A graphical method to evaluate predominant geochemical processes occurring in groundwater systems for radiocarbon dating

Chemical Geology 318–319 (2012) 88–112

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Research paper

A graphical method to evaluate predominant geochemical processes occurring in groundwater systems for radiocarbon dating Liang-Feng Han a,⁎, L. Niel Plummer b, Pradeep Aggarwal a a Isotope Hydrology Section, Division of Physical and Chemical Sciences, Department of Nuclear Sciences and Applications, International Atomic Energy Agency, P.O. Box 100, A-1400 Vienna, Austria b U.S. Geological Survey, National Center, Mail Stop 432, Reston, VA 22092, USA

a r t i c l e

i n f o

Article history: Received 8 September 2011 Received in revised form 10 April 2012 Accepted 5 May 2012 Available online 12 May 2012 Editor: B. Sherwood Lollar Keywords: Groundwater chemistry Radiocarbon dating Graphical method Groundwater

a b s t r a c t A graphical method is described for identifying geochemical reactions needed in the interpretation of radiocarbon age in groundwater systems. Graphs are constructed by plotting the measured 14C, δ13C, and concentration of dissolved inorganic carbon and are interpreted according to specific criteria to recognize water samples that are consistent with a wide range of processes, including geochemical reactions, carbon isotopic exchange, 14C decay, and mixing of waters. The graphs are used to provide a qualitative estimate of radiocarbon age, to deduce the hydrochemical complexity of a groundwater system, and to compare samples from different groundwater systems. Graphs of chemical and isotopic data from a series of previously-published groundwater studies are used to demonstrate the utility of the approach. Ultimately, the information derived from the graphs is used to improve geochemical models for adjustment of radiocarbon ages in groundwater systems. © 2012 Elsevier B.V. All rights reserved.

1. Introduction If the 14C age of dissolved inorganic carbon (DIC) in groundwater was a simple function of radioactive decay of 14C, by knowing the rate of 14C decay and the initial 14C content, the age of the dissolved inorganic carbon in groundwater could be calculated from the measured amount of 14C, according to

t¼−

14 t 1=2 C ln 14 ln2 C

ð1Þ

0

where 14C and 14C0 are the measured and initial 14C, t1/2 is the halflife of 14C, and t is the groundwater age. However, estimation of the initial content, 14C0, is not straightforward, and often physical and chemical processes, in addition to radioactive decay, alter 14C0 in aquifers, so the determination of groundwater 14 C age is, in most cases, not as simple as represented by Eq. (1). In earlier studies Brinkmann et al. (1960), Münnich and Vogel (1962), Vogel and Ehhalt (1963), and others, recognized that the 13C content of DIC reflects the relative amounts of several carbon sources contributing to the groundwater. Ingerson and Pearson (1964) solved a carbon mass-balance model (similar to that developed by Tamers, 1967, 1975) using isotope mass balance equations for 13C and 14C to define the ⁎ Corresponding author. Tel.: + 43 1 2600 21762. E-mail address: [email protected] (L-F. Han). 0009-2541/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2012.05.004

adjusted 14C0. Mook (1972), Tamers (1975), Fontes and Garnier (1979), Evans et al. (1979), Salem et al. (1980), Eichinger (1983), and Fontes (1992) have applied models, termed “traditional adjustment models,” that account for some of the more important geochemical reactions that occur in groundwater systems. The traditional adjustment models are discussed in detail in Fontes and Garnier (1979), Fontes (1983), Fontes (1992), Plummer et al. (1994), Kalin (1999), Geyh (2000, 2005), and Plummer and Glynn (in press). However, many more geochemical and isotopic reactions that occur in groundwater systems are not accounted for by these traditional adjustment models (Glynn and Plummer, 2005; Plummer and Glynn, in press). Further, the traditional adjustment models usually are applied only to the final (geochemically evolved) water, that is, the water sample that is to be dated, in an attempt to estimate the value of 14C0 that has been adjusted for geochemical reactions, but not radioactive decay. This can lead to significant error (usually an old bias) in estimation of adjusted radiocarbon age if other reactions, that are not accounted for in the traditional adjustment model, affect the radiocarbon content of dissolved carbon along the flow path to the final well. To overcome these problems, extended geochemical mass-balance models for 14C dating, termed generalized geochemical adjustment models, have been developed (Plummer, 1977; Wigley et al., 1978, 1979; Plummer et al., 1983; Plummer et al., 1994; Parkhurst and Charlton, 2008; Coetsiers and Walraevens, 2009; El-Kadi et al., 2010; Blaser et al., 2010). These models consider a number of geochemical and isotopic effects such as dissolution of carbonate minerals, oxidation of organic material, mineral precipitation, CO2 degassing, production

