# Stable and radiogenic strontium isotope fractionation during hydrothermal seawater-basalt interaction

Martin Voigt, Christopher R. Pearce, Andre Baldermann, Eric H. Oelkers

► To cite this version:

Martin Voigt, Christopher R. Pearce, Andre Baldermann, Eric H. Oelkers. Stable and radiogenic strontium isotope fractionation during hydrothermal seawater-basalt interaction. Geochimica et Cosmochimica Acta, 2018, 240, pp.131-151. 10.1016/j.gca.2018.08.018. insu-03661390

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# Stable and radiogenic strontium isotope
fractionation during hydrothermal
seawater-basalt interaction

Martin Voigt<sup>a,*</sup>, Christopher R. Pearce<sup>b</sup>, Andre Baldermann<sup>c</sup>, Eric H. Oelkers<sup>a,d</sup>

<sup>a</sup> Géosciences Environnement Toulouse (GET) - CNRS, 14 avenue Édouard Belin, 31400 Toulouse, France (*correspondence: martin.voigt@get.obs-mip.fr)

<sup>b</sup> National Oceanography Centre Southampton, University of Southampton Waterfront Campus, European Way, Southampton SO14 3ZH, United Kingdom

<sup>c</sup> Institute of Applied Geosciences, Graz University of Technology, Graz, Austria

<sup>d</sup> Earth Sciences, UCL, Gower Street, London WC1E6BT, United Kingdom

## Abstract

The fluid-rock interactions occurring in hydrothermal systems at or near mid-oceanic ridges (MOR) were studied experimentally by reacting crystalline and glassy basalt with seawater at 250 and 290 °C while monitoring the fluid phase Sr isotopic evolution (<sup>87</sup>Sr/<sup>86</sup>Sr and δ<sup>88</sup>/<sup>86</sup>Sr). The results indicate that seawater Sr was incorporated into anhydrite during the early stages of seawater-basalt interaction. Fluid <sup>87</sup>Sr/<sup>86</sup>Sr values trend towards the basaltic signature as non-stoichiometric basalt dissolution became the dominant process. This suggests that the interplay between fast Sr incorporation into secondary sulfates versus slow and continuous Sr liberation due to basalt dissolution at intermediate temperatures could partly explain previously identified discrepancies between MOR heat budget constraints and the marine <sup>87</sup>Sr/<sup>86</sup>Sr budget. Late-stage anhydrite re-dissolution, likely caused by the fluid phase becoming more reducing through further basalt dissolution, as well as by quenching of the

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26 experiments, represents a potential explanation for the low amounts of anhydrite found in
27 naturally altered oceanic basalt samples. Relatively strong decreases in fluid δ<sup>88/86</sup>Sr values in
28 experiments with crystalline basalt suggest that isotopically light Sr was preferentially
29 released due to non-stoichiometric dissolution. A slight preference of anhydrite for
30 isotopically heavy Sr (ε<sup>88/86</sup><sub>Anhydrite-Liquid</sub> = 0.33 ± 0.020 ‰) is indicated by the data,
31 suggesting that changes in MOR spreading rates and Sr removal could be recorded in the
32 isotope compositions of authigenic, sedimentary Sr phases. Such insights will help to
33 constrain the influence of hydrothermal systems on the oceanic stable Sr cycle.

34 **1 Introduction**

35 Widespread hydrothermal activity at the mid-ocean ridges is well established (e.g., Stein
36 and Stein, 1994; Alt and Teagle, 2000; Baker and German, 2004; German and Seyfried,
37 2014). Driven by magmatic heat associated with the generation of new crust, seawater is
38 heated while percolating through the oceanic crust, and, after its interaction with this basaltic
39 crust, released again to the ocean (e.g., Sleep, 1991; Bickle et al., 1998; Coogan, 2008). Fluids
40 expelled near mid-ocean ridges, where new oceanic crust is created through submarine
41 volcanic activity, can reach temperatures exceeding 400 °C, while further from active
42 spreading centres, low-temperature hydrothermal fluids, driven by heat transferred from the
43 upper mantle, are released from the crust (e.g., Elderfield and Schultz, 1996). The significance
44 of marine hydrothermal activity on marine geochemical cycles has been widely discussed
45 (e.g., Palmer and Edmond, 1989; Elderfield and Schultz, 1996; Mottl, 2003; Davis et al.,
46 2003; Coogan and Dosso, 2012; German and Seyfried, 2014): high alteration temperatures
47 near mid-ocean ridges, and the large areas and residence times of low-temperature alteration
48 lead to chemical fluxes of similar magnitude as riverine fluxes. However, the details of
49 chemical fluxes from these systems to the oceans, and notably the partitioning of heat
50 transfer, fluid flow and chemical fluxes between high-temperature, near-axis hydrothermal
systems, and low-temperature off-axis venting, remain a subject of active debate (e.g., Mottl, 2003; Davis et al., 2003; Nielsen et al., 2006; German and Seyfried, 2014).

Strontium (Sr) is a particularly useful tracer of the sources and sinks of marine geochemical cycles, as the presence of both radiogenic and stable isotope systems enables different processes to be resolved (Faure and Hurley, 1963; Krabbenhöft et al., 2010; Pearce et al., 2015). For example, due to ⁸⁷Sr production by ⁸⁷Rb decay combined with the higher Rb/Sr ratios and age of the continental crust (Armstrong, 1968), the continental crust has relatively high ⁸⁷Sr/⁸⁶Sr ratios (where ⁸⁷Sr/⁸⁶Sr designates the traditional radiogenic Sr ratio corrected to a constant ⁸⁶Sr/⁸⁸Sr value of 0.1194; cf. Nier, 1938) compared to the unradiogenic mantle and oceanic crust. These large isotopic differences enable the use of seawater ⁸⁷Sr/⁸⁶Sr compositions to assess variations in marine Sr sources (Nier, 1938; Faure and Hurley, 1963). The major Sr sources to the oceans are river and groundwater discharge from the continents, which provide radiogenic ⁸⁷Sr/⁸⁶Sr signatures to the oceans, and marine hydrothermal fluids that have unradiogenic ⁸⁷Sr/⁸⁶Sr compositions derived from the basaltic oceanic crust (Davis et al., 2003; Krabbenhöft et al., 2010; Pearce et al., 2015). It has been suggested that high-temperature hydrothermal circulation is not sufficient to balance the other Sr input fluxes given the observed increase of ~5.4×10⁻⁵ Myr⁻¹ (Hodell et al., 1990) in oceanic ⁸⁷Sr/⁸⁶Sr over the last few million years (Palmer and Edmond, 1989; Butterfield et al., 2001; Davis et al., 2003). Various explanations have been proposed to provide the missing unradiogenic ⁸⁷Sr/⁸⁶Sr flux, such as island arc weathering (Allègre et al., 2010; Li and Elderfield, 2013), ridge-flank hydrothermal systems (Butterfield et al., 2001), low-temperature alteration and carbonate precipitation (Coogan and Dosso, 2015), temporal variability in continental weathering (Vance et al., 2009; Krabbenhöft et al., 2010), and the dissolution of river derived basaltic particles (Jones et al., 2012b; Jones et al., 2012a; Jones et al., 2014; Jeandel and Oelkers, 2015). Furthermore, it has been suggested that variations of paleoseawater compositions significantly affected the exchange of Sr in hydrothermal systems (Antonelli et al., 2017).
There are, nevertheless, considerable uncertainties and debate in the magnitude of these processes and fluxes, a well as the question if the marine radiogenic Sr system is presently in steady state (Mokadem et al., 2015).

In contrast to ⁸⁷Sr/⁸⁶Sr, stable Sr isotope ratios (δ⁸⁸/⁸⁶Sr) are affected by both mixing of sources with distinct δ⁸⁸/⁸⁶Sr values and isotopic fractionation during precipitation. At isotopic steady state, the δ⁸⁸/⁸⁶Sr composition of the marine Sr input and output fluxes should be equal. However, it has been demonstrated that the δ⁸⁸/⁸⁶Sr signature of marine carbonates (the dominant Sr sink in the oceans) is significantly lower than the flux-weighted input δ⁸⁸/⁸⁶Sr composition (Krabbenhöft et al., 2010; Pearce et al., 2015). This implies either an imbalance in the seawater δ⁸⁸/⁸⁶Sr (i.e. the modern system is not in isotopic steady state), and/or that the δ⁸⁸/⁸⁶Sr composition of oceanic sources and sinks, such as marine hydrothermal systems, is not well quantified at present. In general, stable Sr isotopes potentially provide an additional tool to resolve ambiguities arising from the ⁸⁷Sr/⁸⁶Sr data such the reason for increasing seawater ⁸⁷Sr/⁸⁶Sr (e.g., changes in the riverine flux vs. its composition; Krabbenhöft et al., 2010). However, persisting uncertainties and inconsistencies in both ⁸⁷Sr/⁸⁶Sr and δ⁸⁸/⁸⁶Sr marine budgets compels further assessment of the marine Sr sources and sinks. In this context the reactions occurring in seafloor hydrothermal systems are of critical importance. They are thought to represent the dominant source of unradiogenic ⁸⁷Sr/⁸⁶Sr to seawater, yet the size of this Sr flux to the oceans from high- and especially low-temperature systems remains poorly quantified (Davis et al., 2003; Coogan and Dosso, 2015). Furthermore, the potential effects of Sr sequestration by precipitation of secondary sulfates/carbonates in such systems on oceanic δ⁸⁸/⁸⁶Sr are currently unconstrained (Pearce et al., 2015).

The reactions leading to changes in the chemical composition of seawater during hydrothermal alteration have been extensively studied using experimental systems (e.g., Bischoff and Dickson, 1975; Seyfried and Bischoff, 1977; Mottl and Holland, 1978; Seyfried and Bischoff, 1979; Mottl et al., 1979; Seyfried and Bischoff, 1981; Seyfried and Mottl, 1982;
Seewald and Seyfried, 1990; Luhmann et al., 2017a; Luhmann et al., 2017b). Experimental constraints on the extent of isotopic fractionation during such hydrothermal reactions are, however, sparse (Menzies and Seyfried, 1979; Berndt et al., 1988; James et al., 2003). This limits the interpretation of the behaviour of isotopic tracers such as radiogenic $^{87}$Sr/$^{86}$Sr ratios or $\delta^{88/86}$Sr in natural systems and introduces significant uncertainties in marine Sr mass balance models as described above. This study was designed to overcome some of these limitations through a series of experiments reacting seawater with basaltic glass and crystalline basalt at temperatures of 250 °C and 290 °C. The interpretation of measured fluid phase Sr isotopic ratios during these experiments provides new insight into the behaviour of Sr and its isotopic ratios during seawater-basalt interaction.

