Natural analogue monitoring to estimate the hydrochemical change of groundwater by the carbonating process from the introduction of CO2

Accepted Manuscript Research papers Natural Analogue Monitoring to Estimate the Hydrochemical Change of Groundwater by the Carbonating Process from the Introduction of CO2 Hanna Choi, Nam Chil Woo PII: DOI: Reference:

S0022-1694(18)30306-8 https://doi.org/10.1016/j.jhydrol.2018.04.053 HYDROL 22756

To appear in:

Journal of Hydrology

Received Date: Revised Date: Accepted Date:

17 January 2018 19 April 2018 22 April 2018

Please cite this article as: Choi, H., Woo, N.C., Natural Analogue Monitoring to Estimate the Hydrochemical Change of Groundwater by the Carbonating Process from the Introduction of CO2, Journal of Hydrology (2018), doi: https:// doi.org/10.1016/j.jhydrol.2018.04.053

This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.

Author Information of the Manuscript Ref #: HYDROL_22756 1st Author: Name: Hanna Choi Affiliations: 1

Department of Earth System Sciences, Yonsei University 50 Yonsei-ro, Sudaemoon-gu, Seoul, Korea

2

School of Earth and Environmental Sciences, Seoul National University 1, Gwanak-ro, Gwanak-gu, Seoul, Korea

e-mail: [email protected] Fax: 82-2-873-3647 2nd Author: Name: Nam Chil Woo Affiliations: 1

Department of Earth System Sciences, Yonsei University 50 Yonsei-ro, Sudaemoon-gu, Seoul, Korea

e-mail: [email protected] Fax: 82-2-2123-8169

Corresponding Author: Nam Chil Woo,

e-mail:

[email protected]

Order of Author:

Hanna Choi, Nam Chil Woo

1. Introduction

Among various carbon capture and sequestration (CCS) technologies, the geologic carbon storage (GCS) is accepted as one of the most feasible techniques. The GCS is the technique injecting CO2 into inland or coastal geologic medium in depths of about 800 m to 1,000 m below surface level (b.s.l) (Metz et al, 2005; Benson and Cole, 2008; DOE, 2011). The stored CO2 is supposed to be retained in reservoirs for a long time (more than 99% of the injected CO2 will be maintained for more than 1,000 years) and the acceptable leakage level has been suggested as less than 0.1% of storage volume per year (Harvey et al., 2012; Song and Zhang, 2012; Miocic et al., 2013). However, when the CO2 leakage from reservoir is above the acceptable level, the water quality of the upper aquifer can be degraded due to the accelerated water-rock interaction by the intruded CO2 gas (Kharaka et al., 2009; Zoback and Gorelick, 2012). Moreover, once the excess CO2 plume is leaked into the atmosphere through the preferential flow path such as fault or fracture, it could result in human inhalation problems depending on the concentration and exposure time (Schütze et al., 2012; Gasparini et al., 2016). Therefore, the GCS sites should be monitored intensively and regularly to track the movement of the injected CO2 plume in reservoirs and to detect the hydrochemical changes in the aquifer above the cap rock (Terwel et al., 2011; Wallquist et al., 2012; Liebscher et al., 2013).

If the excess CO2 plumes intrude into the aquifer above the cap rock, the pressure will most likely rise, or the hydrochemistry will change regardless of the masking effect of the other geologic matrix (Lewicki et al., 2007; Vishal and Singh, 2016). Hovorka et al. (2013) measured pressure change to detect the CO2 leakage with a baseline monitoring in the deep aquifer which was called as ‘above-zone monitoring interval (AZMI)’. The AZMI section placed in the 400 m to 700 m below surface level (b.s.l) has the advantage of receiving minimal impact from artificial factors and facilitate an immediate response to physicochemical reactions occurring in the event of CO2 leakage from reservoirs (Kharaka et al., 2009; Song and Zhang, 2012; Tao et al., 2013; Jenkins et al., 2015). Consequently, monitoring and identifying the changes in the physicochemical conditions of the AZMI section will be an effective option for early leakage detection at the CCS sites. However, in the stage of developing GCS facility without any direct injection and monitoring practice, understanding of potential reactions of leaked CO2 with groundwater in various natural conditions is indispensable to predict and evaluate the risk. For example, the natural analogue studies on geothermal groundwater (hot spring) and carbonated groundwater (soda spring) could provide information on the background condition of the AZMI and CO2 intruded aquifers (Wikinson et al., 2009; Kaszuba et al., 2011; Becker and Lynds, 2012; Gal et al., 2012; Schütze et al., 2012; Do et al., 2017). Accordingly, several studies have been performed to monitor the groundwater systems with stable conditions and to predict the GCS impact by the CO2 injection. Wikinson et al. (2009) reported geochemical changes of groundwater from the analogue site near the sea using the variation of ion species and stable C and O isotopes. Gilfillan et al., (2011) monitored He, Ne gases and C isotope in the natural CO2 rich spring and groundwater to trace the CO2 source where deeper mantle origin or shallower GCS reservoir origin. Gal et al. (2012) monitored gases such as CO2, 222Rn and 4He from the soil and water in the natural analogue site to figure out spatial influence of gas leakage under the ground and to suggest mixing processes as leakage from GCS site. This study will focus on the hydrogeochemical data of geothermal groundwater (TW), carbonated groundwater (CW), and shallow non-carbonated groundwater (GW) to construe the natural data as a part of the carbonating process. Fig. 1 shows a schematic diagram related to the two possible pathways

of the CW development from fresh groundwater: One is the deep reservoir located in the AZMI section (PATH 1) and the other is the shallow reservoir located in the 50 m to 250 m b.s.l (PATH 2).

2. Geological background Fig. 2 shows the location and the geology of the monitoring sites, and their detailed descriptions are as follows (Kim and Choi, 1998).

2.1. Geology of geothermal groundwater sites

This study conducted the TW monitoring from two geologic groups. One group is Suanbo and Yuseong sites in the Okcheon metamorphic belt region and the other group is Deoksan and Seokmodo sites in the Mesozoic granite region (Fig. 2 (a) to (d)). The depth of the TW wells ranged from 300 m to 700 m b.s.l, similar to the AZMI depth of the GCS. These sites are known for high geothermal gradient, which originated from the decay energy of radioisotopes such as U, Th, K, etc. (Lee et al., 1998; Kim, 2007). The Suanbo geothermal site is located on the Seochangni Formation of the Okcheon metamorphic belt, alternately composed of dark gray colored phyllite and black shale (Kim and Lee, 1965). The Yuseong site consists mainly of Jurassic two-mica granite and schistose granite of Hyangsan-ri Formation intruded quartz porphyry and quartz veins (Park et al., 1977). The Deoksan site is mainly consisted of Jurassic Daebo granite over Gayabong and Sangwangsan area (Shin et al., 1989). The Seokmodo site has bedrocks of Precambrian metamorphic rock and Mesozoic biotite granite intruded hornblende granodiorite along the southwest coast and also TW monitoring well is located on the biotite granite region (Hwang and Kihm, 2005). Suanbo, Yuseong and Deoksan sites are located in the inland area whereas Seokmodo site is located in the coastal area within 500 m from the beachfront.

2.2. Geology of carbonated groundwater sites

The CW in the study sites are generally produced from the fault zone along the geological boundaries (Fig. 2 (e) to (h)). The bedrocks of Chojeong site consist of age-unknown arenaceous phyllite layer and metasedimentary rocks of the Okcheon metamorphic belt (Lee and Kim, 1971). The Bugang site shows bedrocks of age-unknown two-mica schist crossing to the southwest direction and Jurassic granites (Kwon and Jin, 1974). Jurassic granites are composed of porphyritic granite and biotite granite with various directions of foliation. The Shin and Bangadari sites are located close to each other – approximately 2 km in distance – and have the same geologic background of the Precambrian gneiss intruded by age-unknown gneiss, Jurassic sedimentary layer, and Mesozoic granite (Cheong et al., 1975; Jeong et al., 2005).

3. Methods

3.1. Sampling and analysis for hydrochemical elements

This study conducted seasonal monitoring on the TW, GW, and the surface water (SW) of the Suanbo, Yuseong, Deoksan and Seokmodo sites and the CW and the SW in the Chojeong, Bugang, Shin and Bangadari sites from July 2013 to May 2015. It should be noted that the Chojeong site coincided with the GW collection. Table 1 shows brief geographic and hydrogeologic information of the study sites since the hydrogeochemical data will give information on inherent aquifer characteristics at different depths.

At the sampling sites, the in-situ measurements including temperature, pH, electric conductivity (EC), and redox potential (Eh) were taken. Laboratory analysis was performed on cations (Ca, K, Mg, Na, Si, Sr, Rb), anions (F, Cl, Br, NO3, SO4, HCO3) and stable isotopes (13C,

18

O and 2H). Alkalinity

titration was carried out using 0.01N hydrochloric acid and Metrohm 796 Titroprocessor on the sampled day. All water samples were filtered through 0.45 µm membrane filters (Advantec, Mixed cellulose ester), separately collected in polypropylene bottles for cation, anion, alkalinity, oxygen-hydrogen isotopes, and kept in the refrigerator under 4℃ for chemical analyses. The Cation samples were acidified with concentrated nitric acid to lower the pH to less than 2. For stable carbon isotope sampling, target water was injected into a 10-4 torr vacuum glass bottle (100 mL volume bottle filled with 10 mL phosphoric acid) by using a pre-cleaned syringe with 0.45 μm pore filter for CO2 gas production (Capasso et al., 2005). The radiocarbon samples were collected through 1L HDPE bottles using flow cells to prevent mixing with atmospheric carbon gases. Then, the samples were treated with 10 mL of NaOH solution and 2 mL of BaCl2 solution in an anaerobic chamber to obtain DIC (Dissolved Inorganic Carbon) precipitation in BaCO3 form. The NaOH and the BaCl2 solutions were purged by reacting with N2 gas one day before the field experimentation in order to eliminate the dissolved CO2 gas. The cations and anions were analyzed using the Inductively Coupled Plasma Optical Emission Spectrometer (ICP-OES, OPTIMA 8300, Perkin Elmer) and the Ion Chromatography (IC, Metrohm, 883 Basic IC plus) at the Hydrogeology laboratory at Yonsei University. The oxygen (δ18O) and the hydrogen (δ2H) isotopes of the water samples were analyzed using the Stable Isotope Ratio Mass Spectrometry (S-IRMS, GV instruments, Limited) with ±0.1 ‰ and ±1.0 ‰ precision and then data reported relative to the standardized Standard Mean Ocean Water (SMOW) at the Korea Basic Science Institute in Ochang. The stable carbon (δ13C) isotope of the DIC was analyzed using the IRMS (Delta V Plus, Thermo Scientific, Germany) with ±0.1 ‰ precision relative to the Pee Dee Belemnite (PDB) reference at the Korea Basic Science Institute in Ochang. The BaCO3 samples for radiocarbon dating (Δ14C) were combusted to transform graphite by means of an elemental analyzer with Fe catalyst and hydrogen gas at 600℃. These samples were measured at the KIGAM (Korea Institute of Geosciences and Mineral Resources) through the Accelerator Mass Spectrometer (AMS) system with pelletized graphite samples.

