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SOIL CHEMISTRY

Iron Isotope Fractionation during Pedogenesis in Redoximorphic Soils

^a Institute of Biogeochemistry and Pollutant Dynamics, ETH Zurich, CHN, 8092 Zurich, Switzerland, and Institute of Isotope Geochemistry and Mineral Resources, ETH Zurich, NW, 8092 Zurich, Switzerland
^b Institute of Isotope Geochemistry and Mineral Resources, ETH Zurich, NW, 8092 Zurich, Switzerland, and Institute of Earth Sciences, Hebrew Univ. of Jerusalem, 91904 Jerusalem, Israel
^c Institute of Biogeochemistry and Pollutant Dynamics, ETH Zurich, CHN, 8092 Zurich, Switzerland, and Department of Environmental Geosciences, Univ. of Vienna, Althanstrasse 14, 1090 Vienna, Austria
^d Institute of Isotope Geochemistry and Mineral Resources, ETH Zurich, NW, 8092 Zurich, Switzerland, and Dep. of Earth Sciences, Univ. of Oxford, Parks Road, Oxford, OX1 3PR, UK
^e Institute of Biogeochemistry and Pollutant Dynamics, ETH Zurich, CHN, 8092 Zurich, Switzerland

^* Corresponding author (wiederhold{at}env.ethz.ch).

	ABSTRACT

TOP NOTES ABSTRACT INTRODUCTION MATERIALS AND METHODS RESULTS AND DISCUSSION CONCLUSIONS REFERENCES

 
Stable Fe isotopes provide a potential new tool for tracing the biogeochemical cycle of Fe in soils. Iron isotope ratios in two redoximorphic soils were measured by multicollector inductively coupled plasma mass spectrometry to study the relationships between pedogenic Fe transformation and redistribution processes, and mass-dependent fractionation of Fe isotopes. Redoximorphic Fe depletion and enrichment zones were sampled in addition to the bulk soil samples. A three-step sequential extraction procedure was used to separate different Fe pools, which were examined in addition to total soil digests. Significant enrichments of heavy Fe isotopes of about 0.3 in ⁵⁷Fe were found in total soil digests of Fe-depleted zones compared with bulk soil samples and were explained by the preferential removal of light isotopes, presumably during microbially mediated Fe oxide dissolution under anoxic conditions. Accordingly, pedogenic Fe enrichment zones were found to be slightly enriched in light Fe isotopes. Distinct Fe isotope variations of >1 in ⁵⁷Fe were found between different Fe pools within soil samples, specifically enrichments of light isotopes in pedogenic oxides contrasting with heavy isotope signatures of residual silicate-bound Fe. Our data demonstrate that pedogenic Fe transformations in redoximorphic soils are linked to isotope fractionation, revealing greater mobility of lighter Fe isotopes compared with heavier isotopes during pedogenesis. No simple quantitative relationship between Fe depletion and isotope fractionation could be inferred, however. Our findings provide new insights into the behavior of Fe isotopes in soils and contribute to the development of Fe isotopes as a tracer for the biogeochemical Fe cycle.

	INTRODUCTION

TOP NOTES ABSTRACT INTRODUCTION MATERIALS AND METHODS RESULTS AND DISCUSSION CONCLUSIONS REFERENCES

 
Iron is not only an essential nutrient element for almost all organisms, but it also exerts a major influence on the mobility of nutrient and pollutant elements in soils (Stucki et al., 1988). Redox transformations between Fe(II) and Fe(III) play an important role in the biogeochemical cycle of Fe. Moreover, Fe transformations and translocations are key processes in soil formation and soil classification (van Breemen and Buurman, 2004). Iron in soils occurs in a variety of different phases such as primary silicate minerals, clay minerals, Fe (oxyhydr)oxide minerals of different crystallinity, as well as organically bound Fe (Stucki et al., 1988). Under atmospheric conditions, Fe(III) is the thermodynamically stable oxidation state. The weathering of primary silicate minerals in soils releases Fe(II), which is rapidly oxidized, followed by the precipitation of poorly crystalline Fe (oxyhydr)oxide minerals such as ferrihydrite. A transformation to better crystallized Fe oxide minerals, such as goethite and hematite, takes place during further soil development. Under atmospheric conditions, these phases are characterized by a very low solubility (Cornell and Schwertmann, 2003). Under anaerobic conditions in water-saturated soils, however, Fe(III) is used as an alternate electron acceptor by dissimilatory Fe-reducing bacteria (Lovley et al., 2004). This process results in the formation of soluble Fe(II) in soil solution, which is transported within the soil by advective and diffusive processes. Reprecipitation of Fe (oxyhydr)oxide minerals occurs as soon as this mobile Fe(II) pool comes into contact with O₂. This reprecipitation happens, depending on the temporal and spatial variability of the soil water regime, in characteristic horizons of the soil profile and results in the formation of typical Fe depletion and enrichment zones.

The corresponding soils are generally referred to as redoximorphic or hydromorphic soils (Schlichting and Schwertmann, 1973). They are formed by pedogenic processes that are summarized by the term redoximorphosis, and they are characterized by typical redoximorphic features (Vepraskas, 1996). In soils that are permanently water saturated, mainly at locations with a high groundwater table, Fe enrichments occur in a distinct horizon corresponding to the capillary fringe, whereas Fe depletion zones are found predominantly in soil horizons below the groundwater table. In contrast, soils that are only seasonally water saturated, typically because of stagnant water above a dense substrate layer that inhibits vertical drainage, exhibit a different pattern of redoximorphic features. In this case, the depletion zones are mainly found in the vicinity of larger pores and along preferential flow paths where water infiltrates rapidly and reducing conditions are generated. The enrichment zones are then mainly found in the interior of soil aggregates and in the soil matrix, where remaining O₂ results in the precipitation of Fe (oxyhydr)oxide minerals, often in the form of nodules or concretions (Vepraskas, 1996). These pedogenic redistribution processes of Fe within hydromorphic soils not only are interesting in terms of soil morphology and classification but also influence the fate of other elements of environmental interest such as As (Cummings et al., 1999) and P (Szilas et al., 1998). Despite extensive research during the last few decades on biogeochemical Fe cycling in soils, there are still many unsolved issues. These include, for instance, the importance of different Fe (oxyhydr)oxide minerals as electron acceptors for bacterial respiration in natural environments or mineral dissolution kinetics and mechanisms that influence nutrient and pollutant element cycling.

