cover sequence has been deformed by upward movements of the step-like tilted blocks in the basement. There seem to be three
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The step-like tilted block structures within the basement provide clear evidence of the buried Miocene rift system formed during the spreading of the Japan Sea, as suggested by Kato et al. (2009). Most aftershocks in the studied area occurred near the contacts between the sedimentary cover sequence and the basement. These near-orthogonal faults, corresponding roughly to the pre-existing boundaries of the tilted blocks, were both reactivated as reverse faults. However, the steep NW-dipping faults were far from optimally oriented for reactivation in the inferred regional stress field with horizontal maximum compressive stress (Kato et al., 2006; Townend and Zoback, 2006), the implication being that those faults were mechanically weakened by high pore fluid pressures or clay minerals with low frictional coefficients within the fault zones (e.g., Kato et al., 2006; Sibson, 2007). 4. Reflection image
Fig. 4. (a) Results of the checkerboard resolution test of the Vp, and the superimposed resolution diagonal elements value contoured by solid line (0.4). (b) Depth section through the Vp model with superimposed relocated aftershocks distributed within ± 4 km from the section of Y = 0 km (gray circles observed in 2004, and red circles in 2005). Masked areas correspond to low model resolution. White curves denote iso-velocity contours of Vp = 5.7 km/s, and white broken lines show faults suggested from aftershock alignments and the top surface geometries of the basement. Three tilted blocks are labeled as A, B and C. Eastward- and westward-dipping alignments of aftershocks near the top surface of the basement are indicated by black (ED) and white (WD) arrows. The top figure shows a topography variation along the profile. Red lines correspond to the three plateaus in the topography.
plateaus within the topography to the west of the Muikamachi fault trace (red horizontal lines in the top graph in Fig. 4b). We therefore hypothesize that these plateaus might be created by both upward movements of the three tilted blocks and compressive deformations of the sedimentary strata in the hanging wall through compressional inversion.
Fig. 5a shows the final NERP depth section with the relocated aftershocks. Because most aftershocks occurred at depths less than 9 km, no reflection images at shallow depths were obtained. The observed reflective zones were delineated within the middle to lower crust, with relatively strong reflectors at depths of 10 km, 13–14 km, 17 km, 19–21 km, and 23 km, characterizing it as a pervasively reflective lower crust (e.g., Gough, 1986; Ito, 1999). Although some of these reflectors had been reported previously (Matsumoto et al., 2005), the detailed distribution of multiple reflectors had not previously been imaged. A similar reflective lower crust associated with a large inland earthquake was identified below the source regions of the 2000 western Tottori earthquake (Nakagawa et al., 2005) and the 1896 Riku-u earthquake in Japan (Sato et al., 2002). In addition, clear S-wave reflectors appear to be concentrated below the focal area of the 1962 northern Miyagi earthquake in the northeast Japan (e.g., Nakajima and Hasegawa, 2003). It thus seems likely that these reflectors exist below any regions where large intraplate earthquakes occur. Previous regional tomography studies (Nakajima and Hasegawa, 2008; Kato et al., 2009) have suggested that anomalously slow velocities are localized in the lower crust beneath the mainshock source area of the mid-Niigata Prefecture earthquake. These slow anomalies correlate well with the reflective layers delineated in the present study. Fig. 5b shows a depth section of S-wave perturbations along the same profile as the present study, interpolated from a regional (largescale) tomography model (Nakajima and Hasegawa, 2008). Although the spatial resolution of the regional model is fairly coarse (grid nodes with horizontal spacing of 0.1° and depth spacing of 0, 5, 10, 15, and 25 km), the lowermost crust beneath the seismogenic zone is characterized by anomalously slow velocities (shown as SLV in Fig. 5b). Note that the large-scale tomographic model was not able to delineate the ancient rift structure beneath the thick sediments at shallow depths, as imaged in the present study; it showed only low and high velocity layers at shallow depths. Additionally, a wideband magnetotelluric survey (Uyeshima et al., 2005) found a high electrical conductive body in the lower crust (deeper than 15 km) beneath the source region. The combination of anomalously low seismic velocities with high electrical conductivity in the lower crust is most plausibly explained by accumulations of overpressured hydrothermal fluids (e.g., Sibson, 2009). From geological observations of exhumed vein systems (Foxford et al., 2000; Tunks et al., 2004), it seems likely that fluids overpressured to approximately lithostatic values would occupy arrays of flat-lying extension fractures developed by hydraulic fracturing in a compressional stress regime where fluid pressure exceeded the vertical stress (Sibson, 2009). Because the minimum principal compressive stress is vertical in such a regime, the hydrothermally infilled extension fractures would be subhorizontal and elongated parallel to the profile in Fig. 5. The effect of strong fluid overpressuring to near-lithostatic values is to enhance connectivity between pore space and fractures, thereby increasing electrical conductivity and lowering seismic velocities in the rock mass (e.g., Gough, 1986).