## 2 Methods

### 2.1 Experimental design

Basaltic glass and crystalline basalt were reacted with seawater in 400 mL Ti-autoclaves (Fig. 1) at temperatures of 250 °C and 290 °C for 53 and 92 days, respectively. Two experiments were conducted at each temperature using either the basaltic glass or the crystalline basalt as solid starting materials (Table 1). As the autoclaves were not pressurised before heating, pressure at the experimental temperature was constrained by the liquid-gas phase relations. As pressure was not directly monitored, it was estimated approximating the seawater as a H₂O-NaCl solution with a molar NaCl fraction of 0.0106 (i.e. allocating all Cl⁻ to NaCl), allowing the thermodynamic equilibrium state of this H₂O-NaCl system to be predicted using the relationships in Driesner and Heinreich (2007) and Driesner (2007). Using the temperature, volume and mass constraints from our experiments, a NaCl bearing liquid phase and minor amounts of a low-NaCl vapour phase were predicted to coexist. For our experiments, pressures of 39 bar and 73 bar were calculated for the 250 °C and 290 °C experiments, respectively, which is similar to pure H₂O vapour pressures. Liquid sampling
during the experiments did not significantly lower the pressures in the reactors (for example, pressures are predicted to drop from 73.13 bar to 72.94 bar during the 290 °C experiments), but the volume changes did affect the concentration of aqueous species in some cases (see section 3.1.4).

All reactor parts in contact with the experimental materials were constructed from Ti (UT 40) with exception of a polytetrafluoroethylene (PTFE) gasket between the bottom of the reactor and the closure of the autoclave (Fig. 1). Before each experiment, the autoclaves were mechanically cleaned by re-machining the inner walls, filling with ultrapure water (18.2 MΩ×cm at 25 °C, Millipore) and heating to 200 °C for several days, then repeating this procedure (including the heating of the reactor) with ~1 mol/kg bi-distilled HNO₃ and finally with ultrapure water.

At the beginning of each experiment, the reactors were filled with basalt and seawater so that the seawater to basalt mass ratio was ~10 (Table 1) consistent with previous experiments and similar to conditions in submarine hydrothermal systems (Seyfried and Bischoff, 1979; Seyfried and Bischoff, 1981). The mass of the reactants were lower for the 290 °C experiment to account for the greater thermal expansion at this temperature. Before sealing the reactors, the head space was flushed with N₂ to avoid oxidation reactions caused by an uptake of O₂ from the gas phase. The experiments were started by placing the reactors into the furnaces, starting the rocking motion (~1 cycle/min) and heating the furnaces to the desired temperature. The desired experimental temperature was reached within approximately 2 h in all experiments. The temperature inside the reactors was continuously monitored with a type K thermocouple inserted into a thermowell in the reactors with an uncertainty of approximately ± 2 °C.

Samples of the liquid phase in the reactors were taken during all experiments using a valve with an attached water-cooling apparatus (Fig. 1). The rocking motion was stopped prior to sampling and a Ti-frit, located at the sampling outlet in the reactor (pore size 5 µm),
prevented solids from leaving the autoclave. The liquid was filtered through a 0.45 µm
cellulose acetate membrane at the outlet (replaced after each sampling) and collected in three
separate polypropylene (PP) vials: The first ~3 mL collected during each sampling were
discarded to avoid cross-contamination from the sampling device (which was cleaned using
dilute bi-distilled HNO₃ and ultrapure water before and after each sampling). Subsequently, a
non-acidified sample (~3 mL) was taken for pH measurements, followed by a ~15 mL sample
for later compositional and isotopic analysis collected in a vial containing ~0.45 g of bi-
distilled HNO₃ to avoid precipitation of solid phases due to the temperature and pressure
changes after quenching of the sample.

The experiments were terminated by quenching the reactors in water so that they
attained ambient temperature within approximately 30 minutes. After opening the reactors,
the solid reaction products were separated from the liquid by vacuum filtration through a
0.45 µm cellulose acetate membrane, rinsed with ultrapure water, and dried at 50 °C.

## 2.2 Reactants

Two different types of basalt were used as reactants in the experiments: basaltic glass, and crystalline basalt (Fig. 2a, b). Both rock types were collected from a basaltic dyke on Stapafell Mountain, SW Iceland, and have been previously described by Oelkers and Gislason (2001) and Gudbrandsson et al. (2011). The chemical compositions of these samples (Table 2) are similar to that of mid-ocean ridge basalt (MORB). Their Sr content and isotopic compositions were determined in this study, and correspond to typical basaltic values (Table 3). The crystalline basalt is composed of approximately 43 wt.% plagioclase, 47 wt.% pyroxene and 9 wt.% olivine, while only minor amounts of olivine (~3 wt.%) are present in the basaltic glass (~97 wt.% amorphous phase) according to quantitative XRD analyses. Both materials were ground using a jaw crusher and an agate disc mill before the 40-80 µm size fraction was obtained by sieving. Ultra-fine particles were removed from these powders by repeated ultrasonic cleaning in ultrapure water, followed by gravitational settling and
decantation of the supernatant water. The resulting powder was dried at 60 °C. The specific surface areas of these powders were determined to be 0.07 m²/g and 0.4 m²/g for the basaltic glass and crystalline basalt, respectively, by the BET technique (Brunauer et al., 1938) using Kr gas and a Quantachrome Autosorb 1-MP automated gas sorption analyser with a relative uncertainty (one standard deviation) of approximately 10%. The difference in the surface area of the two materials likely arises from differences in surface roughness, as the grain sizes are similar for both powders.

Natural seawater sampled from the equatorial Pacific Ocean from a depth of 350 m (EqPac cruise November 1992; Murray et al., 1997) was used as the initial fluid reactant for all experiments. The seawater was acidified upon collection using distilled concentrated HCl (to a pH of ~ 1.5) to prevent the precipitation of solids and growth of organic material. The pH of the seawater was re-adjusted to a natural value of ~8.14 by the gradual addition of 25 % ammonia solution (EMD Millipore Suprapur) prior to its use in the experiments (Table 4), but was not equilibrated with atmospheric CO₂ as it was filled into the reactors immediately whose headspace was subsequently flushed with N₂, so that its dissolved inorganic carbon concentrations remained low (most likely on the order of 10⁻⁵ to 10⁻⁴ mol/kg). The total ammonia and HCl concentrations were approximately 0.04 mol/kg and 0.05 mol/kg, respectively, the former potentially impacting pH evolution (cf. section 4.1).

## 2.3 Analytical methods

All liquid samples were analysed for their pH, elemental compositions and Sr isotopic composition. Fluid pH was measured using a Metrohm microelectrode at 22±1 °C. The electrode was calibrated against three NIST traceable pH buffer solutions (pH 4.0, 6.9 and 9.2) to achieve an analytical precision of ± 0.07 pH units based on replicate analyses of buffers and samples.

Liquid samples from the 250 °C experiments were analysed for their total Na, Mg, Si, K, Ca, Sr, and Fe contents using a ThermoFisher inductively coupled plasma atomic emission
spectrometer (ICP-AES), while Al contents were determined by furnace atomic adsorption spectrometry (AAS) using a Perkin Elmer AAnalyst 600 (analytical uncertainties and detection limits are reported in Electronic Supplementary Table 1). Concentrations of Mg, Si, K, Ca, Sr, Fe and Al in the liquid samples from the experiments at 290 °C were determined by inductively coupled plasma mass spectrometry (ICP-MS) using an Agilent 7500ce after adding In and Re as internal standards. For these samples, Na contents were determined using a Perkin Elmer AAnalyst 400 flame AAS. The seawater used as reactant was analysed by all of the above mentioned techniques, and the concentrations determined by the different techniques are in agreement with each other within the uncertainties. Aqueous SO₄²⁻ and Cl⁻ of all liquid samples were determined using a Dionex ICS 2000 ion chromatography (IC) system. All calibration standards were prepared from aqueous Merck Certipur solutions.

The ⁸⁷Sr/⁸⁶Sr and δ⁸⁸/⁸⁶Sr ratios of both solid and liquid samples were determined by thermal ionisation mass spectrometry (TIMS) using a ThermoFisher Triton at the National Oceanography Centre, Southampton following the sample preparation and analytical procedures described in Pearce et al. (2015). For each sample, an unspiked and a spiked aliquot (containing an additional ⁸⁴Sr-⁸⁷Sr solution of known isotopic composition) were analysed, allowing correction for mass fraction during sample preparation and analysis. Radiogenic ⁸⁷Sr/⁸⁶Sr ratios were determined by correcting the data from the measurements of the unspiked sample for instrument mass fractionation assuming the exponential mass fractionation law and using an ⁸⁶Sr/⁸⁸Sr ratio of 0.1194. The raw mean ⁸⁷Sr/⁸⁶Sr ratio of NIST SRM 987 during this study was 0.710251 (2 standard deviations (SD) = 7 ppm), though all reported ⁸⁷Sr/⁸⁶Sr compositions have been corrected to an external NIST SRM 987 ⁸⁷Sr/⁸⁶Sr value of 0.710248 (McArthur et al., 2012) for comparison with published data. Over the course of this study, the NIST SRM 987 standard was analysed 16 times, with 4 measurements at the beginning of each session with a total of 15 to 21 measurements (half of these standards were processed using the same preparation methods as for the samples). No
significant drift for stable Sr isotope measurements of standards was observed over the course
of these measurements, while radiogenic standard measurements decreased slightly by ~ 7
ppm towards the end of the measurements. Stable Sr isotope ratios were determined from
deconvolution of the spiked and unspiked measurements using the exponential mass
fractionation law and a Newton-Raphson optimisation (e.g., Ypma, 1995). The δ<sup>88/86</sup>Sr values,
defined as δ<sup>88/86</sup>Sr = (δ<sup>88Sr/86Sr</sup>)<sub>Sample</sub> - 1 and given in units of per mil (‰, i.e. multiplied by
1000), are given relative to the mean of the repeated analyses of the NIST SRM 987 standard.

The two basalt powders used as reactants and the solid reaction products recovered
from the reactors were analysed by scanning electron microscopy (SEM) using a Jeol JSM
6360LV equipped with an electron dispersive spectroscope (EDS). Powder X-ray diffraction
(PXRD) patterns of the unreacted basalt powders and of the solids recovered from the reactors
after the experiments were recorded on a PANalytical X'Pert PRO diffractometer (Co-Ka
radiation) operated at 40 kV and 40 mA. This diffractometer is outfitted with a high-speed
Scientific X'Celerator detector, 0.5° antiscattering and divergence slits, spinner stage, primary
and secondary soller and automatic sample changer. The top loading technique was used for
the preparation of the samples. The specimens were examined over the range 4-85° 2-θ using
a step size of 0.008° 2-θ and a count time of 40 s per step. Phase identification and
quantification was performed using the PANalytical X'Pert HighScore Plus software and a
pdf-2 database.

2.4 Thermodynamic calculations

The saturation states (Ω) of solid phases were calculated using PHREEQC (Parkhurst
and Appelo, 2013) together with its Lawrence Livermore National Laboratory (LLNL)
database (Johnson et al., 2000) at the experimental temperature and saturated vapour pressure.
The saturation index (SI) is defined as the base-10 logarithm of the saturation state Ω, which
is the ratio of the ion activity product to the equilibrium constant of the precipitation reaction.
The enthalpy of the aqueous MgSO₄ complex dissociation reaction was taken from the minteq.v4 database, as it is consistent with the results of Allison et al. (1991). In these calculations, the standard state for solid phases and H₂O is the pure phase, while a hypothetical 1 mol/kg aqueous solution referenced to infinite dilution is chosen for unit activity (a) of aqueous species, both at the temperature and pressure of interest. The Livermore National Laboratory aqueous model (Daveler and Wolery, 1992), which uses the extended Debye-Hückel equation (Helgeson, 1969) was used to calculate activity coefficients of the aqueous species. As the redox state of the fluids was not specifically analysed, calculations were carried out assuming oxidising conditions (i.e. aqueous S is in the form of S⁶⁺), limiting the applicability to Fe-bearing minerals. The quality of the thermodynamic database at temperatures of 200 to 300 °C under seawater-like salinities is limited such that the saturation indexes calculated for the reactive fluid samples collected in this study are somewhat uncertain. For anhydrite and quartz, minerals that contain few components, the uncertainties associated with their calculated saturation indexes are likely no more than 0.2, in contrast, for feldspars and clay minerals, which have more complex hydrolysis reactions, such uncertainties likely exceed 1 (e.g., Voigt et al., in review).

