Pathways of coupled arsenic and iron cycling in high arsenic groundwater of the Hetao basin, Inner Mongolia, China: An iron isotope approach

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Geochimica et Cosmochimica Acta 112 (2013) 130–145 www.elsevier.com/locate/gca

Pathways of coupled arsenic and iron cycling in high arsenic groundwater of the Hetao basin, Inner Mongolia, China: An iron isotope approach Huaming Guo a,b,⇑, Chen Liu b, Hai Lu c, R.B. Wanty d, Jun Wang c, Yinzhu Zhou b a

State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences, Beijing 100083, PR China b School of Water Resources and Environment, China University of Geosciences, Beijing, PR China c The National Institute of Metrology, Beijing 100013, PR China d U.S. Geological Survey, MS 964d Denver Federal Center, Denver, CO 80225, USA Received 20 November 2012; accepted in revised form 25 February 2013; available online 7 March 2013

Abstract High As groundwater is widely distributed all over the world, which has posed a significant health impact on millions of people. Iron isotopes have recently been used to characterize Fe cycling in aqueous environments, but there is no information on Fe isotope characteristics in the groundwater. Since groundwater As behavior is closely associated with Fe cycling in the aquifers, Fe isotope signatures may help to characterize geochemical processes controlling As concentrations of shallow groundwaters. This study provides the first observation of Fe isotope fractionation in high As groundwater and evaluation of Fe cycling and As behaviors in shallow aquifers in terms of Fe isotope signatures. Thirty groundwater samples were taken for chemical and isotopic analysis in the Hetao basin, Inner Mongolia. Thirty-two sediments were sampled as well from shallow aquifers for Fe isotope analysis. Results showed that groundwater was normally enriched in isotopically light Fe with d56Fe values between 3.40& and 0.58& and median of 1.14&, while heavier d56Fe values were observed in the sediments (between 1.10& and 0.75&, median +0.36&). In reducing conditions, groundwaters generally had higher d56Fe values, in comparison with oxic conditions. High As groundwaters, generally occurring in reducing conditions, had high d56Fe values, while low As groundwaters normally had low d56Fe values. Although sediment d56Fe values were generally independent of lithological conditions, a large variation in sediment d56Fe values was observed in the oxidation–reduction transition zone. Three pathways were identified for Fe cycling in shallow groundwater, including dissimilatory reduction of Fe(III) oxides, re-adsorption of Fe(II), and precipitation of pyrite and siderite. Dissimilatory reduction of Fe(III) oxides resulted in light d56Fe values (around 1.0&) and high As concentration (>50 lg/L) in groundwater in anoxic conditions. Re-adsorption of isotopically heavy Fe(II) produced by microbially mediated reduction of Fe(III) oxides led to further enrichment of isotopically light Fe in groundwater (up to 3.4& of d56Fe) in anoxic–suboxic conditions. Arsenic re-adsorption was expected to occur along with Fe(II) re-adsorption, decreasing groundwater As concentrations. In strongly reducing conditions, precipitation of isotopically light Fe-pyrite and/or siderite increased groundwater d56Fe values, reaching +0.58& d56Fe, with a subsequent decrease in As concentrations via co-precipitation. The mixed effect of those pathways would regulate As and Fe cycling in most groundwaters. Ó 2013 Elsevier Ltd. All rights reserved.

1. INTRODUCTION ⇑ Corresponding author at: School of Water Resources and

Environment, China University of Geosciences, Beijing 100083, PR China. Tel.: +86 10 8232 1366; fax: +86 10 8232 1081. E-mail address: [email protected] (H. Guo). 0016-7037/$ - see front matter Ó 2013 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.gca.2013.02.031

High As groundwater is widespread in many countries, including Argentina, Bangladesh, Cambodia, China, Hungary, India, Mexico, Nepal, Pakistan, Romania, Vietnam, and USA (Ravenscroft et al., 2009). Both oxic and reducing

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aquifers host high As concentrations (Smedley and Kinniburgh, 2002). High As concentrations in oxic groundwater have been attributed to both oxidation of pyrite and desorption of Fe oxyhydroxide-adsorbed As (Baker et al., 1998; Peters et al., 1999; Smedley and Kinniburgh, 2002; Jones and Pichler, 2007; Pili et al., 2012). In anoxic groundwaters, dissimilatory reduction of Fe/Mn oxide minerals is believed to be the major mechanism for As mobilization (Nickson et al., 1998; Stu¨ben et al., 2003; Islam et al., 2004; McArthur et al., 2004; Charlet and Polya, 2006). In either scenario, Fe is a key element for As cycling, although As(V) reduction, independent of Fe cycling, has also been shown to mobilize As. In comparison with oxic conditions, high As groundwater related to reducing conditions is more common in South and Southeast Asia (Ravenscroft et al., 2009).These high As groundwaters typically also have high Fe concentrations, ranging between 0.01 and 29 mg/L in Bangladesh (Nickson et al., 1998), 3–13.7 mg/L in West Bengal (Mukherjee et al., 2009), <0.05–44.3 mg/L in Vietnam (Berg et al., 2008), <0.1–26.5 mg/L in Cambodia (Rowland et al., 2008), and 0.01–5.9 mg/L in Inner Mongolia (Guo et al., 2008a). On the other hand, in these anoxic systems, Fe levels as low as 0.1 mg/L have been detected in many places, which is often but not always associated with low As. Although reductive dissolution of Fe-oxides has been proposed to be the major mechanism for As release, low As has been found in both high and low Fe groundwater, and vice-versa (Smedley et al., 2003; Berg et al., 2008; Guo et al., 2008a; Rowland et al., 2008; van Geen et al., 2008; Mukherjee et al., 2009; Kim et al., 2012). Other processes including pyrite and siderite precipitation (Lowers et al., 2007; Reza et al., 2010; Wang et al., 2012) and Fe(II) adsorption (Ligier et al., 1999; Appelo et al., 2002; van Geen et al., 2004) have been used to explain this poor correlation. However, characteristics of Fe isotopes in these groundwater systems are still unknown. It remains unclear whether Fe isotopes can be useful for revealing mechanisms of As mobilization in aquifers. Iron isotope systematics can help constrain the pathways and mechanisms of Fe redox transformations (Beard et al., 2003a). Terrestrial igneous rocks are very homogenous in isotopic composition with d56Fe of 0.00 ± 0.05& (relative to the UW-Madison average igneous standard; Beard et al., 2003b). Reductive dissolution of Fe(III) oxides, adsorption and re-adsorption of Fe(II) are proposed to result in 56Fe depletion in residual dissolved Fe(II), while Fe(II) mineral precipitation evidently has the opposite effect and enriches 56Fe in residual dissolved Fe(II). Fractionation of Fe isotopes is closely related to dissimilatory Fe(III) reduction, in which the produced Fe(II) is isotopically lighter (Brantley et al., 2004; Crosby et al., 2005, 2007; Emmanuel et al., 2005; Wu et al., 2009; Tangalos et al., 2010; Percak-Dennett et al., 2011). In addition, microbial Fe(II) oxidation led to positive Fe isotope fractionations between Fe(III) oxide/hydroxide precipitates and aqueous Fe(II) (Balci et al., 2006; Kappler et al., 2010). In addition, heavier Fe(II) is preferentially adsorbed onto Fe(III) (hydro)oxides surfaces, leaving isotopically light Fe(II) in the solution (Skulan et al., 2002; Welch et al.,

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2003; Icopini et al., 2004; Teutsch et al., 2005; Beard et al., 2010; Wu et al., 2010, 2011, 2012a). Precipitation of Fe minerals (i.e. ferrihydrite, mackinawite, pyrite and siderite) also may result in Fe isotopic fractionation, with isotopically light Fe reacting fastest and becoming enriched in the solid phases (Wiesli et al., 2004; Butler et al., 2005; Clayton et al., 2005; Severmann et al., 2006; Guilbaud et al., 2011a,b; Sivan et al., 2011). Therefore, characteristics of Fe isotopes should aid in understanding biogeochemical and geochemical processes of Fe and As cycling in high As groundwater systems. The objectives of this study are to (1) characterize Fe isotopes in high As groundwaters and aquifer sediments, (2) evaluate the linkage between Fe isotope and redox-sensitive components, and (3) provide new insight into biogeochemical processes in high As groundwaters. 2. HYDROGEOLOGICAL SETTING The study area is located north of the Yellow River, east of the Wuranbuh Desert, and south of the Langshan Mountains (Fig. 1a). The Langshan Mountains are mainly composed of a metamorphic complex (slate, gneiss and marble) generally of Jurassic to Cretaceous age, which is folded and fractured. Continuous tectonic subsidence of the basin and uplift of the mountains has led to thick accumulation of sediments; the basin thickness of Quaternary sediments ranges between 400 and 4000 m. The aquifer dips towards the northwest. Aquifer sediments were composed of alluvial-pluvial sand, sandy silt, lacustrine and fluviallacustrine sandy silt, silty clay and clay rich in organic matter in the central part of the basin, fluvial sand-fine sand on river banks, and alluvial sand in fan areas. Lacustrine sediments sporadically occur at depths between 5–40 m below land surface (BLS) in the flat plain. The climate is semiarid-arid, with average annual precipitation of 130–220 mm (mainly during July to September) and annual evaporation rates of about 2000–2500 mm. Average annual temperatures range from 5.6 to 7.8 °C (Guo et al., 2008a). Groundwater is recharged by downward infiltration of meteoric water in the basin and lateral flow of water from fractures in marble, slate and gneiss along the mountain front, as well as some leaked water from natural lakes and drainage canals, and irrigation return flow from farmland, and discharged mainly via evapotranspiration and artificial abstraction (Guo et al., 2008a). The general direction of groundwater flow in the northern part is from north to south, while that in the southern part is from south to north, but the flow rate is very slow. To investigate variations in chemical and isotopic compositions of groundwater along the flow path, a typical hydrogeological unit was selected (Fig. 1c and d), where the general direction of groundwater flow is from northwest to southeast. Measurement of groundwater levels along with an elevation survey of wells confirmed the groundwater flow direction, which was consistent with the report by Zhang et al. (2012). Hydraulic conductivity decreases by a factor of 10 along the flow path ranging from 20.0 m/d along the mountain front to 2.2 m/d in the downdip region (Guo et al., 2010). Shallow groundwaters mainly occur

