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Potential geologic osmotic pressure in the Wakkanai Formation: Preliminary estimation based on the dynamic equilibrium between chemical osmosis and advection
Journal of Hydrology 579 (2019) 124166
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Potential geologic osmotic pressure in the Wakkanai Formation: Preliminary estimation based on the dynamic equilibrium between chemical osmosis and advection
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Mikio Takeda , Mitsuo Manaka, Kazumasa Ito National Institute of Advanced Industrial Science and Technology (AIST), Central 7, 1-1-1 Higashi, Tsukuba, Ibaraki 305-8560, Japan
A R T I C LE I N FO
A B S T R A C T
This manuscript was handled by Corrado Corradini, Editor-in-Chief, with the assistance of Patrick Lachassagne, Associate Editor
Geologic osmotic pressure has been proposed as a cause of excess fluid pressures in argillaceous formations in which dynamic geologic processes have little effect on pressure. However, the identification of geologic osmotic pressure has rarely been attempted, thus its potential effect on groundwater systems is a longstanding question in hydrogeology. Herein we present a quantitative analysis of excess fluid pressures of unknown origin in the Wakkanai Formation, assuming osmotically generated overpressure. Potential geologic osmotic pressures are estimated from profiles of pore water salinity and rock physical properties using a one-dimensional model that describes geologic osmotic pressures in a state of dynamic equilibrium between chemical osmosis and advection. The model results indicate that the Wakkanai Formation has the requisite conditions of both salinity gradient and semipermeability to generate geologic osmotic pressures comparable to observed excess pressures, up to 230 kPa. The applicability of the dynamic equilibrium model is also confirmed by using a steady-state model with the same pore water salinity and rock physical properties data. These results suggest that chemical osmosis is a potential cause for the excess fluid pressures observed in the Wakkanai Formation.
Keywords: Chemical osmosis Osmotic pressure Semipermeability Argillite Excess fluid pressure
1. Introduction Chemical osmosis that arises from the semipermeability of argillite and salinity differences in groundwater generates geologic osmotic pressure within argillaceous formations, which is observed as overpressure or underpressure relative to hydrostatic or hydrodynamically equilibrated fluid pressure (Greenberg et al., 1973; Neuzil, 2000; Wilson et al., 2003; Gonçalvès et al., 2004; Neuzil and Provost, 2009). Chemical osmosis occurs from low- to high-salinity regions of an argillaceous formation, inducing fluid pressure buildup in the high-salinity region and drawdown in the low-salinity region. The pressure gradient that develops within the formation causes advection that counters the chemical osmotic flow and causes the osmotically induced pressure to dissipate. However, because of the low permeability of argillites, the osmotically induced pressure is maintained in a quasiequilibrium state as the counteracting flows ultimately reach a state of dynamic equilibrium (Neuzil, 2000; Gonçalvès et al., 2004; Bader and Kooi, 2005; Takeda et al., 2014). Since the salinity profile controlling osmotic flow is generally in a transient state and is mainly in the diffusion-dominated regime of argillaceous formations (e.g., Mazurek et al., 2011), the osmotically induced pressure (i.e., the geologic
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osmotic pressure) lasts over geologically long timescales (Neuzil, 2000; Gonçalvès et al., 2004). Geologic osmotic pressure is relatively small in comparison to excess fluid pressures caused by dynamic geologic processes such as tectonic loading or compaction disequilibrium (e.g., Osborne and Swarbrick, 1997; Gonçalvès et al., 2010); however, ignoring or underestimating the role of geologic osmotic pressure may result in misinterpretation of groundwater flow directions and associated solute transport (e.g., Marine and Fritz, 1981). Therefore, when excess fluid pressures are encountered in hydrogeologic investigations, the possibility of geologic osmotic pressure needs to be identified, particularly in the case of nuclear waste isolation in argillaceous formations (Neuzil, 2000, Wilson et al., 2003; Neuzil and Provost, 2009). Geologic osmotic pressure has been proposed as a cause of excess fluid pressures in argillaceous formations in which dynamic geologic processes are thought to be too minor to affect pressures (e.g., Young and Low, 1965; Marine and Fritz, 1981). However, the idea of geologic osmotic pressure has long been controversial because of the lack of data directly supporting its existence in situ. Since evidence for the semipermeability of argillites to produce geologic osmotic pressure was first obtained from in situ borehole experiments on the Pierre Shale, performed by Neuzil (2000), semipermeability has been investigated
Corresponding author. E-mail address: [email protected] (M. Takeda).
https://doi.org/10.1016/j.jhydrol.2019.124166 Received 24 June 2019; Received in revised form 28 August 2019; Accepted 19 September 2019 Available online 20 September 2019 0022-1694/ © 2019 Elsevier B.V. All rights reserved.
Journal of Hydrology 579 (2019) 124166
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denudated (Iwatsuki et al. 2009). Although many borehole investigations have been conducted in this area (Matsui et al., 2007), detailed semipermeability measurements have been only performed on Wakkanai mudstones retrieved from the SAB-2 borehole (Fig. 1a) (Takeda and Manaka, 2018). Accordingly, the data obtained from the SAB-2 borehole were used in the present study. Fig. 1b shows the lithology of the SAB-2 borehole. The Wakkanai Formation mainly consists of siliceous mudstone, and its shallow section (5–75 m depth) is a transition zone between siliceous and diatomaceous mudstones. Fig. 2 shows the data obtained from the SAB-2 borehole (Amano et al., 2012; Suko et al., 2014). Fig. 2a shows that the dominant ions in pore water are sodium and chloride. Their concentrations tend to increase with depth, suggesting that the salinity gradient within the Wakkanai Formation may induce downward chemical osmosis to generate geologic osmotic pressures at depth. Fig. 2b shows the temperature measured during drilling of the SAB-2 borehole (Suko et al., 2014). The temperature gap at 450 m depth is due to separate drilling stages for the shallow and deep sections, which occurred in different seasons. Fig. 2c shows that the observed fluid pressures are close to hydrostatic fluid pressures determined from the varying solute concentration and temperature. However, the observed fluid pressures tend to deviate from hydrostatic with increasing depth, with a maximum deviation of 230 kPa at a depth of 660 m.
extensively in other argillites, including the Opalinus Clay (Noy et al., 2004; Horseman et al., 2007), Boom Clay (Garavito et al., 2007), Callovo-Oxfordian argillite (Gonçalvès et al., 2004; Gueutin et al., 2007, Cruchaudet et al., 2008; Rousseau-Gueutin et al., 2008, 2009, 2010), and Wakkanai mudstone (Takeda et al., 2014; Takeda and Manaka, 2018; Manaka and Takeda, 2018). So far, extensive studies of the Callovo-Oxfordian formation have revealed that geologic osmotic pressure can explain 0.1–0.15 MPa of the 0.5–0.6 MPa overpressures observed in the formation (Gonçalvès et al., 2004; Rousseau-Gueutin et al., 2009). A one-dimensional, steady-state analysis performed by Tremosa et al. (2012) also suggests that the excess fluid pressures within the argillaceous formation at Tournemire, France may be partly due to geologic osmotic pressures. Neuzil and Provost (2009) examined data in the existing literature on the semipermeability of argillites and suggested that all analyzed media exhibited at least a modest ability to generate osmotic pressure. However, geologic osmotic pressures quantitatively identified based on field and laboratory data are still limited, thus the potential effects of geologic osmotic pressures on groundwater systems are still largely unknown. Here we report on a quantitative analysis of excess fluid pressures of unknown origin, observed in the Wakkanai Formation in the Horonobe region of Japan, as osmotically generated. In the Wakkanai Formation, excess fluid pressures up to 400 kPa and salinities up to half that of seawater have been observed at depth (Kurikami, 2007; Iwatsuki et al., 2009). The excess fluid pressures at depth in this region can be explained by topography-driven flow, but needs the assumption of a caprock in hydrogeological modeling that has not been observed by borehole investigations nor by permeability experiments (Karasaki et al., 2011). In addition, although the Horonobe region is located near the North American-Eurasian plate boundary, the strain rate in this region is too low to increase fluid pressures as a result of relatively high permeability (Tokiwa and Sugimoto, 2012; Neuzil, 2015). Furthermore, a recent study on the semipermeability of the Wakkanai mudstone suggests that the mudstone may act as a weak geologic membrane (Takeda and Manaka, 2018). In this study, we use a one-dimensional model that describes geologic osmotic pressures in a state of dynamic equilibrium between chemical osmosis and advection, which was originally developed by Neuzil (2000) for theoretically screening favorable conditions for generating geologic osmotic pressures. In order to estimate geologic osmotic pressures from salinity distribution and rock physical properties, both of which are obtained from field and laboratory investigations, we use a first-order approximation of the model, the effectiveness of which was demonstrated by Yu (2017) in analyses of geologic osmotic pressures in the Opalinus Clay of Mont Terri in northwestern Switzerland. We also compare our results to the geologic osmotic pressure estimated from experimentally determined semipermeabilities of the Wakkanai mudstone. The applicability of the dynamic equilibrium model is also tested against a steady-state model using additional hypothetical data and boundary conditions. Finally, we discuss the potential causes affecting fluid pressure distribution in the Horonobe region to assess the relevance of geologic osmotic pressure to the overpressures observed in this region.