3.2. Simulating hydrochemical variation of groundwater on the CO2 intrusion

To evaluate potential changes in water chemistry when groundwater reacts with dissolved CO2 under various pressure and temperature conditions, this study simulated different hydrochemical variations using a geochemical inverse model – the PHREEQC version 3 (Parkhurst and Appelo, 2013) which calculates the solubility of gases in the high-pressure condition (Peng and Robinson, 1976). The simulations follow equilibrium reactions among CO2water-mineral phases according to the temperature, pressure and groundwater composition. The aquifer pressure through the geologic matrix was related to the pressure condition of groundwater system, as increasing 10 bar per 100 m b.s.l (Perrin et al., 2009; Wertz et al., 2013). The unit conversion was conducted for the pressure unit of 1 bar in order to calculate with 1 atm (about 1.01 bar). Fig. 3 shows the experimental temperature and calculated pressure conditions of the deep aquifer (50~71℃, 32~76 atm) as a PATH 1 and the shallow aquifer (13~22℃, 6~25 atm) as a PATH 2 which are suggested in Fig. 1. For the modeled condition, this study set the minimum and the maximum concentrations of CO2 from 0.0001 to 10 mol L-1 intruded into the noncarbonated reservoir. These models could provide the analyses on reservoir depth of carbonated groundwater as well as the chemical changes of noncarbonated groundwater by CO2 intrusion by using the PHREEQC with ‘EQUILIBRIUM PHASES’ and ‘REACTION’ calculation on each site.

3.3. Estimating carbonate reservoir using geothermometer Some dissolved elements are sensitive to temperature for precipitation or dissolution during water-rock interaction. Accordingly, a geothermometer was used to estimate the ion species from the reservoir temperature of the formed groundwater and there are two assumptions for the geothermometer

calculations. The first assumption is that temperature indicating elements have been formed through the hydrochemical equilibrium at the reservoir and the second assumption is that no mixing or dilution occurred when the groundwater rose to the surface (Arnórsson, 1975; Giggenbach, 1988; Choi, 2003). Even though the reservoir temperature estimation of CW by geothermometer still poses a challenge, many studies have attempted to do geothermometer application. Thomas et al. (2016) applied the chalcedony silica geothermometer to estimate the subsurface temperature in volcanic areas of Iceland with temperatures of 110℃ and higher. Choi (2003) studied the CO2-rich groundwater of Gangwon Province in Korea to calculate the carbonated reservoir temperature of the Na-HCO3 type groundwater ranging between 140 and 160℃. Based on the thermodynamic calculation and empirical formula from the Islandic geothermal groundwater (Fournier, 1977; Arnórsson et al., 1983), the silica geothermometer could be applied to low-temperature reservoir estimation under 200℃ (Edmiston and Benoit, 1984). At the Islandic volcanic rock region, chalcedony and quartz maintains equilibrium with groundwater at temperatures 20 to 110℃ and above 180℃ (Arnórsson, 1975). The chalcedony geothermometer suggested by Fournier (1977) can be applied 0 to 250℃ temperature range and temperature equation is the following:

Estimated reservoir temperature (℃) = [1032/(4.69-logS)] - 273.15

(1)

where, S means SiO2 concentration (mg/kg). Michard et al. (1986) reported that the groundwater ranging between 70 and 140℃ reaches the equilibrium state with both quartz and chalcedony at a granite rock region. Meanwhile, Kharaka et al. (1977) reported that in a sedimentary rock region, groundwater below 70℃ equilibrates with chalcedony, and that above 70℃ equilibrates with quartz. Thus, this study applied the chalcedony geothermometer to simulate the changes of the carbonated groundwater at the TW sites.

4. Results and discussion

4.1. Field measurement data

Table A1 specified the groundwater and surface water data of the TW sites. The average temperatures of the TW of Deoksan and Seokmodo remained steady at 48.5 and 71℃, respectively. Meanwhile, the Suanbo TW retained a consistent average temperature of 53.0℃ (SA-1 to SA-6, see Table A1) but showed a large difference of approximately 6℃ in SA-7 despite Eh uniformity. It is considered that GW (SAS-5, temperature 12℃) could be mixed from the maintenance construction in the nearby stream. The Suanbo, Yuseong and Deoksan TW showed weak alkaline properties with pH 7.91~8.88. The pH range of coastal TW and Seokmodo varies from 6.44 to 6.11, which classify as weakly acidic to neutral. The Seokmodo TW was reported as sea-water intruded TW (Lee et al., 2011) and it can also be confirmed by the total dissolved solids (TDS) value of measured field data, approximately 21,000 to 23,000 mg/L. The Seokmodo TW showed lower alkalinity of 35.25~36.89 mg/L despite being affected by sea-water intrusion (alkalinity 113.11 mg/L as CaCO3). These results indicate that the TW was influenced relatively more by GW (alkalinity 18.85 mg/L as CaCO3) and surface water than sea-water during the mixing process. Redox potential (Eh) of the Seokmodo TW ranged between 88.7 and 102.1 mV, implying reducing condition (Nordstrom and Wilde, 2005). In the Suanbo site and the Yuseong site, the transitional conditions of 255.1~276.6 mV and 247.3 mV indicated the mixing of surface water.

Table A2 represents the groundwater and surface water data of the CW sites. Eh of the CW samples ranged from 263.7 to 488.7 mV implying oxidizing condition. This oxidizing condition can be interpreted as the electron pair exchange reactions between the dissolved oxygen within the CW and mineral of the bedrock (McBride, 1994; Stumm and Morgan, 1996). Based on the findings by previous studies in Korea (Koh et al., 2002; Choi, 2003; Jeong et al., 2005; Koh et al., 2008), this study regarded the samples over -0.5 Log PCO2 as carbonated groundwater and those below -0.7 Log PCO2 as non-carbonated groundwater. The CO2 partial pressure for the CJ-1 sample was -1.28, which was especially low compared with other carbonated waters showing partial pressures from -0.03 to -0.49. The unusually low CO2 partial pressure was attributed to the mixing of shallow non-carbonated groundwater resulted from excessive pumping, which in turn induced pH increase and CO2 partial pressure decrease (Choi, 2003; Choi et al., 2005). Log PCO2 in Table A1 and Table A2 present the values calculated using the PHREEQC version 3.3. Temperatures of the Bugang CW site and the Bangadari CW site were found to be relatively consistent. However, the Sh-4 and Sh-5 samples from the Shin spring show especially low temperature that the sample was potentially exposed to the atmosphere during the sampling process. The EC values for the CW samples ranged from 130 μS/cm to 1600 μS/cm and these observations are higher than those of the GW and the SW at each site. The EC differences could be originating from acidic pH of the dissolved CO2, which accelerates chemical reactions with bedrock minerals (Stumm and Morgan, 1996; Clark and Fritz, 1997).

Dissolved inorganic carbon (DIC) ⇄ CO2 (aq) + H2CO30 + HCO3- + CO32-

(2)

H2CO30 + CO2 (aq) ⇄ H2CO3* ⇄ H+ + HCO3-

(3)

Reaction (2) shows the main components of DIC in carbonated groundwater. Superscript * refers to the dissolved CO2 which is comprised of 99.85% of aqueous (aq) CO2 phase and 0.15% of H2CO3 phase.

4.2. Hydrochemical characteristics of geothermal and carbonated groundwaters

4.2.1. Hydrochemical evolution

Fig. 4 presents the GW plots and the SW plots of the TW and the CW sites. Major ion compositions of all samples are presented in Table A3 and Table A4. In Fig. 4(a), the TW samples have Na-HCO3 type (as seen inside black solid lines) and identified as alkaline water with prevailing HCO3 (Ravikumar et al., 2015). The GW samples (red plots in Fig. 4(a)) show similar chemical properties to SW as earth alkaline water with prevailing HCO3, which may infer that permanent stream, SW, consists of dominant GW contribution (Fetter, 2000). The dotted circle located on the upper right of the Piper diagram (Fig. 4 (a)) indicates the Seokmodo samples from coastal areas. They show a markedly higher Na-Cl ratio compared to other water bodies. This means that anion ratios for the Seokmodo samples are maintained consistently, at approximately 90:5:5 ratios for Cl: SO4: HCO3 whereas the Na ratio in sea-water was approximately 80% and showed a high concentration of cations. On the contrary, the Ca ratios have an increasing trend from approximately 45% to 60% in the ascending order of geothermal groundwater, surface water, and shallow groundwater. The CW samples have normal earth alkaline water with prevailing HCO3 type. Especially, the Bugang, Shin, and Bangadari sites were categorized as Ca-HCO3 type maintaining a consistent chemical composition regardless of seasonal climate changes (Fig. 4 (b) and Table A4). On the other hand, the

carbonated groundwater in the Chojeong site was categorized as Ca-Na-HCO3 type. The plots inside the black dotted circle in Fig. 4 (b) are for the Sh-6 and BA-6 samples showing the large increase of Na-ratio but the decrease of Ca- and Mg- ratio. This possibly resulted from the influence of summer tourists getting into the surface water next to the sampling site.

4.2.2. Sea-water intrusion in coastal area

Cl and Br ion concentrations are generally low in groundwater and these elements are strongly conservative. Accordingly, these ions have been used as the indicators of sea-water intrusion or anthropogenic influence (Hem, 1991; Fetter, 2000). Fig. 5 (a) shows the Br/Cl ratio of the Seokmodo samples with the reference line for the sea-water ratio of sea-water, 0.003 (drawn by dotted line). The fact that all samples were plotted close to the reference line indicated that sea-water mixing is a dominant process. They may be categorized into two groups depending on the mixing ratio. One is the SW and GW group (Type I) and the other is the TW group (Type II), which is classified based on the degree of mixing dependence. Fig. 5 (b) reveals the Na/Cl ratio variation due to ion exchange via the intruding sea-water. Recent sea-water intrusion would be expected to decrease the Na/Cl ratio of the advancing front. In Fig. 5 (b) and (c), the Type II water group is situated between the dashed lines of sea-water and the Type I group but the dotted line, indicating mixing, shows a curved formation that implies the influence of cation exchange from water-rock interactions. The chemical equations for the reactions are as follows (Kim et al., 2003):

2Na+ (or K+) + Ca-X2 ⇄ 2 Na+ (or K+)-X + Ca2+

(4)

2Na+ (or K+) + Mg-X2 ⇄ 2 Na+ (or K+)-X + Mg2+

(5)

where, X refers to ion exchange sites in the aquifer matrix. In Fig. 5 (d), Type I and Type II are clearly separated by chemical properties. The general Na/Cl molar ratio for sea-water is known as 0.85 (Richter and Kreitler, 1993; Bear et al., 1999). Meanwhile, the molar ratio of the Seokmodo samples were calculated as 0.70 for sea-water, 0.54 for geothermal groundwater, 0.38 for surface water, and 0.29 for shallow groundwater (Table A3). Thus, it seems to be affected by the mixing with fresh water and by the exchange reactions between Na ions and (Ca + Mg) ions within bedrock minerals (Mondal et al., 2010). Fig. 5 (d) shows that shallow groundwater has high (Ca+Mg)/SO4 molar ratio compared to sea-water. This is because halite (NaCl) deposits are often mixed with gypsum (CaSO4•2H2O) and anhydrite (CaSO4) beds (Richter and Kreitler, 1993). Mg2+ is usually replaced in the position of Ca2+, which is why these minerals frequently exhibit admixtures of magnesium sulfates. Therefore, the molar ratios of (Ca+Mg)/SO4 versus Na/Cl might be able to estimate sea-water and fresh water interaction.

