4.1 Interpretation of the resistivity model
The C-2 anomaly lies directly below the outcrop of the Shirikomadake serpentinite complex (Fig. 3). A similar conductive anomaly has been identified under a serpentinite body in the southern part of the Kamuikotan zone (Ogawa et al. 1994; Ichihara et al. 2019). In a general sense, these relationships can be considered to imply an association between serpentinite bodies and underlying conductive anomalies. However, the resistivity (3–100 Ω m) in the shallow part of our 3-D model of the outcropping serpentinite complex (Fig. 3b) is not particularly low. Near-surface resistivity surveys (0–200 m depth) conducted in the south of the Shirikomadake serpentinite complex (Fig. 1b) by Okazaki et al. (2011) detected a range of resistivities consistent with our results. They also showed that the resistivities of their massive, foliated, and clay-rich serpentinite core samples (massive, foliated, and clay-rich forms) were 250–1,500, 200, and 20–40 Ω m, respectively. Therefore, the resistivities of serpentine group minerals are not the sole control of conductive anomalies in serpentinite, such as our C-2 anomaly.
Candidates for the source of conductive anomaly C-2 include the presence of (1) inter-connected pore fluids, (2) interconnected conductive minerals (e.g., magnetite) that are commonly found in serpentinite, and (3) a high-temperature zone or melt that decreases resistivity. The first of these is the most likely candidate as previous studies of conductive anomalies in subduction zones (e.g., Aizawa et al. 2021; Ichihara et al. 2016; Evans et al. 2014) and an experimental study for conductive antigorite (Reynard et al. 2011) suggested. The formation of serpentinite in mantle wedges requires an abundant supply of fluids from the subducting plate or possibly from repeated serpentinite dehydration (e.g., Hyndman and Peacock 2003). Large volumes of interconnected fluids are common in mélange zones because of their numerous fractures and high permeability. In addition, the resistivity of dehydrated saline fluid is low at depth (Sakuma and Ichiki 2016; Sinmyo and Keppler 2017). Therefore, serpentine mélange in subduction zones is suitable to the presence of interconnected conductive aqueous pore fluids resulting in low-resistivity anomaly.
The second candidate (interconnected conductive minerals) has been considered for mid-crustal conductive areas (e.g., Myer et al. 2013) because magnetite contained in serpentinites is highly conductive and can produce a conductive anomaly (Stesky and Brace 1973). However, it is difficult to attribute the C-2 anomaly to magnetite alone because both the conductivity of borehole samples collected by Okazaki et al. (2011) and the surface-measured resistivity of the Shirikomadake serpentinites do not indicate a sufficient anomaly, although the presence of magnetite may contribute to the low resistivity of interconnected pore fluids. The third candidate (high-temperature zone or melt) is unlikely to be a direct cause of the C-2 anomaly because the study area is a considerable distance from the nearest area of volcanic activity and no high-temperature phenomena, such as hot springs, are known in the study area.
There are two possible explanations for the relatively high resistivity zone between the serpentinite outcrop and the C-2 anomaly (Figs. 3 and 4). It might represent resistive rocks that contain little or no pore fluid, such as, massive serpentinite, basaltic lavas (Katoh et al. 1979), or metamorphic rocks of the Kamuikotan zone that do not outcrop in the study area. Alternatively, because the resistivities of aqueous fluids increase with decreasing temperature (i.e., with decreasing depth) (Sakuma and Ichiki 2016; Sinmyo and Keppler 2017), the resistivity remains high in the shallow sequence.
Borehole SK-1 near the western margin of the study area (Fig. 1b) initially penetrated Paleogene–Neogene sedimentary rocks at 3,030 m depth and, was within late Cretaceous sedimentary rocks at its total depth (4,505 m) (Ogura and Kamon 1992). Borehole logs indicate that resistivities in the Paleogene–Neogene and late Cretaceous sequences were 1–4 and 4–10 Ω m, respectively (Kanekiyo 1999). The modeled resistivity within anomaly C-1 (Fig. 5) is consistent with that logging data. It is noteworthy that the surface extent of the C-1 anomaly is consistent with the mapped extent of the Cretaceous (Yezo Group) and Paleogene–Neogene sedimentary rocks (Figs. 1b and 3a). Hence, we interpret the C-1 anomaly to represent Cretaceous and Paleogene–Neogene conductive sedimentary rocks. Similar associations between resistivity anomalies and outcrops have been recognized in other parts of middle–western Hokkaido (Ichihara et al. 2008, 2016; Yamaya et al. 2017; Ichihara et al. 2019).
The C-1 anomaly deepens toward the western boundary of the study area (Fig. 5) and the steepest dip of its lower boundary coincides with thrust faults (Fig. 5) that uplifted the eastern part of the study area. The C-1 anomaly is partly overlain by a layer of moderate resistivity (about 100 Ω m) that is similar to a near-surface layer of moderate resistivity identified by Ueda et al. (2014) near sites B13 and B14. On the basis of seismic reflection survey data, they interpreted this layer to be the Quaternary Sarabetsu Formation, which contains relatively fresh groundwater. Thus, we interpret the layer of moderate resistivity that we modeled overlying anomaly C-1 to represent Quaternary sediments.
4.2 Uplift of serpentinite and its implications for intra-plate SSEs
Previous geological studies have identified faults at the boundaries between the Shirikomadake serpentinite complex and surrounding sedimentary rocks (Igi 1959; Katoh et al. 1979; Oka 1985). The early Pleistocene Sarabetsu Formation near the center of the study area (Oka and Igarashi 1993; Niwa et al. 2020) contacts with the serpentinite complex by fault and is vertically inclined near the contact, indicating that uplift of the complex occurred during the Pleistocene or later. In addition, tectonic fault blocks within the serpentinite complex contain Cretaceous sedimentary rocks (Katoh et al. 1979), suggesting that Cretaceous rocks from the surrounding area were dragged into the serpentinite complex during the Quaternary (or earlier) uplift. The shape of the C-2 anomaly indicates that the faults bounding the serpentinite complex extend to deep area (Fig. 5). Previous researchers (e.g., Guillot et al. 2015) have proposed buoyancy-driven diapiric ascent of serpentinites from depth. Therefore, serpentinite complexes in the Kamuikotan zone may have been uplifted in the surrounding rocks and thus contribute to the transport of aqueous fluids derived from the subducting plate.
Recent seismic studies have proposed that intra-plate SSEs occur in subduction zones when the presence of high-pressure pore fluids weakens the plate interface (e.g., Shelly et al. 2006; Kato et al. 2010). In our study area, an intra-plate SSE occurred near the western edge of the Shirikomadake serpentinite complex during 2012–2013 (Ohzono et al. 2015) (Fig. 1). Because we interpreted the C-2 anomaly to represent upwelling aqueous pore fluid, such fluid may have been available to the area where that SSE occurred (Fig. 5). Ohzono et al. (2015) indicated that the SSE might have occurred in a detachment fault at the base of the Cretaceous–Neogene sedimentary rocks. Moreover, if the sedimentary rocks contained impermeable clay minerals and thus acted as cap rocks, they may have trapped (Fig. 5) and elevated the pressure of the pore fluids within the detachment fault, thus causing an intra-plate SSE. Therefore, the plate-interface SSE reported by Ohzono et al. (2015) might well have been caused by high-pressure fluids associated with the Shirikomadake serpentinite complex. However, our study area does not fully cover the fault zone and the earthquake swarm that accompanied the SSE occurred mainly beyond the southern margin of the MT array (Ichiyanagi et al. 2015). Thus, additional MT soundings south of our study area are required to better understand the fault rupture processes in the study area.