Chemical Geology 360–361 (2013) 32–40
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Iron isotope geochemistry of biogenic magnetite-bearing sediments from the Bay of Vidy, Lake Geneva Elizabeth M. Percak-Dennett a,b, Jean-Luc Loizeau c, Brian L. Beard a,b, Clark M. Johnson a,b, Eric E. Roden a,b,⁎ a b c
Department of Geoscience, University of Wisconsin-Madison, 1215 West Dayton Street, Madison, WI 53706, USA NASA Astrobiology Institute, University of Wisconsin-Madison, Madison, WI 53706, USA Institut F.-A. Forel et Centre d'Études en Sciences Naturelles de l'Environnement, Université de Genève, 10 route de Suisse, CH-1290 Versoix, Switzerland
a r t i c l e
i n f o
Article history: Received 16 May 2013 Received in revised form 4 October 2013 Accepted 5 October 2013 Available online 15 October 2013 Editor: Michael E. Böttcher Keywords: Bay of Vidy Sediment Microbial iron reduction Iron isotope geochemistry Magnetite formation
a b s t r a c t Dissimilatory microbial iron oxide reduction (DIR) has been hypothesized to be an important respiratory pathway on early Earth, potentially generating significant quantities of Fe(II) that have been preserved in Proterozoic and Archean sedimentary rocks. In particular, DIR has been implicated in the formation of magnetite in Precambrian marine sediments. To date, however, only one modern sedimentary environment exists where in situ magnetite formation has been linked to DIR: the Bay of Vidy in Lake Geneva, Switzerland. Previous work at this locality has characterized a magnetic susceptibility anomaly that reflects the presence of fine-grained magnetite produced via microbial reduction of amorphous Fe(III) oxides that enter the Bay of Vidy from a nearby sewage treatment plant. In this study, we report on the Fe isotope composition of aqueous and solid-phase Fe in the Bay of Vidy sediments. Extensive Fe(III) reduction has occurred, resulting in the conversion of nearly all reactive (non-silicate) Fe(III) to a variety of Fe(II)-bearing phases, with mixed Fe valence magnetite being a minor but easily detectable component (0.5–8 wt.%). Very little Fe isotope variation was observed in any solid phase Fe components, including magnetite, although significant fractionation was observed between aqueous and solid-phase Fe(II). Because Fe mass-balance was dominated by the solid phase, little net change in δ56Fe values for Fe(II)-bearing components was produced despite clear evidence for DIR. This study provides a basis for interpreting instances in the rock record where DIR was the driving force for Fe(II) production and magnetite formation, yet no significant deviations in δ56Fe values were preserved. A key implication of the results is that Fe isotope homogeneity is not sufficient to rule out a biological mechanism for magnetite formation, and this should be taken into account when examining the Precambrian rock record. © 2013 Elsevier B.V. All rights reserved.
1. Introduction Dissimilatory iron reduction (DIR) is a microbial metabolism that couples the enzymatic reduction of Fe(III) oxides to oxidation of H2 or organic compounds (Lovley et al., 2004). This metabolic pathway is widespread in modern marine and freshwater sedimentary environments and plays a significant role in redox cycling of Fe (i.e., Lovley, 1991). The ferrimagnetic mineral magnetite (Fe3O4) is a common end product of microbial reduction of poorly crystalline Fe(III) oxides in laboratory studies, and can be produced by a wide range of microorganisms (Lovley and Phillips, 1987; Roden and Lovley, 1993; Fredrickson et al., 1998; Zhang et al., 1998; Zachara et al., 2002; Roh et al., 2006; Zavarzina et al., 2006; Borch et al., 2007; Perez-Gonzalez et al., 2010; Salas et al., 2010). Some thermophilic dissimilatory iron-
⁎ Corresponding author at: 1215 West Dayton St., University of Wisconsin-Madison, Madison, WI 53706, USA. Tel.: +1 608 831 3680. E-mail address:
[email protected] (E.E. Roden). 0009-2541/$ – see front matter © 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.chemgeo.2013.10.008
reducing microorganisms are deeply branching based on 16S rRNA molecular phylogeny (Vargas et al., 1998), indicating that microbial DIR could be an ancient metabolic pathway. Fe-bearing minerals, such as magnetite, in Archean and Proterozoic sedimentary rocks are hypothesized to preserve evidence for microbial redox cycling of Fe (Cloud, 1974; Klein, 1974; Walker, 1984; Nealson and Myers, 1990; Vargas et al., 1998; Konhauser et al., 2005; Yamaguchi et al., 2005). Numerous studies have examined the Fe isotope composition of magnetite and other Fe(II) minerals in Archean and Proterozoic banded iron formations as a means to understand Fe redox cycling and atmospheric development (Dauphas et al., 2007; Johnson et al., 2008; Craddock and Dauphas, 2011; Czaja et al., 2013). Microbial reduction of poorly crystalline Fe(III) oxides results in the production of up to millimolar levels of low-δ56Fe aqueous Fe(II) (Johnson et al., 2005). Incorporation of this aqueous Fe(II) into magnetite and other Fe(II)bearing minerals has been identified as a potential mechanism for the origin of low-δ56Fe signatures in Archean and Proterozoic Fe-rich rocks (Johnson et al., 2008; Tangalos et al., 2010). Mobilization and mass transport of low-δ56Fe aqueous Fe(II) can also result in changes to δ56Fe values on macroscopic scales and sediments as residual Fe(III) becomes
E.M. Percak-Dennett et al. / Chemical Geology 360–361 (2013) 32–40
