Impact of ambient conditions on the Si isotope fractionation in marine pore fluids during early diagenesis

Benthic fluxes of dissolved silicon (Si) from sediments into the water column are driven by the dissolution of biogenic silica (bSiO₂) and terrigenous Si minerals and modulated by the precipitation of authigenic Si phases. Each of these processes has a specific effect on the isotopic composition of silicon dissolved in sediment pore fluids, such that the determination of pore fluid δ³⁰Si values can help to decipher the complex Si cycle in surface sediments. In this study, the δ³⁰Si signatures of pore fluids and bSiO₂ in the Guaymas Basin (Gulf of California) were analyzed, which is characterized by high bSiO₂ accumulation and hydrothermal activity. The δ³⁰Si signatures were investigated in the deep basin, in the vicinity of a hydrothermal vent field, and at an anoxic site located within the pronounced oxygen minimum zone (OMZ). The pore fluid δ³⁰Si_pf signatures differ significantly depending on the ambient conditions. Within the basin, δ³⁰Si_pf is essentially uniform, averaging $+\mathrm{1.2}\pm \mathrm{0.1}$ ‰ (1 SD). Pore fluid δ³⁰Si_pf values from within the OMZ are significantly lower (0.0±0.5 ‰, 1 SD), while pore fluids close to the hydrothermal vent field are higher ($+\mathrm{2.0}\pm \mathrm{0.2}$ ‰, 1SD). Reactive transport modeling results show that the δ³⁰Si_pf is mainly controlled by silica dissolution (bSiO₂ and terrigenous phases) and Si precipitation (authigenic aluminosilicates). Precipitation processes cause a shift to high pore fluid δ³⁰Si_pf signatures, most pronounced at the hydrothermal site. Within the OMZ, however, additional dissolution of isotopically depleted Si minerals (e.g., clays) facilitated by high mass accumulation rates of terrigenous material (MAR_terr) is required to promote the low δ³⁰Si_pf signatures, while precipitation of authigenic aluminosilicates seems to be hampered by high water ∕ rock ratios. Guaymas OMZ δ³⁰Si_pf values are markedly different from those of the Peruvian OMZ, the only other marine OMZ setting where Si isotopes have been investigated to constrain early diagenetic processes. These differences highlight the fact that δ³⁰Si_pf signals in OMZs worldwide are not alike and each setting can result in a range of δ³⁰Si_pf values as a function of the environmental conditions. We conclude that the benthic silicon cycle is more complex than previously thought and that additional Si isotope studies are needed to decipher the controls on Si turnover in marine sediment and the role of sediments in the marine silicon cycle.

1 Introduction

Silicon (Si) is one of the key macronutrients in the ocean, mainly utilized by siliceous organisms, such as diatoms, radiolarians, or sponges (see the recent review by Sutton et al., 2018). The marine Si cycle is closely linked to the carbon (C) cycle by marine siliceous organisms, which transport C to the sediment and thus exert a strong control on C export from the atmosphere impacting present and past climate (e.g., Lewin, 1961; Tréguer and Pondaven, 2000; Tréguer and De La Rocha, 2013, and recent reviews by Frings et al., 2016, and Sutton et al., 2018). Studies of Si isotopes (δ³⁰Si) have revealed complex uptake and dissolution processes of siliceous organisms, which have a dominant control on the δ³⁰Si distribution in ocean waters (e.g., De La Rocha et al., 1997; Varela et al., 2004; Cardinal et al., 2005; Beucher et al., 2008; Fripiat et al., 2011; Ehlert et al., 2012; Grasse et al., 2013; Sutton et al., 2013; de Souza et al., 2014, 2015). Diatoms constitute the largest part of the Si cycling fluxes in the ocean (Ragueneau et al., 2000) and discriminate between its isotopes during Si uptake, whereby the light isotopes are preferentially incorporated into the diatom frustules (e.g., De La Rocha et al., 1997). Si isotope fractionation during Si uptake is dependent on, e.g., the diatom species, the availability of Fe, and the degree of Si utilization. Fractionation factors between −0.5 ‰ and −2.1 ‰ have been derived from regional water mass mixing and laboratory studies (De La Rocha et al., 1997; Varela et al., 2004; Cardinal et al., 2005; Beucher et al., 2008; Sutton et al., 2013; Meyerink et al., 2017).

After a planktonic bloom, whereupon the nutrients are exhausted, biogenic silica (bSiO₂, mainly diatoms) sinks through the ocean, partially dissolves, and accumulates on the seafloor, where its preservation and recycling is controlled by dissolution and Si re-precipitation processes. The dissolution of bSiO₂ mainly controls the accumulation of silicic acid (Si(OH)₄) in pore fluids, although the in situ concentration remains below the equilibrium concentration of dissolved bSiO₂, which has been explained by simultaneous formation of authigenic silicates (e.g., McManus et al., 1995; Van Cappellen and Qiu, 1997a, b; Rickert et al., 2002). This process is termed reverse weathering, given that the authigenic precipitates are rich in seawater-derived cations, like Na, K, or Mg (Mackenzie et al., 1981; Michalopoulos and Aller, 1995). Experimental studies of bSiO₂ dissolution kinetics revealed a dependence of the bSiO₂ reactivity on sediment depth as well as the ratio between terrigenous material and bSiO₂ (Michalopoulos and Aller, 1995; Van Cappellen and Qiu, 1997a, b; Dixit et al., 2001; Rickert et al., 2002). Marine weathering of terrigenous material (primary silicates like feldspars and secondary silicates like clays) was found to release cations, such as aluminum (Al) or iron (Fe), which reduce the solubility and dissolution rate of bSiO₂ and induce aluminosilicate precipitation (Michalopoulos and Aller, 1995; Van Cappellen and Qiu, 1997a, b; Michalopoulos et al., 2000; Dixit et al., 2001; Rickert et al., 2002; Loucaides et al., 2010). The accumulation of Si in pore fluids and the Si reflux into bottom waters are controlled by three interdependent processes, namely opal dissolution, dissolution of terrigenous solids, and precipitation of authigenic minerals.

Early silica diagenesis has been shown to fractionate Si isotopes as a function of the crystallization state, seawater Si concentration, sedimentation rate, and terrigenous mineral content (Tatzel et al., 2015; Geilert et al., 2016). Since the light ²⁸Si isotope is more reactive compared to the heavier isotopes (²⁹Si, ³⁰Si), processes, such as reverse weathering, adsorption, and direct Si precipitation from saturated solutions, show low δ³⁰Si values in the reaction product and high δ³⁰Si values in its substrate (i.e., fluids) (e.g., Georg et al., 2006a, 2009; Delstanche et al., 2009; Opfergelt et al., 2013; Geilert et al., 2014; Roerdink et al., 2015; Ehlert et al., 2016). The few modeling and experimental studies addressing Si isotope fractionation during formation of secondary phases, report isotope fractionation factors between −1.6 ‰ and −2 ‰ (Ziegler et al., 2005; Méheut et al., 2007; Dupuis et al., 2015; Ehlert et al., 2016).