## 2.5 Uncertainties

Unless otherwise stated, uncertainties in this study represent the 95% confidence interval (CI). The ± symbol is used to denote the half-width of the confidence interval (for normally distributed values, this corresponds to approximately 1.96 standard errors (SE)). Where possible, uncertainties were propagated through functions using linearisations. For non-algebraic or highly non-linear functions, Monte Carlo simulations were used to estimate the transformed probability distribution and the confidence interval is calculated from its quantiles (e.g., Ogilvie, 1984).
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# 3 Results

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## 3.1 Evolution of the reactive fluid compositions during the experiments

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### 3.1.1 Magnesium, sulfate, calcium and potassium

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The chemical composition of the fluids sampled from the reactors changed significantly during the experiments (Figs. 3 and 4, Table 4). Magnesium and $SO_4^{2-}$ were almost completely removed from the starting seawater solution within one week of reaction with basaltic glass at both studied temperatures. This consumption was slower in the crystalline basalt experiments, such that near complete Mg removal was only achieved after three weeks at 250 °C. Calcium concentrations remained stable or decreased slightly at the beginning of the experiments, but increased during the later stages in all experiments, attaining values of approximately 42 mmol/kg (the only exception being the experiment with basaltic glass at 290 °C, where 79 mmol/kg Ca was present in the final liquid sample). The highest increases of Ca coincided with greatest Mg removal, i.e. an inverse correlation of Ca and Mg trends can be observed. Similarly to Ca, aqueous K concentrations increased during all experiments, reaching values of up to 24 mmol/kg in the 290 °C basaltic glass experiment, and approximately 17 mmol/kg in the other experiments (Table 4).

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### 3.1.2 pH

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The ambient temperature pH in the basaltic glass experiment only showed a slight decrease at 250 °C before its steady increase to approximately 9.5 at both temperatures (Figs. 3 and 4, Table 4). Conversely, the pH attained acidic values of 3.7 and 4.9 during the first month of the crystalline basalt experiments at 250 °C and 290 °C, respectively, before increasing to a final pH of approximately 9.2.

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### 3.1.3 Strontium

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After dropping during the first 24 h, aqueous Sr concentrations increased throughout the experiments, similar to the behaviour of aqueous Ca (Fig. 5, Table 4). In the basaltic glass
experiments, the radiogenic ⁸⁷Sr/⁸⁶Sr isotope ratios decreased from the initial seawater value (0.709173) and attained values similar to that of the basaltic glass (0.703193) within one day. This process was slower in the crystalline basalt experiment at 290 °C, where similar ⁸⁷Sr/⁸⁶Sr ratios were reached within 4 days. In the corresponding 250 °C experiment the largest decrease in ⁸⁷Sr/⁸⁶Sr occurred between days 7 and 21, but values showed a higher variability and remained above 0.704210 until the end of the experiment. During the later stages of all experiments (i.e. after 1 month), the ⁸⁷Sr/⁸⁶Sr ratios increased slightly to final values of 0.704276 and 0.704114 for the basaltic glass experiments and of 0.704862 and 0.705116 for the crystalline basalt experiments, respectively, at 250 and 290 °C. Stable δ⁸⁸/⁸⁶Sr values decreased from the initial seawater composition (0.38 ‰) within the first days in all experiments (Fig. 5, Table 4). However, the magnitude of this change varied among the experiments. In the basaltic glass experiments, δ⁸⁸/⁸⁶Sr values similar to that of the bulk basalt (0.23 ‰) were attained. Conversely, in the crystalline basalt experiments, seawater δ⁸⁸/⁸⁶Sr decreased to levels significantly below that of bulk crystalline basalt (0.26 ‰) after the first week. The δ⁸⁸/⁸⁶Sr decrease was most pronounced in the 290 °C experiment, where a value of -0.06 ‰ was observed temporarily. Afterwards, the δ⁸⁸/⁸⁶Sr ratios increased steadily towards the crystalline basaltic signature but always remained below this basalt value during the further evolution of these experiments.

### 3.1.4 Sodium and chloride

The dissolved Na and Cl concentrations increased steadily within the first weeks and more rapidly during the late stages of the experiments due to fluid sampling and related formation of a low-NaCl vapour phase (see section 2.2). This interpretation is validated by the composition of the quenched liquid recovered from the basaltic glass experiment at 250 °C after the experiment was terminated, showing Cl⁻ concentrations returning to similar values as those taken during the early stages of the experiment (see Table 4). Due to this process, the salinity increased significantly in the last two fluid samples collected from the basaltic glass
experiment at 250 °C (≈17 % in terms of Cl⁻ compared to seawater, cf. Figs. 3 and 4, Table 4) and of the last samples collected from both experiments at 290 °C. No significant Cl was released from the solids in our experiments. Only small amounts of H₂O were taken up by hydrated solid alteration products; however, this process further enriches the Na and Cl concentrations in the fluids. The uptake of H₂O by the vapour phase on the Na and Cl concentration of the liquid phase was calculated using the relationships of Driesner and Heinreich (2007) and Driesner (2007), which was accounted for by the mass balance calculations and models discussed below.

### 3.1.5 Silica

Dissolved silica concentrations increased dramatically from the below detection quantities of the initial seawater during the first day of the experiments (Figs. 3 and 4, Table 4). At 250 °C, silica concentrations increased to 14 mmol/kg in the basaltic glass experiment and to 20 mmol/kg in the reaction with crystalline basalt, before stabilising at approximately 10 mmol/kg in both experiments. Similar trends were observed at 290 °C, but higher absolute concentrations (approximately 25 mmol/kg) were attained before a final decrease to 10 mmol/kg and 15 mmol/kg in the basaltic glass and crystalline basalt experiments, respectively.

### 3.1.6 Iron

The basaltic glass experiment at 250 °C reached an Fe concentration of 0.27 mmol/kg after one day before dropping to non-detectable levels, while in the corresponding experiment at 290 °C no dissolved Fe was detectable in any fluid samples (Figs. 3 and 4, Table 4). Higher Fe concentrations were reached in the experiments with crystalline basalt (2.4 mmol/kg and 4.7 mmol/kg at 250 °C and 290 °C, respectively), and concentrations above 1 mmol/kg were sustained in the lower temperature experiment for two weeks.
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**3.2 Solid alteration products**

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The solids recovered at the end of the experiments were substantially altered relative to the starting materials. The SEM images of reacted basalt from the 250 °C experiments (Fig. 2c-f) show that the initially clean and relatively smooth grain surfaces were completely covered by micrometre-scale, honeycomb-like structures typical for neo-formed smectite (e.g., Fiore et al., 2001; Baldermann et al., 2012; Baldermann et al., 2017). While the alteration features observed for the 290 °C experiments are very similar, the clayey textures seem to be more regular for both crystalline basalt experiments (Fig. 2c, d). In some cases, the formation of a protective smectite layer covering the underlying basalt grains (Fig. 2e) was observed. It seems plausible that the separation of the protective layer from the underlying basalt grain results from quenching and the sample preparation. Apart from the ubiquitous clay minerals, abundant anhydrite formed in all experiments (Fig. 2f). The size of the precipitated anhydrite crystals ranged from approximately 1 µm to 200 µm in length. A separate experiment with basaltic glass at 250 °C that was terminated after only one day showed that both of these alteration minerals were already present after this period, although the honeycomb-like smectite structures were less developed.

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The formation of smectite (probably ferrous saponite) and anhydrite was confirmed by XRD analyses of the reacted basalt glass and crystalline basalt samples at both temperatures (Supplementary Fig. 1), whereas no indication for the formation of secondary carbonates such as aragonite and strontianite was found. Minor amounts of zeolite also formed during the two basaltic glass experiments, in addition to plagioclase (close to albitic composition) and serpentine minerals at 290 °C.
378 **4 Discussion**

379 **4.1 Chemical exchange**

380 The changes in the chemical composition of the seawater and the alteration mineral
381 assemblage observed during the experiments in this study are comparable to those seen in
382 natural marine hydrothermal vent fluids (e.g., German and Seyfried, 2014). Consistent with
383 previous experimental studies (Seyfried and Bischoff, 1981; Seyfried and Mottl, 1982), Mg
384 was found to be transferred from seawater into secondary, trioctahedral clay minerals such as
385 saponite and serpentine that formed rapidly during the experiments, leading to the almost
386 complete removal of Mg from the liquid (Figs. 3 and 4). Furthermore, the kinetics of Mg
387 removal and the Ca trends at 250 °C fit are consistent with experiments at 150 °C and 300 °C
388 reported by Seyfried and Bischoff (1979; 1981), as illustrated for the crystalline basalt
389 experiments in Fig. 6. The rapid formation of saponite is consistent with the saturation indices
390 for Mg-saponite calculated with PHREEQC, which indicate that the reactive liquid is
391 extremely oversaturated with respect to this phase (Supplementary Fig. Error! Reference
392 source not found.2). The formation of anhydrite consumed dissolved SO4²⁻, and to a lesser
393 extent Ca, including some of the Ca released from the basalt dissolution (see section 4.2).
394 Sodium, Cl and K are not significantly incorporated into alteration phases, with the latter
395 being progressively enriched in the fluid. In contrast, Si and Fe released by basalt dissolution
396 were likely scavenged by authigenic (ferrous) clay minerals (Baldermann et al., 2014;
397 Baldermann et al., 2015; Baldermann et al., 2017) and/or in the case of Si precipitated on
398 minor amounts of pre-existing quartz which was potentially introduced while grinding the
399 basalt (indicated by XRD spectra; Supplementary Fig. 1) at least during the initial stage of
400 seawater-basalt interaction (Supplementary Fig. Error! Reference source not found.2). The
401 absence of Fe³⁺-bearing clay minerals such as nontronite and Fe-O-OH phases in the reaction
402 products (Supplementary Fig. 1) indicates that reducing conditions had been established
during the experiments, albeit traces of ferric phases (i.e. ferrihydrite or goethite), which could have been formed during quenching of the experiments. Our thermodynamic calculations also suggest that the reactive fluids were slightly undersaturated with respect to low-albite, highlighting the uncertainties inherent in current thermodynamic data at the experimental conditions as the formation of albite-rich plagioclase during the basaltic glass experiment at 290 °C is clearly indicated by XRD (cf. section 3.2).