increasingly isotopically heavy relative to starting materials (Severmann et al., 2006, 2008; Tangalos et al., 2010). Although poorly crystalline Fe(III) oxides are ubiquitous in modern sedimentary environments, diagenetic formation of magnetite through reduction of Fe(III) oxides is rare in recent marine sediments, due to fundamental changes in ambient geochemistry (e.g., oxygen, sulfate) relative to those in the Archean and Paleoproterozoic. Formation of significant quantities of magnetite associated with DIR has only been reported for one modern location on Earth: the Bay of Vidy in Lake Geneva, Switzerland (Gibbs-Eggar et al., 1999). The Bay of Vidy is a unique environment where a mixture of amorphous Fe(III) oxides, heavy metals, and high levels of nutrients and organic carbon is delivered to the bottom of Lake Geneva through the outflow of a municipal sewage treatment plant. Unlike modern marine environments, the Bay of Vidy is a low-S environment that provides an important analogy to Archean marine environments. Sewage de-phosphatization has utilized ferric chloride since 1971, resulting in elevated Fe concentrations in treated run off (Loizeau et al., 2004; Pardos et al., 2004; Wildi et al., 2004). Previous investigations in the Bay of Vidy have identified significant quantities of fine-grained magnetite, mineralogically and magnetically similar to that formed as a by-product of DIR in laboratory cultures, which demonstrate temporal correlations to the use of ferric chloride for wastewater treatment (Gibbs-Eggar et al., 1999). Additional work has determined the presence significant microbial activity (Thevenon et al., 2011), and identified DNA from Fe(III)reducing bacteria in Bay of Vidy sediments (Snoeyenbos-West et al., 2000; Glass-Haller, 2010; Haller et al., 2011). Due to the constant influx of sewage treatment run off, ongoing work has focused on contaminant transport (Goldscheider et al., 2007), and the environmental impacts of this discharge (Loizeau et al., 2004; Pardos et al., 2004; Pote et al., 2008). Garcia-Bravo (2011) studied aspects of aqueous and solid-phase Fe and S geochemistry of Bay of Vidy sediments in the context of mercury methylation. No additional work, however, has been conducted on characterizing the solid-phase Fe geochemistry or isotope composition in these sediments, and the potential use of these deposits as a proxy for organic-rich, low-S marine sediments on ancient Earth. In this paper we present geochemical and isotopic data obtained from magnetite-bearing Bay of Vidy sediments. Careful mineralogical analyses of the magnetic portions of the sediment were conducted, and a sequential extraction procedure was used to isolate several different solid-phase Fe pools for Fe isotope analysis. The results yield new insights into Fe isotope behavior in sediments that have undergone extensive microbial Fe(III) reduction accompanied by limited Fe mobilization, which has implications for understanding the Fe isotope variations recorded in Fe(II)-bearing minerals in Proterozoic and Archean sedimentary rocks.
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all material collection was done as quickly as possible (ca. 2 min per sampling interval), and subsamples were immediately placed in N2-flushed 50-mL plastic centrifuge tubes. The tubes were filled completely and immediately centrifuged at ca. 7000 ×g for 20 min. Pore fluids were recovered in a N2-purged glove bag, filtered through 0.2 μm nylon filters, and acidified to a final concentration of 0.5 M HCl. After pore fluid extraction, the sediment pellets were immediately frozen and remained frozen until analysis. In total, 17 samples were collected from core LPD3; we report here on four depth intervals, which were selected to encompass the range in magnetic susceptibility as a function of depth.
2.2. Sequential Fe extraction A sequential extraction method was employed to determine the relative concentrations of amorphous and crystalline Fe(II)- and Fe(III)-containing phases. All Fe concentrations were measured using Ferrozine (Stookey, 1970), with and without hydroxylamine to determine Fe(II) and total Fe concentrations, respectively. Fe(III) contents were determined from the difference of Fe(II) and total Fe. Fractions of Fe that had the potential to undergo oxidation during extraction (bulk Fe, HF and HNO3 digestions) were quantified only for total Fe. The frozen sediment was thawed from each depth interval in an anaerobic chamber (Coy Products, Grass Lake, MI, USA). Magnetic separates were removed using a small handheld rare-earth magnet from an anoxic sediment–water slurry. After removal of magnetic particles, sediment slurry from each depth interval was portioned into two aliquots for replicate extraction and analysis. The sediment was first subjected to a 1 h digestion with 0.5 M HCl, during which time acid-volatile sulfide (AVS) was trapped using zinc acetate and measured using Cline's reagent (Cline, 1969). This extraction was followed by a 23 h 0.5 M HCl digestion of the residual sediment. The 1 h and 23 h HCl extracts were analyzed for Fe, Ca, and P concentrations on an inductively coupled plasma optical emission spectrometer (TJA IRIS Advantage ICP-OES; University of Wisconsin Soil and Plant Analysis Laboratory, Madison, WI; accuracy = ±5%). A 16 h 29 M HF digestion was then used to liberate Fe-associated with silicates (Severmann et al., 2006), followed by a 16 h 14 M HNO3 digestion to remove any remaining Fe-phases, inferred to be pyrite (Huerta-Diaz and Morse, 1990). All values were normalized to units of μmol/g dry sediment. Total bulk Fe from the sediment was determined by taking a separate aliquot from the core prior to removal of magnetic separates, heating to 850 °C for 12 h followed by total digestion in 29 M HF and 14 M HNO3. Total elemental composition of the sediment was also measured by heating and acidification with HNO3 and HF, followed by ICP-OES analysis (University of Wisconsin Soil and Plant Analysis Laboratory, Madison, WI).