The δ³⁰Si data for marine pore fluids from the Peruvian margin upwelling region, which is characterized by very high diatom productivity and bSiO₂-rich sediments (Abrantes et al., 2007; Bruland et al., 2005), agree with these findings and clearly indicate Si isotope fractionation as a consequence of authigenic aluminosilicate precipitation accompanied by a Si isotope fractionation factor of −2.0 ‰ (Ehlert et al., 2016). Pore fluids from the Greenland margin and Labrador Sea also reflect early diagenetic reactions detected by pore fluid δ³⁰Si values (between +0.76 ‰ and +2.08 ‰), and the δ³⁰Si values were interpreted as the product of reverse weathering reactions (Ng et al., 2020). In this study, the Guaymas Basin in the Gulf of California was chosen as study area, as it is characterized by a relatively high diatom productivity and sediments that are predominantly composed of diatomaceous muds (up to 50 % diatoms; Kastner and Siever, 1983). Moreover, the Guaymas Basin in the Gulf of California is influenced by hydrothermal activity (e.g., Von Damm, 1990). These bSiO₂-rich sediments are thus ideal for dedicated studies of early diagenesis under the influence of different thermal and redox conditions. We investigated the processes controlling Si isotope fractionation during early diagenesis based on pore fluid and bSiO₂ data from three fundamentally different environmental settings within the Guaymas Basin, including the deep basin, a hydrothermal site, and a site within the oxygen minimum zone (OMZ) on the slope of the basin (Fig. 1). In addition, the Si isotope composition of the water column, bottom waters, and hydrothermal fluids was determined. A numerical transport–reaction model was applied to the OMZ setting to constrain marine weathering processes and to compare the results to the Peruvian margin. The aim of this study was to constrain the factors controlling Si isotope fractionation during early diagenesis and to identify processes influencing bSiO₂ dissolution and authigenic silicate precipitation.

Figure 1 (a) Location map of the sampling stations in the Guaymas Basin, Gulf of California. The black square in the overview map indicates the sampling area. (b) Sedimentary bSiO₂ content at each sampling station. Water column stations were above MUC15-02 (VCTD02) in the basin, at the hydrothermal site (VCTD06, 09), and at the OMZ site (VCTD07).

2 Geological setting, sampling, and methods

2.1 Geological setting

The Guaymas Basin in the Gulf of California is a currently opening continental rifting environment with two graben systems (northern and southern trough), which are offset by a transform fault and reaches spreading rates of up to 6 cm yr⁻¹ (Calvert, 1966). High biological productivity and terrigenous matter input result in high sediment accumulation rates and have produced thick sequences of organic-rich sediments (DeMaster, 1981). Siliceous sediments at the Guaymas slope show fine laminations within the OMZ (dissolved oxygen < 10 µM at ca. 500–900 m water depth; Campbell and Gieskes, 1984) due to the absence of burrowing organisms (Calvert, 1964). The Guaymas Basin is characterized by vigorous hydrothermal activity represented by Black Smoker type vents discovered in both the northern and southern troughs (Berndt et al., 2016; Von Damm et al., 1985). Hydrothermal plumes spread horizontally and mix with deep basin water up to 300 m above the seafloor, resulting in a fraction of hydrothermal fluids of ∼0.1 % in the deep waters of the Guaymas Basin (Campbell and Gieskes, 1984). Hydrothermal sills and dikes intruding into the sediments were found to accelerate early diagenetic reactions (due to the released heat) and change pore fluid geochemistry significantly (Gieskes et al., 1982; Kastner and Siever, 1983; Von Damm et al., 1985; Von Damm, 1990). At present, pore fluids in surface sediments show a seawater composition (Geilert et al., 2018), and the absence of diagenetic high-temperature processes render these pore fluids suitable for studying recent early diagenetic processes.

2.2 Sampling

Sediments were sampled via multicorer (MUC) deployment during RV Sonne cruise SO241 in summer 2015, as described in detail in Geilert et al. (2018). In total, six stations have been investigated: four within the basin (termed basin sites: MUC33-11, MUC22-04, MUC23-05, MUC15-02), one in the vicinity of a hydrothermal vent field (termed hydrothermal site: MUC66-16), and one within the OMZ (termed OMZ site: MUC29-09) (Fig. 1, Table 1). The coring locations within the basin were sampled in water depths between 1726 and 1855 m below sea level (m b.s.l.). The hydrothermal site (MUC66-16) was sampled at 1842 m b.s.l. and was located at a distance of ∼500 m from the active hydrothermal mound described by Berndt et al. (2016). The OMZ site (MUC29-09) was sampled on the Guaymas Basin slope on the Mexican mainland at 665 m b.s.l.

Table 1 Pore fluid Si concentration (µM); δ³⁰Si_pf values (‰) and biogenic silica weight fraction (bSiO₂ in wt %); Al∕Si ratio (mM M⁻¹); ${\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}$ values (‰); porosity (Ø); and Al and K contents (wt %) for the basin sites, hydrothermal site, and OMZ site.

^* Porosity corrected (see Supplement for details).

After core retrieval, bottom water above the sediment was sampled and filtered immediately using 0.2 µm cellulose acetate membrane filters. Bottom water from MUC22-04 may have been contaminated with surface waters during core retrieval, as indicated by Si and Mn concentrations (53.8 and 0.05 µM, respectively) being lower than at the remaining sites within the deep basin, which show distinct anomalies caused by mixing with hydrothermal plume fluids (e.g., MUC15-02: Si = 177.8 µM and Mn = 0.34 µM). Processing of sediments was conducted in a cool laboratory in an argon-flushed glove bag immediately after core retrieval. Sampling intervals were 1–5 cm, with the highest resolution close to the sediment surface and increasing distance down-core. Pore fluids were separated from sediments by centrifugation (20 min at 4500 rpm) and subsequently filtered (0.2 µm cellulose acetate membrane filters) for further analyses.

Water column samples were taken using a video-guided Niskin Water sampler CTD (conductivity, temperature, and depth profiler; Sea-Bird 911plus) rosette system. Water samples were taken in the basin above MUC15-02 at 1844 m b.s.l. (VCTD02), within the hydrothermal plume between 1781 and 1800 m b.s.l. (VCTD06 and 09) above the Black Smoker mound, as described in Berndt et al. (2016), and above the OMZ site (MUC29-09) at 586 m b.s.l. (Fig. 1).

2.3 The bSiO₂ separation and digestion

The bSiO₂ mass fractions of sediment samples were determined using an automated leaching method following Müller and Schneider (1993) at the GEOMAR Helmholtz Centre for Ocean Research Kiel. The sample material was treated with 1 M NaOH at 85 ^∘C to extract the opal fraction. The increase in dissolved Si was monitored and evaluated using a method described by DeMaster (1981). The precision of the mass fraction determination was 5 % to 10 % (1 SD). The bSiO₂ was separated from the sediment for Si isotope analyses following the method of Morley et al. (2004). About 500 mg of a freeze-dried sediment sample was transferred into a centrifuge tube. Organic matter and carbonate material was removed by adding H₂O₂ (Suprapur) and HCl, respectively. Clay particles (grain size <2 µm) were separated from the remaining sediment by the Atterberg method following Stoke's law (Müller, 1967). The remaining sediment (bSiO₂ and heavy minerals) was filtered using a 5 µm filter and subsequently separated from the remaining detritus using heavy liquid separation (sodium–polytungstate solution). The heavy liquid purification method was repeated until clean bSiO₂ samples (>95 %) examined via light microscopy were obtained. Light microscopy revealed that the bSiO₂ fraction essentially consisted of diatoms and only traces of radiolarians and sponges were present (<5 %). The bSiO₂ samples were stored in Milli-Q water (MQ water). The bSiO₂ sample of the OMZ site stems from a nearby gravity core (GC07), which is described in detail in Geilert et al. (2018). The bSiO₂ samples were dissolved following a method by Reynolds et al. (2008). Aliquots of the cleaned bSiO₂ samples were transferred into Teflon vials and dried on a hot plate. Drying of the samples was shown by Ehlert et al. (2012) to have no effect on the Si isotope composition of the samples. Subsequently, 1 mL of 0.1 M NaOH was added and the samples were placed on a hot plate at 130 ^∘C for 24 h. After sample digestion, the supernatant and residue (undissolved traces of radiolarians and sponges) were separated via centrifugation. The supernatant was treated with 200 µL H₂O₂ (Suprapur) in order to remove remaining organic matter and then dried and redissolved in 1 mL 0.1 M NaOH at 130 ^∘C for 24 h. After the digestion procedure, the samples were diluted with MQ water and neutralized with 1 M HCl.

