Geochemistry of High-Ph Waters from Serpentinites of the Gruppo Di Voltri

Applied Geochemistry 19 (2004) 787–802 www.elsevier.com/locate/apgeochem

Geochemistry of high-pH waters from serpentinites of the Gruppo di Voltri (Genova, Italy) and reaction path modeling of CO2 sequestration in serpentinite aquifers Francesco Cipolli, Barbara Gambardella, Luigi Marini*, Giulio Ottonello, Marino Vetuschi Zuccolini Laboratorio di Geochimica, DipTeRis, Universita` di Genova, Corso Europa 26, I-16132 Genova, Italy

Received 24 March 2003; accepted 10 October 2003 Editorial handling by H. A´ rmannsson

Abstract The large number of geochemical data gathered on the Gruppo di Voltri springs confirm that progressive interaction of meteoric waters with ultramafic rocks variably affected by serpentinization leads initially to the formation of Mg– HCO3 waters when the system is open to CO2, and Na–HCO3 and Ca–OH type water upon further interaction with the 3 rock, under highly reducing closed-system conditions with respect to CO2. As indicated by H data, these high-pH waters have had long residence times underground in deep aquifers hosted by serpentinitic rocks. These waters are the only available evidence of the presence of such deep aquifers. High-pressure injection of CO2 into these deep aquifers was simulated by reaction path modeling. Results indicate that this is a feasible methodology to reduce the inputs of anthropogenic CO2 into the atmosphere. Serpentinitic rocks have a high capacity for CO2 sequestration, mainly through formation of carbonate minerals. Dissolution of serpentinitic rocks and precipitation of magnesite and silica minerals occurs naturally in areas of high terrestrial CO2 fluxes such as in southern Tuscany, corroborating the feasibility of this methodology of CO2 sequestration. However, this process causes a progressive decrease in the porosity of the aquifer, at least under closed-system conditions. These side effects must be carefully evaluated by means of further laboratory tests and field activities. # 2003 Elsevier Ltd. All rights reserved.

1. Introduction the last century (Bryant, 1997), has not yet been proven. Nevertheless, humans cannot wait for a definite answer Since the beginning of the industrial revolution, the on this topic. Anthropogenic CO2 inputs to the atmo- CO2 concentration in the atmosphere has been increas- sphere must be drastically reduced, and several strate- ing. A rise from315 ppm in 1958 to370 ppm in 2001 gies have to be undertaken for this purpose. has been documented (Keeling and Whorf, 2002). This Sequestration of CO2 in deep geological reservoirs is global-scale phenomenon has been attributed chiefly to one disposal option. It is currently under way in the the use of fossil fuels and to a lesser extent to concrete Norwegian sector of the North Sea (Korbol and Kad- production (e.g., Vernadsky, 1924; Marland et al., dour, 1995) and its feasibility has been evaluated for 2001). The causative relationship between this CO2 other sites, such as the Alberta sedimentary basin, increase in the atmosphere and global warming, with an Canada (Gunter et al., 1993, 1996, 1997; Bachu et al., approximate temperature increase of 0.4–0.6 C during 1994; Law and Bachu, 1996). According to Hitchon (1996), geological storage of CO2 can be carried out in 3 ways: (1) trapping as gas or * Corresponding author. supercritical fluid below a low-permeability caprock, a E-mail address: [email protected] (L. Marini). process termed hydrodynamic trapping; (2) dissolution

0883-2927/$ - see front matter # 2003 Elsevier Ltd. All rights reserved. doi:10.1016/j.apgeochem.2003.10.007 788 F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 into deep waters (solubility trapping); (3) precipitation Voltri, variably affected by serpentinization, are covered of carbonate minerals brought about by dissolution of by the comparatively impervious sedimentary rocks of primary silicates upon injection of CO2 into a deep the Piedmont Tertiary Basin potentially acting as an aquifer (mineral sequestration). effective caprock (Bozzo et al. 1986, 1992). Mineral sequestration is especially effective as it These aquifers host waters with unusual chemistry, 2+ would fix CO2 forever, preventing its return to the dominated by Ca and OH ions, and high pH, typi- atmosphere. It is well known that CO2 is extracted from cally in the 11.3–11.9 range at the surface outlet (Bruni the atmosphere by silicate weathering and consequent et al., 2001, 2002). Other characteristics of these waters deposition of carbonate minerals (Lasaga, 1981). The are the very low concentrations of total dissolved car- 2 balance between this process and the CO2 release to the bonate, mainly present as CO3 ion, and the deposition atmosphere by Earth degassing, controls the CO2 con- of travertines from surface discharges through absorp- centration in the atmosphere on a geological time scale tion of atmospheric CO2 (Bruni et al., 2001, 2002; of 1 million years or more (Walker et al., 1981; Berner et Marini and Ottonello, 2002). al., 1983; Kerrick and Caldeira, 1993). To investigate the feasibility of CO2 sequestration The appeal of permanent CO2 fixation in carbonates through injection into deep aquifers hosted by the in deep geological formations (Seifritz, 1990; Bachu et ultramafic-serpentinitic rocks of the Gruppo di Voltri it al., 1994) has stimulated the experimental investigation was decided to gather additional geochemical data on of mineral trapping. Pearce et al. (1996) and Rochelle et local springs, especially those of high-pH and to carry al. (1996) have injected supercritical CO2, at pressures of out reaction path modeling of CO2 sequestration. The 90 and 200 bar, into a sandstone reservoir at tempera- results achieved so far are presented below. tures of 105 and 80 C, for periods of 1 and 8 months. Significant alteration of pre-existing calcite and dolo- mite was observed, as well as dissolution of anhydrite 2. The high-pH waters accompanied by precipitation of calcite and possible corrosion of detrital feldspars with precipitation of Na- 2.1. Geological background smectite, although the evidence supporting this last process was not conclusive. Gunter et al. (1997) performed The Gruppo di Voltri includes (Chiesa et al., 1975): 1-month-long experimental tests at 105 C and 90 bar of (1) antigoritic serpentinites, less frequent eclogitic meta- PCO2 on glauconitic sandstones from the Alberta sedimen- gabbros, and infrequent metarodingites, produced tary basin. Under these conditions, only a small amount through metamorphism of original associations of of CO2 was trapped through reactions with Al-silicate ultramafic rocks, prevailingly peridotites, and gabbros; minerals, whereas dissolution of mineral carbonates (2) prasinites and calc-schists, derived by metamorphism took place. Geochemical modeling suggests that a per- of original basalts and associated oceanic sediments. iod of much longer than one month, probably 6–40 a, is High-pH, Ca–OH springs are situated in zones where needed in this experimental system to attain chemical serpentinites and related rocks crop out, generally equilibrium. Completion of CO2 fixation in the natural along faults and fractures. The physical and chemical system would require even longer periods of time, of the characteristics of these waters suggest that they come order of hundreds of years. However, these periods are from relatively deep aquifers, which are mainly hosted shorter than the residence times of fluids in deep aquifers. in serpentinites and related rocks (Bruni et al., 2001, From these considerations, it is clear that geochemical 2002). modeling is an appropriate tool in projects aimed at evaluating CO2 sequestration through carbonate 2.2. Sampling and analyses mineral deposition, at least during the pre-feasibility stage, because of the slow kinetics of alteration reactions A total of 25 samples were collected from the 15 of silicate and Al-silicate minerals and the difficulties springs of high pH identified in the Gruppo di Voltri and cost of field tests. area. Field characteristics, including location, are As pointed out by Lackner et al. (1995), disposal of described by Marini and Ottonello (2002). In the field, large amounts of CO2 through mineral carbonation outlet temperature, pH, Eh, sulfide (by the methylene requires involvement of silicate rocks rich in Mg and blue colorimetric method), and total alkalinity (by Ca. Ultramafic rocks, such as peridotites and serpenti- acidimetric titration) were determined. Spring water was nites, are especially suitable for this purpose as they filtered through 0.45 mm cellulose acetate membranes. A contain 40–50 weight% of MgO on average. filtered portion was stored as such, for Ion Chromato- A very promising geological setting for CO2 seques- graphy (IC) analysis, whereas concentrated HCl was tration through injection into deep aquifers hosted by added before storage to a second filtered portion, for ultramafic-serpentinitic rocks is the southern Piedmont, Atomic Absorption Spectrophotometry (AAS) and SiO2 Italy, where the ultramafic rocks of the Gruppo di analysis. New polyethylene bottles were used. F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 789

