Dissolved inorganic carbon evolution and stable carbon isotope fractionation in acid mine drainage contaminated streams: Insights from a laboratory study

Applied Geochemistry 23 (2008) 2634–2648

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Dissolved inorganic carbon evolution and stable carbon isotope fractionation in acid mine drainage contaminated streams: Insights from a laboratory study Ernest W. Fonyuy a, Eliot A. Atekwana b,* a b

Department of Geological Sciences and Engineering, 129 McNutt Hall, Missouri University of Science and Technology, Rolla, MO 65409, USA Boone Pickens School of Geology, 105 Noble Research Center, Oklahoma State University, Stillwater, OK 74078, USA

a r t i c l e

i n f o

Article history: Received 15 September 2007 Accepted 25 May 2008 Available online 25 July 2008 Editorial handling by R. Fuge

a b s t r a c t Samples of groundwater, spring water and stream water contaminated by acid mine drainage (AMD), and uncontaminated stream water were collected and allowed to evolve in contact with air in the laboratory for 15–88 days. The objective of this study was to (1) document temporal changes in dissolved inorganic C (DIC) concentrations and stable isotopic composition (d13CDIC) and (2) to determine the reaction mechanism and resulting isotopic fractionation (13C/12C) accompanying the chemical evolution of AMD. The contaminated spring and stream samples and one groundwater sample (with no HCO 3) showed temporal decreases in pH, Fe2+, alkalinity, and DIC, and enrichment in d13CDIC. One contaminated groundwater sample (with HCO 3 between 529 and 630 mg/L) showed a temporal increase in pH despite observed decreases in Fe2+, alkalinity and DIC, and enrichment in d13CDIC. The uncontaminated stream samples showed a continuous temporal increase in pH, relatively constant alkalinity and DIC, and enrichment in d13CDIC. The results suggest that proton production related to Fe2+ transformation is the driving force for DIC loss in AMD-contaminated samples, and that DIC loss can be described by first order kinetics. The C isotope enrichment rates associated with DIC loss in the contaminated samples varied between 1.0‰ and 1.8‰ for stream water, 2.1‰ and 2.6‰ for the spring, 1.0‰ and  1.2‰ for groundwater with no HCO 3 , and 7.6‰ and 9.3‰ for groundwater with high HCO3 . 13 Variations in C enrichment in the contaminated samples are attributed to differences in 13 C enrichment in the AMDthe initial Fe2þ :HCO 3 ratio. The effect of proton production on contaminated samples was modeled as a Rayleigh-type distillation, whereby isotope fractionation was constant and occurred in an ‘‘equilibrium closed system”. In the uncontaminated stream samples, C exchange between DIC and atmospheric CO2 resulted in an overall enrichment in d13CDIC of 6‰. It is concluded that C isotope enrichment induced by the chemical evolution of AMD in contaminated streams should range from 1.0‰ to 3.0‰ in the absence of in-stream processes that may affect DIC. Ó 2008 Elsevier Ltd. All rights reserved.

1. Introduction The acid (H+) in acid mine drainage (AMD) and H+ subsequently generated from hydrolysis of dissolved metals (e.g., Fe3+, Al3+, and Mn2+; e.g., Morin et al., 1988; Berger  et al., 2000) in surface waters dehydrates CO2 3 and HCO3 * Corresponding author. Tel.: +1 405 744 9247; fax: +1 405 744 7841. E-mail address: [email protected] (E.A. Atekwana). 0883-2927/$ - see front matter Ó 2008 Elsevier Ltd. All rights reserved. doi:10.1016/j.apgeochem.2008.05.012

and produces CO2(aq) that is lost from solution as CO2(g). Thus, a unique attribute of hydrologic systems contaminated by AMD is the coupling of metal and C cycling. Numerous studies have examined the production and fate of metals in AMD and streams contaminated by AMD (e.g., Chapman et al., 1983; Sullivan and Drever, 2001; Yu and Heo, 2001; Kim and Kim, 2004; Olias et al., 2004; Gunsinger et al., 2006; Sanchez-Espana et al., 2006; Canovas et al., 2007; Cravotta, 2008a,b). However, there remains a

E.W. Fonyuy, E.A. Atekwana / Applied Geochemistry 23 (2008) 2634–2648

paucity of published data on dissolved inorganic C (DIC) concentrations and stable isotopic compositions (d13CDIC) in AMD-contaminated hydrologic systems (e.g., Toran, 1990; Mayo et al., 1992, 2000; Perez del Villar et al., 2005; Fonyuy and Atekwana, 2008). Carbonate dehydration reactions and diffusive loss of CO2(g) driven by AMD acidification are accompanied by C isotope fractionation that results in the enrichment of d13CDIC (Fonyuy and Atekwana, 2008). The acid in AMD also enhances photooxidation of organic C in surface waters (Gennings et al., 2001; Friese et al., 2002; Wu et al., 2005). Additionally, a better understanding of the isotopic effects of AMD on ‘‘natural” DIC provides constraints that are requisite to fully understand the sources of sedimentary organic matter that drive key biogeochemical processes (e.g., Blodau et al., 2000; Chabbi et al., 2006). In a field study of an AMD-contaminated stream, Fonyuy and Atekwana (2008) showed that metal concentrations decreased in the downstream direction, concomitant with decreases in DIC and enrichment in d13CDIC. The authors suggest that DIC cycling was in part coupled to metal cycling. The chemical model for this process is that (1) reductions in aqueous metal concentrations and accompanying increases in H+ result from metal hydrolysis, and (2) decreases in DIC concentrations and enrichment in d13CDIC result directly from variations in proton concentrations. In AMD-contaminated streams, hydrolysis and precipitation of dissolved metals reduce their aqueous concentrations (e.g., Berger et al., 2000; Sullivan and Drever, 2001; Kim et al., 2003; Kim and Kim, 2004). This decrease, however, could also result from dilution by uncontaminated groundwater and tributaries (e.g., Yu and Heo, 2001; Olias et al., 2004; Canovas et al., 2007). Thus, the extent to which the decrease in metal concentration is due to hydrolysis in AMD-contaminated streams must be quantified to fully assess this effect on DIC and d13CDIC. Quantifying the effect of these reactions in natural streams is difficult because of several competing chemical and physical processes that affect DIC and d13CDIC. Streams continuously exchange CO2 with the atmosphere (e.g., Kling et al., 1991; Telmer and Veizer, 1999; Worrall and Lancaster, 2005). If the partial pressure of CO2 (pCO2) in stream water is higher than atmospheric, DIC will decrease due to CO2 loss and the d13CDIC will become more positive because of a preferential loss of 12C (Pawellek and Veizer, 1994; Telmer and Veizer, 1999; Finlay, 2003). A decrease in the DIC concentration of stream water may also result from aquatic photosynthetic drawdown (O’Leary, 1988), which is accompanied by the enrichment of d13CDIC. Increases in the DIC concentrations can result from the dissolution of carbonates (d13C 0‰; Wigley et al., 1978) in streambed sediments or channels (Wicks and Groves, 1993; Webb and Sasowsky, 1994). This process will enrich d13CDIC. On the other hand, aquatic respiration and oxidation of organic matter may provide a source of isotopically depleted CO2 that will increase stream water DIC and deplete d13CDIC (e.g., Clark and Fritz, 1997). In addition, hydrologic factors such as spatial and seasonal groundwater discharge and input from tributary streams may also contribute to variations in DIC concentration and d13CDIC. Because these processes can affect the DIC pool over different spatial and temporal scales