The analysis of stable isotope ratios represents an important method to elucidate biogeochemical reactions and element cycles in the environment (Hoefs, 2004). Until recently, however, this approach was confined to isotope ratios of lighter elements (e.g., C, O, H, N, and S), which can be measured in the gas phase. The development of new analytical methods, mainly multiple collector inductively coupled plasma mass spectrometry (Halliday et al., 1998), has expanded this range to heavier elements such as Fe and opened up a new field of isotope geochemistry (Johnson et al., 2004a). Iron isotopes may provide a new tool to trace the biogeochemical Fe cycle of soils. Iron has four stable isotopes (natural abundance percentage): ⁵⁴Fe (5.84%), ⁵⁶Fe (91.76%), ⁵⁷Fe (2.12%), and ⁵⁸Fe (0.28%). The notation is commonly used to describe Fe isotope fractionation relative to the international Fe isotope standard IRMM-014 and is defined as

[1]

or

[2]

The two values can be easily converted into each other by the approximation ⁵⁷Fe = 1.5 ⁵⁶Fe because the observed fractionation effects are mass dependent (Dauphas and Rouxel, 2006). Variations of ⁵⁶Fe in bulk igneous rocks were found to be very small (Beard and Johnson, 1999, 2004). In contrast, significant deviations from this average terrestrial background ratio of about 4 occur in various low-temperature environments such as sediments or soils (Brantley et al., 2001; Wiederhold and von Blanckenburg, 2002; Fantle and DePaolo, 2004; Matthews et al., 2004; Emmanuel et al., 2005; Thompson et al., 2007). It has been shown that Fe isotopes can be fractionated by kinetic and equilibrium isotope effects during both biotic and abiotic reactions (Anbar, 2004; Johnson et al., 2004b; Dauphas and Rouxel, 2006). Laboratory studies have shown that dissolution of Fe (oxyhydr)oxide minerals by dissimilatory Fe-reducing bacteria results in an enrichment of light Fe isotopes in solution (Beard et al., 1999; Crosby et al., 2005, 2007). The abiotic dissolution of goethite in the presence of oxalate by both a photochemical reductive and a ligand-controlled mechanism has been shown to enrich light Fe isotopes in the first dissolved fractions (Wiederhold et al., 2006). In contrast, the abiotic oxidative precipitation of ferrihydrite has been shown to favor heavy Fe isotopes in the reaction product (Bullen et al., 2001). Furthermore, fractionation of Fe isotopes during adsorption processes in reduced groundwater systems was investigated by Teutsch et al. (2005). They found a preferential adsorption of heavy Fe isotopes during the reaction of dissolved Fe(II) and Fe(III) (oxyhydr)oxide minerals in a natural aquifer. Iron isotope fractionations along redox gradients in marine sediments were recently reported by Severmann et al. (2006) and Staubwasser et al. (2006), with enrichments of light Fe isotopes in extracts of the reactive Fe mineral pools. The mechanisms, however, that govern the distribution of Fe isotopes during redox processes in natural environments are still poorly understood. So far, there are very few data available on Fe isotope fractionation in redoximorphic soils.

Our objective was to investigate Fe isotope variations in natural soil environments with distinct Fe redox dynamics and to relate Fe isotope fractionation to pedogenic processes in redoximorphic soils. Therefore, we performed a detailed study of the spatial distribution of Fe isotope ratios in relation to hydromorphic properties in two soils exhibiting contrasting soil pH and seasonal water regimes. The first soil was seasonally saturated by stagnant water and exhibited an acidic soil pH, while the second soil was permanently water saturated in the subsoil due to groundwater and had a neutral soil pH. Separation of Fe depletion and enrichment zones during sampling enabled a detailed investigation of small-scale variations during redox transformations. In addition, the separation of different Fe pools by sequential extractions and their subsequent Fe isotope analysis in addition to total soil digestions provided further insights into the variability and fractionation of Fe isotopes in redoximorphic soils.

	MATERIALS AND METHODS

TOP NOTES ABSTRACT INTRODUCTION MATERIALS AND METHODS RESULTS AND DISCUSSION CONCLUSIONS REFERENCES

 
Sampling Sites
Soil samples were collected at two sites representing an acidic, seasonally water-saturated soil and a neutral soil that is permanently water saturated in the subsoil horizons. The two soils are classified as Typic Epiaquept and Typic Humaquept, respectively, according to the U.S. Soil Taxonomy (Soil Survey Staff, 2006). According to the World Reference Base for Soil Resources (International Union of Soil Scientists, 2006), the profiles are classified as Stagnic Cambisol and Haplic Gleysol, respectively. The World Reference Base nomenclature for soil types and horizon designations will be used. The two sampled profiles are depicted in Fig. 1 along with designations of the pedogenic horizons. Table 1 lists selected soil properties including soil colors of the two investigated profiles.