4.2.3. Isotopes related to groundwater evolution

Stable isotopes (18O, 2H and 13C) of TW The oxygen-hydrogen isotope composition from the South Korean precipitation at the IAEA/WMO station (Pohang) presented the Local Meteoric Water Line (LMWL) as δ2H 8.06δ18O+12.87 (Lee and Kim, 2007). The Global Meteoric Water Line (GMWL) proposed by Rozanski et al. (1993) was δ2H = 8.13δ18O+10.8 (Fig. 6). All water samples, except those from coastal sites, are plotted from -7 to -11‰ for δ18O and -50 to -75‰ for δ2H (Fig. 6), indicating the water bodies of the deep reservoir and the shallow reservoir GW of the study sites originated from meteoric water. In Fig. 6 (a) and (d),

oxygen-hydrogen isotopes values cluster in a relatively small range for both TW and CW samples, reflecting their unique water-rock reaction characteristics. By referencing the altitude data on Table 1 and Fig. 6 (b), isotopes become deplete with the altitude rising of the geothermal groundwater sites, indicating that precipitation infiltrated into the deep aquifers without significant isotope fractionations. On the other hand, the Seokmodo isotopes equation in Fig. 6 (a) is calculated δ2H = 6.13δ18O - 6.77 (R2 = 0.95) and its slope is lower than that of the LMWL with 8.06. The d-value (d = δ2H - 8δ18O) less than 10‰ indicates an evaporative enrichment of river water (Lee and Kim, 2007). The low correlation of Seokmodo samples with the LMWL is attributed to vapor mass formed continuously and early stages of precipitation were reflected (Kim, 2010). In the δ2H- δ18O diagram, the Seokmodo water samples aligned to the sea-water value and the TW (δ18O = -4.39 to -4.50‰ and δ2H = -30.5 to 35.6‰) were more enriched than the GW (Table A5) because deeper groundwater wells are more affected by the intrusion of isotopically heavy sea-water (Fetter, 2000). 13

C analysis can be utilized to distinguish respiration and metabolic processes of microbe, reactions with atmosphere and carbonates, and determining

for the inflow of deep origin CO2 (Maher et al., 2015; Kim, 2010; Becker et al., 2012). During the study period, δ13C values of the Suanbo, Deoksan, and Seokmodo TW samples (Table A5) remained at –11.93±0.31‰, -17.40±0.26‰, -15.73±0.68‰ respectively. The δ13C values of these sites including Yuseong (δ13C = -18.0‰) are implied by the impacts of freshwater carbonates and organic materials (Clark and Fritz, 1997; Hoefs, 2009).

Radioactive carbon isotope (14C) of TW In order to estimate the residence time of the hot spring sites, this study carried out 14C analysis on the TW and GW samples. The NETPATH adjustment modeling (Plummer et al., 1994) was conducted to correct the AMS results simulating 14C residence time of the total dissolved carbon (TDC) based on 14C pMC values. Time decay and carbon isotope mass-balance equation are as follows (Plummer et al., 1994; Suckow et al., 2013):

Δt (years) = (5730/ln 2) ln(And/A)

(6)

ΔRmDIC =

(7)

R*dIi –

Rαi-sdOi

The modern half-life of a 14C is 5730 years, And is the initial 14C content (pMC value) after adjustment for geochemical reactions except radioactive decay and A is the measured 14C content in the sample. In the equation (7), R is the carbon isotope ratio, mDIC the molality ratio of DIC, I and O incoming and outgoing carbon likewise dissolution and precipitation, respectively, superscript * to the ith carbon-bearing phases, and αi-s the fractionation factor between the ith phase and the aqueous solution. The SA-7, YS-1, DS-7, SM-7, SAG, YSG, DSG and SMG samples collected in April 2015 were used for radiocarbon dating. The deviation between the AMS measured value and the adjusted NETPATH value showed relatively small differences of approximately 100 years. Following calibrations, the age of water bodies in each study area ranged as follows: Seokmodo (6,400 years), Suanbo (5,580 years), Yuseong (1,980 years), and Deoksan (3,240 years) BP (before present, present means 1950s) (Table A5). The GW and the TW have a relationship of depth to residence time with the following formula (Edmands and Smedly, 2000):

pMC (percent modern carbon) = - 0.06 Depth + 98 (R2 = 0.93)

(8)

It was interpreted that the longer flow path of meteoric water would result in the development of high-temperature geothermal groundwater. On the other hand, the Deoksan shallow groundwater has a (-) value, indicating that mixing with present water such as surface water (Godwin, 1962; Fetter, 2000;

Suckow et al., 2013).

Stable isotopes (18O, 2H and 13C) of CW The oxygen-hydrogen isotopes of the CW samples are clustered in the right side of the LMWL (Fig. 6 (c) and (d)). This phenomenon was interpreted to have resulted from two factors. First, the dissolved CO2 might have accelerated water-rock interactions between groundwater and bedrock minerals. Thus, the enriched oxygen isotopes within the minerals could have leached into the groundwater. Second, the carbonated groundwater with high CO2 partial pressure could have been exposed to the surface environment, which in turn would have undergone degassing process. The relatively depleted oxygen isotopes would have diffused into the atmosphere first, which in turn enriched groundwater isotopes (Koh et al., 2002; Kim, 2010). There are possible provenances of CO2 in groundwater with three types: mantle degassing, biogenic decomposition in the soil, and metamorphic devolatilization of carbonate minerals (Jeong et al., 2005; Koh et al., 2008). Many researchers have studied the origins of CO2 in the carbonated groundwater by using various geochemical tools. Based on δ13C and Log PCO2 partial pressure calculations (Table A2 and Table A5), the CO2 provenance in these study sites were regarded to be from a deep mantle origin (Koh et al., 2002; Jeong et al., 2005; Koh et al., 2008). Fig. 7 (a) shows CW plots with constant δ18O value but δ13C is distributed at -3 to -10‰ range. It was considered that the former reflects site-specific characteristics owing to geologic minerals and the latter reflects evidence of mantle origin CO2 mixture (Clark and Fritz, 1997). Fig. 7 (b) presents consistent total dissolved solids (TDS) concentration for each site and CO2 partial pressure (Log PCO2) spreading -0.03 to -0.49 range. The TDS value is generally controlled by water-rock interaction such as geology, flow velocity, and residence time (Güler and Thyne, 2004). Meanwhile, the CJ-1 sample (in the dotted circle) of Fig. 7 (b) is lower than the reference pCO2 of CW (Log pCO2 over -0.50) but it does not show significant TDS and δ18O differences. This phenomenon seems to be related to the reflection of site specific characteristics implied through the TDS concentration and the implication of

meteoric water origin through δ18O data (Kim, 2007). Average pH and TDS of CW samples have proportional relations in Fig. 7 (c). The Bugang samples had the highest TDS (1,614 mg/L) and Chojeong the lowest TDS (329 mg/L) but pH was (6.12 and 5.6, respectively) showed approximately 0.5 order difference. The cause is ascribed to weak acidic pH (5.09 ~ 6.72) of the CW samples not providing enough acceleration for the TDS concentration increase through water-rock interaction.

4.3. Carbonated reservoir estimation to predict CO2 leakage process The carbonated groundwater had reached three-component equilibrium by undergoing hydrochemical reactions with the groundwater body, bedrock minerals, and dissolved CO2. Therefore, the PHREEQC model with the geologic consideration can reflect the gas-water-rock interactions in the actual groundwater system when CO2 leakage occurs from the GCS sites.

4.3.1. Temperature estimation of CW reservoir using geothermometer

The chalcedony geothermometer can be applied to groundwater of granite area and sedimentary rock region to estimate the reservoir temperature which was equilibrated with SiO2 species (Kharaka et al., 1977; Michard et al., 1986). Calculated temperatures using equation (1), which is the chalcedony geothermometer, are presented in Table 2. These results can be further applied to estimate reservoir depth by multiplying average temperature and geothermal gradient values. Especially, calculated temperature range distribute from 43.08 to 83.13℃, which are in accordance with the geothermal groundwater of these study sites. Blavoux et al. (1982) are reported the δ13C values of CO2 from deep origin such as mantle and magma having -8.0 to 4.0‰ range. In comparison with this, CW samples present the δ13C range between -8.0 and -3.2‰ (Table 4). Based on the δ13C and geothermometer

results, the CW sites of this study can be assumed to evolve from the TW reservoir by mixing of mantle CO2 gases. Geothermometer is applied to the geothermal groundwater in the volcanic area rather than carbonated groundwater, due to the silicate minerals highly sensitive to temperature (Arnórsson, 1975; Giggenbach, 1988; Choi, 2003). When carbonated groundwater is compared with thermal water, excessive SiO2 can be released from bedrock minerals after the dissolution of CO2. Thus, calculated reservoir depths using geothermometer can be slightly overestimated. On the other hand, average temperatures of the CW samples are distributed at a range of 8.77 to 17.55℃, which is expected from a CW reservoir at a shallower depth than a TW aquifer. In addition, Koh and Chae (2008) reported that the mean residence time (MRT) of the soda springs in Chungcheong Province – the southern part of Korea – had turned out to have a shorter residence time than the TW samples, of 15 to 50 years. The tritium analysis study by Vuataz and Goff (1986) suggested the cold springs of Jemez Mountain – the northern part of New Mexico – to have 20 to 75 years of MRT despite consisting of meteoric water. They estimated the short MRT because of mixing among deep thermal fluids, near-surface groundwater, and relatively old but cold groundwater. Therefore, there is another possibility in the mixing process such as GW and CW mixing but there are many complex factors to consider in binary mixing, such as mixing ratio or end member decision. For that reason, this study focuses on the CW evolution at the deep reservoirs (300 m to 700 m as PATH 1) and the shallow reservoirs (below 300 m as PATH 2) while excluding the mixing of GW.