2.2. Analysis based on the state of dynamic equilibrium between chemical osmosis and advection The present analysis assumes a horizontally uniform argillaceous formation in which lateral hydrodynamic fluxes are negligibly small. Based on this assumption, one-dimensional, vertical pore water flux through the argillaceous formation is described by a generalized Darcy’s law equation (e.g., Neuzil and Provost, 2009; Tremosa et al., 2012):
∂π k ∂ q = − ⎧ (p − ρgz ) − σ ⎫ ∂z ⎬ μ⎨ ⎭ ⎩ ∂z
(1)
where q is the pore water flux (m/s), k is the intrinsic permeability (m2), μ is the dynamic viscosity (Pa·s), p is the fluid pressure (Pa), ρ is the pore water density (kg/m3), g is gravitational acceleration (m/s2), z is the depth below the ground level (m), σ is the osmotic efficiency (-), and π is the theoretical osmotic pressure (Pa). σ is a dimensionless parameter that ranges between 0 and 1 by definition and represents semipermeability that is controlled by the pore size between clay minerals, the electrochemical properties of clay minerals, and the solute concentration of the pore water (e.g., Bresler, 1973). The osmotic pressure π has the same dimension as pressure, but is a measure of the decrease in the chemical potential of water (Neuzil and Provost, 2009). The first and second terms on the right side of Eq. (1) represent advection driven by pressure and gravity potentials and chemical osmosis driven by chemical potential, respectively. When a state of dynamic equilibrium is established between chemical osmosis and advection, the net water flux through the argillaceous formation becomes almost zero, and geologic osmotic pressure emerges as the excess fluid pressure relative to the hydrostatic or hydrodynamic fluid pressures. This situation was originally considered by Neuzil (2000) for theoretically screening favorable conditions for generating geologic osmotic pressures, and it gives Eq. (1) the constraint of q = 0. Further introducing p = phy + pos, where phy is the hydrostatic or hydrodynamic pressure (Pa) and pos is the geologic osmotic pressure (Pa), Eq. (1) can be rewritten as
2. Estimation of geologic osmotic pressure in the Wakkanai formation 2.1. Geological setting and data suggesting geologic osmotic pressures The Wakkanai Formation is located in the Horonobe area of Hokkaido prefecture, Japan (Fig. 1a). According to Iwatsuki et al. (2009), the Wakkanai Formation and overlying formations were originally formed in a marine environment during the Neogene to Quaternary periods, then were compacted during subsidence prior to 1 Ma. The maximum burial depth of the top of the Wakkanai Formation is estimated to be 1000 m (Ishii et al., 2011a). The strata in this area have been uplifted due to folding and thrusting, and the top layers have been
∂p ∂π k ∂ 0 = − ⎧ (phy − ρgz ) + ⎛ os − σ ⎞ ⎫ μ⎨ ∂z ⎠ ⎬ ⎝ ∂z ⎭ ⎩ ∂z ⎜
⎟
(2)
Assuming that hydraulic flux, which is represented by the first term on the right side of Eq. (2), is negligible, owing to hydrostatic or 2
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Fig. 1. Location and stratigraphic section of the SAB-2 borehole (based on Ishii et al., 2011b; Suko et al., 2014). (a) Location map, geologic map, and geologic cross section of the Horonobe area in Hokkaido, showing the location of the SAB-2 borehole. The bold line in the geologic cross section indicates the approximate projection of the borehole along the faults. (b) Stratigraphic section of the SAB-2 borehole.
Fig. 2. Data collected from the SAB-2 borehole. Vertical distributions of (a) the concentrations of major groundwater solute components (Amano et al., 2012), (b) temperature, and (c) pore fluid pressure measured during the course of drilling (Suko et al., 2014). Hydrostatic fluid pressures in (c) are determined from the total amount of solutes dissolved in groundwater collected from the SAB-2 borehole and its varying temperature.
osmosis, represented by the right side of Eq. (3), is equal to the counteracting advection driven by the geologic osmotic pressure gradient when the system is at dynamic equilibrium. Upon integration, Eq. (3) can be solved over the vertical domain when σ and π are given as smooth functions of z and/or solute
hydrodynamic equilibrium, Eq. (2) can be rewritten as (e.g., Neuzil and Provost, 2009)
∂pos ∂π =σ ∂z ∂z
(3)
Although k and μ are eliminated, Eq. (3) states that chemical 3
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concentration (Neuzil, 2000; Neuzil and Provost, 2009). However, the present analysis addresses field and laboratory data that were discretely obtained from a borehole (Fig. 2), so pos is calculated using the discretized form of Eq. (3)
Δpos (z i ) = σi Δπ (z i )
(4)
where i is the index for the discretized formation element (i = 1 for the top element); zi is the vertical coordinate of the center of the ith element; Δpos and Δ π are the changes in the geologic and theoretical osmotic pressure, respectively, along the ith element; and σ i is the osmotic efficiency of the ith element. Summing Eq. (4) from the top to the nth element, the geologic osmotic pressure at the nth element can be estimated as n
pos (z n ) =
∑ Δpos (z i) i=1
(5)
The above approach for assigning osmotic efficiencies and osmotic pressure is essentially the same as the first-order approximation of the generalized Darcy’s law adopted by Yu (2017). 2.3. Osmotic efficiency calculation Fig. 3. Osmotic efficiency σ as a function of the nominal half distance between clay surfaces b and pore fluid concentration C for monovalent solutes. The curve is the Kemper–Bresler model (Bresler, 1973). The plots represent the results of fitting the Kemper–Bresler model to osmotic efficiencies of 0.021 and 0.081, based on samples from the Wakkanai mudstone at 401 and 698 m in the SAB-2 borehole, respectively (Takeda and Manaka, 2018). The gray lines indicate salinity ranges of 0.6–0.1 M NaCl for each measurement.
The osmotic efficiency, σ in Eq. (4), is a measure of the anion exclusion effect that is exerted by the overlap of electrical double layers (EDLs) at pore spaces between negatively charged clay surfaces (e.g., Marine and Fritz, 1981; Appelo and Postma, 2005). Accordingly, σ varies by location in argillaceous formations in accordance with the factors affecting the EDLs overlap, such as pore size between clay minerals, electrochemical properties of the clay minerals, and solute concentrations of the pore water. Among a number of models developed for the anion exclusion effect (e.g., Kemper and Rollins, 1966; Bresler, 1973; Marine and Fritz, 1981; Fritz, 1986; Revil and Leroy, 2004; Gonçalvès et al., 2007; Tremosa et al., 2012; Revil, 2017), the Kemper–Bresler model for monovalent solutes (Bresler, 1973) was used in our analysis because the model has an advantage of describing osmotic efficiency as a single curve relating σ to b C , where b is the nominal half distance between clay surfaces in Å (Fig. 3). In addition, the Kemper–Bresler model has previously been successfully applied to describe the osmotic efficiencies of undisturbed argillaceous formation media (Cey et al., 2001; Neuzil and Provost, 2009; Oduor et al., 2009; Rousseau-Gueutin et al., 2008; Takeda et al., 2014). Although the groundwater recovered from the SAB-2 borehole contains divalent cations, which compress the EDLs more than monovalent cations (e.g., Bresler, 1970; Bresler, 1973; Tremosa et al., 2012), the concentration of divalent cations in the groundwater is considerably lower and accounts for only 0.3–2.7% of the total cations and 1.3% on average (Fig. 2a and Amano et al., 2012). These values are close to that of the multi-component solution adopted in the experiments performed by Cey et al. (2001), in which the osmotic efficiencies were well characterized by the Kemper–Bresler model for a NaCl solution. Parameter b in the Kemper–Bresler model has been related to rock physical properties by Neuzil (2000) and is expressed as
b=
n ρb As
The b values were also estimated from the experimental values of σ: 0.021 and 0.081, which were measured in the cores retrieved from depths of 401 and 698 m, respectively, in the SAB-2 borehole, at a salinity difference of 0.6–0.1 M NaCl (Takeda and Manaka, 2018). These experimental σ values are those measured at effective confining stresses for individual cores after applying their past maximum effective burial stresses. This is because the Wakkanai Formation is currently overconsolidated as a result of the denudation of overlying strata. For these experimental σ values, the b values were adjusted to fit the Kemper–Bresler model with the measured σ values (Fig. 3). The estimated b values are 31 and 21 Å for the cores retrieved from depths of 401 and 698 m, respectively (Fig. 4d). These values are relatively close to those calculated from physical properties data. However, the measured σ values were obtained at a relatively large salinity difference of 0.6–0.1 M NaCl, so the representativeness of the estimated b values is ambiguous (Fig. 3). Also, the number of experimentally derived b values were not sufficient to delineate a vertical profile for b. Accordingly, these experimentally derived b values are used as a reference in the analysis, using a linear function with respect to z for calculating the vertical profiles of osmotic efficiency (Fig. 4d).
2.4. Theoretical osmotic pressure approximation The theoretical osmotic pressure, π in Eq. (4), is the decrease in the chemical potential of water due to the presence of solutes, thus it is affected by the concentration of each solute present in the groundwater. Although the groundwater of the Wakkanai Formation contains various solutes, its dominant anionic and cationic species are chloride and sodium, respectively (Fig. 2a). We therefore assumed that the groundwater in the Wakkanai Formation was a NaCl solution. Fig. 5 shows the equivalent NaCl concentrations estimated from the total amounts of solutes dissolved in groundwater collected from the SAB-2 borehole (Amano et al., 2012). The equivalent NaCl concentrations ranged from 0.01 to 0.16 M. For solutions with concentrations lower than 1 M, the theoretical osmotic pressure can be approximated according to the van’t Hoff relation (Fritz, 1986): π = νRTC, where ν is the number of
(6)
where n is the porosity, ρb is the dry bulk density, and As is the specific surface area. In order to obtain the vertical profile of b in the SAB-2 borehole, we measured n and ρb using mercury intrusion porosimetry, and As using the Brunauer-Emmett-Teller method (Brunauer et al., 1938) on the cores retrieved from the borehole. Fig. 4 shows the measured n, ρb, and As values and the estimated b values. The profile of b exhibits fluctuations at depths of 649 and 698 m. However, b values for the other depths are relatively constant and within a range of 29–31 Å. The profile of σ, which is necessary for calculating Eq. (4), was evaluated using the Kemper–Bresler model and the profiles of b and the solute concentrations. 4
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Fig. 4. Physical properties measured from samples of the Wakkanai mudstone retrieved from the SAB-2 borehole. (a) Porosity and (b) medium dry bulk density were measured using mercury intrusion porosimetry. (c) Specific surface area was measured using the Brunauer-Emmett-Teller method (Brunauer et al., 1938). (d) The nominal half distance b between clay surfaces was determined from porosity, medium dry bulk density, and specific surface area using Eq. (6). The experimental values are those determined from the osmotic efficiencies measured from samples of the Wakkanai mudstone retrieved from 401 and 698 m depths in the SAB-2 borehole (Takeda and Manaka, 2018). The straight line on the experimental values represents the linear approximation of b with respect to z.