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Fig. 1. The study area (a), locations of sampling wells and drilling sites (b), sampling sites approximately along the flow path (c), and a profile of hydraulic heads approximately along the flow path (d).

at depths between 5 and 10 m below land surface (BLS) in alluvial sediments in the recharge area (around Well Nos. 46 and 47) and between 1.5 and 2.0 m in alluvial-lacustrine sediments in the discharge area (between Well No. 32 and Well No. 28). The flow rate is generally slow, but a little faster in the alluvial fan than in the flat plain. There is a groundwater drainage canal, which is used to lower the groundwater table and reduce groundwater evaporation discharge (Guo et al., 2011c), but which has only a local effect on the groundwater flow system in the unit (Fig. 1d). The canal water recharges shallow groundwater during the rainy season when there is a high water level in the canal, but the recharge does not occur during dry seasons. The groundwater table is very shallow in the flat plain (averagely 2.0 m BLS), especially immediately down-gradient of the canal (around Well Nos. 32 and 38), and strong evaporation occurs resulting in increased soil salinity. In this region, aeration takes place due to good permeability of sediments. 3. METHODS AND MATERIALS According to our previous investigations (Guo et al., 2008a, 2010, 2011c), thirty representative groundwater samples were collected from hand-pumped tube wells and electric-powered public water supply wells (including 11 samples in July 2009 and 19 samples in May 2010) for

chemical and isotopic analysis. Sampling locations are shown in Fig. 1b. Among 30 samples, nine are located approximately along a groundwater flow path near the Langshan Mountains (Fig. 1c and d). The depths of these wells ranged between 9.0 and 100 m. Prior to groundwater sampling, wells were pumped for more than 20 min until water temperature, electrical conductivity (EC), pH, and Eh were stable. All samples were filtered through 0.20 lm membrane filters in the field. Water samples for major, trace element and Fe isotope analysis were collected in 100 mL HNO3-washed polyethylene bottles, followed by addition of 6 M ultrapure HNO3 to pH <2. Samples for analysis of As species were preserved with 0.25 M EDTA. Samples for anion analysis were not acidified. At the time of groundwater sampling, parameters including water temperature, EC, pH, and Eh, were measured using a multiparameter portable meter (HANNA, HI 9828), which was calibrated using standard solutions before use. Concentrations of S2 and Fe(II) were measured using a portable spectrophotometer (HACH, DR2800) with methylene blue and 1, 10 phenanthroline methods, respectively. Alkalinity was measured using a Model 16900 digital titrator (HACH) using bromocresol green–methyl red indicator. The Eh readings are reported relative to the standard hydrogen electrode. Concentrations of major cations and trace elements were determined by ICP-AES and ICP-MS, respectively. Unaci-

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dified aliquots were analyzed for F, Cl, NO3, and SO42 concentrations by Ion Chromatography with a Dionex DX120. Arsenic species in groundwater samples were analyzed by HPLC-ICP-MS (Guo et al., 2011a). Thirty-two sediment samples were taken from two boreholes at two sites (SH09-1 and SH09-2) (Fig. 1) at depths up to 22 m BLS. Immediately after extraction from the borehole, sediments were put into sealed N2-filled plastic bags and preserved in anaerobic boxes (Biomerieux, France) with an O2 adsorbent (AnaeroPack, Mitsubishi) and an anaerobic indicator (Oxoid, England), and transported to the laboratory at 4 °C. They were stored at 20 °C in the laboratory. Sediment subsamples were dried and disaggregated, and milled for XRD analysis. Wet sediments were also leached in 1.2 mol/L HCl at 80 °C for 30 min to determine the Fe(II)/FeTot ratio of the leachable Fe fraction (Horneman et al., 2004). Iron(II) in the leachate was determined by colorimetry with ferrozine. Total Fe concentrations were analyzed by means of ICP-AES. Iron isotopes were analyzed for groundwater samples and sediment samples. Sediment samples were completely dissolved (total digestion) in SavillexÒ containers in ultrapure concentrated HCl, HF and 6 M HClO4 (Matthews et al., 2004). Solutions for Fe isotope analysis were pretreated according to Borrok et al. (2007). Briefly, columns were washed in 7 M HNO3, fitted with a frit (0.70 lm pores), and loaded with 0.65 mL of precleaned, 100– 200 mesh AG MP-1 anion-exchange resin (Bio-RadÒ) under 10 M HCl, to separate Fe from other matrix elements. The evaporated sample (3 lg of Fe dissolved in 1 mL of 10 M HCl) was loaded onto the column. After that, 4 mL of 10 M HCl was passed through the column to elute most matrix cations and anions, which was followed by an additional 8 mL of 5 M HCl to elute the Cu fraction. Finally, 4 mL of 1 M HCl was loaded to rapidly elute Fe from the column. The Fe fraction was evaporated, dissolved in 2% HNO3, and then re-evaporated and re-dissolved in 2% HNO3 to drive off chlorine prior to Fe isotope analysis. Iron isotope analysis was performed at the isotope laboratories of the National Institute of Metrology (China) using IsoProbe MC-ICP-MS (GV Instruments), a single focusing instrument with a sector-style magnet, which includes a collision cell. A collision gas (Ar), flowing into the cell at a flow rate of 1.0 mL/min, thermalized the ion beam, reducing the ion energy spread to 1 eV, and causing the mass analyzer to produce flat-topped peaks. Another reactive collision gas (H2) was used to eliminate or greatly reduce molecular isobars such as ArN. Plasma and collision cell gases were optimized to completely remove argide interferences on the Fe mass spectrum. Iron solutions were introduced in a weak HNO3 acid matrix (2%) using a self-aspirating, low-inflow desolvating nebulizer. Iron concentrations in the samples were approximately 800 lg/L. Isotopes of masses 54, 56, and 57 were simultaneously measured in a hard extraction mode. Each sample was bracketed by concentration-matched IRMM14 Fe standard solution, and the isotopic results are expressed in standard delta notation (&) relative to the average of the bracketing standards, IRMM14ave, as shown in Eq. (1) (Borrok et al., 2007).

# ð56 Fe=54 FeÞsample d Fe ¼ 56  1  1000 ð Fe=54 FeÞIRMM14ave

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"

56

ð1Þ

The validity of the standard-sample bracketing approach was confirmed through mass bias tests similar to those described by Beard et al. (2003a). We doped “Fe free” groundwaters with the IRMM-14 Fe standard and then ran the samples through the column separation process to evaluate both the accuracy and precision of the isotopic measurements. The results show that the doped samples had d56Fe values of 0 ± 0.06& (2r, n = 12). Moreover, a plot of measured d56Fe versus d57Fe shows a good correlation (r2 = 0.96) with a slope of 1.5 (S2 in Supporting Materials), reflecting the expected mass-dependent fractionation. Replicate mass spectrometric analyses were made in different sessions (days). The average standard deviation (2r) calculated from replicate measurements of unknowns over multiple analytical sessions was 0.06& in this study. 4. RESULTS 4.1. Groundwater chemistry Groundwater pH ranged from 7.1 to 8.9 (median 8.2), indicating a neutral-weak alkaline environment. HCO3 and Cl were the major anions, with ranges between 191 and 1290 mg/L (median 531 mg/L) and between 35 and 1240 mg/L (median 351 mg/L), respectively. Sodium was the major cation, generally accounting for >40% of total meq concentration. Most groundwater samples had high salinity, with a range of electrical conductivity (EC) between 0.53 and 4.16 mS/cm (median 1.82 mS/cm) (S1 in Supporting Materials). Fluoride concentration ranged between 0.35 and 2.2 mg/L, with 23% of samples exceeding the WHO guideline of 1.5 mg/L. Redox potentials (Eh) ranged from 35 to 290 mV (median 77 mV). Nitrate concentrations were mostly less than 5.0 mg/L. There were four groundwater samples having NO3 concentrations between 6 and 20 mg/L, which was possibly due to pollution from agricultural sources. Groundwater had moderate NH4–N concentration, ranging between <0.1 and 12.5 mg/L (median 2.3 mg/L). Concentration of sulfide ranged between <1.0 and 20 lg/L. Sulfate concentrations in the collected samples span a large range from <0.01 to 708 mg/L. The occurrence of both S2 and SO42 was observed in most groundwater samples. Dissolved Fe concentration had a range from 0.03 to 3.02 mg/L. Ferrous Fe was detected in most samples (median 0.24 mg/L). Concentrations of dissolved organic carbon (DOC) were quite high, ranging between 1.3 and 44 mg/L (median 7.5 mg/L). Arsenic concentrations ranged between 0.4 and 720 lg/ L (median 48.7 lg/L). Among 30 samples, 22 contained >10 lg/L As, which is the drinking water guideline in China, European countries and the USA (Guo et al., 2011b). No organic As species were detected. Inorganic As(III) was the major As species, generally amounting to >70% of total As. There were negative correlations between As concentration and Eh, SO42, and NO3 (Fig. 2a–c). In groundwaters with Eh values > 150 mV, As concentrations