A separate aliquot was collected into an evacuated In the d18O vs. dD correlation plot (Fig. 1), the 184 glass bottle, containing concentrated HCl in excess with spring waters from the Gruppo di Voltri area are dis- respect to total alkalinity to determine total dissolved tributed, independent of their compositions, into a inorganic C, which is also called total dissolved carbon- unique alignment, which is conveniently described by ate or total dissolved CO2 in the geochemical literature. the linear regression line (R=0.986): It consists of the sum of the molal concentrations of 18 2 D ¼ 7:64 Â O þ 12:4: ð3Þ CO2, HCO3 ,CO3 and related aqueous complexes and is here indicated with the acronym TDIC. For a given 18O/16O ratio, the dD values of spring All the high-pH waters were analyzed for: (1) Na, K, waters are 3–5 % units less negative than those of the Mg, Ca, by AAS; (2) Cl, SO4,NO3 by IC; (3) molyb- meteoric waters from the Genova-Sestri Ponente IAEA- date-reactive SiO2 by the heteropoly blue colorimetric WMO station. Although this shift is, for unknown rea- method; (4) dissolved CH4 by Gas Chromatography sons, larger than the analytical uncertainty for dD, 1 % (GC); (5) TDIC, according to the method reported in unit, there is little doubt of the meteoric origin of all the Appendix. In addition, the 18O/16O and D/1H ratios groundwaters sampled, independent of their composition. and 3H activity were determined in water from selected Tritium activity of a groundwater discharging at the springs. All chemical data, dD and d18O values for high- surface can be used to estimate its residence time in the pH waters are reported in Table 1, whereas 3H activity aquifer system of provenance, with reference to the two is listed in Table 2. limiting models of piston flow and a well-mixed reservoir In addition to the high–pH springs, approximately (Pearson and Truesdell, 1978). The first model assumes 600 neutral springs in the surveyed area were sampled, absence of mixing along the entire circuit, from the the samples analysed chemically, and 18O/16O and D/1H infiltration to the emergence, whereas the second model isotope ratios determined in selected samples. Analytical is based on the hypothesis of perfect mixing between the data for these neutral springs are given in Marini and water entering the reservoir and the water already stored Ottonello (2002). within it. It is also assumed that this mixture is fully representative of the water leaving the reservoir. Knowl- 2.3. Isotope geochemistry edge of time changes in 3H activity of local meteoric waters is essential to both models. Tritium activity has The IAEA–WMO station of Genova-Sestri Ponente is been determined on a regular basis since 1961 for the rain located within the study area. Here, rain water samples waters collected at the IAEA–WMO station of Genova- are periodically collected and analysed for the D/1H and Sestri Ponente and a time series of 359 data points is 18 16 3 O/ O isotope ratios of H2O and H activity (IAEA/ available for this site (IAEA/WMO, 1998). Similar to WMO, 1998). what has been observed elsewhere in the world, 3H Available dD and d18O data on local rain waters increased rapidly in 1961–1963, reaching the peak value cover the time intervals 1961–1965 and 1972–present, of 4473 TU in May 1963, with a subsequent gradual for a total of 304 pairs of dD and d18O values. Large decrease. The reason for this is the input of 3H into the fluctuations have been recorded around the average atmosphere in 1954–1962 through thermonuclear explo- values of 31.4 % for dD and 5.07 % for d18O, sions and cessation of this activity afterwards, accom- with minima and maxima of 91 and +17.6 % for panied by 3H decay with a half-life of 12.26 a (Faure, dD and 12.7 and +2.58 % for d18O, respectively. 1986). In spite of large short-term fluctuations, which These data define the following linear regression line are probably reduced by mixing in the aquifer system, (R=0.962): the 3H data for the Genova-Sestri Ponente station, from 1963 to present, fit the regression equation (R=0.909): D ¼ 7:28 Â 18O þ 5:54 ð1Þ 3H ¼ 344:85 EXP½0:1397 Â ðÞy1963:4 ; ð4Þ which is representative of local rain waters. This relationship is somewhat different from the world meteoric where y indicates time in years. Eq. (4) and those water line, initially proposed by Craig (1961), and derived by Pearson and Truesdell (1978) were used to recently updated by Rozanski et al. (1993), who pro- obtain the theoretical relations between residence time posed this relation: and 3H(Fig. 2) which are needed to interpret the 3H analytical data for high-pH springs (Table 1). Samples D ¼ 8:17 Â 18O þ 10:4 ð2Þ collected from springs C11 and S70 in December 1999 and based on the average isotopic data collected from 206 from springs PIO14, LER20, and GOR34 in February– stations of the IAEA/WMO grid. However, differences April 2001 have very low 3H activities. These indicate between Eqs. (1) and (2) are small in the dD and d18O residence times > 37 a based on the piston-flow model, ranges of spring waters from the study area, and the two whereas residence times between 100 and 10,000 a are lines intersect at dDof40.6 % and d18Oof6.32 % generally evaluated based on the well-mixed model, (Fig. 1). which is the most likely reference frame for these waters 790

Table 1 Concentrations of chemical components, dD and d18 O values, and other relevant data for the high-pH waters of the surveyed areaa

18 Sample Date X m Ym Hm T CEh pH Ca Mg Na K AIkt SO4 CL SiO2 NO3 CH4 HS TDIC PCO 2(*) dD % d O % code (UTM) (UTM) asl mV mg/L mg/L mg/L mg/L mg/L mg/L mg/L mg/L mg/L mg/L mg/L mg/L bar vs. vs. as HCO3 as HCO3 V V SMOW SMOW