4.3.2. PHREEQC simulations of hydrochemical change with CO2 mixing

The carbonation modeling of the TW and GW samples at hot spring sites gives the information on the hydrochemical changes of deep reservoirs and shallow reservoirs during the carbonation processes. In comparison with the hot spring concept, the GW at the Chojeong site can be used to predict the changes of the shallow reservoirs located directly under the soda spring site. The groundwater velocity was assumed to be slow enough (less than Reynold

number 10) to apply the Darcy’s laws, and there is sufficient time between the groundwater and the introduced CO2 to achieve chemical equilibrium similar to the closed batch condition. The abovementioned geothermometer calculations were applied to the PHREEQC simulation (Parkhurst and Appelo, 2013) and Geochemist’s workbench® student edition (Bethke and Yeakel, 2009). This modeling process conjectured that different pressure and temperature conditions accommodate different CO2 solubility for each condition. 0.0001 to 10 mol L-1 of CO2 was introduced to simulate the changes of the two water groups, which are the TW with higher temperature and pressure (50~71℃, 32~76 atm) and the GW with lower temperature and pressure (13~22℃, 6 ~25 atm). The GW of Chojeong varied from Ca-Mg-HCO3 to Ca-HCO3 type (Fig. 8 (a)). In comparison, the water properties of the TW of the inland sites (Suanbo, Yuseong and Deoksan) changed from alkali-bicarbonate type (Na-HCO3) to earth alkaline with prevailing bicarbonate type (Ca-HCO3) due to CO2 intrusion (Fig. 8 (b), (d), (f)). CO2 partial pressure is calculated to be below -0.5 when 0.01 mol L-1 to 0.1 mol L-1 of CO2 is intruded into the reservoir but if 0.1 mol L-1 of CO2 flow into the groundwater, carbonating process sufficiently appears in all TW sites with Ca-HCO3 type. The GW samples of these sites increase in the Ca and HCO3 proportion from Ca-Na-HCO3 type (Fig. 8 (c), (e), (g)). The introduced CO2 concentration increases following the direction of the dotted arrow. The GW of Suanbo, Yuseong and Deoksan show carbonated water conditions under 0.01 mol L-1 of CO2 mixing, which is lower than TW transition point. It can be accounted for the fact that TW samples have higher alkalinities and higher buffer capacity (Table A1). The interactive formula for the buffer capacity is as follows (Stumm and Morgan, 1996):

CO2 (aq) + H2O ⇄ H2CO3 (aq) ↔ H+ + HCO3- (induce pH decreasing)

(9)

Dissolved HCO3- (groundwater) + released H+ ⇄ H2CO3 (aq) (pH buffer)

(10)

The TW and GW samples of the Seokmodo coastal site have high Na-Cl and Ca-SO4 ratios, as earth alkaline water bodies have high concentrations of SO4 or Cl. The TW is changed to earth alkaline water with HCO3 type but the GW move to alkaline water with HCO3 type due to differences in cation behavior (Fig. 8 (h) and (i)). The piper diagram shows only major ionic ratios; thus, the determination of carbonated water has to be checked using HCO3 ion concentration in relation to CO2 moles. Fig. 9 shows the changes of the dissolved ion concentration in accordance with the CO2 intrusion per each site. The Chojeong GW simulation showed that Si ion and SO4 ion increase but also Ca ion decreases (Fig. 9 (a)). In the inland TW, Ca and HCO3 ion concentrations increase but SO4 ion decreases when the moles of CO2 increase (Fig. 9 (b), (d), (f)). Likewise, the inland GW increases Ca, SO4 and HCO3 ion concentration but Na ions and K ions decrease along with CO2 concentration (Fig. 9 (c), (e), (g)). It is inferred that CO2 mixing accelerates a catalytic water-rock interaction. Also, when the mixing reaction occurs, secondary minerals can be precipitated into the reservoir. Coastal TW starts to change water quality with 0.01 mol L-1 of CO2, which is higher concentration compared with the inland TW samples (Fig. 9 (h) and (i)). It is considered that high salinity of various ion species dissolved in the coastal groundwater is crucial factor to restrain the change of water quality.

5. Conclusion

To discover the details of the leaked CO2-plume induced carbonating process of groundwater with the various depth, this natural analogue study was carried out at the geothermal groundwater (TW) sites and the carbonated groundwater (CW) sites. The TW samples show consistent water type of NaHCO3 in inland areas (Suanbo, Yuseong and Deoksan) and Na-Cl in a coastal area (Seokmodo). The δ2H-δ18O data of the inland TW indicated that the meteoric water has different isotopic values depending on the altitude infiltrated into the thermal aquifer. The Δ14C data of the TW samples collected from

deep aquifer revealed long residence times. The CW samples showed Ca-HCO3 type regardless of seasonal variations but the CW of Chojeong has been found to be partially mixed with shallow groundwater (GW). The dissolved inorganic carbon of the CW appeals to be affected not by dissolution of carbonate minerals but by the mantle degassing process owing to the δ13C range. Based on the hydrochemical analysis, the study sites of the TW and the CW were considered to be adequately closed systems. Thus, it is reasonable to consider them as the batch conditions for the simulation of the carbonating processes using the PHREEQC model. The chalcedony geothermometer was considered suitable for application on groundwater equilibrated with SiO2 in the granite area and the sedimentary bedrock region, such as the ones found on the study sites. The reservoir temperatures of the CW sites have been estimated to range between 43.08 and 83.13℃, which are similar temperature range with the TW sites. In addition to these results, reservoir depths (km) can estimate reservoir temperature (℃) divide by geothermal gradient (℃ per km). Also, the calculated reservoir depths of the study sites indicate the geothermal groundwater flowed out from the above zone monitoring interval (AZMI), which depth is 400 m to 700 m below surface level (b.s.l)). On the other hand, CW samples were monitored to be in 3.8 to 18.7℃ range and 263.7 to 488.7 mV of redox potential which mean shallow reservoirs can be taken into account in the GW carbonating process. Therefore, two possibilities can be considered to explain the carbonated reservoir of CW samples. One is the deeper reservoirs (300-700 m b.s.l) from the chalcedony geothermometer and geothermal gradient calculation and the other is the shallower reservoirs (below 300 m b.s.l) from the temperature and redox potential data. In a bid to simulate the carbonating processes of deep reservoirs and shallow reservoirs using the PHREEQC model, pCO2 of groundwater data exceeds -0.5 in the inland TW and GW more than 0.1 mol L-1 and 0.01 mol L-1 of CO2, respectively. During the CO2 mixing, Ca and HCO3 ions increase but SO4 decreases in inland TW samples and Ca, SO4 and HCO3 increase in inland GW samples, at the same time, Na and K ions decrease. The coastal TW can be reflecting the carbonating process of the coastal AZMI section containing large quantities of salts. The carbonation of these samples is triggered by 0.01

mol L-1 of CO2 introduced into reservoir but the Na-Cl water type was kept. The anion ratios of coastal water have changed due to HCO3 increase but cations were still consisted of high Na and K ratio during simulation. The concentrated salts in the coastal groundwater are considered to increase the buffer capacity to respond to the CO2 mixing. This study focuses on both deep reservoirs and shallow reservoirs, which identically have a Ca-HCO3 type with carbonated condition, despite that they have undergone different ionic variations. Furthermore, if the same amount of CO2 is mixed, GW will be reached through a faster carbonated condition. Based on all GW data, temperature range of discharge water from TW reservoir was thought to coincide with in-situ CW samples due to the uplift-cooling. In order to hypothesize a more realistic carbonation condition, selection of mixing components and mixing ratio should be considered in future studies.

Acknowledgements This research was supported by the Energy Efficiency & Resources Core Technology Program of the Korea Institute of Energy Technology Evaluation and Planning (KETEP) (No. 2012T100100395), and Basic Science Research Program by the National Research Foundation of Korea (NRF) funded by the Ministry of Education (NRF-2017R1A6A3A01008897). The authors deeply thank Dr. Yoon, Yoon-Yeol, Dr. Park, Junghun and Dr. Hong, Wan of Korea Institute of Geoscience and Mineral Resources (KIGAM). This paper is a part of Ms. Hanna Choi’s Ph.D. thesis research at Yonsei University.

References Arnórsson, S., Gunnlaugsson, E., Svavarsson, H., 1983. The chemistry of geothermal waters in Iceland. II. Mineral equilibria and independent variables controlling water compositions. Geochimica et Cosmochimica Acta, 47(3): 547-566. Arnorsson, S., 1975. Application of the silica geothermometer in low temperature hydrothermal areas in Iceland. Am. J. Sci.;(United States), 275(7). Bear, J., Cheng, A.H.-D., Sorek, S., Ouazar, D., Herrera, I., 1999. Seawater intrusion in coastal aquifers: concepts, methods and practices, 14. Springer Science & Business Media. Becker, T.P., Lynds, R., 2012. A geologic deconstruction of one of the world's largest natural accumulations of CO2, Moxa arch, southwestern Wyoming. AAPG bulletin, 96(9): 1643-1664. Benson, S.M., Cole, D.R., 2008. CO 2 sequestration in deep sedimentary formations. Elements, 4(5): 325-331. Bethke, C., Yeakel, S., 2009. Geochemist's Workbench: Release 8.0 Reference Manual. RockWare Incorporated. Blavoux, B., Dazy, J., Sarrot-Reynauld, J. 1982. Information about the origin of thermomineral waters and gas by means of environmental isotopes in eastern Azerbaijan, Iran, and southeast France. Journal of hydrology, 56(1-2): 23-38. Capasso, G., Favara, R., Grassa, F., Inguaggiato, S., Longo, L., 2005. On-line technique for preparingand measuring stable carbon isotopeof total dissolved inorganic carbonin water samples (d13CTDIC). Annals of Geophysics. Cheong, C.H., Won, C.K., Cha, M.S., Kang, K.W., Lee, Y.C., 1975. Explanatory text of the geological map of Odaesan sheet (1:50,000), Geological and Mineral institute of Korea. Choi, H.S., 2002. Hydrogeochemical and environmental isotope studies of CO2-rich groundwaters in the Kangwon Province, Korea: Water–rock interaction, origin and evolution. Unpub, Ph. D. Thesis, Korea Univ., Seoul, Republic of Korea, 208p.

Choi, H.S., Koh, Y.K., Bae, D.S., Park, S.S., Hutcheon, I., Yun, S.T., 2005. Estimation of deep-reservoir temperature of CO2-rich springs in Kangwon district, South Korea. Journal of volcanology and geothermal research, 141(1): 77-89. Clark, I.D., Fritz, P., 1997. Environmental isotopes in hydrogeology. CRC press. Dansgaard, W., 1964. Stable isotopes in precipitation. Tellus, 16(4): 436-468. Do, H.K., Kim, K.H., Yun, S.T., 2017. Hydrochemical and Isotopic Characteristics of CO2-rich Groundwater in the Gyeongsang Sedimentary Basin, South Korea: A Natural Analogue Study on the Potential Leakage of Geologically-stored CO2. Energy Procedia, 114: 3805-3811. DOE, U.S., 2011. Report of the interagency task force on carbon capture and storage, U.S. Department of Energy, Washington DC. Edmiston, R.C., Benoit, W.R., 1984. Characteristics of basin and range geothermal systems with fluid temperatures of 150℃ to 200℃. Geothermal Resources Council Transactions, 8: 417-424. Fetter, C.W., 2000. Applied hydrogeology. Prentice hall. Fournier, R.O., 1977. Chemical geothermometers and mixing models for geothermal systems. Geothermics, 5(1-4): 41-50. Gal, F., Brach, M., Braibant, G., Bény, C., Michel, K., 2012. What can be learned from natural analogue studies in view of CO2 leakage issues in Carbon Capture and Storage applications? Geochemical case study of Sainte-Marguerite area (French Massif Central). International Journal of Greenhouse Gas Control, 10: 470-485. Gasparini, A., Sainz-García, A., Grandia, F., Bruno, J., 2016. Atmospheric dispersion modelling of a natural CO2 degassing pool from Campo de Calatrava (northeast Spain) natural analogue. Implications for carbon storage risk assessment. International Journal of Greenhouse Gas Control, 47: 38-47. Giggenbach, W.F., 1988. Geothermal solute equilibria. derivation of Na-K-Mg-Ca geoindicators. Geochimica et cosmochimica acta, 52(12): 2749-2765. Gilfillan, S. M. V., Wilkinson, M., Haszeldine, R. S., Shipton, Z. K., Nelson, S. T., Poreda, R. J., 2011. He and Ne as tracers of natural CO2 migration up a