2.4 XRD measurements

X-ray diffraction (XRD) analyses of the dried clay samples were performed at the Central Laboratory for Crystallography and Applied Material Science, ZEKAM, Dept. of Geosciences, University Bremen, using a Philips X'Pert Pro multipurpose diffractometer. The diffractometer was equipped with a Cu tube, a fixed divergent slit of 0.25^∘, a secondary Ni filter, and a X'Celerator detector system. A continuous scan from 3 to 85^∘ 2θ was applied for the measurements with a calculated step size of 0.016^∘ 2θ (calculated time per step this was 50 s). Quantification of mineral phases were based on the Philips software X'Pert High Score™, the freely available X-ray diffraction software MacDiff 4.25 (Petschick et al., 1996), and the QUAX full-pattern method, following Vogt et al. (2002). The standard deviation is ±1 %–3 % for well-crystallized minerals (see also Vogt et al., 2002) and ±5 % for the remaining mineral phases.

2.5 Geochemical analyses of fluid and solid phases

Analyses of major and trace element concentrations of pore fluids from the basin sites, the hydrothermal site, and the water column are described in Geilert et al. (2018). OMZ site pore fluids were treated in the same way. In brief, the pore fluids were analyzed on board by photometry (NH₄) and on shore for dissolved anions (Cl) and cations (Si, K, Na, Mg) using ion chromatography (IC, METROHM 761 Compact, conductivity mode) and inductively coupled plasma optical emission spectrometry (ICP-OES, VARIAN 720-ES), respectively. Analytical precision was constrained using the IAPSO seawater standard for all chemical analyses (Gieskes et al., 1991) and was found to be <1 % for Cl; <2 % for K, Na, and Mg; and <5 % for Si.

Freeze dried and ground sediment samples were digested in HF (40 % Suprapur), HNO₃ (Suprapur), and HClO₄ (60 % p.a.) for major element analyses. The accuracy of the method was tested by method blanks and the reference standards SDO-1 (Devonian Ohio Shale, USGS) and MESS-3 (Marine Sediment Reference Material, Canadian Research Council). The digested samples were measured for their K and Al contents by ICP-OES (VARIAN 720-ES) and reproducibility was ≤5 %. Total carbon (TC) and total organic carbon (TOC) were measured in freeze-dried and ground sediment samples by flash combustion using the Carlo Erba Element Analyzer (NA-1500). Carbonate carbon (CaCO₃) was calculated by subtracting TOC from TC.

The digested bSiO₂ samples were analyzed for their Al and Si contents using the Agilent 7500 series quadrupole ICPMS at GEOMAR to provide information about potential clay contamination of the separated bSiO₂ fraction (Shemesh et al., 1988), whereby Al∕Si ratios below 50 mM M⁻¹ are considered negligible clay contamination (van Bennekom et al., 1988; Hurd, 1973). Al∕Si ratios in the studied bSiO₂ ranged between 15 and 39 mM M⁻¹, with three exceptions in MUC22-04, MUC15-02, and MUC66-16 yielding Al∕Si ratios of 71, 57, 50 mM M⁻¹, respectively. However, all ${\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}$ values agreed well with surrounding ${\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}$ values, and clay contamination is thus considered unimportant. All other bSiO₂ samples are considered clay-free.

2.6 Sample purification and Si isotope measurements

Fluid and digested bSiO₂ samples were prepared for Si isotope measurements following the purification method of Georg et al. (2006b). The concentration of the samples was adjusted and loaded (1 mL with ∼64 µM Si) onto 1 mL pre-cleaned cation exchange resin (Biorad AG50 W-X8) and subsequently eluted with 2 mL MQ water. Matrix effects originating from dissolved organic compounds and anions, which cannot be separated by this purification method, have previously been found to potentially influence Si isotope measurements (van den Boorn et al., 2009; Hughes et al., 2011). However, no influence of the matrix effects on pore fluid Si isotope measurements has been found during our measurements following several tests described in Ehlert et al. (2016). Briefly, Ehlert et al. (2016) removed organic compounds via H₂O₂ and SO₄ via Ba addition, yielding δ³⁰Si values identical to untreated samples, within error. Therefore, our samples were not treated with H₂O₂ or Ba before sample purification and Si isotope measurements.

Si isotope samples were measured in medium resolution on a NuPlasma MC-ICPMS (Nu InstrumentsTM, Wrexham, UK) at GEOMAR using the Cetac Aridus II desolvator. Sample Si concentrations of about 21 µM resulted in a ²⁸Si intensity of 3 to 4 V. The MQ blank was ≤3 mV, resulting in a blank to signal ratio <0.1 %. The measurements were performed using the standard-sample bracketing method to account for mass bias drifts of the instrument (Albarède et al., 2004). Si isotopes are reported in the δ³⁰Si notation, representing the deviation of the sample ³⁰Si∕²⁸Si from that of the international Si standard NBS28 in per mill (‰). Long-term δ³⁰Si values of the reference materials Big Batch ($-\mathrm{10.6}\pm \mathrm{0.2}$ ‰; 2 SD; n=49), IRMM018 ($-\mathrm{1.5}\pm \mathrm{0.2}$ ‰; 2 SD; n=48), Diatomite ($+\mathrm{1.3}\pm \mathrm{0.2}$ ‰; 2 SD; n=44), and BHVO-2 ($-\mathrm{0.3}\pm \mathrm{0.2}$ ‰; 2 SD; n=13) are in good agreement with δ³⁰Si values in the literature (e.g., Reynolds et al., 2007; Zambardi and Poitrasson, 2011). The seawater intercalibration standard Aloha (1000 m) resulted in $+\mathrm{1.3}\pm \mathrm{0.2}$ ‰ (2 SD; n=8) in very good agreement with Grasse et al. (2017). Additionally, two in-house matrix standards have been measured. The pore fluid matrix standard yielded an average δ³⁰Si value of $+\mathrm{1.3}\pm \mathrm{0.2}$ ‰ (2 SD; n=17) and the diatom matrix standard (E. rex) $-\mathrm{1.0}\pm \mathrm{0.2}$ ‰ (2 SD; n=22), which agrees well with earlier reported values (Ehlert et al., 2016). All samples were measured two to four times on different days and the resulting δ³⁰Si values have uncertainties between 0.1 ‰ and 0.4 ‰ (2 SD, Table 1). The δ³⁰Si values of pore fluids, bSiO₂, and bottom water are given as δ³⁰Si_pf, ${\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}$, and δ³⁰Si_BW, respectively. Error bars in the figures indicate the uncertainty of the individual sample measurements (2 standard deviations, 2 SD).

2.7 Numerical model

A numerical reactive-transport model was applied to simulate Si turnover within OMZ site sediments. The model was based on a previously published Si isotope model (Ehlert et al., 2016) and was extended to consider the dissolution of additional phases. A detailed description of the model can be found in the Supplement.