V18 Jan 01 485187 4921800 140 14.0 482 11.37 22.5 0.012 16.6 2.45 88.5 4.10 11.7 22.1 <0.001 0.18 0.62 1.85 4.07E10 38.0 6.45 BR1 Jan 01 482500 4922000 120 12.0 515 11.86 47.4 0.020 23.7 3.10 172 0.14 21.2 1.44 <0.001 1.84 0.14 1.26 2.29E11 n.d. n.d. .Cplie l ple eceity1 20)787–802 (2004) 19 Geochemistry Applied / al. et Cipolli F. L43 Jan 01 481715 4922985 150 23.0 516 11.52 49.3 0.008 28.3 5.07 203 0.44 18.6 1.97 <0.001 4.00 0.31 3.09 2.05E10 44.3 7.38 S70 Jan 01 472650 4931900 430 12.0 479 11.42 36.4 0.039 5.5 1.78 83.6 18.3 23.3 0.14 <0.001 0.16 1.34 0.84 1.26E10 51.6 8.28 C11 Jan 01 475775 4924210 315 10.5 390 10.50 3.2 5.80 12.8 1.54 34.2 25.3 23.2 0.17 <0.001 2.02 1.03 20.4 2.21E07 37.1 6.46 LER20 Jan 01 472550 4918775 150 13.7 493 11.57 44.1 0.164 12.7 1.03 117 12.8 26.5 0.18 <0.001 0.32 1.03 2.35 1.54E10 41.2 7.22 BR2 Sept 01 482780 4922237.5 210 20.3 525 11.73 61.9 0.001 41.1 5.51 260 0.10 30.5 3.12 <0.001 13.9 0.22 0.53 1.22E11 39.2 6.94 ERR20 Sept 01 460200 4928700 360 15.6 481 11.36 20.5 0.032 16.5 1.78 79.3 3.08 15.4 3.31 <0.001 0.01 0.13 1.97 4.51E10 62.8 9.60 GOR34 Oct 01 483025 4938400 400 18.5 516 11.68 59.0 0.55 18.5 4.22 212 0.10 17.3 1.64 <0.001 3.27 0.11 5.47 1.70E10 48.0 8.09 GOR34A Oct 01 478515 4938500 400 19.0 502 11.55 46.9 1.81 18.3 4.15 185 0.55 15.0 1.64 <0.001 0.98 0.03 17.1 1.12E09 46.3 8.07 LERI8A Sept 01 472700 4919225 160 14.0 483 11.38 37.8 0.60 6.8 0.51 98.6 1.18 18.9 0.86 <0.001 0.28 0.19 15.4 2.70E09 39.9 7.14 LER2 Sept 01 469625 4917175 230 14.2 460 11.11 20.0 0.17 10.3 0.63 58.0 3.78 19.7 1.25 <0.001 0.05 0.09 10.8 7.98E09 39.1 6.90 LER20 Sept 01 472550 4918775 150 17.0 502 11.53 45.0 0.033 12.7 1.08 127 1.53 27.9 0.47 <0.001 0.33 0.58 2.08 1.47El0 n.d. n.d. LER2I Sept 01 472875 4919225 200 18.0 503 11.49 43.5 0.015 9.9 1.01 124 1.11 23.1 0.47 <0.001 0.34 0.57 0.96 8.34E11 41.1 7.18 ORB101 Sept 01 464925 4929900 540 12.6 423 10.59 26.2 15.4 3.9 0.74 139 3.34 14.7 11.2 1.51 0.02 0.06 68.9 3.96E07 54.2 8.38 PIO14 Oct 01 478525 4938500 300 14.0 441 10.69 0.6 4.71 53.0 10.8 150 1.68 17.2 0.09 <0.001 5.94 0.58 78.6 4.16E07 54.1 8.69 S70 Oct 01 472650 4931900 430 13.8 493 11.48 36.6 0.041 5.4 0.55 86.4 2.14 23.9 0.76 <0.001 0.23 1.07 0.81 8.76E11 n.d. n.d. V18 Oct 01 485187 4921800 140 15.6 481 11.31 22.2 0.001 16.1 2.54 92.5 4.60 11.2 22.3 <0.001 0.17 0.56 2.36 6.58E10 n.d. n.d. GOR35 Oct 01 483125 4938350 415 19.0 490 11.44 25.0 0.026 6.7 1.92 111 0.10 8.96 1.15 <0.001 nd. 0.06 n.d. n.d. 51.7 8.31 L43 May 02 481715 4922985 150 22.4 518 11.55 49.5 0.002 27.7 5.05 201 0.45 20.8 1.87 <0.001 4.16 0.47 0.57 3.30E11 n.d. n.d. BR1 May 02 482500 4922000 120 14.0 514 11.79 47.3 0.004 23.5 3.11 178 0.20 20.8 1.43 <0.001 1.80 0.12 1.69 4.06E11 n.d. n.d. BR3 May 02 482837.5 4922175 150 13.0 504 11.72 40.2 0.002 18.4 2.59 149 1.50 17.4 10.1 <0.001 1.03 0.39 1.56 5.69E11 n.d. n.d. PIO14 May 02 478525 4938500 300 13.1 426 10.49 0.6 4.50 53.6 10.7 145 2.52 19.6 0.09 <0.001 6.37 0.62 85.9 9.70E07 n.d. n.d. GOR36 May 02 479425 4939875 350 13.0 388 9.95 3.9 8.80 84.0 8.32 215 2.95 46.1 2.23 <0.001 5.60 0.22 160 1.02E05 n.d. n.d. V99 May 02 486712.5 4924087 210 16.4 480 11.28 25.2 0.044 68.1 3.55 79.3 20.4 97.4 1.10 <0.001 3.72 1.81 5.33 1.59E09 n.d. n.d.

a n.d.=not determined. (*)PCO 2 was computed by means of EQ3NR. F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 791

2 of deep provenance (see below). The only exception is (computed on the basis of the SO4 /HS redox couple) spring LER20, for which two different solutions are of 460 to 520 mV, apart from sample ORB101 which possible based on the well-mixed model. However, the is affected by mixing with surface waters. These waters lower one (5 a) is considered to be highly unlikely. have low Mg concentrations (as low as 0.001 mg/kg), Summing up, dD and d18O data indicate that all the low total dissolved carbonate (most between 0.5 and 17 springs sampled are of meteoric origin, independent of mg/kg as HCO3 ) and, consequently, very low PCO2 their composition. Tritium activities of high-pH waters values, from 108.1 to 1010.9 bar. suggest that they circulate for long intervals of time in Four samples (GOR36, PIO14a, PIO14b, and C11) relatively deep aquifers. have Na-HCO3 chemical composition, and with respect to the Ca-OH waters described above, they have some- 2.4. Chemical characteristics of waters interacting with what lower pH values (9.95 to 10.7), less negative Eh ultramafic rocks and serpentinites (390 to 440 mV), somewhat higher total dissolved 6.7 carbonate (20 to 160 mg/kg as HCO3) and PCO2 (10 Twenty of the 25 samples of high-pH waters have Ca– OH chemical composition, pH of 11.3 to 11.9 and Eh

Table 2 Tritium activity, referred to the sampling date, for some high- pH waters of the study area (1 TU corresponds to an atomic 3H/1H ratio of 1018)

