New insight on Li and B isotope fractionation during serpentinization derived from batch reaction investigations
Introduction
Serpentinization is a key process in the geochemical cycling between seawater and lithosphere in oceanic spreading environments and subduction zones. In particular along slow and ultraslow spreading mid-ocean ridges, ultramafic rocks are exposed to seawater at the seafloor in oceanic core complexes that form as a consequence of detachment faulting and exhumation (Smith et al., 2008, Smith et al., 2012). Serpentinization is also ubiquitous in rifted continental margin settings (e.g. Beard and Hopkinson, 2000) and it might even affect the lithospheric mantle in bent faults of the outer rise of subduction zones (Ranero et al., 2003), although the extent of the latter is currently debated (Korenaga, 2017). Finally, serpentinization of the mantle wedge and deserpentinization of serpentinites in the down going slab in subduction zone settings have been assigned a key role in element cycling (Bostock et al., 2002, Benton et al., 2004, Savov et al., 2005, Deschamps et al., 2011, Pabst et al., 2012).
A critical consequence of serpentinization is the profound modification of the lithologies’ physical properties, such as rheology, density, and magnetic susceptibility, as well as of heat budgets. The process is also important as a potential energy source for deep life as it releases significant portions of H2 and CH4 (Sleep et al., 2004).
The formation and breakdown of serpentine is also key to the cycling of water in the Earth’s lithosphere (Rüpke et al., 2004). Recent estimates suggest that serpentinites with up to 13 wt.% H2O might account for as much as 25 vol.% of the seafloor associated with slow spreading ridges (Cannat et al., 2010, Alt et al., 2012, Alt et al., 2013). With serpentine minerals as a water carrier, considerable quantities of water could potentially be transferred deep down into subduction zones, where different serpentine modifications destabilize and liberate water at individual depth (Deschamps et al., 2011, Korenaga, 2017). The upward migration of this water triggers partial melting or secondary serpentinization within the overlying mantle wedge and contributes to subduction zone fluxing and arc volcanism (Marschall et al., 2007). Related to this comprehensive transfer of water is also an exchange of fluid-mobile elements like chlorine (Cl), caesium (Cs), lithium (Li) and boron (B), between these different reservoirs (Straub and Layne, 2002, Straub and Layne, 2003).
In recent years, B and Li received increasing attention as potential tracers for mass transfer processes that involve hydration and dehydration reactions (Marschall et al., 2007, Vils et al., 2008). The strong elemental partitioning between rock and water is critical in making these tracers useful. Moreover, Li and B show characteristic isotope fractionations that can be used as proxies for water-rock interactions. A quantitative treatment of B and Li element partitioning and associated isotope effects are often difficult because it is hard to deconvolute the effects of compositional and physical (P, T, fluid flux) variability. In approaching this problem, clues about fundamental characteristics of the respective element cycles and the nature of fractionation mechanisms have been gained by investigating samples of vent fluids, altered oceanic crust and abyssal serpentinites collected in the framework of ocean drilling programs (Foustoukos et al., 2004, Vils et al., 2008, Vils et al., 2009, Deschamps et al., 2011, Brant et al., 2012, Yamaoka et al., 2015).
It is clear that B is highly concentrated in the seawater reservoir (4.5 μg/g; Spivack and Edmond, 1987), whereas mantle peridotite is depleted in B (0.19 µg/g for primitive mantel to 0.077 μg/g for depleted mantle; Marschall et al., 2017). Boron can become highly enriched in serpentines formed in the course of low to moderate temperature (<300 °C) alteration of peridotites and associated serpentine formation (Seyfried and Dibble, 1980, Bonatti et al., 1984). Serpentine mineral phases recovered from the seafloor show B contents > 100 µg/g. Similarly high B contents were determined in serpentinized peridotites from drill holes in the Mid-Atlantic Ridge (MAR) 15°20′N Fracture zone area (10–91 μg/g in whole rock; Vils et al., 2008, Boschi et al., 2008, Harvey et al., 2014; and up to 138 μg/g in serpentine mineral spot analyses; Vils et al., 2008) (see Table 1).
Boron isotopic composition is uniform in seawater (δ11BSeawater = +39.61‰; Foster et al., 2010) but highly variable in serpentinites, ranging between compositions from close to seawater down to those of the fresh precursor mantle peridotite (δ11BPeridotite = −7.1‰; Marschall et al., 2017). Given the exceedingly low initial B concentration of peridotites and the commonly high B concentrations (10–112.0 µg/g; Spivack and Edmond, 1987, Deschamps et al., 2013) and increased δ11B values (δ11BSerpentinite values of +8.3 to +40.7‰; Spivack and Edmond, 1987, Boschi et al., 2008, Vils et al., 2009, Harvey et al., 2014) of abyssal serpentinites, it is widely accepted that boron in abyssal serpentinite is almost exclusively derived from seawater. This assertion implies, that the serpentine mineral phases formed preferentially incorporate 10B. Nevertheless, this wide range in B concentrations and isotope compositions in serpentinized peridotites seems to reflect a large variability in the extent of B partitioning and isotope fractionation in the course of serpentinization. Temperature and pH variations were assigned an important role in creating this variability (Bonatti et al., 1984, Spivack and Edmond, 1987, Foustoukos et al., 2008). Important factors are also the integrated water – rock ratios (W/R), the total degree of hydrothermal alteration and the effect of evolving compositions along the fluids’ flow paths within the lithosphere (Boschi et al., 2008, Vils et al., 2009) (see Table 1).
