Self-potential distribution on active volcano controlled by three-dimensional resistivity structure in Izu-Oshima, Japan

Transcription

1 Geophys. J. Int. (2009) 178, doi: /j X x Self-potential distribution on active volcano controlled by three-dimensional resistivity structure in Izu-Oshima, Japan Shin ya Onizawa, 1 Nobuo Matsushima, 1 Tsuneo Ishido, 1 Hideaki Hase, 1,2 Shinichi Takakura 1 and Yuji Nishi 1 1 Geological Survey of Japan, AIST, Central 7, Higashi 1-1-1, Tsukuba, Ibaraki , Japan. s-onizawa@aist.go.jp 2 Graduate School of Science, Hokkaido University, N10W8, Kita-ku, Sapporo, Hokkaido , Japan GJI Volcanology, geothermics, fluids and rocks Accepted 2009 April 1. Received 2009 March 6; in original form 2008 October 15 1 INTRODUCTION Izu-Oshima is an active island volcano in Japan. Associated with the latest eruptions in , various phenomena were observed: volcanic tremors originating around the groundwater table below the summit (Watanabe 1987); changes in the geomagnetic intensities (which suggested the expansion of demagnetized regions; Hamano et al. 1990); appearance of new fumaroles at the summit area; temporal changes in the relative contributions of magmatic, meteoric and sea water-source fluids in the fumarole gases (Kazahaya et al. 1993); remarkable temperature rise of the groundwater near the shoreline to the northwest of the summit (Takahashi et al. 1991) and so on. These phenomena suggested the interaction of magma and/or volcanic gases with groundwater and the development of a hydrothermal system beneath the summit area. Between 1989 March and 1994 March, annual self-potential (SP) surveys were carried out on the Izu-Oshima island (Ishido et al. SUMMARY A self-potential (SP) survey is conducted in Izu-Oshima volcano to reveal subsurface fluid flows during a dormant period. At the southwestern slope outside the caldera, terrain-related SP is observed. Higher potentials are observed in the southern caldera, where fumarolic and thermal activities are found. Contrarily, lower potentials are observed in the northern caldera. This region coincides with an area where lavas pile thicker than that on the outside. 3-D simulations of the groundwater flow due to meteoric water infiltration and induced SP are performed with a heterogeneous electrical resistivity model, to understand the fundamental groundwater flow regime and causes of the observed SP. The observed SP variation is strongly affected by the subsurface resistivity structure and streaming potential within the unsaturated zone, and an overall pattern of the observed SP is reproduced without including an effect of thermally driven upward flows. The terrain-related SP at the southwestern slope is caused by the downward flow of the meteoric water in the unsaturated zone as proposed by previous researchers. The primary cause of the high SP observed in the southern caldera is shallow conductors connecting with a lower conductive layer. Contribution from the upward fluid flows due to the thermal activities, which was believed to be the main cause of the high SP at active areas, is thought to be smaller. The lower potential at the northern caldera is interpreted by the greater streaming potential as compared to that of the surroundings. This can be caused by the difference in the electrical and/or hydraulic parameters of lavas compared with those of pyroclastic rocks. One of the probable reasons of the greater streaming potential is the greater magnitude of zeta potentials of lavas. Probably, lower vertical permeabilities of lavas also lead to the greater streaming potential. Key words: Electrical properties; Magnetotelluric; Hydrogeophysics; Hydrothermal systems; Permeability and porosity; Volcano monitoring. 1997). A terrain-related SP distribution of about 1 mv m 1 of elevation was observed outside the caldera in all of the five surveys. Inside the caldera, SP increased from about 350 mv to almost 0 mv (relative to the coastline), as the summit crater was approached. The SP inside the caldera decreased by about 100 mv between 1989 March and 1990 March surveys, which appeared to be correlated with a significant decline in the degassing rate from the summit crater. After 1990, the SP distribution was quite steady along the entire survey line, which extends from the west coast through the southern part of the caldera and ends at the eastern flank. High-potential anomalies at the summit and thermally active areas were also observed on a number of volcanoes (e.g. Zablocki 1976; Sasai et al. 1997; Michel & Zlotnicki 1998; Kanda and Mori 2002; Aizawa 2004; Hase et al. 2005), and remarkable changes in the SP were detected in association with the volcanic activities at volcanoes, such as Kilauea (Zablocki 1976), Unzen (Hashimoto & Tanaka 1995) and Piton de la Fournaise (Michel & Zlotnicki 1998; 1164 C 2009 The Authors Journal compilation C 2009RAS

2 SP controlled by resistivity structure in Izu-Oshima volcano 1165 Revil et al. 2003). In most cases, the streaming potential associated with the thermally driven upflow was believed to be the primary cause of these positive anomalies. Recently, the SP measurement was also applied to reveal the location and pattern of groundwater flow in volcanoes such as Stromboli in combination with CO 2 soil gas degassing and temperature measurements (Finizola et al. 2003) and further with electrical resistivity tomography (e.g. Revil et al. 2004; Finizola et al. 2006). Revil et al. (2008) performed a fluid flow simulation, incorporating these multiparameter observations, to reveal the hydrothermal system of La Fossa di Vulcano. To quantitatively clarify the SP generation mechanisms in a volcanic edifice, Ishido et al. (1997) performed numerical modelling for the first time. They first carried out numerical simulations to reproduce the subsurface fluid flow and the distribution of pertinent parameters, such as temperature, salinity, liquid-phase saturation and so on, and then calculated the SP from the underground conditions computed via the first-step fluid flow simulation by using the so-called EKP postprocessor (Ishido & Pritchett 1999). Ishido (2004) applied the same technique to investigate the mechanisms responsible for the generation of W -shaped SP profiles observed on volcanoes such as Miyakejima (Sasai et al. 1997), Izu-Oshima (Ishido et