JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 111, B03410, doi: /2005jb003656, 2006

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1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 111,, doi: /2005jb003656, 2006 Broadband hydroseismograms observed by closed borehole wells in the Kamioka mine, central Japan: Response of pore pressure to seismic waves from 0.05 to 2 Hz Yasuyuki Kano 1 and Takashi Yanagidani 1 Received 30 January 2005; revised 7 November 2005; accepted 20 December 2005; published 23 March [1] We obtained broadband hydroseismograms by monitoring the pore pressure changes of a rock mass in the Kamioka mine, using borehole wells. The wellhead was sealed to maintain an undrained condition, under which there is no flow of water through the interface between the well and the rock mass. This reduces the wellbore storage effect, which can cause a high-frequency cutoff response for systems of conventional open wells and rock mass. Using these closed borehole wells, 16 hydroseismograms were recorded for earthquakes in a range of magnitudes of and epicentral distances of Direct P waves, SV waves converted to P, and Rayleigh phases are clearly observed on the hydroseismograms. The similarity between hydroseismograms and seismograms reveals a clear relationship between radial ground velocity and pore pressure. The relationship is expressed as a zero-order system, which is characterized by no distortion or time lag between the input and output, and the pore pressure has no coupling with shear deformation. These results are consistent with an undrained constitutive relation of linear poroelastic theory and confirm that the relation is valid for the seismic frequency range. We determined in situ values of pore pressure sensitivity to volumetric change of the rock mass, which were then used to estimate in situ Skempton coefficients with values of Citation: Kano, Y., and T. Yanagidani (2006), Broadband hydroseismograms observed by closed borehole wells in the Kamioka mine, central Japan: Response of pore pressure to seismic waves from 0.05 to 2 Hz, J. Geophys. Res., 111,, doi: /2005jb Introduction [2] Water level fluctuations of open wells are sometimes observed due to the passage of teleseismic earthquake waves. In the past, there have been many reports that these oscillations of the water level resemble seismograms, and Roeloffs [1996] refers to the traces of well level changes as hydroseismograms. In this paper, we regard the traces of pore pressure changes, which we observed using closed wells, as hydroseismograms and examine the performance of the closed wells over a range of frequencies. The coseismic steps and persistent changes in well level occasionally observed at wells close to earthquakes are not addressed in this paper. [3] The first published observations of hydroseismograms were by Blanchard and Byerly [1935]. They referred to the apparatus that recorded the hydroseismogram as a phreatic seismograph and recognized the phase arrivals of P, S, and Rayleigh waves on their hydroseismogram. They suggested a simple model by assuming that the aquifer was an open cavity, which is below the ground surface and filled with 1 Disaster Prevention Research Institute, Kyoto University, Kyoto, Japan. Copyright 2006 by the American Geophysical Union /06/2005JB003656$09.00 water, and that a volumetric change of the cavity induced a change in the well level. They also suggested that the phase that corresponded to the arrival of the S wave can be attributed to the mode conversion of a shear wave to a compressional wave at the free surface. Further, they explained the absence of Love waves on the basis of the theoretical prediction that there is no dilatation of the cavity due to these waves. [4] Several observations and theoretical investigations to model the response of hydroseismograms have followed Blanchard and Byerly [1935]. In recent years, inversions for the hydraulic characteristics of the rock mass have focused upon using hydroseismograms from wells that penetrate the region [Eaton and Takasaki, 1959; Rexin et al., 1962; Cooper et al., 1965; Bredehoeft et al., 1965; Sterling and Smets, 1971; Liu et al., 1989; Ohno et al., 1997; Brodsky et al., 2003]. The response of open wells was modeled by taking into account the following two aspects: (1) pore pressure buildup in an aquifer, i.e., the formulation of a relationship between the total deformation of the rock mass and the accompanying change in pore pressure, and (2) wellaquifer interaction, i.e., the characterization of the change in well level caused by the pore pressure buildup in the aquifer, considering the water flow between the well and the aquifer. Cooper et al. [1965] formulated the pore pressure buildup and the well-aquifer interaction and 1of11