The drop in pH observed at the beginning of the crystalline basalt experiments was predominantly caused by the incorporation of Si, Fe and Mg complexes and $OH^-$ into the neo-formed clay minerals (Bischoff and Dickson, 1975; Seyfried and Bischoff, 1977; Seyfried and Bischoff, 1979; Seyfried and Bischoff, 1981; Seyfried and Mottl, 1982). These initially low pH are less prominent in the basaltic glass experiments, where alkaline conditions prevailed throughout the experiments (Figs. 3 and 4), because the dissolution rate of this material is faster compared to its crystalline analogue. While pH in natural hydrothermal fluids is commonly slightly acidic, which can be explained by buffering with a plagioclase-chlorite-tremolite-quartz assemblage (Seyfried et al., 2010; Pester et al., 2012), pH increased to slightly alkaline values at the end of all experiments in this study; these final pH values exceeds those seen in previous studies (Menzies and Seyfried, 1979; Seyfried and Bischoff, 1981). These high pH values are potentially related to uptake of HCl by the vapour, which is increased at high salinities (Simonson and Palmer, 1993; Palmer et al., 2004), or to the presence of low amounts of ammonium, which potentially affects pH due to its influence on aqueous speciation.

## 4.2 Strontium isotopic exchange

### 4.2.1 Initial stage (<24 hours)

The most significant Sr elemental and isotopic exchange occurred during the first 24 h. Aqueous Sr concentrations decreased by more than 70% over this time, as the $^{87}$Sr/$^{86}$Sr ratio of the fluid decreased from the seawater towards the basalt value (Fig. 5). Because changes in
429

$^{87}$Sr/$^{86}$Sr reflect the mixing between different phases and not isotopic fractionation, both the release of Sr from the basalt and the removal of Sr into alteration phases are necessary to explain these observations.

430

431

432

The presence of anhydrite in the solids recovered after the experiments coupled with the decrease of $SO_4^{2-}$ during the first 24 h (Fig. 3) suggests that anhydrite was the main phase taking up Sr (Shikazono and Holland, 1983). This is consistent with the anhydrite saturation indices calculated with PHREEQC for these experiments (Supplementary Fig. Reference source not found.2) and past observations of anhydrite precipitation when heating seawater to temperatures of 150 to 200 °C (Bischoff and Seyfried, 1978). It is worth noting, however, that the existence of minor $Mg^{2+}$, $Ca^{2+}$ or $Sr^{2+}$ in the interlayer sites of ferrous saponite is possible, because divalent cations preferentially substitute for $Na^+$ (Baldermann et al., 2014), although the predominance of interlayer $Na^+$ is likely due to the high Na concentrations (~500-600 mmol/kg, Table 4) of the reactive fluids.

433

434

435

436

437

438

439

440

441

442

Assuming that anhydrite was the only sink for aqueous Ca and $SO_4^{2-}$ during this period, mass balance constraints allow the calculation of the mass of basalt dissolved and of anhydrite precipitated from the compositions of the initial seawater and first fluid sample, the liquid mass, and the basalt composition. Note that this calculation also assumes that the basalt dissolution was stoichiometric, which was likely not the case as inferred from Sr isotopic trends as described below. Therefore, the calculated amounts of basalt dissolved are approximate, but could be regarded as a proxy for Ca release rates if Ca release rates can be assumed to be proportional to those of Sr. The results of these calculations suggest that 4.3±0.3 g and 4.7±0.4 g of basaltic glass were dissolved at 250 °C and 290 °C, respectively, within the first day. Consistent with its aforementioned lower reactivity, only 0.62±0.06 g and 1.7±0.2 g of crystalline basalt dissolved at 250 °C and 290 °C over this time frame. The corresponding masses of anhydrite precipitated are 1.029±0.002 g and 0.85±0.02 g for the basaltic glass experiments at 250 °C and 290 °C, respectively, and 0.513±0.001 g and

443

444

445

446

447

448

449

450

451

452

453

454
0.74±0.02 g for the crystalline basalt experiments. The relative timing of basalt dissolution and anhydrite precipitation, and the evolution of the Sr isotopic system during the initial 24 h of each experiment were determined using an iterative mass balance model. In this model, the mass of precipitated anhydrite ($m_{Anhydrite}$), and the fluid Sr concentration, $^{87}$Sr/$^{86}$Sr and $\delta^{88/86}$Sr were predicted as functions of the mass of basalt dissolved ($m_{Basalt}$) (see discussion below and Appendix A for details). The rate of anhydrite precipitation (relative to basalt dissolution) was assumed to be proportional to ($\Omega_{Anhydrite}$ − 1), where $\Omega_{Anhydrite}$ stands for the saturation state of the liquid with respect to anhydrite at each model step. The amount of Sr incorporated into anhydrite was calculated assuming a constant partitioning coefficient between Sr in anhydrite and liquid $D_{Sr} = (dm_{Anhydrite}^{Sr}/dm_{Anhydrite}^{Ca})/(m_{Liquid}^{Sr}/m_{Liquid}^{Ca})$ for each experiment, where $m_{p}^{i}$ denotes the mass of element i in phase p. The best-fit values for $D_{Sr}$ range from 0.56 to 0.84, which is similar to the range of experimentally determined values of 0.27 to 0.73 (Shikazono and Holland, 1983) and partition coefficients of 0.60 (2 s.d. = 0.28) derived from the TAG hydrothermal mound (Teagle et al., 1998a). It should be noted, however, that partitioning of Sr between fluid and anhydrite is potentially affected by kinetic effects due to the fast precipitation in our experiments: it is well established that non-equilibrium partitioning can lead to significant variability in partitioning coefficients (e.g., DePaolo, 2011), suggesting that they might be different in natural hydrothermal systems, where anhydrite precipitation is likely slower in most cases. The stable Sr isotopic evolution was constrained by two fractionation factors, defined as

$$ \alpha_{\text{Anhydrite-Liquid}}^{88/86} = \frac{d^{88}\text{Sr}_{\text{Anhydrite}}/d^{86}\text{Sr}_{\text{Anhydrite}}}{^{88}\text{Sr}_{\text{Liquid}}/^{86}\text{Sr}_{\text{Liquid}}} \quad (4.1) $$

and

$$ \alpha_{\text{CrytBasalt-Liquid}}^{88/86} = \frac{^{88}\text{Sr}_{\text{CrystBasalt}}/^{86}\text{Sr}_{\text{CrystBasalt}}}{d^{88}\text{Sr}_{\text{Liquid}}/d^{86}\text{Sr}_{\text{Liquid}}} \quad (4.2) $$
where $ ^{4}Sr_{p} $ stands for the amount of the Sr isotope with mass number $ A $ in phase $ p $. The use of the differentials (d) signifies that isotopic fractionation was assumed to be irreversible. The latter factor, describing Sr isotopic fractionation between dissolving crystalline basalt (d$Sr_{Liquid}$) and the bulk crystalline basalt ($Sr_{CrystBasalt}$) was introduced, to account for the systematically larger $ \delta^{88/86}Sr $ variations in the experiments with crystalline basalt compared to basaltic glass (Fig. 5). This apparent fractionation is likely the result of the preferential dissolution of minerals having a stable Sr composition distinct from that of the bulk basalt. The effect of non-stoichiometric dissolution has also been suggested to explain stable Sr isotopic signatures in continental weathering environments (Halicz et al., 2008; Wei et al., 2013). The model was applied to all four experiments simultaneously to obtain the best-fit values for unknown parameters ($D_{Sr}$, $\alpha_{i}$, ...), but distinct values of these parameters for different experiments or temperatures were allowed as described above and in Appendix A.

Model predictions are shown in Fig. 7 and Fig. 8, and reveal a good fit to the experimental data. Accordingly, fast anhydrite precipitation (compared to basalt dissolution) from seawater is predicted at the experimental temperatures (Bischoff and Seyfried, 1978), leading to a drop in aqueous Sr concentrations, while $ ^{87}Sr/^{86}Sr $ remains relatively constant at this early stage. Subsequently, basalt dissolution becomes the dominant process and fluid $ ^{87}Sr/^{86}Sr $ decreases towards the basalt signature, which is promoted by the lower Sr content of the fluid. Afterwards, aqueous Sr concentrations remain stable or increase slightly due to ongoing basalt dissolution (Fig. 8). At this stage, anhydrite precipitation is limited by the low $ SO_{4}^{2-} $ concentrations in the fluids while $ ^{87}Sr/^{86}Sr $ remain low (Fig. 8). This result is similar to the proposed reaction path of Berndt et al. (1988), and consistent with the trends inferred from anhydrite found in altered basalts from the Costa Rica Rift, eastern Pacific (ODP Hole 504B, Teagle et al., 1998b), while evidence for the last stage of fluid evolution was missing in these natural anhydrite samples.
The model results further suggest that changes in aqueous δ<sup>88/86</sup>Sr are caused by isotopic fractionation between the reactive fluid and the precipitating anhydrite, as well as by the non-stoichiometric dissolution of crystalline basalt. The best-fit anhydrite-liquid fractionation factor is ε<sub>Anhydrite-Liquid</sub><sup>88/86</sup> = (α<sub>Anhydrite-Liquid</sub><sup>88/86</sup> - 1) = 0.034 ± 0.019 ‰. This value is slightly positive suggesting that anhydrite preferentially incorporates heavier Sr isotopes. The ε<sub>CrystBasalt-Liquid</sub><sup>88/86</sup> values are 0.27±0.17 ‰ and 0.34±0.03 ‰ at 250 °C and 290 °C, respectively (Fig. 9) suggesting that the light Sr isotopes are preferentially released during non-stoichiometric crystalline basalt dissolution. This can also be seen from the changes in fluid Sr concentration and δ<sup>88/86</sup>Sr in the experiment with crystalline basalt at 290 °C between samples 01 and 02 (Fig. 5): Strontium concentrations increase from very low levels to 37 µmol/kg, indicating that almost all Sr in fluid sample 02 originates from the dissolving basalt, because anhydrite re-dissolution is unlikely at this point as <sup>87</sup>Sr/<sup>86</sup>Sr remained low. Therefore, the difference in δ<sup>88/86</sup>Sr of sample 02 and the bulk crystalline basalt, approximately 0.18 ‰, is likely attributable to the fractionation of Sr during crystalline basalt dissolution; this fractionation is of similar magnitude as the fractionation factors obtained by the mass balance model. Fig. 9 shows that the uncertainties of modelled ε<sub>Anhydrite-Liquid</sub><sup>88/86</sup> and ε<sub>CrystBasalt-Liquid</sub><sup>88/86</sup> are inversely correlated, so that if a lower degree of fractionation during non-stoichiometric crystalline basalt dissolution in our experiments could be inferred from independent constraints, the best fit ε<sub>Anhydrite-Liquid</sub><sup>88/86</sup> would increase.

It should be noted that the primary incentive of this mass balance model is to explain the variation of Sr concentrations and isotopic composition in the fluid samples of the experiments in this study and to extract isotopic fractionation factors from these data. While the general trends seen here are likely to occur in natural marine hydrothermal systems as well, a quantitative, predictive model of Sr isotope variations in such systems is beyond the
scope of this model (e.g. due to differences in rates of anhydrite precipitation as explained above).