(Fig. 3). Using Eqs. (5) and (7) and the vertical profiles of T, C, and σ (Figs. 2b, 5 and 6a), the potential geologic osmotic pressures in the formation were estimated. Fig. 6b shows the vertical profiles of the estimated geologic osmotic pressures and the fluid overpressures observed at the SAB-2 borehole. The observed fluid overpressures were determined by subtracting the hydrostatic fluid pressures from the observed fluid pressures (Fig. 2c). The estimated geologic osmotic pressure profiles replicate the downward increase in fluid overpressure observed in the SAB-2 borehole. However, they fail to delineate the fluid overpressures observed at depths shallower than 370 m. In contrast, the potential geologic osmotic pressures estimated for depths below 490 m are relatively close to observed fluid overpressures. In particular, potential geologic osmotic pressures estimated from the linear approximation of experimental b values are almost equivalent to the observed fluid overpressures (Fig. 6b) despite the sparse semipermeability measurements (Fig. 4d) and the relatively high salinity conditions (Takeda and Manaka, 2018). On the other hand, the analysis performed using the b values determined from the physical properties of retrieved cores slightly overestimates the fluid overpressures (Fig. 6b). 4. Discussion
Fig. 5. Vertical profile of equivalent NaCl concentration estimated from the total amount of solutes dissolved in pore fluids in the SAB-2 borehole (Amano et al., 2012).
This study assumed chemical osmosis as the process that deviates pore fluid pressures from a hydrostatic distribution and exemplified the potential of the Wakkanai Formation to generate geologic osmotic pressures comparable to the observed fluid overpressures of 230 kPa. However, our analyses do not replicate observed fluid overpressures particularly well at depths shallower than 370 m in the SAB-2 borehole. This suggests that other factors still need to be addressed in order to determine the relevance of geologic osmotic pressure in the fluid overpressures observed in this region.
ionic species (e.g., ν = 2 for NaCl), R is the gas constant (8.314 m3 Pa/ K/mol), and T is the absolute temperature (K). Introducing this relationship, Eq. (4) can be rewritten as
Δp (z i ) = σi νRT ΔC (z i )
(7)
where ΔC is the change in NaCl concentration along the i element. The temperature within the Wakkanai Formation increases with depth (Fig. 2b), so we assumed that the temperature was a linear function of z (median line in Fig. 2b). th
4.1. Applicability of dynamic equilibrium model
3. Results
The dynamic equilibrium model adopted in this study enables the calculation of geologic osmotic pressure from limited salinity and basic rock physical properties data, independently from permeability variations. This is an advantage of the dynamic equilibrium model compared
Fig. 6a shows the vertical distribution of the osmotic efficiency calculated from b and C (Fig. 4d and 5) using the Kemper–Bresler model 5
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Fig. 6. Vertical profiles of (a) osmotic efficiency calculated using the Kemper–Bresler model (Bresler, 1973), the NaCl concentration profile (Fig. 5), and the profile of b values (Fig. 4d); and (b) observed fluid overpressure and potential geologic osmotic pressure estimated using Eqs. (5) and (7) and NaCl concentration and osmotic efficiency profiles (Figs. 5 and 6a) in the SAB-2 borehole.
permeabilities reported for the Wakkanai mudstone by Kurikami et al. (2008), which range from 10−20 to10−18 m2 and tend to decrease with increasing depth. Fig. 7b shows the profiles of geologic osmotic pressure estimated from the steady-state analyses with a constant hydrostatic pressure at the bottom boundary. The geologic osmotic pressures estimated from the steady-state analyses are slightly smaller than those estimated from the dynamic equilibrium model. This indicates that the magnitude of geologic osmotic pressure is dependent on osmosis-driven downward advection, which is controlled by permeability variations. However, the geologic osmotic pressures estimated from the steady-state analyses are still large and agree well with the observed overpressures at depths below 490 m, suggesting that the intrinsic permeabilities assumed for the Wakkanai Formation (Fig. 7a) are sufficiently small to maintain geologic osmotic pressures (e.g., Tremosa et al., 2012). Fig. 7c shows the profiles of geologic osmotic pressure estimated from the steadystate analyses with a no-flux boundary at the bottom, for reference. As expected, the estimated geologic osmotic pressures are almost the same as those from the dynamic equilibrium model. This explicitly indicates that the analysis based on the dynamic equilibrium model is equivalent to the steady-state analysis with the profile of k and a no-flux bottom boundary, and can be used to estimate the maximum potential geologic osmotic pressure for given salinity and rock physical property profiles. This is the advantage of using the dynamic equilibrium model when adequate data for performing steady-state or transient state modeling are not available, especially during the early stages of hydrogeologic investigations.
to the steady-state and transient models, both of which require additional data regarding transport parameters, such as permeability and diffusivity, and initial and boundary conditions at formation boundaries. However, Eq. (3), on which the dynamic equilibrium model is based, implicitly assumes a no-flux boundary condition in the direction that chemical osmosis occurs, and consequently could lead to the overestimation of geologic osmotic pressures where a no-flux boundary condition cannot be assumed at a formation boundary. If this is the case, the permeability variation also exerts an influence on the estimation of the geologic osmotic pressure profile (e.g., Tremosa et al., 2012). In order to examine the applicability of dynamic equilibrium model, we compared our results with those estimated from steady-state analyses without a no-flux boundary condition using the same data shown in Figs. 4 and 5. The analysis used the continuity equation at steady state (Tremosa et al., 2012)
∂ ⎧ k⎡ dπ ρ (p − ρgz ) − σ ⎤ ⎫ = 0 ∂z ⎨ μ dz ⎦ ⎬ ⎣ ⎩ ⎭
(8)
The thickness of the Wakkanai formation at the SAB-2 borehole has not been investigated, therefore the hydraulic boundary condition at the bottom of the formation cannot be explicitly defined. Thus, the formation thickness was assumed as 1000 m, referring to the nearby geologic profile (Ota et al., 2011), and the fluid pressure at the bottom of the formation was assumed as a constant hydrostatic pressure at that depth, which is the most conservative boundary condition for the emergence of geologic osmotic pressure. Field and laboratory data necessary for the analyses are also not available for depths below 700 m (Figs. 4 and 5), so they were assumed to be the same as those at 700 m. The vertical profile of k can be estimated from rock physical properties, assuming Poiseuille’s law for a plane, parallel pore geometry (Tremosa et al., 2012)
k=
b2 3n−2.3
4.2. Topography-driven flow The present analyses are based on a one-dimensional vertical model, so they cannot reflect topography-driven flow that governs the threedimensional groundwater flow regime, in accordance with the threedimensional permeability structure. The Horonobe region contains hilly and sloping terrain that could store higher- or lower-elevation water tables than that of the SAB-2 borehole (Fig. 1). In addition, the hydrogeologic environment in the Horonobe region has been affected by tectonic deformation (Fig. 1 and Ishii et al., 2011b), so the connectivity of conductive fractures and major faults makes the permeability structure of this region complex on both local and regional scales (Kurikami et al., 2008; Ishii, 2018). With these topography-driven effects and a complex permeability structure, the groundwater flow regime, and thus the fluid pressure distribution in this region, is very uncertain. Indeed, it is difficult for topography-driven flow to
(9)
Fig. 7a shows the estimated profile for k. The profile of k was also evaluated using the experimental values of k: 1.41 × 10−19 and 4.29 × 10−20 (m2), which were measured by the same experiments as the osmotic efficiencies at depths of 401 and 698 m in the SAB-2 borehole (Takeda and Manaka, 2018). From these experimentally derived k values, the profile of k was estimated, assuming a linear relationship of logk - z (Kurikami et al., 2008). Fig. 7a shows that the estimated intrinsic permeability ranges from 1.3 × 10−20 to 7.0 × 10−19 (m2), which corresponds approximately to the intrinsic 6
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Fig. 7. Vertical profiles of (a) intrinsic permeability determined from the porosity and nominal half distance between clay surfaces using Eq. (9) and measured in chemical osmosis experiments performed by Takeda and Manaka (2018), and potential geologic osmotic pressures estimated from steady-state models with bottom boundary conditions of (b) constant pressure equivalent to hydrostatic pressure and (c) no flux.
have equilibrated with hydrostatic fluid pressure in the open portion of the borehole above the packer. Consequently, the excess pressures might have been measured at lower values than those present prior to drilling. This effect should be ascertained by using a steady-state analysis that describes the three-dimensional configuration of fluid measurement and permeability distribution within the formation for further study.
comprehensively explain the fluid overpressures observed in this region, and therefore the calibration of permeability distributions (e.g., Suko et al., 2014) or the assumption of a caprock (Karasaki et al., 2011) have been used to reproduce the observed fluid pressures in groundwater flow modeling. Now that the potential for geologic osmotic pressure in the Wakkanai formation is indicated, incorporating its effect into conventional modeling may improve the misfit between observed and predicted fluid pressures in this region. Also, the local fluid pressure disparity observed in our analyses may be improved by using simulations that also account for topography-driven flow.