2 derived equations to model the frequency response curve of the hydroseismogram to the ground motion associated with seismic surface waves. In their formulation of the pore pressure buildup, they adopted an early version of linear poroelasticity. They approximated that the material constituents are incompressible; however, material constituents can be compressible in modern formulations of linear poroelasticity. The formulation of the well-aquifer interaction described by Cooper et al. [1965] successfully described the characteristics of observed hydroseismograms, such as the resonance of the water level to surface waves and the high-frequency cutoff response of hydroseismograms. It facilitated the estimation of diffusivity and the storage coefficient of the aquifer for water resource research. Although the response of the well was accurately modeled by the well-aquifer interaction, this interaction is not desirable for monitoring the total deformation of the poroelastic rock mass via pore pressure because the well-aquifer interaction masks the pore pressure buildup in the aquifer. [5] Detailed studies on the response of open wells have been carried out, and currently, the response is analyzed on the basis of a developed theory of linear poroelasticity, as described by Rice and Clearly [1976], Van der Kamp and Gale [1983], Rojstaczer and Agnew [1989], Roeloffs [1996], Kümpel [1997], and Wang [2000]. According to the theory, total stress changes that are introduced in the rock mass are shared by the skeletal framework and pore fluid when the pores are fully saturated. Furthermore, when the aquifer is under an undrained condition, i.e., when there is no pore water flow, the pore pressure change induced by a deformation is proportional to the mean stress change. The proportional coefficient, B, is called the Skempton coefficient. The validity of the undrained condition extends to high-frequency bands, if the wavelength of the deformation is much larger than the volume in which the pore pressure change is considered. Such a poroelastic system must have a constant gain and zero phase delay. From the viewpoint of system dynamics, the fully saturated system of the porous rock mass under an undrained condition can be expressed as a zero-order system [see Doebelin, 2004] which has no distortion or time lag between the input and the output. Here, the mean stress change is represented as the input, while the pore pressure change is represented as the output. [6] The passage of the seismic waves produces a stress change and therefore a pore pressure change of the rock mass. For a plane seismic wave, the total stress change is proportional to the particle velocity [Love, 1944]. In general, the deformation produced by a seismic wave occurs too rapidly for fluid flow to take place and the aquifer is assumed to be under an undrained condition. The pore pressure changes proportional to the total stress change and is therefore proportional to the particle velocity. [7] The well-aquifer interaction must be considered for open wells. Even if the aquifer itself is maintained in an undrained condition, a well drilled in an aquifer behaves as a drain path of the pore water. The pore pressure in the aquifer and the water head of the well equilibrate at the interface between the well and aquifer. Water must be squeezed into or out of the aquifer for aquifer pore pressure changes to be reflected in water level changes in open wells. The flow of water through the interface between the well and the aquifer is necessary, and it reduces the sensitivity along with causing a phase delay between the water level and pore pressure change in the aquifer at high frequency. The pore pressure change in the aquifer at high frequency cannot be accurately observed by the water level of the well since the water flow cannot keep up with the deformation. This reduction of sensitivity and phase delay is called the wellbore storage effect. In past studies that recorded seismic waves in open wells, the observed water level change, or hydroseismograms, consisted of only monotonic oscillations [e.g., Cooper et al., 1965; Bredehoeft et al., 1965; Liu et al., 1991; Ohno et al., 1997]. The seismic waves most likely to be observed by open wells are surface waves, especially Rayleigh waves. Body waves are characterized by much smaller amplitudes and much higher frequencies than surface waves. Since low-frequency surface waves were often recorded on hydroseismograms, reflecting the frequency response of the well, it was incorrectly concluded that the wells respond only to surface waves. This selection of recorded wave types is mainly the result of the frequency response caused by the wellbore storage effect. [8] The resonance and reduction of sensitivity at higher frequencies caused by the well-aquifer interaction can be avoided if the pore pressure of the rock mass is directly measured without an open well, which acts as a drain path and disturbs the undrained condition of the aquifer. In practice, however, it is impossible to measure the pore pressure without a well. One solution is to cap the well with an impermeable and rigid material such as a solid packer. If an artesian borehole, from which water flows out, is available, there is an easy way to make a closed well without using a packer: we can simply close the borehole at the wellhead. In this study, we monitor the pressure of an artesian well in order to reduce the wellbore storage effect. We assume that artesian wells and airtight boreholes behave as closed borehole wells, and pore pressure measurements of the rock mass are made with a thin needle sensor. If the pressure in the borehole is equal to the pore pressure in the aquifer, the observed hydroseismograms can be interpreted with a simple undrained constitutive relationship of linear poroelasticity. High-frequency body waves, such as P waves, can be observed by the closed borehole well, since the system has an expected flat response to relatively high frequencies. [9] With these considerations, we monitored the pressure in closed borehole wells from which water had flowed before the wellheads were closed, and we tried to observe pore pressure changes caused by the passage of seismic waves. We successfully recorded several hydroseismograms, which include rarely seen body waves. We characterize the performance of the closed borehole wells, analyzing both the hydroseismograms and velocity seismograms recorded at the ground surface. We will demonstrate a proportional relationship between the pore pressure of the aquifer and the ground velocity, considering the type of waves, such as P, S, and surface waves. These field observations will provide estimation for a range of in situ values of B. 2. Response of a Poroelastic Rock Mass to Seismic Waves [10] We predict the response of the pore pressure of the aquifer to teleseismic waves, by combining two expressions 2of11