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(Wachniew, 2006), it becomes difficult to assign particular shifts in the DIC concentrations and changes in d13CDIC to metal cycling reactions induced by AMD pollution. However, under controlled conditions in the laboratory, the effects of AMD on DIC and d13CDIC can be inferred with greater certainty. In this study, the results of a laboratory experiment in which AMD-contaminated stream, groundwater and spring samples, and uncontaminated stream samples, were allowed to react under ambient conditions (in contact with the atmosphere) for 15–88 days, are reported. It was hypothesized that Fe2+ in the contaminated samples would react with O2 to form Fe3+, which is subsequently hydrolyzed to precipitate Fe oxides and produce H+. The H+  would then dehydrate CO2 3 and HCO3 and produce CO2(aq). Therefore the objectives were to (1) document temporal changes in DIC and d13CDIC, and (2) determine the reaction mechanism and the fractionation of C isotopes (13C/12C) accompanying the evolution of AMD.

2. Methods 2.1. Sample selection AMD-contaminated and uncontaminated water samples for the experiments were collected from Huntsville, Missouri, USA which was an area of intense coal mining from 1891 to 1960 (Fig. 1). Mining occurred in coal seams interbedded with Pennsylvanian-aged shale and clay, which is successively overlain by limestone and sandstone, and covered by glacial drift and/or alluvium (Blevins and Ziegler, 1992). AMD was generated in the waste rock and mine cavities both during and after mining activities, and has contaminated groundwater and streams in the area (Blevins and Ziegler, 1992; Christensen, 2005). The AMD-contaminated groundwater samples were collected from two United States Geological Survey monitoring wells (GW-1 and GW-2) in Huntsville Gob. The contaminated spring (AMS-1) occurred in the headwater of Michelle Creek. Contaminated stream water was collected 60 m (MC-2), 800 m (MC-5), and 1220 m (MC-7) downstream from Michelle Creek headwater. Uncontaminated stream water was collected from two locations: a stream at Highway O (HWO) in Huntsville with a similar geologic and climatic setting as the contaminated stream samples and from Little Piney Branch (LP) in Rolla, Missouri (37°570 0500 N; 91°460 1600 W). The uncontaminated stream samples contained SO2 and Fe2+ concentrations below 4 the United States National Secondary Drinking Water limit of 250 mg/L and 0.3 mg/L, respectively (US Environmental Protection Agency, 2005). To account for the natural variability in hydrologic conditions, water samples were collected in winter, spring and summer of 2005. The samples were collected in 20 L FortPaksÒ plastic containers, which were subsequently used as reactors. Samples for the winter sampling period were collected from AMS-1, MC-5, MC-7 and GW-1. In the spring, samples were collected from the same stations as in the winter, with an additional groundwater sample collected from GW-2. The GW-2 sample was included because it

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92 0 34’W

920 31’ 58” W

92 0 29’ 56” W

39 0 32’ N

39 0 30’ N

Mi ch 39 0

28’ N

e ll

e

GW-2

Cr ee k

-7 MC

ville Hunts

-5 MC

Gob

GW-1

-2 MC 1 AMS-

39 0 26’ N USA

0

study site

Missouri

1Km

Watershed boundary

AMD spring sampling station

N

Stream course Lagoon

Stream water sampling station

Road

Groundwater sampling station

Railroad

Fig. 1. Map showing sample collection sites for AMD-contaminated stream water and groundwater in Huntsville, Missouri, USA.

had a greater extent of AMD contamination (e.g., higher 2+ SO2 4 , specific conductance, and Fe , and lower pH) compared to GW-1 (Table 1). During the summer, samples were collected from the same stations as in spring, with an additional stream sample from MC-2. The MC-2 sample was included because the sampling site is located 5 m downstream of AMS-1, so that this water was assumed not to be as influenced by in-stream processes as samples collected from MC-5 and MC-7 further downstream. 2.2. Water sampling and analysis The water samples collected at each location were unfiltered composite samples. The reactors were filled with no head space to minimize exposure of the samples to air during transit. Upon return to the laboratory, the reactor lids were removed to expose the samples to air and start the experiments. Oxygen diffused into the samples and reacted with Fe2+ to produce Fe3+ followed by hydrolysis of Fe3+, which resulted in the production of protons and precipitates (e.g., Fonyuy and Atekwana, 2008). The effect of the protons on DIC and d13C was determined from temporal physical measurements and sampling of each reactor for chemical and isotopic analyses. The frequency of measurement and sampling of each reactor was based on a combi-

nation of factors such as rate of pH change, DIC and Fe2+decrease, and when Fe2+ concentration decreased to below the detection limit of <0.01 mg/L. Prior to collecting samples both in the field and from the reactors, temperature, specific conductance (SPC), dissolved O2 (DO) and Eh were measured with a Yellow Spring Instruments (YSI) multi-probe field meter. The YSI multiprobe meter was calibrated according to manufacturer instructions. Water for chemical analyses was filtered using 0.45 lm syringe filters. A portion of the filtered water was immediately used to determine alkalinity by acid titration (Hach, 1992) and Fe2+ using the Phenanthroline method (CHEMetrics, 2004). Unacidified aliquots of the filtered water were stored for anion analysis. Aliquots for cation and metal analysis were acidified with high purity HNO3 prior to storage. The samples for major ions and metal concentrations were stored at 4 °C until analysis. SO2 was analyzed by ion chromatography and Ca, Mg, 4 Mn and Al by inductively coupled plasma-optical emission spectrometry (ICP-OES). Variable amounts (10–150 mL) of water for CO2 extraction and C isotope determination were directly filtered into evacuated vials pre-loaded with 85% phosphoric acid and magnetic stir bars. The vials consisted of 16  100 mm glass septum tubes (VACUTAINERÒ Serum Tubes, Becton Dickson & Company, Franklin Lakes, NJ

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Table 1 Results of the initial physical, chemical and C isotope analyses of AMD-contaminated and uncontaminated samples collected for the laboratory experiment Sample ID