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of methane, and isotopic fractionation between solid, gas, and aqueous phases. Given the available chemical, isotopic, and mineralogic data, computer models, such as NETPATH (Plummer et al., 1994), NetpathXL (Parkhurst and Charlton, 2008), and NETPATH-WIN (El-Kadi et al., 2010) find all the geochemical adjustment models that are consistent with the available data and compute the corresponding adjusted radiocarbon ages. Pearson (1992) introduced an iterative procedure for estimating uncertainty in adjusted radiocarbon age that results from uncertainty in chemical and isotopic data. These approaches provide the most complete analysis of potential models and therefore produce the best estimates for the initial 14C value (14C0) that has been adjusted for geochemical water–rock reactions occurring in the aquifer. As one begins to use the geochemical adjustment models, a conceptualization of the possible geochemical reactants in a groundwater system is needed. Thus, a qualitative examination of chemical and isotopic data from groundwater systems can help in recognizing reactions and processes that need to be included in generalized geochemical adjustment models, such as in NETPATH. Graphs are particularly suitable for recognizing complex relations among water samples, and for recognizing those samples that are consistent or inconsistent with these relations. They are superior to tables of values because clusters and trends in data can be readily distinguished. Graphs also are useful when several data sets are to be compared. Graphical methods for analysis of carbon isotopic and HCO3− data have been used previously in the study of groundwater systems. Most previous work used the relationship of δ 13C and 14C,

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some used the relationship of δ 13C and/or 14C relative to the quantity 1/[DIC] (e.g. Pearson and Swarzenki, 1974; Stute and Deak, 1989; Taylor, 1997). However, these earlier works often focused on the analysis of relatively simple reactions using the graphs. In this paper we present a graphical method for studying the changes in carbon isotopic composition and DIC concentration that can occur in reacting natural water systems. This method allows qualitative inspection of the measured data to recognize a number of effects that can modify 14C composition in water, and, in combination with the results obtained by the more advanced dating models, e.g. geochemical adjustment models found by using NETPATH, provides a better understanding of the systems under investigation and their radiocarbon ages. 2. The graphical method The basic elements and arrangement of the graphs are shown in Fig. 1. Graph I shows the relationship between δ 13C and reciprocal of dissolved inorganic carbon concentration ([DIC]), graph II shows the relationship between 14C and 1/[DIC], and graph III shows the relationship between 14C and δ 13C. The method of constructing the graphs is described in Appendix A. In Fig. 1, the points (A, O and B) represent end members in the system. At point A the water contains only CO2(aq) equilibrated with soil CO2 in an open system; at point O (the ‘crossing point’) the water has completely reacted with biogenic (soil) CO2 and carbonates 0 B (I) Z

-5

b 2

-10 Y -15

O

-20

a

3

1

A

-25 X -30 0

0.005

0.01

0.015

0.02

1/[DIC] 120

120 (III)

X 100

X

(II)

100

A a

80

A

80

a

4

5 1

1

60

60

Y

Y O

40

Z

Z

3

20

b

6

6 B

-25

-20

2

20

0 -30

O

40

-15

-10

-5

0

b

B

0 0

0.005

0.01 1/[DIC]

Fig. 1. Graphs for carbon data interpretation.