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Fig. 2. Variation of As concentrations with Eh (a), SO42 (b), NO3 (c), S2 (d), Fe(II) (e), and dissolved Fe (f) of groundwater samples.

were less than 10 lg/L. High As concentrations (>50 lg/L) were found in groundwaters with Eh < 150 mV, suggesting that As mobilization occurred in reducing conditions. In high As groundwaters, both SO42 and NO3 were lower than those in low As groundwaters. However, S2 concen-

tration generally showed a positive correlation with As concentration (Fig. 2d). Negative correlations between As and Fe(II) or FeTot were observed in studied groundwaters (Fig. 2e and f). Generally, groundwaters with Fe(II) > 0.8 mg/L contained As

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concentrations of <100 lg/L. In groundwaters with Fe(II) < 0.8 mg/L, As concentrations were normally <100 lg/L at S2 < 10 lg/L, and >100 lg/L at S2 > 10 lg/L. Saturation indices (SI), calculated using the hydrogeochemical code PHREEQC (Parkhurst and Appelo, 1999), showed that all groundwaters were oversaturated with respect to pyrite and mostly saturated with respect to siderite and mackinawite (data not shown).

4.2. Iron isotopes in groundwaters It has been known that terrestrial igneous rocks are very homogenous in isotopic composition with d56Fe of 0.00 ± 0.05& (relative to the UW-Madison average igneous standard; Beard et al., 2003b). In our study area, groundwater had d56Fe values between 3.40& and 0.58& with a median of 1.14&, and d57Fe values between 4.95& and 0.77& with a median of 1.74&. With the exception of three samples, all the dissolved d56Fe values are less than 0.0&, showing that most groundwater samples were depleted in 56Fe.

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In reducing conditions, groundwaters generally had higher d56Fe values, in comparison with moderate oxic conditions (Fig. 3a). In groundwaters with Eh values 50 mV, d56Fe values were between 1.50& and +0.58&, which are higher than that of groundwater (around 3.40&) with Eh of 270 mV. High As groundwaters, generally occurring in reducing conditions, had high d56Fe values (Fig. 3b), while low As groundwaters normally had low d56Fe values. A positive correlation was observed between d56Fe values and S2 concentrations (Fig. 3c), so that groundwaters with low S2 concentration generally had low d56Fe values. Three end-members were observed in the diagram of d56Fe versus dissolved Fe (Fig. 3d), one with low Fe concentrations and low d56Fe values in the lower left (EM1), one with high Fe concentrations and moderate d56Fe values in the upper right (EM2), and another with low Fe concentrations and high d56Fe values in the upper left (EM3). Groundwaters in EM1 had Eh values between 100 and 290 mV, in EM2 between 50 and 60 mV, and in EM3 between 35 and 50 mV. Those samples with Eh between 60 and 100 mV plot within the triangle with the end-members as corners, which may indicate mixing of those endmembers.

Fig. 3. Variation of d56Fe values with Eh (a), As (b), S2 (c), and dissolved Fe (d) of groundwater samples.

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Fig. 4. Plot of d56Fe versus dissolved Fe for illustrating pathways for Fe cycling in the groundwaters.

In groundwaters with Fe concentrations less than 1.0 mg/L and d56Fe values greater than 1.50&, As concentrations were mostly higher than 50 lg/L (Fig. 4). However, groundwaters depleted in 56Fe (with d56Fe values between 1.50& and 3.50&) generally had low to the intermediate As levels (0.4–50 lg/L). Groundwaters in EM1 had low As concentrations (<50 lg/L), while in EM3 As was mostly >100 lg/L. In EM2, As concentrations

were higher in groundwater samples with higher Fe concentrations (Fig. 4). 4.3. Iron isotopes in sediments Sediments had d56Fe values between 1.10& and 0.75& with a median of 0.36&. In SH09-1, sediment d56Fe values showed a relatively greater variation (between 1.10& and

Fig. 5. Depth profiles of color and lithologic sketches, bulk Fe content, 1.2 mol/L hot HCl extracted sediment Fe(II)/Fe ratio, and sediment d56Fe and porewater d56Fe in SH09-1. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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Fig. 6. Depth profiles of color and lithologic sketches, bulk Fe content, 1.2 mol/L hot HCl extracted sediment Fe(II)/Fe ratio, and sediment d56Fe and porewater d56Fe in SH09-2. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

0.75& with a median of 0.34&) in comparison with sediments from SH09-2 (between 0.16& and 0.57& with a median of 0.44&). XRD did not quantitatively detect the presence of Fe minerals (e.g., iron (hydr)oxides and iron sulfides) in most sediment samples (data not shown). Hot HCl leaching tests showed that yellowish brown sediments had lower Fe(II)/FeTot ratios (0.40–0.45) than grey sediments (0.55–0.60) (Figs. 5 and 6). Although Fe isotope fractionation between aqueous Fe(II) and ferric oxide/ hydroxides is closely related to Fe minerals (Beard et al., 2010; Wu et al., 2010, 2011, 2012a), sediment d56Fe values were generally independent of lithological conditions in this study. In SH09-1, although sediments changed from yellowish silt to yellowish brown clay at depths between 1.0 and 5.6 m BLS, d56Fe values were mostly around 0.40 ± 0.10& (Fig. 5). Changes in lithologic characteristics from yellowish brown clay, through yellowish brown silty fine sand, to yellowish brown fine sand, did not lead to a significant variation in d56Fe values (0.50 ± 0.07&) at depths between 5.0 and 9.0 m BLS in SH09-2 (Fig. 6). In the oxidation–reduction transition zone, a big variation in sediment d56Fe values was observed in SH09-1. This zone occurred between yellowish brown silt and light grey silty fine sand at a depth around 9.0 m, where d56Fe values showed a decreasing trend and Fe(II)/FeTot ratio an increasing trend, although bulk Fe content kept around 2.0% (Fig. 5). At a depth of 10 m BLS, the sediment (light grey silty fine sand) had the lowest d56Fe value (1.10&). Below this zone, d56Fe values increased to 0.75& at a depth of 11.5 m BLS, and mostly kept around 0.40 ± 0.20& in grey fine sand at depths between 14.5 and 21.1 m BLS. Three porewater samples were taken from aquifer sediments, including two at depths of 10 and 20 m BLS of SH09-1, and one at a depth of 10.6 m BLS of SH09-2. At the depth of 10 m of SH09-1, porewater d56Fe value was

0.70&, being higher than that of the sediment (1.10&) (Fig. 5). However, the other two porewater samples had lower d56Fe values than the host sediments (Figs. 5 and 6). 5. DISCUSSION 5.1. Implications for dissolved Fe sources Groundwater d56Fe values generally decreased with increasing Eh values (Fig. 3a), while less variation in sediment d56Fe values was observed. This may mirror the differences in the amounts of sediment Fe and dissolved Fe. If the porosity is 20%, a 1-L volume of saturated aquifer material would contain about 0.2 mg dissolved Fe, while the sediment would contain about 80 g Fe, which is more than 5 orders of magnitude more Fe in the solid than in the water. Therefore, in the water–rock systems, Fe-oxide isotopic composition in sediments would change little during partial Fe-oxide reduction, whereas dissolved Fe isotopic composition could be largely impacted by small amounts of reductive dissolution (Roy et al., 2012). The average groundwater d56Fe value was 1.56& lighter than the average sediment d56Fe value. Both Fe-oxide reduction and oxidative precipitation of dissolved Fe can produce lighter groundwater d56Fe values (e.g., Bullen et al., 2001; Johnson et al., 2002, 2004, 2008; Skulan et al., 2002; Beard et al., 2003a; Brantley et al., 2004; Croal et al., 2004). High As groundwaters were generally anoxic and contained no NO3 (Fig. 2), and therefore prevented oxidation of dissolved Fe(II) in high As groundwater aquifers. Therefore, the observed isotopically light d56Fe values likely resulted from reductive Fe-oxide dissolution. Mineral dissolution in the presence of Fe-reducing bacteria preferentially released light Fe from silicates and