Temp. (°C)

SPC (lS/cm)

DO (mg/L)

Eh (mV)

pH

Alkalinitya (mg/L)

DIC (mg C/L)

d13CDIC (‰)

SO2 4 (mg/L)

Fe2+ (mg/L)

AMD-contaminated samples Winter AMS1-1 MC5-1 MC7-1 GW1-1

13.6 9.6 7.1 13.9

2057 1750 1550 3684

0.3 9.6 9.8 0.4

9 67 18 4

5.83 6.29 6.66 6.04

157 58 76 630

85.8 15.7 13.3 199.9

11.5 6.9 3.9 10.5

1688 1646 1536 3012

146.98 87.21 57.90 61.00

Spring AMS1-1 MC5-1 MC7-1 GW1-1 GW2-1

13.5 15.2 14.7 14.7 13.7

2740 2856 2725 4645 8194

0.1 8.2 8.9 0.4 0.4

16 100 104 32 199

5.74 5.46 5.36 6.02 3.76

165 13 8 575 bdl

115.4 11.8 1.9 299.8 212.2

11.7 7.7 3.9 12.1 12.9

1823 2181 1960 3254 9640

100.90 117.40 59.20 48.55 1882.44

Summer AMS1-1 MC2-1 MC5-1 MC7-1 GW1-1 GW2-1

14.3 15.2 22.0 19.7 15.1 14.8

2223 2261 2787 2728 3857 6115

0.3 7.7 7.3 8.0 0.2 0.9

32 4 208 351 5 143

5.57 5.85 4.76 3.28 5.98 3.62

136 100 4 bdl 529 bdl

119.7 51.8 4.8 0.7 339.9 172.1

11.3 7.6 9.1 nd 12.1 12.9

1353 1592 1616 1787 2726 7127

84.25 107.50 85.17 57.63 31.75 1969.38

464 383

5.7 8.8

190.0 153.6

7.14 7.93

155 177

32.3 40.1

12.6 11.6

Uncontaminated samples HWO-1 22.6 LP-1 30.3

133.0 12.6

0.28 0.02

bdl = below detection limit. a Concentration expressed in mg/L CaCO3.

07417) or 25, 50, 100, or 250 mL PYREX volumetric flasks. The necks of the volumetric flasks were modified by adding a short length of glass tubing that would fit the rubber septa from the vacutainer tubes. CO2 was extracted from the vials on a vacuum line and its concentration was determined manometrically following cryogenic separation as described by Atekwana and Krishnamurthy (1998). The extracted CO2 was stored in 6 mm Pyrex tubes and later analyzed for its isotopic composition by isotope ratio mass spectrometry at Western Michigan University, Kalamazoo, Michigan. The isotopic composition is expressed in ‰ units relative to the Vienna Pee Dee Belemnite (VPDB) C standard, using the standard delta notation. The isotopic composition has a precision of better than 0.1‰. 3. Results and discussion 3.1. Temporal evolution of AMD-contaminated samples The initial physical, chemical and isotopic parameters measured in field samples are shown in Table 1, and the results of the time-series experiments are presented in Table S-1. Chemical reactions generating AMD at the sampling locations have been described previously (Blevins and Ziegler, 1992; Christensen, 2005) and are initiated by oxidation of Fe sulfide minerals (e.g., pyrite) in waste rock and abandoned mine cavities. Pyrite oxidation is described by the following equation (e.g., Singer and Stumm, 1970): 2þ

FeS2 þ 3:5O2 þ H2 O ! Fe

þ

þ 2H þ 2+

2SO2 4 SO2 4

ð1Þ

The seasonal variability of Fe and in the contaminated groundwater and spring samples used in the labora-

tory experiment are related to the extent of pyrite oxidation, water–rock interactions, and dilution by uncontaminated groundwater (e.g., Toran, 1990; Mayo et al., 1992; Webb and Sasowsky, 1994). Reaction (1) indicates that the production of protons should result in very low pH in the AMD-contaminated samples. However, very low pH was only measured in GW-2 with a pH of 3.6–3.8 (Table 1). The higher pH values (5.6–6.0) measured for GW-1 and AMS-1 are likely due to neutralization of protons by carbonate minerals in the waste rock and aquifer (e.g., Webb and Sasowsky, 1994; Blowes et al., 2003; Gunsinger et al., 2006):

CaCO3 þ Hþ ! Ca2þ þ HCO3

ð2Þ

When AMD with high metal concentrations (e.g., Fe2+) is exposed to the atmosphere or enters and mixes with water in receiving streams, oxidation and subsequent hydrolysis and precipitation of the metals occurs. This reduces the dissolved metal concentrations and produces protons that result in acidic pH values characteristic of AMD-contaminated surface waters (e.g., Morin et al., 1988; Blodau, 2006). The following reaction pathway linking Fe2+ and Fe3+ results in the generation of acidity (e.g., Singer and Stumm, 1970):

Fe2þ þ 0:25O2 þ Hþ ! Fe3þ þ 0:5H2 O Fe

3þ

þ

þ 3H2 O ! FeðOHÞ3ðsÞ þ 3H

ð3Þ ð4Þ

The chemical evolution of the contaminated samples (e.g., via reactions (3) and (4)) in the reactors is evident in the temporal water chemistry (Fig. 2; Table S-1). Fig. 2A–C shows that Fe2+ in all of the contaminated samples decreased over time due to precipitation of Fe oxyhydroxide

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2500

10

Winter

2000

DO (mg/L)

Fe

2+

(mg/L)

8

1500 1000 150 100

6 4 2

50

Winter

0

0

0

100 200 300 400 500 600

1200

1800

0

2400

100 200 300 400 500 600

B

F

2500

Spring

1000 150 100

10

6 4

Spring

0

0

0

100 200 300 400 500 600

1200

1800

2400

0

100 200 300 400 500 600

Time (hours)

DO (mg/L)

(mg/L) 2+ Fe

2400

8

1500 1000 150 100

6 4 2

Summer 0

0 0

100 200 300 400 500 600

1200

1800

0

2400

100 200 300 400 500 600

H

2500

HWO and LP

2000

1200

1800

2400

Time (hours)

Time (hours) 10 8

1500

DO (mg/L)

(mg/L)

1800

G 10

Summer

50

2+

1200

Time (hours)

2500 2000

Fe

2400

2

50

D

1800

8

1500

DO (mg/L)

Fe

2+

(mg/L)

2000

C

1200

Time (hours)

Time (hours)

1000 150 100

6 4 2

50

HWO and LP

0

0

0

100

200

300

400

Time (hours)

0

100

200

300

400

500

600

700

800

Time (hours)