5. Conclusions The quantitative analysis presented herein reveals that the fluid overpressures observed in the Wakkanai Formation, which cannot be explained by gravity-driven flow due to topographic effects (Karasaki et al., 2011) or pressure-driven flow due to dynamic geologic processes (Neuzil, 2015), may be explained as geologic osmotic pressures. However, further research is needed to ascertain whether geologic osmotic pressures are the primary cause of the observed fluid overpressure in the Wakkanai Formation in this region. The effects of topographydriven lateral flow should be also addressed to improve further detailed studies. Also, scrutinizing measured fluid pressures using numerical modeling and detailed semipermeability modeling that considers varying groundwater compositions and additional semipermeability measurements may improve the misfit between the calculated and observed values. Generally, geologic osmotic pressure is relatively small in comparison to fluid overpressures caused by other geologic processes such as tectonic loading or compaction disequilibrium (e.g., Osborne and Swarbrick, 1997; Gonçalvès et al., 2010); however, it could be a dominant process that affects groundwater flow and associated solute transport over geologically long timescales in stable argillaceous
4.3. Errors in fluid pressure measurement Except the natural process described above, artificial errors in fluid pressure measurements can result in a disparity between observed and estimated values. Errors in borehole measurements can be introduced by insufficient recovery of fluid pressures after drilling and/or by the short-circuiting of fluid from a measurement interval to the open hole above or below the measurement interval (e.g., Quinn et al., 2012). Although the fluid pressure measurements in the SAB-2 borehole were performed during the course of drilling, the measurements were made after the fluid pressure in the packed off interval reached equilibrium (Suko et al., 2014). Therefore, the possibility of insufficient pressure recovery can be ruled out. However, short-circuiting of fluid through the surrounding formation to the open borehole can occur, particularly at shallow depths in the SAB-2 borehole where the observed overpressures are below our estimated pressures. Because the permeability of the shallow part of the Wakkanai formation is relatively high (Kurikami et al., 2008; Takeda et al., 2014; Takeda & Manaka, 2018), the pressure of the fluid isolated below the packer in the borehole might 7
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formations (e.g., Neuzil, 2000). Therefore, the identification of geologic osmotic pressure is also crucial for investigating nuclear waste isolation in argillaceous formation sites (Neuzil, 2000; Wilson et al., 2003; Neuzil and Provost, 2009). The approach used in this study can only be applied to laterally uniform argillaceous formations. However, an approach using profiles of salinity and basic rock physical properties can be conducted without initial and boundary conditions or transport parameters, such as semipermeability, permeability, and diffusivity, which are necessary for the type of conventional simulation that fully describe transport processes within an argillaceous formation. Accordingly, the approach used in this study may help identify potential geologic osmotic pressures in argillaceous formations, especially during the early stages of hydrogeologic investigations when adequate data for simulating transport may not be available.
Oxfordian at the Bure site. Phys. Chem. Earth. 32, 434–440. Horseman, S.T., Harrington, J.F., Noy, D.J., 2007. Swelling and osmotic flow in a potential host rock. Phys. Chem. Earth 32 (1–7), 408–420. https://doi.org/10.1016/j. pce.2006.05.003. Ishii, E., Sanada, H., Iwatsuki, T., Sugita, Y., Kurikami, H., 2011a. Mechanical strength of the transition zone at the boundary between opal-a and opal-CT zones in siliceous rocks. Eng. Geol. 122 (3–4), 215–221. https://doi.org/10.1016/j.enggeo.2011.05. 007. Ishii, E., Sanada, H., Funaki, H., Sugita, Y., Kurikami, H., 2011b. The relationships among brittleness, deformation behavior, and transport properties in mudstones: an example from the Horonobe Underground Research Laboratory, Japan. J. Geophys. Res. 116, B09206. https://doi.org/10.1029/2011JB008279. Ishii, E., 2018. Assessment of hydraulic connectivity of fractures in mudstones by singleborehole investigations. Water Resour. Res. 54, 3335–3356. https://doi.org/10. 1029/2018WR022556. Iwatsuki, T., Ishii, E., Niizato, T., 2009. Scenario development of long-tern evolution for deep hydrochemical conditions in Horonobe area, Hokkaido, Japan. J. Geogr. 118, 700–716. https://doi.org/10.5026/jgeography.118.700. Karasaki, K., Ito, K., Wu, Y.S., Shimo, M., Sawada, A., Maekawa, K., Hatanaka, K., 2011. Uncertainty reduction of hydrologic models using data from surface-based investigation. J. Hydrol. 403 (1–2), 49–57. https://doi.org/10.1016/j.jhydrol.2011.03. 039. Kemper, W.D., Rollins, J.B., 1966. Osmotic efficiency coefficients across compacted clays. Soil Sci. Soc. Am. J. 30 (5), 529–534. https://doi.org/10.2136/sssaj1966. 03615995003000050005x. Kurikami, H., 2007. Groundwater flow analysis in the Horonobe Underground Research Laboratory project – Recalculation based on the investigation until fiscal year 2005 – (JAEA-Research-2007-036). Ibaraki, Japan: Japan Atomic Energy Agency (JAEA). Kurikami, H., Takeuchi, R., Yabuuchi, S., 2008. Scale effect and heterogeneity of hydraulic conductivity of sedimentary rocks at Horonobe URL site. Phys. Chem. Earth. 33, S37–S44. https://doi.org/10.1016/j.pce.2008.10.016. Manaka, M., Takeda, M., 2018. An experimental examination of the effect of the salt type on the chemo-osmotic, diffusive, and hydraulic parameters of Wakkanai mudstones. J. Hydrol. 564, 1058–1073. https://doi.org/10.1016/j.jhydrol.2018.07.061. Marine, I.W., Fritz, S.J., 1981. Osmotic model to explain anomalous hydraulic heads. Water Resour. Res. 17, 73–82. https://doi.org/10.1029/WR017i001p00073. Matsui, H., Kurikami, H., Kunimaru, T., Morioka, H., Hatanaka, K., 2007. Horonobe URL project: present status and future plans. In: Eberhardt, E., Stead, D., Morrison, T. (Eds.), Rock Mechanics: Meeting Society’s Challenges and Demands. London, U. K, Taylor and Francis Group, pp. 1193–1201. Mazurek, M., Alt-Epping, P., Bath, A., Gimmi, T., Waber, H.N., Buschaert, S., et al., 2011. Natural tracer profiles across argillaceous formations. Appl. Geochem. 26 (7), 1035–1064. https://doi.org/10.1016/j.apgeochem.2011.03.124. Neuzil, C.E., 2000. Osmotic generation of ‘anomalous’ fluid pressures in geological environments. Nature 403 (6766), 182–184. https://doi.org/10.1038/35003174. Neuzil, C.E., Provost, A.M., 2009. Recent experimental data may point to a greater role for osmotic pressures in the subsurface. Water Resour. Res. 45, W03410. https://doi. org/10.1029/2007WR006450. Neuzil, C.E., 2015. Interpreting fluid pressure anomalies in shallow intraplate argillaceous formations. Geophys. Res. Lett. 42, 4801–4808. https://doi.org/10.1002/ 2015GL064140. Noy, D.J., Horseman, S.T., Harrington, J.F., Bossart, P., Fisch, H.R., 2004. An experimental and modelling study of chemico-osmotic effects in the Opalinus Clay of Switzerland. No. 6 In: Heitzmann, P. (Ed.), Mont Terri project—Hydrogeological synthesis, osmotic flow, Geology Series. Federal Office for Water and Geology, Bern, Switzerland, pp. 95–126. Oduor, P.G., Santos, X.T., Casey, F.X., 2009. Solute exclusionary properties of porous shale wafers. J. Porous Media 12 (6), 501–518. https://doi.org/10.1615/JPorMedia. v12.i6.20. Osborne, M.J., Swarbrick, R.E., 1997. Mechanisms for generating overpressure in sedimentary basins: a reevaluation. Bull. Am. Assoc. Petrol. Geol. 81, 1023–1041. Ota, K., Abe, H., Kunimaru, T., 2011. Horonobe underground research laboratory project synthesis of phase I investigations 2001–2005 volume ‘‘Geoscientific Research.’’ (JAEA-Data/Code 2010-068). Ibaraki, Japan: Japan Atomic Energy Agency (JAEA). Quinn, P., Cherry, J.A., Parker, B.L., 2012. Hydraulic testing using a versatile straddle packer system for improved transmissivity estimation in fractured-rock boreholes. Hydrogeol. J. 20, 1529–1547. https://doi.org/10.1007/s10040-012-0893-8. Revil, A., Leroy, P., 2004. Constitutive equations for ionic transport in porous shales. J. Geophys. Res. 109, B03208. https://doi.org/10.1029/2003JB002755. Revil, A., 2017. Transport of water and ions in partially water-saturated porous media. Part 2. Filtration effects. Adv. Water Resour. 103, 139–152. https://doi.org/10.1016/ j.advwatres.2016.07.016. Rousseau-Gueutin, P., de Greef, V., Gonçalvès, J., Violette, S., Chanchole, S., 2009. Experimental device for chemical osmosis measurement on natural clay-rock samples maintained at in situ conditions: implications for formation pressure interpretations. J. Colloid Interface Sci. 337 (1), 106–116. https://doi.org/10.1016/j.jcis.2009.04. 092. Rousseau-Gueutin, P., Gonçalvès, J., Cruchaudet, M., de Marsily, G., Violette, S., 2010. Hydraulic and chemical pulse tests in a shut-in chamber imbedded in an argillaceous formation: numerical and experimental approaches. Water Resour. Res. 46, W08516. https://doi.org/10.1029/2008WR007371. Rousseau-Gueutin, P., Gonçalvès, J., Violette, S., 2008. Osmotic efficiency in CallovoOxfordian argillites: experimental vs. theoretical models. Phys. Chem. Earth 33, S106–S113. https://doi.org/10.1016/j.pce.2008.10.064. Suko, T., Takano, H., Uchida, M., Seki, Y., Ito, K., Watanabe, Y., et al., 2014. Research on Validation of the Groundwater Flow Evaluation Methods Based on the Information of