fault from a deep reservoir. International Journal of Greenhouse Gas Control, 5(6): 1507-1516. Godwin, H., 1962. Half-life of radiocarbon. Nature, 195(4845): 984. Güler, C., Thyne, G.D., 2004. Hydrologic and geologic factors controlling surface and groundwater chemistry in Indian Wells-Owens Valley area, southeastern California, USA. Journal of Hydrology, 285(1): 177-198. Harvey, O. R., Qafoku, N. P., Cantrell, K. J., Lee, G., Amonette, J. E., Brown, C. F., 2012. Geochemical implications of gas leakage associated with geologic CO2 storage-A qualitative review. Environmental science & technology, 47(1): 23-36. Hem, J.D., 1991. Study and interpretation of the chemical characteristics of natural water, 3rd edition, 2254. Department of the Interior, US Geological Survey. Hovorka, S. D., Meckel, T. A., Trevino, R. H., Lu, J., Nicot, J.-P., Choi, J.W., Freeman, D., Cook, P., Daley, T. M., Ajo-Franklin, J. B., 2011. Monitoring a large volume CO2 injection: year two results from SECARB project at Denbury’s Cranfield, Mississippi, USA. Energy Procedia, 4: 3478-3485. Hwang, J.H., Kihm, Y.H., 2005. Geological report of the Ganghwa·Onsuri sheet (1:50,000), Korea institute of geoscience and mineral resources, Daejeon. Jenkins, C., Chadwick, A., Hovorka, S.D., 2015. The state of the art in monitoring and verification—ten years on. International Journal of Greenhouse Gas Control, 40: 312-349. Jeong, C.H., Kim, H.J., Lee, S.Y., 2005. Hydrochemistry and genesis of CO2-rich springs from Mesozoic granitoids and their adjacent rocks in South Korea. Geochemical Journal, 39(6): 517-530. Kaszuba, J.P., Navarre-Sitchler, A., Thyne, G.D., Chopping, C., Meuzelaar, T., 2011. Supercritical carbon dioxide and sulfur in the Madison Limestone: A natural analog in southwest Wyoming for geologic carbon–sulfur co-sequestration. Earth and Planetary Science Letters, 309(1): 131-140. Kharaka, Y.K., Callender, E., Carothers, W.W., Meriwether, J., 1977. Geochemistry of geopressured geothermal waters from the Texas Gulf Coast.

Kharaka, Y. K., Thordsen, J. J., Hovorka, S. D., Nance, H. S., Cole, D. R., Phelps, T. J., Knauss, K. G., 2009. Potential environmental issues of CO2 storage in deep saline aquifers: geochemical results from the Frio-I Brine Pilot test, Texas, USA. Applied Geochemistry, 24(6): 1106-1112. Kim, K.H., 2007. Hot springs in Korea. Ewha Womans University Press, Seoul. (in Korean). Kim, K.H., 2010. Isotope geochemistry. Sigmapress, Seoul. (in Korean). Kim, K.H., Choi, H.J., 1998. A geochemical study on the thermal water and groundwater in the hot spring area, South Korea. Journal of the Korean earth science society, 19(1): 22-34. (in Korean with English abstract). Kim, K.W., Lee, H.K., 1965. Geological map of Chungju Sheet (1:50,000), Geological survey of Korea. Kim, Y.J., Lee, K.S., Koh, D.C., Lee, D.H., Lee, S.G., Park, W.B., Koh, G.W., Woo, N.C., 2003. Hydrogeochemical and isotopic evidence of groundwater salinization in a coastal aquifer: a case study in Jeju volcanic island, Korea. Journal of Hydrology, 270(3): 282-294. Koh, D.C., Chae, G.K., 2008. Estimation of mixing properties and mean residence time using 3H for groundwater in typical geothermal and CO2-rich areas in South Korea. Journal of Geological Society of Korea, 44(4): 507-522. (in Korean with English abstract). Koh, Y.K., Choi, B.Y., Yun, S.T., Choi, H.S., Mayer, B., Ryoo, S.W., 2008. Origin and evolution of two contrasting thermal groundwaters (CO2-rich and alkaline) in the Jungwon area, South Korea: Hydrochemical and isotopic evidence. Journal of Volcanology and Geothermal Research, 178(4): 777786. Koh, Y.K., Kim, C.S., Bae, D.S., Han, K.W., 2002. The isotopic and chemical compositions of the CO2-rich waters in Korea. Geofísica Internacional, 41(4): 491-498. Kramer, J.H., Dohms, P., 1997. ASTM Standards on Ground Water and Vadose Zone Investigations: Drilling, Sampling, Well Installation and Abandonment Procedures. Eos, Transactions American Geophysical Union, 78(3): 23-23.

Kwon, Y.I., Jin, M.S., 1974. Explanatory text of the geological map of Cheongju sheet (1:50,000), Geological and Mineral institute of Korea. Lee, C.H., Kim, J.H., 1971. Explanatory text of the geological map of Jeungpyeong sheet (1:50,000), Geological survey of Korea. Lee, J.S., Chon, H.T., Kim, J.S., Kim, K.W., Lee, H.K., Moon, H.S., 1998. Enrichment of potentially toxic elements in areas underlain by black shales and slates in Korea. Environmental Geochemistry and Health, 20(3): 135-147. Lee, K.S., Kim, Y.J., 2007. Determining the seasonality of groundwater recharge using water isotopes: a case study from the upper North Han River basin, Korea. Environmental Geology, 52(5): 853-859. Lee, S.G., Kim, T.K., Lee, T.J., 2011. Strontium isotope geochemistry and its geochemical implication from hot spring waters in South Korea. Journal of Volcanology and Geothermal Research, 208(1): 12-22. Lewicki, J.L., Oldenburg, C.M., Dobeck, L., Spangler, L., 2007. Surface CO2 leakage during two shallow subsurface CO2 releases. Geophysical Research Letters, 34(24). Liebscher, A., Möller, F., Bannach, A., Köhler, S., Wiebach, J. Schmidt-Hattenberger, C., Weiner, M., Pretschner, C., Ebert, K., Zemke, J., 2013. Injection operation and operational pressure–temperature monitoring at the CO2 storage pilot site Ketzin, Germany—Design, results, recommendations. International Journal of Greenhouse Gas Control, 15: 163-173 Maher, D.T., Cowley, K., Santos, I.R., Macklin, P., Eyre, B.D., 2015. Methane and carbon dioxide dynamics in a subtropical estuary over a diel cycle: Insights from automated in situ radioactive and stable isotope measurements. Marine Chemistry, 168: 69-79. McBride, M.B., 1994. Environmental chemistry of soils. Oxford University Press, New York. Metz, B., Davidson, O., De Coninck, H., Loos, M., Meyer, L., 2005. IPCC special report on carbon dioxide capture and storage, Intergovernmental Panel on Climate Change, Geneva (Switzerland). Working Group III.

Michard, G., Sanjuan, B., Criaud, A., Fouillac, C., Pentcheva, E. N., Petrov, P. S., Alexieva, R., 1986. Equilibria and geothermometry in hot alkaline waters from granites of SW Bulgaria. Geochemical Journal, 20(4): 159-171. Miocic, J.M., Gilfillan, S.M.V., McDermott, C., Haszeldine, R.S., 2013. Mechanisms for CO2 leakage prevention–A global dataset of natural analogues. Energy Procedia, 40: 320-328. Navarre-Sitchler, A.K., Maxwell, R.M., Siirila, E.R., Hammond, G.E., Lichtner, P.C., 2013. Elucidating geochemical response of shallow heterogeneous aquifers to CO2 leakage using high-performance computing: Implications for monitoring of CO2 sequestration. Advances in Water Resources, 53: 45-55. Nordstrom, D.K., Wilde, F.D., 2005. Reduction-6.5 oxidation potential (electrode method), US Geological Survey. Odeh, N.A., Cockerill, T.T., 2008. Life cycle GHG assessment of fossil fuel power plants with carbon capture and storage. Energy Policy, 36(1): 367-380. Oh, J.H., Kim, K.Y., Han, W.S., Kim, T.H., Kim, J.C., Park, E.G., 2013. Experimental and numerical study on supercritical CO2/brine transport in a fractured rock: Implications of mass transfer, capillary pressure and storage capacity. Advances in Water Resources, 62: 442-453. Park, H.I., Lee, J.D., Cheong, J.G., Min, K., 1977. Geological map of Yuseong Sheet (1:50,000), Korea research institute of geoscience and mineral Resources. Parkhurst, D.L., Appelo, C.A.J., 2013. Description of input and examples for PHREEQC version 3: a computer program for speciation, batch-reaction, one-dimensional transport, and inverse geochemical calculations. 2328-7055, US Geological Survey. Peng, D.Y., Robinson, D.B., 1976. A new two-constant equation of state. Industrial & Engineering Chemistry Fundamentals, 15(1): 59-64. Perrin, J.-C., Krause, M., Kuo, C.-W., Miljkovic, L., Charoba, E., Benson, S. M., 2009. Core-scale experimental study of relative permeability properties of CO2 and brine in reservoir rocks. Energy Procedia, 1(1): 3515-3522.

Plummer, L.N., Prestemon, E.C., Parckhurst, D.L., 1994. User’s guide to NETPATH-A computer program for mass-balance calculations: an interactive code (NETPATH) for modelling NET geochemical reactions along a flow PATH-Version 2.0. US Geological Survey, Water Resources Investigations Report: 94-4169. Ravikumar, P., Somashekar, R.K., Prakash, K.L., 2015. A comparative study on usage of Durov and Piper diagrams to interpret hydrochemical processes in groundwater from SRLIS river basin, Karnataka, India. Elixir International Journal, 80: 31073-31077. Richter, B.C., Kreitler, C.W., 1993. Geochemical techniques for identifying sources of ground-water salinization. CRC press, New York. Rozanski, K., Araguás‐ Araguás, L., Gonfiantini, R., 1993. Isotopic patterns in modern global precipitation. Climate change in continental isotopic records: 1-36. Saibi, H., Ehara, S., 2010. Temperature and chemical changes in the fluids of the Obama geothermal field (SW Japan) in response to field utilization. Geothermics, 39(3): 228-241. Schütze, C., Sauer, U., Beyer, K., Lamert, H., Bräuer, K., Strauch, G., Flechsig, C. H., Kämpf, H., Dietrich, P., 2012. Natural analogues: a potential approach for developing reliable monitoring methods to understand subsurface CO2 migration processes. Environmental Earth Sciences, 67(2): 411-423. Shin, B.W., So, C.S., Park, B.S., Lee, S.H., 1989. Geological report of the Haemi sheet (1:50,000), Korea institute of energy and resources. Song, J., Zhang, D., 2012. Comprehensive review of caprock-sealing mechanisms for geologic carbon sequestration. Environmental science & technology, 47(1): 9-22. Stumm, W., Morgan, J.J., 1996. Aquatic Chemistry: Chemical Equilibria and Rates in Natural Waters, 3rd edition, 126. John Wiley & Sons, New York. Suckow, A., Aggarwal, P.K., Araguas-Araguas, L., 2013. Isotope methods for dating old groundwater. International Atomic Energy Agency (IAEA),