### 4.2.2 Sr evolution at elapsed times >24 hours

Following the initial period of coupled anhydrite precipitation and basalt dissolution, the Sr concentrations in the fluid phase increased. Aqueous $^{87}$Sr/$^{86}$Sr also increased slightly during the later stages of all experiments (Fig. 5 and solid lines in Fig. 8), this suggest this evolution is not driven by basalt dissolution, as this would be expected to drive aqueous $^{87}$Sr/$^{86}$Sr towards the lower basalt signature. Since this phenomenon also occurs in the basaltic glass experiments (which is assumed to have a homogeneous $^{87}$Sr/$^{86}$Sr composition), the dissolution of a primary mineral with elevated $^{87}$Sr/$^{86}$Sr compared to the bulk basalt is unlikely to be the cause of this process. It seems likely, therefore, that the increase in $^{87}$Sr/$^{86}$Sr is driven by the dissolution of previously precipitated alteration minerals that incorporated the higher $^{87}$Sr/$^{86}$Sr ratios of the initial seawater. Note that the degree of supersaturation with respect to anhydrite calculated with PHREEQC decreased during all experiments, and undersaturation was predicted for some of the later fluid samples in the crystalline basalt experiments (Supplementary Fig. Error! Reference source not found.2). Assuming a mean $^{87}$Sr/$^{86}$Sr ratio of 0.708938 for the precipitated anhydrite (as estimated by the mass balance model described earlier), mass balance calculations suggest that up to 15 % of the anhydrite may have been re-dissolved. Since $SO_4^{2-}$ did not increase significantly during the later stages of the experiments (Fig. 3), a possible mechanism for anhydrite re-dissolution is the reduction of aqueous $SO_4^{2-}$ by the oxidation of $Fe^{2+}$ released by basalt dissolution, leading to an undersaturation with respect to anhydrite and its subsequent re-dissolution. Although the reduced aqueous S concentrations are not available for our experiments (and PHREEQC saturation states in this study were calculated assuming that sulfur is present as $S^{6+}$), calculations by Mottl et al. (1979) under similar conditions suggest that anhydrite became metastable during their experimental runs due to the reducing nature of the final fluids.
Furthermore, it is possible that the fast precipitation of anhydrite at the beginning of the experiments lead to the incorporation of higher amounts of Sr than at equilibrium due to kinetic effects, so that Sr was released from anhydrite during the later stages due to its re-crystallization (Syverson et al., 2016; Scheuermann et al., 2018), contributing to the release of radiogenic Sr.

Reactive fluid $\delta^{88/86}\text{Sr}$ remained relatively constant after the early anhydrite precipitation (Fig. 5) with the exception of the crystalline basalt experiment at 290 °C, where the extremely low aqueous $\delta^{88/86}\text{Sr}$ gradually increase towards the basaltic $\delta^{88/86}\text{Sr}$ value, likely due to continued basalt dissolution and/or possible anhydrite re-dissolution.

## 4.3 Implications for the marine Sr budget

The contribution of hydrothermal systems to the marine $^{87}\text{Sr}/^{86}\text{Sr}$ budget can be estimated by comparing the heat budget of mid-ocean ridges with the cooling of new oceanic lithosphere. Estimates of the latent heat of crystallisation and the required amount of cooling of newly created oceanic crust places upper limits on the amount of heat transferred to the oceans through hydrothermal circulation at mid ocean-ridges. Assuming a certain distribution of fluid temperatures, the corresponding hydrothermal water flux can be calculated (Morton and Sleep, 1985; Palmer and Edmond, 1989; Elderfield and Schultz, 1996; Mottl, 2003; Davis et al., 2003). This, in turn, allows estimation of the contribution of these hydrothermal systems to the marine $^{87}\text{Sr}/^{86}\text{Sr}$ budget, assuming the change of $^{87}\text{Sr}/^{86}\text{Sr}$ and Sr concentration in the seawater during hydrothermal alteration is known. Based on such calculations, Davis et al. (2003) estimated that high-temperature fluids (400 °C), transferring 50 % of the available heat and cooling the lower crust to 860 °C, can only supply $1/6^{th}$ of the unradiogenic Sr required to balance the other marine Sr fluxes (mainly the high-$^{87}\text{Sr}/^{86}\text{Sr}$ river input). Our results show that fluids with low $^{87}\text{Sr}/^{86}\text{Sr}$ ratios and Sr concentrations comparable to that of seawater can be generated by seawater-basalt interactions at 250 °C within 2 months, and even faster at 290 °C. This is consistent with the experimental data of James et al. (2003),
who observed significant Sr exchange with basalt even at 150 °C. The time required for Sr exchange in our experiments is well within the residence time of ≤ 3 years estimated for fluids with temperatures of above 200 °C in the Juan de Fuca Ridge hydrothermal system by Kadko and Butterfield (1998) using Ra isotopes. It is also within the range of residence times on the order of years estimated independently by Fisher (2003). Assuming, as an example and upper limit, that the 50 % of the available heat not removed by high-temperature (400 °C) fluids in the Davis et al. (2003) estimate is lost through intermediate-temperature (250 °C) fluids, the calculated basaltic Sr flux to the oceans increases by a factor of ~3.2, reducing the missing basaltic Sr flux in the marine $^{87}$Sr/$^{86}$Sr budget by half. This scenario assumes that such intermediate-temperature fluids interact with basalts of relatively unradiogenic (i.e. unaltered) Sr composition, which is likely as alteration at lower temperatures in the ridge flanks is concentrated in the volcanic section, where higher-temperature circulation focuses on the underlying sheeted dikes and gabbros (e.g., Staudigel, 2014). The large uncertainties in both the amount of heat transferred from the crust to the oceans and the partitioning of hydrothermal fluid temperatures (German and Seyfried, 2014) limit a more detailed estimation of the basaltic Sr flux from hydrothermal systems. However, their potential to help reconcile the marine $^{87}$Sr/$^{86}$Sr budget means that intermediate-temperature hydrothermal fluids could play a significant role in the global heat flux at MORs, supporting a similar notion of Nielsen et al. (2006), and complementing other proposed unradiogenic $^{87}$Sr/$^{86}$Sr sources such as island arc weathering (Allègre et al., 2010), low-temperature alteration and carbonate precipitation (Coogan and Dosso, 2015) and dissolution of river derived basaltic particles (Jones et al., 2012a).

Apart from heat flux constraints discussed above, Davis et al. (2003) noted that the extent of ocean crust alteration is inconsistent with the current marine $^{87}$Sr/$^{86}$Sr budget. This argument is based on the assumption that basalt alteration occurs solely by Sr isotopic exchange with the fluids, replacing the low MORB $^{87}$Sr/$^{86}$Sr by high seawater $^{87}$Sr/$^{86}$Sr while
leaving the Sr concentrations unchanged. Our experimental results and Sr isotopic modelling indicate that isotopic exchange occurs mainly during the early anhydrite precipitation and subsequent release of unradiogenic ⁸⁷Sr/⁸⁶Sr by basalt dissolution. If anhydrite is re-dissolved at a later stage, the ⁸⁷Sr/⁸⁶Sr composition of the altered basalt may be similar to the fresh basalt, such that the ⁸⁷Sr/⁸⁶Sr ratios of the recovered altered oceanic rocks (e.g., Staudigel et al., 1995; Alt and Teagle, 2000; Harris et al., 2015) may not accurately reflect the degree of isotopic Sr exchange. In the present experiments, re-dissolution of anhydrite was most likely caused by the establishment of reducing conditions due to the release of Fe²⁺ (cf. section 4.2.2). Furthermore, increased SO₄²⁻ in the quenched liquid compared to the final sample taken at experimental temperature (Table 4) indicates that anhydrite rapidly re-dissolves when it becomes undersaturated at lower temperatures (as the solubility of anhydrite is retrograde). In natural systems, anhydrite re-dissolution might occur due to interactions with circulating low-temperature fluids at a later time (Mottl et al., 1979). Such processes limit the use of altered basalt compositions to infer the hydrothermal Sr flux to the oceans and provide a potential explanation for the Sr imbalances arising from models such as the marine Sr budget calculations of Davis et al. (2003) and the global oceanic high-temperature hydrothermal chemical exchange model of Coogan and Dosso (2012). It should also be noted that anhydrite has been found where expected in undisturbed oceanic crustal sections that have been sampled, but that there is a lack of samples from undisturbed basement sites (Alt, 2003).

Similar to the radiogenic ⁸⁷Sr/⁸⁶Sr imbalance in the marine budget, several studies reported a disequilibrium between marine Sr input and output fluxes by analysing stable δ⁸⁸/⁸⁶Sr budgets (Krabbenhöft et al., 2010; Pearce et al., 2015). In these studies, the estimated δ⁸⁸/⁸⁶Sr of the combined marine Sr inputs, dominated by rivers and hydrothermal systems (~0.3 ‰), is significantly heavier than the carbonate sink (0.15 to 0.21 ‰), suggesting that seawater δ⁸⁸/⁸⁶Sr is currently not at steady state. However, only a few δ⁸⁸/⁸⁶Sr data of hydrothermal fluids are available at present, with Pearce et al. (2015) deriving a value of
δ<sup>88/86</sup>Sr=0.24 ‰ by extrapolation of eleven hydrothermal fluid measurements to a MORB
629
hydrothermal endmember composition. Our experiments suggest that significantly lighter
630
fluid δ<sup>88/86</sup>Sr compositions may be reached in reactions with crystalline basalt. Although there
631
is no requirement for modern seawater δ<sup>88/86</sup>Sr to be at steady state (and indeed, glacial-
632
interglacial weathering effects would imply that this is unlikely to be the case; Vance et al.,
633
2009), a lower integrated hydrothermal δ<sup>88/86</sup>Sr composition, or a larger net flux of this
634
composition, would shift the combined marine input δ<sup>88/86</sup>Sr towards lighter values, reducing
635
the suggested imbalance between current input and output fluxes. Furthermore, as Sr is
636
incorporated into anhydrite during precipitation from temperate seawater, hydrothermal
637
systems may also represent an important but currently overlooked Sr sink in the marine Sr
638
cycle. The relatively small δ<sup>88/86</sup>Sr isotopic fractionation between anhydrite and liquid
639
observed during the early stage of hydrothermal alteration in our experiments produces a
640
δ<sup>88/86</sup>Sr composition similar to that of seawater. Therefore, marine hydrothermal systems may
641
shift the flux-weighted δ<sup>88/86</sup>Sr output composition towards heavier values, further reducing
642
the potential imbalance between global Sr sources and sinks. However, the global significance
643
of hydrothermal sinks relative to the large carbonate burial Sr flux (Krabbenhöft et al., 2010;
644
Pearce et al., 2015) is not well constrained at present.
645

Importantly, the possibility that hydrothermal fluid δ<sup>88/86</sup>Sr compositions may be
646
isotopically lighter than previously inferred provides new insight into the processes that
647
govern marine Sr exchange, including its fractionation during incorporation into anhydrite
648
and non-stoichiometric basalt dissolution. The resulting lower hydrothermal δ<sup>88/86</sup>Sr input flux
649
would also affect models of past seawater δ<sup>88/86</sup>Sr, and therefore put further constraints on
650
past oceanic carbonate fluxes as aragonite and calcite can take up significant amounts of Sr
651
(Vollstaedt et al., 2014). Furthermore, the magnitude of isotopic fractionation at equilibrium
652
increases with decreasing temperature (Urey, 1947; Bigeleisen and Mayer, 1947), so that Sr
653
isotopic fractionation between anhydrite and liquid may likely be larger in low-temperature
654
systems, unless equilibrium fractionation is cancelled out by opposing kinetic isotope effects. 655