0.015

0.02

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in a closed system; and at point B the DIC in water is enriched in δ 13C and/or has very low 14C compared with the DIC at point O. In many cases the DIC at point B has δ 13C and 14C values close to zero (e.g. HCO3− from the reaction of CaCO3 with non-carbonic acids). Waters plotting around point O often are relatively young (young water in closed system). These waters normally contain detectable 3H (see Figs. 2, 6, 10, 14, and 18). Lines X and Y separate the graphs into different regions (see below). Line Z represents mixing of two carbon reservoirs (DIC) (Pearson and Swarzenki, 1974):
 ½DICA ⋅δA þ ½DICm −½DICA ⋅δB ¼ ½DICm ⋅δm

ð2Þ

where, [DIC] denotes the concentration of carbon isotope-bearing DIC, δ denotes carbon isotope content (amount of 14C or 13C), and subscripts A, B, m denote end member A, B, and mixing product m. After rearrangement Eq. (2) becomes: δm ¼

1 ½DICA ðδA −δB Þ þ δB : ½DICm

ð3Þ

A simple mixing effect (without 14C decay) can be displayed by plotting the measured carbon isotopic composition ( 14C or 13C content, δm) against the reciprocal of the measured DIC concentration. Thus, a sample with mixed carbon sources O and B in a closed system will plot on line Z in Fig. 1. Lines X and Z in Fig. 1 are process lines. A process line in the graph represents a process or several processes that have similar effects on the dissolved carbon concentration and carbon isotopic composition. Line X includes processes that occur at constant 1/[DIC] or constant δ 13C with or without 14C decay. In the graphs, a point on line Z may have very low 14C content, but because of its very high δ 13C value, this water could be relatively young. Although data around this line do not necessarily represent young waters, we refer to this line as the “zero-age” line because the variations in 14C, δ 13C and 1/[DIC] due solely to geochemical reactions without 14C decay can be represented by line Z. A set of samples that have undergone similar chemical and/or physical processes may plot in Fig. 1 in different regions marked with 1–6. Data plotting in region 1 may indicate water systems under open conditions with respect to CO2. These waters usually have zero age and will evolve through different paths (illustrated by arrows) to equilibrium. Coupled with geochemical modeling, the

0 B

(I) Z

-5

-10

Y B'

-15

-20 X -25 0

0.005

0.01

0.015

0.02

1/[DIC]

120

120 (III)

80

(pmc)

80

60 Y

60 Y 40

40

B

B'

0 -20

-15

-10

-5

Z

20

Z

20

-25

(II)

X 100

14C

14C

(pmc)

X 100

B'

0 0

0

0.005

0.01

0.015

0.02

1/[DIC] Fig. 2. The three samples represented by empty circles contain tritium. The arrows indicate chemical and isotopic evolution of the waters represented by solid circles. Letter B′ represents HCO3− from reaction of fossil organic matter with marine carbonates. Pearson and Hanshaw estimated the δ13C values for the CO2(aq) and mineral carbonate to be − 22 and 0.5‰, respectively. Data from Pearson and Hanshaw (1970).

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range of data points that plot in region 1 tells us something about the evolution of the water system under open system conditions (dotdashed arrows) and closed system conditions (full-line arrows) (see Sections 3.1.1 and 4.6). Data plotting in region 2 include processes that would decrease [DIC] and 14C, and increase δ 13C, such as dedolomitization involving dissolution of dolomite enriched in 13C and calcite precipitation. Data plotting in region 3 may indicate additional carbon introduced by old organic matter and the additional carbon is more depleted in 13C than the current soil CO2. In some cases data plotting in region 3 may indicate weathering of silicates by additional carbonic acid. These processes would increase [DIC] and, decrease δ 13C and 14C. Data plotting in region 4 may indicate additional carbon introduced by 14C-containing organic matter. Similarly, data plotting in region 5 may indicate the effect of fermentation of organic matter and the product methane contains 14C. Region 6 is typical for old waters. If some sample points plot in this region, these waters in most cases could have undergone 14C decay. In special cases, if the 14C–δ 13C relationship is a curve (see Appendix B, Fig. B2), this may indicate that during decay of 14C, δ 13C has more or less steadily increased with time. The increase in δ 13C can be caused by processes which can add carbon enriched in δ 13C such as further carbonate dissolution, adding CO2 of geogenic (abiogenic or thermogenic) origin, isotope exchange with solid carbonate, incongruent dissolution of carbonate, methanogenesis, etc. Many groundwater systems, including geothermal and/or carbonate aquifers, show a curved 14C vs. δ 13C relationship (e.g. Gonfiantini and Zuppi, 2003; Dennis et al., 1997; Deak, 1978, see Appendix B for explanation). At point A the water contains predominantly CO2(aq), further towards point O in region 1, the water contains both CO2(aq) and HCO3− and, at point O the predominant carbon species is HCO3−. Other carbon species (e.g. CO2(g), H2CO3, CO32 −) are negligibly small fractions compared with these two species. Therefore, for data points plotting left of line X in graphs I and II, [HCO3−] and [DIC] are nearly identical. For data points plotting right of line X, [DIC] data should be used. In some cases the data may plot along or close to line Y in subdiagram I with different [DIC]. Such patterns may occur if: 1) the samples are collected from different areas with different CO2 contents in Table 1 Processes represented by graphical elements and regions and their effects on [DIC],