2. Materials and methods 2.1. Sample collection
2.3. Mineralogical analysis
Six cores, three sets of duplicates, were collected from the Bay of Vidy using a gravity corer in September 2010. Details of the core collection procedures are given in Loizeau et al. (1997). One core (LPD3) was selected as a representative sample and used for all analyses. Core LPD3 was collected from a site previously investigated due to its proximity to a sewage treatment plant outflow pipe (Goldscheider et al., 2007) (Fig. 1). The core was recovered from 18.7 m depth at GPS coordinates X = 534760, Y = 151711. Magnetic susceptibility was measured in 1-cm increments prior to sampling using a MS2C susceptibility meter (Bartington Instruments, Oxford, England). Subsampling was done by extruding the core and pooling 2.5 cm depth intervals, discarding the top 1–2 cm that became jumbled during core collection. In order to limit oxidation of core materials,
X-ray diffraction (XRD) analyses were carried out for bulk sediment and magnetic separates from each of the examined depth intervals. Diffraction data were acquired using a Rigaku Rapid II Xray diffraction system with a two-dimensional image plate detector (Mo Kα radiation). The two-dimensional images were integrated to produce conventional patterns using the Rigaku 2DP software. Anoxically dried magnetic separates underwent SEM and TEM analysis. SEM analyses were performed with a Hitachi S-3400 variable pressure SEM; TEM analyses were performed a FEI Titan Aberration Corrected scanning/transmission electron microscope. Secondary electron (SE) mode was used to examine particle morphology. Energy-dispersive Xray spectra (EDS) were used to ascertain elemental composition of the magnetic separates.
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2.4. Fe isotope analysis Iron isotope compositions were determined for the HCl, HF, and HNO3 extracts, as well as pore fluid aqueous Fe(II) and 10 M HCl digests of the magnetic separates. Previous studies with synthetic Fe(III) oxides (Skulan et al., 2002; Wiederhold et al., 2006) and Fe(II)/Fe(III) oxide mixtures (Wu et al., 2010, 2011) have shown that proton-promoted dissolution with HCl does not cause Fe isotope fractionation. To our knowledge the potential for Fe isotope fractionation during HF and HNO3 extraction has not been evaluated. HF extraction can cause minor dissolution of pyrite (Lord, 1982), and hence the possibility exists that the measured Fe isotope composition of the silicate and pyrite pools could have been altered slightly during the sequential extraction procedure. However, this extraction scheme has been used successfully in sediment Fe isotope studies where the Fe isotope composition of the pyrite pool could be well-rationalized in terms of sediment Fe–S geochemistry and environmental setting (Severmann et al., 2006, 2008). Moreover, given the near-complete uniformity of solid-phase Fe isotope compositions in Bay of Vidy sediments (see below), any Fe isotope exchange that may have occurred during partial phase dissolution would not change the conclusions of this study. All Fe aliquots were processed through anion-exchange chromatography (Beard et al., 2003), followed by analysis by MulticollectorInductively Coupled Plasma Mass Spectrometry (MC-ICP-MS; Micromass IsoProbe) to determine iron isotope compositions. Yields after column chemistry for all components were 95 ± 5%, as determined using Ferrozine and colorimetric analysis. Measured data are reported as the 56 Fe/54Fe ratio relative to the average of igneous rocks using standard δ notation (Table A.1), in units of per mil (‰): 56
δ Fe ¼
56
54 Fe= Fe
sample
56 54 = Fe= Fe
IgRx
–1 1000
Isotopic fractionation between two phases A and B is expressed using standard notation: 56
56
56
Δ FeA‐B ¼ δ FeA −δ FeB : External precision of δ56Fe values was ±0.08‰ (2-SD) based on replicate standard analysis. The measured Fe isotope composition of the IRMM-014 standard was δ56Fe = −0.07 ± 0.09 (2SD, n = 4), which lies in the long-term values of δ56Fe = −0.09‰ for the lab. Iron isotope compositions of additional in-house standards were: HPS I: δ56Fe = 0.48 ± 0.02 (2SD, n = 3), J-M Fe: δ56Fe = 0.33 ± 0.16 (2SD, n = 2). The accuracy of the Fe isotope analyses was evaluated by analysis of two test solutions made by combining 75 μg Fe of a known isotopic composition (HPS-1 standard) with acid flushed through anion exchange columns during sample purification. The acid contained the same concentrations of major elements as the initial samples, and this process was done to rule out potential interferences and matrix factors that could affect the accuracy of isotopic measurements. These samples were processed identically to all other samples, and the results were within error of the pure HPS-1 Fe solution, at δ56Fe = 0.51 ± 0.09 (2SD, n = 3). In total, 74 samples were analyzed, 11 sample duplicates were run (2SD = 0.06‰), as were nine standards, and two test solutions. Additional data on laboratory procedures can be found in Beard et al. (2003). 3. Results