View larger version (113K): [in this window] [in a new window]   Fig. 1. Photographs of the Stagnic Cambisol (near Rafz, Switzerland) and the Haplic Gleysol (near Tettnang, Germany). Scale bars indicate 10-cm sections. Both profiles exhibit clear redoximorphic features. Horizon designations correspond to the collected samples and follow FAO (2006): O = organic surface layer, i = slightly decomposed organic material, e = moderately decomposed organic material, Ah = organic-rich mineral horizon, B = subsurface mineral horizon altered by pedogenic processes, w = weathered, C = soil substrate, c = concretions or nodules, g = stagnic conditions, l = capillary fringe mottling (gleying), r = strong reduction; prefix numbers indicate lithogenic discontinuities, suffix numbers indicate vertical subdivision of horizon.

The second soil profile (the Haplic Gleysol) was sampled in southern Germany near Tettnang, Upper Swabia (47°39'N, 9°33'E). It formed on carbonate-rich sandy sediments that were deposited by fluvial processes during the Pleistocene and later covered by a layer of loamy sediments that makes up the upper part of the soil profile (Kösel and Vogl, 1997). The vegetation at this site is dominated by deciduous trees [Fraxinus excelsior L., Alnus glutinosa (L.) Gaertn., and Quercus robur L.]. No organic surface layer exists because the high biological activity and easily degradable plant litter prevent accumulation of organic matter. Soil pH(CaCl₂) ranges from 5.9 in the upper mineral soil horizon (Ah) to 7.7 at a depth of 120 cm, a range that is classified as moderately acidic to slightly alkaline (Soil Survey Division Staff, 1993). Carbonate is present below a depth of about 40 cm, buffering the soil pH to values >7.5. The deeper soil horizons are permanently saturated by groundwater and the capillary fringe is situated at a depth of about 50 cm, with little seasonal variation (Kösel and Vogl, 1997). As a consequence, the deeper soil horizons are subjected to reducing conditions, which has led to significant Fe mobilization and subsequent development of matrix depletion features. Iron enrichment zones are prominent at depths around 50 to 70 cm, where the reaction of mobilized Fe(II) with atmospheric O₂ results in oxidative precipitation of Fe(III) (oxyhydr)oxide minerals.

Soil Sampling and Sample Preparation
The soil profiles were sampled according to pedogenic horizons (eight depths per profile, Fig. 1). Additionally, two horizons were chosen in each profile for detailed sampling of distinct redoximorphic features. In the Stagnic Cambisol profile, the Bcg2 horizon (80 cm) and the CBg horizon (120 cm) exhibiting distinct Fe depletion and enrichment zones were selected. At these depths, separate samples of the gray depleted zones and the brown enriched zones were collected in addition to the bulk soil samples. For the Haplic Gleysol profile, the brown redoximorphic enrichment zones of the 2Blc1 horizon (50 cm) and the 2Blc2 horizon (70 cm) were separated from the bulk soil matrix samples. No specific Fe depletion zones could be sampled because the whole soil matrix was depleted with respect to Fe in these horizons.

The soil samples were dried at 40°C in the laboratory and passed through a 2-mm sieve. An aliquot of the samples was ground to a fine powder with a rotary disk mill. From this powder, wax pellets were produced to measure total elemental concentrations by energy-dispersive x-ray fluorescence analysis (Spectro-X-Lab 2000, Spectro, Kleve, Germany). A small amount of the powdered samples (100–500 mg depending on Fe content) was dissolved completely in a microwave digestion with HF–HNO₃–HCl (1:2:2 mixture). These samples represent the total Fe pool of the soil samples. Samples rich in organic matter were pretreated with 30% H₂O₂ before the digestion. The clear solutions after the digestion were evaporated in Teflon beakers on a hot plate to remove excess HF. The residue was redissolved in concentrated HNO₃ to ensure complete oxidation of Fe(II) to Fe(III). This solution was again evaporated and the residue taken up in 6 mol L^–1 HCl. All reagents used during sample preparation were at least reagent grade and prepared with ultrapure water (>18 M cm, Milli-Q, Millipore, Billerica, MA). Hydrochloric and nitric acids were further purified by subboiling distillation. Hydrofluoric acid used during digestions was suprapure quality (Merck, Darmstadt, Germany).

In parallel, another aliquot of the dried and sieved soil samples was subjected to a three-step sequential extraction procedure. The three steps were designed to separate operationally defined Fe pools from the soil samples, which are dominated by the following mineral phases: (i) poorly crystalline Fe (oxyhydr)oxides, (ii) crystalline Fe (oxyhydr)oxides, and (iii) silicate-bound Fe.

In the first extraction step, 2 g of the soil material was weighed into 50-mL centrifuge tubes, 40 mL of 0.5 mol L^–1 HCl was added to the tubes, and the samples were placed on an end-over-end shaker for 24 h at room temperature. Afterward, the tubes were centrifuged (3400 x g, 15 min) and the supernatants decanted and syringe filtered through 0.45-µm nylon membrane filters (Opti-Flow, WiCom GmbH, Heppenheim, Germany). This solution was defined as the Fe_HCl fraction of the soil, representing mainly poorly crystalline Fe oxyhydroxides.

In the second extraction step, 40 mL of a 1 mol L^–1 NH₂OH–HCl solution in 1 mol L^–1 HCl was added to the residue in the centrifuge tubes and shaken vigorously. The tubes were then placed in a hot water bath at 90°C with an integrated horizontal shaker for 4 h. During this period, the samples were additionally manually shaken and turned upside down several times. Afterward, the tubes were centrifuged (3400 x g, 15 min) and the supernatants decanted and filtered through 0.45-µm nylon membrane filters. This solution was defined as the Fe_NH2OH-HCl fraction of the soil, representing mainly crystalline Fe oxyhydroxide minerals.

The residue was washed twice with water, dried overnight at 105°C, and then ground to a fine powder with a rotary disk mill. An aliquot of this powder, between 100 and 500 mg depending on the Fe content, was then completely dissolved in a microwave digestion with HF–HNO₃–HCl (1:2:2 mixture) and further processed like the total Fe samples. This solution was defined as the Fe_residue fraction of the soil, representing mainly silicate-bound Fe.