Declaration of Competing Interest The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper. Acknowledgments The main part of this research project was conducted as regulatorysupporting research funded by the Secretariat of the Nuclear Regulation Authority, Japan. The authors wish to thank the Japan Atomic Energy Agency for providing mudstone samples and Kazuya Miyakawa and Eiichi Ishii for fruitful comments and discussions. References Amano, Y., Yamamoto, Y., Nanjyo, I., Murakami, H., Yokota, H., Yamazaki, M., et al., 2012. Data of Groundwater from Boreholes, River Water and Precipitation for the Horonobe Underground Research Laboratory Project (2001-2010) (JAEA-Data/Code 2011-023). Japan Atomic Energy Agency (JAEA), Ibaraki, Japan. Appelo, C.A.J., Postma, D., 2005. Geochemistry, Groundwater and Pollution, 2nd ed. Balkema Publishers, Netherlands. Bader, S., Kooi, H., 2005. Modelling of solute and water transport in semi-permeable clay membranes: comparison with experiments. Adv. Water Resour. 28, 203–214. https:// doi.org/10.1016/j.advwatres.2004.11.001. Bresler, E., 1970. Numerical solution of the equation for interacting diffuse layers in mixed ionic systems with nonsymmetrical electrolytes. J. Colloid Interface Sci. 33, 278–283. https://doi.org/10.1016/0021-9797(70)90030-5. Bresler, E., 1973. Anion exclusion and coupling effects in nonsteady transport through unsaturated soils: I. Theory. Soil Sci. Soc. Am. Proc. 37 (5), 663–669. https://doi.org/ 10.2136/sssaj1973.03615995003700050013x. Brunauer, S., Emmett, P.H., Teller, E., 1938. Adsorption of gases in multimolecular layers. J. Am. Chem. Soc. 60 (2), 309–319. https://doi.org/10.1021/ja01269a023. Cey, B.D., Barbour, S.L., Hendry, M.J., 2001. Osmotic flow through a cretaceous clay in southern Saskatchewan, Canada. Canad. Geotech. J. 38 (5), 1025–1033. https://doi. org/10.1139/t01-033. Cruchaudet, M., Croisé, J., Lavanchy, J.M., 2008. In situ osmotic experiment in the Callovo-Oxfordian argillaceous formation at the Meuse/Haute-Marne URL (France): Data and analysis. Phys. Chem. Earth. 33, S114–S124. https://doi.org/10.1016/j.pce. 2008.10.045. Fritz, S.J., 1986. Ideality of clay membranes in osmotic processes: a review. Clays Clay Miner. 34 (2), 214–223. https://doi.org/10.1346/CCMN.1986.0340212. Garavito, A.M., De Cannière, P., Kooi, H., 2007. In situ chemical osmosis experiment in the Boom Clay at the Mol underground research laboratory. Phys. Chem. Earth. 32 (1–7), 421–433. https://doi.org/10.1016/j.pce.2006.01.004. Gonçalvès, J., Violette, S., Wendling, J., 2004. Analytical and numerical solutions for alternative overpressuring processes: application to the Callovo-Oxfordian sedimentary in the Paris basin, France. J. Geophys. Res. 109, B02110. https://doi.org/10. 1029/2002JB002278. Gonçalvès, J., Rousseau-Gueutin, P., Revil, A., 2007. Introducing interacting diffuse layers in TLM calculations: a reappraisal of the influence of the pore size on the swelling pressure and the osmotic efficiency of compacted bentonites. J. Colloid Interface Sci. 316, 92–99. https://doi.org/10.1016/j.jcis.2007.07.023. Gonçalvès, J., Rousseau-Gueutin, P., de Marsily, G., Cosenza, P., Violette, S., 2010. What is the significance of pore pressure in a saturated shale layer? Water Resour. Res. 46, W04514. https://doi.org/10.1029/2009WR008090. Greenberg, J.A., Mitchell, J.K., Witherspoon, P.A., 1973. Coupled salt and water flows in a groundwater basin. J. Geophys. Res. 78 (27), 6341–6353. https://doi.org/10.1029/ JC078i027p06341. Gueutin, P., Altmann, S., Gonçalvès, J., Cosenza, P., Violette, S., 2007. Osmotic interpretation of overpressures from monovalent based triple layer model, in the Callovo-
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Agency (JAEA), Ibaraki, Japan: Japan. Tremosa, J., Gonçalvès, J., Matray, J.M., 2012. Natural conditions for more limited osmotic abnormal fluid pressures in sedimentary basins. Water Resour. Res. 48, W04530. https://doi.org/10.1029/2011WR010914. Wilson, A.M., Fenstemaker, T., Sharp Jr., J.M., 2003. Abnormally pressured beds as barriers to diffusive solute transport in sedimentary basins. Geofluids 3 (3), 203–212. https://doi.org/10.1046/j.1468-8123.2003.00060.x. Young, A., Low, P.F., 1965. Osmosis in argillaceous rocks. Bull. Am. Assoc. Petrol. Geol. 49 (7), 1004–1008. Yu, C.J., 2017. Comparative Study of Convective and Diffusive Transport Phenomena within the Opalinus Clay of Mont Terri. Doctoral dissertation. Aix-Marseille University. https://www.irsn.fr/FR/Larecherche/Formation_recherche/Theses/ Theses-soutenues/PRP-DGE/Documents/2017-These-Yu.pdf.
Geological Environment In and Around Horonobe Underground Research Area (JNES-RE-2013-9032). Nuclear Energy Safety Organization (JNES), Tokyo, Japan: Japan. Takeda, M., Hiratsuka, T., Manaka, M., Finsterle, S., Ito, K., 2014. Experimental examination of the relationships among chemico-osmotic, hydraulic, and diffusion parameters of Wakkanai mudstones. J. Geophys. Res. 119, 4178–4201. https:// doi:10.1002/2013JB010421. Takeda, M., Manaka, M., 2018. Effects of confining stress on the semipermeability of siliceous mudstones: implications for identifying geologic membrane behaviors of argillaceous formations. Geophys. Res. Lett. 45, 5427–5435. https://doi.org/10. 1029/2018GL078591. Tokiwa, T., Sugimoto, S., 2012. Crustal Deformation Analysis by Using GPS Observation Data in the Northern Hokkaido, Japan. (JAEA-Data/Code 2012–024). Atomic Energy
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Research papers
Potential geologic osmotic pressure in the Wakkanai Formation: Preliminary estimation based on the dynamic equilibrium between chemical osmosis and advection
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Mikio Takeda , Mitsuo Manaka, Kazumasa Ito National Institute of Advanced Industrial Science and Technology (AIST), Central 7, 1-1-1 Higashi, Tsukuba, Ibaraki 305-8560, Japan
A R T I C LE I N FO
A B S T R A C T
This manuscript was handled by Corrado Corradini, Editor-in-Chief, with the assistance of Patrick Lachassagne, Associate Editor
Geologic osmotic pressure has been proposed as a cause of excess fluid pressures in argillaceous formations in which dynamic geologic processes have little effect on pressure. However, the identification of geologic osmotic pressure has rarely been attempted, thus its potential effect on groundwater systems is a longstanding question in hydrogeology. Herein we present a quantitative analysis of excess fluid pressures of unknown origin in the Wakkanai Formation, assuming osmotically generated overpressure. Potential geologic osmotic pressures are estimated from profiles of pore water salinity and rock physical properties using a one-dimensional model that describes geologic osmotic pressures in a state of dynamic equilibrium between chemical osmosis and advection. The model results indicate that the Wakkanai Formation has the requisite conditions of both salinity gradient and semipermeability to generate geologic osmotic pressures comparable to observed excess pressures, up to 230 kPa. The applicability of the dynamic equilibrium model is also confirmed by using a steady-state model with the same pore water salinity and rock physical properties data. These results suggest that chemical osmosis is a potential cause for the excess fluid pressures observed in the Wakkanai Formation.
Keywords: Chemical osmosis Osmotic pressure Semipermeability Argillite Excess fluid pressure
1. Introduction Chemical osmosis that arises from the semipermeability of argillite and salinity differences in groundwater generates geologic osmotic pressure within argillaceous formations, which is observed as overpressure or underpressure relative to hydrostatic or hydrodynamically equilibrated fluid pressure (Greenberg et al., 1973; Neuzil, 2000; Wilson et al., 2003; Gonçalvès et al., 2004; Neuzil and Provost, 2009). Chemical osmosis occurs from low- to high-salinity regions of an argillaceous formation, inducing fluid pressure buildup in the high-salinity region and drawdown in the low-salinity region. The pressure gradient that develops within the formation causes advection that counters the chemical osmotic flow and causes the osmotically induced pressure to dissipate. However, because of the low permeability of argillites, the osmotically induced pressure is maintained in a quasiequilibrium state as the counteracting flows ultimately reach a state of dynamic equilibrium (Neuzil, 2000; Gonçalvès et al., 2004; Bader and Kooi, 2005; Takeda et al., 2014). Since the salinity profile controlling osmotic flow is generally in a transient state and is mainly in the diffusion-dominated regime of argillaceous formations (e.g., Mazurek et al., 2011), the osmotically induced pressure (i.e., the geologic
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osmotic pressure) lasts over geologically long timescales (Neuzil, 2000; Gonçalvès et al., 2004). Geologic osmotic pressure is relatively small in comparison to excess fluid pressures caused by dynamic geologic processes such as tectonic loading or compaction disequilibrium (e.g., Osborne and Swarbrick, 1997; Gonçalvès et al., 2010); however, ignoring or underestimating the role of geologic osmotic pressure may result in misinterpretation of groundwater flow directions and associated solute transport (e.g., Marine and Fritz, 1981). Therefore, when excess fluid pressures are encountered in hydrogeologic investigations, the possibility of geologic osmotic pressure needs to be identified, particularly in the case of nuclear waste isolation in argillaceous formations (Neuzil, 2000, Wilson et al., 2003; Neuzil and Provost, 2009). Geologic osmotic pressure has been proposed as a cause of excess fluid pressures in argillaceous formations in which dynamic geologic processes are thought to be too minor to affect pressures (e.g., Young and Low, 1965; Marine and Fritz, 1981). However, the idea of geologic osmotic pressure has long been controversial because of the lack of data directly supporting its existence in situ. Since evidence for the semipermeability of argillites to produce geologic osmotic pressure was first obtained from in situ borehole experiments on the Pierre Shale, performed by Neuzil (2000), semipermeability has been investigated
Corresponding author. E-mail address: [email protected] (M. Takeda).
https://doi.org/10.1016/j.jhydrol.2019.124166 Received 24 June 2019; Received in revised form 28 August 2019; Accepted 19 September 2019 Available online 20 September 2019 0022-1694/ © 2019 Elsevier B.V. All rights reserved.