On the other hand, isotopic fraction is likely smaller at higher temperatures (> 400 °C), 656

suggesting that δ<sup>88/86</sup>Sr potentially provides a tool to distinguish high- vs. low-temperature Sr 657

exchange in altered samples from the oceanic crust: while mass balances yield Sr isotopic 658

compositions of the bulk solids at the end of the experiments ranging in <sup>87</sup>Sr/<sup>86</sup>Sr from 659

0.705706 to 0.706541, which is similar to values of natural samples at varying temperatures, 660

mass balances indicate δ<sup>88/86</sup>Sr of 0.28 ‰ and 0.29 ‰ for the final solids of experiments with 661

basaltic glass at 250 and 290 °C, respectively, and 0.35 ‰ and 0.37 ‰ for the corresponding 662

experiments with crystalline basalt. This indicates that basalt altered at moderate temperatures 663

may be significantly heavier compared to fresh basalt (δ<sup>88/86</sup>Sr of 0.23-0.26 ‰), which is 664

unlikely to occur at higher temperatures due to smaller magnitudes of fractionation. 665

666

667 **5 Conclusions**

Basaltic glass and crystalline basalt were reacted with seawater at 250 °C and 290 °C 668

and at saturated vapour pressure. Chemical exchange was monitored by liquid phase 669

sampling, and the results agree well with those obtained from direct measurements of mid- 670

ocean ridge vent fluid compositions. Minor differences, such as an increase of pH after near 671

complete removal of aqueous Mg by authigenic ferrous saponite and serpentine, may be 672

related to the absence of epidote from the alteration mineral assemblage. 673

The Sr concentration, radiogenic <sup>87</sup>Sr/<sup>86</sup>Sr ratio and stable δ<sup>88/86</sup>Sr composition of the 674

fluid phase changed considerably during the experiments. An initial drop in <sup>87</sup>Sr/<sup>86</sup>Sr towards 675

basaltic values coupled with a decrease in Sr concentrations can be modelled by assuming Sr 676

removal via anhydrite precipitation, followed by an increased influence of Sr release from 677

basalt dissolution over time as the liquid becomes depleted in SO₄²⁻. A decrease in δ<sup>88/86</sup>Sr 678

observed during the early stage of basalt-seawater interaction suggests that isotopically light 679
680 Sr may be preferentially released during non-stoichiometric crystalline basalt dissolution. In 681 contrast, anhydrite shows a slight preference for isotopically heavy Sr with $\epsilon_{Anhydrite-Liquid}^{88/86}$ = 682 0.034 ± 0.019 ‰. The later evolution of the Sr isotopic composition of the fluids suggests 683 that up to 15 % of the precipitated anhydrite was re-dissolved, potentially resulting from 684 reducing conditions and quenching of the fluids at the end of the experiments.

685 Given the rapid achievement of near basaltic $^{87}$Sr/$^{86}$Sr signatures and significant 686 changes in the Sr concentrations even at 250 °C, intermediate-temperature oceanic 687 hydrothermal fluids could be the source of some of the ‘missing basaltic Sr input’ that has 688 been postulated in marine radiogenic Sr budgets (e.g., Davis et al., 2003). Furthermore, the 689 inconsistency of observed MORB alteration patterns with the marine Sr budget and low 690 amounts of anhydrite found in ocean crust profiles (Teagle et al., 1998b) may be related to the 691 re-dissolution of anhydrite induced during later circulation of reducing or low-temperature 692 fluids. However, the persisting uncertainties in the temperature distribution of the global 693 hydrothermal water flux complicate the quantitative balancing of the marine Sr budget, which 694 should also include other proposed low-$^{87}$Sr/$^{86}$Sr sources such as island arc weathering 695 (Allègre et al., 2010), dissolution of riverine particulate materials (Jones et al., 2012a) and 696 low-temperature alteration and carbonate precipitation (Coogan and Dosso, 2015).

697 Finally, our experimental data suggests that marine hydrothermal fluids with an 698 isotopically light $\delta^{88/86}$Sr signature can be generated by isotopic fractionation during both 699 anhydrite precipitation and non-stoichiometric crystalline basalt dissolution. Such insights 700 into the processes that potentially affect seawater $\delta^{88/86}$Sr highlight the need for detailed data 701 on the $\delta^{88/86}$Sr of the marine hydrothermal outputs, which will help to further constrain models 702 of the current and past marine Sr cycle (Krabbenhöft et al., 2010; Vollstaedt et al., 2014; 703 Pearce et al., 2015) and thus its relationship to global changes in plate tectonics, weathering, 704 and climate.
705 Acknowledgements

706 This research was supported by the CO2-REACT Marie Curie Actions Initial Training

707 Network which was funded by the European Union. The authors would like to thank Frederic

708 Caudaudap, Aurélie Marquet, Alain Castillo, Pascal Gisquet, Thierry Aigouy, Stephanie

709 Mounic, Mathieu Benoit, Carole Causserand, Manuel Henry, Jonathan Prunier, Matt Cooper,

710 and Agnes Michalik for help with analyses and technical assistance. We would also like to

711 thank Deirdre Clark, Jan Přikryl, and Sigurður Reynir Gíslason for providing the Stapafell

712 basalt, as well as Catherine Jeandel for providing the seawater. Reviews by Nicholas Pester,

713 Sune Nielsen, and an anonymous reviewer were greatly appreciated by the authors, and we

714 thank Jeffrey Alt for editorial handling of the manuscript.

715 References

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717 fundamental role of island arc weathering in the oceanic Sr isotope budget. *Earth*

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722 U.S. Environmental Protection Agency, Athens, Georgia.

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984

985 **Appendix A Model details**

986 The rate of anhydrite precipitation is constrained using the relationship

$$
\frac{dm_{\text{Anhydrite}}}{dm_{\text{Basalt}}} = k(\Omega_{\text{Anhydrite}} - 1) \qquad (\text{A.1})
$$

987 where ΩAnhydrite stands for the saturation state of the liquid with respect to anhydrite, and k is
988 a fitted constant. This corresponds to a simple form of the empirical rate law (Morse, 1978)
989 assuming that dm<sub>Basalt</sub>/dt does not vary significantly with liquid composition and that basalt
990 dissolution and anhydrite precipitation have the same relative dependency on temperature
991 during the short initial heating period. The anhydrite saturation state is calculated from the
992 aqueous solution composition using PHREEQC; for this calculation, the aqueous solution
993 composition other than Ca, Sr, and SO<sub>4</sub><sup>2-</sup> is estimated by a linear mixture between the known
994 initial seawater composition and first sample compositions as function of m<sub>Basalt</sub>.

995 The amount of Sr incorporated into anhydrite is calculated using the partition
996 coefficient

$$
D_{\text{Sr}} = \frac{\frac{dm_{\text{Anhydrite}}^{\text{Sr}}}{dm_{\text{Anhydrite}}^{\text{Ca}}}}{\frac{m_{\text{Liquid}}^{\text{Sr}}}{m_{\text{Liquid}}^{\text{Ca}}}} \quad (\text{A.2})
$$

997 with $m_p^i$ denoting the mass of element $i$ in phase $p$. Its value is determined by finding the best
998 fit to the boundary values (as for $k$), and both of these parameters are allowed to differ for the
999 individual experiments to account for the different reaction kinetics and temperatures.
The evolution of the Sr isotopic composition is modelled using the two fractionation factors between anhydrite and liquid and during non-stoichiometric basalt dissolution. The values for both fractionation factors are estimated by fitting the model to the boundary values, assuming distinct $\alpha_{CrystBasalt-Liquid}^{88/86}$ for the two temperatures used. Aqueous $^{87}$Sr/$^{86}$Sr is calculated from the stable Sr isotope ratios by re-normalisation to a $^{86}$Sr/$^{88}$Sr value of 0.1194 using the exponential fractionation law.

In addition to the isotopic fractionation factors, the rate of anhydrite precipitation and the bulk Sr partition coefficient $D_{Sr}$, the following relationships define the evolution of the Sr isotope ratios in the liquid. For each isotope of $^{A}$Sr with mass number A,

$$ \frac{\mathrm{d}m_{\mathrm{Liquid}}^{A\mathrm{Sr}}}{\mathrm{d}m_{\mathrm{Basalt}}} = c_{\mathrm{Anhydrite}}^{A\mathrm{Sr}} \frac{\mathrm{d}m_{\mathrm{Anhydrite}}}{\mathrm{d}m_{\mathrm{Basalt}}} + c_{\mathrm{Basalt}}^{A\mathrm{Sr}} \quad (\mathrm{A.3}) $$

where $m_p^i$ and $c_p^i$ denote the mass and concentration of i in phase p. Furthermore, basic mass balance implies $\sum_A m_p^{A\mathrm{Sr}} = m_p^{\mathrm{Sr}}$ and $\sum_A c_p^{A\mathrm{Sr}} = c_p^{\mathrm{Sr}}$. Due to its negligible abundances, $^{84}$Sr was ignored in the model calculations. Based on a given $\alpha^{88/86}$, the $\alpha^{87/86}$ in the model is calculated according to the kinetic mass fractionation law (Young et al., 2002)

$$ \alpha^{A/86} = (\alpha^{B/86})^{\beta} \quad (\mathrm{A.4}) $$

where $\beta$ is defined as

$$ \beta = \frac{\ln \frac{M_{A\mathrm{Sr}}}{M^{86\mathrm{Sr}}}}{\ln \frac{M_{B\mathrm{Sr}}}{M^{86\mathrm{Sr}}}} \quad (\mathrm{A.5}) $$

where $M_{X_{Sr}}$ denotes the atomic mass of isotope $^{X}$Sr. The kinetic mass fractionation equation was chosen as opposed to the equilibrium mass fractionation equation due to the relatively fast Sr precipitation rates. However, due to the negligible difference of the $\beta$ values in both equations, this does not significantly impact the results.
The solutions to the differential equation system are computed numerically using the NDSolve function of Wolfram Mathematica in combination with the IPhreeqc modules (Charlton and Parkhurst, 2011), essentially taking a sequence of steps in the independent variable $m_{\text{Basalt}}$ starting at the initial values, and using an adaptive procedure to determine the size of the steps depending on the variation of the solution in a particular region. The values for the model parameters ($k$, $D_{\text{Sr}}$ and the different $\alpha^{88/86}$) are found by numerically minimising the sum of squared errors between the model predictions and the first sample composition, weighted by the reciprocal of the variance and using the Nelder-Mead method (Nelder and Mead, 1965) implementation in Wolfram Mathematica.