The extraction solutions (Fe_HCl and Fe_NH2OH-HCl) were also evaporated in Teflon beakers on a hotplate and oxidized with HNO₃ and H₂O₂ to destroy organic matter and hydroxylamine and to convert Fe(II) to Fe(III). Their residue was then also taken up in 6 mol L^–1 HCl.

Iron concentrations of all samples were measured by atomic absorption spectrometry (SpectrAA 220, Varian, Melbourne, Australia). The Fe concentrations of the reagents used during extractions and digestions were found to be negligible compared with the Fe content of all samples. The organic surface layer samples of the Stagnic Cambisol profile were not subjected to the entire sequential extraction procedure because the method is only applicable for mineral soil samples. For these samples, only the total digestion and the first extraction step with 0.5 mol L^–1 HCl were performed. Further sample preparation took place in a clean chemistry laboratory. Teflon microcolumns filled with about 1 mL of anion exchange resin (Bio-Rad AG1 X4, 200–400 mesh, Bio-Rad Laboratories, Hercules, CA) were used to separate Fe in the samples from matrix elements (Strelow, 1980). In 6 mol L^–1 HCl, Fe(III) is present as the FeCl₄^– anion. The Fe complex was retained on the resin while the sample matrix was washed out by repeated additions of 6 mol L^–1 HCl. Quantitative elution of Fe from the columns was achieved with 0.05 mol L^–1 HCl. The eluates were again evaporated and finally taken up in 0.05 mol L^–1 HCl as solution matrix for the Fe isotope measurement.

Iron Isotope Measurement
Iron isotope ratios were determined by multiple collector inductively coupled plasma mass spectrometry (Nu Plasma, Nu Instruments, Wrexham, North Wales, UK). Comprehensive descriptions of Fe isotope analytical methods were recently published by Schoenberg and von Blanckenburg (2005) and de Jong et al. (2007). The analytical procedures for Fe isotope measurement in our laboratory have been previously described in detail (Williams et al., 2004; Teutsch et al., 2005). Briefly, a standard-bracketing approach was used to correct for machine drift and instrumental mass bias. A membrane desolvation system (MCN-6000, Cetac Technologies, Omaha, NE) was used to minimize argide interferences (ArN⁺, ArO⁺, ArOH⁺) to insignificant levels (background/signal ratio typically <0.001). The ⁵⁷Fe/⁵⁴Fe and ⁵⁶Fe/⁵⁴Fe ratios were measured simultaneously and all data plotted on the theoretical mass fractionation line, demonstrating the absence of isobaric interferences. Chromium was monitored at mass 52 or 53 to calculate the potential influence of ⁵⁴Cr on ⁵⁴Fe; however, our purified solutions did not contain significant amounts of Cr. Hence, the Cr-corrected and uncorrected ⁵⁷Fe values differed by less than ±0.02 for all samples. All masses were collected in Faraday cups equipped with 10¹¹ resistors except mass 56, which was collected in a Faraday cup equipped with a 10⁹ resistor. This allowed analysis of solutions with relatively high Fe concentrations (8 mg L^–1) but affected the precision of the ⁵⁶Fe/⁵⁴Fe measurement. Therefore, due to the smaller analytical error for the ⁵⁷Fe/⁵⁴Fe ratio, the results are expressed as ⁵⁷Fe. Samples were only measured after several stable isotope measurements of an internal laboratory standard (ETH hematite or ETH Fe salt standard, prepared in an identical matrix as the samples) against IRMM-014. This standard was again measured after every six samples and at the end of the analytical run. The long-term reproducibility of ⁵⁷Fe of our internal house standards is better than ±0.15 (2 SD). Samples were measured several times and the error bars for ⁵⁷Fe represent either the reproducibility (2 SD) of replicate sample measurements or, in case of fewer measurements (n < 3), the reproducibility (2 SD) of our internal house standard during the same analytical session. A mass balance approach was used to verify the Fe concentration and Fe isotope results of the different Fe fractions. The isotope mass balance was calculated according to the following equation, where [Fe]_n is the Fe concentration in pool n:

[3]

The error bars of the calculated total Fe value were adapted with the following equation:

[4]

Iron isotope fractionation of depleted or enriched zones relative to bulk soil is denoted as ⁵⁷Fe_{depleted–bulk} = ⁵⁷Fe_depleted – ⁵⁷Fe_bulk or ⁵⁷Fe_{enriched–bulk} = ⁵⁷Fe_enriched – ⁵⁷Fe_bulk. Error bars of ⁵⁷Fe values were adapted according to the following equations, where 2 SD[⁵⁷Fe_n] indicates the double standard deviation of the isotope ratio of pool n:

[5]

or

[6]