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denudated (Iwatsuki et al. 2009). Although many borehole investigations have been conducted in this area (Matsui et al., 2007), detailed semipermeability measurements have been only performed on Wakkanai mudstones retrieved from the SAB-2 borehole (Fig. 1a) (Takeda and Manaka, 2018). Accordingly, the data obtained from the SAB-2 borehole were used in the present study. Fig. 1b shows the lithology of the SAB-2 borehole. The Wakkanai Formation mainly consists of siliceous mudstone, and its shallow section (5–75 m depth) is a transition zone between siliceous and diatomaceous mudstones. Fig. 2 shows the data obtained from the SAB-2 borehole (Amano et al., 2012; Suko et al., 2014). Fig. 2a shows that the dominant ions in pore water are sodium and chloride. Their concentrations tend to increase with depth, suggesting that the salinity gradient within the Wakkanai Formation may induce downward chemical osmosis to generate geologic osmotic pressures at depth. Fig. 2b shows the temperature measured during drilling of the SAB-2 borehole (Suko et al., 2014). The temperature gap at 450 m depth is due to separate drilling stages for the shallow and deep sections, which occurred in different seasons. Fig. 2c shows that the observed fluid pressures are close to hydrostatic fluid pressures determined from the varying solute concentration and temperature. However, the observed fluid pressures tend to deviate from hydrostatic with increasing depth, with a maximum deviation of 230 kPa at a depth of 660 m.
extensively in other argillites, including the Opalinus Clay (Noy et al., 2004; Horseman et al., 2007), Boom Clay (Garavito et al., 2007), Callovo-Oxfordian argillite (Gonçalvès et al., 2004; Gueutin et al., 2007, Cruchaudet et al., 2008; Rousseau-Gueutin et al., 2008, 2009, 2010), and Wakkanai mudstone (Takeda et al., 2014; Takeda and Manaka, 2018; Manaka and Takeda, 2018). So far, extensive studies of the Callovo-Oxfordian formation have revealed that geologic osmotic pressure can explain 0.1–0.15 MPa of the 0.5–0.6 MPa overpressures observed in the formation (Gonçalvès et al., 2004; Rousseau-Gueutin et al., 2009). A one-dimensional, steady-state analysis performed by Tremosa et al. (2012) also suggests that the excess fluid pressures within the argillaceous formation at Tournemire, France may be partly due to geologic osmotic pressures. Neuzil and Provost (2009) examined data in the existing literature on the semipermeability of argillites and suggested that all analyzed media exhibited at least a modest ability to generate osmotic pressure. However, geologic osmotic pressures quantitatively identified based on field and laboratory data are still limited, thus the potential effects of geologic osmotic pressures on groundwater systems are still largely unknown. Here we report on a quantitative analysis of excess fluid pressures of unknown origin, observed in the Wakkanai Formation in the Horonobe region of Japan, as osmotically generated. In the Wakkanai Formation, excess fluid pressures up to 400 kPa and salinities up to half that of seawater have been observed at depth (Kurikami, 2007; Iwatsuki et al., 2009). The excess fluid pressures at depth in this region can be explained by topography-driven flow, but needs the assumption of a caprock in hydrogeological modeling that has not been observed by borehole investigations nor by permeability experiments (Karasaki et al., 2011). In addition, although the Horonobe region is located near the North American-Eurasian plate boundary, the strain rate in this region is too low to increase fluid pressures as a result of relatively high permeability (Tokiwa and Sugimoto, 2012; Neuzil, 2015). Furthermore, a recent study on the semipermeability of the Wakkanai mudstone suggests that the mudstone may act as a weak geologic membrane (Takeda and Manaka, 2018). In this study, we use a one-dimensional model that describes geologic osmotic pressures in a state of dynamic equilibrium between chemical osmosis and advection, which was originally developed by Neuzil (2000) for theoretically screening favorable conditions for generating geologic osmotic pressures. In order to estimate geologic osmotic pressures from salinity distribution and rock physical properties, both of which are obtained from field and laboratory investigations, we use a first-order approximation of the model, the effectiveness of which was demonstrated by Yu (2017) in analyses of geologic osmotic pressures in the Opalinus Clay of Mont Terri in northwestern Switzerland. We also compare our results to the geologic osmotic pressure estimated from experimentally determined semipermeabilities of the Wakkanai mudstone. The applicability of the dynamic equilibrium model is also tested against a steady-state model using additional hypothetical data and boundary conditions. Finally, we discuss the potential causes affecting fluid pressure distribution in the Horonobe region to assess the relevance of geologic osmotic pressure to the overpressures observed in this region.
2.2. Analysis based on the state of dynamic equilibrium between chemical osmosis and advection The present analysis assumes a horizontally uniform argillaceous formation in which lateral hydrodynamic fluxes are negligibly small. Based on this assumption, one-dimensional, vertical pore water flux through the argillaceous formation is described by a generalized Darcy’s law equation (e.g., Neuzil and Provost, 2009; Tremosa et al., 2012):
∂π k ∂ q = − ⎧ (p − ρgz ) − σ ⎫ ∂z ⎬ μ⎨ ⎭ ⎩ ∂z
(1)
where q is the pore water flux (m/s), k is the intrinsic permeability (m2), μ is the dynamic viscosity (Pa·s), p is the fluid pressure (Pa), ρ is the pore water density (kg/m3), g is gravitational acceleration (m/s2), z is the depth below the ground level (m), σ is the osmotic efficiency (-), and π is the theoretical osmotic pressure (Pa). σ is a dimensionless parameter that ranges between 0 and 1 by definition and represents semipermeability that is controlled by the pore size between clay minerals, the electrochemical properties of clay minerals, and the solute concentration of the pore water (e.g., Bresler, 1973). The osmotic pressure π has the same dimension as pressure, but is a measure of the decrease in the chemical potential of water (Neuzil and Provost, 2009). The first and second terms on the right side of Eq. (1) represent advection driven by pressure and gravity potentials and chemical osmosis driven by chemical potential, respectively. When a state of dynamic equilibrium is established between chemical osmosis and advection, the net water flux through the argillaceous formation becomes almost zero, and geologic osmotic pressure emerges as the excess fluid pressure relative to the hydrostatic or hydrodynamic fluid pressures. This situation was originally considered by Neuzil (2000) for theoretically screening favorable conditions for generating geologic osmotic pressures, and it gives Eq. (1) the constraint of q = 0. Further introducing p = phy + pos, where phy is the hydrostatic or hydrodynamic pressure (Pa) and pos is the geologic osmotic pressure (Pa), Eq. (1) can be rewritten as
2. Estimation of geologic osmotic pressure in the Wakkanai formation 2.1. Geological setting and data suggesting geologic osmotic pressures The Wakkanai Formation is located in the Horonobe area of Hokkaido prefecture, Japan (Fig. 1a). According to Iwatsuki et al. (2009), the Wakkanai Formation and overlying formations were originally formed in a marine environment during the Neogene to Quaternary periods, then were compacted during subsidence prior to 1 Ma. The maximum burial depth of the top of the Wakkanai Formation is estimated to be 1000 m (Ishii et al., 2011a). The strata in this area have been uplifted due to folding and thrusting, and the top layers have been
∂p ∂π k ∂ 0 = − ⎧ (phy − ρgz ) + ⎛ os − σ ⎞ ⎫ μ⎨ ∂z ⎠ ⎬ ⎝ ∂z ⎭ ⎩ ∂z ⎜
⎟
(2)
Assuming that hydraulic flux, which is represented by the first term on the right side of Eq. (2), is negligible, owing to hydrostatic or 2
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Fig. 1. Location and stratigraphic section of the SAB-2 borehole (based on Ishii et al., 2011b; Suko et al., 2014). (a) Location map, geologic map, and geologic cross section of the Horonobe area in Hokkaido, showing the location of the SAB-2 borehole. The bold line in the geologic cross section indicates the approximate projection of the borehole along the faults. (b) Stratigraphic section of the SAB-2 borehole.
Fig. 2. Data collected from the SAB-2 borehole. Vertical distributions of (a) the concentrations of major groundwater solute components (Amano et al., 2012), (b) temperature, and (c) pore fluid pressure measured during the course of drilling (Suko et al., 2014). Hydrostatic fluid pressures in (c) are determined from the total amount of solutes dissolved in groundwater collected from the SAB-2 borehole and its varying temperature.
osmosis, represented by the right side of Eq. (3), is equal to the counteracting advection driven by the geologic osmotic pressure gradient when the system is at dynamic equilibrium. Upon integration, Eq. (3) can be solved over the vertical domain when σ and π are given as smooth functions of z and/or solute
hydrodynamic equilibrium, Eq. (2) can be rewritten as (e.g., Neuzil and Provost, 2009)
∂pos ∂π =σ ∂z ∂z
(3)
Although k and μ are eliminated, Eq. (3) states that chemical 3
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concentration (Neuzil, 2000; Neuzil and Provost, 2009). However, the present analysis addresses field and laboratory data that were discretely obtained from a borehole (Fig. 2), so pos is calculated using the discretized form of Eq. (3)
Δpos (z i ) = σi Δπ (z i )
(4)
where i is the index for the discretized formation element (i = 1 for the top element); zi is the vertical coordinate of the center of the ith element; Δpos and Δ π are the changes in the geologic and theoretical osmotic pressure, respectively, along the ith element; and σ i is the osmotic efficiency of the ith element. Summing Eq. (4) from the top to the nth element, the geologic osmotic pressure at the nth element can be estimated as n
pos (z n ) =
∑ Δpos (z i) i=1
(5)
The above approach for assigning osmotic efficiencies and osmotic pressure is essentially the same as the first-order approximation of the generalized Darcy’s law adopted by Yu (2017). 2.3. Osmotic efficiency calculation Fig. 3. Osmotic efficiency σ as a function of the nominal half distance between clay surfaces b and pore fluid concentration C for monovalent solutes. The curve is the Kemper–Bresler model (Bresler, 1973). The plots represent the results of fitting the Kemper–Bresler model to osmotic efficiencies of 0.021 and 0.081, based on samples from the Wakkanai mudstone at 401 and 698 m in the SAB-2 borehole, respectively (Takeda and Manaka, 2018). The gray lines indicate salinity ranges of 0.6–0.1 M NaCl for each measurement.