13

the soil (in some regions, e.g. extremely arid, the primary carbon source in water may be atmospheric rather than soil CO2); 2) the samples are collected from different systems (e.g. in silicate aquifers the carbonate concentration may be low and in some karst systems the water moves fast into a groundwater system without contact with soil air); 3) the contribution of open and closed system conditions to chemical evolution of the water is different; 4) the samples may be mixtures of waters with different DIC contents; and 5) the CO2 content of the soil air has varied with time (in some cases, it may be necessary to construct Fig. 1 for modern and paleo conditions, given that the principal lines and crossing point could move with climatic conditions). Except for 14C decay, geochemical, physical and isotopic exchange processes will affect [DIC], 14C and 13C in groundwater; however, the processes may affect [DIC] and carbon isotope contents differently. Processes that can cause elevated [DIC] include: a) elevated soil CO2 partial pressure which is a function of temperature and precipitation (Brook et al., 1983); b) aerobic oxidation of organic matter forming CO2; c) dissolution of solid mineral carbonate due to decrease in alkaline-earth metal ions, caused by ion exchange between clay mineral and water; d) dissolution of solid mineral carbonate due to reaction of protons with carbonate, caused by ion exchange between organic matter and water; e) adding of CO2 of geogenic origin; and f) adding CO2 from organic matter through methanogenesis reaction. Processes c, d, e and f will have the same effect on the chemistry and carbon isotopes of the DIC: the 14C concentration is diluted by addition of carbonate which has a 14C value close to 0 pmc and enriched in 13C. In the case of processes a and b, however, the CO2 could contain 14C and be depleted in 13C. Within the variables 14C and δ 13C, there are four combinations: (i) increase in δ 13C and decrease in 14C content; (ii) decrease in δ 13C and decrease in 14C content; (iii) decrease in δ13C and increase in 14C content; and (iv) increase in δ13C and increase in 14C content. Besides these effects some processes may have variable change in δ 13C and 14C content. Table 1 is a summary of selected chemical and physical processes in groundwater systems represented by graphical elements and regions and their effects on [DIC], δ 13C and 14C content. (Most of the processes are taken from Plummer and Glynn, in press).

C and

14

C. Representative graphical elementsa

Process

1 2 3

4 5 6 7 8 9 10

11 12 13 14 15 16 a

91

Dissolution of marine carbonates by soil CO2 (open system) Dissolution of marine carbonates by geogenic CO2 (closed system) Oxidation of fossil organic matter (FOM) Case 1: δ13CFOM = δ13C of CO2,soil Case 2: δ13CFOM b δ13C of CO2,soil Case 3: incomplete reaction of CO2,FOM Methanogenesis involving old organic matter Ion exchange on clay mineral (carbonate dissolution) Dissolution of marine carbonates by non-carbonic acids Isotopic exchange between water and carbonate Isotopic exchange between water and soil CO2 (open system) Precipitation of carbonates Weathering of silicates (complete transformation of CO2 into HCO3−) Case 1: CO2 contains 14C Case 2: fossil CO2 Incongruent dissolution Dissolution of CaSO4 and dolomite with calcite precipitation (closed system). (Dedolomitization) Methanogenesis involving organic matter containing 14C (e.g. landfill) Isotopic exchange with carbonate that has higher δ13C and 14C content Loss of CO2 gas 14 C decay