3 for deformation: (1) the relationship between pore pressure change and stress change and (2) the relationship between stress change and ground motion, represented by ground velocity. We derive first the relationship between the stress change and the pore pressure change of an aquifer based on the theory of linear poroelasticity. The total stress change produced by the seismic waves is shared by the skeletal framework of the rock and pore water. If the water flow into or out of a reference element volume of the water-saturated rock mass is negligible with respect to the timescale of the passage of seismic waves, i.e., an undrained condition is maintained, the pore pressure change is proportional to the applied mean stress. The ratio of the pore pressure to the applied mean stress is defined as the Skempton coefficient, B [e.g., Roeloffs, 1996]. The relation between the undrained pore pressure change, P p, and the mean stress change, s m = (s xx + s yy + s zz )/3, is given by P p ¼ Bs m ; where the stress component is positive for extension and the pore pressure is positive for an increase. The minus sign implies that pore pressure increases correspond to compression. It is known that B is constant for small disturbances in the stress and that it depends on the type of rock and the effective confining pressure [Green and Wang, 1986; Lockner and Stanchits, 2002]. The total change in shear stress is solely supported by the skeletal framework of the rock since water cannot support shear stress. [11] The constitutive relationship between the stress, s ij, and the strain, e ij, for the isotropic, homogeneous, and undrained poroelastic rock mass is derived as follows: ð1þ n u s ij ¼ 2G d ij e kk þ e ij ; ð2þ 1 2n u where G is the shear modulus and n u is the undrained Poisson s ratio. The subscripts i and j represent the three directions x, y, and z; d ij is the Kronecker delta; and the repetition of subscripts implies the summation convention. [12] The ground surface is characterized by a free surface boundary condition; thus the stress component along the z direction is always zero. The free surface condition is applicable only to near-surface aquifers. Since we are considering a plane wave, the strain component e yy =0. Substituting these conditions into equation (2) yields s m ¼ 2G 3 1 þ n u 1 n u e xx : ð3þ Substituting equation (3) into equation (1) gives the undrained pore pressure buildup: P p ¼ B 2G 3 1 þ n u 1 n u e xx : ð4þ [13] We now turn to the second relationship that represents stress changes in an isotropic and homogeneous rock caused by seismic waves. We consider the total motion of the ground surface caused by seismic waves incident on the ground surface and in their reflections. We assume that the traveling seismic wave is a plane wave because this study analyzes waves at great distances from their sources. For a Cartesian coordinate system, the x axis represents the direction of the surface projection of the propagation of the seismic plane wave, or the radial direction. The z axis represents the depth direction perpendicular to the surface, and a plane defined by z = 0 represents the ground surface. The total displacement of the ground surface, u(x, t), at a point x for a time t is given by u ¼ u 0 exp½iwðpx tþš; ð5þ where u 0 is the sum of the amplitudes of displacements consisting of the incident and reflected waves, w is the angular frequency, and p is the horizontal slowness. The time and space derivatives of u are ¼ e xx ; e xx is eliminated from equation (4) by using equation (6) as follows: P p ¼ B 2G 3 1 þ n u 1 n : The sensitivity, c p (p), of pore pressure change P p to the ground velocity, v r u/@t, is c p ðpþ ¼ P p ¼ 2Ggp; v r where loading efficiency, g, is given by ð6þ ð7þ ð8þ g ¼ B ð 1 þ n uþ 31 ð n u Þ : ð9þ Equation (9) shows a proportional relation between v r and P p, i.e., a zero-order system, under an undrained condition. c p depends on p, which implies that c p varies for waves that have different slowness, even within the same hydroseismogram. [14] For the case of an incident plane P wave, the wave is a propagation of volumetric changes and thus produces pore pressure changes. Although incident P waves produce reflected P waves and SV waves, only the incident and reflected P waves produce pore pressure changes. [15] An S wave is a propagation of shear deformation. There are no volume changes, and thus no pore pressure changes are induced in the isotropic poroelastic rock mass associated with an S wave. The case for incident SH waves is simple. All of the energy of an incident SH wave is reflected as an SH wave without conversions. SH waves induce no pore pressure change at the ground surface. In contrast to the case of incident SH waves, SV waves induce volumetric changes at the surface. The incident SV wave is partly converted to a reflected P wave at the free surface to satisfy the boundary conditions, i.e., s zx = s zz =0onz =0. The arrivals corresponding to the SV waves induce pore pressure changes due to volumetric changes caused by the converted P waves. S phases are propagation of shear 3of11