Sample Sampling Tritium date activity TU

PIO14 04/04/01 1.20.5 C11 21/01/01 0.40.5 S70 20/01/01 0.00.5 LER20 14/03/01 3.10.5 GOR34 05/05/01 0.60.5

Fig. 1. Plot of d18O vs. dD, both in % vs. V-SMOW, for the 184 spring waters coming from the Gruppo di Voltri area, also showing the linear best fit for these data, the local meteoric water line, and the world meteoric water line. Sym- bols are as follows: triangles=high-pH waters (this study); Fig. 2. Theoretical relationships between residence time and 3H diamonds=neutral Mg–HCO3 waters from Marini and Otto- for the (a) piston-flow model and (b) the well-mixed reservoir 3 nello (2002); squares=neutral Ca–HCO3 waters from Marini model. The H analytical data of the high-pH springs of the and Ottonello (2002). Gruppo di Voltri area are also shown (triangles). 792 F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 to 105.0 bar), and larger Mg concentrations of 4.5 to All these values are typical of many natural ground- 8.8 mg/kg. As shown by Bruni et al. (2001; 2002) for waters circulating in comparatively shallow aquifers. sample C11, reaction path modeling suggests that these Analytical data for both high-pH waters and neutral waters appear to be relatively less evolved than those pH Mg-HCO3 waters are reported in the Eh-pH, total belonging to the Ca-OH facies (see below). carbonate-pH, Ca-pH, and Mg-pH plots of Fig. 3, also Sample V99 is unique for its mixed Na-Cl/Ca-OH showing the theoretical paths referring to dissolution of characteristics, but pH (11.3), Eh (480), total dissolved a local serpentinite simulated by means of the EQ3/6 8.8 carbonate (5.3 mg/kg as HCO3 ), and PCO2 (10 bar) software package (Wolery, 1992; Wolery and Daveler, are comparable to those of Ca–OH waters. 1992). Following Bruni et al. (2002) and Marini and The chemical characteristics of the high-pH waters Ottonello (2002) this process was simulated in two are markedly different from those of the neutral pH separate steps. Mg-HCO3 waters, whose pH is 6.0 to 8.5, Eh 100 to 350 In a first step, the dissolution of a local serpentinite 3.5 1.5 mV, and PCO2 values mostly between 10 to 10 bar. (Table 3) was simulated under open-system conditions

Fig. 3. Correlation plots of (a) Eh–pH, (b) total carbonate–pH, (c) Ca–pH, and (d) Mg–pH showing the analytical data for both high- pH waters (triangles) and neutral Mg–HCO3 waters (diamonds) from Marini and Ottonello (2002), as well as the theoretical paths 41.5 referring to dissolution of a local serpentinite under both (i) open-system conditions with respect to CO2,PO2 of 10 bar, and 73.6 temperature of 12 C, and (ii) closed-system conditions with respect to CO2,PO2 of 10 bar, and temperature of 12 and 25 C. F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 793

1.5 2.5 with respect to CO2 by setting PCO2 at 10 ,10 , important. Modeling of this process is complicated by and 103.5 bar in separate runs, with a temperature of redox disequilibrium. In fact, for high-pH Ca–OH 41.5 2 12 C, and PO2 of 10 bar, in order to reproduce the waters, the log PO2 computed based on the HS /SO4 average T-PCO2-redox conditions present in the shallow redox couple (which was chosen as reference to fix redox aquifers hosting Mg–HCO3 waters. The PCO2 values of conditions in reaction path modeling) are 3–4 units 1.5 3.5 10 and 10 bar represent limiting values for these higher than the log PO2 calculated on the basis of the 2.5 waters, whereas the PCO2 of 10 bar is close to the CH4(aq)/HCO3 redox couple. It is not surprising, there- average value of this log-normally distributed variable. fore, that the concentrations of aqueous CH4 predicted 41.5 The PO2 of 10 bar reproduces the average redox by the model are several orders of magnitude lower than conditions of these shallow aquifers, as shown in Fig. 3a, the analytical data. Moreover, to model reduction of in spite of the large scatter of analytical data. Separate carbonate-C to CH4, the open system constraint with runs were performed with different PCO2 because this respect PO2 must be abandoned and a proper reductant variable has a huge impact on the distribution of many must be chosen. Further research is needed on this chemical components of interest in the considered multi- subject. phase system. On the other hand, PO2 was not varied up to the limiting values for these waters (1037 to 1046 bar) as these changes have no influence on the multi- 3. The dissolution kinetics of serpentine valent components of interest. In a second step, the same serpentinitic rock was Since serpentinites are almost monomineralic rocks reacted with a Mg–HCO3 water produced in the pre- primarily constituted by serpentine, CO2 sequestration vious step for incipient calcite saturation, at a PCO2 of through mineral carbonation of serpentinites is expected 3.5 10 bar. The system was closed with respect to CO2, to be largely governed by the dissolution kinetics of 73.6 PO2 was decreased to 10 bar (Fig. 3a) and tempera- serpentine which, consequently, is a subject of utmost ture was fixed to 12 C (25 C in a separate run), in importance for this study. Unfortunately the dissolution order to reproduce the T-PCO2-redox conditions pre- kinetics of serpentine has been the subject of a relatively sumably present in the relatively deep aquifers hosting limited number of experimental investigations (Luce et high-pH Na–HCO3 and Ca–OH waters. al., 1972; Lin and Clemency, 1981; Bales and Morgan, The general agreement between theoretical paths and 1985). The available experimental results are critically analytical data indicates that progressive interaction of reviewed below. meteoric waters with serpentinites leads to the forma- Luce et al. (1972) carried out batch dissolution tion of the Mg–HCO3 chemical waters, under open-sys- experiments on lizardite, one of the 3 main polymorphs tem conditions with respect to CO2, whereas the of serpentine, and on forsterite and enstatite, at 25 C subsequent evolution towards the Na–HCO3 and Ca–OH and different pH values. At all pHs, the Mg/Si molal compositions occurs in a system closed to CO2 and ratio in the aqueous solution was significantly different under strongly reducing conditions (Pfeifer, 1977; Bruni from the stoichiometric value of 1.5, indicating incon- et al., 2002; Marini and Ottonello, 2002). gruent dissolution. According to Luce et al. (1972), This chemical evolution is due to C depletion in a during the first dissolution step of all these solid phases, system closed with respect to C sources. This C deple- Mg2+ ions, which are present at the surface of the solid tion was attributed to calcite precipitation (Bruni et al., phase, exchange with H+ ions. Extraction of Mg and Si 2001, 2002), although the possible reduction of carbon- from the crystal lattice would take place subsequently + ate-C to organic-C and ultimately to CH4 might also be and, again, exchange with H ions. During this second step, the amounts of Mg and Si entering the aqueous solution are linearly correlated with the square root of Table 3 time, t, i.e., the kinetics is parabolic. Only in strongly Composition of the serpentinite of Case Giutte (Cortesogno et acidic solutions (pH 1.65), was linear kinetics al., 1979) used as solid reactant in reaction path modeling observed. According to Berner (1981), the dissolution rates of SiO2 42.28 TiO2 0.04 silicate minerals under weathering conditions exhibit Al2O3 0.41 a linear dependence on t, if solid phases are suitably Fe2O3 9.60 pre-treated to remove highly reactive, ultrafine particles. MnO 0.13 ‘‘We believe that the parabolic kinetics obtained by MgO 34.13 other workers [e.g., Luce et al. (1972)] is the result of CaO 1.74 varying rates of dissolution of particles produced during Na2O 0.05 grinding’’ (Berner, 1981). K O 0.13 2 Even though these lizardite dissolution rates are not H O 10.81 2 linear, except in strongly acidic solutions, and are affected 794 F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 by non-stoichiometric release of Mg and Si, the experi- stoichiometric release of Mg and Si is attained in some mental data of Luce et al. (1972) were used to derive the hours in acidic solutions and in 1 week in basic solu- dissolution rate of lizardite as a function of pH. Luce et tions. The logarithm of the dissolution rate of forsterite al. (1972) report their results in moles/cm2 of Mg and Si determined by Pokrovsky and Schott (2000) under stoi- extracted from each mineral as a function of time. The chiometric conditions correlates linearly with pH with a pH of the aqueous solution is also given as a function of slope close to 0.5 in the pH interval 1 to 8 (Fig. 5). time. These data were used to compute the ratio ÁMg/ The dissolution rates measured at pH > 8 are not 3Át [moles m2 s1], where ÁMg is the increase in the taken into account here as they are beyond the pH amount of dissolved Mg [moles m2] during the time range of interest of this study. Fig. 5 also shows that interval Át [s]. The coefficient 3 is used to convert the the data of Luce et al. (1972) are also proportional to 0.5 moles of Mg in moles of lizardite, whose chemical for- aH+ but differ in the intercept which is 0.78 log units mula is Mg3Si2O5(OH)4. The logarithm of the ratio greater than that of the Pokrovsky and Schott (2000) ÁMg/3Át (which approximates the dissolution rate of data. Assuming that the rates of lizardite dissolution lizardite) is plotted against the average pH measured in have the same shift in the intercept, the following the corresponding time interval, Át,inFig. 4. Although equation is obtained for lizardite: data are somewhat scattered, the logarithm of the dis- ÂÃ log Rate moles m2 s1 ¼0:31 Â pH 7:08 ð5Þ solution rate of lizardite decreases linearly with increasing pH with a slope close to 0.3, which is the minimum Lin and Clemency (1981) measured the dissolution value generally observed for silicates and Al-silicates rate of antigorite (a serpentine polymorph), brucite, and (Drever, 1994). talc at 25 C at a constant PCO2 of 1 bar. During the Now the dissolution rates of forsterite measured by dissolution of antigorite, the Mg/Si molal ratio in the Luce et al. (1972) are compared with those recently aqueous solution was always greater than 1.5, the determined by Pokrovsky and Schott (2000), who used expected value for congruent dissolution. According to an open-system mixed-flow reactor and performed a Lin and Clemency (1981) the reasons for incongruent large number of steady-state dissolution experiments at dissolution might be either preferential removal of pH from 1 to 12, varying ionic strength, and in both Mg2+ ion from the octaedral layer with respect to that 4+ CO2-free and CO2-bearing aqueous solutions. Pok- of Si ion from the tetraedral layer or the precipitation rovsky and Schott (2000) have shown that the initial of a secondary, Si-bearing phase, e.g., amorphous silica. stages of forsterite dissolution are always non-stoichio- This second possible explanation is supported by equi- metric, with preferential release of Mg and formation of librium calculations, using EQ3, which show that a Si-rich layer at pH 48, whereas at pH 510 the pre- amorphous silica saturation was attained in the experi- ferential expulsion of Si generates a Mg-rich layer. The ments and maintained after 3 h. The analyses reported after 1, 3, 7, and 15 h indicate that magnesite saturation was also attained, but precipitation of this solid phase is