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Fe-oxides (Brantley et al., 2001, 2004; Crosby et al., 2005, 2007; Wu et al., 2009; Tangalos et al., 2010; Percak-Dennett et al., 2011). Teutsch et al. (2005) found a slight fractionation (around 0.3& between dissolved Fe and sedimentary Fe) with microbially mediated reduction of Fe oxyhydroxides in an anoxic aquifer. Experimental studies have shown that Fe isotope fractionation during dissimilatory iron reduction (DIR) was related to reduction rates (Beard et al., 2003a; Icopini et al., 2004; Johnson et al., 2004). Faster reduction rates produced greater Fe isotope fractionations, with rapid formation and adsorption of Fe(II) onto Fe(III) oxides producing fractionations as large as 2.3&. In less extreme cases, Fe isotope fractionation during DIR around 1.0& to 1.5& between biogenic Fe(II) and Fe(III) oxides was more representative (Beard et al., 2003a; Icopini et al., 2004; Rouxel et al., 2008). Furthermore, the low d56Fe values for aqueous Fe(II) produced by microbial dissimilatory iron reduction (DIR) is controlled by coupled electron and Fe atom exchange among three Fe inventories: aqueous Fe(II), sorbed Fe(II), and a reactive Fe(III) component at iron oxide surfaces (Crosby et al., 2005, 2007). Although the presence of Si altered Fe isotope fractionation factors between aqueous Fe(II) or sorbed Fe(II) and reactive Fe(III) at the hematite surfaces, DIR produced aqueous Fe(II) with negative d56Fe values under conditions of variable pH and dissolved Si (Wu et al., 2009). In the case of amorphous Fe(III) oxide–silica coprecipitate, DIR produced large quantities of low d56Fe Fe(II) (aqueous and solid phase), and led to rapid and near-complete isotope exchange among all Fe(II) and Fe(III) components (Percak-Dennett et al., 2011). Based on sediment core analyses and incubation experiments, Tangalos et al. (2010) observed that partial reduction of amorphous Fe(III) by DIR produced isotopically light (d56Fe  1.5& to 0.5&) aqueous Fe(II). An average Fe isotope fractionation around 1.5& was expected to represent the predominance of dissimilatory reduction of Fe(III) oxides in the studied aquifers (Pathway 1 in Fig. 4). High DOC concentrations of studied groundwater (between 1.3 and 44 mg/L, median 7.5 mg/L) would also support the domination of microbial Fe-oxide reduction. The isotopically light groundwater d56Fe values would be associated with OC diagenesis. Hot HCl leaching tests showed that more than 30% Fe would be present as Fe(III) in the studied sediments, which may indicate that Fe(III) was partially reduced. Although groundwater Fe isotopic composition may reflect that of the sediment if the sediment bulk Fe can mostly be reductively dissolved, partial dissimilatory reduction of Fe-oxides was expected to lead to Fe fractionation around 1.5& between dissolved Fe in groundwaters and sedimentary Fe in aquifer solids. 5.2. Geochemical controls on dissolved Fe isotopic compositions 5.2.1. Adsorption processes Of 30 groundwater samples analyzed, ten had d56Fe values less than 1.50&, which was believed to be the lower limit for representative Fe isotope fractionation during

DIR (Beard et al., 2003a; Icopini et al., 2004; Rouxel et al., 2008). Therefore, DIR processes alone would be unlikely to produce d56Fe values as low as 3.40& in groundwaters with low Fe concentrations. Groundwaters with low d56Fe values (<1.50&) also had low Fe concentrations (<1.0 mg/L) (Fig. 4). The simultaneous decreases in dissolved Fe concentration and d56Fe values can result from adsorption of dissolved Fe onto Fe-oxides and/or oxidative Fe(II) precipitation, although we cannot rule out other mechanisms. Hot HCl leaching tests showed that although they were present in reducing conditions, Fe(III), possibly as Fe-oxides, remained in the grey sediments (>30% of total Fe). In SH09-1, sediments had Fe(II)/FeTot ratios of 0.52 and 0.60 at depths of 10 and 20 m, with corresponding porewater Fe(II) concentrations of 1.65 and 4.27 mg/L, respectively. In SH09-2, sediments had a Fe(II)/FeTot ratio of 0.42 at a depth of 10 m, with a corresponding porewater Fe(II) concentration of 0.30 mg/L. This result indicates that the higher porewater Fe(II) concentration was associated with the higher Fe(II)/FeTot ratio in the sediments. These data suggested that Fe(II) would be re-adsorbed in the aquifer systems after being reduced under moderately reducing conditions (Ligier et al., 1999; Appelo et al., 2002; van Geen et al., 2004). In groundwater samples with Eh values between 80 and 150 mV, adsorption of dissolved Fe(II) onto residual Fe-oxides would be the most likely process responsible for Fe isotope fractionation. The strong affinity of Fe(II) for Fe oxyhydroxide surfaces has been well documented (Appelo et al., 2002; van Geen et al., 2004; Handler et al., 2009). It has been experimentally demonstrated that around one-third of dissolved Fe(II) produced by microbial Feoxide reduction was re-adsorbed onto the mineral-surface of sedimentary Fe(III) oxides (Icopini et al., 2004). This re-adsorption resulted in a decrease in dissolved d56Fe values by 0.7–1.2& relative to d56Fe of solid Fe(III) oxide (Icopini et al., 2004; Crosby et al., 2005, 2007). A two-step process, combining Fe-oxide reduction and dissolved Fe(II) re-adsorption (shown as Pathway 1 and Pathway 2 in Fig. 4, respectively), is typical of dissimilatory Fe-oxide reduction. On the one hand, Fe isotope fractionation between aqueous Fe(II) and iron oxides was related to the iron minerals, with fractionation factors between aqueous Fe(II) and hematite of 3.16& (Skulan et al., 2002; Welch et al., 2003; Wu et al., 2010), between aqueous Fe(II) and goethite of 1.05& (Beard et al., 2010), between aqueous Fe(II) and amorphous Fe(III) oxide–Si coprecipitates of 2.58& to 3.99& depending on Fe:Si molar ratios (Wu et al., 2011, 2012a). On the other hand, Fe fractionation may occur between aqueous Fe(II) and adsorbed Fe(II), and between adsorbed Fe(II) and surface Fe(III) due to electron and atom exchange during adsorption of Fe(II) on Fe-oxide minerals, although the lighter 56Fe/54Fe ratios of aqueous Fe(II) were usually observed in comparison with that of bulk Fe-oxides (Handler et al., 2009). In the case of goethite, the most common in lacustrine sediments, there was a kinetic fractionation associated with electron transfer into the goethite structure, which produced an Fe(II) material with a d56Fe value higher than that of aqueous Fe(II), in addition to adsorption of isotopically heavy

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Fe(II) (Jang et al., 2008). It was experimentally demonstrated that aqueous Fe(II) was 1.05& lighter in d56Fe than coexisting goethite at 22 °C, and 1.24& lighter than the adsorbed Fe(II) (Beard et al., 2010). Iron fractionation between aqueous Fe(II) and surface Fe(III) was found to be between 1.63& and 2.62& (Crosby et al., 2007; Beard et al., 2010). The isotopic fractionation due to exchange depends on the molar ratios of aqueous Fe(II) and Fe(III) oxide/hydroxide, with the largest shifts in Fe isotope compositions of the aqueous Fe(II) at the low ratios (Beard et al., 2010). In a similar field site, Teutsch et al. (2005) reported strong Fe isotope fractionation (up to 3&) in groundwater Fe(II) resulting from rapid adsorption of Fe(II) on Fe-oxyhydroxides formed during injection of O2-containing water. By either of these mechanisms, the combination of Fe-oxide reduction and dissolved Fe(II) re-adsorption contributes to the lower left end-member (EM1) in Fig. 3d. Samples with Eh values >270 mV and d56Fe value <3.05& showed that oxidative Fe(II) precipitation may predominantly regulate Fe isotope fractionation. Assuming that oxidation of dissolved Fe(II) and precipitation of Fe(III) oxyhydroxide were associated with moderately oxic conditions, the low d56Fe value may result from the precipitation of isotopically heavy Fe oxyhydroxides. Rouxel et al. (2008) also found that strongly negative values of Fe isotopes (d56Fe values down to 5.0&) were mostly the result of the oxidative pathways of the Fe cycle and the sequestration of heavy Fe isotopes in Fe-oxides. 5.2.2. Fe isotope fractionation due to Fe-sulfide precipitation Most groundwater samples were near saturation with respect to mackinawite and greatly oversaturated with respect to pyrite (Fig. 7a and b). Precipitation of Fe-sulfides likely occurred in the aquifers, which is indicated also by the negative correlation between S2 and dissolved Fe in groundwaters. Generally, groundwaters with higher SImackinawite and SIpyrite values had heavier d56Fe values (Fig. 7a and b). Precipitation of mackinawite (FeSm) produces Fe fractionation between solid and solution. Butler et al. (2005) showed that precipitation of FeSm led to 0.9& (0.6& to 1.6&) of d56Fe value in the solid phase relative to the dissolved phase at a pH of 4, which was believed to be a kinetic fractionation (Guilbaud et al., 2010). However, Wu et al. (2012b) showed that at neutral pH, isotope ex-