al. 1997) and so on by taking into account the presence of a thick unsaturated zone. He showed that the primary cause of the SP pattern is the drag current associated with downward liquid flows in the unsaturated and underlying saturated layers within the volcanic edifice, which decrease and increase the electric potential in the shallow and deep regions, respectively. If a conductive structure extending to deep levels is present below the summit area, the SP around the summit is substantially increased, resulting in the characteristic W -shaped SP profile. The calculated high SP amplitude around the summit is sensitive to the conductivity structure, which is thought to change over time due to volcanic activities such as magma ascent, development of hydrothermal convection and so on. We conducted SP and audiofrequency magnetotelluric (AMT) measurements at the Izu-Oshima island in 2006 and 2007, as a part of field studies to monitor the volcanic activities by continuous and/or repetitive SP and AMT measurements. In this paper, we present a SP map over the central part of the Izu-Oshima island and resistivity sections crossing the caldera region, both of which were taken under the quiet condition of the Izu-Oshima volcano, and then discuss the subsurface hydrology based upon these data as well as other geophysical and geological data. We develop a 3-D numerical model of the subsurface fluid flow and calculate the SP distribution on the ground surface from the underground conditions as computed from the fluid flow simulation. 2 SP AND AMT OBSERVATIONS Izu-Oshima island is located 120 km south-southwest of Tokyo, Japan (Fig. 1a). Oshima volcano a current active volcano is underlain by late Pliocene or Pleistocene volcanoes, which are exposed from the eastern to northern cliffs. Throughout its history, most of the erupted materials have been basalt or basaltic andesite. Above sea level, the edifice is mainly composed of the alternation of Figure 1. (a) Location of Izu-Oshima island. (b) Topographic relief map of Izu-Oshima island and observation sites. A rectangle indicates the area of (c). Purple dots: SP measurement points; red circles with a numeral: audiofrequency magnetotelluric surveys in (Takakura et al. 2007); yellow circles: ELF and VLF magnetotelluric surveys in (Utada and Shimomura 1990); light blue inverse triangles: boreholes where the groundwater level is observed (Takahashi et al. 1988; Nakada et al. 1999); green triangles: springs (Takahashi et al. 1988). (c) Close-up map of the caldera area. Journal compilation C 2009 RAS

3 1166 S. Onizawa et al. scoria, lava and volcanic ash layers. Tuff breccias become dominant, approximately, below the sea level by reflecting its growth under the sea. The basement layer comprises altered volcanic and volcaniclastic sedimentary rocks found in boreholes at about 400 m below sea level (BSL) in the northwestern part and at about 250 m BSL in the central part of the island (e.g. Nakamura 1964; Isshiki 1984a; Nakada et al. 1999). In the central part of the island, there is a caldera with a diameter of 3 km or more (Figs 1b and c). A central cone, called Mt. Mihara, with the highest altitude of 764 m, rests on the southern part of the caldera. During more than one century at least, most eruptions occurred on Mt. Mihara (Isshiki 1984b), although the latest eruptions in occurred from two newly opened fissures inside and outside the caldera as well as from Mt. Mihara. 2.1 SP mapping SP mapping was conducted in 2006 March and 2007 April and June. The main target was the caldera region where most of the eruptions occurred historically and fumarolic and thermal activities are cur- rently found. Further, a flank area outside the caldera was measured as the reference. The measurements were made with silver silver chloride non-polarizing electrodes and a high-impedance voltmeter (internal resistance >1 G ). For each survey line, the maximum wire length (from a fixed base electrode) and the measurement intervals were typically 1 2 km and m, respectively. The locations of the electrodes were measured with a handy GPS (accuracy was ±6 m). The total number of measurement points was 613. Closure offsets were relatively small, for example, an 1850 m loop traversing the central cone was closed with an offset of less than 10 mv. An offset of a caldera-scale loop enclosed within 5 d was 25 mv. The observed SP distribution is shown in Fig. 2(a). A common ground reference is chosen near the southwestern coast of the island. The main features of the observed distribution are as follows. As commonly observed in many volcanoes, a terrain-related SP is seen on the southwestern slope outside the caldera. The SP decreases from about 100 to 350 mv, corresponding to the ground surface elevation increase from 150 to 450 m (Fig. 2b). This terrainrelated SP is quite steady since similar correlation was repetitively Figure 2. Observed SP. (a) Map view. Purple dots: observation points; red star: location of the potential reference; red arrow: a region where the correlation between the SP and topographic altitude is shown in (b). (b) The correlation between SP and topographic altitude. Journal compilation C 2009RAS

4 SP controlled by resistivity structure in Izu-Oshima volcano 1167 observed on this slope in the annual surveys carried out from 1990 to 1994 (Ishido et al. 1997). The magnitude of the terrain-related SP is about 0.8 mv m 1, which can be produced by streaming potentials associated with the downward percolation of meteoric water from higher elevations (Ishido 1989, 2004). In the present survey, a similar SP elevation correlation is also seen along the eastern part of the survey lines (between 2000 and 3500 m of the E W distance in Fig. 2a). Inside the caldera, the SP does not decrease further with an elevation increase. In the southern part of the caldera, the SP is generally high. The highest potential of 69 mv is present near a fumarole at Kengamine, the eastern peak of Mt. Mihara (Fig. 1c). High potentials are also observed around the pit crater and distributed at the southwestern, southeastern and eastern caldera rims. Most of these high potentials around the central cone were observed by the previous survey in 1990 (Ishido et al. 1997) and do not seem to have been changed largely during the 16 yr since They are