Fig. 4. Dependence of the logarithm of the dissolution rate of serpentine polymorphs on pH at 25 C. Triangles=data of Luce et al. (1972) on lizardite; circles=data of Lin and Clem- Fig. 5. Dependence of the logarithm of the dissolution rate of ency (1981) on antigorite; crosses=data of Bales and Morgan forsterite on pH at 25 C. Squares=data of Pokrovsky and (1985) on chrysotile. Schott (2000); triangles=data of Luce et al. (1972). F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 795 considered unlikely due to its relatively slow kinetics, at laboratory for reaction path modeling of CO2 seques- least at 25 C(Chou et al., 1989). As the analytical data tration in serpentinite aquifers. referring to 1, 3, 7 and 15 h are those closest to the condition of congruent dissolution, they were used to compute the ratio ÁMg/3Át (see above), whose loga- 4. Simulation of CO2 sequestration through high-pressure rithm is plotted against pH in Fig. 4. Interestingly, these injection into a deep aquifer hosted in serpentinitic data plot relatively close to those obtained by Eq. (5). rocks Bales and Morgan (1985) determined the kinetics of dissolution of chrysotile, through 5-day-long batch tests, The irreversible mass exchanges presumably taking at constant temperature, 25 C, and pH, in the range 7 place during high-pressure CO2 injection into a deep to 10. The measured Mg/Si molal ratio in the aqueous aquifer hosted in serpentinitic rocks were modeled by solution was close to 2, i.e., significantly greater than the means of the software package EQ3/6 version 7.2b values of 1.4 to 1.5 expected for stoichiometric dissolu- (Wolery, 1992; Wolery and Daveler, 1992). As anticipated, tion of chrysotile used in the experiments. this investigation is aimed at evaluating: (1) the amounts According to Bales and Morgan (1985), the release of of CO2 that can be stored through both solubility trapping Mg takes place at constant rate, after the first day, and and mineral sequestration, and (2) the volume changes 0.24 is proportional to aH+. The dissolution rates of chryso- occurring upon dissolution of serpentinite and precipita- tile evaluated on the basis of the Mg release at pH 7, 7.5, tion of secondary phases, and related implications. and 8 are smaller than those computed by means of equation (5), by 1–1.5 log units (Fig. 4). 4.1. Setting up the water–rock interaction model Assuming that the more reliable data on the dissolution kinetics of serpentine, irrespective of the poly- To set up the model, the CO2 fugacity and aquifer morphic form, are those of Luce et al. (1972) at pH temperature were assumed to be 250 bar and 60 C, 1.63–1.67 (as they are the only ones of this dataset respectively. It was hypothesized that the aqueous solu- showing linear kinetics) and those of Bales and Morgan tion hosted in the serpentinitic reservoir is represented (1985), the dependence of the dissolution rate of ser- by the Ca–OH spring water BR2 (Table 1). Heating of pentine on pH is described by the following relation this aqueous solution from the emergence temperature, (Fig. 4): 20.3 C, to the 60 C reservoir temperature was simu- ÂÃ lated by means of EQ3/6. log Rate moles m2 s1 ¼0:70 Â pH 6:38 ð6Þ Since serpentinites are almost monomineralic rocks, Finally, it is valuable to recognize that all these dis- stoichiometric serpentine [Mg3Si2O5(OH)4] was con- solution rates of serpentine measured in laboratory sidered to be the only solid phase under dissolution. experiments are several orders of magnitude greater Simulation was carried out in kinetic mode, assuming than those obtained by Freyssinet and Farah (2000), constant values for the dissolution rate of serpentine either 13.85 or 15.76 log (moles m2 s1), depending and its reactive surface. Instantaneous partial equili- on the reactive surface area. These values were esti- brium was assumed for precipitating, secondary solid mated by means of elemental mass balances at the phases, which means to assume that the dissolution of watershed scale for an Amazon watershed with out- reactants is the rate limiting step of the overall process. cropping ultramafic rocks. A large difference between Results are presented against both time and the reaction dissolution rates measured in the field and rates deter- progress variable, x, which was expressed in g of dis- mined in the laboratory has been found in several stud- solved serpentine per kg of water, for simplicity. ies of silicate dissolution kinetics (e.g., Appelo and In the previous sections two different relations Postma, 1996 and references therein). In the authors’ expressing the dependence of the serpentine dissolution opinion, this is an ill-posed problem, because in these rate on pH were derived. The serpentine dissolution studies the field and laboratory data are considered rates computed by means of Eqs. (5) and (6) gradually antithetical rather than synergistic. A synergistic, suc- diverge with increasing pH, but at the pH of interest, cessful approach is the kinetic dissolution model of 4.5 (see below), differences are of 1 log-unit only. Sverdrup and Warfvinge (1988) and Sverdrup (1990). Therefore the average value, 109 moles m2 s1, was This model uses the kinetic rate laws determined in the taken as the most probable dissolution rate of serpen- laboratory to compute the dissolution of different tine at 25 C at the pH of interest. Assuming that the minerals considering the effects of the soil solution activation energy for dissolution of serpentine Ea is composition, available surface area, and temperature. 7010 kJ mol1 (Thomassin et al., 1977), the serpentine Results of the kinetic model are comparable with results dissolution rate at the temperature of interest, T, was from mass balance calculations at the watershed scale. computed by the integrated Arrhenius equation: Based on these encouraging findings, the authors will RateT ¼ Rate25 Â exp½ððÞEa=R ðÞ1=T 1=298:15 7Þ use the dissolution rate of serpentine determined in the 796 F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 obtaining the value of 1.93Â108 moles m2 s1 at 4.3. The aqueous solution 60 C. To estimate the reactive surface of serpentine refer- During CO2 injection, the pH of the aqueous solution ence was made to Mg–HCO3 waters, with average pH changes from 3.3 to 4.5 (Fig. 7a). Most of this pH of 7.240.66, average Mg concentration of 16.811.4 increase occurs between the onset of chalcedony pre- mg/kg, and average temperature of 10.72.9 C. Under cipitation and the beginning of magnesite precipitation, these temperature-pH conditions, the mean dissolution after which pH is buffered at large m (Fig. 7b)in HCO3 rate of serpentine is 9.9Â1012 moles m2 s1 (from Eq. accordance with the following equilibria: 5, 6, and 7). Assuming that dissolved Mg is entirely MgCO þ Hþ ¼ Mg2þ þ HCO ð10Þ contributed by stoichiometric dissolution of serpentine, 3 3 its reactive surface, s, can be computed by means of the and following equation: þ CO2ðÞg þ H2O ¼ HCO3 þ H ð11Þ mMg ¼ vMg Â Rate Â s Â Dt; ð8Þ