## Tables

<table><caption>**Table 1:** Experimental parameters of the basalt-seawater experiments.</caption><thead><tr><th rowspan="2">Experiment</th><th rowspan="2">T / °C</th><th rowspan="2">Basalt type</th><th colspan="2">Starting materials / g</th><th rowspan="2">Duration / days</th></tr><tr><th>Seawater</th><th>Basalt</th></tr></thead><tbody><tr><td>BSW04</td><td>250</td><td>Glassy</td><td>275.0</td><td>27.50</td><td>53</td></tr><tr><td>BSW05</td><td>250</td><td>Crystalline</td><td>275.0</td><td>27.50</td><td>53</td></tr><tr><td>BSW06</td><td>290</td><td>Glassy</td><td>220.0</td><td>22.01</td><td>92</td></tr><tr><td>BSW07</td><td>290</td><td>Crystalline</td><td>220.0</td><td>22.00</td><td>93</td></tr></tbody></table>

<table><caption>**Table 2:** Composition (wt. %) of the basaltic glass and crystalline basalt used as starting materials.</caption><thead><tr><th>Sample</th><th>SiO<sub>2</sub></th><th>Al<sub>2</sub>O<sub>3</sub></th><th>Fe<sub>2</sub>O<sub>3</sub><sup>c</sup></th><th>MgO</th><th>CaO</th><th>Na<sub>2</sub>O</th><th>K<sub>2</sub>O</th><th>TiO<sub>2</sub></th><th>MnO</th><th>P<sub>2</sub>O<sub>5</sub></th><th>Total</th></tr></thead><tbody><tr><td>Glass<sup>a</sup></td><td>48.1</td><td>14.6</td><td>12.0</td><td>9.1</td><td>11.8</td><td>2.0</td><td>0.3</td><td>1.6</td><td>0.2</td><td>0.2</td><td>99.9</td></tr><tr><td>Crystalline<sup>b</sup></td><td>47.9</td><td>13.4</td><td>12.3</td><td>10.0</td><td>12.2</td><td>1.5</td><td>0.3</td><td>1.6</td><td>0.2</td><td>0.2</td><td>99.1</td></tr></tbody></table>

<table><caption>**Table 3:** Sr content and isotopic composition of the basalt powders used as reactants in the experiments. The uncertainties are the 95 % confidence intervals (≈ 1.96 standard errors).</caption><thead><tr><th></th><th>Basaltic glass</th><th>Crystalline basalt</th></tr></thead><tbody><tr><td>Sr / ppm</td><td>138 ± 20</td><td>126 ± 20</td></tr><tr><td><sup>87</sup>Sr/<sup>86</sup>Sr</td><td>0.7031929 ± 0.0000028</td><td>0.7031887 ± 0.0000026</td></tr><tr><td><sup>88</sup>Sr/<sup>86</sup>Sr</td><td>0.2304 ± 0.0075</td><td>0.2592 ± 0.0075</td></tr></tbody></table>

<footnote>a From Oelkers and Gislason (2001)</footnote>

<footnote>b From Gudbrandsson et al. (2011)</footnote>

<footnote>c Most iron is $Fe^{2+}$</footnote>
1038

1039

1040

1041

<table><caption>**Table 4:** Chemical compositions of the seawater used as reactant and of the liquid samples taken during the experiments. Rows labelled "Quenched" show compositions of the liquid in the reactor after terminating the experiment by quenching. The uncertainties (±) designate the 95 % confidence interval (~2 standard errors), while '< LOD' indicates values that were below the detection limits (listed below the table). Uncertainties for all data are provided in Electronic Supplementary Table 1.</caption><thead><tr><th rowspan="2">Experiment</th><th rowspan="2">Sample</th><th rowspan="2">Duration / days</th><th rowspan="2">pH</th><th rowspan="2">Fluid mass / g</th><th colspan="10">Concentration / mmol/kg</th><th colspan="2"><sup>87</sup>Sr/<sup>86</sup>Sr</th><th colspan="2">δ<sup>88</sup>Sr / ‰</th></tr><tr><th>Na</th><th>Mg</th><th>Si</th><th>K</th><th>Ca</th><th>Sr</th><th>Fe</th><th>Cl</th><th>SO<sub>4</sub><sup>2-</sup></th><th>Value</th><th>± / ppm</th><th>Value</th><th>±</th></tr></thead><tbody><tr><td></td><td>Seawater</td><td></td><td>8.14</td><td></td><td>474</td><td>55</td><td>&lt; LOD</td><td>11.8</td><td>10.3</td><td>0.0896</td><td>0.040</td><td>597.4</td><td>29.05</td><td>0.7091729</td><td>2.8</td><td>0.372</td><td>0.007</td></tr><tr><td rowspan="8">BSW04<br>250 °C<br>basaltic glass</td><td>BSW04-01</td><td>1.0</td><td>7.71</td><td>275</td><td>489</td><td>17.0</td><td>14.4</td><td>12.9</td><td>16.8</td><td>0.0210</td><td>0.27</td><td>599.9</td><td>0.6039</td><td>0.7034758</td><td>2.7</td><td>0.214</td><td>0.007</td></tr><tr><td>BSW04-02</td><td>3.8</td><td>8.72</td><td>257</td><td>491</td><td>0.56</td><td>13.8</td><td>13.6</td><td>35.8</td><td>0.0435</td><td>0.0069</td><td>606.5</td><td>0.3053</td><td>0.7034811</td><td>2.5</td><td>0.248</td><td>0.006</td></tr><tr><td>BSW04-03</td><td>6.8</td><td>9.09</td><td>235</td><td>493</td><td>0.106</td><td>13.4</td><td>13.7</td><td>39.7</td><td>0.0497</td><td>0.0008</td><td>619.0</td><td>0.2867</td><td>0.7035802</td><td>2.6</td><td>0.273</td><td>0.006</td></tr><tr><td>BSW04-04</td><td>13.9</td><td>9.48</td><td>212</td><td>500</td><td>&lt; LOD</td><td>10.5</td><td>13.8</td><td>43.6</td><td>0.0633</td><td>&lt; LOD</td><td>620.4</td><td>0.2951</td><td>0.7043252</td><td>2.8</td><td>0.280</td><td>0.007</td></tr><tr><td>BSW04-05</td><td>21.0</td><td>9.58</td><td>190</td><td>504</td><td>&lt; LOD</td><td>9.8</td><td>13.8</td><td>45.8</td><td>0.0737</td><td>&lt; LOD</td><td>624.5</td><td>0.2893</td><td>0.7048168</td><td>2.8</td><td>0.311</td><td>0.007</td></tr><tr><td>BSW04-06</td><td>33.9</td><td>9.65</td><td>169</td><td>575</td><td>&lt; LOD</td><td>9.2</td><td>15.9</td><td>42.8</td><td>0.0566</td><td>&lt; LOD</td><td>684.6</td><td>0.3454</td><td>0.7045957</td><td>2.7</td><td>0.289</td><td>0.008</td></tr><tr><td>BSW04-07</td><td>52.9</td><td>9.52</td><td>148</td><td>601</td><td>&lt; LOD</td><td>10.5</td><td>16.9</td><td>36.9</td><td>0.0593</td><td>&lt; LOD</td><td>704.1</td><td>0.3501</td><td>0.7042755</td><td>2.6</td><td>0.298</td><td>0.006</td></tr><tr><td>Quenched</td><td></td><td></td><td></td><td>513</td><td>&lt; LOD</td><td>4.25</td><td>9.6</td><td>46.9</td><td>0.0433</td><td>&lt; LOD</td><td>621.9</td><td>4.003</td><td></td><td></td><td></td><td></td></tr><tr><td rowspan="8">BSW05<br>250 °C<br>crystalline basalt</td><td>BSW05-01</td><td>1.0</td><td>3.74</td><td>275</td><td>484</td><td>46.7</td><td>18.2</td><td>14.7</td><td>1.28</td><td>0.0074</td><td>1.37</td><td>608.0</td><td>14.854</td><td>0.7074845</td><td>2.7</td><td>0.194</td><td>0.006</td></tr><tr><td>BSW05-02</td><td>3.8</td><td>4.01</td><td>254</td><td>479</td><td>42.9</td><td>20.4</td><td>15.2</td><td>1.43</td><td>0.0066</td><td>2.0</td><td>603.8</td><td>12.448</td><td>0.7075882</td><td>3.0</td><td>0.225</td><td>0.007</td></tr><tr><td>BSW05-03</td><td>6.8</td><td>4.83</td><td>232</td><td>490</td><td>40.4</td><td>20.6</td><td>15.8</td><td>1.63</td><td>0.0098</td><td>2.4</td><td>610.8</td><td>10.150</td><td>0.7083240</td><td>2.6</td><td>0.254</td><td>0.006</td></tr><tr><td>BSW05-04</td><td>14.0</td><td>5.89</td><td>212</td><td>480</td><td>20.9</td><td>17.1</td><td>16.1</td><td>10.3</td><td>0.0196</td><td>1.43</td><td>597.5</td><td>0.9019</td><td>0.7062790</td><td>2.7</td><td>0.158</td><td>0.007</td></tr><tr><td>BSW05-05</td><td>21.0</td><td>9.03</td><td>192</td><td>489</td><td>0.84</td><td>6.0</td><td>16.7</td><td>35.7</td><td>0.0678</td><td>0.014</td><td>608.3</td><td>0.2906</td><td>0.7042098</td><td>2.7</td><td>0.168</td><td>0.007</td></tr><tr><td>BSW05-06</td><td>33.9</td><td>9.01</td><td>172</td><td>494</td><td>0.31</td><td>10.4</td><td>16.7</td><td>37.7</td><td>0.0779</td><td>0.007</td><td>620.1</td><td>0.2631</td><td>0.7044848</td><td>2.7</td><td>0.165</td><td>0.007</td></tr><tr><td>BSW05-07</td><td>52.9</td><td>9.08</td><td>151</td><td>502</td><td>0.172</td><td>9.9</td><td>17.1</td><td>40.6</td><td>0.0926</td><td>0.004</td><td>628.9</td><td>0.2544</td><td>0.7048623</td><td>2.7</td><td>0.189</td><td>0.007</td></tr><tr><td>Quenched</td><td></td><td></td><td></td><td>482</td><td>&lt; LOD</td><td>3.97</td><td>16.1</td><td>44.5</td><td>0.0982</td><td>&lt; LOD</td><td>622.0</td><td>4.003</td><td></td><td></td><td></td><td></td></tr></tbody></table>

1042

Detection limits (LOD) in mmol/kg were approximately: Mg: 0.026; Si: 0.018; Fe (BSW04 and BSW05): 0.0001; Fe (BSW06 and BSW07): 0.02