2.8 Calculation of the amount of terrigenous material and mass accumulation rate

The amount of terrigenous material (%) for the Guaymas Basin was calculated as the total mass subtracted by the carbonate content (CaCO₃), the organic matter content (OMC), and the bSiO₂ content (Sayles et al., 2001):

$$\begin{array}{}\text{(1)}& \mathrm{terrigenous}\mathrm{material}=\mathrm{100}-({\mathrm{CaCO}}_{\mathrm{3}}+\mathrm{OMC}+{\mathrm{bSiO}}_{\mathrm{2}}).\end{array}$$

The mass accumulation rate (MAR) was calculated as follows:

$$\begin{array}{}\text{(2)}& \mathrm{MAR}=S\cdot d\cdot (\mathrm{1}-\mathrm{\Phi }),\end{array}$$

with S the sedimentation rate as 0.18 cm yr⁻¹ (Thunell et al., 1994) and d the bulk dry density as 2.5 g cm⁻³. The porosity (Φ) was taken at the deepest part of the core at 0.925.

3 Results

3.1 Water chemistry and sediment composition

All water column, pore fluid, and hydrothermal Si concentration data; bSiO₂ weight fractions; and Si isotope values are reported in Tables 1 and 2. A detailed description of the water column properties, pore fluids, and hydrothermal fluid chemistry can be found in Berndt et al. (2016) and Geilert et al. (2018). Pore fluids predominantly show a seawater composition at all sampling sites and are not influenced by high-temperature processes related to sill intrusions or mixing with hydrothermal fluids. Pore fluid geochemistry of major elements in the OMZ resembles that of the remaining sampling sites with the exception of a strong enrichment in NH₄ (Table S1 and Geilert et al., 2018) and high Fe and low Mn concentrations (Scholz et al., 2019). The porosity-corrected K (see Supplement) and Al contents in the sediments ranged between 0.9 wt % and 21.2 wt % and 2.9 wt % and 66.6 wt %, respectively (Table 1). The TOC contents ranged between 0.3 % and 7.8 % and are shown in Table S1 in the Supplement.

Table 2 Water column and hydrothermal plume Si concentration (µM) and Si isotope values (‰). Additionally, the share of hydrothermal fluids within the hydrothermal plume is given based on the calculation provided by Berndt et al. (2016) in their Supplement.

^* Calculation in Berndt et al. (2016).

3.2 Bottom water, water column, and hydrothermal fluid Si concentrations and δ³⁰Si values

The bottom water Si concentration ranged between 173 and 254 µM for all basin sites and the hydrothermal site (between 1726 and 1855 m b.s.l.) with the exception of MUC22-04, where Si concentrations were as low as 54 µM (possible surface water contamination; see Sect. 2.2 and Table 1). The bottom water δ³⁰Si_BW signatures ranged between +1.5 ‰ and +2.0 ‰ for all basin sites and the hydrothermal site and overlap within error (average δ³⁰Si_BW: $+\mathrm{1.8}\pm \mathrm{0.2}$ ‰, 1 SD; highest δ³⁰Si_BW for the surface-contaminated sample, MUC22-04). Bottom water Si concentration for the OMZ site was 31 µM (665 m water depth). The bottom water within the OMZ site had a distinctly lower δ³⁰Si_BW value of +0.8 ‰. Here, a potential contamination with surface waters can be excluded given that they are characterized by high δ³⁰Si values (from 1.7 ‰ to 4.4 ‰; Ehlert et al., 2012; Grasse et al., 2013), due to the preferential biological uptake of ²⁸Si (De La Rocha et al., 1997).

The basin water (VCTD02), which was sampled about 1 m above the seafloor, had a Si concentration of 163 µM and a δ³⁰Si_deepBasin value of +1.5 ‰. Hydrothermal plume Si concentrations ranged between 253 and 690 µM, and δ³⁰Si values ranged from +0.7 ‰ (VCTD09-06) to +1.4 ‰ (VCTD06-06). The water column within the OMZ (586 m b.s.l.) had a Si concentration of 78 µM and a δ³⁰Si_OMZ value of +1.5 ‰ (Table 2).

3.3 Pore fluid Si concentration and δ³⁰Si_pf values

Pore fluid Si concentrations generally increased asymptotically with depth from bottom water values until reaching maximum average concentrations between 605 and 864 µM Si (Table 1, Fig. 2). For the basin sites MUC33-11 and MUC15-02 and the OMZ site, the Si concentrations asymptotically increased until average values of 742 (≥ 8 cm below seafloor, cm b.s.f.), 729 (≥ 9 cm b.s.f.), and 765 µM (≥9 cm b.s.f.) were reached, respectively. Hydrothermal site Si concentrations were higher and increased to 864 µM (≥4.5 cm b.s.f.). MUC22-04 and MUC23-05 Si concentrations initially increased to values of 605 µM (5.5–11 cm b.s.f.) and 735 µM (3.5–9 cm b.s.f.) on average and then decreased again to 364 µM (26 cm b.s.f.) and 640 µM (22 cm b.s.f.), respectively.

Figure 2 Depth (cm b.s.f.) profiles for all stations for pore fluid Si concentration (Si(OH)₄) in µM (grey symbols) and δ³⁰Si_pf values (colored symbols) and biogenic opal weight fraction (bSiO₂) in wt % (grey symbols) and ${\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}$ values (colored symbols). The dashed line is the δ³⁰Si value of the deep basin (VCTD02), and the dotted line represents the δ³⁰Si value of the water column in the OMZ (VCTD07). The uppermost Si isotope data point in the pore fluid diagrams refers to the bottom water (BW). Note the different depth scale for the OMZ site. The brackets around the MUC22-04 bottom water Si concentration value indicate possible surface water contamination. Error bars not indicated are within symbol size. The long-term error (2 SD) of international standards is indicated in the upper right δ³⁰Si_pf depth profile.

Pore fluid δ³⁰Si_pf values for all basin sites ranged from +0.9 ‰ to +1.5 ‰ (Table 1, Fig. 2), which is lower than the respective bottom water δ³⁰Si_BW values (δ³⁰Si_BW from +1.6 ‰ to +2.0 ‰). The decrease in Si concentration at MUC22-04 and MUC23-05 below 13 and 11 cm b.s.f., respectively, is not reflected in a significant change in δ³⁰Si_pf. The hydrothermal site showed the highest δ³⁰Si_pf values, ranging from +1.8 ‰ and +2.2 ‰, which is higher than the bottom water δ³⁰Si_BW (+1.5 ‰). In contrast, the OMZ site had the lowest δ³⁰Si_pf values, between −0.5 ‰ and +0.8 ‰, which was also characterized by very low δ³⁰Si_BW (+0.8 ‰; Table 1).

3.4 The bSiO₂ content and ${\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}$ values

The bSiO₂ content of the sediments (Table 1, Fig. 2) varied between 4.7 wt % and 47.6 wt %. The lowest contents were present at the basin site MUC22-04 (7.6 wt %–13.1 wt %) and the hydrothermal site (4.7 wt %–14.6 wt %). The remaining sampling sites showed higher bSiO₂ contents of 23±7 wt % (1 SD) on average. The ${\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}$ signatures ranged between +0.4 ‰ and +1.0 ‰ and did not vary systematically with depth or sampling site within error. The small variability in ${\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}$ signatures most likely stems from natural variability within the Guaymas Basin.