4 Figure 1. (a) Map of the observation sites. (b) Map of the Mozumi-Sukenobu fault (MSF). The Atotsugawa fault system consists of the Atotugawa fault (AF), MSF, and the Ushikubi fault (UF). The pore pressure is measured at MOZ and the ground velocity is measured at KTJ. (c) Map view of the MOZ tunnel (d) Cross section oriented in a N30 W S30 E direction. Solid circles show the wellheads of the closed borehole wells, and the open circle shows the site for barometric pressure measurement. The fault crushed zone is shown by arrows. deformation, but if the S phase contains an SV component, they induce pore pressure changes and thus can be observed in hydroseismograms. [16] To apply equation (8) to Rayleigh waves, p is replaced by the reciprocal of the phase velocity. Following the analysis of Cooper et al. [1965], the response of the well water level to a Rayleigh wave has been compared to the vertical component of ground velocity [Bredehoeft et al., 1965; Liu et al., 1989; Ohno et al., 1997; Brodsky et al., 2003]. Cooper et al. [1965] derived the relationship between ground velocity and volumetric strain based on elastic theory. The approach of Cooper et al. [1965] is identical to our derivation of the relation between particle velocity and pore pressure change, employing the proportionality of particle velocity and strain, although the proportionality constant is different. For the case of Love waves, no pore pressure change is induced for an assumed isotropic poroelastic rock mass. Love waves are propagation of pure shear deformation, similar to SH waves, and produce no volumetric change of the aquifer. [17] Summarizing the derivation, changes of pore pressure for an aquifer are associated with seismic waves that produce volumetric changes of the aquifer. The change of pore pressure, P p, is proportional to the product of the radial ground velocity, v r, and slowness, p, of the wave. The pore pressure has no shear coupling. Hydroseismograms obtained by closed borehole wells should resemble the radial component of velocity seismograms. The amplitude ratios of the seismic waves in the hydroseismogram to the radial seismogram depend on the slowness of the specific wave, so the relative amplitudes between the hydroseismograms and the radial velocity seismogram differ for different types of waves. The relation between vertical ground velocity and pore pressure change can be derived by combining the relationship between vertical ground velocity and e zz. The proportionality constant is a function of not 4of11

5 only p but also the P and S wave velocities at the surface, unlike the case for the radial component, which only depends on p. To avoid increasing the number of unknown parameters, such as P and S wave velocities, we do not discuss the vertical ground velocity. 3. Observation Sites [18] We selected the Mozumi observation (MOZ) tunnel in the Kamioka mine, central Japan, as an optimal site for measurement of pore pressure, using a closed borehole well, to obtain broadband hydroseismograms. Several artesian boreholes appropriate for closed borehole wells have been drilled in the MOZ tunnel. During the Active Fault Probe of the Earthquake Frontier Research Project [Ando, 2002], in 1996, the MOZ tunnel was excavated through the Mozumi- Sukenobu fault, which is a part of the Atotsugawa fault system. The tunnel is located at N and Eatan elevation of 360 m. The tunnel extends horizontally in a N24 W direction. The overburden of the observation tunnel is around m (Figure 1d). The bedrock surrounding the tunnel consists of alternating sandstone and shale. Two zones of intense deformation, or fault crushed zones, were identified by geologists from the Mitsui Mining and Smelting Co. [19] We carried out continuous pore pressure measurements using two artesian boreholes in the MOZ tunnel. The borehole indicated by A in Figures 1c and 1d, is drilled into the wall of the tunnel, 20 m from the fault crushed zone. The axis of the borehole is almost horizontal (strike N100 W, dip 5 ) and its length and diameter are 15.5 m and 76 mm, respectively. The borehole is uncased along its entire length, except for the 1-m-long pipe at the wellhead. Another borehole, indicated by C in Figures 1c and 1d, is tapped vertically at the old mine tunnel, 100 m from the fault crushed zone. The rock where borehole C is tapped is more competent in contrast to where borehole A is tapped. The direction of the borehole is almost vertical (strike N90 W, dip 70 ) and its length and diameter are 600 m and 140 mm, respectively. The borehole is cased with steel pipe, but no information about the shape and depth of the screen, at which the borehole is connected to the aquifer, is reported. Before closing the borehole for pore pressure measurements, the outflow of water from boreholes A and C was 25 and 375 L min 1, respectively. The boreholes can be referred to as artesian wells. [20] We closed these wells at the wellheads and installed pressure transducers (Honeywell PPTR0300AP). The transducers consist of a sensing element that is a semiconductor strain gauge welded on a diaphragm and contain an excitation circuit of the strain gauge and an analog-to-digital (A/D) converter. The digital data are transmitted via the RS232C port. The resolution of the sensor is 16 bit, relative to a full scale of 2.07 MPa, and the response