	RESULTS AND DISCUSSION

TOP NOTES ABSTRACT INTRODUCTION MATERIALS AND METHODS RESULTS AND DISCUSSION CONCLUSIONS REFERENCES

 
Iron Concentration Profiles in Total Soil Digests
Both soil profiles exhibited pronounced redoximorphic features such as Fe mottling and Mn concretions (Fig. 1). The depth profiles of total Fe concentration in the two soil profiles are presented in Fig. 2 . The Stagnic Cambisol profile displayed higher total Fe concentrations in the subsoil horizons than the topsoil horizons. This difference was probably not caused by pedogenic processes and may instead be explained by a lower initial Fe content of the topsoil substrate resulting from dilution by loess deposition. The lowest total Fe concentrations were present in the two organic surface layer samples, which is consistent with the low Fe content of plant litter compared with the mineral soil material. Significant Fe concentration differences were found in the enriched and depleted zones compared with the bulk soil samples in the seasonally water-saturated Bcg2 and CBg horizons, as expected from visual observation in the field. The gray depletion zones contained considerably less Fe than the bulk soil samples, while the brown enrichment zones exhibited slightly higher Fe concentrations. The concentration data indicated that about 30% of the total Fe in the bulk soil has been leached from the depleted zones. We were not able to perform a quantitative mass balance of Fe fluxes within the profile because the relative proportions of depleted and enriched zones were difficult to quantify and because the soil is an open system from which some Fe may have been lost during soil formation. Some Fe that was mobilized during anoxic periods has presumably left the soil profile by vertical or lateral transport. In any case, the significant concentration differences were clear indications of pedogenic Fe redistribution as a result of Fe redox transformations. Other elements besides Fe were also strongly influenced by the described redox cycling, which was confirmed by x-ray fluorescence (XRF) analysis of depleted and enriched zones (data not shown). Elements such as Mn, As, P, and Pb showed similar depletion and enrichment patterns as Fe, corresponding either to their own redox chemistry or because they are bound to Fe minerals. In contrast, no significant concentration differences between Fe-depleted and Fe-enriched zones were found for redox-insensitive elements such as Si, Al, Ca, Na, and Ti.

View larger version (16K): [in this window] [in a new window]   Fig. 2. Depth profiles of total Fe concentration in the two soil profiles. The redoximorphic features of the Bcg2 and the CBg horizon of the Stagnic Cambisol and of the 2Blc1 and the 2Blc2 horizon of the Haplic Gleysol were sampled separately. Open symbols indicate the calculated sum of the three Fe fractions that were separated by sequential extractions (Fig. 4). Horizon designations follow FAO (2006): B = subsurface mineral horizon altered by pedogenic processes, C = soil substrate, c = concretions or nodules, g = stagnic conditions, l = capillary fringe mottling (gleying); suffix numbers indicate vertical subdivision of horizon.

View larger version (20K): [in this window] [in a new window]   Fig. 4. Depth profiles of Fe concentrations for the three Fe fractions poorly crystalline Fe oxyhydroxides (FeHCl), crystalline Fe oxyhydroxides (FeNH2OH-HCl), and silicate-bound Fe (Feresidue) for bulk soil samples and depleted and enriched zones. Note different concentration scale bars.

Samples of unweathered parent materials would have been helpful for assessing element depletions and to investigate the open-system behavior of the soils with respect to element losses by vertical and lateral transport. Unfortunately, it was not possible to obtain such unweathered materials because of the geologic settings of the two sampling sites. The parent materials, the glacial moraine deposit for the Stagnic Cambisol and the fluvial deposits for the Haplic Gleysol, were sediments that already contained partially weathered mineral components at the time of deposition. In this study, we focused on relative differences among soil horizons and between redoximorphic depletion and enrichment zones within the soil profiles.

Iron Isotope Ratios in Total Soil Samples
The results of the Fe isotope measurements of the total soil digests of the bulk soil samples exhibited a relatively narrow range for the Stagnic Cambisol profile, with ⁵⁷Fe values of 0.0 to 0.2 relative to IRMM-014 (Fig. 3 ). The depleted zones of the Bcg2 and CBg horizons, however, had a significantly heavier Fe isotope signature, with ⁵⁷Fe values of 0.4 and 0.5, respectively. The enriched zones of the two horizons were not significantly fractionated relative to the bulk soil samples. These findings corresponded to the Fe concentration data (Fig. 2), where the depleted zones also showed more pronounced concentration changes than the bulk soil samples. The heavier ⁵⁷Fe values in the depleted zones are consistent with the concept of the preferential removal of light Fe isotopes during microbial dissimilatory Fe reduction (Beard et al., 1999; Crosby et al., 2005, 2007), which results in a relative enrichment of heavy isotopes and thus higher ⁵⁷Fe values. The bulk soil data of the Haplic Gleysol profile showed a very similar range of ⁵⁷Fe values of 0.0 to 0.2, except for the two top horizons, Ah1 and Ah2, which exhibited ⁵⁷Fe values of 0.3 and 0.4, respectively. As mentioned above, these horizons consisted of a genetically different parent material and cannot be directly compared with the soil material below. The sample collected from the enriched zone of the 2Blc2 horizon at 70-cm depth had a significantly lighter Fe isotope signature of –0.15. This effect was probably caused by the preferential mobilization of light Fe isotopes from the subsoil and quantitative oxidative precipitation of this light Fe pool in the enrichment zone of the soil profile. The enriched zone of the second horizon (2Blc1), however, did not exhibit a significant Fe isotope fractionation relative to the bulk soil. Iron isotope ratios of the bulk soil samples increased slightly with increasing depth from 0.01 at 50 cm to 0.18 at 110 cm. This is consistent with the upward transport of isotopically light Fe during pedogenic Fe transformation under reducing conditions in the subsoil.

View larger version (15K): [in this window] [in a new window]   Fig. 3. Depth profiles of 57Fe values in total soil digests of the two soil profiles relative to the international Fe isotope standard IRMM-014. Error bars represent 2 SD of replicate measurements. The redoximorphic features of the Bcg2 and the CBg horizon of the Stagnic Cambisol and of the 2Blc1 and the 2Blc2 horizon of the Haplic Gleysol were sampled separately. Open symbols indicate the calculated isotope mass balance of the three Fe fractions that were separated by sequential extractions (Fig. 4 and 5). Horizon designations follow FAO (2006): B = subsurface mineral horizon altered by pedogenic processes, C = soil substrate, c = concretions or nodules, g = stagnic conditions, l = capillary fringe mottling (gleying); suffix numbers indicate vertical subdivision of horizon.

View larger version (20K): [in this window] [in a new window]   Fig. 5. Depth profiles of 57Fe for the three Fe fractions poorly crystalline Fe oxyhydroxides (FeHCl), crystalline Fe oxyhydroxides (FeNH2OH-HCl), and silicate-bound Fe (Feresidue) for bulk soil samples and depleted and enriched zones relative to the international Fe isotope standard IRMM-014. Error bars represent 2 SD of replicate measurements. Note different 57Fe scale in Fig. 5C.