3.5 XRD results of the clay fraction

The main silicate mineral phases of all samples were phyllosilicates (16 wt %–59 wt %), primary silicates (quartz, plagioclase, potassium feldspar; 15 wt %–38 wt %), and amorphous SiO₂ (4 wt %–43 wt % including abiogenic and biogenic opal) (Table S2). The phyllosilicates were mainly composed of variable fractions of smectite, illite, montmorillonite, and kaolinite. Apart from silicate minerals, minor amounts of Fe-(hydr)oxides (≤10 wt %, most between 2 wt % and 3 wt %), pyroxenes (≤8 wt %), and carbonates (≤6 wt %) were present. Biogenic opal fragments were identified via light microscopy to be the dominating amorphous silicate phase at all sites besides the hydrothermal site and MUC23-05 in the basin basin. At the hydrothermal site, the abiogenic amorphous silica fraction was the dominating silica phase in the uppermost and lowermost core sections, with only minor occurrences of biogenic opal fragments. Abiogenic amorphous silica was also found in the uppermost and lowermost core sections of MUC23-05.

4 Discussion

Pore fluid Si concentration and δ³⁰Si_pf signatures vary significantly between sampling sites and appear to depend strongly on ambient conditions. The Si concentration and isotope compositions are proposed as being affected by dissolution of bSiO₂, the dissolution of terrigenous phases, and the formation of authigenic aluminosilicates; the latter process is defined as reverse weathering. Dissolution of bSiO₂ is most effective in the reactive surface layer (≤10 cm b.s.f.) where the degree of Si undersaturation is highest. When Si is released into solution via bSiO₂ dissolution, certain amounts of Si re-precipitate as authigenic aluminosilicates as a function of the availability of reactive metals made available by dissolution of terrigenous material (e.g., Michalopoulos and Aller, 1995; Van Cappellen and Qiu, 1997a, b; Loucaides et al., 2010). In the course of this process, authigenic silicate precipitation induces δ³⁰Si fractionation, whereby the ²⁸Si is preferentially incorporated into the solid phase, enriching the remaining fluid in ³⁰Si (Ehlert et al., 2016). In the following sections, the processes during early diagenesis influencing pore fluid δ³⁰Si_pf signatures under different ambient conditions are discussed. For the OMZ site, we quantify these processes using a reactive transport model and compare the results to the only other OMZ site where pore fluid δ³⁰Si_pf data are available, the Peruvian margin.

4.1 Influences on δ³⁰Si_pf due to source mixing

In the open ocean, a strong correlation between the inverse Si concentration (1∕Si) and Si isotope composition in intermediate and deep waters exists, showing low δ³⁰Si values with high Si concentrations, and this can be used to identify water mass mixing between two endmembers with distinct Si characteristics (e.g., de Souza et al., 2012). Here, we use the mixing of two endmembers in Eq. (3) to calculate the mixing between the deep water column and fluids originating from bSiO₂ dissolution according to the following equation:

$$\begin{array}{}\text{(3)}& {\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{\mathrm{mix}}={\displaystyle \frac{{\displaystyle \begin{array}{c}({\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{\mathrm{water}\mathrm{column}}\cdot {\left[\mathrm{Si}\right]}_{\mathrm{water}\mathrm{column}}\cdot f)\\ +({\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}\cdot {\left[\mathrm{Si}\right]}_{{\mathrm{bSiO}}_{\mathrm{2}}}\cdot (\mathrm{1}-f\left)\right)\end{array}}}{\begin{array}{c}\left({\left[\mathrm{Si}\right]}_{\mathrm{water}\mathrm{column}}\cdot f\right)\\ +({\left[\mathrm{Si}\right]}_{{\mathrm{bSiO}}_{\mathrm{2}}}\cdot \left(\mathrm{1}-f\right))\end{array}}},\end{array}$$

with δ³⁰Si_{water column} and [Si]_{water column} as the respective water column Si isotope composition and concentration (Table 2) and ${\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}$ as the average bSiO₂ value (+0.8 ‰) of all sites. The equilibrium concentration in respect to bSiO₂ dissolution was derived from an experimental study by Van Cappellen and Qiu (1997a) with [Si]${}_{{\mathrm{bSiO}}_{\mathrm{2}}}=\mathrm{900}$ µM. Mixing fractions are represented by f, varied over 100 % water column and 0 % fluids affected by bSiO₂ dissolution and vice versa.

Pore fluid δ³⁰Si_pf values of all sites deviate from mixing curves between the deep water column and fluids originating from bSiO₂ dissolution and are obviously affected by additional processes (Fig. 3). Si isotope fractionation during bSiO₂ dissolution is not well constrained and ranges between −0.55 ‰ and 0 ‰, while most studies argue against a significant isotope effect (Demarest et al., 2009; Egan et al., 2012; Wetzel et al., 2014). Therefore, we will also exclude Si isotope fractionation in dependence of bSiO₂ dissolution as a process affecting pore fluid δ³⁰Si_pf values. The δ³⁰Si_pf values of the basin sites and hydrothermal site are higher (+0.9 ‰ to +2.2 ‰) compared to the respective mixing curves, while δ³⁰Si_pf values of the OMZ site are lower (−0.5 ‰ to +0.8 ‰). While a shift to lower δ³⁰Si_pf values points to dissolution of an isotopically light phase, higher δ³⁰Si_pf values indicate that precipitation processes are important, given that the light Si isotope is preferentially incorporated in authigenic secondary phases (Georg et al., 2009). This indicates that processes governing the pore fluid Si isotope composition differ significantly between the individual sites.

Figure 3 Pore fluid δ³⁰Si values are displayed versus the inverse Si concentration (1∕Si) for the basin sites, the hydrothermal site, and the OMZ site. Error bars not indicated are within symbol size. Mixing curves are calculated following Eq. (3) between the respective water column and the average bSiO₂ δ³⁰Si value for all sites (see text for details). The grey-shaded area indicates the uncertainty on the equilibrium solubility of bSiO₂ assuming 900±150 µM following Van Cappellen and Qiu (1997a).

4.2 Influences on δ³⁰Si_pf from terrigenous and biogenic material

The terrigenous ∕ bSiO₂ ratio was found to be the main mechanism controlling asymptotic pore fluid Si concentration and the benthic Si flux (Van Cappellen and Qiu, 1997a, b). Maximum Si concentrations were reached asymptotically at four out of six sampling sites (Fig. 2). Two sites within the basin (MUC22-04 and MUC23-05) show lower Si concentrations in the deep core sections (Fig. 2), which are most likely related to the decrease in reactive silica with depth, caused by the formation of less-soluble silica phases (Van Cappellen and Qiu, 1997a). At these sites, the asymptotic Si concentration is defined as the maximum concentration values in the center of the core. The amount of terrigenous material for the Guaymas Basin was calculated according to Eq. (1) and accounts for 75 %.

Asymptotic Si concentrations plotted against the terrigenous ∕ bSiO₂ ratio fall on the global trend except for the hydrothermal site (Fig. 4a). Here, high geothermal gradients are likely responsible for the higher Si concentrations with respect to the global trend (see also Sect. 4.3.2). In contrast to the asymptotic Si concentration, no strong correlation of the pore fluid δ³⁰Si_pf values with the terrigenous ∕ bSiO₂ ratio exists (Fig. 4b). In order to identify processes responsible for the different pore fluid δ³⁰Si_pf values and to facilitate comparison between the three settings within the Guaymas Basin, only average δ³⁰Si_pf values will be discussed in the following.

Figure 4 Asymptotic Si concentration (a) and the pore fluid δ³⁰Si_pf values (b) as a function of the terrigenous ∕ bSiO₂ ratio for the basin sites, the hydrothermal site, and the OMZ site in the Guaymas Basin. An exponential increase in Si concentrations with decreasing terrigenous ∕ bSiO₂ ratio is observed, which is not reflected by corresponding systematic changes in δ³⁰Si_pf. The values for the terrigenous ∕ bSiO₂ ratio defining the global trend (grey dots) are from the Southern Ocean, Scotia Sea, Norwegian Sea, northeastern Atlantic Ocean, Juan de Fuca Ridge, Arabian Sea, and the Peru Basin (Van Cappellen and Qiu, 1997a; Rabouille et al., 1997; Rickert, 2000).