delay is 17 ms. We set the sampling rate at 20 Hz. After closing the wellhead of the two boreholes, the pressure inside the borehole immediately reached 200 kpa for borehole A and 700 kpa for borehole C. The pressure then gradually increased to 1.0 MPa for borehole A and 1.4 MPa for borehole C over about a half year and has likely reached a final equilibrium state. Figure 2. Waveforms of the 28 June 2002 Vladivostok earthquake. Hydroseismograms of (a) borehole A and (b) borehole C and seismograms for the (c) radial, (d) transverse, and (e) vertical components. [21] We simultaneously measured barometric pressure, because it is well known that the pore pressure of an aquifer responds to barometric pressure changes. The barometric response is also expected in pore pressure measured by closed borehole wells. Although the details of the barometric response is discussed in a companion paper (Y. Kano and T. Yanagidani, manuscript in preparation, 2006), the effect can be neglected for measuring seismic waves because the frequency band of the barometric pressure changes ( Hz) is much less than that of the of seismic waves. [22] The ground velocity is recorded using a three-component Streckeisen STS-1 seismograph installed at Kamitakara observatory (KTJ, latitude N, longitude E, altitude 760 m), located 20 km east of MOZ. Seismograms are recorded at a sampling rate of 20 Hz. The full scale is 7.2 m s 1 and the resolution of the A/D conversion is 24 bit. [23] The sites at which the hydroseismograms and seismograms were obtained are not the same but are separated by a distance of 20 km; however, this distance is within a wavelength of the seismic waves analyzed. There are 5of11

6 Figure 3. Traces produced by applying a low-pass Butterworth filter with a 40 db decade 1 roll-off at 0.1 Hz to data in Figure 2. The SP phase that arrives faster than the S phase on the transverse component is observed on the radial and vertical components. difficulties in keeping accurate time in the MOZ tunnel, so the traces of the hydroseismograms are aligned by matching the arrivals of the P waves to the arrivals in the seismograms. [24] We obtained 16 hydroseismograms during our observation period from 3 April 2002 to 30 November The magnitudes, M w, of the earthquakes that produced the hydroseismograms at MOZ ranged from 4.5 to 7.9 and the epicentral distance, D, ranged from 1.0 to It should be noted that no coseismic steps were observed on the hydroseismograms for this period. 4. Results [25] We characterize the hydroseismograms by examining two typical records from 16 earthquakes that produced hydroseismograms at MOZ. One is a deep-focus earthquake, which contains mainly body waves, and the other is a shallow earthquake, which is dominated by surface waves. The hydroseismograms and seismograms of the M w 7.3 deep-focus Vladivostok earthquake on 28 June 2002 (depth 565 km, epicentral distance 9.1 ) are shown in Figure 2. Figures 2a and 2b depict hydroseismograms of boreholes A and C, and Figures 2c to 2e illustrate threecomponent seismograms, where the traces of the horizontal motions are converted to radial and transverse components. The hydroseismograms closely resemble the radial components and differ significantly from the transverse and vertical components. The P and S phases are clearly observed on both the hydroseismograms and the seismograms. This comparison between the hydroseismograms and seismograms is consistent with the proportionality between pore pressure and ground velocity mentioned in section 2. The absence of surface waves on both the hydroseismograms and the seismograms is explained by the fact that the Vladivostok event is a deep-focus earthquake and the epicentral distance is small. The shapes of the waveforms of the hydroseismograms of boreholes A and C are nearly the same, although the amplitude of the hydroseismogram of borehole A is smaller than that of borehole C. So, the two hydroseismograms are characterized by the same frequency response with a difference in gain. [26] Figure 3 depicts traces that are produced by applying a low-pass Butterworth filter with a 40 db decade 1 roll-off at 0.1 Hz to the traces in Figure 2 in order to reduce the high-frequency variations produced by small-scale geological structures. The filter also attenuates the seismic waves with wavelengths less than the distance between MOZ and KTJ. [27] The phases of the P and S arrivals on the filtered hydroseismograms and seismograms are in good agreement; the pore pressure tracks the radial ground velocity. The ratio of the peak-to-peak amplitudes of the P and S phases in the hydroseismograms differs slightly from that of the radial seismogram. To explain the different amplifications of the pore pressure change to different phases, such as P and S, we calculate, c p (equation (8)). We extract one cycle of the P and S waves from the hydroseismograms and velocity seismograms and estimate c p by comparing the amplitudes of the two waveforms point by point, and we use a least squares fit to find the best proportionality constant between the two waveforms. To verify the goodness of fit, we calculate the correlation coefficient, R, which is equal to unity if a hydroseismogram completely agrees with the product of the seismogram and c p. Table 1 lists the values of c p and R for boreholes A and C. We determine p from the depth and epicentral distance of the event, referring to Table 1. The Parameters c p and R Determined by Comparison of the Hydroseismograms and Seismograms Phase P, skm 1 A c p, kpa mm 1 s 1 R C c p, kpa mm 1 s 1 Vladivostok P S Mariana P S Rayleigh R 6of11