In contrast to B, Li concentrations in unaltered peridotite (1.39 μg/g for primitive and 1.20 µg/g for depleted mantle; Marschall et al., 2017) or mantle olivine (0.52–1.3 μg/g; Vils et al., 2008) are high relative to seawater (0.18 μg/g; Li, 1982). Consequently, mixing of Li derived from the two isotopically distinct rock and seawater reservoirs (δ7LiSeawater = +30.8‰, δ7LiMantle = +3.4‰; Rosner et al., 2007, Tomascak et al., 2008) plays a crucial role in setting Li isotopic compositions of serpentinites. Altered peridotites can be enriched or depleted in Li (0.07–3.37 μg/g; Vils et al., 2008 – or potentially even up to 15 µg/g; Marschall et al., 2017 – Supplementary), but on average they have a concentration (0.67 μg/g) which is about half of what is found in fresh mantle rocks (Vils et al., 2008). Serpentinites originated from pyroxene-poor harzburgites or dunites, are usually depleted in Li, but rocks in which most of the Li is bound to clinopyroxene retain elevated contents of Li, because clinopyroxene often does not break down during serpentinization and may even scavenge Li (Vils et al., 2008). Leaching of Li from the underlying lithology is however indicated by high-temperature vent fluids in the Logatchev and Rainbow hydrothermal fields that have roughly ten times higher Li concentrations than seawater (1.7–2.4 µg/g, Charlou et al., 2002 and references therein). The heavier 7Li is preferentially leached during serpentinization of olivine-rich rocks, resulting in a relative enrichment of 6Li and δ7LiSerp.Peridotite values as low as −28.46‰ (Vils et al., 2009).
In previous studies of Li and B systematics of abyssal serpentinites (Vils et al., 2008, Vils et al., 2009, Boschi et al., 2008, Foustoukos et al., 2008), mechanisms of partitioning and isotope fractionation were in part derived empirically from rock and fluid data (Seyfried and Dibble, 1980, Janecky, 1982, Janecky and Seyfried, 1986). Employing this approach requires a comprehensive understanding of the specific characteristics and the history of the studied systems. The problem is that the temperature evolution of a multi-stage alteration process as well as integrated water-rock ratios are often not well known and have to be deduced from other data (δ18O, 87Sr/86Sr).
Several studies investigated the effect of B speciation in aqueous solutions and its dependency on the pH and salinity of fluids (Kakihana et al., 1977, Palmer et al., 1987, Pokrovski et al., 1995). In aqueous solutions, boron is either trigonally or tetrahedrally coordinated by hydroxyl-groups. This speciation is pH-dependent, such that the trigonal B(OH)3(aq) species is more abundant at pH < 8, and the tetrahedral species (B(OH)4−) predominates at higher pH (Kakihana et al., 1977). Differences in inter-atomic bond strength translating to variations in vibrational energy and symmetry between the two species and lead to a preferred incorporation of 11B in the trigonal site, whereas 10B is enriched in the tetrahedrally coordinated species (Pokrovski et al., 1995, Chacko et al., 2001). This fractionation in isotopes between different aqueous B species is then transferred to mineral phases (such as carbonates, clays or serpentine), when B is incorporated or adsorbed (Liu and Tossell, 2005). Additional and coordination-unrelated fractionation is associated with the transition from an aqueous species to mineralogically bound B. Fractionation factors for both mechanisms were derived from ab initio molecular orbital calculations (Liu and Tossell, 2005) and vibrational spectroscopic data (Sanchez-Valle et al., 2005). While these studies addressed the effect of pH, P, T, and salinity to some extent, predicting the connected effects of all variations in all these parameters is not possible based on existing models. Moreover, other mineral-solution interactions during consecutive adsorptive and structural uptake (e.g. exchange between several intermediate species) are insufficiently understood and hence not considered in these models.
Additionally, the application of the existing theoretical fraction models to serpentinization processes is problematic, as the system involves many different mineral phases (e.g., different serpentine polymorphs, brucite, olivine, etc.). In minerals formed during serpentinization, trigonally coordinated B is adsorbed onto brucite surfaces (Pokrovsky et al., 2005), whereas tetrahedrally coordinate B is structurally incorporated into tetrahedral sites in serpentine (Früh-Green et al., 2004, Pabst et al., 2011). These differences, combined with the isotope effects outlined above, result in variably mineral-bound fractions of B that are isotopically fractionated in opposite directions.