Vienne. Tao, Q., Bryant, S.L., Meckel, T.A., 2013. Leakage fingerprints during storage: modeling above-zone measurements of pressure and temperature. Energy Procedia, 37: 4310-4316. Terwel, B.W., Harinck, F., Ellemers, N., Daamen, D.D.L., 2011. Going beyond the properties of CO2 capture and storage (CCS) technology: how trust in stakeholders affects public acceptance of CCS. International Journal of Greenhouse Gas Control, 5(2): 181-188 Thomas, D.L., Bird, D.K., Arnórsson, S., Maher, K., 2016. Geochemistry of CO2-rich waters in Iceland. Chemical Geology, 444: 158-179. Vishal, V., Singh, T.N., 2016. Geologic Carbon Sequestration: Understanding Reservoir Behavior. Springer. Vuataz, F.D., Goff, F., 1986. Isotope geochemistry of thermal and nonthermal waters in the Valles Caldera, Jemez Mountains, northern New Mexico. Journal of Geophysical Research: Solid Earth, 91(B2): 1835-1853. Wallquist, L., Seigo, S.L., Visschers, V.H.M., Siegrist, M., 2012. Public acceptance of CCS system elements: a conjoint measurement. International Journal of Greenhouse Gas Control, 6: 77-83. Wertz, F., Gherardi, F., Blanc, P., Bader, A.-G., Fabbri, A., 2013. Cement CO2-alteration propagation at the well–caprock–reservoir interface and influence of diffusion. International Journal of Greenhouse Gas Control, 12: 9-17. Wiese, B., Zimmer, M., Nowak, M., Pellizzari, L., Pilz, P., 2013. Well-based hydraulic and geochemical monitoring of the above zone of the CO2 reservoir at Ketzin, Germany. Environmental earth sciences, 70(8): 3709-3726. Wilkinson, M., Haszeldine, R. S., Fallick, A. E., Odling, N., Stoker, S. J., Gatliff, R. W., 2009. CO2–mineral reaction in a natural analogue for CO2 storage—implications for modeling. Journal of Sedimentary Research, 79(7): 486-494. Zoback, M.D., Gorelick, S.M., 2012. Earthquake triggering and large-scale geologic storage of carbon dioxide. Proceedings of the National Academy of

Sciences, 109(26): 10164-10168.

Fig. 1. Schematic diagram of the carbonating processes to estimate the carbonated reservoir depth. Deep (300 m to 700 m below surface level, bsl) and shallow (50 m to 250 m) aquifers with non-carbonated fresh water could be undergone different hydrochemical changes when CO2 plume intrudes into the aquifer.

Fig. 2. Sampling sites of geothermal groundwater (black circle) and carbonated groundwater (empty circle) in South Korea (modified from Kim and Choi, 1998). Geologic maps are for (a) Suanbo, (b) Yuseong, (c) Deoksan, (d) Seokmodo, (e) Chojeong, (f) Bugang, (g) Shin and (h) Bangadari sites, respectively (modified from Kim and Lee, 1965; Park et al., 1977; Shin et al., 1989; Hwang and Kihm, 2005; Lee and Kim, 1971; Kwon and Jin, 1974; Cheong et al., 1975, respectively). Scale unit in the geologic map is 1-km in length.

Fig. 3. Schematic diagram for PHREEQC simulation to estimate carbonating processes under various pressure and temperature conditions of deep and shallow reservoirs.

Fig. 4. Chemical water type of groundwater and surface water samples in (a) the hot spring sites and (b) the carbonated spring sites. Each site abbreviates as follows: Suanbo (SA), Yuseong (YS), Deoksan (DS), Seokmodo (SM) samples plotted on (a) and Chojeong (CJ), Bugang (BG), Shin (Sh), Bangadari (BA) samples plotted on (b). The water samples are classified as groundwater (G), surface water (S) and sea water (sea).

Fig. 5. Changes of water chemistry between the surface water and the groundwater of Seokmodo samples: Type I and II denote samples of the surface water and the shallow groundwater and those of the geothermal groundwater, respectively. Dashed lines of (a) and (d) indicate the references of sea water ratio, but those in (b) and (C) show the mixing lines of the sea water and the surface water. GW means the abbreviation of the groundwater.

Fig. 6. δ2H-δ18O isotope signatures from the water samples: (a) geothermal groundwater, (b) altitude effect in hot spring sites, (c) carbonated groundwater and (d) its enlarged diagram. LMWL and GMWL refer to Lee et al. (2003) and Rozanski et al. (1993), respectively.

Fig. 7. Relations among isotope signatures, TDS and pH carbonated groundwater samples: (a) δ18O-δ13C isotope signatures, (b) TDS-Log PCO2, and (c) average pH-TDS. Dotted circle in (b) shows CJ-1 sample as an outlier.

Fig. 8. Changes of water chemistry corresponding to the concentrations of CO2 leaked into the deep geothermal groundwater (TW) and shallow groundwater (GW) samples simulated using PHREEQC and Geochemist’s workbench programs with various pressure and temperature condition.

Fig. 9. Ionic variations of deep geothermal groundwater (TW) and shallow groundwater (GW) for the CO2 inflow with varying concentrations from 0.0001 to 10 mol L-1, simulated using the PHREEQC program.

Fig. A1. Saturation indices (SIs) of geothermal groundwater (TW) and shallow groundwater (GW) imply the stability of minerals following the injection of various CO2 concentrations from 0.0001 to 10 mol L-1.

Table 1. Geographic and hydrogeologic information on the sampling sites.

Table 2. Calculated reservoir temperatures and depths of carbonated groundwater based on the chalcedony geothermometer.

Table A1. Field measured data from the geothermal groundwater, nearby shallow (cold) groundwater and surface water including sea-water samples.

Table A2. Field measured data from carbonated groundwater, nearby shallow (non-carbonated) groundwater and surface water samples.

Table A3. Major ion analysis data gained from geothermal groundwater, nearby shallow groundwater and surface water (river and Seokmodo sea-water).

Table A4. Major ion compositions from carbonated groundwater, nearby non-carbonated shallow groundwater (Chojeong site only) and surface water (Bugang, Shin and Bangadari sites).

Table A5. Stable isotopes (13C, 18O and 2H) and radioactive isotope (14C) data of groundwater samples.

Fig. 1.

Fig. 2.

Fig. 3.

Fig. 4.

Fig. 5.

Fig. 6.

Fig. 7.

Fig. 8.

Fig. 9.

Appendix

Fig. A1.

Table 1. Water types

Study site Suanbo

Geothermal

Yuseong

groundwater

Deoksan Seokmodo Chojeong

Carbonated

Bugang

groundwater

Shin Bangadari

Shallow (cold) groundwater

Lithology Gneissic terrain

Granitic terrain

Gneissic terrain

Granitic terrain

Latitude

Longitude

Altitude (m) Drilling depth (m)

36° 50’ 48.83” N

127° 59’ 17.94” E

195

750

36° 21’ 21.51” N

127° 20’ 41.13” E

49

320

36° 41’ 49.18” N

126° 39’ 49.18” E

38

580

37° 40’ 17.35” N

126° 20’ 09.84” E

5

755

36° 43' 19.24'' N

127° 36' 07.03'' E

99

50

36° 31' 37.10'' N

127° 22' 43.51'' E

45

Artesian well

37° 40' 38.82'' N

128° 29' 54.50'' E

855

Artesian well

37° 41' 33.77'' N

128° 30' 17.12'' E

905

Artesian well

Chojeong*

Gneiss & granite

36° 43’ 21.20” N

127° 35’ 15.80” E

50

250

Suanbo

Limestone & granite

36° 50’ 18.00” N

128° 02’ 28.00” E

306

100

Yuseong

Gneiss & granite

36° 21’ 22.78” N

127° 20’ 52.55” E

49

150

36° 41’ 15.63” N

126° 39’ 03.03” E

47

100

37° 39’ 05.00” N

126° 20’ 00.60” E

8

60

Deoksan Seokmodo

Granite

* Shallow (cold) groundwater in Chojeong site is non-carbonated groundwater.

Table 2. Water type

Site

Chojeong

Bugang

Carbonated groundwater Shin

Bangadari

Sample CJ 1 CJ 2 CJ 3 CJ 4 BG 1 BG 2 BG 3 BG 4 BG 5 BG 6 Sh 1 Sh 2 Sh 3 Sh 4 Sh 5 Sh 6 BA 1 BA 2 BA 3 BA 4 BA 5 BA 6

In-situ temperature (℃) 18.7 18.3 16.9 16.3 15.2 15.9 15.3 14.2 14.7 15.5 11.6 13.5 10.2 3.8 6.3 11.4 8.9 10.0 9.3 7.7 7.8 8.9

Estimated* temperature (℃) 57.64 55.75 49.10 50.02 89.58 61.03 83.31 81.83 90.94 92.08 46.96 50.78 23.13 40.18 49.41 48.01 52.57 56.49 31.76 42.70 53.65 51.71

Average temperature (℃)

Geothermal gradient** (℃/km)

Estimated depth (Km)

53.13

26.98

1.97

83.13

21.33

3.90

43.08

32.13

1.34

48.15

32.13

1.50

NonCJG-1 16.2 34.58 carbonated Chojeong 40.29 26.98 1.49 CJG-2 16.6 46.00 groundwater * Chalcedony geothermometer: estimated reservoir temperature (℃) = [1032/(4.69-logS)] - 273.15, S means SiO2 concentration (mg/kg) ** Geothermal gradient data are derived from New and Renewable Energy Data Center, http://kredc.kier.re.kr.