The osmotic efficiency, σ in Eq. (4), is a measure of the anion exclusion effect that is exerted by the overlap of electrical double layers (EDLs) at pore spaces between negatively charged clay surfaces (e.g., Marine and Fritz, 1981; Appelo and Postma, 2005). Accordingly, σ varies by location in argillaceous formations in accordance with the factors affecting the EDLs overlap, such as pore size between clay minerals, electrochemical properties of the clay minerals, and solute concentrations of the pore water. Among a number of models developed for the anion exclusion effect (e.g., Kemper and Rollins, 1966; Bresler, 1973; Marine and Fritz, 1981; Fritz, 1986; Revil and Leroy, 2004; Gonçalvès et al., 2007; Tremosa et al., 2012; Revil, 2017), the Kemper–Bresler model for monovalent solutes (Bresler, 1973) was used in our analysis because the model has an advantage of describing osmotic efficiency as a single curve relating σ to b C , where b is the nominal half distance between clay surfaces in Å (Fig. 3). In addition, the Kemper–Bresler model has previously been successfully applied to describe the osmotic efficiencies of undisturbed argillaceous formation media (Cey et al., 2001; Neuzil and Provost, 2009; Oduor et al., 2009; Rousseau-Gueutin et al., 2008; Takeda et al., 2014). Although the groundwater recovered from the SAB-2 borehole contains divalent cations, which compress the EDLs more than monovalent cations (e.g., Bresler, 1970; Bresler, 1973; Tremosa et al., 2012), the concentration of divalent cations in the groundwater is considerably lower and accounts for only 0.3–2.7% of the total cations and 1.3% on average (Fig. 2a and Amano et al., 2012). These values are close to that of the multi-component solution adopted in the experiments performed by Cey et al. (2001), in which the osmotic efficiencies were well characterized by the Kemper–Bresler model for a NaCl solution. Parameter b in the Kemper–Bresler model has been related to rock physical properties by Neuzil (2000) and is expressed as
b=
n ρb As
The b values were also estimated from the experimental values of σ: 0.021 and 0.081, which were measured in the cores retrieved from depths of 401 and 698 m, respectively, in the SAB-2 borehole, at a salinity difference of 0.6–0.1 M NaCl (Takeda and Manaka, 2018). These experimental σ values are those measured at effective confining stresses for individual cores after applying their past maximum effective burial stresses. This is because the Wakkanai Formation is currently overconsolidated as a result of the denudation of overlying strata. For these experimental σ values, the b values were adjusted to fit the Kemper–Bresler model with the measured σ values (Fig. 3). The estimated b values are 31 and 21 Å for the cores retrieved from depths of 401 and 698 m, respectively (Fig. 4d). These values are relatively close to those calculated from physical properties data. However, the measured σ values were obtained at a relatively large salinity difference of 0.6–0.1 M NaCl, so the representativeness of the estimated b values is ambiguous (Fig. 3). Also, the number of experimentally derived b values were not sufficient to delineate a vertical profile for b. Accordingly, these experimentally derived b values are used as a reference in the analysis, using a linear function with respect to z for calculating the vertical profiles of osmotic efficiency (Fig. 4d).
2.4. Theoretical osmotic pressure approximation The theoretical osmotic pressure, π in Eq. (4), is the decrease in the chemical potential of water due to the presence of solutes, thus it is affected by the concentration of each solute present in the groundwater. Although the groundwater of the Wakkanai Formation contains various solutes, its dominant anionic and cationic species are chloride and sodium, respectively (Fig. 2a). We therefore assumed that the groundwater in the Wakkanai Formation was a NaCl solution. Fig. 5 shows the equivalent NaCl concentrations estimated from the total amounts of solutes dissolved in groundwater collected from the SAB-2 borehole (Amano et al., 2012). The equivalent NaCl concentrations ranged from 0.01 to 0.16 M. For solutions with concentrations lower than 1 M, the theoretical osmotic pressure can be approximated according to the van’t Hoff relation (Fritz, 1986): π = νRTC, where ν is the number of
(6)
where n is the porosity, ρb is the dry bulk density, and As is the specific surface area. In order to obtain the vertical profile of b in the SAB-2 borehole, we measured n and ρb using mercury intrusion porosimetry, and As using the Brunauer-Emmett-Teller method (Brunauer et al., 1938) on the cores retrieved from the borehole. Fig. 4 shows the measured n, ρb, and As values and the estimated b values. The profile of b exhibits fluctuations at depths of 649 and 698 m. However, b values for the other depths are relatively constant and within a range of 29–31 Å. The profile of σ, which is necessary for calculating Eq. (4), was evaluated using the Kemper–Bresler model and the profiles of b and the solute concentrations. 4
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Fig. 4. Physical properties measured from samples of the Wakkanai mudstone retrieved from the SAB-2 borehole. (a) Porosity and (b) medium dry bulk density were measured using mercury intrusion porosimetry. (c) Specific surface area was measured using the Brunauer-Emmett-Teller method (Brunauer et al., 1938). (d) The nominal half distance b between clay surfaces was determined from porosity, medium dry bulk density, and specific surface area using Eq. (6). The experimental values are those determined from the osmotic efficiencies measured from samples of the Wakkanai mudstone retrieved from 401 and 698 m depths in the SAB-2 borehole (Takeda and Manaka, 2018). The straight line on the experimental values represents the linear approximation of b with respect to z.
(Fig. 3). Using Eqs. (5) and (7) and the vertical profiles of T, C, and σ (Figs. 2b, 5 and 6a), the potential geologic osmotic pressures in the formation were estimated. Fig. 6b shows the vertical profiles of the estimated geologic osmotic pressures and the fluid overpressures observed at the SAB-2 borehole. The observed fluid overpressures were determined by subtracting the hydrostatic fluid pressures from the observed fluid pressures (Fig. 2c). The estimated geologic osmotic pressure profiles replicate the downward increase in fluid overpressure observed in the SAB-2 borehole. However, they fail to delineate the fluid overpressures observed at depths shallower than 370 m. In contrast, the potential geologic osmotic pressures estimated for depths below 490 m are relatively close to observed fluid overpressures. In particular, potential geologic osmotic pressures estimated from the linear approximation of experimental b values are almost equivalent to the observed fluid overpressures (Fig. 6b) despite the sparse semipermeability measurements (Fig. 4d) and the relatively high salinity conditions (Takeda and Manaka, 2018). On the other hand, the analysis performed using the b values determined from the physical properties of retrieved cores slightly overestimates the fluid overpressures (Fig. 6b). 4. Discussion
Fig. 5. Vertical profile of equivalent NaCl concentration estimated from the total amount of solutes dissolved in pore fluids in the SAB-2 borehole (Amano et al., 2012).
This study assumed chemical osmosis as the process that deviates pore fluid pressures from a hydrostatic distribution and exemplified the potential of the Wakkanai Formation to generate geologic osmotic pressures comparable to the observed fluid overpressures of 230 kPa. However, our analyses do not replicate observed fluid overpressures particularly well at depths shallower than 370 m in the SAB-2 borehole. This suggests that other factors still need to be addressed in order to determine the relevance of geologic osmotic pressure in the fluid overpressures observed in this region.
ionic species (e.g., ν = 2 for NaCl), R is the gas constant (8.314 m3 Pa/ K/mol), and T is the absolute temperature (K). Introducing this relationship, Eq. (4) can be rewritten as
Δp (z i ) = σi νRT ΔC (z i )
(7)
where ΔC is the change in NaCl concentration along the i element. The temperature within the Wakkanai Formation increases with depth (Fig. 2b), so we assumed that the temperature was a linear function of z (median line in Fig. 2b). th
4.1. Applicability of dynamic equilibrium model
3. Results
The dynamic equilibrium model adopted in this study enables the calculation of geologic osmotic pressure from limited salinity and basic rock physical properties data, independently from permeability variations. This is an advantage of the dynamic equilibrium model compared
Fig. 6a shows the vertical distribution of the osmotic efficiency calculated from b and C (Fig. 4d and 5) using the Kemper–Bresler model 5
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Fig. 6. Vertical profiles of (a) osmotic efficiency calculated using the Kemper–Bresler model (Bresler, 1973), the NaCl concentration profile (Fig. 5), and the profile of b values (Fig. 4d); and (b) observed fluid overpressure and potential geologic osmotic pressure estimated using Eqs. (5) and (7) and NaCl concentration and osmotic efficiency profiles (Figs. 5 and 6a) in the SAB-2 borehole.
permeabilities reported for the Wakkanai mudstone by Kurikami et al. (2008), which range from 10−20 to10−18 m2 and tend to decrease with increasing depth. Fig. 7b shows the profiles of geologic osmotic pressure estimated from the steady-state analyses with a constant hydrostatic pressure at the bottom boundary. The geologic osmotic pressures estimated from the steady-state analyses are slightly smaller than those estimated from the dynamic equilibrium model. This indicates that the magnitude of geologic osmotic pressure is dependent on osmosis-driven downward advection, which is controlled by permeability variations. However, the geologic osmotic pressures estimated from the steady-state analyses are still large and agree well with the observed overpressures at depths below 490 m, suggesting that the intrinsic permeabilities assumed for the Wakkanai Formation (Fig. 7a) are sufficiently small to maintain geologic osmotic pressures (e.g., Tremosa et al., 2012). Fig. 7c shows the profiles of geologic osmotic pressure estimated from the steadystate analyses with a no-flux boundary at the bottom, for reference. As expected, the estimated geologic osmotic pressures are almost the same as those from the dynamic equilibrium model. This explicitly indicates that the analysis based on the dynamic equilibrium model is equivalent to the steady-state analysis with the profile of k and a no-flux bottom boundary, and can be used to estimate the maximum potential geologic osmotic pressure for given salinity and rock physical property profiles. This is the advantage of using the dynamic equilibrium model when adequate data for performing steady-state or transient state modeling are not available, especially during the early stages of hydrogeologic investigations.