(Gibbs-Eggar et al., 1999; Loizeau et al., 2003). The four depth intervals indicated by stars in Fig. 2 correspond to those analyzed for Fe speciation and isotope composition, including intervals corresponding to the three peaks in magnetic susceptibility, as well as an interval with lower susceptibility near the sediment surface (ca. 5 cm depth). 3.2. Sediment geochemistry and mineralogy The solid-phase geochemical data for core LPD3 are presented as the average of replicate samples (Table 1 and Fig. 3). Total bulk Fe varied from 322 to 687 μmol Fe/g dry sediment. The 1 h 0.5 M HCl and 29 M HF digestions comprised the greatest proportions of Fe (Fig. 3). Given that nearly all of the 0.5 M HCl extractable Fe was recovered as Fe(II) (100% in the 1 h extraction, and more than 95% in the 23 h extraction), these results indicate a system dominated by amorphous Fe(II) (see below) and silicate-associated Fe(II). Acid volatile sulfide (AVS) ranged from 22 to 81 μmol S/g dry sediment, comprising 8–19 wt.% of total Fe, whereas non-AVS Fe(II) comprises 17–76 wt.% of total iron. Pyrite (recovered in the HNO3 digestion) and magnetic separates were minor contributors to the overall Fe budget. The abundance of magnetic separates (8% of total Fe) peaked at 36–39.5cm depth, corresponding to the depth of greatest magnetic susceptibility. ICP-OES analysis of bulk sediment digestions and 0.5 M HCl extracts (combined values for the 1 and 23 h extractions) showed very high Ca concentrations (Table 2), most likely due to significant calcite abundance in core materials (see XRD data below). Phosphorous increased as a function of depth, where HCl-extractable P comprised between 14 and 45% of total P. Fe/P ratios were calculated using data from the HCl digestions as well as overall bulk data. Significantly elevated Fe/P ratios were observed at the shallowest depth while ratios remained relatively staple in deeper depths (Table 2). The dominant mineral phases in bulk sediment were quartz, calcite, and plagioclase in decreasing abundance as determined by XRD analysis. XRD analysis of the magnetic separates revealed the presence of magnetite, maghemite, and hematite (Fig. 4). SEM investigation demonstrated that typical magnetic separates exhibited a rough surface texture and irregular morphology (Figs. A.1 and A.2). This irregular surface was due to small-scale particle aggregation. These particle-like features ranged from 50 to 200 nm in diameter, compared to total aggregate sizes between 0.5 and 60 μm (with most in the 1–2 μm range). EDS showed that the elemental composition of these particles was predominantly Fe and O, although many regions contained a mixture of Ca, P, and O, with very little Fe. Selected-area electron diffraction (SAED) indicated the presence of both magnetite and hematite in the separates, consistent with XRD results and possible oxidation of magnetite. 3.3. Fe isotope compositions The average δ56Fe value of bulk sediment was 0.09 ± 0.05‰ (1-SD, n = 7). A narrow range of δ56Fe values was observed among the various sequential extracts and magnetic separates (Fig. 5; Table 3). With the exception of the pore fluids, most δ56Fe values were within 1-SD of the near-zero bulk composition, including the magnetic separates, indicating very little solid-phase Fe isotope fractionation. δ56Fe values for aqueous Fe(II) averaged −1.07‰, and varied from −1.81‰ to − 0.61‰. Iron recovered in the HNO3 digestion, interpreted to be associated with pyrite, showed the highest variation among sediment solid phases, with a range in δ56Fe values from −0.73 to 0.38‰. No direct relation between δ56Fe values and magnetic susceptibility was evident.
3.1. Magnetic susceptibility 3.4. Isotopic composition of iron-sulfides The magnetic susceptibility measurements revealed peaks at approximately 10, 40, and 60 cm depth (Fig. 2), similar to those observed previously for cores collected in the vicinity of LPD3
Sequential HCl extractions allowed for quantification of amorphous Fe, however physical separation and measurement of Fe(II) and Fe(III)
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Fig. 1. Sampling location in the Bay of Vidy, Switzerland.
components was not possible. Likewise, it was not possible to directly determine δ56Fe values for acid-volatile FeS phases, as these co-dissolve with non-sulfide-associated amorphous Fe compounds during HCl extraction. These values can, however, be calculated using a massbalance approach, which allowed us to solve for δ56Fe values of different HCl-extractable Fe phases (e.g., Severmann et al., 2006; Percak-Dennett et al., 2011) (see Supplementary Text A.1). δ56Fe values calculated for FeS (Fig. 5, Table A.2) showed a range of 0.29 to −0.60‰ (1-SD = 0.12 to 0.30‰). Given this range, and corresponding consideration of errors, the FeS component is not significantly different in isotopic composition from bulk Fe. Similarly, the pyrite fraction showed no significant isotopic differences from bulk Fe, and a similar range in δ56Fe values (−0.73 to 0.38‰). The overall very low abundance of pyrite-Fe (b6% of total Fe), together with the geological history of this region and low sulfide concentrations, indicates that the pyrite may be detrital in origin. Thus, the observed heterogeneities in δ56FePyrite-Fe values may be inherited from source pyrite, rather than a reflection of Fe isotope fractionation during authigenic pyrite formation. These nuances are outside the main focus of the paper, and we therefore do not consider this component further in the discussion below.