7 Figure 4. Waveforms of the 26 April 2002 Mariana earthquake. Hydroseismograms of (a) borehole A and (b) borehole C and seismograms for the (c) radial, (d) transverse, and (e) vertical components. Earth model iasp91 [Kennett and Engdahl, 1991]. The dependence of c p on the slowness p is clear, with c p increasing proportionally to p, as predicted by the theory. [28] Figure 4 illustrates the traces of the M w 7.1 Mariana earthquake on 26 April 2002 (depth 55 km, epicentral distance 24.0 ) with the horizontal components oriented in the radial and transverse directions. In contrast to the Vladivostok earthquake, the maximum amplitudes are observed during the passage of the surface waves. The surface waves are dominant on the hydroseismograms and seismograms because of strong excitation caused by the shallow hypocenter of the Mariana earthquake and reduction of high-frequency content of the body waves because of attenuation effects from the propagation over a large epicentral distance. The phase arrivals of the P and S waves are also seen on both the hydroseismograms and seismograms, although they are smaller than the surface waves. The waveforms of the hydroseismograms and seismograms are similar for the P and S phases, as was pointed out for the Vladivostok event. Figure 5 shows the traces of Figure 4 7of11

8 Figure 5. Traces produced by applying a low-pass Butterworth filter with a 40 db decade 1 roll-off at 0.1 Hz to the data in Figure 4. that were low-pass filtered in order to compare the hydroseismograms with the seismograms. [29] There is very good agreement of the waveforms of the hydroseismograms and the radial seismogram for the surface waves. The hydroseismograms contain Rayleigh waves that are observed on the radial component of the seismogram but do not contain Love waves that are observed only on the transverse component of the seismogram. Love waves are pure shear deformation and produce no volumetric changes. The absence of Love waves on the hydroseismogram is consistent with the theoretical expectations, which show that the pore pressure has no response to shear deformation. The agreement in waveforms of the Rayleigh waves between the hydroseismograms and the radial seismogram confirms that pore pressure responds as a zero-order system to the volumetric change of the rock mass caused by Rayleigh waves, similar to the response to the P and S phases. The values of c p and R were calculated for the P, S and Rayleigh waves, as described for the Vladivostok earthquake, and are listed in Table 1. Among the 16 8of11

9 Figure 6. (a) Amplitude spectrum of hydroseismogram and (b) amplitude spectrum of seismogram. The solid line shows the amplitude spectrum calculated for 25 s of data from arrival time of the P phase. The dashed line shows noise level, which is the amplitude spectrum calculated for the data before the arrival of seismic waves. (c) Gain of the frequency response function of pore pressure change to radial ground velocity obtained by dividing the amplitude spectrum of the hydroseismogram by that of the radial component seismogram. The gain is not calculated above 2 Hz because of low S/N ratio in the hydroseismogram. recorded hydroseismograms, the Vladivostok and Mariana events are the only events that have the data quality and well-separated phase arrivals necessary to calculate c p. [30] Water level measurements in open wells to observe pore pressure changes of the rock mass suffer from a highfrequency cutoff response that is caused by the wellbore storage effect. Consequently, it is difficult to observe body waves usually dominated by high-frequency waves. The closed borehole well has a zero-order response to the stress induced by the passage of the seismic wave, as shown by direct comparison between the hydroseismograms and seismograms. A closed borehole well theoretically has no limit in response to the pore pressure changes of the aquifer for the higher frequency range, if it measures the pore pressure of a small volume of the rock mass in which the flow of water is negligible. [31] We examine the high-frequency cutoff of pore pressure measured in the closed borehole well, by comparing frequency spectra of the hydroseismograms and seismograms. The analysis is limited by the frequency content of the incoming seismic waves and the frequency responses of the measurement systems. The Nyquist frequency is 10 Hz for the sampling rate of 20 Hz of both the pressure gauges and the seismometer. The STS-1 seismometer installed at KTJ and the pressure gauge have flat responses between about and 7.5 Hz. [32] Figures 6a and 6b illustrate the amplitude spectra of the borehole C hydroseismogram and the radial component of the seismogram, respectively, for the Vladivostok earthquake. The amplitude spectra of 25 s of the P arrivals are shown by the solid lines, and the amplitude spectra of 25 s before the first arrivals are shown by dashed lines to indicate the noise levels of the hydroseismogram and seismogram. The amplitude spectra of both the hydroseismogram and the seismogram have a similar shape and fall off with the same slope at higher frequencies. The examination of the frequency response of the hydroseismogram and seismogram is limited to the frequency band below 2 Hz because of the insufficient signal-to-noise ratio of the amplitude spectrum of the hydroseismogram above 2 Hz. The hydroseismogram has no high-frequency cutoff response in this frequency band. Figure 6c illustrates the frequency response of the hydroseismogram to the radial component of the ground motion, which is obtained by dividing the amplitude spectra of the hydroseismogram (Figure 6a) by the radial component of the seismogram (Figure 6b). The frequency response of the hydroseismogram to the radial component of the ground motion is flat below 2 Hz. These observations confirm that the pore pressure and radial ground velocity are proportional at frequencies below 2 Hz. The gain, or spectral ratio, of the hydroseismogram to the seismogram gives a value of c p 4 kpa mm 1 s 1 below 0.1 Hz, which 9of11