47

Table A1. Site

Sample

Sampling date

Temp. (℃)

pH

EC (µS/cm)

Alkalinity (mg/L as CaCO3)

Eh (mV)

Log PCO2*

Geothermal groundwater SA-1 Jul. 2013 52.0 8.60 367.0 104.0 -3.37 SA-2 Aug. 2013 52.0 8.31 373.0 123.8 -2.97 SA-3 Oct. 2013 54.5 8.35 386.0 118.0 -3.01 Suanbo SA-4 Jan. 2014 53.0 7.91 379.0 119.7 -2.56 SA-5 Apr. 2014 53.0 8.30 554.0 118.9 255.1 -2.96 SA-6 Jul. 2014 53.5 8.19 577.0 119.7 270.9 -2.84 SA-7 Apr. 2015 47.4 8.04 477.0 115.6 276.6 -2.75 Yuseong YS-1 May. 2015 53 8.14 326.0 93.44 247.3 -2.90 DS-1 Jul. 2013 47.5 8.88 271.5 85.00 -3.79 DS-2 Aug. 2013 47.9 8.83 272.7 104.9 -3.63 DS-3 Oct. 2013 48.3 8.84 269.8 89.34 -3.71 Deoksan DS-4 Jan. 2014 49.0 8.80 301.0 94.26 -3.64 DS-5 Apr. 2014 49.0 8.77 326.0 91.80 313.4 -3.61 DS-6 Jul. 2014 49.5 8.36 355.0 92.62 164.3 -3.15 DS-7 Apr. 2015 48.0 8.50 397.0 95.08 245.1 -3.31 SM-1 Jul. 2013 71.0 6.82 47,000 40.00 -1.96 SM-2 Aug. 2013 72.0 6.61 40,200 36.89 -1.74 SM-3 Oct. 2013 72.0 6.70 45,400 37.70 -1.84 Seokmodo SM-4 Jan. 2014 70.0 6.63 46,080 35.25 -1.80 SM-5 Apr. 2014 70.5 6.70 38,500 34.43 88.7 -1.90 SM-6 Jul. 2014 71.0 6.44 62,550 35.25 97.8 -1.55 SM-7 Apr. 2015 71.0 6.35 41,360 35.25 102.1 -1.45 Shallow (cold) groundwater Suanbo SAG Apr. 2015 12.5 7.27 140.1 27.87 425.3 -1.63 Yuseong YSG May. 2015 21.5 7.12 494 145.9 300.0 -1.58 Deoksan DSG Apr. 2015 14.7 6.56 288.1 30.33 465.9 -0.87 Seokmodo SMG Apr. 2015 13.0 6.47 5,810 18.85 532.1 -0.82 Surface water around the geothermal groundwater well Suanbo SAS-1 Oct. 2013 15.4 9.16 231.3 91.80 -4.27 SAS-2 Jan. 2014 1.5 6.67 140 80.33 -1.80 SAS-3 Apr. 2014 14.6 8.16 196 86.07 320.9 -3.21 SAS-4 Jul. 2014 24.9 8.56 261.5 104.9 283.7 -3.49 SAS-5 Apr. 2015 12.0 7.37 58.9 13.11 422.8 -3.23 Yuseong YSS-1 Apr. 2015 17.9 6.97 308 57.38 444.7 -2.17 YSS-2 May. 2015 22.0 6.76 532 64.75 452.0 -1.88 Deoksan DSS-1 Aug. 2013 31.0 7.46 381 51.00 -2.62 DSS-2 Oct. 2013 15.1 7.05 115.5 28.69 -2.55 DSS-3 Jan. 2014 4.8 7.78 79.7 41.80 -3.17 DSS-4 Jul. 2014 27.3 7.50 159.8 46.72 357.1 -2.72 DSS-5 Apr. 2015 12.7 7.28 163.5 24.59 375.0 -2.86 Seokmodo SMS-1 Aug. 2013 30.4 7.85 626 21.31 -3.42 SMS-2 Oct. 2013 12.7 7.11 1243 41.80 -2.50 SMS-3 Apr. 2014 15.3 7.39 1321 38.52 312.0 -2.81 SMS-4 Apr. 2015 19.5 7.00 4440 22.95 349.4 -2.65 Seokmodo sea-water Seokmodo SMsea Apr. 2015 20.4 7.17 53160 113.11 434.7 -2.22 * Log PCO2 calculated from physicochemical data using PHREEQC program (ver. 3.3; Parkhurst and Appelo, 2013).

48

Table A2. Temp. Sampling EC Alkalinity Eh pH Log PCO2* date (℃) (µS/cm) (mg/L as CaCO3) (mV) Carbonated groundwater CJ-1 Jul. 2013 18.7 6.72 493 256.0 -1.28 CJ-2 Aug. 2013 18.3 5.51 407 136.1 -0.34 Chojeong CJ-3 Oct. 2013 16.9 5.09 363 104.9 -0.03 CJ-4 Jan. 2014 16.3 5.09 276 103.3 -0.04 BG-1 Jul. 2013 15.2 6.20 1613 940.0 -0.25 BG-2 Aug. 2013 15.9 6.18 1764 967.2 -0.20 BG-3 Oct. 2013 15.3 6.11 1732 990.2 -0.14 Bugang BG-4 Jan. 2014 14.2 6.24 1214 936.1 -0.30 BG-5 Apr. 2014 14.7 6.04 1256 929.5 263.7 -0.10 BG-6 Jul. 2014 15.5 5.96 1250 934.4 488.7 -0.10 Sh-1 Jul. 2013 11.6 6.15 1074 580.0 -0.41 Sh-2 Aug. 2013 13.5 5.98 1114 623.0 -0.20 Sh-3 Oct. 2013 10.2 6.00 1098 604.9 -0.24 Shin Sh-4 Jan. 2014 3.8 6.09 752 619.7 -0.36 Sh-5 Apr. 2014 6.3 6.24 675 629.5 314.4 -0.49 Sh-6 Jul. 2014 11.4 5.99 760 588.5 484.0 -0.25 BA-1 Jul. 2013 8.9 5.95 884 360.0 -0.43 BA-2 Aug. 2013 10.0 5.86 842 477.0 -0.34 BA-3 Oct. 2013 9.3 5.81 859 472.1 -0.16 Bangadari BA-4 Jan. 2014 7.7 5.80 561 470.5 -0.17 BA-5 Apr. 2014 7.8 5.92 534 480.3 332.5 -0.28 BA-6 Jul. 2014 8.9 5.74 549 452.5 348.9 -0.12 Non-carbonated groundwater CJG-1 Apr. 2014 16.2 6.30 291 89.30 508.9 -1.32 Chojeong CJG-2 Jul. 2014 16.6 6.28 270 78.70 408.9 -1.35 Surface water around the carbonated spring BGS-1 Aug. 2013 31.2 8.03 344 104.1 -2.90 BGS-2 Oct. 2013 15.0 7.61 490 112.3 -2.54 Bugang BGS-3 Jan. 2014 4.8 7.42 210 109.8 -2.41 BGS-4 Jul. 2014 25.9 7.47 416 124.6 431.2 -2.29 ShS-1 Aug. 2013 18.8 7.10 59 23.00 -2.67 ShS-2 Oct. 2013 8.2 7.20 61 23.80 -2.82 Shin ShS-3 Apr. 2014 6.7 6.68 48 19.70 364.8 -2.38 ShS-4 Jul. 2014 17.2 6.95 99 30.30 482.8 -2.41 BAS-1 Aug. 2013 16.6 7.06 34 13.10 -2.89 BAS-2 Oct. 2013 8.6 7.51 36 12.30 -3.41 BangaBAS-3 Jan. 2014 0.7 5.97 21 12.30 -1.90 dari BAS-4 Apr. 2014 5.8 7.06 23 8.20 421.4 -3.15 BAS-5 Jul. 2014 14.9 6.26 44 14.8 364.0 -2.04 * Log PCO2 calculated from physicochemical data using PHREEQC program (ver. 3.3; Parkhurst and Appelo, 2013). Site

Sample

49

Table A3. Cl

NO3

SO4

22.7 14.5 22.6 14.9 14.0 14.1 12.7 16.6 23.6 14.4 23.3 14.3 13.7 13.4 13.0 13,500 13,300 13,400 13,400 12,800 12,400 12,500

0.5> 0.5> 0.5> 0.5> 0.5> 0.5> 0.5> 1.84 0.5> 0.5> 0.5> 0.5> 0.5> 0.72 0.5> 0.5> 0.5> 0.5> 0.5> 0.5> 0.5> 0.5>

27.4 30.4 30.1 27.0 34.6 25.6 26.1 17.7 17.3 15.8 20.3 16.1 22.5 15.6 15.5 1,010 1,070 1,140 1,060 1,000 1,010 1,090

127 151 144 146 145 146 141 114 104 128 109 115 112 113 116 48.8 45.0 46.0 43.0 42.0 43.0 43.0

11.0 8.24 7.57 7.07 9.64 7.33 8.39 8.12 3.38 3.44 3.07 2.81 4.84 2.79 3.90 3,580 3,400 3,230 3,250 3,440 3,420 3,150

Suanbo SAG 1.03 4.84 6.34 4.30 Yuseong YSG 0.99 25.2 6.67 26.7 Deoksan DSG 0.12 28.6 110 5.17 0.05> Seokmodo SMG 1740 10.1 112 Surface water around the geothermal groundwater well

34.0 178 37.0 23.0

Site

F

Sample

HCO3

Ca K unit: mg/L

Mg

Na

SiO2

TDS

3.46 2.68 1.96 1.90 1.32 2.12 2.02 0.96 1.84 1.84 1.39 1.39 2.42 1.88 1.33 138 146 82.0 80.0 95.0 102 57.0

0.1 > 0.1 > 0.1 > 0.1 > 0.1 > 0.1 > 0.1 > 0.1 > 0.1 > 0.1 > 0.1 > 0.1 > 0.1 > 0.1 > 0.1 > 263 253 271 269 293 299 247

83.1 65.9 69.2 70.4 48.1 49.3 79.1 62.6 49.1 49.2 53.2 55.9 36.6 42.8 61.3 4,690 4,430 4,580 4,630 4,910 4,910 3,880

105 73.8 79.6 81.0 67.4 71.5 70.0 45.8 55.0 53.2 56.1 58.4 46.9 49.2 47.8 30.1 57.4 62.6 47.0

386 354 362 355 327 323 347 273 258 269 270 267 242 243 262 23,200 22,700 22,700 22,700 22,600 22,200 21,000

11.6 36.6 32.7 514

0.26 0.61 1.16 7.66

0.73 0.30 5.93 102

4.87 36.0 23.8 327

23.3 33.2 27.1 21.8

91.3 344 272 2,860

38.6 37.0 38.5 41.5 8.20 24.7 26.3 12.9 11.1 11.5 13.7 12.2 44.9 87.8 118 368

1.74 1.34 1.89 1.61 0.33 6.86 12.3 2.73 2.06 1.32 2.56 1.18 4.13 5.17 4.43 7.24

5.08 4.98 4.57 3.94 1.26 4.53 5.44 2.57 2.55 2.63 1.73 2.56 5.85 12.5 16.3 43.7

6.02 7.03 5.38 9.81 3.28 20.0 30.3 8.75 8.97 10.8 12.6 9.91 55.3 109 120 320

15.6 11.2 12.6 15.7 9.77 5.35 6.16 13.0 16.6 14.7 12.6 11.8 12.5 9.33 1.48 1.36

217 207 187 236 61.0 218 295 117 105 109 128 107 346 623 810 2,320

Geothermal groundwater Suanbo

SA-1 7.33 SA-2 7.56 SA-3 7.16 SA-4 7.20 SA-5 7.32 SA-6 7.21 SA-7 7.54 Yuseong YS-1 5.14 Deoksan DS-1 3.68 DS-2 3.48 DS-3 3.58 DS-4 3.37 DS-5 3.46 DS-6 3.27 DS-7 3.58 Seokmodo SM-1 0.05> SM-2 0.05> SM-3 0.05> SM-4 0.05> SM-5 0.05> SM-6 0.05> SM-7 0.05> Shallow (cold) groundwater

Suanbo

Yuseong Deoksan

Seokmodo

SAS-1 SAS-2 SAS-3 SAS-4 SAS-5 YSS-1 YSS-2 DSS-1 DSS-2 DSS-3 DSS-4 DSS-5 SMS-1 SMS-2 SMS-3 SMS-4

0.29 0.28 2.45 0.31 0.19 0.24 0.25 0.13 0.12 0.10 0.16 0.18 0.30 0.55 0.05> 0.05>

8.88 12.4 4.67 11.5 6.23 37.6 66.2 10.6 10.6 14.9 14.5 12.6 181 320 468 1,460

12.3 17.5 0.5 > 7.91 9.88 16.8 29.6 6.24 10.5 14.1 5.01 19.7 0.5 > 0.5 > 0.5 > 0.5 >

16.7 17.2 11.5 15.3 5.81 31.9 39.6 9.09 7.31 6.27 8.47 7.24 16.3 28.0 35.1 94.0

112 98.0 105 128 16.0 70.0 79.0 51.0 35.0 33.0 57.0 30.0 26.0 51.0 47.0 28.0

Seokmodo sea-water Seokmodo SMsea 0.05> 17,400 0.5 > 2,400 138 297 260 910 7,890 8.79 29,300 * Seokmodo samples are diluted 400 times for geothermal groundwater, 40 times for shallow groundwater, 10 times for surface water and 500 times for sea water, and thus low-concentration ions such as F and NO3 were not detected using IC analysis.