4. Discussion 4.1. Sediment Fe geochemistry The high concentrations of Fe(II) recovered in 0.5 M HCl extractions indicate that DIR is active in Bay of Vidy sediments (Lovley and Phillips, 1987), consistent with previous molecular biological work at this site sediments (Snoeyenbos-West et al., 2000; Glass-Haller, 2010; Haller et al., 2011). AVS and pyrite-associated Fe values were small compared to total 0.5 M HCl-extractable Fe(II), comprising an average of 12% and 3% of total Fe, respectively. These results suggest that Fe(III) reduction dominates over bacterial sulfate reduction with respect to Fe diagenesis, as observed in other S-limited freshwater sediments (Suess, 1979; Wallmann et al., 1993; Roden and Wetzel, 1996; Thomsen et al., 2004). Bay of Vidy sediments contain high concentrations of P, which is not surprising given that these sediments are impacted by effluent from a sewage treatment plant (Dorioz et al., 1998). Phosphate immobilization in Fe(II)-rich non-sulfidic sediments typically involves precipitation of mineral phases such as vivianite (Fe3(PO4)2), or sorption of PO3− to 4 ferrous hydroxides (Einsele, 1936; Patrick and Khalid, 1974; Suess,
Table 1 Results of sequential Fe extractions. All values are averages of duplicate extractions, in units of μmol/g dry sediment. See text for a discussion of the extraction procedure for each component. Depth interval
Aqueous Fe (μmol/L)
Acid volatile sulfide-associated Fe
1 h 0.5 M HCl extractable Fe(II) (non AVS)a
23 h 0.5 M HCl extractable Fe(II)
23 h 0.5 M HCl extractable: Fe(III)
HF extractable Fe silicates
HNO3 extractable pyrite
Magnetic separates
Total Feb
Bulk Fec
1–4.5 cm 8–11.5 cm 36– 39.5 cm 57– 60.5 cm
0.026 0.062 0.023
30.57 66.87 22.37
65.06 121.08 157.48
54.6 49.2 20.6
2.5 2.6 1.5
247.9 149.6 63.4
21.9 18.1 1.8
1.6 4.4 21.8
393.7 345.0 266.6
475.5 573.7 322.0
0.013
81.30
441.71
44.0
2.0
86.7
2.9
5.5
582.8
686.8
a b c
1 h HCl extracted Fe(II) minus AVS. Total Fe is the sum of all Fe pools presented except for bulk Fe. Bulk Fe was measured as total Fe recovered in core sample using a bulk HF/HNO3 extraction.
E.M. Percak-Dennett et al. / Chemical Geology 360–361 (2013) 32–40
0 3 6 9 12 15 18 21 24 27 30 33 36 39 42 45 48 51 54 57 60 63
0 3 7 10
Depth (cm)
Depth (cm)
36
14
AVS
17
1 hour non-AVS HCl: Fe(II)
21
23 hour HCl: Fe(II)
24
23 hour HCl: Fe(III)
28
HF digestion: Fe(tot)
31
HNO3 digestion: Fe(tot)
35
Magnetic Seperates Fe(tot)
38 42 45 49 52 56 59 63
0
100
200
0
300
200
400
600
800
µmol Fe/g dry sediment
Magnetic Susceptibility Fig. 2. Magnetic susceptibility in core LPD3. Stars indicate samples analyzed for Fe geochemistry and isotope composition.
1979; Roden and Edmonds, 1997). In light of the relatively low Fe/P ratios in the 0.5M HCl extracts compared to bulk values (Table 2), it is likely that Fe(II)-P phases comprise a significant portion of the reactive Fe in Bay of Vidy sediments. The size and morphology of magnetic particles in Bay of Vidy sediments indicate that they are authigenic rather than detrital in origin. The magnetic separates from all depths had a very rough, irregular surface texture indicating aggregation of very small (nm-scale) particles, and sub-μm grain size (Figs. A.1 and A.2). Detrital magnetite, in contrast, has rounded grain boundaries and is typically larger than 200 μm in diameter (Grimley et al., 2004; Horneman et al., 2004). These results confirm the previous findings of Gibbs-Eggar et al. (1999), who documented the presence of superparamagnetic (SP) magnetite, with an estimated grain size of 20–100 nm in Bay of Vidy sediments. The hummocky surface texture of magnetite from this sediment is similar in morphology to previously described instances of framboidal magnetite (Aldana et al., 1999). Although the mechanism behind framboidal magnetite formation is unexplored in natural systems, these “framboids” have been inferred to be authigenic precipitates. Previous investigations of magnetic separates from this location described SP magnetite (Gibbs-Eggar et al., 1999), thus limiting the possible modes of formation. Our data support the claim of a biological origin for this material. In addition to microbial formation from DIR, SP magnetite has been detected in environmental samples, and has been ascribed due to pedogenesis, combustion of precursor Fe minerals, bacterial magnetosomes, as well the less common processes of fuel combustion (Dearing, 1994). We can rule out pedogenesis due to the
Fig. 3. Pools of Fe recovered by sequential extraction of sediment from different depths in core LPD3. Non-AVS Fe(II) recovered by 1 h 0.5 M HCl extraction represents total Fe(II) in the extract minus the amount of reduced S recovered as AVS. Aqueous Fe values were too low (b0.1 umol Fe/g dry sediment) to be visible on this scale.