10 Table 2. The Parameter 2Gg Determined From the Observed c p 2Gg, GPa Event Phase A C Vladivostok P S Mariana P S Rayleigh Average Standard deviation 3 6 is consistent with the values estimated in the time domain (Table 1). The increase of gain at higher frequencies may correspond to a difference in the small-scale geological structure, which causes amplification of the high-frequency content between MOZ and KTJ. 5. Discussion 5.1. Advantages of a Closed Borehole Well Over an Open Well [33] In previous studies, the detection of both body and surface waves on hydroseismograms from open wells has been rare. It had been thought that the hydroseismograms of open wells responded only to low-frequency Rayleigh waves. We showed from the pore pressure measurements at MOZ that the hydroseismograms obtained by closed borehole wells record all phases that produce volumetric changes of the aquifer, such as P phases, Rayleigh waves, and S phases which consist of SV and converted P waves. The resemblance of hydroseismograms and radial seismograms is predicted by the theoretical consideration that the pore pressure change induced by passage of seismic waves is proportional to the product of the radial ground velocity and slowness of the phase. This result was confirmed by our observations. [34] Closed borehole wells detect a variety of phases compared to open wells, and past studies concluded that the water level responded mainly to surface waves, especially Rayleigh waves [Blanchard and Byerly, 1935; Rexin et al., 1962; Eaton and Takasaki, 1959; Brodsky et al., 2003]. In those studies, surface waves were the most significant phase, although body waves were sometimes observed. They focused on surface waves to simplify the treatment and avoided using various phase velocities. Hydroseismograms from both closed borehole wells and open wells represent the pore pressure changes of the aquifer caused by seismic waves. These pore pressure changes are predicted by linear poroelasticity, as suggested in many previous studies. Differences in the processes of transmitting pore pressure changes of the aquifer are probably the cause of the discrepancy between closed borehole wells and open wells. The pore pressure change can be directly measured by closed borehole wells, while water levels of open wells are not a simple representation of the pore pressure of aquifer. This is because of the wellbore storage effects, which are a function of the transmissibility between the well and aquifer, in addition to the geometry of the well [Cooper et al., 1965; Liu et al., 1989]. [35] Wellbore storage effects introduce a particular frequency response to the well-aquifer system, which is inherently included in hydroseismograms obtained by water level measurements of open wells. This causes oscillations in the open well and the resonant frequency usually falls in the frequency range of seismic surface waves [Bredehoeft et al., 1965; Liu et al., 1989; Ohno et al., 1997]. Also, the sensitivity reduction at the frequencies above the resonance frequency limits the observations of body waves. In contrast to open wells, closed borehole wells should have a flat response even at high frequencies. As shown in this study, body waves can be observed by measurements in closed borehole wells. Also, the comparison of amplitude spectra of recorded hydroseismograms and seismograms confirms the flatness of the frequency response of the closed borehole up to 2 Hz In Situ Determination of Loading Efficiency [36] We can determine the loading efficiency, g, by in situ measurements of pore pressure and ground velocity according to equation (8). In situ values of g derived from the values of c p and assumed p (Table 1) for several phases, are listed in Table 2. The average value of 2Gg is 25 GPa for borehole A and 33 GPa for borehole C. The standard deviation is 3 GPa for borehole A and 6 GPa for borehole C. This is the first determination of g in this high-frequency range of 0.1 Hz. [37] In laboratory experiments, the Skempton coefficient, B, rather than g, is usually measured. B is a measure of the pore pressure buildup coefficient for isotropic loading, which is rarely achieved in the field. On the other hand, g is a measure of the pore pressure buildup coefficient under uniaxial strain, which is difficult to maintain in the laboratory. If we assume an undrained Poisson ratio, n u, and shear modulus, G, g determined above provides an estimation of B. B is estimated to be for borehole A and for borehole C, assuming n u takes a value between 0.30 and 0.35, which is typical for the rocks at sites similar to ours. G is 8 GPa for borehole A and 15 GPa for borehole C. Considering the uncertainty of G, 10% changes of G result in 10% changes of B. In situ B estimated at MOZ is larger than values determined for sandstones in the laboratory [Wang, 2000, Table C.1], although the difference of rock type should be taken into account. This may be