2+ where nMg is the stoichiometric coeﬃcient of the Mg ion in the dissolution reaction of serpentine. Eq. (8) is obtained hypothesizing that s, rate, and nMg do not change during the considered interval of time Át. Plug- ging in Eq. (8) the average residence time of HCO3 waters, 1.8 a based on 3H data (Marini and Ottonello, 1997; Marini et al., 2000), it turns out that the reactive surface of serpentine is 0.41 m2, although it might range from 0.061 to 1.5 m2 taking into account the changes in pH and Mg concentrations. Although this uncertainty in the reactive surface might appear rather large, it is not surprising considering that the main dif- ﬁculty in kinetic geochemical modeling is the character- ization of reactive surface areas of solid phases (e.g., Lichtner, 1996, 1998).

4.2. Solid product phases

The only solid phases which are considered to be produced through high-pressure (f CO2 250 bar) CO2 injection into a deep aquifer hosted in serpentinitic rocks are chalcedony, which precipitates at x 50.07 g of dissolved serpentine/kg of water, and magnesite, which precipitates at x 54.6 g of dissolved serpentine/ kg of water (Fig. 6a). For x > 20–30 g/kg the stoi- chiometry of the process is adequately described by the reaction:

Mg3Si2O5ðÞOH 4þ3CO2 ¼ 3MgCO3 þ 2SiO2 þ 2H2O: ð9Þ

In other words, 3 moles of magnesite and 2 moles of chalcedony are produced for each mole of serpentine dissolved. The masses of magnesite and chalcedony increase linearly with time (Fig. 6b). In particular, the magnesite mass is 625 g/kg water after 10 a and 1250 g/kg water after 20 a, and so on, indicating the Fig. 6. Moles of secondary solid phases produced during high CO2 sequestration capacity of this process. A time high-pressure CO2 (f CO2 250 bar) injection into a deep aquifer of tens of years is apparently long, but it is actually (temperature 60 C) hosted in serpentinitic rocks against (a) short if compared with the residence times of high-pH the reaction progress variable, expressed as g of dissolved waters in deep aquifers (see above). serpentine/kg of water, and (b) time. F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 797

The main solute is CO2 throughout the process, with trapping) and through incorporation in magnesite concentrations close to 4 moles/kg. The saturation indi- (mineral fixation) are plotted against both the mass of ces of chrysotile and antigorite [Mg48Si24O85(OH)62] dissolved serpentine (Fig. 8a) and time (Fig. 8b). The vary from 44.055 to 17.818 and from 648.060 to total mass of sequestered CO2, corresponding to the 226.155, respectively, indicating that the aqueous sum of these two terms, is also shown in both figures. solution is strongly undersaturated with respect to ser- Inspection of Fig. 8 shows that CO2 sequestration pentine minerals throughout the process. through dissolution in the aqueous solution is instantaneous, but relatively limited, as the maximum amount of 4.4. CO2 sequestration CO2 which can be dissolved in 1 kg of water is 180 g, at the specified conditions of 250 bar fCO2 and 60 C. The masses of CO2 sequestered during the process The mass of CO2 incorporated in precipitating magne- through dissolution in the aqueous solution (solubility site increases linearly with time, attaining values much greater than the mass of CO2 dissolved in the aqueous