Average δ³⁰Si_pf values show distinct variations between the individual settings. The δ³⁰Si_pf values of the OMZ site are lower (0.0±0.5 ‰, 1 SD, n=6) than those of the basin sites ($+\mathrm{1.2}\pm \mathrm{0.1}$ ‰, 1 SD, n=17) and the hydrothermal site, which shows the highest δ³⁰Si_pf values ($+\mathrm{2.0}\pm \mathrm{0.2}$ ‰, 1 SD, n=3). The homogeneity of the individual δ³⁰Si_pf profiles indicates that possible effects of bioirrigation at the basin sites and hydrothermal site are quickly compensated for by isotopic exchange. Similarly, bioturbation has also little impact on the sediment composition given the homogeneity of the bSiO₂ content profiles with depth (Fig. 2).

4.3 Influence of the ambient environmental conditions on the δ³⁰Si_pf values

4.3.1 Basin sites

Pore fluid δ³⁰Si_pf values of the basin sites deviate from the mixing curve between the deep water column and fluids originating from bSiO₂ dissolution and are shifted to higher values ($+\mathrm{1.2}\pm \mathrm{0.1}$ ‰; Fig. 3). This shift to higher δ³⁰Si_pf values can be explained by Si re-precipitation as authigenic aluminosilicates, which preferentially incorporate the light ²⁸Si isotope (Fig. 5a). Alteration of terrigenous material leads to mobilization and re-precipitation of Al and the uptake of K from seawater in the authigenic phase (Michalopoulos and Aller, 2004). The sedimentary K∕Al ratio can thus be used to detect these early diagenetic reactions in addition to pore fluid δ³⁰Si_pf values. Authigenic aluminosilicates formed during alteration of terrigenous material were found to have K∕Al ratios of 0.32 (Michalopoulos and Aller, 2004), which is higher than the pristine K∕Al ratio of terrigenous material carried by rivers ($\mathrm{K}/\mathrm{Al}=\mathrm{0.19}$; Viers et al., 2009). The average K∕Al ratio of the basin sites is 0.34±0.01 (1 SD), which is in the same range as K∕Al ratios indicative of authigenic aluminosilicate formation (Michalopoulos and Aller, 2004). Thus, at basin sites both K∕Al ratios and δ³⁰Si_pf values are in agreement, recognizing bSiO₂ dissolution followed by authigenic clay formation as significant processes taking place.

Figure 5 Conceptual model of the processes influencing pore fluid δ³⁰Si_pf values in the Guaymas Basin (a) and the Peruvian margin following Ehlert et al. (2016) (b). Bold values in the sediment show the average pore fluid δ³⁰Si_pf values. Arrow length indicates the dominating process (dissolution versus precipitation). The δ³⁰Si values in the hydrothermal plume indicate dilution with seawater (see Sect. 4.5). (c) The average pore fluid δ³⁰Si_pf values are shown, indicating the dominance of precipitation or dissolution processes for the three settings in the Guaymas Basin and the Peruvian OMZ.

4.3.2 Hydrothermal site

The δ³⁰Si_pf values from the hydrothermal site are higher ($+\mathrm{2.0}\pm \mathrm{0.2}$ ‰) than the mixing curve between the deep water column and fluids originating from bSiO₂ dissolution and are also much higher than pore fluid δ³⁰Si_pf values from the basin (Figs. 3, 5a). The high δ³⁰Si_pf values indicate that precipitation plays a significant role at this site. Sedimentary K∕Al ratios are equivalent to basin values (K∕Al = 0.34) and thus show the formation of authigenic aluminosilicates. Consequently, the higher pore fluid δ³⁰Si_pf values compared to the basin sites can either be explained by a different Si isotope fractionation factor or by the precipitation of additional silica phases. The hydrothermal site is located in close proximity to a hydrothermal vent field and hydrothermal deposits are mainly composed of Fe sulfides (Berndt et al., 2016), while nearby sediments are dominated by amorphous silica, quartz, and Fe–Si silicates (e.g., ferrosilite, fayalite) (Kastner, 1982; Von Damm et al., 1985). Si adsorption to iron (oxyhydr)oxide and incorporation into Fe–Si gels can create substantial Si isotope fractionation, with ²⁸Si being preferentially enriched in the solid phase (Delstanche et al., 2009; Zheng et al., 2016). Pore fluids show high Fe²⁺ concentrations (up to 190 µM; Scholz et al., 2019), and the precipitation of Fe–Si silicates is thus likely, shifting pore fluid δ³⁰Si_pf to the observed high values. Saturation indices calculated based on the program PHREEQC (Parkhurst and Appelo, 1999) indicate that quartz is supersaturated and amorphous silica is close to saturation at about >6 cm b.s.f. (Fig. S1 in the Supplement). Due to the dynamics of hydrothermal systems, this can be subject to changes and supersaturation of amorphous silica is likely to be obtained occasionally, due to the ascent of Si enriched fluids from greater depth. Gieskes et al. (1988) reported on amorphous silica cement in the hydrothermally influenced sediments of the Guaymas Basin, which is supported by findings of this study and likely explains the high amorphous SiO₂ contents identified by XRD (∼35 wt %; see Sect. 3.5; Table S2). In the Guaymas Basin, high thermal gradients (up to 11 K m⁻¹; Geilert et al., 2018) caused by igneous sill intrusions near the active spreading center significantly influence diagenetic reactions at depth and accelerate Si dissolution and precipitation (Fig. 5a) (e.g., Kastner and Siever, 1983). This can also explain the high pore fluid δ³⁰Si_pf values, given that deep Si saturated fluids ascend and Si precipitates (likely along with Fe) over a large temperature range, whereby lower temperatures are associated with larger Si isotope fractionation (Geilert et al., 2014; Zheng et al., 2016) shifting pore fluid δ³⁰Si_pf to the observed high values.

4.3.3 OMZ site

At the OMZ site, the δ³⁰Si_pf values are significantly lower (on average 0.0±0.5 ‰) than the water column δ³⁰Si_{water column} value ($+\mathrm{1.5}\pm \mathrm{0.2}$ ‰) and also lower than the ${\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}$ value (+0.8 ‰) (Figs. 2, 3). Interestingly, the only other δ³⁰Si_pf values from an OMZ were obtained at the Peruvian margin (Ehlert et al., 2016), where the δ³⁰Si_pf values in the upper 10 cm b.s.f. are slightly higher than the water column δ³⁰Si_BW values (+1.8 ‰ and +1.5 ‰, respectively). As Ehlert et al. (2016) concluded, the Peruvian δ³⁰Si_pf values are influenced by bSiO₂ dissolution and precipitation of authigenic aluminosilicates, the latter process shifting pore fluid δ³⁰Si_pf to values higher than those of the water column. Consequently, in order to explain the extremely low δ³⁰Si_pf values at the Guaymas Basin OMZ, an additional process must occur. We hypothesize that a phase enriched in ²⁸Si needs to dissolve in order to shift the pore fluid δ³⁰Si_pf values and this phase might be (1) iron (oxyhydr)oxides with adsorbed ²⁸Si or (2) terrestrial clays.