probably associated with the hydrothermal activity beneath the summit crater (Ishido et al. 1997) or the resulting high electrical conductivities (Ishido 2004) in the southern part of the caldera. A numerical simulation to reproduce these high potentials will be described in Section 3. In contrast to the southern part, no high potential is seen in the northern part of the caldera. Instead, a low-sp anomaly with a large magnitude (less than 600 mv) is observed on the caldera floor. The spatial extent is probably about 1.5 km and more than 2 km in the north south and east west directions, respectively. The western end of this anomaly was also observed in the 1990 survey (Ishido et al. 1997). The area of this low potential coincides with a region where the lava flows cover the surface. Because the region shows a prominent high gravity anomaly (Yokoyama & Tajima 1957; Ando et al. 1994), it is thought that dense lavas pile thicker than that in the outside area (Yokoyama 1969). The observed low SP is probably caused by the different lithology between the inside and outside of the region. The possible mechanism will be discussed in Section AMT measurements AMT measurements were conducted along two ENE WSW lines crossing the northern caldera and Mt. Mihara (see Fig. 1b) in 2006 and 2007 (Takakura et al. 2007). One of the purposes of this AMT survey is to provide an electrical conductivity structure required to perform quantitative modelling of the SP generation in the volcano. Fig. 3 shows the 2-D resistivity models along the two ENE WSW lines. Along both of the lines, the resistivity is divided into two layers of deeper low resistivity (less than several tens of ohm-metres) and shallower high resistivity (more than several hundreds of ohmmetres) (This is a common feature of the Izu-Oshima volcano observed in earlier observations, e.g. Ono et al. 1961; Ogawa and Takakura 1990; Utada & Shimomura 1990.). Along the northern caldera line (Fig. 3a), the conductive layer lies below the sea level outside the caldera, whereas the top of this layer is shallower at about m above sea level (ASL) inside the caldera. Along the line crossing Mt. Mihara (Fig. 3b), the top of the conductive layer rises to shallow levels like 500 m ASL beneath the central cone. 1-D vertical resistivity profiles obtained by the ELF VLF magnetotelluric observations (Utada & Shimomura 1990) are overlaid on the cross-sections (Fig. 3). The observations were conducted in , before the latest eruption in In spite of the differences in the observation times, the frequency bands and data analysis methods, the ELF VLF results show a similar feature with regard to the depth variations of the boundary between the resistive and conductive layers. Because these measurements were carried out over the entire island (Fig. 1b), the data are useful to extrapolate the present AMT results and construct a 3-D resistivity model, which will be explained in Section 3. The shallower resistive and deeper conductive layers are thought to correspond to the unsaturated and saline water-saturated layers, respectively. Below the high-elevation areas inside the caldera, the saline water probably intrudes the shallow levels near the water table due to hydrothermal activity and brings about high conductivity together with the formation of clay minerals. The shallow conductors beneath Mt. Mihara are located well above the water table (which is about m ASL, as shown in the next section). Probably, the high conductivity is caused by alteration minerals in part. Since these conductors were present before the last eruptions in , they are thought to be formed by long-term geothermal activities around the volcanic conduit. 2.3 Hydrogeological setting The elevation of the water table was measured in many wells in the island (Fig. 1b). In the wells located in the low-elevation areas, the water tables are almost at the sea level and the chemical compositions indicate the mixing of meteoric water and sea water, even though they are more than 1 km away from the coastline. These waters are thought to be the basal groundwater (Takahashi et al. 1987, 1988). In a deep well about 1 km from Mt. Mihara in the southwestern caldera (Fig. 1b), the water table is 36 m ASL (Nakada et al. 1999). The measured elevation of the water table coincides well with the estimated level from the AMT survey (Fig. 3b). The temperature logging of this well shows homogeneous temperature slightly less than 20 C in the upper 500-m interval above the water table (H. Watanabe, personal communication), which suggests a substantial downward flow of meteoric water in the unsaturated zone. The data also shows steep temperature increase of about 0.3 Cm 1 in the lower 500-m interval below the water table, which shows that the influence of hydrothermal activity beneath Mt. Mihara reaches the southwestern caldera rim. In the wells drilled near the northern caldera rim (Fig. 1b), the top of the saturated groundwater layer is as high as about 200 m ASL and varies from well to well. Chemical compositions are isolated from the basal groundwater. Takahashi et al. (1988) suggested that the features are similar to the dyke-impounded water, which was proposed for groundwater systems in Hawaiian volcanoes (e.g. Ingebritsen & Scholl 1993). Subparallel scattered dykes host locally isolated groundwater compartments and the groundwater level discontinuously changes across the dykes. In the Izu-Oshima volcano, flank volcanoes and eruptive fissures are aligned in the NNW SSE direction, and subsurface solidified dykes are imaged as an intense magnetization and high seismic velocity belt with a width of several kilometres (Makino et al. 1988; Onizawa et al. 2002; Ueda 2007). Since the wells near the northern caldera rim are located within the belt, the dyke-impounded water is probably responsible for the relatively high elevation of groundwater. The annual precipitation at Izu-Oshima is as much as about 3myr 1. The evapotranspirations in Japan are generally several tens of percentages of the precipitation, although no direct observations in Izu-Oshima are available. Since surface waters rarely appear except for shortly after squalls, it is thought that permeabilities are Journal compilation C 2009 RAS