We decided to use a newly developed three-step sequential extraction procedure in our study, rather than the classical oxalate and dithionite methods (Borggaard, 1988). The first extraction step by HCl (0.5 mol L^–1, 24 h, 25°C) dissolves poorly crystalline Fe (oxyhydr)oxide minerals such as ferrihydrite as well as adsorbed and some organically bound Fe (Heron et al., 1994). While the dissolution of goethite in the presence of oxalate can result in significant Fe isotope fractionation (Wiederhold et al., 2006), no Fe isotope fractionation was observed during the stepwise dissolution of goethite by 0.5 mol L^–1 HCl (Wiederhold et al., 2006). This is in agreement with previous studies on hematite dissolution in HCl (Skulan et al., 2002). The drying procedure at 40°C before the extractions presumably resulted in the oxidation of Fe(II) phases in the soil samples and their precipitation in the form of poorly crystalline Fe (oxyhydr)oxide mineral phases. These phases, however, were probably redissolved quantitatively in the first extraction step.

The second extraction step with NH₂OH–HCl (1 mol L^–1, 4 h, 90°C, modified after Ribet et al., 1995) was calibrated to achieve a complete reductive dissolution of crystalline Fe (oxyhydr)oxides such as goethite and hematite. It is possible, however, that some Fe that was bound to clay minerals or other silicate minerals was extracted as well during this extraction step. Thus, a small bias may have been introduced into the isotope data of the last two steps of the sequential extraction procedure.

In the third step, all remaining soil material, which still contained Fe in silicate minerals, was totally dissolved by a microwave digestion procedure with HF–HNO₃–HCl. The sequential extractions were done in parallel with the total soil digests, which allowed us to perform an isotope mass balance between the sum of the three extractions and the total digests.

Iron Concentration Profiles in Sequential Extraction Samples
The results of the Fe concentration measurements of the three-step sequential extraction procedure are displayed in Fig. 4 . In both soil profiles, the crystalline Fe oxide fraction (Fe_NH2OH-HCl) exhibited the strongest variability within the profile, reflecting the pedogenic Fe transformation processes (Fig. 4B and 4E). The Fe enrichment at the capillary fringe of the Haplic Gleysol profile was clearly dominated by this pool. In the Stagnic Cambisol profile, the Fe_NH2OH-HCl fraction represented the dominant Fe pool in all horizons. The relatively small proportion of silicate-bound Fe (Fe_residue, Fig. 4C) compared with oxide-bound Fe can be explained by the intense weathering in the Stagnic Cambisol profile due to the strong acidity and the age of the soil. In contrast, silicate-bound Fe still represented the dominant Fe pool in the subsoil of the Haplic Gleysol (Fig. 4F), where the extent of weathering was less intense due to the neutral pH and the younger age of the soil. The Fe_residue fractions of both soil profiles showed increasing concentrations with soil depth, which can be explained by more intense weathering in the upper soil horizons. In contrast to this trend, the two topsoil horizons of the Haplic Gleysol profile exhibited a higher Fe_residue content. This is probably caused, however, by the different parent material compared with the underlying sandy horizons. The observed lesser degree of silicate weathering is consistent with the younger age of the cover sediment. The Fe_HCl fraction, which mainly consists of poorly crystalline Fe(oxyhydr)oxide minerals such as ferrihydrite, was the smallest Fe pool in most soil samples. It showed almost no variation with soil depth in the Stagnic Cambisol profile (Fig. 4A) except for the uppermost organic surface layers. In contrast, the pedogenic Fe enrichment at about 50-cm soil depth of the Haplic Gleysol profile (Fig. 4D) was clearly apparent in the Fe_HCl fraction, indicating that poorly crystalline Fe (oxyhydr)oxides are partly responsible for the higher Fe content of this horizon. Mass balance considerations dictate that the sum of the Fe concentrations of the three sequential extractions corresponds to the Fe content of the total soil digestion. The open symbols in Fig. 2 represent the calculated sum of the three separated Fe pools. The close match with the measured total Fe values demonstrated the excellent agreement of the two data sets, confirming the validity of our separation procedure.

The Fe concentrations of the three extracted soil Fe pools in the separated depletion and enrichment zones of the selected horizons are also depicted in Fig. 4. The strongest redoximorphic effect is apparent in the Fe_HCl and Fe_NH2OH-HCl fractions of the Stagnic Cambisol profile, with distinctly lower Fe concentrations in the depleted zones than in the bulk soil. The enriched zones of the Haplic Gleysol profile exhibited higher Fe concentrations in the Fe_HCl and Fe_NH2OH-HCl fractions than in the bulk soil. No significant concentration differences between depletion or enrichment zones and the bulk soil samples were found in the Fe_residue fraction, which indicates that silicate-bound Fe does not play a significant role in the redoximorphic Fe redistribution within this soil profile.