Silicon exhibits a strong affinity to iron (oxyhydr)oxides (see also Sect. 4.3.2; Davis et al., 2002) and Si isotopes fractionate significantly during Si adsorption and coprecipitation (Delstanche et al., 2009; Zheng et al., 2016). Dissolved Fe²⁺ in pore fluids can be transferred across the sediment–water interface via diffusion and re-precipitates as iron (oxyhydr)oxides, where it subsequently dissolves again in the reducing sediment. This process can repeat, resulting in multiple cycles of Fe dissolution and re-precipitation on the Guaymas Basin slope (Scholz et al., 2019). We hypothesize that the light ²⁸Si adsorbs to iron (oxyhydr)oxides in the water column and that, upon reductive dissolution of Fe minerals in the surface sediment, the light ²⁸Si isotope is re-released into the pore fluids, adding to the observed low δ³⁰Si_pf values. However, the quantification of this Fe–Si shuttle and the contribution to the low δ³⁰Si_pf values in OMZ pore fluids remains difficult given that Fe undergoes multiple cycles of dissolution and re-precipitation. Furthermore, the exact process of complexation, Si isotope fractionation, and coprecipitation is unknown and requires further investigations (see also Fig. S3). We can only speculate that the transport of ²⁸Si via the Fe–Si shuttle is only of minor importance given that the MAR of bSiO₂ and terrigenous material are dominating the Si supply to Guaymas OMZ sediments (Calvert, 1966; DeMaster, 1981).

The low δ³⁰Si_pf values can also be explained by dissolution of terrigenous clay particles, which are generally enriched in ²⁸Si, showing a large range in δ³⁰Si with the majority between −3 ‰ to 0 ‰ (Frings et al., 2016 and references therein). Primary minerals like feldspars or olivine, which are generally considered to control marine weathering reactions (e.g., Wallmann et al., 2008) have higher δ³⁰Si values (e.g., feldspars with −0.15 ‰; Georg et al., 2009), and their dissolution alone cannot create the low pore fluid δ³⁰Si_pf values (see Supplement and Fig. S6). Clays are usually considered to be the stable end product of silicate weathering. However, fine clay particles and highly reactive surface sites of clays, such as montmorillonite, smectite, and illite, may dissolve in natural waters (Cappelli et al., 2018; Golubev et al., 2006; Köhler et al., 2005). The dissolution is promoted by organic ligands and the reduction of structural iron of clay minerals under reducing conditions (Anderson and Raiswell, 2004). Humic substances that are abundant in OMZ sediments enriched in organic matter may catalyze the microbial reduction of structural iron in clays (Lovley et al., 1998) and their dissolution (Liu et al., 2017). Clays are abundant in OMZ sediments, given that fine-grained terrigenous material is transported downslope from the shelf to the basin (Scholz et al., 2019). Furthermore, the microbial oxidation of ferrous Fe in these fine-grained silicate minerals and its subsequent conversion to reactive iron minerals was also found to contribute to the Fe cycling on the Guaymas Basin slope (Scholz et al., 2019).

In order to constrain the possibility of terrigenous clay dissolution and the related shift to low δ³⁰Si_pf values in the Guaymas OMZ, a reactive transport model was applied, based on our previously published δ³⁰Si model (Ehlert et al., 2016). The data obtained at the Guaymas OMZ were used to model the turnover of Si in these sediments, and the previously published model was extended to consider additional processes (Fig. 6). A full description of the model is presented in the Supplement. The model was fit to dissolved Si concentrations and δ³⁰Si values measured in pore fluids and biogenic opal and K∕Al ratios determined in the solid phase (Fig. 6). High rates of terrigenous clay dissolution were applied at the sediment surface to reproduce the observed minima in δ³⁰Si_pf pore fluid values and K∕Al ratios in a model run best fitting our data set (Fig. 6, best fit). Since the terrigenous phases deposited at the sediment surface contain potassium (K_terr=1.7 wt %, $\mathrm{K}/\mathrm{Al}=\mathrm{0.19}$; Viers et al., 2009) and are depleted in ³⁰Si (clay δ³⁰Si (late stage weathering product) = −2 ‰; Opfergelt et al., 2010), the dissolution of these phases induces a decline in pore fluid δ³⁰Si and solid-phase K∕Al (Supplement). The precipitation of authigenic phases that are depleted in ³⁰Si (Si isotope fractionation: ${\mathrm{\Delta }}^{\mathrm{30}}{\mathrm{Si}}_{\mathrm{au}}=-\mathrm{2}$ ‰; Ehlert et al., 2016) and characterized by high K contents ($\mathrm{K}/\mathrm{Al}=\mathrm{0.32}$; Michalopoulos and Aller, 2004) induces a down-core increase in pore fluid δ³⁰Si_pf and solid-phase K∕Al below the surface minimum. Consequently, terrigenous clay dissolution under the reducing conditions of the OMZ and subsequent authigenic aluminosilicate precipitation can explain the low δ³⁰Si_pf values detected in Guaymas OMZ pore fluids (Fig. 5). However, our model results should be regarded with caution because we applied strongly simplifying assumptions (e.g., steady state, simple rate laws). Moreover, our estimates of solid-phase reactivity and isotopic composition are preliminary and not supported by independent data.

Figure 6 Data and model results for OMZ core: (a) porosity, (b) biogenic opal concentration in solid phase, (c) K∕Al ratio in solid phase, (d) dissolved Si concentration in pore fluids, (e) dissolved potassium in pore fluids, (f) isotopic composition of dissolved Si, (g) rate of biogenic opal dissolution, (h) rate of authigenic phase precipitation, and (i) rate of terrigenous-phase dissolution. Sensitivity tests of the model are indicated by setting the dissolution of terrigenous clay phases to zero (dashed red line) and setting the precipitation of authigenic phases to zero (dashed black line). The model values agree best with the measured data when both processes take place (blue line, best fit).

Additional simulations were conducted to investigate how the solid phase and pore fluid composition are affected by the dissolution of terrigenous clay phases and the precipitation of authigenic phases (Fig. 6) and how varying δ³⁰Si values of the dissolving terrigenous phase and Si isotope fractionation factors impact pore fluid δ³⁰Si_pf values (Fig. S6). Furthermore, varying K∕Al ratios and bSiO₂ solubilities are tested (see Supplement). The surface minima in pore fluid δ³⁰Si_pf and solid-phase K∕Al disappear when the dissolution rate is set to zero (R_terr=0), while the ongoing precipitation of authigenic phases leads to a strong down-core increase and high values at depth that are not consistent with the data. Pore fluid δ³⁰Si_pf and solid-phase K∕Al values strongly decrease with depth when the rate of authigenic-phase precipitation is set to zero (R_au=0), such that the model yields values that are significantly lower than the measured values. Dissolved Si concentrations cannot be used to further constrain R_terr and R_au because they are largely controlled by the dissolution of biogenic opal (R_opal). Dissolved K concentrations show a much lower sensitivity to R_terr and R_au than solid-phase K∕Al ratios due to the high porosity of the OMZ sediments. Changes in dissolved K are largely eliminated by molecular diffusion that is favored by the high porosity, while the effect of the solid-phase reactions R_terr and R_au on the pore fluid composition is diminished by the low solid-phase contents and the high background concentration of dissolved K in ambient bottom waters. However, the model runs show that the more sensitive pore fluid δ³⁰Si_pf and solid-phase K∕Al can be used to constrain the balance between the dissolution of terrigenous phases and the precipitation of authigenic phases and that both reactions are required to model the low δ³⁰Si_pf values measured in the Guaymas OMZ. Additionally, δ³⁰Si values for the dissolving terrigenous material higher than −2 ‰ cannot reproduce the measured δ³⁰Si values in the OMZ pore fluids. Only if the fractionation factor is lowered to −1 ‰, can terrigenous material with δ³⁰Si values of −1.7 ‰ produce the observed values (Fig. S6). Thus, low pore fluid δ³⁰Si values in the Guaymas OMZ can only be reproduced by dissolving terrigenous clay particles highly depleted in ³⁰Si.