5 1168 S. Onizawa et al. Figure 3. 2-D resistivity model revealed by AMT observation (Takakura et al. 2007). (a) Cross-section across the northern caldera. (b) Cross-section across Mt. Mihara. Purple inverse triangles with a numeral indicate observation sites shown in Fig. 1b. 1-D models at nearby sites by ELF VLF MT observations (Utada and Shimomura 1990) are overlaid. as high as that much of the precipitation infiltrates into the subsurface. According to an estimate by Ministry of Agriculture, Forestry and Fisheries of Japan (1986), 58 per cent of the precipitation infiltrates in Izu-Oshima. Since springs are found at higher elevations than the sea level (Fig. 1b), there must be local impermeable layers hosting local perched waters. However, the total of the discharge rate from these springs is thought to be two orders of magnitude smaller than the recharge rate. Therefore, most of the infiltrated water should reach the basal groundwater. Considering the low elevation of the water table and the substantial amount of downward flow of meteoric water, the global permeability of volcanic edifice is thought to be quite high in Izu- Oshima. Although the measurements of hydraulic conductivity are sparse to discuss the global permeability, very high permeabilities from 10 9 to m 2 are estimated in the foot areas on the basis of pumping tests and tidal response of the water table elevation (Ministry of Agriculture, Forestry and Fisheries of Japan 1980, 1986; Koizumi et al. 1998). These high permeabilities probably reflect loose scoria deposits, fractures in lavas and perhaps in tuff breccias. In the numerical simulation of groundwater flow described in the next section, the global permeability is regarded as an adjustable parameter to reproduce the observed water table elevation for a given meteoric water infiltration rate. 3 NUMERICAL SIMULATION OF ELECTROKINETIC POTENTIAL We apply the so-called EKP postprocessor (Ishido & Pritchett 1999) to a 3-D model to interpret the observed SP data from Izu- Oshima. After outlining the EKP postprocessor (Section 3.1), a numerical simulation of the groundwater flow based upon the hydrogeological data mentioned in the previous section is described in Section 3.2. Then, the results of the SP calculations from the underground conditions computed by the groundwater flow simulation are described in Sections EKP postprocessor The EKP postprocessor (Ishido & Pritchett 1999) calculates the space time distributions of the electrokinetic potentials resulting from the histories of underground conditions (pressure, Journal compilation C 2009RAS

6 SP controlled by resistivity structure in Izu-Oshima volcano 1169 temperature, salt concentration, liquid-phase saturation, etc.) computed by multiphase multicomponent unsteady thermohydraulic simulations. The basic equation solved by the postprocessor is I cond = I drag. (1) Eq. (1) represents sources of conduction current that are required for the appearance of electric potential at the surface (Sill 1983). I cond = ( L ee φ) is the conduction current density caused by the electric potential gradient ( φ) andi drag [ = L ev (p + ρ L gz)] is the drag current density caused by charges moved by the fluid flow through electrokinetic coupling (p is the pore-fluid pressure, ρ L is the density of liquid-phase water, g is the acceleration due to gravity and z is the elevation). The electrical conductivity L ee is given as the user-specified value or calculated by the following equation: L ee = (1 η)σ R + F 1 S n L (σ L + m 1 S ), (2) where η and F denote the porosity and formation factor, respectively; S L is the liquid-phase saturation; σ R and σ L are the electrical conductivities of the rock matrix and pore liquid, respectively; and m and S are the hydraulic radius of the pore and surface conductance, respectively. (Here m is defined as the ratio of the total pore volume to the total internal pore surface area, so e.g. m = a/2 if the pore has slit shape with an aperture a.) The coupling coefficient L ev is given as follows based upon the capillary model described by Ishido & Mizutani (1981) L ev = ηt 2 R ev (S L )εζ μ L = F 1 R ev (S L )εζ μ L, (3) where t is the tortuosity of the porous medium (ηt 2 equals the reciprocal of the formation factor F); R ev (S L ) is a correction factor ( 1) for liquid/gas two-phase flow and a function of the liquid-phase saturation; and ε, μ L and ζ are the (absolute) dielectric permittivity, dynamic viscosity and zeta potential of the liquid phase, respectively. If the upward conduction current balances the vertically downward drag current associated with the liquid water downflow in the unsaturated zone, the isoelectric potential surfaces are horizontal (Ishido 2004). In such situations, the vertical potential gradient is given as follows from I cond = I drag : φ z = F 1 R ev (S L )ζ ερ Lg. (4) L ee μ L If we approximate eq. (2) as L ee = F 1 S n L σ L, eq. (4) becomes φ z = R ev(s L )εζ SL nσ ρ L g = R ev(s L ) Lμ L SL n Cρ L g, (5) where C is the streaming potential coefficient ( L ev /L ee ) for S L = 1, and the term R ev /S n L represents the liquid-phase saturation dependency of the streaming potential coefficient. 3.2 Groundwater flow simulation To simulate a steady groundwater flow due to topographic relief on the Izu-Oshima island, we used the STAR generalpurpose geothermal reservoir simulator coupled with the BRN- GAS equation-of-state (Pritchett 1995, 2002), which is applicable to a system of three pore components (H 2 O, NaCl and air) and three pore phases (liquid, gas/vapour and solid halite precipitate). We used a 3-D Cartesian grid, which includes the entire island and surrounding sea area and extends and m in the east west and north south directions, respectively (Fig. 4). The area is subdivided into blocks ranging in size from 200 m 200 m in the central 6000 m 6000 m region to 500 m 500 m in the outer region. The grid extends vertically from 5000 to 725 m relative to sea level (RSL). Each of the uppermost 29 layers is 25 m thick, and the thickness ranges from 50 to 1000 m for the deepest 23 layers. The upper surface of the grid is uneven to represent the topographic relief (Fig. 4). The entire region is divided into two subregions, one of which represents sea and has very high porosity (0.99) and permeability (10 10 m 2 ). In the main subregion, the rock is homogeneous; the porosity is 0.35 and the permeability