Symbolb

Graph I

Graph II

Graph III

Region 1 Z

Region 1 Z

Region 1 Z

↑ ↑ ↑c ↑↑↓

Left of O Region 3 Region 3 Z Z Z X(b) X(a) Region 1

Z Z Z Z Z Z X(b) X(a) Region 2

X(b) Region Region Z Z Z Z Region Region

↑–↓ ↑↓↓ ↑↓↓ ↑↑↓d ↑↑↓ ↑↑↓ –↑↓ –↓↑ ↓↓↓

Region 3 Region 3 X(b) Region 2 Z X(b) Region 2 Point O

Region 4 Z X(b) Region 2 Region 4 X(a) Region 1 X(b)

Region Region Z Z Region Region Region X(b)

3 3

1 3 1 3

5 5 5

↑↓↑ ↑↓↓ –↑↓ ↓↑↓ ↑↑↑ –↑↑ ↓↑↑ ––↓

Assuming that only a single process has occurred. Change in [DIC], δ13C and 14C value, respectively (↑: increase; ↓: decrease; –: no change). Unless specifically stated the changes are relative to O point (i.e. assuming that the sample point was at Point O before changes in [DIC], δ13C and 14C value). c The changes are relative to Point A. d The line will cross at δ13C > 0 with the ordinate in graph I and abscissa in graph III. b

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Finally, it should be pointed out that in natural systems, usually multiple processes/reactions occur together, and then the trends in 13 C and 14C with DIC can go in many directions depending on extent of reactions. Resort to forward reaction modeling aids in interpreting systems in which multiple reactions occur simultaneously (see for example Parkhurst and Appelo, 1999). 3. Examples of data interpretation In this section the processes that may be affecting the 14C content of the water sample are briefly described followed by examples of graphical interpretations with geochemical and isotopic data taken from the literature. In all the graphs, the unit of dissolved inorganic carbon concentration is expressed as mg/L HCO3−, although in the literature the measured data are expressed differently (e.g. as alkalinity, [DIC] or [HCO3−]). For simplicity, in some figures the letters (X, Y, Z, O, B, etc.) and line Z are omitted. 3.1. Simple geochemical evolution 3.1.1. Effect of “dead carbon” introduced by dissolution of calcite and/or dolomite (processes 1 and 2 in Table 1) This is the most important process among all the processes that can change 14C concentration. The solubility of calcium carbonate in pure water is low (12.3 mg CaCO3/kg water at 25 °C). Given the common presence of soil gas CO2, water containing soil CO2 is acidic (point A in Fig. 1). Because calcium/magnesium carbonate will react with water that is saturated with CO2 to form soluble calcium/magnesium bicarbonate, the following reactions can introduce carbon into water: þ

−

CO2ðgÞ þ H2 O↔CO2ðaqÞ ↔H þ HCO3 :

ð4Þ

The weak acid will react with carbonate mineral: þ

−

2þ

H þ HCO3 þ MeCO3 ↔Me

−

þ 2HCO3

ð5Þ

where Me is calcium or magnesium. Because limestones and dolostones typically are old and no longer contain 14C, this process typically introduces 14C-free DIC into the system and increases δ 13C. There are of course always exceptions, such as dissolution of recently formed carbonates that contain 14C, but for illustration purposes, this discussion is limited to the more common cases. From point A water containing CO2(aq) can react with carbonate minerals through different paths to equilibrium. In Fig. 1, the long full-line arrows show the evolution path for a ‘pure’ closed system (i.e. the closed system conditions are encountered at point A). Under open conditions, if carbonate minerals are dissolved, more soil CO2 will go into water and the water will dissolve more carbonates (dot-dashed-line arrows) and, due to isotopic exchange with soil CO2 the isotopic composition of DIC depends only on isotopic fractionation between DIC and soil CO2. Because the relative content of the main carbon species in the DIC, i.e. the ratio CO2(aq)/[HCO3−], changes with progress of carbonate dissolution, the isotopic composition of DIC will also change. When an initially open system becomes closed to soil CO2, the water will be “neutralized” through reaction 5 (i.e. the short full-line arrows). The position of equilibration point (point O) depends on the point where closed system conditions are encountered. For a pure closed system, at point O, the DIC should have δ 13C and 14C values of about ½ of the initial CO2(aq), provided no other processes including 14C decay have taken place. On the other hand, if the open system phase is predominant during dissolution of carbonate minerals, due to isotopic exchange with soil CO2, the DIC at equilibrium may have lower δ 13C and higher 14C values. A water at point O can further evolve by receiving additional carbon through carbonate dissolution or losing carbon through calcite precipitation under closed system