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Fig. 5. (a) Final NERP reflection profile along the Y = 0 km with an assumption of PxP reflection mode. Arrows denote depths of relatively strong reflectors. Circles show the relocated aftershocks distributed within ± 4 km from the cross-section of Y = 0 km (gray circles observed in 2004, and red circles in 2005). The heavy line corresponds to the iso-velocity contours of Vp = 5.7 km/s in Fig. 4b. (b) Depth section of S-wave velocity perturbations along the same profile as in Fig. 5a, interpolated from the regional (large-scale) tomography model constructed by Nakajima and Hasegawa (2008). SLV means slow velocity anomaly in the lowermost crust.
5. Reactivations of boundary faults by ductile creeping of weak shear zones in the overpressured lower crust The fine-scale tomography image (Fig. 4b) suggests that the seismogenesis of the mid-Niigata Prefecture earthquakes was due primarily to compressional inversion tectonics involving reactivation of weakened boundary faults related to the Miocene rift system. Furthermore, the reflection study reveals the potential existence of overpressured hydrothermal fluids beneath the ancient rift system. Since the strength of the ductile creeping that is a dominant deformation process in the lower crust is lowered by crustal fluids (e.g., Carter and Tsenn, 1987), weak zones could develop beneath the ancient rift system (Iio et al., 2002; Kato et al., 2009). Thus, ductile creeping of weak shear zones in the lower crust may have loaded faults in the upper crust inherited from the Miocene rifting episode, leading to reactivation of this system in the form of intraplate earthquakes (e.g., Kenner and Segall, 2000; Kato et al., 2009). Because the lower curst was affected by Miocene rifting, the possibility also exists that some of the reflectors in the lower crust arise from mafic intrusions in the form of dyke–sill complexes (Hirata et al., 1989). Although regional tomographic studies (Nakajima and Hasegawa, 2008; Kato et al., 2009) suggest that the lower crust has slow anomalies on average beneath the source region, higher resolution studies are needed to resolve the fine-scale heterogeneity within the lower crust and determine the origin of the reflectors.
Given these considerations, we propose that inherited boundary faults within the buried ancient rift system were reactivated through activation of ductile shear zones in the overpressured lower crust. Hydrothermal overpressuring in the lower crust would likely lead to episodic invasion of overpressured fluids into the lower seismogenic zone through fault-valve action (e.g., Sibson, 2007, 2009). The present study provides additional evidence supporting such behavior. In central Italy in 1997, a shallow normal-faulting sequence occurred with accompanying clear earthquake migration and multiple mainshocks (Mw 5.2–6). It has been proposed that this sequence could be driven by the propagation of overpressured fluids trapped below the seismogenic layer through fault-damaged zones (e.g., Miller et al., 2004; Chiarabba et al., 2009). Overpressured fluids below the seismogenic layer are likely to trigger intraplate earthquakes even in different tectonic environments. It is therefore necessary to detect and monitor overpressured fluids below the seismogenic layer during an interseismic cycle. 6. Conclusions We deployed a dense portable seismic array within the source region of the 2004 mid-Niigata Prefecture Earthquake to understand fine-scale crustal heterogeneity. We conducted double-difference tomographic analysis with the first arrival data, and performed the