k is adjusted to reproduce the water table elevation (as described later). The relative permeabilities to the liquid and gas phases are assumed to depend linearly on the liquid-phase saturation with residual saturation of S RL = 0.3 and S RG = Capillary effects are neglected. All exterior boundaries except the bottom surface are open. The pressure and temperature are maintained at 1 bar and 20 C along the top boundary; any fluid that flows downward into the grid through the top surface is assumed to be air. To represent meteoric water recharge, a source of freshwater (H 2 O with mass fraction of NaCl) equivalent to I m of rainfall per year is uniformly applied to the uppermost grid blocks in the island region. Along the four vertical side surfaces, the pressure is maintained at the hydrostatic sea water pressure and the temperature is fixed at 20 C. Any fluid that flows into the grid through the side surface is assumedtobe seawater (H 2 O with mass fraction of NaCl). The initial conditions are at a uniform temperature of 20 C in the entire region, and the regions below and above the sea level are fully saturated with sea water and partially saturated (S L = 0.3) with freshwater, respectively. In the first series of simulations, we assume isotropic permeability for the entire region. Starting from the initial condition, groundwater flows from higher to lower elevations were computed. The system reached a quasi-steady state after several thousands of years of simulated time. Since the water table depth depends mainly on the ratio of the recharge rate of meteoric water (I) to the permeability (k) of the volcanic edifice, we performed simulations for a number of cases by changing I and k. As a result, the computed water table elevation matches fairly well with the observation mentioned in the previous section in the two cases when I/k = (m yr) 1 (one is a combination of I = 3myr 1 = annual precipitation and k = m 2 and the other is I = 0.3myr 1 and k = m 2 ). Since the subhorizontal alternation of relatively permeable and impermeable layers induces anisotropy, we assumed anisotropic permeability in the second series of simulations. The best result was obtained for cases with I/k H = (m yr) 1 and I/k V = (m yr) 1 (if I = 0.3 m yr 1, k H = and k V = m 2 ); the distributions of mass flux and those of the NaCl mass fraction in the fluid are shown for 10,000 yr in Fig. 5. The water table is located near m RSL and a thick unsaturated zone develops above that level in the volcanic edifice. A structure with a freshwater layer overlying the sea water layers is produced. A vertically downward flow takes place in the unsaturated zone and almost horizontally radial flows dominate the freshwater layer. As a result of a number of simulations, it is revealed that the computed water table elevation is not sensitive to the vertical permeability k V but to the horizontal permeability k H.Thisisbecause the major groundwater flow below the water table is restricted within the freshwater layer and the horizontal flow direction dominates there. As for the vertical permeability, the minimum value is obtained from the requirement that the downward liquid flow rate in the unsaturated zone is larger than the recharge rate when the Journal compilation C 2009 RAS

7 1170 S. Onizawa et al Distance [m] Altitude [m] 0 0 Distance [m] Figure 4. Grid block configuration for the 3-D groundwater flow simulation. Light green and light blue blocks indicate rock and sea regions, respectively. Solid and broken contours indicate the on-land and submarine topographies, respectively. The topographic contour interval is 100 m. liquid-phase saturation is close to unity: k V ρ L g/μ L > I (m s 1 ). This corresponds to I/k V < (m yr) 1 (here, I is in m yr 1 ). 3.3 SP calculation resistivity model and coupling coefficient Next, the EKP postprocessor was applied to the results of the STAR simulation as described in the previous section. First, it calculates the distributions of the pertinent parameters based upon eqs (2) and (3). As for the electrical conductivity, it is also possible to assign user-specified distribution instead of calculating L ee based upon eq. (2). We adopted this approach in this study and constructed a two-layered model (Fig. 6) based upon the results of the AMT survey (Section 2.2) and the observations by Utada and Shimomura (1990). In Fig. 6, the location of the boundary between the shallower resistive (1 k -m) and deeper conductive (10 -m) layers is defined by the isoresistivity surface of 50 -m. With regard to the coupling coefficient L ev, we adjust the zeta potential to reproduce the terrain-related SP observed on the southwestern slope outside the caldera: 0.8 mv m 1 (Fig. 2). This correlation is produced by the downward water flow in the unsaturated zone, since the water table is very deep in Izu-Oshima. By substituting φ/ z = 0.8 mv m 1, L ee = S m 1 and the values at 20 Cand1barforε, μ L and ρ L into eq. (4), we get F 1 R ev (S L )ζ = 0.11 mv. In this study, we assume R ev (S L )tobe the same as the liquid-phase relative permeability R L. According to the result of the groundwater flow simulation in the previous section, the mean S L in the unsaturated zone is and R L (S L = 0.323) = 0.032; therefore, R ev (S L = 0.323) is Further, if we take F = 10, ζ becomes about 35 mv. In the actual numerical simulation using the resistivity model with an uneven boundary surface, we set F 1 R ev (S L )ζ = 0.13 mv by adjusting the formation factor and zeta potential to be F 1 = ηt 2 = (F = 8.6) and 35 mv for freshwater at 20 C, respectively. The zeta potential is dependent on the salinity based upon Ishido and Mizutani s model (1981), so the magnitude is much smaller in the sea water layer. In the second step, the EKP postprocessor is used to calculate the electric potential distribution by solving eq. (1) within a finitedifference grid, which is the same as the grid used for the fluid flow simulation in this study. Boundary conditions on the potential are zero normal gradient on the ground surface (upper surface) and zero potential along the bottom and vertical sides of the grid. Journal compilation C 2009RAS