AMS-1: AMD spring – Michelle Creek

GW-1: Groundwater – Huntsville Gob

MC-2: 5 m downstream of AMD confluence with Michelle Creek MC-5: 800 m downstream of Michelle Creek headwater MC-7: 1200 m downstream of Michelle Creek headwater

GW-2: Groundwater – Huntsville Gob Stream water – Highway O (HWO) Stream water – Little Piney (LP)

Fig. 2. Temporal variation of Fe2+ and dissolved O2 (DO) of AMD-contaminated and uncontaminated samples.

phases (e.g., Fe(OH)3(s); FeOOH). As expected, the pH of the contaminated samples (except GW-1) decreased with de-

creases in Fe2+ due to proton production (reaction (4); Fig. 3A–C). Unlike the other contaminated samples, the

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E.W. Fonyuy, E.A. Atekwana / Applied Geochemistry 23 (2008) 2634–2648

A

D

10

10

Winter

Winter 8

pH

pH

8

6

6

4

4

2

2 0

50

100

0

150 1000 1500 2000 2500

2000

6000

8000

E

10

10

Spring

Spring 8

pH

pH

8

6

4

6

4

2

2

0

50

100

150 1000

1500

2000

2500

0

2000

4000

6000

8000

10000

SO42- (mg/L)

Fe2+ (mg/L)

C 10 Summer

F

10

Summer

8

8

pH

pH

10000

SO42- (mg/L)

Fe2+ (mg/L)

B

4000

6

4

6

4

2

2 0

50

100

150 1000 1500 2000 2500

Fe2+ (mg/L) AMS-1: AMD spring – Michelle Creek MC-2: 5 m downstream of AMD confluence with Michelle Creek MC-5: 800 m downstream of Michelle Creek headwater

0

2000

4000

6000

8000

10000

SO42- (mg/L) MC-7: 1200 m downstream of Michelle Creek headwater GW-1: Groundwater – Huntsville Gob GW-2: Groundwater – Huntsville Gob

Fig. 3. Cross plots of Fe2+ vs. pH and SO2 4 vs. pH for AMD-contaminated samples.

pH of GW-1 increased over time to a value of about 8.0 as the concentration of Fe2+decreased below the detection limit (<0.01 mg/L). Oxidation of Fe2+ in the contaminated samples was accompanied by a continuous increase in DO (Fig. 2E–G) and Eh (Table S-1). Decreased consumption of O2 with declining Fe2+ concentrations resulted in the accumulation of dissolved O2 and an increase in Eh. With nearly complete removal of Fe2+ from solution, DO concentrations approached stable values and Eh began to decrease late in the experiment. In contrast to the contaminated samples, uncontaminated samples (HWO and LP) with very low concentrations of Fe2+ (<0.3 mg/L;

Table 1) showed only slight increases in DO (Fig. 2H) and pH, and decreases in Eh (Table S-1). Water samples collected for the experiment were unfiltered, thus the oxidation of colloidal Fe-sulfide particles (e.g., reaction (1)) could also cause a decrease in pH (e.g., Chapman et al., 1983; Perez del Villar et al., 2005). The relationships between Fe2+ vs. pH and SO2 4 vs. pH (Fig. 3) were examined to evaluate if observed decreases in pH in the contaminated samples were due to reactions (3) and (4). With the exception of GW-1, all contaminated samples showed a decrease in Fe2+ that was accompanied by decreases in pH (Fig. 3A–C). The SO2 4 is nearly invariant with

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E.W. Fonyuy, E.A. Atekwana / Applied Geochemistry 23 (2008) 2634–2648

pH changes (Fig. 3D–F) suggesting that SO2 4 is conserved during the chemical evolution of AMD-contaminated samples (e.g., Sullivan and Drever, 2001; Yu and Heo, 2001). The relationships between Fe2+ vs. pH and SO2 4 vs. pH suggest that the decrease in pH for the contaminated samples can not be explained by the oxidation of pyrite or other colloidal Fe-sulfide particles. If this had been the case, decreases in pH should have been concomitant with in2+ 3+ creases in SO2 hydrolysis is used 4 and Fe . Although Fe to illustrate acid production in the contaminated samples, hydrolysis involving Al3+ and Mn2+ (e.g., Morin et al., 1988; Blodau, 2006) also produces protons that decrease pH, producing a similar effect on DIC as Fe3+. 3.2. Reaction mechanism and isotopic fractionation during DIC loss in contaminated samples 3.2.1. Reaction mechanism during DIC evolution The initial alkalinity and DIC concentrations in contaminated and uncontaminated samples are shown in Table 1. It is noted that GW-2 had a pH of 3.6–3.8, and despite no measurable alkalinity, the DIC concentration was relatively high (172–212 mg C/L). Alkalinity and DIC were lower in contaminated stream samples compared to groundwater and uncontaminated stream water (Table 1). Compared to uncontaminated stream samples, contaminated groundwater had up to 3 times more alkalinity and almost 9 times as much DIC. The higher alkalinity and DIC in the contaminated groundwater and spring samples resulted from the dissolution of carbonate minerals in waste rock, abandoned mine cavities, and in the aquifer (see reaction (2)). In the reactors, both alkalinity and DIC for the contaminated samples decreased over time (Fig. 4A–C and E–G). This is in contrast to uncontaminated stream samples HWO and LP that showed relatively little variation in alkalinity and DIC (Fig. 4D and H). When the carbonate ion concentration is low (based on the pH of the samples), decreases in DIC by CO2(g) loss to the atmosphere can be described by the following equation:

HCO3 þ Hþ $ H2 CO3 $ H2 O þ CO2ðgÞ

ð5Þ

In the contaminated samples, reaction (5) is driven to the right by the consumption of protons produced from the hydrolysis of metals (e.g., reaction (4)). The loss of DIC as CO2(g) is dependent on sample pH, temperature and atmospheric pCO2, as well as the degree of agitation and aeration of the sample. Time-series plots of alkalinity and DIC for the contaminated samples show an exponential decrease (Fig. 4A–C and E–G), whereas the uncontaminated samples are relatively linear (Fig. 4D and H). The reaction mechanism for exponential decreases in DIC concentration over time is governed by first order kinetics:

ln f ¼ lnðC t =C 0 Þ ¼ kt

ð6Þ

where f is the fraction of residual DIC since the start of DIC transformation, Ct is the residual concentration of DIC at any time after the start of the experiment, C0 is the concentration of DIC at start of the experiment, k is the rate constant, and t is time.