Fig. 8. Masses of CO2 sequestered through solubility trapping Fig. 7. Changes in (a) the pH of the aqueous solution and (b) and mineral fixation and cumulative mass of CO2 sequestered chemical speciation, during high-pressure CO2 (f CO2 250 bar) during high-pressure CO2 (f CO2 250 bar) injection into a deep injection into a deep aquifer (temperature 60 C) hosted in ser- aquifer (temperature 60 C) hosted in serpentinitic rocks pentinitic rocks. The reaction progress variable, expressed as g against (a) the reaction progress variable, expressed as g of of dissolved serpentine/kg of water, is reported on the X-axis. dissolved serpentine/kg of water, and (b) time. 798 F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 solution, but this process requires comparatively long that: (1) the natural system might behave as an open, time intervals. For instance, the contribution of mineral flow-through system (rather than as a closed system), in fixation attains the same level of solubility trapping only which case secondary solid phases might deposit outside after 5.5 a. However, mineral carbonation becomes the reference volume affected by serpentine dissolution; more and more important beyond this time threshold, (2) fracturing induced by injection of pressurized CO2 whereas solubility trapping does not change with time. might occur, with consequent increases in the effective Again, the sequestration capacity of the process is large porosity and permeability; (3) in southern Piedmont, and time is less than the residence times of high-pH CO2 injection could be performed in the conglomerate waters in deep aquifers. layers, chiefly made up of serpentinitic clasts, which occur at the top of the serpentinites and below the rela- 4.5. Changes in the porosity of aquifer rocks tively impervious sedimentary rocks of the Piedmont Tertiary Basin (Roberto Bencini, written communi- To evaluate the changes in the porosity of aquifer rocks in response to high-pressure (fCO2 250 bar) CO2 injection the authors continue to make reference to 1 kg of water. In this way, the model is kept completely independent of the effective porosity of the aquifer. For the present purpose, it is useful to take into account the ratio of the solid volume of products to reactant, ÁV/V, equal to: ÀÁ DV=V ¼ 100 Â ÆnpVp ÆnrVr =ÆnrVr; ð12Þ where Vi and ni represent the molar volume and the moles of the i-th solid phase, respectively, and subscripts p and r indicate products and reactants, respectively. In this simplified case, the term ÆnrVr represents the volume of serpentine dissolved. During the first stages of the process (for x <0.07 g/ kg) the ÁV/V is 100% as serpentine dissolves con- gruently, i.e., without precipitation of secondary solid phases (Fig. 9a). Formation of chalcedony brings about an increase in ÁV/V to 58%. When magnesite begins to precipitate ÁV/V increases again and finally approa- ches +19%, which is the limiting value constrained by dissolution of serpentine accompanied by stoichiometric precipitation of chalcedony and magnesite (reaction 9). Values of ÁV/V close to 16–17% are attained in 1a (Fig. 9b). Obviously, for ÁV/V=0, the serpentine dissolved is substituted by an equal volume of secondary solid phases and the porosity of the system does not change. Unfortunately, in the case under consideration, the ÁV/V is positive and significantly different from zero. Consequently, serpentine dissolution accompanied by precipitation of magnesite and chalcedony causes a progressive reduction of the effective porosity of the aquifer. This would drop to zero if its initial value were 19%, whereas effective porosities >0 are only possible for initial values >19%. If precipitation of amorphous silica takes place instead of chalcedony, the reduction in porosity could be even larger, with a ÁV/V of 30.9%. Hydromagnesite precipitation instead of magnesite is Fig. 9. Variations of the ratio of the solid volume of products unlikely, as magnesite is the thermodynamically stable to reactant, ÁV/V during high-pressure CO2 (f CO2 250 bar) 9.98 Mg-carbonate for PCO2 >10 bar at 60 C. injection into a deep aquifer (temperature 60 C) hosted in Although an initial porosity of 19% seems to be too serpentinitic rocks against (a) the reaction progress variable, high a value for serpentinites, it must be emphasized expressed as g of dissolved serpentine/kg of water, and (b) time. F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 799 cation); clogging of pore spaces should not be a major dominated by serpentine can sequester more CO2 than a problem in these conglomerates, due to their high por- system of forsteritic olivine. osity and permebility. As a matter of fact, magnesite veins with nodules of opal are common in the serpentinitic rocks of Southern 6. Conclusions Tuscany (Arisi Rota et al., 1971), a region characterized by high thermal fluxes (>100 mW m2, Baldi et al., The geochemical data on spring waters, gathered 1994) and high CO2 fluxes (Marini and Chiodini, 1994), through an extensive survey of the Gruppo di Voltri indicating that reaction 9 occurred naturally in these area, confirm that progressive interaction between environments ultramafic rocks variably affected by serpentinization and meteoric waters produces Mg–HCO3 waters first, in shallow aquifers open to CO2 exchange, followed by the 5. Discussion development of Na–HCO3 and Ca–OH type waters, under closed-system conditions with respect to CO2. Reactive transport modeling of CO2 injection into 3 Tritium data indicates that these evolved, high-pH different geological frameworks, including a dunite, was groundwaters experience prolonged circulation in com- carried out by Xu et al. (2000), by means of TOUGH- paratively deep aquifers hosted by serpentinites. These REACT (Xu and Pruess, 1998), a code for modeling groundwaters constitute the only evidence of the exis- geochemical reactive transport in non-isothermal multi- tence of the deep aquifers they come from. phase systems. The dunite case is relatively similar to the In the light of reaction path modeling of high-pres- present case and is briefly presented here. Xu et al. sure CO2 injection into deep serpentinitic aquifers, this (2000) have considered a system made up of forsterite methodology of CO2 sequestration appears to represent (85% by volume) and fayalite (9.5 by volume) with a a feasible action to reduce anthropogenic CO2 inputs porosity of 5%. They assumed dissolution rates at 25 C into the atmosphere. Serpentinitic rocks and ultra- 13 2 1 of 10 moles m s for both forsterite and fayalite mafites have a high CO2 sequestration capacity, mainly and precipitation rates at 25 Cof1012 moles m2 s1 through mineral fixation as magnesite and subordinately for all non-carbonate secondary minerals and of through solubility trapping. The natural occurrence of 0.6Â108 moles m2 s1 for magnesite and siderite. dissolution of ultramafic rocks, variably affected by ser- Activation energy, Ea, was set to 62.76 kJ/mole for all pentinization, accompanied by precipitation of magne- solid phases except magnesite and siderite, whose Ea site and silica minerals, in areas of high terrestrial CO2 was assumed to be 41.87 kJ/mole. Initial surface areas fluxes such as Southern Tuscany, represents field evi- (in m2/dm3) were fixed to 8.55 for forsterite and 0.95 for dence supporting the feasibility of this methodology of fayalite, which corresponds to assuming a total surface CO2 sequestration. However, dissolution of serpentine area of 10 m2/dm3 and to fractionate it based on the accompanied by precipitation of magnesite and chal- volumetric abundances of these two primary phases. cedony brings about a progressive reduction in the Upon injection of CO2 with a pressure of 260 bar, the effective porosity of the aquifer, at least under closed- pH of the aqueous solutions attains a stable value of 4.8 system conditions, and the situation could be much (which compares favourably with the authors’ value, worse if amorphous silica precipitates instead of chal- 4.5, see above), forsterite and fayalite dissolve and cedony. In addition, as the magnesite and silica pre- magnesite, siderite, talc and amorphous silica precipitate in the serpentinitic aquifer, these minerals will 3 cipitate. Referring to a system of 1 m , 100 kg CO2 likely armor the remaining reactants, thereby further are fixed as carbonate minerals, mainly magnesite, in a decreasing effective surface area in addition to decreas- lapse of time of 1 ka. This process determines a sig- ing porosity. nificant decrease in porosity from 5 to 0.6% and this These effects could represent serious obstacles for porosity reduction could bring about the end of mineral the implementation of this methodology of CO2 alteration and CO2 sequestration. However, the domi- sequestration and their importance must be evaluated nant reaction considered in the model by Xu et al. by means of laboratory experiments first and field tests (2000): afterwards.