8 SP controlled by resistivity structure in Izu-Oshima volcano Distance [m] Altitude [m] Distance [m] Liquid Phase Saturation Figure 5. Result of the groundwater flow simulation. The colour scale indicates water saturation. Yellow arrows indicate groundwater flow directions. In the horizontal map, water saturation and horizontal component of flow vector for blocks at 0 25 m ASL are shown. Vertical to horizontal scale exaggeration is three times. Blue contours in the cross-sections show NaCl concentration. Note that the top of the saturated region is higher than sea level in the central island, indicating higher basal groundwater level compared to that of coastal areas. 3.4 Effect of conductive layer relief Before adopting the resistivity model shown in Fig. 6, we calculated the SP by assuming a simple flat two-layered resistivity structure (the boundary between layers of 10 3 and 10 -m is horizontal at the sea level). In this case, the equipotential surfaces are horizontal (Figs 7c and e) and the potential distribution on the ground surface is negatively correlated with elevation even in the summit area (Figs 7a, b and d), which effectively reproduces the terrain-related SP observed on the southwestern slope outside the caldera. However, the calculated result does not agree with the observed SP inside the caldera (Fig. 2). The distribution of the difference between the observed and calculated SP is shown in Fig. 9(a). The average and rms values of the differences (called residuals hereafter) over all of the measurement points are 68 and 204 mv, respectively. The result of the SP calculation taking into account the resistivity model (Fig. 6) is shown in Fig. 8. In this case, high potentials appear in the southern part of the caldera (Figs 8a, b and d). In particular, the locations of high potentials on Mt. Mihara and at the southeastern rim of the caldera are well reconstructed. As seen in the cross-sections (Figs 8c and e), these high surface potentials are caused by an upward shift in the higher-potential regions associated with the rise in the top of the conductive layer (see also Fig. 6). As seen in Fig. 9(b), the positive residuals in the southern caldera are largely suppressed as compared to the case with the flat two-layered structure (Fig. 9a). However, negative residuals in the northern caldera are slightly enhanced due to the shallower resistivity boundary. The average and rms values of the residuals are 87 and 174 mv, respectively. The observed SP distribution is effectively reconstructed by the simulation in which the electrical resistivity structure is considered. This indicates that the characteristic SP distribution observed at the Izu-Oshima volcano is mainly caused by meteoric water infiltration, and the electrical resistivity should be considered carefully in the analysis of the SP distribution. The high conductivity of the shallow water-saturated layer beneath Mt. Mihara is probably brought about by hot saline water drawn upward near the water table due to hydrothermal convection Journal compilation C 2009 RAS

9 1172 S. Onizawa et al Distance [m] Altitude [m] Top Altitude of Conductive Layer [m] Resistive Conductive Distance [m] Figure 6. Uneven two-layered resistivity model used for the electrokinetic postprocessor calculations. Map view shows the depth variation of the boundary between the surface resistive and the underlying conductive layers. Dark and light grey shades in the cross-sections indicate the resistive and conductive layers, respectively. Vertical to horizontal scale exaggeration is twice. around the volcanic conduit (Section 2.2). In the present fluid flow model, this hydrothermal convection is not reproduced; instead, freshwater flows almost horizontally outward in the shallow layer, as shown in Fig. 5. However, the potential difference caused by the drag current is very small in the region where the high conductivity is assigned, particularly for saline water flow due to smaller magnitude of the coupling coefficient. Therefore, the entire high conductivity region is nearly iso-potential, irrespective of the fluid flow directions. This is confirmed by Ishido (2004) based upon an axisymmetrical model. 3.5 Low SP in northern caldera The spatial extent of the low-sp anomaly observed in the northern caldera coincides well with the area of high-gravity anomaly, as mentioned in Section 2.1. Fig. 10(a) shows the high-pass-filtered Resistive 0 Conductive Bouguer anomaly map (Ando et al. 1994). Distinct high-gravity anomalies were observed as enclosed by the caldera rim and the ratios of dense lavas to pyroclastic rocks in these areas were thought to be higher as compared to those in the surrounding areas (e.g. Yokoyama 1969). If the lithology is different, the coupling coefficient is probably different from that in the outside areas. If the parameter F 1 R ev (S L )ζ has a magnitude greater than 0.13 mv (the value adopted to reproduce the terrain-related SP of 0.8 mv m 1, as mentioned in Section 3.3), the downward flow of meteoric water will produce a negative SP anomaly of larger magnitude in this area. In this series of SP calculations, a different value of the coupling coefficient was assumed for the northern caldera region shown by dark green in Fig. 10(b) (the vertical extension of this region is from the ground surface to the sea level). Only the parameter F 1 R ev (S L )ζ (actually F for simplicity of parameter setting) in this Journal compilation C 2009RAS

10 SP controlled by resistivity structure in Izu-Oshima volcano 1173 Figure 7. Calculated SP for the flat two-layered resistivity model. (a) Map view. (b) SP profile along the E W red line. (c) Cross-section of the subsurface resistivity structure and the electrical potentials along the E W red line. Contour interval for the electrical potential is 100 mv. (d) SP profile along the N S red line. (e) Cross-section of the subsurface resistivity structure and the electrical potentials along the N S red line. Vertical to horizontal scale exaggeration of (c) and (e) is twice. high-gravity region was adjusted (all the remaining parameters were unchanged) to minimize the rms value of the residuals between the measured and computed SP values. The best result was obtained when F 1 R ev (S L )ζ = 0.27 mv for the high-gravity region. The SP residuals were fairly suppressed over the entire area (Fig. 9c); the average and rms values of the residuals were reduced to 10 and 111 mv, respectively. As shown in Fig. 11, the