Fig. 5 shows plots of ln(Ct/C0) vs. time for the contaminated and uncontaminated samples. The ln(Ct/C0) calculated for some samples measured in the laboratory were greater than zero due to the slight variability between DIC values measured in the field when the reactors were first loaded, and the first samples collected from the laboratory reactors. Despite this, the ln(Ct/C0) for contaminated samples are negatively correlated with time, and linear for uncontaminated samples. These results suggest that DIC transformation in the contaminated samples is governed by first order kinetics. A compilation of the least squares regression equations of ln(Ct/C0) vs. time and the rate constants (k) are presented in Table 2. The correlation coefficients (R2) for the regression fit to data for contaminated samples ranged from 0.83 to 0.99. The overall rates of DIC loss from contaminated stream samples decrease in the downstream direction (MC2 < MC-5 < MC-7) relative to AMS-1, which was the source of stream contamination (Table 2). The decreasing slopes of the regression equations (Table 2) can be attributed to smaller contributions of protons from Fe2+ transformation as the concentration of Fe2+ decreases in the downstream direction (Table 1). Alternatively, higher relative HCO 3 concentrations that consume protons produced from Fe3+ hydrolysis could result in lower slopes. The lowest slope of the regression line was observed for sample GW-1, which had the lowest Fe2+ concentration (32–61 mg/L) and the highest alkalinity (529–630 mg/L; Fig. 5E). The highest slope was observed for GW-2 which had the highest Fe2+ content (1882–1969 mg/L) and undetectable levels of alkalinity (Fig. 5F). The slopes of lines for contaminated stream samples are between these two extremes. In fact, critical examination of the slopes of the spring (AMS-1), stream (MC-2 and MC-5) and groundwater (GW-1) samples show a lower slope in the first 100–200 h which are followed by steeper slopes (Figs. 5A–C and F). It is deduced from these results that during the chemical evolution of AMD, the rate of DIC loss increases with higher relative Fe2+ concentrations and/or lower alkalinity, and decreases with lower Fe2+ concentrations and/or higher alkalinity (Fig. 5). The uncontaminated stream (HWO and LP) samples with low concentrations of Fe2+ (<0.3 mg/L) and initial alkalinity of 155–177 mg/L showed little change in DIC and nearly flat slopes (Fig. 5G). The effect of alkalinity on the rate of DIC loss is best illustrated by groundwater GW-1 with lower Fe2+ content and a slope which approaches that of uncontaminated samples. 3.2.2. Carbon isotope fractionation The d13CDIC for the contaminated and uncontaminated samples ranged between 12.9‰ and 3.9‰ and 12.7‰ and 11.6‰, respectively (Table 1). Despite different reaction mechanisms controlling DIC transformation (Fig. 5), the contaminated samples (except MC-5 and MC7) and uncontaminated samples are characterized by temporal enrichment in d13CDIC (Fig. 6). The d13CDIC of GW-1 was enriched by 15‰, GW-2 and AMS-1 were enriched by 1.0–4.0‰ and sample MC-2 was enriched by 1.3‰. In contrast, samples from MC-5 (except the winter samples) and MC-7 were generally characterized by depletion in d13CDIC of up to 7.0‰. For the uncontaminated stream

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800 700

E

400

DIC ( mg C/L)

A

Alkalinity (mg/L as CaCO3)

E.W. Fonyuy, E.A. Atekwana / Applied Geochemistry 23 (2008) 2634–2648

300

Winter

Winter

600 500 400 300 200

200

100

100 0

0

0

100 200 300 400 500 600

1200

1800

2400

0

100 200 300 400 500 600

800

F

400

300

Spring

700

500 400 300 200

100

100 0

0 100 200 300 400 500 600

1200

1800

2400

0

100 200 300 400 500 600

600

DIC (mg C/ L)

Alkalinity (mg/L as CaCO3)

G 400

Summer

700

1200

1800

2400

Time (hours)

800

500 400 300 200

Summer

300

200

100

100 0

0 0 100 200 300 400 500 600

1200

1800

2400

0 100 200 300 400 500 600

Time (hours)

H 400

800

1800

2400

HWO and LP

HWO and LP

700

1200

Time (hours)

600

DIC (mg C/L)

Akalinity (mg/L as CaCO3)

2400

200

Time (hours)

D

1800

Spring

600

0

C

1200

Time (hours)

DIC (mg C/L)

B

Alkalinity (mg/L as CaCO3)

Time (hours)

500 400 300 200

300

200

100

100

0

0 0

100

200

300

400

500

600

700

800

Time (hours)

0

100

200

300

400

500

600

700

800

Time (hours)

AMS-1: AMD spring – Michelle Creek

GW-1: Groundwater – Huntsville Gob

MC-2: 5 m downstream of AMD confluence with Michelle Creek MC-5: 800 m downstream of Michelle Creek headwater MC-7: 1200 m downstream of Michelle Creek headwater

GW-2: Groundwater – Huntsville Gob Stream water – Highway O (HWO) Stream water – Little Piney (LP)

Fig. 4. Temporal variations of alkalinity and DIC for AMD-contaminated and uncontaminated samples.

samples, the d13CDIC was enriched by 5.9‰ for HWO and 5.5‰ for LP.

Shifts in d13C are controlled by isotopic fractionation which depends on whether DIC transformation and/or loss

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E.W. Fonyuy, E.A. Atekwana / Applied Geochemistry 23 (2008) 2634–2648

A

E

1

AMS-1

0

1

-1

ln(Ct/C0)

-1

ln(Ct/C0)

GW-1

0

-2 -3 -4

-2 -3 -4

-5

-5

-6

-6 -7

-7 0

100

200

300

400

0

500

100 200 300 400 500 600

B

1

ln(Ct/C0)

ln(Ct/C0)

-1

-2 -3 -4

-2 -3 -4

-5

-5

-6

-6

-7

-7 0

100

200

300

400

500

0

100 200 300 400 500 600

Time (hours)

G

2400

0 1

ln (Ct/C0)

ln(Ct/C0)

1800

-1

MC-5

-1 -2 -3 -4

2 3 4

-5

5

-6

6

-7

HWO and LP

7

0

100

200

300

400

500

0

100

Time (hours)

200

300

400

500

Time (hours)

1

AMD-contaminated spring

MC-7

0

and stream water samples

-1

ln(Ct/C0)

1200

Time (hours)

1 0

D

2400

GW-2

0

-1

C

1800

1

F

MC-2

0

1200

Time (hours)

Time (hours)

AMD-contaminated groundwater (sample GW-2)

Winter samples

Spring

-2

Spring samples

Summer

-3

Summer samples

-4

AMD-contaminated groundwater (sample GW-1)

Uncontaminated stream water

-5 Winter samples

-6

LP HWO

Spring samples

-7 0

100

200

300

400

500

Summer samples

Time (hours) Fig. 5. Time-series plot of ln(Ct/C0) for AMD-contaminated and uncontaminated samples.