0:8Mg2SiO4 þ CO2 þ 0:2H2O

¼ MgCO3 þ 0:2Mg3Si4O10ðÞOH 2 ð13Þ Acknowledgements is characterised by a ÁV/V of +57.8%, which is much The authors gratefully acknowledge Halldo´ r higher than that of serpentine dissolution accompanied A´ rmannsson, Stefa´ n Arno´ rsson, and Mark H. Reed for by precipitation of magnesite and chalcedony (reaction helpful comments which greatly improved this paper. 9). Therefore, for the same initial porosity, an aquifer This work has been carried out within the framework of 800 F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 the MIUR Project of Relevant National Interest P V n ¼ CO2 L ðA1Þ ‘‘GeoCO ’’. CO2;L 2 KH

where KH represents Henry’s constant of CO2 in bars/ Appendix. Analytical method for the determination of (mol/kg) at the laboratory temperature TA, which was TDIC obtained from the solubility data of Wilhelm et al.

(1977). The moles of CO2 in the head space gas, nCO2;G Due to the very low TDIC concentrations of high-pH, are calculated by the equation (pressure in bars, volume Ca–OH waters, it is impossible to compute this para- in liters, temperature in Kelvin): meter from pH and alkalinity, as it is usually done for P V 273:15 n ¼ CO2 G ðA2Þ most natural waters (Bruni et al., 2001; 2002). The need CO2;L 1:013 22:414 T to change the consolidated procedure of computing A TDIC from pH and alkalinity, by means of suitable Then the TDIC molality is obtained as follows: speciation calculations, was already underscored by nCO2;L þ nCO2;G Wolery (1992). In this way, in fact, significant errors are mTDIC ¼ ðA3Þ VL sometimes introduced in the computed TDIC and, consequently, in the subsequent interpretation of geochem- TDIC concentration is expressed as mg of HCO3 ion/l ical data. These errors are huge for high-pH, Ca–OH in Table 1. waters, since alkalinity virtually represents a measure- ment of the concentration of hydroxyl ion, or a sort of duplicate pH determination. Speciation calculations References carried out by means of EQ3NR (Wolery, 1992) show, in fact, that in high-pH Ca-OH waters, total alkalinity is Appelo, C.A.J., Postma, D., 1996. Geochemistry, Ground- mainly explained by hydroxyl ion (61–98%) and sub- waters and Pollution. A.A. Balkema, Rotterdam. 2 ordinately by HCO3 and CO3 ions (0.4–36%), whereas Arisi Rota, F., Brondi, A., Dessau, G., Branzini, M., Stea, B., the contributions of silicate-, borate-, and sulfide-species Vighi, L., 1971. I giacimenti minerari. In ‘‘La Toscana Mer- idionale. Fondamenti geologico-minerari per una prospettiva is generally negligible, apart from the SiO2-rich sample 2 di valorizzazione delle risorse naturali’’. Rendiconti S.I.M.P. V18. In this case the H3SiO4 and H2SiO4 ions con- tribute 24% of total alkalinity. The difficulty in deter- 27, 357–544. Bachu, S., Gunter, W.D., Perkins, E.H., 1994. Aquifer disposal mining TDIC of high-pH (about 11–12) Ca-OH waters of CO2: hydrodynamic and mineral trapping. Energy Con- is confirmed by the lack of TDIC in old analyses vers. Mgmt. 35, 269–279. (Barnes et al., 1967, 1972, 1978; Barnes and O’Neil, Baldi, P., Bellini, S., Ceccarelli, A., Fiordalisi, A., Squarci, P., 1969; Pantazis, 1976; Papastamataki, 1977; Pfeifer, Taffi, L., 1994. Correlazioni fra le anomalie termiche ed altri 1977; Neal and Stanger, 1983), preceeding the work of elementi geofisici e strutturali della Toscana Meridionale. Bruni et al. (2001, 2002). Studi Geol. Camerti. 1, 139–149. Following Bruni et al. (2001, 2002), to measure TDIC Bales, R.C., Morgan, J.J., 1985. Dissolution kinetics of chryso- in a reliable way, water was sampled into evacuated tile at pH 7 to 10. Geochim. Cosmochim. Acta 49, 2281–2288. glass bottles (equipped with a 3-way valve) containing a Barnes, I., O’Neil, J.R., 1969. The relationship between fluids few ml of concentrated HCl, but in excess with respect in some fresh alpine-type ultramafics and possible modern serpentinization, western United States. Geol. Soc. Am. Bull. to total alkalinity. In the laboratory, bottles were pre- 80, 1947–1960. pared and weighed before sampling. In the field, after Barnes, I., LaMarche Jr, V.C., Himmelberg, G., 1967. Geo- flushing the 3-way valve with spring water, this was chemical evidence of present-day serpentinization. Science made to enter the bottle, leaving a head space of 156, 830–832. approximately 1/4 of the bottle volume. Hydrochloric Barnes, I., Rapp, J.B., O’Neil, J.R., Sheppard, R.A., Gude, acid converts all the carbonate species to CO2, which A.J., 1972. Metamorphic assemblages and the direction of accumulates in the head space, together with other gas flow of metamorphic fluids in four instances of serpentiniza- constituents. In the laboratory, each bottle was weighed tion. Contrib. Mineral. Petrol. 35, 263–276. Barnes, I., O’Neil, J.R., Trescases, J.J., 1978. Present-day to determine the volume of sampled water, VL. Know- ing V , the head space gas volume, V , is obtained by serpentinization in New Caledonia, Oman and Yugoslavia. L G Geochim. Cosmochim. Acta 42, 144–145. difference between the volume of the empty bottle and Berner, R.A., 1981. Kinetics of weathering and diagenesis. In: V . Then, the P in the head space gas was measured L CO2 Lasaga, A.C., Kirkpatrick, R.J. (Eds.), Kinetics of Geo- by gas-chromatography. Assuming equilibrium dis- chemical Processes, Reviews in Mineralogy, 8, 111-134. tribution of CO2 between the gas and liquid phases, the Berner, R.A., Lasaga, A.C., Garrels, R.M., 1983. The carbon- moles of CO2 dissolved in the liquid, nCO2;L are com- ate-silicate geochemical cycle and its effect on atmospheric puted by means of the relation (pressure in bars, volume carbon dioxide over the past 100 million years. Am. J. Sci. in liters): 283, 641–683. F. Cipolli et al. / Applied Geochemistry 19 (2004) 787–802 801

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