calculated SP distribution becomes much closer to the observed one. 4 DISCUSSIONS 4.1 Coupling coefficient in the unsaturated zone Aubert and Atangana (1996) proposed a linear relation between SP and the thickness of the unsaturated zone in case that the unsaturated zone is homogeneous and has much higher resistivity than that of the underlying saturated zone. These prerequisites are satisfied on the southwestern slope outside the caldera in Izu-Oshima. As mentioned in Section 3.3, the terrain-related SP of 0.8 mv m 1 is thought to be produced by the downward water flow in the unsaturated zone. To reproduce this correlation in the numerical simulation, the parameter set F 1 R ev (S L )ζ in eq. (4) needs to be about 0.13 mv. To satisfy this, we set the parameters as F = 8.6 and ζ = 35 mv, since R ev (S L ) was assumed to equal to R L (S L ), which is a linear function of S L and corresponding to I/k V = (m yr) 1 assumed in the groundwater flow simulation. As mentioned in Section 3.2, however, I/k V is not much constrained, because the water table elevation is rather sensitive to I/k H. Further, it is uncertain whether or not the assumed parameters and relationships are valid for Izu-Oshima volcano. Here, we check the zeta potential by the streaming potential measurements for Izu-Oshima samples, and discuss how the parameter and model uncertainties affect our simulation results mentioned earlier Zeta potential We collected five samples around Mt. Mihara and carried out streaming potential measurements by using the EKA system (manufactured by Anton Paar GmbH, Graz, Austria). Each rock sample was grained into particles with sizes from 255 to 350 μm, cleaned Journal compilation C 2009 RAS

11 1174 S. Onizawa et al. Figure 8. Calculated SP for the uneven two-layered resistivity model. Notations are the same as those in Fig. 7. adequately with distilled water and stored in an aqueous solution of 10 3 mol L 1 KCl for more than 3 months before they were introduced into the sample holder. Fig. 12 shows the zeta potential of five samples as a function of ph for 10 3 mol L 1 KCl solution at room temperature (25 ± 3 C). The ph was varied between 2.5 and 9.5 with a step of about 0.5 by using HCl and NaOH. At each ph value, the ratio of the potential difference to the pressure difference between the two ends of the sample plug (i.e. the streaming potential coefficient) was repeatedly measured by alternating the fluid flow direction to obtain at least six stable values, which took about 1 h in a typical case. The zeta potential was calculated from the measured streaming potential coefficients without and with surface conductivity correction (see, e.g. Ishido & Mizutani 1981; Hase et al. 2003). The correction of surface conductivity is small in the present measurement and the standard deviation becomes several times larger near the lowest and highest ph conditions. Measured parameters and deduced zeta potential for each ph are summarized in Table 1. As seen in Fig. 12, the samples can be categorized into two groups: the lava samples have lower isoelectric point (IEP, see e.g. Ishido & Mizutani 1981) of 4 ± 0.6 and the scoria and ash samples have higher IEP of 6.8 ± 0.5. At neutral ph (7), the zeta potential ranges 16 to 12 mv and 5 to+1 mvforthelavaand scoria/ash samples, respectively. Considering that the IEP of silicate minerals ranges from about 2 to 5 (Ishido & Mizutani 1981), the observed IEP values for the scoria/ash samples are quite high. Such weak zeta potentials were measured for scoriae and breccias from Mount Pelée ( 4 to 19 mv at ph = ) by Jouniaux et al. (2000). Further, Hase et al. (2003) found samples of high IEP values above 6 for volcanic rocks from Aso volcano. Since the zeta potential of volcanic rocks is very sensitive to the weak mineralogical composition variations related to the eruption mechanisms (Jouniaux et al. 2000), we need further studies including more detailed mineralogical analysis. The ph and electrical conductivity of freshwater springs on the flank of Izu-Oshima volcano were reported as and S m 1, respectively (Takahashi et al. 1987). These conductivities are close to that of 10 3 mol L 1 KCl solution used for the present measurement at room temperature (about S m 1 ). Therefore, the in situ zeta potential is estimated to range from about 15 to +5 mv from the present results (Fig. 12), if we assume that the zeta potentials measured for the crushed granular samples represent the values for the in situ rocks. As far as the laboratory measurements were concerned, the zeta potential of 35 mv assumed for the freshwater condition in the numerical simulation was thought to be too large in magnitude. If we take into account the Journal compilation C 2009RAS

12 SP controlled by resistivity structure in Izu-Oshima volcano 1175 Figure 9. Residuals between the observed and calculated SP for (a) the flat two-layered model (Fig. 7), (b) the uneven two-layered resistivity model (Fig. 8) and (c) the uneven two-layered resistivity model with a region of different coupling coefficient in the northern caldera (Fig. 11). results of the zeta potential measurement and assume ζ = 7 mv instead of 35 mv, F 1 R ev (S L ) should be five times larger to obtain the same distribution of the calculated SP Hydraulic parameters The formation factors for volcanic rocks of Mount Pelée were determined by laboratory experiments (Jouniaux et al. 2000; Bernard et al. 2007). Bernard et al. (2007) showed that the formation factor ranges from 7 to The larger values are for lavas and indurated block-and-ash flow deposits, whereas the smaller values less than 20 correspond to pumices and scoriae. Comparing with these experimental results, 8.6 of the formation factor adopted for our simulation belongs to the smallest value group. Therefore, it is preferable to increase R ev (S L ) rather than to decrease F further, to make F 1 R ev (S L ) five times larger. If we assume F being constant and correct the difference between the assumed and measured zeta potentials solely by R ev (S L ), R ev (S L ) should be 0.16 (five times larger than used for the calculations in Section 3.3). According to this correction, hydraulic parameters such as the liquid-phase saturation and vertical permeability should also be changed. In our simulations, we adopted the simple relationship