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change was much more rapid and resulted in equilibrium fractionation between aqueous Fe(II) and FeSm, in which FeSm had d56Fe value greater than the aqueous Fe(II). Iron fractionation during FeSm formation would depend on the degree of FeSm precipitation, the degree of isotope exchange during equilibrium, the availability of S2, the solution pH, and the role of aqueous Fe-S speciation (Guilbaud et al., 2011b; Wu et al., 2012b). Therefore, it seems unlikely that FeSm precipitation would produce the high d56Fe values in the studied groundwater with neutral-weak alkaline pH. The heaviest d56Fe value (+0.58&) was found in sample No. 39 with SI pyrite = 19. The increase in groundwater d56Fe values could have resulted from preferential enrichment of light Fe isotopes in pyrite (Guilbaud et al., 2011a). In modern anoxic basins, diagenetic pyrite displayed isotopic compositions between 0.4& and 1.2& (Severmann et al., 2006), which were comparable to d56Fe values of synthetic pyrite at 40 and 100 °C at pH = 6 (between 0.28 and 1.75&) (Guilbaud et al., 2011a). Sivan et al. (2011) observed that FeS2 precipitation resulted in porewater with a d56Fe value about +1.0& in Lake Kinneret, which is slightly heavier than groundwater d56Fe values in EM3 of the Hetao basin. Although we were unable to collect discrete sedimentary pyrite to show light-isotope enrichment there, the proposed Pathway 3 in Fig. 4 would model the reaction described here. The result would be enrichment of the heavy isotopes of Fe remaining in the groundwater, with lower Fe concentrations, at the lowest Eh values (<80 mV; also see Archer and Vance, 2006). This fractionation pathway also was supported by limited data of both aquifer sediment d56Fe and related groundwater d56Fe in drilling boreholes of this study. Those results showed that the maximum difference between groundwater and sediment d56Fe values (D56Fegroundwater–sediment) was around +0.39& in the grey reducing sediment (at a depth of 10 m BLS of SH09-1), which is higher than a D56Fegroundwater–sediment of 2.70& in the yellow sediment (at a depth of 10 m BLS of SH09-2) (Figs. 5 and 6). In addition, around 30% of studied groundwaters were near saturation with respect to siderite (Fig. 7c). The precipitation of siderite could occur in groundwater systems. It has been experimentally demonstrated that isotopically light Fe is preferentially enriched in abiotic siderite, resulting in Fe isotope fractionation between Fe(II)aq and siderite

Fig. 7. Variation of d56Fe values with SImackinawite (a), SIpyrite (b) and SIsiderite (c).

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of +0.26& to 0.70& (Wiesli et al., 2004). In biological systems, experiments showed that d56Fe values of Fe(II)aq produced from dissimilatory Fe(III) reduction and subsequent precipitation of siderite was around 0& relative to siderite under equilibrium conditions (Johnson et al., 2005). Either way, moderately negative d56Fe values were generally observed in many siderites of Archean and Proterozoic sedimentary rocks (Matthews et al., 2004; Yamaguchi et al., 2005; Heimann et al., 2010). Therefore, precipitation of siderite in the studied groundwater also would result in the increase in d56Fe values with decreasing dissolved Fe concentration (Fig. 4). Groundwaters were mostly oversaturated with respect to magnetite (data not shown). However, there is no solid observation on Fe isotope fractionation during magnetite formation. In long-term experiments at low Fe(III) reduction rates, Johnson et al. (2005) found that the aqueous Fe(II)–magnetite fractionation is 1.3&, whereas no detectable fractionation of Fe isotopes in the bacterial magnetites was observed by Mandernack et al. (1999). Therefore, we do not propose the contribution of magnetite precipitation to the high d56Fe values of the studied groundwaters in EM3.

5.3. Implications for geochemical processes controlling As concentrations 5.3.1. Controls on As concentrations Characteristics of Fe isotope fractionation suggested three pathways of Fe cycling in the aquifers of the Hetao basin: microbially mediated reduction of Fe oxyhydroxides; Fe(II) re-adsorption after reduction; and pyrite and siderite precipitation (Fig. 4). Those pathways were closely associated with As cycling in groundwater systems. Although Fe isotope fractionation showed the occurrence of dissimilatory reduction of Fe(III) oxides in reducing conditions of the Hetao aquifers, roles of this reduction in Fe and As cycling varied (Fig. 4). Groundwater samples near Pathway 1 were dominated by microbial Fe(III) reduction (Fig. 4). Along this pathway, groundwater mostly had As concentrations greater than 50 lg/L, suggesting that the high As concentration was directly related to microbial Fe(III) reduction. It has been well-known that reductive dissolution of Fe(III) oxides may trigger As mobilization in reducing aquifers (Nickson et al., 1998; Islam et al., 2004; Charlet and Polya, 2006; Guo et al., 2008b). This is the first investigation revealing microbially mediated reduction of Fe oxyhydroxides as the major mechanisms for As mobilization using Fe isotope fractionation. Pathway 2 in Fig. 4 shows the effects of re-adsorption of Fe(II) produced by Fe(III) oxide reduction. This re-adsorption requires the presence of remnant Fe(III) oxides in the aquifer materials. In this system, As being released from Fe-oxides via reductive dissolution would be re-adsorbed onto the residual Fe(III) oxides along with Fe(II). It has been well documented that both Fe(III) oxides and Fe(II)–Fe(III) components are good scavengers for dissolved As (Sun and Doner, 1998; Jessen et al., 2005; Cundy et al., 2008; Hohmann et al., 2010; Guo et al., 2013). There-

fore, low As groundwater (As < 50 lg/L) was found along Pathway 2. Pathway 3 in Fig. 4 represents pyrite and siderite precipitation. Although reductive dissolution of Fe-oxides has been well-known to be the cause of As mobilization in reducing conditions (Nickson et al., 1998; Islam et al., 2004; McArthur et al., 2004; Charlet and Polya, 2006), the fact that As concentrations were normally less than 100 lg/L at S2 < 10 lg/L, and greater than 100 lg/L at S2 > 10 lg/L in groundwaters with Fe(II) < 0.8 mg/L, indicated controls of both Fe(II) and S2 cycling on As concentration in groundwater. Formation of pyrite and siderite also scavenges dissolved As by co-precipitation in strongly reducing environments (Sengupta et al., 2004; Lowers et al., 2007; Kim et al., 2009; Reza et al., 2010). One sample saturated with respect to both pyrite and siderite along this pathway had As concentration less than 10 lg/L, likely resulting from co-precipitation of As with pyrite and siderite. There were several samples with high As concentrations along Pathway 3 (Fig. 4). The possible reason is that the amount of As released from Fe(III) oxide reduction was much greater than that subsequently scavenged by pyrite and siderite precipitation, leaving high As concentration in groundwaters, although we cannot rule out other possibilities. Importantly, the three idealized pathways contributed in varying proportions to the overall groundwater chemical evolution, since most groundwaters are not located along any one pathway. Arsenic concentrations in the samples between Pathways 1 and 2 were mainly controlled by microbial Fe(III) reduction and re-adsorption, while those between Pathways 1 and 3 predominantly by microbial Fe(III) reduction and pyrite and siderite precipitation. The mixed effect of those pathways would regulate As and Fe cycling in most groundwaters. 5.3.2. Iron cycling and As behaviors along the flow path Nine groundwater samples were selected approximately along a flow path to assess Fe cycling and As behaviors in groundwaters associated with different hydrogeological settings (Fig. 1c). Generally, groundwater Eh was higher in the recharge zone than further down-gradient along the flow path, with exception of Wells No. 46 and 32 (Table 1, Fig. 8a). Although Well No. 46 had a depth of 100 m, groundwater had a relatively high Eh value (130 mV). Well No. 32 is located in the well-aerated evaporation discharge zone, as evidenced by its greater d18O values. Groundwater Eh dramatically increased from 56 mV in Well No. 36 to 150 mV in Well No. 32 due to aeration. This attributed to local decreases in H2S and NH4–N between Wells No. 39 and 28, although S2 concentration reached around 17 lg/L in Well No. 39. In the recharge zone (Wells No. 46 and 47), groundwaters occurred in oxic–suboxic conditions with Eh values of 130–290 mV (Fig. 8a), with low dissolved Fe and Fe(II) concentrations. Groundwater was enriched in light Fe isotopes with d56Fe values of 3.05& to 1.90&. Light d56Fe values and low dissolved Fe concentration indicated adsorption of isotopically light Fe into Fe-oxides or oxidative Fe(II) precipitation of isotopically light Fe as Fe-oxi-

69 68 This study 67 66 78 73 72 70 1.89 3.05 0.66 0.58 0.58 0.58 1.78 0.97 0.71 72 70 77 77 80 64 67 75 87 0.17 0.03 0.46 0.68 0.54 3.02 0.13 1.61 0.99 0.09 0.02 0.31 0.47 0.23 2.07 0.03 0.28 0.50 6.11 0.40 50.0 97.5 364 123 0.37 35.9 720 2.9 <1.0 39.7 92.4 328 99.9 <1.0 26.5 643 49.9 169 71.9 73.3 59.3 60.9 120 33.7 19.4 16.0 65.5 22.7 21.2 30.7 86.7 138 63.6 26.6 38.0 87.0 60.0 30.0 44.0 555 746 596 308 2.4 4.0 3.1 2.2 3.3 19.3 5.3 5.2 6.8 191 620 212 214 293 524 798 847 600 <0.01 124 1.04 4.16 6.70 <0.01 <0.01 2.08 <0.01 90.6 200 150 107 35.5 304 708 36.8 2.30 35.0 120 51.5 48.8 90.7 823 929 557 192 2.20 12.5 1.50 2.60 6.80 4.25 3.55 3.00 4.80 2.0 1.0 2.0 18.0 17.0 4.0 3.0 2.0 3.0 0.53 1.56 0.66 0.62 0.72 3.13 4.16 2.60 1.39 100 40 50 41 53 20 28 20 23 0.0 1.1 2.5 3.4 4.8 6.0 6.3 8.1 10 46 47 45 44 39 36 32 28 27

m km

7.9 7.1 8.3 8.0 8.3 8.2 7.9 8.0 8.6

130 290 67 61 48 56 150 60 52

lg/ L mS/ cm mV

3.3 4.2 15 3.1 3.3 4.3 6.0 11 15

10.3 9.6 10.4 10.7 11.1 8.8 8.0 9.3 11.6

& & & mg/ L mg/ L mg/L lg/ L lg/L mg/ L mg/L mg/ L mg/ L mg/L mg/L mg/L mg/ L

Ca2+ Mg2+ Na+ K+ HCO3 NO3 SO42 Cl

NH4– N mg/L H2S EC Eh pH Depth Distance No.