50

Table A4. Site

Sample

F

Carbonated groundwater CJ-1 1.03 CJ-2 0.47 Chojeong CJ-3 0.67 CJ-4 0.37

Cl

NO3

SO4 HCO3

Ca

K unit: mg/L

Mg

Na

29.7 23.0 29.6 23.4

15.8 17.2 18.7 19.1

24.9 23.9 27.2 23.8

312 166 128 126

55.6 41.6 33.8 34.7

3.59 1.75 1.50 1.46

11.7 9.30 8.84 8.93

31.8 31.7 27.8 29.2

SiO2

Fe

TDS

37.2 35.7 30.7 31.4

0.05 > 0.41 0.05 > 0.05 >

Bugang

BG-1 BG-2 BG-3 BG-4 BG-5 BG-6

3.96 3.46 3.80 2.78 3.35 3.29

28.7 4.74 27.1 4.76 4.53 4.43

0.5 > 0.5 > 0.5 > 0.5 > 0.5 > 0.5 >

14.7 9.36 20.4 9.54 11.2 9.31

1,147 1,180 1,208 1,142 1,134 1,140

294.1 153.2 282.9 293.8 296.4 306.4

6.04 4.07 2.99 2.72 2.60 3.33

15.8 8.99 18.9 18.7 18.8 18.5

73.9 39.2 71.3 71.1 48.6 62.9

70.0 40.0 62.4 60.6 71.7 73.2

3.50 13.6 14.3 14.0 14.9 15.3

Shin

Sh-1 Sh-2 Sh-3 Sh-4 Sh-5 Sh-6

0.99 0.60 0.73 0.56 0.53 0.59

15.1 4.21 8.77 4.23 3.98 3.90

2.85 0.5 > 0.5 > 0.5 > 0.5 > 0.5 >

8.78 6.34 8.51 6.07 7.41 6.33

708 760 738 756 768 718

164.5 163.6 177.8 162.2 168.0 159.0

4.02 2.83 3.04 2.26 1.50 2.31

34.8 34.8 39.3 36.3 33.8 33.9

37.8 38.7 37.7 37.1 23.4 30.9

29.2 31.9 16.1 24.9 30.9 30.0

6.85 5.36 11.1 11.2 10.7 9.06

BA-1 0.71 13.1 0.5 > 11.9 BA-2 0.21 1.93 0.5 > 9.77 BA-3 0.37 6.55 0.5 > 11.8 Banga -dari BA-4 0.29 1.93 0.5 > 9.38 BA-5 0.22 1.66 0.5 > 12.4 BA-6 0.26 1.73 0.5 > 9.51 Non-carbonated groundwater CJG-1 0.35 26.8 17.6 39.9 Chojeong CJG-2 0.88 22.7 22.8 36.6 Surface water around the carbonated groundwater site BGS-1 0.80 26.2 8.96 23.7 BGS-2 0.65 21.1 10.2 22.5 Bugang BGS-3 0.71 25.0 13.8 23.4 BGS-4 0.99 29.3 7.91 21.9

439 582 576 574 586 552

131.8 128.5 133.2 118.6 125.2 118.6

3.22 2.05 2.26 1.47 1.01 1.24

32.7 31.9 35.2 30.5 29.5 28.1

16.0 16.8 14.7 12.2 8.43 13.8

33.2 36.2 20.2 26.5 34.0 32.6

14.0 13.8 18.2 16.6 16.8 15.9

523 350 306 298 1,650 1,440 1,700 1,600 1,590 1,620 1,000 1,040 1,030 1,030 1,040 985 682 809 800 775 798 758

109 96.0

33.3 34.7

0.76 1.63

4.33 3.51

20.8 24.9

21.7 28.6

0.05 > 0.05 >

275 272

127 137 134 152

35.7 38.9 40.7 45.7

5.22 4.01 4.23 3.82

4.34 5.41 6.02 5.04

20.5 18.7 22.3 25.2

19.3 19.9 18.2 21.7

0.05 > 0.05 > 0.05 > 0.05 >

272 278 288 314 60.9 62.0 64.7 88.4 40.8 38.9 38.8 37.0 53.0

Shin

ShS-1 ShS-2 ShS-3 ShS-4

0.09 0.08 0.05 > 0.10

3.63 4.04 8.36 10.2

4.40 3.87 6.54 4.45

3.23 3.21 3.08 3.19

28.0 29.0 24.0 37.0

7.24 7.99 9.94 11.0

0.63 0.56 0.52 1.36

1.03 1.32 1.45 0.1 >

2.82 3.09 2.97 11.1

9.84 8.81 7.86 9.97

0.05 > 0.05 > 0.05 > 0.10

Banga -dari

BAS-1 BAS-2 BAS-3 BAS-4 BAS-5

0.08 0.08 0.09 0.05 > 0.08

1.89 1.90 2.07 1.83 1.84

4.14 4.38 6.26 8.30 5.29

3.31 3.26 3.03 2.81 2.98

16.0 15.0 15.0 10.0 18.0

4.16 4.25 4.04 5.08 5.15

0.49 0.42 0.32 0.78 1.41

0.71 0.79 0.79 0.78 0.1 >

1.64 1.72 1.47 1.48 10.0

8.38 7.07 5.70 5.97 8.26

0.05 > 0.05 > 0.05 > 0.05 > 0.05 >

51

Table A5. Sample

δ13CDIC (‰)

δ18O (‰)

δ2H (‰)

SA-1 SA-2 SA-3 SA-4 SA-5 SA-6 SA-7 YS-1 DS-1 DS-2 DS-3 DS-4 DS-5 DS-6 DS-7 SM-1 SM-2 SM-3 SM-4 SM-5 SM-6 SM-7

-10.9 -11.6 -11.5 -12.7 -12.5 -12.0 -12.2 -18.0 -16.0 -17.3 -16.7 -18.5 -18.1 -17.7 -17.2 -14.5 -15.2 -14.4 -14.8 -17.1 -16.5 -15.5

-9.78 -9.58 -9.73 -9.62 -9.57 -9.20 -9.67 -8.55 -7.90 -7.67 -7.86 -7.78 -7.71 -7.47 -7.86 -4.42 -4.41 -4.40 -4.39 -4.50 -4.39 -4.41

-71.3 -70.8 -66.0 -66.1 -66.7 -66.1 -67.6 -58.2 -56.9 -56.5 -53.3 -53.2 -52.7 -52.4 -52.8 -35.6 -35.0 -32.9 -33.1 -31.3 -30.8 -30.5

SAG

-18.2

-9.47

YSG DSG SMG

-18.9 -19.1 -18.1

-7.85 -7.89 -7.38

Sample

δ13CDIC (‰)

δ18O (‰)

δ2H (‰)

-6.8 -6.7 -6.4 -7.7 -6.7 -5.5 -5.2 -7.5 -8.0 -7.9 -3.8 -3.4 -3.8 -5.0 -5.7 -5.4 -3.8 -3.2 -3.3 -4.0 -5.3 -5.4

-7.99 -7.80 -7.89 -7.86 -9.13 -8.98 -9.13 -9.07 -8.91 -8.94 -10.74 -10.57 -10.66 -10.65 -10.44 -10.49 -10.72 -10.53 -10.65 -10.59 -10.28 -10.22

-58.1 -58.6 -54.7 -56.3 -65.4 -65.9 -61.6 -62.8 -61.7 -61.4 -75.1 -76.9 -72.5 -73.0 -72.1 -71.3 -75.5 -75.3 -71.0 -71.6 -70.5 -70.3

-62.5

6400±30** 1290±30

CJ-1 CJ-2 CJ-3 CJ-4 BG-1 BG-2 BG-3 BG-4 BG-5 BG-6 Sh-1 Sh-2 Sh-3 Sh-4 Sh-5 Sh-6 BA-1 BA-2 BA-3 BA-4 BA-5 BA-6 CJG-1

-18.3

-8.03

-55.5

-54.2

410±30

CJG-2

-18.3

-8.06

-55.5

-52.9

-60±20

BGS-1

-13.3

-7.77

-56.6

-50.4

390±40

BGS-2

-11.9

-7.98

-54.6

14

C age

(yrs BP*)

5580±30 1980±30

3240±30

SAS-1 -11.5 -8.95 -60.6 BGS-3 -10.8 -7.93 -55.3 SAS-2 -8.9 -8.93 -62.1 BGS-4 -13.2 -7.56 -53.4 SAS-3 -13.8 -8.40 -57.4 ShS-1 -8.7 -10.16 -73.3 SAS-4 -14.4 -8.39 -56.4 ShS-2 -7.0 -9.98 -67.4 SAS-5 -8.52 -56.4 ShS-3 -7.5 -9.74 -64.1 YSS-1 -12.3 -7.86 -52.4 ShS-4 -8.3 -9.82 -65.2 YSS-2 -12.3 -7.74 -54.6 BAS-1 -13.0 -9.97 -71.4 DSS-1 -13.9 -7.58 -54.9 BAS-2 -10.1 -10.03 -65.7 DSS-2 -13.6 -8.15 -53.5 BAS-3 -6.9 -10.04 -68.0 DSS-3 -9.7 -8.09 -55.6 BAS-4 -9.9 -9.80 -64.4 DSS-4 -13.6 -7.02 -48.0 BAS-5 -15.0 -9.37 -64.3 DSS-5 -14.6 -7.79 -50.0 SMS-1 -7.4 -6.73 -49.9 SMS-2 -10.4 -6.89 -47.9 SMS-3 -7.86 -5.45 -40.1 SMS-4 -8.0 -3.95 -32.1 SMsea -2.6 0.11 -3.2 * BP means the abbreviation of ‘before present’, and the ‘present’ of 14C data reflects the atmospheric 14C concentration in 1950’s (Suckow et al., 2013). ** Personal communication with Yoon Y. Y. of KIGAM.

52

Highlights • Hydrochemical monitoring of geothermal groundwater and carbonated groundwater to interpret carbonating processes as natural analogues. • Carbonated reservoir depths using chalcedony geothermometer and geothermal gradient were estimated to be 1.34-3.90 km below surface level (b.s.l) in this natural analogue site. • Carbonating simulation using the PHREEQC modelling implies that groundwaters of a deep reservoir and a shallow reservoir have undergone different reaction paths to the final Ca-HCO3 type.

53