location of magnetite as part of the sediment core sequence. Burning of other iron oxides such as goethite under reducing conditions would create SP magnetite, yet no goethite was seen in XRD or SEM/TEM analysis of dried sediment. Similarly, morphological examination rules out fly ash deposits and bacterial magnetosomes. The environmental setting and spatial distribution of magnetic susceptibility measurements throughout the Bay of Vidy thus rule out these alternative mechanisms for magnetite formation. This leaves magnetite created through DIR as the only viable explanation for the presence of SP magnetite. Microscopic analysis by SEM, TEM, and XRD indicate that the magnetic separates were dominated by magnetite with some hematite; however, Ca and P were both also present. XRD spectra revealed no hydroxyapatite (Ca5(PO4)3(OH)) or other Ca or P-based minerals. Previous work by Czaja et al. (2012) has demonstrated sensitivity between 0.1 and 0.5 wt.% for iron-bearing phases for the XRD methods used here, and we infer the sensitivity to hydroxyapatite and phosphate phases to be similar. Laboratory studies have found Ca-phosphate minerals will nucleate around a magnetite seed, although pH is the main driver of these reactions, not the presence of magnetite (Karapinar et al., 2004). Given the elevated levels of Ca and P recovered in the 0.5 M HCl extractions, as well as Ca and P signals in EDS spectra, Ca and P must be intimately associated with the magnetic separates, yet the exact nature of this relation cannot be determined through microscopic examination. The association could reflect sorption to particle surfaces, or alternatively Ca and P could have adsorbed to
Table 2 Ca, Fe, and P inventories for core LPD3. Values are normalized to μmol/g of dry sediment and represent averages of analysis of duplicate subsamples. 1 h + 23 h HCl Extraction (μmol/g sediment)a
Sediment totals (μmol/g sediment)b
% Released in HCl digestion
Fe/P ratio
Depth (cm)
Ca
P
Fe
Ca
P
Fe
P
Fe
HCl extraction
Bulk sediment
1–4.5 8–11.5 36–39.5 57–60.5
322 259 163 655
22 83 72 193
42 22 60 160
1545 1633 1720 1501
49 448 529 569
424 412 289 664
44.6 18.4 13.6 34.0
9.9 5.2 20.9 24.1
1.91 0.26 0.84 0.83
8.63 0.92 0.55 1.17
a b
Summed data from ICP-OES analysis of 1 h and 23 h 0.5 M HCl digestions. Ca and P values determined by ICP-OES analysis of bulk sediment, Fe determined by HF and HNO3 extraction and colorimetric measurement.
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Table 3 Fe isotope data from sequential extractions. Depth interval Aqueous Fe
1–4.5 cm 8–11.5 cm 36–39.5 cm 57–60.5 cm a b c
1 h 0.5 M HCl extractable Fe
δ56Fe
SDa
nb %c
−1.16 −0.71 −1.81 −0.61
0.02 0.02 0.02 0.03
1 1 2 1
δ56Fe
SD
23 h 0.5 M HCl extractable Fe n %
δ56Fe
SD
n %
HF extractable Fe silicates
HNO3 extractable pyrite
Magnetic separates
δ56Fe
δ56Fe
δ56Fe SD
n %
δ56Fe SD
n
0.17 0.15 0.05 0.09
2 2 2 2
0.08 0.11 0.04 0.15
2 3 3 2
SD
0.01 −0.05 0.06 5 22.5 −0.11 0.08 2 13.5 0.05 0.02 0.01 0.27 0.01 2 45.6 0.26 0.07 2 12.6 0.11 0.06 0.01 0.06 0.03 2 62.2 0.15 0.01 3 7.7 −0.01 0.02 0.00 0.15 0.03 2 78.8 0.25 0.01 2 6.9 0.06 0.02
n % 1 2 2 3
SD
n %
58.4 −0.17 0.19 2 5.2 36.3 −0.73 0.11 2 4.4 21.9 −0.22 0.05 2 0.6 13.1 0.38 0.09 2 0.4
0.02 0.04 0.01 0.04
HF/HNO3 extractable bulk Fe
0.4 1.1 7.5 0.8
0.04 0.03 0.02 0.03
1 SD of all replicates and duplicates, if n = 1, then this is 2-SE. Note that all δ56Fe values are relative to igneous rocks. Number of replicates. Overall percent of total Fe for component.
Fe(III) oxide phases prior to microbial Fe(III) reduction. Subsequent formation of magnetite through Fe(III) reduction would then preserve residual Ca and P as a separate amorphous phase, rather than direct incorporation into the magnetite crystal lattice. Either of these mechanisms could produce a Ca- and P-rich magnetic solid, as observed in our cores, although neither has any bearing on the biogenicity of the magnetite. 4.2. Constraints on Fe isotope geochemistry There have been no prior measurements of the Fe isotope composition of DIR-produced magnetite in modern sediments, and the Bay of Vidy is unique as the only known environmental source for such magnetite. Magnetic separates, however, were a minor component of the overall sediment Fe budget, and therefore the discussion of Fe isotope geochemistry will center on the system as a whole. The Fe isotope composition of aqueous Fe(II) could be controlled by either isotopic exchange with Fe phases and/or sorption of aqueous Fe(II) to sediment solids (see Wu et al., 2011, 2012b and references cited therein). In natural freshwater systems where partial microbial reduction of naturally occurring Fe(III) oxides has occurred, fractionation between Fe(II) aqueous and Fe(III) oxides are ~−2.0‰ (Tangalos et al., 2010), and marine environments dominated by DIR have similar fractionations of −1.1 to −3‰ (Severmann et al., 2006). These natural systems are distinct from the Bay of Vidy, because near complete Fe(III) reduction has occurred in Bay of Vidy sediment, thus eliminating Fe(III) as a participant in redox-driven equilibrium fractionation. Instead the ca. −1‰ values for aqueous Fe(II) are likely the result of sorption of a relatively small pool of aqueous Fe(II) to the large pool of solid phase materials in the sediment, as prior experiments have indicated a −0.5 to −1.2‰ fraction between aqueous and sorbed Fe(II) (Wu et al., 2011). Alternatively, equilibrium fractionation associated with Fe atom exchange between aqueous Fe(II) and other sediment Fe(II) phases could account for the negative aqueous Fe(II) δ56Fe values. Fractionation factors on the order of −0.5‰ have recently been determined for equilibrium Fe isotope fractionation factor between aqueous Fe(II) and
Fig. 4. XRD spectra for magnetic separates obtained from different depths of core LPD3. Peaks are labeled based on corresponding mineralogy: Mt (magnetite) or H (hematite).