Table 1 Chemical and isotope compositions in groundwaters approximately along the flow path (shown in Fig. 1c and d).

As(III)

As

Fe(II)

Fe

DOC

d18O

dD

d56Fe

Well Nos. corresponding to Guo et al. (2011c)

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des. Experimental studies have shown that adsorption resulted in a decrease in aqueous d56Fe values by 0.7–1.2& relative to d56Fe of solid Fe(III) oxide (Icopini et al., 2004; Crosby et al., 2005). In this region, As also was adsorbed and immobilized on the Fe-oxides or Fe(II)–Fe(III) components due to their high affinity for As species (Robert et al., 2004; Cundy et al., 2008; Hohmann et al., 2010; Guo et al., 2013). Accordingly, low As groundwater was present with As concentration less than 10 lg/L. Down-gradient, d56Fe values increased to around 0.58& with increasing dissolved Fe concentration from 0.03 to 0.68 mg/L under anoxic conditions (Eh 60 mV). Dissolved Fe(II) produced from reductive dissolution of Fe(III) oxides usually had d56Fe values between 1.00& and 1.50&. Therefore, the heavier d56Fe could be explained by reductive dissolution of Fe(III) oxides between Wells No. 47 and 44. During reductive dissolution of Fe(III) oxides, As release was observed, resulting in an increase in As concentration from 0.4 lg/L in Well No. 47 to 98 lg/L in Well No. 44. Groundwater d56Fe values between Well Nos. 44 and 36 were near 0.6&, indicating that dissimilatory Fe(III) reduction (DIR) would be the major geochemical process controlling Fe cycling. However, these values are slightly higher than expected for the fractionation (around 0.3&) with microbially mediated reduction of Fe oxyhydroxides in the anoxic aquifer (Teutsch et al., 2005), and lower than representative values (around 1.00& to 1.50&) during DIR observed in experimental studies (Rouxel et al., 2008). Arsenic adsorbed on Fe(III) oxides would be released during dissimilatory Fe(III) reduction (DIR) (Nickson et al., 1998; Islam et al., 2004), leading to further increase in As concentration from 98 to 364 lg/L in Well No. 39. Between Well Nos. 39 and 36, dissolved Fe concentration increased and S2 concentration decreased (Fig. 8b and d), indicating possible precipitation of pyrite. Although precipitation of pyrite likely occurred, d56Fe value was less affected, remaining around 0.6&, due to counteracting effects on Fe isotope fractionation between dissimilatory reduction of Fe(III) oxides and Fe-sulfide precipitation. However, a decrease in As concentration was observed from 364 to 123 lg/L, due to pyrite precipitation with coprecipitation of As (Lowers et al., 2007). An abrupt decrease in d56Fe value from 0.58& to 1.78& in the redox transition zone (from reducing to moderately oxic) was observed between Well Nos. 36 and 32, where Eh increased from 56 to 150 mV (Fig. 8d). The change of redox conditions caused Fe(II) oxidation and Fe(III) oxide precipitation, resulting in a decrease in Fe concentrations from 3.02 to 0.13 mg/L. Therefore, the light d56Fe value may result from the precipitation of isotopically heavy Fe oxyhydroxides (Rouxel et al., 2008). The presence of Fe oxyhydroxides was the scavenger for As species, and therefore resulted in low As concentration in Well No. 32 (0.4 lg/L). Further down-gradient, d56Fe increased from 1.78& to 0.71& between Well Nos. 32 and 27, with a decrease in Eh value from 150 to 52 mV (Fig. 8). Dissolved Fe(II) concentration increased with the decrease in Eh. As shown

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Fig. 8. Variations in Eh and d18O values (a), S2 and NH4–N concentrations (b), As and As(III) concentrations (c), and d56Fe value and Fe(II) and dissolved Fe concentrations (d) approximately along the flow path as functions of distance from the recharge zone (Well No. 44).

between Well Nos. 47 and 44, DIR would be the major geochemical process controlling Fe cycling in this region. Subsequently, As was released during DIR, resulting in an increase in As concentration along the flow path from 0.4 to 720 lg/L. 6. CONCLUSION Our results provide a first attempt to characterize Fe isotope fractionation in high As groundwater and evaluating Fe and As cycling together in shallow aquifers using Fe isotope signatures. Studied groundwater had d56Fe values between 3.40& and 0.58& with a median of 1.14&. Positive d56Fe values were generally observed in the sediments (between 1.10& and 0.75&, median +0.36&). Results indicated the predominance of dissimilatory reduction of Fe(III) oxides in the aquifers. Chemical and Fe isotope characteristics showed three pathways of Fe cycling in groundwaters. One is the dissimilatory reduction of Fe(III) oxides, resulting in light d56Fe values (near – 1.0&) and high As concentration (>50 lg/L). Another is the readsorption of Fe(II) on Fe-oxide minerals. Since Fe-oxides would simultaneously adsorb Fe(II) and As species from groundwater, re-adsorption contributed to lighter d56Feaq values and low As concentrations along this reaction path-

way. The third pathway involves the formation of pyrite and siderite, with heavy d56Fe of groundwaters and light Fe being preferentially incorporated in the solid. Near the third pathway, groundwater As was generally low due to co-precipitation of As in pyrite and siderite. The mixed effect of these pathways would regulate As cycling in most groundwaters from shallow aquifer systems. ACKNOWLEDGMENTS The study has been financially supported by National Natural Science Foundation of China (Nos. 41222020 and 41172224), the Program for New Century Excellent Talents in University (No. NCET-07-0770), the Chinese Universities Scientific Fund (No. 2010ZD04), and the Chinese Scholarship Council (CSC). Constructive comments by Dr. Daniel Giammar (Associate Editor), Dr. Alexander van Geen and another three anonymous reviewers are also gratefully acknowledged.

APPENDIX A. SUPPLEMENTARY DATA Supplementary data associated with this article can be found, in the online version, at http://dx.doi.org/10.1016/ j.gca.2013.02.031.

H. Guo et al. / Geochimica et Cosmochimica Acta 112 (2013) 130–145 REFERENCES Appelo C. A. J., van der Weiden M. J. J., Tournassat C. and Charlet L. (2002) Surface complexation of ferrous iron and carbonate on ferrihydrite and the mobilization of arsenic. Environ. Sci. Technol. 36, 3096–3103. Archer C. and Vance D. (2006) Coupled Fe and S isotope evidence for Archean microbial Fe(III) and sulfate reduction. Geology 34(3), 153–156. Baker L. A., Qureshi T. M. and Wyman M. M. (1998) Sources and mobility of As in the Salt River watershed, Arizona. Water Resour. Res. 34(6), 1534–1552. Balci N., Bullen T. D., Witte-Lien K., Shanks W. C., Motelica M. and Mandernack K. W. (2006) Iron isotope fractionation during microbially stimulated Fe(II) oxidation and Fe(III) precipitation. Geochim. Cosmochim. Acta 70, 622–639. Beard B. L., Johnson C. M., Skulan J. L., Nealson K. H., Cox L. and Sun H. (2003a) Application of Fe isotopes to tracing the geochemical and biological cycling of Fe. Chem. Geol. 195, 87– 117. Beard B., Johnson C. M., Von Damm K. L. and Poulson R. L. (2003b) Iron isotope constraints on Fe cycling and mass balance in oxygenated Earth oceans. Geology 31, 629–632. Beard B. L., Handler R., Scherer M. M., Wu L. L., Czaja A. D., Heimann A. and Johnson C. M. (2010) Iron isotope fractionation between aqueous ferrous iron and goethite. Earth Planet. Sci. Lett. 295, 241–250. Berg M., Trang P. T. K., Stengel C., Buschmann J., Viet P. H., Van Dan N., Giger W. and Stu¨ben D. (2008) Hydrological and sedimentary controls leading to arsenic contamination of groundwater in the Hanoi area, Vietnam: The impact of iron arsenic ratios, peat, river bank deposits, and excessive groundwater abstraction. Chem. Geol. 249, 91–112. Borrok D. M., Wanty R. B., Ridley W. I., Wolf R., Lamothe P. J. and Adams M. (2007) Separation of copper, iron, and zinc from complex aqueous solutions for isotopic measurement. Chem. Geol. 242, 400–414. Brantley S. L., Liermann L. and Bullen T. D. (2001) Fractionation of Fe isotopes by soil microbes and organic acids. Geology 29, 535–538. Brantley S., Liermann L. J., Guynn R. L., Anbar A., Icopini G. A. and Barling J. (2004) Fe isotopic fractionation during mineral dissolution with and without bacteria. Geochim. Cosmochim. Acta 68, 3189–3204. Bullen T. D., White A. F., Childs C. W., Vivit D. V. and Schulz M. S. (2001) Demonstration of significant abiotic iron isotope fractionation in nature. Geology 29, 699–702. Butler I. B., Archer C., Vance D., Oldroyd A. and Rickard D. (2005) Fe isotope fractionation on FeS formation in ambient aqueous solution. Earth Planet. Sci. Lett. 236, 430–442. Charlet L. and Polya D. (2006) Arsenic in shallow, reducing groundwaters in Southern Asia: An environmental health disaster. Elements 2, 91–96. Clayton R. E., Hudson-Edwards K. A., Malinovsky D. and Andersson P. (2005) Fe isotope fractionation during the precipitation of ferrihydrite and transformation of ferrihydrite to goethite. Mineral. Mag. 69(5), 667–676. Croal L. R., Johnson C. M., Beard B. L. and Newman D. K. (2004) Iron isotope fractionation by Fe(II)-oxidizing photoautotrophic bacteria. Geochim. Cosmochim. Acta 68, 1227–1242. Crosby H. A., Johnson C. M., Roden E. E. and Beard B. L. (2005) Coupled Fe(II)Fe(III) electron and atom exchange as a mechanism for Fe isotope fractionation during dissimilatory iron oxide reduction. Environ. Sci. Technol. 39, 6698–6704. Crosby H. A., Roden E. E., Johnson C. M. and Beard B. L. (2007) The mechanisms of iron isotope fractionation produced during