attributed to the difference of effective confining pressure between MOZ and the laboratory experiments of Wang [2000]. It is known that B depends on the effective confining pressure [Green and Wang, 1986; Lockner and Stanchits, 2002], and it reaches unity at low confining pressures. At MOZ, the overburden is 350 m and the pore pressure is 1.0 and 1.5 MPa for borehole A and borehole C, respectively; thus the effective confining pressure is MPa, assuming that the density of the rock is 2700 kg m 3. Taking effective confining pressure into account, the estimated value of B at MOZ is consistent with the laboratory measurements for sandstones by Green and Wang [1986] and Lockner and Stanchits [2002] which indicate that B = for this range of confining pressure. [38] Another method to determine in situ values of g is analysis of the barometric response. The relationship between g for the seismic frequency band and for the lower frequency band that contains the barometric pressure response and tidal response ( Hz) is reported in a 10 of 11

11 companion paper (Y. Kano and T. Yanagidani, manuscript in preparation, 2006) S and Love Waves on Hydroseismograms [39] Love waves were not detected on our hydroseismograms. The arrivals of S waves, which appeared on the hydroseismograms, can be successfully explained by SV-P conversion at the ground surface, taking into account the slowness of the S phase and thus the conversion coefficient. These results are consistent with the predictions of isotropic linear poroelasticity (equation (1)), i.e., there is no coupling between shear stress and pore pressure. The idea that the hydroseismograms do not record Love waves has previously been indicated by Blanchard and Byerly [1935], though it has not been clearly confirmed by observation. [40] There are previous reports that the water level measurements recorded Love waves [Rexin et al., 1962; Sterling and Smets, 1971; Brodsky et al., 2003]. Brodsky et al. [2003] mentioned that in fractured rock where the water flow or transmission of pore pressure is dominated by linear formation of fractures, pore pressure changes could be produced by shear waves. They did not confirm the appearance of Love waves on hydroseismograms, but focused solely on Rayleigh waves, which have inherent volumetric changes. [41] Brodsky et al. [2003] also reported that arrivals corresponding to S phases are detected on hydroseismograms obtained in open wells, following the 2002 Denali, Alaska, earthquake. S waves are the propagation of shear deformations and S waves alone produce no volumetric change of a rock mass, and consequently should produce no pore pressure change in an isotropic rock mass. They explained the shear phase coupling observed for this and other earthquakes by an anisotropic poroelastic response. [42] Shear coupling of pore pressure may occur in an anisotropic rock mass for our particular sites. From our observations in the MOZ tunnel, however, we showed that the waves that appeared on hydroseismograms of closed borehole wells correspond to volumetric changes of rock mass produced by P, SV (which converts to P), and Rayleigh waves. Furthermore, no coupling of pore pressure to shear deformation is observed. Although one of the boreholes is drilled close to the fault crushed zone, which is considered to have anisotropic properties, the anisotropy of the rock mass may not cause large effects on the poroelastic deformation at MOZ. The anisotropic effects appear to be negligible and the pore pressure changes can be treated with isotropic poroelasticity. 6. Conclusions [43] Pore pressure measurements in the Mozumi tunnel recorded hydroseismograms showing clear arrivals of P, S, and Rayleigh waves from regional and teleseismic earthquakes. The waveforms of the hydroseismograms are similar to the radial component of ground velocity recorded on a nearby broadband seismometer. The waveforms for these arrivals (and the lack of recorded Love waves) can be explained with linear isotropic poroelastic theory. The pore pressure responds directly to the volumetric strain change, and when the phase velocity is properly considered, the water pressure changes are shown to be proportional to the radial component of ground velocity. The ratio of the amplitude of the pore pressure changes to the isotropic loading give an estimate for the in situ Skempton coefficient With consideration of the overburden effects, the estimated values for the Skempton constant derived from the Mozumi field observations are consistent with laboratory values for sandstone. [44] Acknowledgments. We thank J. Mori for useful comments. We also thank John Townend and two anonymous reviewers for fruitful discussions that facilitated improvement of the manuscript. The mining section of Kamioka mine provided technical support for our measurement. References Ando, M. 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Smets (1971), Study of Earth tides, earthquakes and terrestrial spectroscopy by analysis of the level fluctuations in a borehole at Heibaart (Belgium), Geophys. J. R. Astron. Soc., 23, Van der Kamp, G., and J. E. Gale (1983), Theory of Earth tide and barometric effects in porous formations with compressible grains, Water Resour. Res., 19, Wang, H. F. (2000), Theory of Linear Poroelasticity With Applications to Geomechanics and Hydrogeology, 287 pp., Princeton Univ. Press, Princeton, N. J. Y. Kano and T. Yanagidani, Disaster Prevention Research Institute, Kyoto University, Gokasho, Uji, Kyoto , Japan. (kano@eqh. dpri.kyoto-u.ac.jp) 11 of 11