from samples occur by equilibrium or kinetic processes. Since DIC loss from the contaminated samples occurred via first order kinetics, the isotope fractionation in these samples can be modeled as a Rayleigh distillation process (e.g., Monson and Hayes, 1980; Galimov, 2006):

d13 Ct ¼ d13 C0 þ ð1  f Þe

ð7Þ

where d13Ct is the calculated C isotope ratio, e is the isotopic enrichment factor (given as: [(aproduct/areactant)  1]  103). Substituting (Ct/C0) which is the ratio of DIC

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E.W. Fonyuy, E.A. Atekwana / Applied Geochemistry 23 (2008) 2634–2648 Table 2 Rate constants (K), correlation coefficients (R2), and regression equations for DIC loss from AMD-contaminated and uncontaminated samples K (mg C/L h)

R2

n

Regression equation

AMD-contaminated water samples Winter AMS-1 Spring AMS-1 Summer AMS-1 Summer MC-2 Winter MC-5 Spring MC-5 Summer MC-5 Winter MC-7 Spring MC-7 Summer MC-7 Winter GW-1 Spring GW-1 Summer GW-1 Spring GW-2 Summer GW-2

0.015780 0.018408 0.014530 0.015717 0.011700 0.016735 0.012337 0.009520 0.012092 0.008787 0.001350 0.002013 0.002111 0.017356 0.013512

0.97 0.98 0.95 0.99 0.94 0.95 0.95 0.96 0.99 0.83 0.94 0.97 0.95 0.97 0.95

10 16 10 11 9 14 10 9 13 8 22 25 12 12 13

y = 0.015780  +0.41257 y = 0.018408  +0.30382 y = 0.015300  +0.67774 y = 0.015717  +0.28724 y = 0.011700  +0.28491 y = 0.018400  +0.30812 y = 0.012337  +0.01113 y = 0.009520  0.41475 y = 0.012092  +0.05221 y = 0.008787  0.17756 y = 0.001300  0.20255 y = 0.002013  0.21672 y = 0.002111  0.13423 y = 0.017356  +0.58236 y = 0.013512  +0.72674

Uncontaminated water samples Spring HWO Spring LP

na na

na na

na na

na na

Season

Sample ID

na = not applicable because DIC loss was by equilibrium exchange with atmospheric CO2.

A

C

6

6 3

δ13CDIC (0/00)

δ13CDIC (0/00)

3 0 -3 -6 -9

Winter

-12

0 -3 -6 -9

Summer

-12

-15

-15 0 100 200 300 400 500 600 1200

1800

2400

0 100 200 300 400 500 600

B

D

6 3

1800

2400

6 3

δ13CDIC (0/00)

δ13CDIC (0/00)

1200

Time (hours)

Time (hours)

0 -3 -6 -9

Spring

-12

0 -3 -6 -9 -12

-15

LP and HWO

-15 0

100 200 300 400 500 600

1200

1800

2400

Time (hours)

0

100

200

300

400

500

600

Time (hours)

AMS-1: AMD Spring - Michelle Creek

GW-1: Groundwater- Hunstville Gob

MC-2:

GW-2: Groundwater- Hunstville Gob

5m downstream of AMD confluence with stream water -Michelle Creek

MC-5 : 800m downstream of AMD Spring- Michelle Creek MC-7:

Stream water- Highway O (HWO) Stream water- Little Piney (LP)

1220m down stream of AMD Spring - Michelle Creek

Fig. 6. Temporal variations of d13CDIC in AMD-contaminated and uncontaminated samples.

700

800

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E.W. Fonyuy, E.A. Atekwana / Applied Geochemistry 23 (2008) 2634–2648

at any time to the beginning of the experiment for 1  f, then:

d13 Ct ¼ eðC t =C 0 Þ þ d13 C0

ð8Þ

Eq. (8) defines a straight line where the slope of the line e gives the 13C enrichment factor assuming a constant fractionation between the reactant and product. Although the transformation of DIC in the AMD-contaminated samples can be modeled as a Rayleigh distillation, the reactor system does not completely fulfill the requirements and/or boundary conditions for the classic formulation. For instance, the DIC to CO2(g) loss mechanism involves multiple steps including (1) dehydration of HCO 3 to CO2(aq) and conversion of CO2(aq) to CO2(g), (2) physical mixing of initial CO2(aq)i in the samples before the start of the experiment with CO2(aq)H+ produced by HCO 3 dehydration, and (3) loss of different proportions of CO2(aq)i and CO2(aq)H+ (Fig. 7A). In addition, reactant (HCO 3 ) and product (CO2(aq)H+) measured as DIC were not physically or chemically separate. The existence of product (CO2(aq)i) as part of the system prior to the start of the experiment means that it began with a reactant reservoir that had already undergone some degree of isotopic change due to CO2 production and/or exchange with the atmosphere. Despite these difficulties, AMD-contaminated samples can still be described by a Rayleigh model if it is assumed that (1) the enrichment factors associated with HCO 3 —CO2ðaqÞ ðea ) and CO2(aq)–CO2(g) (eb) are constant (Fig. 7A), and (2) the enrichment rate (distinct from the enrichment factor) of DIC is a simple summation of the enrichment rates of all processes affecting the isotope composition of DIC. The 13C enrichment rates for contaminated samples (Table 3) can be determined by linear regression from cross plots of Ct/C0 vs. d13CDIC (Fig. 8). To develop the least squares regression equations for contaminated samples, d13CDIC data where temporal enrichment of d13CDIC reversed to depletion for AMS-1, MC-2, and MC-5 near the end of the experiment were excluded (Fig. 8A–C). Data

A

AMD contaminated samples

CO2(g)

ε

b

HCO3-

+ 2H+

ε

a

H2O +

CO2(aq)H+ CO2(aq)i

B

Uncontaminated samples

CO2(g)

HCO3-

ε

ε

CO2(aq)i

Fig. 7. Schematic of a conceptual model of DIC species distribution and isotope enrichment steps in AMD-contaminated samples and uncontaminated samples.