1043
<table><caption>1044 Table 4: (continued)</caption><thead><tr><th rowspan="2">Experiment</th><th rowspan="2">Sample</th><th rowspan="2">Duration / days</th><th rowspan="2">pH</th><th rowspan="2">Fluid mass / g</th><th colspan="9">Concentration / mmol/kg</th><th colspan="2"><sup>87</sup>Sr/<sup>86</sup>Sr</th><th colspan="2">δ<sup>88</sup>Sr / ‰</th></tr><tr><th>Na</th><th>Mg</th><th>Si</th><th>K</th><th>Ca</th><th>Sr</th><th>Fe</th><th>Cl</th><th>SO<sub>4</sub><sup>2-</sup></th><th>Value</th><th>± / ppm</th><th>Value</th><th>±</th></tr></thead><tbody><tr><td rowspan="7">BSW06<br>290 °C<br>basaltic glass</td><td>BSW06-01</td><td>0.8</td><td>8.36</td><td>220</td><td>524</td><td>6.9</td><td>23.7</td><td>12.0</td><td>28.6</td><td>0.0342</td><td>&lt; LOD</td><td>636.5</td><td>0.181</td><td>0.7034329</td><td>2.5</td><td>0.234</td><td>0.006</td></tr><tr><td>BSW06-02</td><td>3.0</td><td>9.20</td><td>200</td><td>541</td><td>0.078</td><td>19.7</td><td>12.3</td><td>42.4</td><td>0.0496</td><td>&lt; LOD</td><td>643.5</td><td>0.124</td><td>0.7034302</td><td>2.9</td><td>0.261</td><td>0.011</td></tr><tr><td>BSW06-03</td><td>7.9</td><td>9.40</td><td>180</td><td>545</td><td>0.0154</td><td>13.5</td><td>12.6</td><td>45.4</td><td>0.0539</td><td>&lt; LOD</td><td>654.6</td><td>0.120</td><td>0.7035378</td><td>2.6</td><td>0.278</td><td>0.006</td></tr><tr><td>BSW06-04</td><td>16.0</td><td>9.41</td><td>160</td><td>550</td><td>0.0080</td><td>15.7</td><td>13.5</td><td>46.3</td><td>0.0611</td><td>&lt; LOD</td><td>667.1</td><td>0.123</td><td>0.7036699</td><td>2.8</td><td>0.273</td><td>0.006</td></tr><tr><td>BSW06-05</td><td>33.8</td><td>9.50</td><td>141</td><td>561</td><td>0.0062</td><td>13.9</td><td>14.2</td><td>44.9</td><td>0.0674</td><td>&lt; LOD</td><td>689.6</td><td>0.122</td><td>0.7038923</td><td>2.5</td><td>0.286</td><td>0.006</td></tr><tr><td>BSW06-06</td><td>54.9</td><td>9.56</td><td>122</td><td>582</td><td>0.0105</td><td>12.4</td><td>18.9</td><td>62.4</td><td>0.1114</td><td>&lt; LOD</td><td>736.3</td><td>0.141</td><td>0.7038094</td><td>2.5</td><td>0.276</td><td>0.006</td></tr><tr><td>BSW06-07</td><td>92.0</td><td>9.50</td><td>104</td><td>608</td><td>0.0200</td><td>10.6</td><td>23.9</td><td>78.8</td><td>0.1706</td><td>&lt; LOD</td><td>806.6</td><td>0.132</td><td>0.7041144</td><td>2.6</td><td>0.260</td><td>0.006</td></tr><tr><td rowspan="7">BSW07<br>290 °C<br>crystalline basalt</td><td>BSW07-01</td><td>0.9</td><td>4.93</td><td>219</td><td>512</td><td>30.6</td><td>25.8</td><td>14.3</td><td>2.25</td><td>0.0025</td><td>4.7</td><td>631.3</td><td>3.951</td><td>0.7040288</td><td>4.1</td><td>-0.057</td><td>0.009</td></tr><tr><td>BSW07-02</td><td>3.8</td><td>8.40</td><td>201</td><td>516</td><td>5.0</td><td>9.8</td><td>15.5</td><td>29.9</td><td>0.0369</td><td>0.50</td><td>648.0</td><td>0.175</td><td>0.7035317</td><td>2.9</td><td>0.082</td><td>0.007</td></tr><tr><td>BSW07-03</td><td>8.7</td><td>8.92</td><td>183</td><td>515</td><td>0.55</td><td>14.6</td><td>15.6</td><td>38.0</td><td>0.0556</td><td>0.044</td><td>656.7</td><td>0.132</td><td>0.7036126</td><td>2.7</td><td>0.125</td><td>0.007</td></tr><tr><td>BSW07-04</td><td>16.9</td><td>9.01</td><td>165</td><td>511</td><td>0.41</td><td>15.6</td><td>15.8</td><td>39.3</td><td>0.0619</td><td>0.033</td><td>660.2</td><td>0.124</td><td>0.7038432</td><td>2.9</td><td>0.121</td><td>0.007</td></tr><tr><td>BSW07-05</td><td>34.6</td><td>9.03</td><td>148</td><td>528</td><td>0.29</td><td>16.6</td><td>16.4</td><td>41.7</td><td>0.0742</td><td>0.032</td><td>678.0</td><td>0.062</td><td>0.7043048</td><td>2.6</td><td>0.162</td><td>0.006</td></tr><tr><td>BSW07-06</td><td>55.7</td><td>9.13</td><td>128</td><td>543</td><td>0.201</td><td>16.3</td><td>17.0</td><td>43.8</td><td>0.0903</td><td>&lt; LOD</td><td>700.5</td><td>0.057</td><td>0.7047286</td><td>2.9</td><td>0.202</td><td>0.007</td></tr><tr><td>BSW07-07</td><td>92.8</td><td>9.22</td><td>110</td><td>589</td><td>0.127</td><td>14.8</td><td>18.4</td><td>48.6</td><td>0.1124</td><td>&lt; LOD</td><td>754.2</td><td>0.113</td><td>0.7051157</td><td>2.6</td><td>0.195</td><td>0.006</td></tr></tbody></table>
1045 **Figures**

1046 **Fig. 1:** Schematic design of Ti-autoclaves used for the experiments in this study. The Ti-autoclaves are rocked to ensure mixing of the solids with the seawater. The cooling apparatus attached to a valve allows fluid samples to be taken during the experimental run.

1049 **Fig. 2:** Scanning electron microscopy images of the starting basalt powders used in the 250 °C experiments (a, b) and of the solid alteration products recovered from experiments BSW04 and BSW05 (c-f).

1052 **Fig. 3:** Aqueous concentrations of Ca, Mg, $SO_4^{2-}$, Na, Cl, Fe, and Si as well as pH of seawater reacting with basaltic glass (solid lines) and crystalline basalt (dashed lines) at 250 °C. The errors bars show the 95 % confidence intervals (~2 standard errors), which are smaller than the symbol size where no error bars are shown. Initial Si concentrations (below the detection limit) are shown as zero.

1057 **Fig. 4:** Fluid sample composition of experiments with basaltic glass and crystalline basalt at 290 °C (cf. Fig. 3).

1059 **Fig. 5:** Strontium concentration, radiogenic $^{87}$Sr/$^{86}$Sr and stable $\delta^{88/86}$Sr isotope ratios of seawater reacting with basaltic glass and crystalline at temperatures of 250 °C (left side) and 290 °C (right side). The dashed lines show the isotopic composition of the basalt used in the experiments. The 95 % confidence intervals (~2 standard errors) for all values are smaller than the symbol size.

1064 **Fig. 6:** Evolution of aqueous Mg and Ca concentrations during the experiments in comparison with similar experiments of Seyfried and Bischoff (1979; 1981) conducted at temperatures of 150 °C and 300 °C.

1067 **Fig. 7:** Model predictions (dashed curves) of anhydrite precipitation as well as aqueous Sr concentrations, $^{87}$Sr/$^{86}$Sr ratios and stable $\delta^{88/86}$Sr values plotted as a function of basalt dissolution during the first 24 h. The symbols show the initial seawater and first sample compositions, and the errors bars and shaded areas show the 95 % confidence intervals, which are smaller than the symbol size where no error bars are shown. For the crystalline experiment at 250 °C, the $\delta^{88/86}$Sr model evolution between 7 and 14 days is shown as well.

1073 **Fig. 8:** Model predictions (dashed curves) of aqueous Sr concentrations and $^{87}$Sr/$^{86}$Sr during the first 24 hours for the 250 °C and 290 °C experiments. The symbols connected by solid lines show the later evolution of the liquids. The grey arrows and labels show the processes leading to the changes in the two parameters during the basalt-fluid reactions. Errors bars and shaded areas show the 95 % confidence intervals, which are smaller than the symbol size where no error bars are shown.

1079 **Fig. 9:** Plot of best-fit fractionation factors $\epsilon_{Anhydrite-Liquid}^{88/86}$ vs. $\epsilon_{CrystBasalt-Liquid}^{88/86}$ (fractionation during non-stoichiometric dissolution of crystalline basalt at (a) 250 °C and (b) 290 °C). The blue point shows the maximum likelihood value while the shaded region is the 95 % confidence interval. Note the different scale of $\epsilon_{CrystBasalt-Liquid}^{88/86}$ in (a) and (b).
<figure><img src="image_2.png" /></figure>

Rocking
motion

10 cm
<figure><img src="image_3.png" /><figcaption>Crystalline basalt</figcaption></figure>

<figure><img src="image_4.png" /></figure>

<figure><img src="image_5.png" /></figure>

<figure><img src="image_6.png" /></figure>

<figure><img src="image_7.png" /></figure>

<figure><img src="image_8.png" /></figure>
250 °C

1 day

1 week

1 month

pH

pH

Basaltic glass

Crystalline Basalt

Ca, Mg, SO₄²⁻ / mmol/kg

<figure><img src="image_9.png" /></figure>

Ca, Mg, SO₄²⁻ / mmol/kg

<figure><img src="image_10.png" /></figure>

Na, Cl / mol/kg

<figure><img src="image_11.png" /></figure>

Na, Cl / mol/kg

<figure><img src="image_12.png" /></figure>

<figure><img src="image_13.png" /></figure>

Si / mmol/kg

<figure><img src="image_14.png" /></figure>

Seawater

Duration / h
290 °C

<figure><img src="image_15.png" /></figure>

<figure><img src="image_16.png" /></figure>

<figure><img src="image_17.png" /></figure>

<figure><img src="image_18.png" /></figure>
250 °C

1 day

1 week1 month

<figure><img src="image_19.png" /></figure>

<figure><img src="image_20.png" /></figure>

<figure><img src="image_21.png" /></figure>

290 °C

1 day

1 week1 month

<figure><img src="image_22.png" /></figure>

<figure><img src="image_23.png" /></figure>

<figure><img src="image_24.png" /></figure>
1 day 1 week 1 month

Mg / mmol/kg

(a)

<figure><img src="image_25.png" /></figure>

Ca / mmol/kg

(b)

<figure><img src="image_26.png" /></figure>

Seawater

Duration / h
250 °C

<figure><img src="image_27.png" /></figure>

290 °C

<figure><img src="image_28.png" /></figure>

<figure><img src="image_29.png" /></figure>

<figure><img src="image_30.png" /></figure>

<figure><img src="image_31.png" /></figure>

<figure><img src="image_32.png" /></figure>

<figure><img src="image_33.png" /></figure>

<figure><img src="image_34.png" /></figure>
<figure><img src="image_35.png" /></figure>

<figure><img src="image_36.png" /></figure>
<figure><img src="image_37.png" /></figure>

<figure><img src="image_38.png" /></figure>
Note the presence of amorphous phases in all samples. The hkl- reflections correspond mainly to ferrous saponite. Zincite (20 wt.%; grey bars) was used as an internal standard.

<figure><img src="image_39.png" /></figure>
Note the absence of amorphous phases in all samples. The hkl- reflections correspond to ferrous saponite. Zincite (20 wt.%; grey bars) was used as an internal standard.

<figure><img src="image_40.png" /></figure>
250 °C

<figure><img src="image_41.png" /></figure>

290 °C

<figure><img src="image_42.png" /></figure>