Iron Isotope Ratios in Sequential Extraction Samples
The ⁵⁷Fe values of the three fractions, Fe_HCl, Fe_NH2OH-HCl, and Fe_residue, which were separated by sequential extractions, are presented in Fig. 5 . The range of observed Fe isotope ratios was much wider in this data set than in the total digestion samples. This indicated that Fe isotope fractionation did occur not only between different zones and horizons of the soil but also among different Fe mineral pools. The calculated isotope mass balance from the sum of the three fractions separated by sequential extractions (indicated by open symbols and dashed lines in Fig. 3) showed an excellent agreement with the measured bulk soil data in the case of the Haplic Gleysol profile. The data from the deeper horizons of the Stagnic Cambisol profile were less congruent and displayed a slight offset between measured and calculated values; however, the difference of about 0.2 was relatively small compared with the range of fractionations observed in the sequential extraction samples. The Fe_HCl and Fe_NH2OH-HCl fractions of the Stagnic Cambisol profile had ⁵⁷Fe values close to IRMM-014 (Fig. 5A and 5B). The Fe_HCl fraction was isotopically slightly lighter than the Fe_NH2OH-HCl fraction across the whole soil depth. The Fe_HCl fraction of the organic surface layer samples exhibited very light ⁵⁷Fe values of up to –0.7, which may be explained by the preferential incorporation of light Fe isotopes into the plant biomass constituting the source of the organic surface layer. This hypothesis is consistent with recent findings that plant biomass is enriched in light Fe isotopes (Walczyk and von Blanckenburg, 2002, Guelke and von Blanckenburg, 2007). A strong enrichment of heavy isotopes was found in the Fe_residue fraction, which exhibits ⁵⁷Fe values of up to 1.6 (Fig. 5C). These pronounced enrichments of heavy isotopes point toward an advanced degree of silicate weathering due to the acidity and the age of the soil profile. The preferential transformation of light Fe isotopes during weathering processes resulted in relative enrichments of heavy Fe isotopes in the residue. The Fe concentration data of the separated Fe mineral pools (Fig. 4A–4C) indicated that the largest proportion of the Fe in the soil has already been transformed to secondary mineral phases during pedogenesis.

The Haplic Gleysol data showed a similar succession of increasing ⁵⁷Fe values along the sequential extraction procedure as in the Stagnic Cambisol. The difference between the Fe fractions was most pronounced in the deeper soil horizons, with ⁵⁷Fe values of around –0.5 in the Fe_HCl fraction (Fig. 5D), of around 0.0 in the Fe_NH2OH-HCl fraction (Fig. 5E), and of around 0.5 in the Fe_residue fraction (Fig. 5F). This trend is consistent with the concept of preferential mobilization of light Fe isotopes during pedogenesis, because the light Fe_HCl fraction corresponds to the least stable and youngest Fe pool. In contrast, the heavy Fe_residue fraction is made up of Fe in silicate minerals, which constitute the depleted residual product of weathering and soil formation processes. The younger cover sediment layer, which makes up the two top horizons of the Haplic Gleysol profile (Ah1 and Ah2), exhibited a smaller range in ⁵⁷Fe, with values between 0 and 0.5. This can be explained by the lesser degree of weathering because of the younger age, but also by the absence of redoximorphic Fe transformations in these oxic upper soil horizons. Significant Fe isotope fractionations were also found between the individual Fe mineral pools of the samples from the redoximorphic Fe depletion and enrichment zones (Fig. 5). Similar to the bulk soil samples, a distinct enrichment of heavy Fe isotopes was found in the Fe_residue fraction. The Fe_HCl and Fe_NH2OH-HCl fractions of the Stagnic Cambisol profile exhibited systematic differences in ⁵⁷Fe relative to IRMM-014, with enrichments of heavy Fe isotopes in the samples from the depleted zones. In contrast, an enrichment of light Fe isotopes was found in the Fe_HCl fraction of the enriched zones in the Stagnic Cambisol profile. The same effect was observed in the samples of the Haplic Gleysol profile, with light Fe isotope signatures in the Fe_HCl fraction of the enriched zones of up to –0.6 in ⁵⁷Fe relative to IRMM-014.

Interpretation of Iron Isotope Effects
The comparison of Fe concentration data and Fe isotope ratios in Fe-enriched and Fe-depleted zones relative to the bulk soil data for the Stagnic Cambisol profile revealed an evident link between pedogenic Fe redistribution and Fe isotope fractionation (Fig. 6 ). Iron depletion zones exhibited depletions of light Fe isotopes and consequently relative enrichments of heavy Fe isotopes, whereas Fe enrichment zones tended to be slightly enriched in light Fe isotopes. The Stagnic Cambisol profile showed the strongest Fe depletion effect in the Fe_NH2OH-HCl fraction of the Bcg2 horizon, with concentration differences of 77% relative to the bulk soil. This sample also displayed a significant Fe isotope fractionation effect, with an enrichment of heavy isotopes of 0.26 in ⁵⁷Fe relative to the bulk soil. The strongest Fe isotope effects were found in the Fe_HCl fraction of the depleted zones, with enrichments of heavy isotopes of 0.55 and 0.59 in ⁵⁷Fe corresponding to Fe depletions of 55 and 53%, respectively. Figure 7 presents the relative differences in Fe concentration and Fe isotope ratios between the enrichment zones and the bulk soil samples of the Haplic Gleysol profile. The pedogenic enrichment zones had higher Fe concentrations, relative to the bulk soil, of up to 200% in the Fe_HCl fraction of the 2Blc2 horizon. The corresponding Fe isotope ratio exhibited an enrichment of light Fe isotopes of 0.15 in ⁵⁷Fe. It was apparent, however, that no simple quantitative relationship between Fe concentration differences and Fe isotope ratios in the studied redoximorphic soils existed.

View larger version (18K): [in this window] [in a new window]   Fig. 6. Difference in Fe concentration and Fe isotope ratios in the enriched and depleted zones of the Stagnic Cambisol profile relative to bulk soil samples. 57Fe represents Fe isotope fractionation relative to the bulk soil sample (57Fe = 57Fedepleted – 57Febulk or 57Fe = 57Feenriched – 57Febulk); FeHCl represents poorly crystalline Fe oxyhydroxides, FeNH2OH-HCl represents crystalline Fe oxyhydroxides, and Feresidue represents silicate-bound Fe. Error bars of 57Fe were adapted according to Eq. [4–6]. Horizon designations follow FAO (2006): B = subsurface mineral horizon altered by pedogenic processes, C = soil substrate, c = concretions or nodules, g = stagnic conditions; suffix numbers indicate vertical subdivision of horizon.