for GW-1 and GW-2 where the rate of d13CDIC enrichment increased dramatically at low DIC concentrations were also excluded (Fig. 8E and F). It is suspected that processes unrelated to AMD-DIC chemical evolution dominate changes in the C isotopic composition at such low DIC concentrations. It is not possible to discern the processes occurring at very low DIC concentrations. Nevertheless, if the chemical evolution of AMD controlled DIC loss and 13 C enrichment rate, contaminated stream samples would be expected to show similar relationships as observed for AMS-1. Samples from MC-2 and MC-5 (winter) show well defined d13C enrichment rates, whereas samples from MC-5 in spring and summer and all samples for MC-7 show depletion (Table 3; Fig. 8B–D). The enrichment of 13C in uncontaminated samples results from DIC equilibration with atmospheric CO2 and is defined by e1 and e2 (Fig. 7B). Thus, if DIC evolution in the contaminated samples was solely controlled by equilibration with atmospheric CO2, a nearly invariant change in DIC with increase in d13CDIC would be observed as shown for uncontaminated samples (Fig. 8G). None of the contaminated samples exhibit the behavior of HWO and LP. At the pH range measured for the contaminated samples, DIC consists of HCO 3 and CO2(aq). Equilibrium isotopic fractionation results in a 13C enrichment factor (e1) of 9‰ for e13 CHCO3 —CO2 ðaqÞ and e2 = 1.1‰ for e13 CCO2 ðaqÞ—CO2 ðgÞ at 25 °C (Mook et al., 1974). The rate of isotopic enrichment in contaminated samples showed good positive correlations which ranged from 1.0‰ to 2.6‰ for AMS-1, MC-2, and MC-5 (winter), 1.0‰ to 1.1‰ for GW-2, and 7.6‰ to 9.3‰ for GW-1 (Table 3). The 13C enrichment rate for GW-2 which had no measurable alkalinity was nonetheless controlled by CO2 loss. This continuous DIC loss suggests that CO2 flux in the reactors is unidirectional, and therefore that the CO2(aq) is subject to kinetic isotopic fractionation effects (Usdowski and Hoefs, 1990). The enrichment rate (e2) associated with this sample is 1.1‰ for kinetically controlled CO2 loss. The contaminated samples AMS-1, MC-2 and MC-5 (winter) have relatively small 13C enrichment rates (ea + eb = 1.0–2.6‰), suggesting a lack of isotopic equilibration between HCO 3 and atmospheric CO2. In order for the 13C enrichment rate to be dominated by e2 for diffusive loss of CO2(g), the CO2(aq) likely equilibrates with HCO 3 following its dehydration, thereby minimizing isotope enrichment. Alternatively, the inability to separate the reactant and product measured as DIC only allows for the determination of the isotopic fractionation resulting from diffusive loss from or exchange of CO2(g) with samples in the reactors. Because the 13C enrichment rates are defined by positive relationships with DIC, the enrichment of d13CDIC for the contaminated samples can be described by an equilibrium closed system model (equilibration of reactants and products during reaction) for DIC isotopic evolution (Gat and Gonfiantini, 1981; Kendall and McDonnell, 1998). In the reactors, the HCO 3 converted to CO2(g) was lost from the contaminated samples because of the progressively decreasing pH and there was limited exchange of DIC with atmospheric CO2. In AMD-contaminated samples, the extent of HCO 3 dehydration to CO2(aq) and exchange with atmospheric CO2 is governed by the HCO 3 concentration relative to

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E.W. Fonyuy, E.A. Atekwana / Applied Geochemistry 23 (2008) 2634–2648

Table 3 Carbon isotope enrichment rates, correlation coefficients (R2) and regression equations for Ct/C0 vs. d13CDIC for AMD-contaminated and uncontaminated samples Isotopic enrichment rate (‰)

R2

n

Regression equation

AMD-contaminated water samples Winter AMS-1 Spring AMS-1 Summer AMS-1 Summer MC-2 Winter MC-5 Spring MC-5 Summer MC-5 Winter MC-7 Spring MC-7 Summer MC-7 Winter GW-1 Spring GW-1 Summer GW-1 Spring GW-2 Summer GW-2

2.31 2.57 2.13 1.02 1.82 na na na na na 7.60 9.30 7.84 1.07 0.99

0.84 0.92 0.94 0.94 0.68 na na na na na 0.87 0.95 0.97 0.62 0.86

10 16 10 11 9 na na na na na 22 25 12 12 13

y = 2.3065  9.5661 y = 2.5714  9.0910 y = 2.1351  9.2346 y = 1.0215  6.6625 y = 1.8239  4.4424 na na na na na y = 7.5466  2.6012 y = 9.3205  2.9876 y = 7.8420  4.4447 y = 1.0689x  12.1091 y = 0.9889x  11.6430

Uncontaminated water samples Spring HWO Spring LP

na na

na na

na na

na na

Season

Sample ID

na = not applicable because sample did not evolve via bicarbonate dehydration or sample showed C isotope depletion.

the concentration of protons produced by Fe3+ hydrolysis (Fonyuy and Atekwana, 2008). In all of the contaminated samples with small enrichment rates (<3‰), the HCO 3 reservoir was exhausted before the pool of Fe2+. The opposite was observed in GW-1 (ea + eb = 7.6–9.3‰), where the pool of Fe2+ was exhausted leaving an excess of HCO 3 (Figs. 2 and 4). The higher enrichment rate observed for GW-1 can be explained by invoking a model whereby CO2 is also lost to the atmosphere in response to proton production from Fe3+ hydrolysis. However, the HCO 3 was able to equilibrate with atmospheric CO2 when Fe2+ was exhausted, so that the observed isotopic enrichment was also controlled in part by HCO 3 equilibration with atmospheric CO2. DIC equilibration with atmospheric CO2 in the absence of proton generation by Fe cycling accounts for the sharp positive incursion of d13CDIC late in the experiment (Fig. 8E). This chemical response is similar to trends observed in the uncontaminated samples (Fig. 8G). Temporal depletion in the d13CDIC was observed in some of the contaminated samples (Fig. 6; Table S-1). Enrichment of d13CDIC should occur if the AMD-produced protons were the driving force for DIC decreases. Interestingly, the temporal depletion in the d13CDIC occurred at low DIC (<1 mg C/L) where DIC was entirely CO2(aq). Below the alkalinity stability (pH 4.5) no new CO2(aq) is produced by the chemical evolution of AMD. Thus, DIC loss must proceed by either pH induced loss (i.e., rapidly decreasing pH) or equilibration with atmospheric CO2. Equilibrium exchange of CO2(aq)–CO2(g) should deplete the aqueous phase by 1.1‰ (e13 CCO2 ðaqÞ—CO2 ðgÞ = 1.1‰ at 25 °C; Mook et al., 1974). The negative excursion of d13CDIC was up to 7.0‰, suggesting a cumulative isotopic effect due to changing pH, or that additional mechanisms could be responsible for the depletion of d13CDIC. Possible sources for isotopically depleted CO2 in aquatic systems include microbial respiration or photochemical oxidation of dissolved organic C. Chapman et al. (1983) and Sanchez-Espana et al.