iron monosulfide (FeS) phases (Guilbaud et al., 2011; Wu et al., 2012a). However, such phases accounted for less than 20% of total sediment Fe (see above), and it therefore seems unlikely that exchange with these phases was primarily responsible for controlling aqueous Fe(II) δ56Fe values. Equilibrium fractionation between aqueous Fe(II) and the large quantity of unidentified, amophous solid-phase Fe(II) could have been responsible, although fractionation factors for such exchange are at present unknown. In any case, the relatively small size of the aqueous Fe(II) pool compared to other sediment Fe(II) phases means that aqueous/solid-phase Fe(II) exchange could not have significantly influenced solid-phase Fe(II) δ56Fe values. Near-complete microbial reduction of Fe(III) oxides resulted in no significant variations to solid-phase δ56Fe values. This has been observed in experimental work where microbial reduction of ferric hydroxides resulted in magnetite formation (Johnson et al., 2005), and seems a likely explanation for the constant, near-zero δ56Fe values for all Fe(II) components, including magnetite, in Bay of Vidy sediments (Fig. 5), and consistent fractionation between aqueous Fe(II) and solidphase Fe(II) and Fe(III) (Table 4). 4.3. Implications to Precambrian Fe cycling In order for Fe isotope deviations to be preserved in the rock record, mass transport (i.e. physical redistribution or diffusion under suboxic
Fig. 5. Fe isotope composition for extracted Fe pools. δ56Fe values for FeS were not directly measured, but extrapolated as described in the supplementary text. Magnetic susceptibility is plotted as a black line on the lower X-axis, and δ56Fe data are plotted on the upper X-axis. Note the scale change between left and right panels. Average δ56Fe measured from bulk digestions is represented by the dashed black line, with 1-SD uncertainty indicated by the grey bar. All δ56Fe values are presented relative to igneous rocks. Uncertainties in δ56Fe are 1-SD errors of all replicates and duplicates. When not visible, error bars are smaller than the size of the symbol.
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Table 4 Fe isotope fractionation factors (Δ values) between aqueous Fe(II) and magnetite, 1 h 0.5 M HCl-extractable Fe, or 23 h 0.5 M HCl-extractable Fe. Fractionation factors were calculated from the data in Table 3. Depth interval
ΔFe(II)aq-magnetite
ΔFe(II)aq-1 h 0.5 M HCl
ΔFe(II)aq-23 h 0.5 M HCl
1–4.5 cm 8–11.5 cm 36–39.5 cm 57–60.5 cm Average ±SD
−1.33 −0.87 −1.85 −0.70 −1.19 ± 0.52
−1.11 −0.98 −1.87 −0.76 −1.18 ± 0.48
−1.05 −0.97 −1.96 −0.86 −1.21 ± 0.51
Fe(II)solid is represented by 1 h 0.5 M HCl + 23 h 0.5 M HCl extractions.
conditions) of a significant amount of isotopically fractionated Fe must take place. In modern sedimentary environments, high concentrations (approaching or exceeding mM levels) of low-δ56Fe aqueous Fe(II) are produced by partial microbial reduction of Fe(III) oxides (Severmann et al., 2006; Rouxel et al., 2008; Tangalos et al., 2010). Based on these results, mobilization and export of isotopically light Fe, produced by DIR in suboxic shelf sediments, to deep basins is hypothesized to be a main Fe isotope shuttling mechanism for marine Fe on Precambrian Earth (Severmann et al., 2008), resulting in widespread negative δ56Fe values for carbonates and sulfide minerals (i.e., Czaja et al., 2010). In addition, low-δ56Fe aqueous Fe(II) produced by partial reduction of Fe(III) oxides can undergo redistribution within the sediment column, producing distinct spatial patterns in δ56Fe values (Severmann et al., 2006; Tangalos et al., 2010). Mass transport of isotopically light Fe(II) components may also result in residual sediment developing positive δ56Fe values relative to initial material, and these deviations can be preserved during compaction and diagenesis, serving as a biological marker in the rock record (Beard et al., 2003; Severmann et al., 2008; Heimann et al., 2010; Percak-Dennett et al., 2011). In contrast to the above scenarios, complete reduction of Fe(III) oxides and limited sulfide formation in Bay of Vidy sediments results in no Fe isotope fractionation within the sediment column. In addition, the presence of oxic lake water prevents migration of aqueous Fe(II) into the water column, thus disallowing basin-scale redistribution of isotopically light Fe(II) components. The latter is an important distinction between modern oxic environments and Paleoproterozoic and Archean Earth, where an anoxic or only slightly oxic planet allowed for a greater mobilization of low-δ56Fe aqueous Fe(II), potentially driving significant Fe isotope redistribution in the sediment section. Specifically, in anoxic environments, a flux of low-δ56Fe aqueous Fe(II) into the water column synchronous with early-stage DIR is possible, resulting in sediment-wide changes to δ56Fe values, even if complete reduction of Fe(III) ultimately occurs. This mobilized low-δ56Fe(II) has the potential to then be shuttled into deep water basins, preserving isotopic evidence for DIR. Build up of oxygen on Earth resulted in increasingly oxic waters, with the potential for shelf margins and other shallow regions to become fully oxic, thus eliminating the potential for significant aqueous Fe(II) flux into the water column, and eliminating the potential for widespread mobilization and redistribution of Fe isotopes in marine sediment columns. Bay of Vidy sediments may therefore provide a model for shallow-water, organic-rich, low-S Precambrian marine sediments that were overlain by oxygen-containing waters. In particular, these sediments appear to be analogous, at least superficially, to magnetite-bearing sediments preserved in the 2.69 Ga Lewin Shale Formation within the Hamersley Basin in Western Australia, which have low reduced S contents and near-zero δ56Fe values (Yamaguchi et al., 2005). Although wide-spread oxygenation of the Earth did not take place until after 2.4 Ga (Canfield, 2005; Kump and Barley, 2007), recent rock record and ocean water column modeling analyses indicate that marine surface water contained significant oxygen concentrations (μM levels) as early as 2.7 Ga (Czaja et al., 2012). The results from Bay of Vidy thus provide a plausible explanation for the near-zero δ56Fe values in the Neoarchean Lewin Shale and analogous shallow-water Precambrian units.