The modeled Si isotope composition of the benthic flux is −0.97 ‰, which is lower than the δ³⁰Si value of the bottom water (+0.8 ‰). The higher bottom water δ³⁰Si value, along with the low Si concentration (∼30 µM), which is lower than the ambient water column Si concentration (∼80 µM), indicates that a certain amount of Si must directly re-precipitate at the sediment–water interface. Regardless, the δ³⁰Si of the bottom water is lower compared to the ambient water column, showing a benthic Si flux with low δ³⁰Si values at continental margin settings, which is also in excellent agreement with previously modeled and calculated δ³⁰Si values (Ehlert et al., 2016; Grasse et al., 2016). These findings show that benthic Si fluxes at continental margins are a source of low δ³⁰Si values to the ocean and need to be taken into account in future marine Si budget models.

4.4 Controlling processes and the impact on the global marine Si cycle

Stable and radioactive Si isotope data revealed significant sedimentary import and export processes influencing the marine Si cycle (Ehlert et al., 2013, 2016; Tréguer and De La Rocha, 2013; Grasse et al., 2016; Rahman et al., 2017; Sutton et al., 2018). Diatom burial removes about 9.9 Tmol yr⁻¹ Si from the ocean to the sediments; however, effects of terrigenous silicate dissolution and reverse silicate weathering on bSiO₂ burial, preservation, and the benthic Si flux (and its Si isotope composition) are not well constrained (Sutton et al., 2018). It has previously been shown that silicate minerals dissolve in deep methanogenic sediments where the dissolution process is favored by high CO₂ and organic ligand concentrations in ambient pore fluids (Wallmann et al., 2008). Similar to chemical weathering on land, the dissolution of terrigenous silicate phases in marine sediments leads to a release of cations and the conversion of CO₂ into ${\mathrm{HCO}}_{\mathrm{3}}^{-}$. Moreover, this marine weathering process provides the dissolved Al that is needed for reverse weathering reactions. Our OMZ data show for the first time that marine silicate weathering (dissolution of terrigenous silicates) also occurs in OMZ surface sediments, where it can outpace reverse weathering (precipitation of authigenic silicates). Our study indicates that ambient environmental conditions appear to significantly influence the balance between marine weathering and reverse weathering and thereby the Si flux back to the ocean. Pore fluid δ³⁰Si_pf values depend on a complex interplay between bSiO₂, terrigenous silicate dissolution, and authigenic aluminosilicate precipitation; however, the controlling factors that determine which process dominates are difficult to constrain (Fig. 5). In view of the OMZ settings (Guaymas Basin versus Peruvian margin), the most pronounced difference is the MAR_terr, which is significantly higher in the Guaymas Basin (252 g m⁻² yr⁻¹; calculated by multiplying the terrigenous content derived in Eq. 1 with the MAR from Eq. 2) than at the Peruvian margin (100 g m⁻² yr⁻¹; MAR from Ehlert et al., 2016; terrigenous content calculated following Eq. 1, with 6 wt % bSiO₂, 15 wt % OMC, and 8 wt % CaCO₃) (Fig. 5a, b). The high terrigenous detritus content is supplied via rivers in the Guaymas Basin (Calvert, 1966; DeMaster, 1981). In combination with the high MAR_terr in the Guaymas OMZ, high water ∕ rock ratios (high porosity) additionally promote dissolution processes (Fig. 5). Lower MAR_terr and water ∕ rock ratios found in the Peruvian upwelling margin appear to limit the dissolution rate of terrigenous phases and promote authigenic aluminosilicate precipitation (Fig. 5b, c), shifting pore fluid δ³⁰Si_pf to higher values compared to the corresponding ${\mathit{\delta }}^{\mathrm{30}}{\mathrm{Si}}_{{\mathrm{bSiO}}_{\mathrm{2}}}$ and δ³⁰Si_BW values. This illustrates that the pore fluid δ³⁰Si_pf values of apparently similar settings (e.g., OMZ sites) highly depend on the ambient environmental conditions and are not easily transferable.

4.5 Hydrothermal impact on the marine Si cycle

Findings of this study show that additional Si sources like hydrothermal input appear to affect the oceanic δ³⁰Si values only in close vicinity of the hydrothermal fields. The δ³⁰Si values of the hydrothermal plume (+0.7 ‰ to +1.4 ‰) are highly diluted by seawater (≥94 %, Table 2) and thus deviate from hydrothermal fluid δ³⁰Si values (−0.3 ‰; De La Rocha et al., 2000). However, the currently available data set regarding δ³⁰Si values of hydrothermal fluids is limited (two data points; De La Rocha et al., 2000), even though they are in excellent agreement with oceanic crust δ³⁰Si values ($-\mathrm{0.29}\pm \mathrm{0.08}$ ‰; Savage et al., 2010), the rock through which hydrothermal fluids circulate and gain their Si isotopic signature. In our data set, no correlation exists between the δ³⁰Si values and the Si concentration of the hydrothermal plume (Fig. S2), and instead the δ³⁰Si values are predominantly controlled by Si precipitation, likely in the hydrothermal conduit during ascent or after discharge in contact with colder seawater. Temperature variations and interlinked precipitation rates were found, in addition to coprecipitation with Al or Fe, to cause large Si fractionation, such that precipitates are enriched in ²⁸Si (Geilert et al., 2014, 2015; Oelze et al., 2015; Roerdink et al., 2015; Zheng et al., 2016). The varying impacts of these factors can also explain why the diluted hydrothermal plume δ³⁰Si values with the highest hydrothermal share (Table 2) do not show the lowest δ³⁰Si values, indicative of hydrothermal fluids, given that Si is more reactive compared to Mg, the element on which the hydrothermal share calculations are based (see Supplement to Berndt et al., 2016). The large range in hydrothermal plume δ³⁰Si values, which clearly show high degrees of seawater dilution, illustrates the complexity of precipitation processes when hydrothermal fluids get in contact with cold seawater and that require further investigation, especially with respect to the impact on the global marine Si cycle.

5 Conclusions

Marine silicate weathering and reverse weathering impact the pore fluid isotopic composition of sediments and are key processes of the marine Si cycle. In the Guaymas Basin, these processes have been studied under markedly differing thermal and redox conditions. Si isotope compositions of pore fluids combined with those of biogenic silica and ambient bottom waters helped to decipher marine weathering and reverse weathering reactions, which would have remained undetected by elemental concentrations alone and highlight the importance of Si isotope studies for constraining early diagenetic reactions. Si concentrations and δ³⁰Si_pf signatures are the result of the interplay between silica dissolution and Si precipitation; however, the involved phases differ significantly between the study sites. Large differences in δ³⁰Si_pf values in a regionally constrained basin show that oxic and anoxic conditions, hydrothermal fluids, water ∕ rock ratios and the input of terrigenous material strongly affect the pathways and turnover rates of Si in marine sediments. The light δ³⁰Si_pf and δ³⁰Si_BW values from the Guaymas OMZ confirm earlier studies suggesting a light Si isotope value of the benthic Si flux (Ehlert et al., 2016; Grasse et al., 2016), which need to be taken into account in future oceanic mass balances of Si and in modeling studies concerning the isotopic Si cycle. Environmental settings, in particular the MARs of terrigenous material, water ∕ rock ratios, and redox conditions, appear to be the major factors controlling the balance between marine silicate weathering and reverse weathering and the Si isotope fractionation in pore fluids of marine sedimentary settings and need to be considered in marine Si isotope studies in particular.