Coordination changes between fluid and minerals are also crucial for Li fractionation (Wunder et al., 2011). In aqueous solution, Li commonly occurs in fourfold coordination ([Li(H2O)4]+) (Yamaji et al., 2001), whereas it can be variably coordinated in different mineral phases. In serpentine (and most other secondary phases) Li is octahedrally coordinated (Li[6]; Wenger and Armbruster, 1991). Similar to the B fractionation behavior, the lighter isotope (6Li) preferentially occupies the higher coordinated site, and hence it becomes enriched in the mineral phase. Then again, Li coordination in the fluid varies with pressure, and variations in Li-O bond lengths are key to isotope fractionation (Wunder et al., 2010, Wunder et al., 2011). Moreover, potential storage of Li in chrysotile nanotubes, an individual structural roll-like form of this serpentine polymorph that forms to accommodate the general 1:1 octahedral (outside) to tetrahedral (inside) layer misfit, can further impact the fractionation process (Wunder et al., 2011).
Finally, for both Li and B, the observed fractionation is strongly dependent on the intensity of fluid fluxing. If incoming fluids are externally buffered (high integrated W/R) they maintain a rather constant isotopic composition throughout the alteration process. If, on the other hand, a batch of fluid is isolated from a large external fluid reservoir and allowed to evolve upon deep crustal penetration, Rayleigh-type fractionation occurs and large isotopic variability may develop (Boschi et al., 2008, Foustoukos et al., 2008). The simultaneous interaction of so many factors in natural systems complicates the empirical determination of isotope fractionation as well as the application of theoretical models.
Other studies therefore aimed for an experimental quantification of elemental and isotopic fractionation. Fundamental insights into Li and B behavior could be derived from a number of ultra-high pressure (2–3.5 GPa) and moderately high temperature (500–900 °C) experiments (Wunder et al., 2005, Wunder et al., 2006, Wunder et al., 2007). In these studies, Li or B bearing minerals were first precipitated from solutions and isotope fractionation between minerals and fluid was determined (Δ11BM-F or Δ7LiM-F). The synthetic minerals were then equilibrated with fluids of known composition far from expected equilibrium. Both steps yielded similar Δ11BM-F or Δ7LiM-F values, which allowed determination of the temperature-dependent fractionation for near neutral and alkaline solutions. Marschall et al. (2009) also emphasized the importance of different degrees of B excess in these experiments when investigating Rayleigh-type fractionation. The experimental studies have in common that phases, scales and conditions far from those observed in typical serpentinization environments were used. Seyfried et al., 1984, Seyfried and Dibble, 1980 examined the geochemical behavior of B and Li during seawater – peridotite/basalt interaction and found that both elements are lost from the solution at low temperatures (<150 °C) but leached from minerals at higher temperatures.
To date, an experimental study investigating Li and B partitioning and isotope fractionation during serpentinization does not exist even though it was repeatedly pointed out (Boschi et al., 2008, Foustoukos et al., 2008, Wunder et al., 2010) that such data is needed.
In order to fill this critical gap in knowledge, we conducted multiple batch experiments reacting fresh olivine with seawater-like fluids. Pressure (40 MPa) and temperature (100 and 200 °C) were chosen to reflect typical serpentinization conditions in oceanic core complexes. Evolving fluid composition as well as geochemistry and mineralogy of solid alteration products were evaluated to derive new insights on Li and B partitioning and isotope fractionation associated with serpentinization.
Section snippets
Experimental setup
All experiments were conducted using a custom-built flexible reaction cell setup, initially introduced by Dickson et al. (1963) and later significantly modified by Seyfried et al. (1987). The reacting components were transferred into a flexible gold bag (Vtotal ≈ 100 ml), which was sealed off through a titanium fitting. Prior to experimentation, the fitting was repeatedly treated with 6.02 M HCl and 14.35 M HNO3, and combusted for 3 h at 450 °C to generate an inert Ti-oxide layer. This reaction
Solid reactant and product phase characterization
Solid phases from both experiments were carefully characterized with respect to their mineralogy and geochemistry prior and subsequent to the autoclaving process.
Serpentinization
For the 200 °C experimental run, the observed product mineral phase assemblage, its mineralogical characteristics and the corresponding fluid geochemistry are in good agreement with well-known serpentinization processes established in numerous earlier studies (e.g. Seyfried et al., 2007, McCollom and Bach, 2009, Klein et al., 2009).
The primary olivine is clearly out of equilibrium in the presence of water at experimental conditions (200 °C, 40 MPa) as indicated by the documented distinct
Summary and conclusions
A first experimental investigation on B and Li partitioning and isotope fractionation during serpentinization provides new perspectives on the geochemical behavior of these elements.
Supporting earlier findings (Spivack and Edmond, 1987, Boschi et al., 2008, Vils et al., 2008), B becomes distinctly enriched in serpentinite. While keeping in mind that initial fluid B concentration used in the experiments was approximately 2.4× higher than in seawater, fluid mass balancing indicated up to 55.61
Acknowledgments
The authors would like to thank P. Monien for his assistance with ICP-MS measurements and helpful discussion input. Many thanks also go to P. Witte, T. Frederichs and M. Wendschuh for their guidance with SEM, AGFM and XRD measurements. This study was funded by the DFG Koselleck project grant BA 1605/10-1. We also acknowledge funding from the DFG Research Instrumentation Program grant INST 144/308-1. Insightful comments and suggestions from Dionysis Foustoukos, Horst Marschall and one anonymous
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