5. Conclusions The Bay of Vidy is an organic-rich, low-S sedimentary environment in which near complete reduction of Fe(III) oxides has produced a variety of Fe(II)-bearing phases, including small amounts (0.5–8% of total Fe) of biogenic magnetite (present in the form of framboids composed of nm-sized crystallites) which account for documented magnetic susceptibility anomalies in these sediments (Gibbs-Eggar et al., 1999, this study). Although an isotopically light aqueous Fe(II) component exists in the sediment pore fluids, this pool of Fe is very small (b0.01% of total Fe) and a shift in its isotopic composition is not enough to cause deviations from bulk δ56Fe values in other Fe pools. As a result, the sediments display near-complete homogeneity in solid-phase Fe isotope compositions. This result contrasts with findings in other sedimentary environments, where some degree of fractionation and isotope redistribution has taken place during early digenetic processes (Severmann et al., 2006; Rouxel et al., 2008; Tangalos et al., 2010). If these sediments were to undergo diagenesis (heating, compaction, etc.) as they exist now, no heterogeneity or low-δ56Fe values would be preserved in the rock record, and the role of DIR in the history of these sediments would thus not be recorded. These findings illustrate the inability of Fe isotopes to accurately record DIR activity in systems that have undergone near complete Fe(III) reduction, and by contrast highlight the value of rocks in the geologic record that do show variations in δ56Fe composition. It is important to emphasize, however, that the observation of Fe isotope homogeneity in ancient (or modern) rocks is not sufficient to rule out a biological mechanism for sediment redox cycling and magnetite formation. Acknowledgments This work was supported by the NASA Astrobiology Institute. We gratefully acknowledge the assistance of D. Ortiz and J. Fournelle for assistance with SEM imaging, and H. Konishi for TEM work. Appendix A. Supplementary data Supplementary data associated with this article (including a complete listing of geochemical and Fe isotope data) will be made available online. Supplementary data to this article can be found online at http://dx.doi. org/10.1016/j.chemgeo.2013.10.008. References Aldana, M., Costanzo-Alvarez, V., Vitiello, D., Colmenares, L., Gumez, G., 1999. Framboidal magnetic minerals and their possible association to hydrocarbons: La Victoria oil field, southwestern Venezuela. Geofis. Int. Mex. 38, 137–152. Beard, B.L., Johnson, C.M., Skulan, J.L., Nealson, K.H., Cox, L., Sun, H., 2003. Application of Fe isotopes to tracing the geochemical and biological cycling of Fe. Chem. Geol. 195 (1–4), 87–117. Borch, T., Masue, Y., Kukkadapu, R.K., Fendorf, S., 2007. Phosphate imposed limitations on biological reduction and alteration of ferrihydrite. Environ. Sci. Technol. 41 (1), 166–172. Canfield, D.E., 2005. The early history of atmospheric oxygen: homage to Robert A. Garrels. Annu. Rev. Earth Planet. Sci. 33, 1–36. Cline, J.D., 1969. Spectrophotometric determination of hydrogen sulfide in natural saters. Limnol. Oceanogr. 14 (3), 454–458. Cloud, P., 1974. Evolution of ecosystems. Am. Sci. 62 (1), 54–66. Craddock, P.R., Dauphas, N., 2011. Iron and carbon isotope evidence for microbial iron respiration throughout the Archean. Earth Planet. Sci. Lett. 303 (1–2), 121–132. Czaja, A.D., Johnson, C.M., Beard, B.L., Eigenbrode, J.L., Freeman, K.H., Yamaguchi, K.E., 2010. Iron and carbon isotope evidence for ecosystem and environmental diversity in the similar to 2.7 to 2.5 Ga Hamersley Province, Western Australia. Earth Planet. Sci. Lett. 292 (1-2), 170–180. Czaja, A.D., Johnson, C.M., Roden, E.E., Beard, B.L., Voegelin, A.R., Nagler, T.F., Beukes, N.J., Wille, M., 2012. Evidence for free oxygen in the Neoarchean ocean based on coupled iron–molybdenum isotope fractionation. Geochim. Cosmochim. Acta 86, 118–137. Czaja, A.D., Johnson, C.M., Beard, B.L., Roden, E.E., Li, W., Moorbath, S., 2013. Biological Fe oxidation controlled deposition of banded iron formation in the ca. 3770 Ma Isua Supracrustal Belt (West Greenland). Earth Planet. Sci. Lett. 363, 192–203. Dauphas, N., van Zuilen, M., Busigny, V., Lepland, A., Wadhwa, M., Janney, P.E., 2007. Iron isotope, major and trace element characterization of early Archean supracrustal rocks from SW Greenland: protolith identification and metamorphic overprint. Geochim. Cosmochim. Acta 71 (19), 4745–4770.
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