(2007) have suggested that oxidation of Fe2+ to Fe3+ in AMD-contaminated water is catalyzed by bacteria due to the fast rate of the reaction. Microbes in AMD derive energy by transfer of electrons between dissolved metals and organic C compounds (Rowe et al., 2007) and produce 13 C-depleted CO2 from organic C (e.g., Blair et al., 1985), causing 13C-depletion of DIC in contaminated samples. Alternatively, enhanced photo-oxidation of organic matter in acidic waters (Gennings et al., 2001; Friese et al., 2002; Wu et al., 2005) would likely produce CO2 depleted in 13C. Mass balance constraints indicate that the bulk d13CDIC of the samples would then decrease (e.g., Opsahl and Zepp, 2001). It is unclear from the results to what extent microbial respiration or photo-oxidation might contribute to the 13 C depletion, as these sources are expected to add CO2 to a pool with continuously decreasing DIC and thus requires further experimentation. 3.2.3. Effect of variable AMD contamination on DIC loss and carbon isotope fractionation The purpose of collecting different types of AMD-contaminated samples (groundwater, spring, and stream water) over different seasons was to ensure that the effects of variable chemistry and seasonal hydrology on AMD chemistry were captured. The chemical variability which manifests as variations in specific conductance or the concentrations of major cations did not affect the DIC loss mechanism (Fig. 5). The DIC loss was driven by protons from the chemical evolution of AMD and governed by first order kinetics in all samples. The DIC loss rate constants depend on the relative concentrations of alkalinity and Fe2+ in addition to the dominant form of DIC (HCO 3 or CO2(aq)) in the samples. The effects of AMD on DIC loss and C isotope fractionation resulted in a narrow range of 13C enrichment values (1–3‰). Higher 13C enrichment values (>4–8‰) were observed in samples with higher relative alkalinity to Fe2+, such that DIC was able to exchange C with atmospheric

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E.W. Fonyuy, E.A. Atekwana / Applied Geochemistry 23 (2008) 2634–2648

Winter

AMS-1

Spring

-8

0

Summer

0

δ CDIC ( /00)

E

-6

δ CDIC ( /00)

A

13

13

-10

-12

6

Winter

4

GW-1

2

Spring

0

Summer

-2 -4 -6 -8 -10 -12 -14

-14 1.2

1.0

0.8

0.6

0.4

0.2

1.2

0.0

1.0

0.8

Ct /C0

B

F

-6

0.2

0.0

Spring

0.2

0.0

GW-2

Summer

0

δ CDIC ( /00)

0

δ CDIC ( /00)

0.4

-6

-8

-8

-10

13

13

-10

Winter

-12

-12

Spring

MC-2

Summer

-14

-14 1.2

1.0

0.8

0.6

0.4

0.2

1.2

0.0

1.0

0.8

G

0

-6

0.4

HWO and LP

MC-5

Winter

-2

0.6

Ct/C0

Ct/C0

C

0.6

Ct /C0

-4

δ CDIC ( /00)

Summer

-9

0

-6 -8

13

13

0

δ CDIC ( /00)

Spring

-10

-12

LP HWO

-12 -14

-15 1.2

1.0

0.8

0.6

0.4

0.2

0.0

1.4

1.2

1.0

Ct /C0

D

AMD contaminated spring

0

MC-7

0.4

0.2

0.0

AMD contaminated groundwater (sample GW-2)

and stream water samples

-4

Winter samples

Spring

Spring samples

Summer

Summer samples

-6

Winter

-8

13

0

0.6

Ct /C0

-2

δ CDIC ( /00)

0.8

AMD contaminated groundwater (sample GW-1)

Spring

-10

Winter samples

Summer

Spring samples

-12

Uncontaminated stream water LP HWO

Summer samples

-14 1.2

1.0

0.8

0.6

0.4

0.2

0.0

Ct /C0

Regression line for winter samples Regression line for spring samples Regression line for summer samples

Fig. 8. Cross plots of d13CDIC vs. Ct/C0 in AMD-contaminated and uncontaminated samples.

CO2 (Fig. 8). Samples that had low Fe2+ and no alkalinity primarily showed depletion in 13C, which may in part be attributed to atmospheric exchange, microbial respiration, or photochemical oxidation of organic matter. The reaction time (400–2100 h) for the laboratory samples is much longer than the expected residence time of

water in most headwater watersheds impacted by AMD. The transformation of Fe2+ in the reactors was limited by the rate of O2 supply. If the objective was to increase the reaction rate, the samples could have been agitated continuously to supply more O2. In field settings, the pH of samples 1200 m from the point of AMD discharge can be as low

E.W. Fonyuy, E.A. Atekwana / Applied Geochemistry 23 (2008) 2634–2648

as 3.0 suggesting extensive production of protons from the chemical evolution of AMD (Fonyuy and Atekwana, 2008). Despite greater aeration and agitation by wind in field settings, the laboratory experiment suggests that chemical evolution of AMD would be the driving force for DIC loss and C isotope fractionation. Superimposed on this would be the effects of all other in-stream process that affect metal and C cycling. 4. Conclusions The main observations in this experiment were: (1) AMD-contaminated samples exposed to air had higher rates of CO2 loss compared to uncontaminated samples, and (2) the direct effect of AMD contamination was enrichment of d13CDIC. Contaminated groundwater and stream samples had variable amounts of Fe2+, alkalinity and DIC. The effects of proton production during the chemical evolution of AMD caused loss of DIC in contaminated samples which could be effectively modeled by first order kinetics. The reaction kinetics was distinguishable from uncontaminated samples in which the main process controlling DIC was C exchange with atmospheric CO2. Because the d13CDIC of both contaminated and uncontaminated samples was enriched over time, the measured d13CDIC in contaminated samples is not diagnostic of AMD-induced reactions. However, in combination with the DIC loss, the isotopic enrichment rates varied from 1‰ to 3‰, consistent with kinetic loss mechanisms. The attempt to model isotope fractionation using a simple Rayleigh model was successful although some boundary conditions for the model could not be fulfilled. However, it is clear that 13C enrichment of DIC in the contaminated samples could be modeled as a closed system equilibrium process which is distinct from equilibrium C exchange in DIC with atmospheric CO2 for uncontaminated samples. A depletion of 13C in contaminated samples at low DIC which is not related to HCO 3 dehydration was also observed. Further studies are required to understand the fractionation of C isotopes at low DIC and pH conditions. Acknowledgements This work was funded by the US National Science Foundation Award EAR-0510954. We thank H. Ali, A. Mukherjee, M. Moidaki, A. Hernan, A. Silva, T. Jerris and F. Ekollo for field help. We also thank A.L. Mayo and an anonymous reviewer for useful comments that helped improve this manuscript. Appendix A. Supplementary material Supplementary data associated with this article can be found, in the online version, at doi:10.1016/j.apgeochem. 2008.05.012. References Atekwana, E.A., Krishnamurthy, R.V., 1998. Seasonal variations of dissolved inorganic carbon and d13C of surface waters: application of a modified gas evolution technique. J. Hydrol. 205, 265–278.

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