of R ev (S L ) = R L (S L ) and the linear relationship of the relative permeability to the liquid-phase saturation with the residual saturation of S RL = 0.3. In this case, the increase in R ev (S L ) is simply realized by reducing the vertical permeability k V by 80 per cent whereas the other parameters such as water recharge rate I are unchanged in the fluid flow simulation. The corresponding liquid-phase saturation becomes instead of Another model can be considered. If we assume R ev (S L ) = R L (S L )/S L following a theoretical model proposed by Linde et al. (2007), the type of relative permeability function can be involved in the above discussion. For the same relationship between relative permeability and liquid-phase saturation as used in the present fluid flow simulation, R ev (S L ) = R L (S L )/S L becomes 0.16 when S L = and R L = To obtain this relative permeability, the vertical permeability should be 1.7 times smaller. Further, when R L (S L ) is given by the function (van Genuchten 1980) R L = [ S e 1 (1 S 1/λ e ) λ] 2 S e = S L S RL /1 S RL and λ = 0.4, S RL = 0.2, R ev (S L ) = R L (S L )/S L becomes 0.16 when S L = and R L = Hence, the vertical permeability should be 3.1 times smaller. Therefore, for any model mentioned earlier, the liquid-phase saturation should increase and the vertical permeability should decrease, as associated with the correction of R ev (S L ). If the estimates of the in situ zeta potential, formation factor and so on are available in addition to the field observations of φ/ z and L ee, it is possible to infer R ev from eq. (4). Further, if we know the liquid-phase saturation dependency of R ev, we can estimate the liquid-phase saturation in the unsaturated zone and obtain an independent constraint to infer the liquid-phase relative permeability and vertical (intrinsic) permeability from geoelectrical observations. From this viewpoint, experimental and theoretical studies on R ev or the streaming potential coefficient for two-phase flow condition (e.g. Guichet et al. 2003; Linde et al. 2007) are very important. In contrast to ambiguities of the hydraulic parameters, the same SP distribution is calculated by just adjusting the parameters to satisfy F 1 R ev (S L )ζ = 0.13 mv regardless of the selected model. Therefore, discussions about the SP generation remain unchanged. 4.2 SP in Izu-Oshima volcano High SP in southern caldera Higher potentials are observed in the southern caldera where the thermal activities are found (Fig. 2). In most cases of such high SP close to active craters and thermal anomalies, the streaming Journal compilation C 2009 RAS

13 1176 S. Onizawa et al. Figure 10. (a) Gravity anomaly map. Components of longer wavelength than 4 km were cut from the Bouguer anomalies (Ando et al. 1994). Contour interval is 1 mgal. Blue dots: gravity observation locations. (b) Area partition for the simulations. The dark green area shows the region where the coupling coefficient is different from the surroundings. potential associated with thermally driven upward flow was believed to be the primary cause of the higher potentials. However, introducing the heterogeneous electrical resistivity model, the overall pattern of the observed SP was reproduced by the groundwater flow only due to meteoric water infiltration (Fig. 8). In particular, locations of the high potential at Mt. Mihara and the southeastern rim of the caldera are well reconstructed, even though the effects of the hydrothermal activity are not taken into account. The shallower conductors connecting with the lower conductive layer pull the equipotential surface up to a shallower region (see Figs 8 and 11), which is thought to be the primary cause of the high SP. This effect has already been proposed by numerical simulations for simple axisymmetrical models (Ishido 2004). We showed this phenomenon by the 3-D simulation based on the observed SP, resistivity and actual topography. Since the shallow conductors were detected before the latest eruptions in , they are thought to be long-lived structures. Probably, these are in part caused by alteration minerals formed by long-term geothermal activities. The contribution of the shallower conductors to the higher potentials can be estimated by using SP calculated for the flat two-layered model as a reference (Fig. 7). Difference between the observed SP and the reference is compared to that between the calculated SP for the uneven resistivity model and the reference. If we consider SP data higher than 200 mv, the positive potentials at Mt. Mihara and at southeastern caldera rim are reproduced 62 and 90 per cent, respectively on average, by introducing the shallower conductors. Even for the highest potential close to the fumarole at Kengamine, 47 per cent can be reproduced. However, the resolution of our resistivity model is insufficient to consider the densely observed SP with short wavelength and large amplitude variations. We need more detailed shape of the shallow conductive structure to discuss the short wavelength features. Prominent spatial correlation between high SP and subsurface shallow conductors at summit or active areas has been observed at other volcanoes such as Miyakejima (Sasai et al. 1997) and Piton de la Fournaise (Lénat et al. 2000). Probably the shallow conductors substantially affect the high SP also in these volcanoes. Therefore, we need to be careful to subtract quantitative information of thermally driven upflows from the high SP anomalies at active areas. Although there is a possibility to explain the short wavelength SP features by incorporating more detailed shallow conductive structure, close spatial correlations between the high SP and thermal activities are recognized at Mt. Mihara. Therefore, it is thought that hydrothermal activities also contribute to the high SP to certain extent at Mt. Mihara. The upflow of volcanic gases and vapour can diminish the downward flow of meteoric water (Ishido et al. 1997). According to the numerical modelling by Ishido (2004), this effect induces an increase of about 100 mv in the SP. So this can be a secondary cause of the high SP in the thermal and fumarolic areas. In addition to Mt. Mihara, a weak fumarole is frequently observed at the eastern caldera rim. At the southwestern caldera rim, a Journal compilation C 2009RAS