to the steady-state and transient models, both of which require additional data regarding transport parameters, such as permeability and diffusivity, and initial and boundary conditions at formation boundaries. However, Eq. (3), on which the dynamic equilibrium model is based, implicitly assumes a no-flux boundary condition in the direction that chemical osmosis occurs, and consequently could lead to the overestimation of geologic osmotic pressures where a no-flux boundary condition cannot be assumed at a formation boundary. If this is the case, the permeability variation also exerts an influence on the estimation of the geologic osmotic pressure profile (e.g., Tremosa et al., 2012). In order to examine the applicability of dynamic equilibrium model, we compared our results with those estimated from steady-state analyses without a no-flux boundary condition using the same data shown in Figs. 4 and 5. The analysis used the continuity equation at steady state (Tremosa et al., 2012)
∂ ⎧ k⎡ dπ ρ (p − ρgz ) − σ ⎤ ⎫ = 0 ∂z ⎨ μ dz ⎦ ⎬ ⎣ ⎩ ⎭
(8)
The thickness of the Wakkanai formation at the SAB-2 borehole has not been investigated, therefore the hydraulic boundary condition at the bottom of the formation cannot be explicitly defined. Thus, the formation thickness was assumed as 1000 m, referring to the nearby geologic profile (Ota et al., 2011), and the fluid pressure at the bottom of the formation was assumed as a constant hydrostatic pressure at that depth, which is the most conservative boundary condition for the emergence of geologic osmotic pressure. Field and laboratory data necessary for the analyses are also not available for depths below 700 m (Figs. 4 and 5), so they were assumed to be the same as those at 700 m. The vertical profile of k can be estimated from rock physical properties, assuming Poiseuille’s law for a plane, parallel pore geometry (Tremosa et al., 2012)
k=
b2 3n−2.3
4.2. Topography-driven flow The present analyses are based on a one-dimensional vertical model, so they cannot reflect topography-driven flow that governs the threedimensional groundwater flow regime, in accordance with the threedimensional permeability structure. The Horonobe region contains hilly and sloping terrain that could store higher- or lower-elevation water tables than that of the SAB-2 borehole (Fig. 1). In addition, the hydrogeologic environment in the Horonobe region has been affected by tectonic deformation (Fig. 1 and Ishii et al., 2011b), so the connectivity of conductive fractures and major faults makes the permeability structure of this region complex on both local and regional scales (Kurikami et al., 2008; Ishii, 2018). With these topography-driven effects and a complex permeability structure, the groundwater flow regime, and thus the fluid pressure distribution in this region, is very uncertain. Indeed, it is difficult for topography-driven flow to
(9)
Fig. 7a shows the estimated profile for k. The profile of k was also evaluated using the experimental values of k: 1.41 × 10−19 and 4.29 × 10−20 (m2), which were measured by the same experiments as the osmotic efficiencies at depths of 401 and 698 m in the SAB-2 borehole (Takeda and Manaka, 2018). From these experimentally derived k values, the profile of k was estimated, assuming a linear relationship of logk - z (Kurikami et al., 2008). Fig. 7a shows that the estimated intrinsic permeability ranges from 1.3 × 10−20 to 7.0 × 10−19 (m2), which corresponds approximately to the intrinsic 6
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Fig. 7. Vertical profiles of (a) intrinsic permeability determined from the porosity and nominal half distance between clay surfaces using Eq. (9) and measured in chemical osmosis experiments performed by Takeda and Manaka (2018), and potential geologic osmotic pressures estimated from steady-state models with bottom boundary conditions of (b) constant pressure equivalent to hydrostatic pressure and (c) no flux.
have equilibrated with hydrostatic fluid pressure in the open portion of the borehole above the packer. Consequently, the excess pressures might have been measured at lower values than those present prior to drilling. This effect should be ascertained by using a steady-state analysis that describes the three-dimensional configuration of fluid measurement and permeability distribution within the formation for further study.
comprehensively explain the fluid overpressures observed in this region, and therefore the calibration of permeability distributions (e.g., Suko et al., 2014) or the assumption of a caprock (Karasaki et al., 2011) have been used to reproduce the observed fluid pressures in groundwater flow modeling. Now that the potential for geologic osmotic pressure in the Wakkanai formation is indicated, incorporating its effect into conventional modeling may improve the misfit between observed and predicted fluid pressures in this region. Also, the local fluid pressure disparity observed in our analyses may be improved by using simulations that also account for topography-driven flow.
5. Conclusions The quantitative analysis presented herein reveals that the fluid overpressures observed in the Wakkanai Formation, which cannot be explained by gravity-driven flow due to topographic effects (Karasaki et al., 2011) or pressure-driven flow due to dynamic geologic processes (Neuzil, 2015), may be explained as geologic osmotic pressures. However, further research is needed to ascertain whether geologic osmotic pressures are the primary cause of the observed fluid overpressure in the Wakkanai Formation in this region. The effects of topographydriven lateral flow should be also addressed to improve further detailed studies. Also, scrutinizing measured fluid pressures using numerical modeling and detailed semipermeability modeling that considers varying groundwater compositions and additional semipermeability measurements may improve the misfit between the calculated and observed values. Generally, geologic osmotic pressure is relatively small in comparison to fluid overpressures caused by other geologic processes such as tectonic loading or compaction disequilibrium (e.g., Osborne and Swarbrick, 1997; Gonçalvès et al., 2010); however, it could be a dominant process that affects groundwater flow and associated solute transport over geologically long timescales in stable argillaceous
4.3. Errors in fluid pressure measurement Except the natural process described above, artificial errors in fluid pressure measurements can result in a disparity between observed and estimated values. Errors in borehole measurements can be introduced by insufficient recovery of fluid pressures after drilling and/or by the short-circuiting of fluid from a measurement interval to the open hole above or below the measurement interval (e.g., Quinn et al., 2012). Although the fluid pressure measurements in the SAB-2 borehole were performed during the course of drilling, the measurements were made after the fluid pressure in the packed off interval reached equilibrium (Suko et al., 2014). Therefore, the possibility of insufficient pressure recovery can be ruled out. However, short-circuiting of fluid through the surrounding formation to the open borehole can occur, particularly at shallow depths in the SAB-2 borehole where the observed overpressures are below our estimated pressures. Because the permeability of the shallow part of the Wakkanai formation is relatively high (Kurikami et al., 2008; Takeda et al., 2014; Takeda & Manaka, 2018), the pressure of the fluid isolated below the packer in the borehole might 7
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formations (e.g., Neuzil, 2000). Therefore, the identification of geologic osmotic pressure is also crucial for investigating nuclear waste isolation in argillaceous formation sites (Neuzil, 2000; Wilson et al., 2003; Neuzil and Provost, 2009). The approach used in this study can only be applied to laterally uniform argillaceous formations. However, an approach using profiles of salinity and basic rock physical properties can be conducted without initial and boundary conditions or transport parameters, such as semipermeability, permeability, and diffusivity, which are necessary for the type of conventional simulation that fully describe transport processes within an argillaceous formation. Accordingly, the approach used in this study may help identify potential geologic osmotic pressures in argillaceous formations, especially during the early stages of hydrogeologic investigations when adequate data for simulating transport may not be available.
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Declaration of Competing Interest The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper. Acknowledgments The main part of this research project was conducted as regulatorysupporting research funded by the Secretariat of the Nuclear Regulation Authority, Japan. The authors wish to thank the Japan Atomic Energy Agency for providing mudstone samples and Kazuya Miyakawa and Eiichi Ishii for fruitful comments and discussions. References Amano, Y., Yamamoto, Y., Nanjyo, I., Murakami, H., Yokota, H., Yamazaki, M., et al., 2012. Data of Groundwater from Boreholes, River Water and Precipitation for the Horonobe Underground Research Laboratory Project (2001-2010) (JAEA-Data/Code 2011-023). Japan Atomic Energy Agency (JAEA), Ibaraki, Japan. Appelo, C.A.J., Postma, D., 2005. Geochemistry, Groundwater and Pollution, 2nd ed. Balkema Publishers, Netherlands. Bader, S., Kooi, H., 2005. Modelling of solute and water transport in semi-permeable clay membranes: comparison with experiments. Adv. Water Resour. 28, 203–214. https:// doi.org/10.1016/j.advwatres.2004.11.001. Bresler, E., 1970. Numerical solution of the equation for interacting diffuse layers in mixed ionic systems with nonsymmetrical electrolytes. J. Colloid Interface Sci. 33, 278–283. https://doi.org/10.1016/0021-9797(70)90030-5. Bresler, E., 1973. Anion exclusion and coupling effects in nonsteady transport through unsaturated soils: I. Theory. Soil Sci. Soc. Am. Proc. 37 (5), 663–669. https://doi.org/ 10.2136/sssaj1973.03615995003700050013x. Brunauer, S., Emmett, P.H., Teller, E., 1938. Adsorption of gases in multimolecular layers. J. Am. Chem. Soc. 60 (2), 309–319. https://doi.org/10.1021/ja01269a023. Cey, B.D., Barbour, S.L., Hendry, M.J., 2001. Osmotic flow through a cretaceous clay in southern Saskatchewan, Canada. Canad. Geotech. J. 38 (5), 1025–1033. https://doi. org/10.1139/t01-033. Cruchaudet, M., Croisé, J., Lavanchy, J.M., 2008. In situ osmotic experiment in the Callovo-Oxfordian argillaceous formation at the Meuse/Haute-Marne URL (France): Data and analysis. Phys. Chem. Earth. 33, S114–S124. https://doi.org/10.1016/j.pce. 2008.10.045. Fritz, S.J., 1986. Ideality of clay membranes in osmotic processes: a review. Clays Clay Miner. 34 (2), 214–223. https://doi.org/10.1346/CCMN.1986.0340212. Garavito, A.M., De Cannière, P., Kooi, H., 2007. In situ chemical osmosis experiment in the Boom Clay at the Mol underground research laboratory. Phys. Chem. Earth. 32 (1–7), 421–433. https://doi.org/10.1016/j.pce.2006.01.004. Gonçalvès, J., Violette, S., Wendling, J., 2004. Analytical and numerical solutions for alternative overpressuring processes: application to the Callovo-Oxfordian sedimentary in the Paris basin, France. J. Geophys. Res. 109, B02110. https://doi.org/10. 1029/2002JB002278. Gonçalvès, J., Rousseau-Gueutin, P., Revil, A., 2007. Introducing interacting diffuse layers in TLM calculations: a reappraisal of the influence of the pore size on the swelling pressure and the osmotic efficiency of compacted bentonites. J. Colloid Interface Sci. 316, 92–99. https://doi.org/10.1016/j.jcis.2007.07.023. Gonçalvès, J., Rousseau-Gueutin, P., de Marsily, G., Cosenza, P., Violette, S., 2010. What is the significance of pore pressure in a saturated shale layer? Water Resour. Res. 46, W04514. https://doi.org/10.1029/2009WR008090. Greenberg, J.A., Mitchell, J.K., Witherspoon, P.A., 1973. Coupled salt and water flows in a groundwater basin. J. Geophys. Res. 78 (27), 6341–6353. https://doi.org/10.1029/ JC078i027p06341. Gueutin, P., Altmann, S., Gonçalvès, J., Cosenza, P., Violette, S., 2007. Osmotic interpretation of overpressures from monovalent based triple layer model, in the Callovo-
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