View larger version (15K): [in this window] [in a new window]   Fig. 7. Difference in Fe concentration [%] and Fe isotope ratios in the enriched zones of the Haplic Gleysol profile relative to bulk soil samples. 57Fe represents Fe isotope fractionation relative to the bulk soil sample (57Fe = 57Feenriched – 57Febulk) ; FeHCl represents poorly crystalline Fe oxyhydroxides, FeNH2OH-HCl represents crystalline Fe oxyhydroxides, and Feresidue represents silicate-bound Fe. Error bars of 57Fe were adapted according to Eq. [4–6]. Horizon designations follow FAO (2006): B = subsurface mineral horizon altered by pedogenic processes, C = soil substrate, c = concretions or nodules, l = capillary fringe mottling (gleying); suffix numbers indicate vertical subdivision of horizon.

The pedogenic Fe depletion under seasonally water-saturated conditions is predominantly mediated by dissimilatory Fe-reducing bacteria, which are able to use Fe(III) as an alternate terminal electron acceptor (Lovley et al., 2004). The direction of the observed isotope effect is consistent with the observations from previous studies reporting a light Fe isotope signature of the solution in laboratory experiments with dissimilatory Fe-reducing bacteria (Beard et al., 1999; Crosby et al., 2005, 2007). We are not able to provide direct evidence, however, that the observed Fe isotope effects were caused by microbially mediated Fe oxide dissolution. The magnitude of the Fe isotope fractionation effects between Fe-depleted and -enriched zones of redoximorphic soils presented in this study was about 0.5 in ⁵⁷Fe. This was considerably smaller than the measured fractionation effects between substrate and solution in laboratory studies of about 1.3 in ⁵⁶Fe (Johnson et al., 2004b), which would correspond to about 1.95 in ⁵⁷Fe. It is important to keep in mind, however, that we focused on the Fe isotope variations between solid Fe phases in our field study. Although the depleted zones of our soil constituted the substrate of the microbial Fe reduction, the magnitude of the fractionation effects cannot be directly compared with the laboratory studies. This is due to the fact that the solid mineral Fe phases in a soil sample constitute large Fe pools compared with the small amounts of Fe that are present in the soil solution. Mass balance considerations illustrate that the isotope signature of a big pool changes much less than a small pool during a fractionating reaction. Moreover, dissolution reactions occur predominantly at mineral surfaces and the isotope signature of the depleted residue changes only along this reaction front at the surface (Wiederhold et al., 2006). Therefore, the smaller magnitude of Fe isotope fractionation found in this study does not contradict the findings of previous laboratory studies. Our results, however, highlight the difficulty in quantitatively applying experimentally determined fractionation factors to isotope signatures from natural systems. Nevertheless, the performance of laboratory experiments to elucidate specific Fe isotope fractionation factors and mechanisms is certainly required to further understand the behavior of Fe isotopes in nature.

	CONCLUSIONS

TOP NOTES ABSTRACT INTRODUCTION MATERIALS AND METHODS RESULTS AND DISCUSSION CONCLUSIONS REFERENCES

 
The biogeochemical Fe cycle in hydromorphic soils results in intense transformations of Fe minerals within the soil profile. Our results showed that the pedogenic transformation and redistribution processes of Fe in such systems are linked to significant Fe isotope fractionation. The data indicated that the reductive mobilization of Fe(III) from Fe (oxyhydr)oxide minerals under anoxic conditions favors the light Fe isotopes, which is in good agreement with findings from laboratory studies with microorganisms (Beard et al., 1999; Crosby et al., 2005, 2007) and photoreductive dissolution experiments of Fe oxide minerals (Wiederhold et al., 2006). As a result of these mobilization processes, Fe-depleted zones were found to be enriched in heavy Fe isotopes, whereas Fe-enriched zones exhibited enrichments of light Fe isotopes. The strongest fractionation effects were found in the Stagnic Cambisol profile, which may be explained by the fact that this soil has developed during a longer time than the Haplic Gleysol. In addition, the extent of mineral transformation reactions and thereby the associated isotope fractionation effects were probably enhanced by the acidic pH of the Stagnic Cambisol profile. The smaller magnitude of Fe isotope variations between the separated Fe pools in the topsoil of the Haplic Gleysol profile compared with the subsoil horizons has two possible explanations. The first explanation is the younger age of the topsoil material, corresponding to a lesser degree of weathering and pedogenesis, and hence less Fe isotope fractionation. Moreover, the redoximorphic processes, which proceed only in the subsoil, are the main driving force for Fe transformation and redistribution processes in the profile, thereby increasing the extent of Fe isotope variations between the different Fe pools in the soil. A distinct qualitative relationship between Fe concentration changes and Fe isotope fractionation in redoximorphic soils was revealed in our study, indicating a greater mobility of lighter relative to heavier Fe isotopes during pedogenesis. It was not possible, however, to perform a simple quantitative correlation based only on these two parameters. This can be explained by the characteristics of isotope fractionation effects during dissolution reactions, which occur mainly on surfaces and exert a smaller influence on the isotope distribution of the bulk mineral. The application of Fe isotopes as a tracer in natural systems is still restrained by the incomplete understanding of Fe isotope fractionation mechanisms. Nevertheless, the potential of this new isotopic tool for soil environments was clearly demonstrated.

	ACKNOWLEDGMENTS

 
We thank Kurt Barmettler for support in the soil chemistry laboratory, and the staff of the ETH MC-ICPMS laboratory for excellent maintenance and support. This research was funded by ETH Research Grant no. 01927.

	NOTES

TOP NOTES ABSTRACT INTRODUCTION MATERIALS AND METHODS RESULTS AND DISCUSSION CONCLUSIONS REFERENCES

 
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Received for publication November 6, 2006.

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