143

dissimilatory Fe(III) reduction by Shewanella putrefaciens and Geobacter sulfurreducens. Geobiology 5, 169–189. Cundy A. B., Hopkinson L. and Whitby R. L. D. (2008) Use of iron-based technologies in contaminated land and groundwater remediation: A review. Sci. Total Environ. 400, 42–51. Emmanuel S., Erel Y., Matthews A. and Teutsch N. (2005) A preliminary mixing model for Fe isotopes in soils. Chem. Geol. 222, 23–34. Guilbaud R., Butler I. B., Ellam R. M. and Rickard D. (2010) Fe isotope exchange between Fe(II)aq and nanoparticulate mackinawite (FeSm) during nanoparticle growth. Earth Planet. Sci. Lett. 300, 174–183. Guilbaud R., Butler I. B. and Ellam R. M. (2011a) Abiotic pyrite formation produces a large Fe isotope fractionation. Science 332, 1548–1551. Guilbaud R., Butler I. B., Ellam R. M., Rickard D. and Oldroyd A. (2011b) Experimental determination of the equilibrium Fe isotope fractionation between Feaq2+ and FeSm (mackinawite) at 25 and 2 °C. Geochim. Cosmochim. Acta 75, 2721–2734. Guo H. M., Yang S. Z., Tang X. H., Li Y. and Shen Z. L. (2008a) Groundwater geochemistry and its implications for arsenic mobilization in shallow aquifers of the Hetao basin, Inner Mongolia. Sci. Total Environ. 393, 131–144. Guo H. M., Tang X. H. and Yang S. Z. (2008b) Effect of indigenous bacteria on geochemical behavior of arsenic in aquifer sediments from the Hetao basin, Inner Mongolia: Evidence from sediment incubation. Appl. Geochem. 23, 3267– 3277. Guo H. M., Zhang B., Wang G. C. and Shen Z. L. (2010) Geochemical controls on arsenic and rare earth elements approximately along a groundwater flow path in the shallow aquifer of the Hetao basin, Inner Mongolia. Chem. Geol. 270, 117–125. Guo H. M., Zhang B. and Zhang Y. (2011a) Control of organic and iron colloids on arsenic partition and transport in high arsenic groundwaters in the Hetao basin, Inner Mongolia. Appl. Geochem. 26, 360–370. Guo H. M., Li Y., Zhao K. and Ren Y. (2011b) Removal of arsenite by synthetic siderite: Behaviors and mechanisms. J. Hazard. Mater. 186, 1847–1854. Guo H. M., Zhang B., Li Y., Berner Z., Tang X. H. and Norra S. (2011c) Hydrogeological and biogeochemical constrains of arsenic mobilization in shallow aquifers from the Hetao basin, Inner Mongolia. Environ. Pollut. 159, 876–883. Guo H. M., Ren Y., Liu Q., Zhao K. and Li Y. (2013) Enhancement of arsenic adsorption during mineral transformation from siderite to goethite: Mechanism and application. Environ. Sci. Technol. 47(2), 1009–1016. Handler R. M., Beard B. L., Hohnson C. and Scherer M. M. (2009) Atom exchange between aqueous Fe(II) and goethite: An Fe isotope tracer study. Environ. Sci. Technol. 43, 1102–1107. Heimann A., Johnson C. M., Beard B. L., Valley J. W., Roden E. E., Spicuzza M. J. and Beukes N. J. (2010) Fe, C, and O isotope compositions of banded iron formation carbonates demonstrate a major role for dissimilatory iron reduction in 2.5 Ga marine environments. Earth Planet. Sci. Lett. 294, 8–18. Hohmann C., Winkler E., Morin G. and Kappler A. (2010) Anaerobic Fe(II)-oxidizing bacteria show As resistance and immobilize As during Fe(III) mineral precipitation. Environ. Sci. Technol. 44(1), 94–101. Horneman A., van Geen A., Kent D., Mathe P. E., Zheng Y., Dhar R. K., O’Connell S., Hoque M., Aziz Z., Shamsudduha M., Seddique A. and Ahmed K. M. (2004) Decoupling of As and Fe release to Bangladesh groundwater under reducing conditions. Part I: Evidence from sediment profiles. Geochim. Cosmochim. Acta 68, 3459–3473.

144

H. Guo et al. / Geochimica et Cosmochimica Acta 112 (2013) 130–145

Icopini G. A., Anbar A. D., Ruebush S. S., Tien M. and Brantley S. L. (2004) Iron isotope fractionation during microbial reduction of iron: The importance of adsorption. Geology 32, 205–208. Islam F. S., Gault A. G., Boothman C., Polya D. A., Charnock J. M., Chatterjee D. and Lloyd J. R. (2004) Role of metalreducing bacteria in arsenic release from Bengal delta sediments. Nature 430, 68–71. Jang J. H., Mathur R., Liermann L. J., Ruebush S. and Brantley S. L. (2008) An iron isotope signature related to electron transfer between aqueous ferrous iron and goethite. Chem. Geol. 250, 40–48. Jessen S., Larsen F., Koch C. B. and Avin E. (2005) Sorption and desorption of arsenic to ferrihydrite in a sand filter. Environ. Sci. Technol. 39, 8045–8051. Johnson C. M., Skulan J. L., Beard B. L., Sun H., Nealson K. H. and Braterman P. S. (2002) Isotopic fractionation between Fe(III) and Fe(II) in aqueous solutions. Earth Planet. Sci. Lett. 195, 141–153. Johnson C. M., Beard B. L., Roden E. E., Newman D. K. and Nealson K. H. (2004) Isotopic constraints on biogeochemical cycling of Fe. Rev. Mineral. Geochem. 55, 359–408. Johnson C. M., Roden E. E., Welch S. A. and Beard B. L. (2005) Experimental constraints on Fe isotope fractionation during magnetite and Fe carbonate formation coupled to dissimilatory hydrous ferric oxide reduction. Geochim. Cosmochim. Acta 69(4), 963–993. Johnson C. M., Beard B. L. and Roden E. E. (2008) The iron isotope fingerprints of redox and biogeochemical cycling in modern and ancient earth. Annu. Rev. Earth Planet. Sci. 36, 457–493. Jones G. W. and Pichler T. (2007) Relationship between pyrite stability and arsenic mobility during aquifer storage and recovery in southwest central Florida. Environ. Sci. Technol. 41(3), 723–730. Kappler A., Johnson C. M., Crosby H. A., Beard B. L. and Newman D. K. (2010) Evidence for equilibrium iron isotope fractionation by nitrate-reducing iron(II)-oxidizing bacteria. Geochim. Cosmochim. Acta 74, 2826–2842. Kim K. J., Park S. M., Kim J., Kim S. H., Kim Y., Moon J. T., Hwang G. S. and Cha W. S. (2009) Arsenic concentration in porewater of an alkaline coal ash disposal site: Roles of siderite precipitation/dissolution and soil cover. Chemosphere 77, 222– 227. Kim S. H., Kim K., Ko K. S., Kim Y. and Lee K. S. (2012) Cocontamination of arsenic and fluoride in the groundwater of unconsolidated aquifers under reducing environments. Chemosphere 87, 851–856. Ligier E., Charlet L. and van Cappellen P. (1999) Surface catalysis of uranium (VI) reduction by Fe(II). Geochim. Cosmochim. Acta 63, 2939–2955. Lowers H. A., Breit G. N., Foster A. L., Whitney J., Yount J., Uddin N. and Muneem A. (2007) Arsenic incorporation into authigenic pyrite, Bengal basin sediment, Bangladesh. Geochim. Cosmochim. Acta 71, 2699–2717. Mandernack K. W., Bazylinski D. A., Shanks W. C. and Bullen T. D. (1999) Oxygen and iron isotope studies of magnetite produced by magnetotactic bacteria. Science 285, 1892–1896. Matthews A., Morgans-Bell H. S., Emmanuel S., Jenkyns H. C., Erel Y. and Halicz L. (2004) Controls on iron-isotope fractionation in organic-rich sediments (Kimmeridge Clay, Upper Jurassic, southern England). Geochim. Cosmochim. Acta 68(14), 3107–3123. McArthur J. M., Banerjee D. M., Hudson-Edwards K. A., Mishra R., Purohit R., Ravenscroft P., Cronin A., Howarth R. J., Chatterjee A., Talukder T., Lowry D., Houghton S. and