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natural earthquake reflection profile analysis with the relocated aftershock waveform data. (1) Relocated aftershocks were clearly aligned on both the steeply westward-dipping and the gently eastward-dipping fault planes that separate the sedimentary cover sequence in the hanging wall from the pre-Neogene basement in the footwall. (2) The basement is characterized by clear step-like tilted block structures that gradually deepen to the west. The block width is about 5 km. The step-like tilted block structures of the basement provide evidence of the Miocene rift system buried beneath the thick sediments. (3) Anomalously low seismic velocities combined with high electrical conductivity suggest that stacked reservoirs of overpressured fluids provide a plausible explanation for an array of relatively strong flat-lying reflectors in the lower crust. We propose that pre-existing boundary faults within the buried ancient rift system were reactivated by ductile creeping of weak shear zones in the overpressured lower crust. Acknowledgments We are thankful to H. Zhang for allowing us to use the TomoDD code. We also thank the research team for aftershock observations of the 2004 mid-Niigata Prefecture Earthquake for providing the arrival times and waveform data of aftershocks. We are grateful to one anonymous reviewer, C. Chiarabba and Editor H. Thybo for their constructive comments, which led to substantial improvement to the original manuscript. J. Nakajima kindly provided us with the regional tomography model. We are grateful to Kyoto University, Kyusyu University, the National Research Institute for Earth Science and Disaster Prevention, and the Japan Meteorological Agency for allowing us to use the waveform data collected at online stations. This research was partially supported by the Ministry of Education, Culture, Sports, Science and Technology (MEXT) of Japan, under its Observation and Research Program for Prediction of Earthquakes and Volcanic Eruptions, and by the cooperative research program of the Earthquake Research Institute. References Carter, N.L., Tsenn, M.C., 1987. Flow properties of the continental lithosphere. Tectonophysics 136, 27–63. Chiarabba, C., De Gori, P., Boschi, E., 2009. Pore pressure migration along a normal fault system resolved by time repeated seismic tomography. Geology 37 (1), 67–70. doi:10.1130/G25220A.1. Foxford, K.A., Nicholson, R., Polya, D.A., Hebblethwaite, R.P.B., 2000. Extensional failure and hydraulic valving at Minas da Panasqueira, Portugal: evidence from vein spatial distributions, displacements and geometries. J. Struct. Geol. 22, 1065–1086. Gough, D.I., 1986. Seismic reflectivity, conductivity, water, and stress in the continental crust. Nature 323, 143–144. Hirata, N., Matsu'ura, M., 1987. Maximum-likelihood estimation of hypocenter with origin time eliminated using nonlinear inversion technique. Phys. Earth Planet. Inter. 47, 50–61. Hirata, N., Tokuyama, H., Chung, T.W., 1989. An anomalously thick layering of the crust of the Yamato Basin, southeastern Sea of Japan: the final stage of back-arc spreading. Tectonophysics 165, 303–314. Iio, Y., Sagiya, T., Kobayashi, Y., Shiozaki, I., 2002. Water-weakened lower crust and its role in the concentrated deformation in the Japanese Islands, Earth Planet. Sci. Lett. 203, 245–253. Ito, K., 1999. Seismogenic layer, reflective lower crust, surface heat flow and large inland earthquakes. Tectonophysics 306, 423–433. JNOC, 1988. Report on basic geophysical exploration in onshore area “KubikiTamugiyama”. Japan National Oil Corporation, Tokyo. 45 pp. (in Japanese). Kato, A., the research team of aftershock observations for the 2004 mid-Niigata Prefecture Earthquake, 2007. High-resolution aftershock observations in the source region of the 2004 mid-Niigata Prefecture Earthquake. Earth Planets Space 59, 923–928. Kato, A., Kurashimo, E., Hirata, N., Iwasaki, T., Kanazawa, T., 2005. Imaging the source region of the 2004 Mid-Niigata prefecture earthquake and the evolution of a seismogenic thrust-related fold. Geophys. Res. Lett. 32, L07307. doi:10.1029/ 2005GL022366. Kato, A., Sakai, S., Hirata, N., Kurashimo, E., Iidaka, T., Iwasaki, T., Kanazawa, T., 2006. Imaging the seismic structure and stress field in the source region of the 2004 mid-Niigata Prefecture Earthquake: structural zones of weakness, seismogenic stress concentration by ductile flow. J. Geophys. Res. 111, B08308. doi:10.1029/2005JB004016.
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