Chadha (2004) Natural organic matter in sedimentary basins and its relation to arsenic in anaerobic ground water: The example of West Bengal and its worldwide implications. Appl. Geochem. 19, 1255–1293. Mukherjee A., Bhattacharya P., Shi F., Fryar A. E., Mukherjee A. B., Xie Z. M., Jacks G. and Bundschuh J. (2009) Chemical evolution in the high arsenic groundwater of the Huhhot basin (Inner Mongolia, PR China) and its difference from the western Bengal basin (India). Appl. Geochem. 24, 1835–1851. Nickson R. T., McArthur J. M., Burgess W. G., Ahmed K. M., Ravenscroft P. and Rahman M. (1998) Arsenic poisoning of Bangladesh groundwater. Nature 395, 338. Parkhurst D. L. and Appelo C. A. J. (1999) User’s guide to PHREEQC (version 2) – A computer program for speciation, batch-reaction, one dimensional transport and inverse geochemical calculation. USGS Water Resources Investigation Report 99-4259, U.S. Geological Survey, Denver. p. 312. Percak-Dennett E. M., Beard B. L., Xu H., Konishi H., Johnson C. M. and Roden E. E. (2011) Iron isotope fractionation during microbial dissimilatory iron oxide reduction in simulated Archaean seawater. Geobiology 9(3), 205–220. Peters S. C., Blum J. D., Klaue B. and Karagas M. R. (1999) Arsenic occurrence in New Hampshire drinking water. Environ. Sci. Technol. 33(9), 1328–1333. Pili E., Tisserand D. and Bureau S. (2012) Origin, mobility, and temporal evolution of arsenic from a low-contamination catchment in Alpine crystalline rocks. J. Hazard. Mater. 2012. http://dx.doi.org/10.1016/j.jhazmat.2012.07.004. Ravenscroft P., Brammer H. and Richards K. S. (2009) Arsenic Pollution: A Global Synthesis. Wiley-Blackwell, Singapore, pp. 1–114. Reza A. H. M. S., Jean J. S., Yang H. J., Lee M. K., Woodall B., Liu C. C., Lee J. F. and Luo S. D. (2010) Occurrence of arsenic in core sediments and groundwater in the Chapai-Nawabganj District, northwestern Bangladesh. Water Res. 44, 2021–2037. Robert L. C., Hug S. J., Ruettimann T., Billah M. M., Khan A. W. and Rahman M. T. (2004) Arsenic removal with iron(II) and iron(III) in waters with high silicate and phosphate concentrations. Environ. Sci. Technol. 38, 307–315. Rouxel O., Sholkovitz E., Charette M. and Edwards K. J. (2008) Iron isotope fractionation in subterranean estuaries. Geochim. Cosmochim. Acta 72, 3413–3430. Rowland H. A. L., Gault A. G., Lythgoe P. and Polya D. A. (2008) Geochemistry of aquifer sediments and arsenic-rich groundwaters from Kandal Province, Cambodia. Appl. Geochem. 23, 3029–3046. Roy M., Rouxel O., Martin J. B. and Cable J. E. (2012) Iron isotope fractionation in a sulfide-bearing subterranean estuary and its potential influence on oceanic Fe isotope flux. Chem. Geol. 300–301, 133–142. Sengupta S., Mukherjee P. K., Pal T. and Shome S. (2004) Nature and origin of arsenic carriers in shallow aquifer sediments of Bengal Delta, India. Environ. Geol. 45, 1071–1081. Severmann S., Johnson C. M., Beard B. L. and McManus J. (2006) The effect of early diagenesis on the Fe isotope compositions of porewaters and authigenic minerals in continental margin sediments. Geochim. Cosmochim. Acta 70, 2006–2022. Sivan O., Adler M., Pearson A., Gelman F., Bar-Or I., John S. G. and Eckert W. (2011) Geochemical evidence for iron-mediated anaerobic oxidation of methane. Limnol. Oceanogr. 56, 1536– 1544. Skulan J. L., Beard B. L. and Johnson C. M. (2002) Kinetic and equilibrium Fe isotope fractionation between aqueous Fe(III) and hematite. Geochim. Cosmochim. Acta 66, 2995–3015.

H. Guo et al. / Geochimica et Cosmochimica Acta 112 (2013) 130–145 Smedley P. L. and Kinniburgh D. G. (2002) A review of the source, behaviour and distribution of arsenic in natural waters. Appl. Geochem. 17, 517–568. Smedley P. L., Zhang M., Zhang G. and Luo Z. (2003) Mobilisation of arsenic and other trace elements in fluviolacustrine aquifers of the Huhhot basin, Inner Mongolia. Appl. Geochem. 18, 1453–1477. Stu¨ben D., Berner Z., Chandrasekharam D. and Karmakar J. (2003) Arsenic enrichment in groundwater of West Bengal, India: Geochemical evidence for mobilization of As under reducing conditions. Appl. Geochem. 18, 1417–1434. Sun X. and Doner H. E. (1998) Adsorption and oxidation of arsenite on goethite. Soil Sci. 163, 278–287. Tangalos G. E., Beard B. L., Johnson C. M., Alpers C. N., Shelobolina E. S., Xu H., Konishi H. and Roden E. E. (2010) Microbial production of isotopically light iron(II) in a modern chemically precipitated sediment and implications for isotopic variations in ancient rocks. Geobiology 8(3), 197–208. Teutsch N., Von Gunten U., Porcelli D., Cirpka O. A. and Halliday A. N. (2005) Adsorption as a cause for iron isotope fractionation in reduced groundwater. Geochim. Cosmochim. Acta 69, 4175–4185. van Geen A., Rose J., Thoral S., Garnier J. M., Zheng Y. and Bottero J. Y. (2004) Decoupling of As and Fe release to Bangladesh groundwater under reducing conditions. Part II: Evidence from sediment incubations. Geochim. Cosmochim. Acta 68, 3475–3486. van Geen A., Radloff K., Aziz Z., Cheng Z., Huq M. R., Ahmed K. M., Weinman B., Goodbred S., Jung H. B., Zheng Y., Berg M., Trang P. T. K., Charlet L., Metral J., Tisserand D., Guillot S., Chakraborty S., Gajurel A. P. and Upreti B. N. (2008) Comparison of arsenic concentrations in simultaneously-collected groundwater and aquifer particles from Bangladesh, India, Vietnam, and Nepal. Appl. Geochem. 23, 3244–3251. Wang Y., Jiao J. J. and Cherry J. A. (2012) Occurrence and geochemical behavior of arsenic in a coastal aquifer–aquitard system of the Pearl River Delta, China. Sci. Total Environ. 427– 428, 286–297. Welch S. A., Beard B. L., Johnson C. M. and Braterman P. S. (2003) Kinetic and equilibrium Fe isotope fractionation

145

between aqueous Fe(II) and Fe(III). Geochim. Cosmochim. Acta 67(22), 4231–4250. Wiesli R. A., Beard B. L. and Johnson C. M. (2004) Experimental determination of Fe isotope fractionation between aqueous Fe(II), siderite and “green rust” in abiotic systems. Chem. Geol. 211, 343–362. Wu L. L., Beard B. L., Roden E. E. and Johnson C. M. (2009) Influence of pH and dissolved Si on Fe isotope fractionation during dissimilatory microbial reduction of hematite. Geochim. Cosmochim. Acta 73(19), 5584–5599. Wu L. L., Beard B. L., Roden E. E., Kennedy C. B. and Johnson C. M. (2010) Stable Fe isotope fractionations produced by aqueous Fe(II)–hematite surface interactions. Geochim. Cosmochim. Acta 74, 4249–4265. Wu L. L., Beard B. L., Roden E. E. and Johnson C. M. (2011) Stable iron isotope fractionation between aqueous Fe(II) and hydrous ferric oxide. Environ. Sci. Technol. 45, 1847–1852. Wu L. L., Percak-Dennett E. M., Beard B. L., Roden E. E. and Johnson C. M. (2012a) Stable iron isotope fractionation between aqueous Fe(II) model Archean ocean Fe–Si coprecipitates and implications for iron isotope variations in the ancient rock record. Geochim. Cosmochim. Acta 84, 14–28. Wu L. L., Druschel G., Findlay A., Beard B. and Johnson C. M. (2012b) Experimental determination of iron isotope fractionations among Feaq2+–FeSaq–Mackinawite at low temperatures: Implications for the rock record. Geochim. Cosmochim. Acta 89, 46–61. Yamaguchi K. E., Johnson C. M., Beard B. L. and Ohmoto H. (2005) Biogeochemical cycling of iron in the Archean–Paleoproterozoic Earth: Constraints from iron isotope variations in sedimentary rocks from the Kaapvaal and Pilbara Cratons. Chem. Geol. 218, 135–169. Zhang Y. L., Cao W. G., Wang W. Z. and Dong Q. Y. (2012) Distribution of groundwater arsenic and hydraulic gradient along the shallow groundwater flow-path in Hetao plain, Northern China. J. Geochem. Explor.. http://dx.doi.org/ 10.1016/j.gexplo.2012.12.004. Associate editor: Daniel E. Giammar