conditions. Deines et al. (1974) discuss the evolution of 13C in carbonate systems open and closed to soil gas CO2.

3.1.2. Example Pearson and Hanshaw (1970) investigated sources of dissolved carbonate species in groundwater and their effects on 14C dating for a Floridan limestone system. In the study area there are many microclimates; for example, there are areas covered with sand dunes into which rainfall infiltrates rapidly, as well as extensive areas of lakes and swamps with large quantities of organic dieback. Two sets of samples were collected. From the recharge area, samples contain high tritium. A second group of samples were collected from wells farther down gradient in the confined portion of the aquifer and contain low 14C. The data are shown in Fig. 2. It is important to determine the location of the crossing point because from this point further development of water geochemistry can be recognized. The crossing point can be determined by the group of samples containing high tritium represented by empty circles (see Appendix A). These samples present a simple closed system evolution. Compared to the three young waters (empty circles) the three samples represented by solid circles have higher DIC and lower 14C concentrations, and they are enriched in 13C. This indicates that for these samples additional dead carbon enriched in 13C has been added to the system under closed system conditions. The addition of HCO3− thus would cause the solid circles to move along line Z to the areas enclosed by dashed ellipse (as indicated by the full-line arrows, symbol ‘↑ ↑ ↓’ in Table 1). Because in graph I these points are not on line Z, a second process can be assumed, namely, some additional carbon may have been added to the system that has a δ 13C value close to the soil CO2 (e.g. through sulfate reduction driven by oxidation of organic carbon; see Table 1, process 3). The additional HCO3− produced in the second process combined with carbonate-mineral dissolution contains no 14C and has a δ 13C value about 1/2 of the organic carbon value, as represented by letter B′. In graph I, this process would cause the points to move along the dashed line (as represented by the dashed-line arrow). Similarly, the effects of reacting organic matter are represented by the dashed line arrows in graphs II and III, showing that the sample points move toward B′. Without 14C decay these samples would plot in the shaded ellipses. Apparently, these samples have undergone 14C decay (as represented by dot-dashed line arrows) (symbol ‘– – ↓’). The ages of these waters could be estimated based on the corrected initial 14C content (initial 14C contents of about 20 pmc are assumed from the shaded areas in Fig. 2).

3.2. Oxidation of organic matter (process 3 in Table 1) 3.2.1. Effect of dead carbon introduced by organic matter Except for carbon from soil CO2, CO2 of geogenic origin, a water sample may contain carbonate species from oxidation of fossil organic material (Pearson and Hanshaw, 1970; Plummer, 1977): 2−

2CH2 O þ SO4

−

¼ H2 S þ 2HCO3

ð6Þ −

þ

6nO2 þ ðC6 H10 O5 Þn þ nH2 O→6nHCO3 þ 6nH 2−

−

−

SO4 þ CH4 ¼ HS þ HCO3 þ H2 O:

ð7Þ ð8Þ

Because carbon from these sources has a 14C value close to zero, adding all the above reaction products (reactions 6, 7, and 8) would lower 14C contents while increasing the HCO3− concentration. Data plotting in region 3 in Fig. 1 may indicate addition of old organic matter with δ 13C values lower than current soil CO2, or, due to weathering of silicates by CO2 produced from oxidation of fossil organic matter (see Section 3.7.1).

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3.2.2. Example 1 Buckau et al. (2000a) discussed the impact of climatic and vegetation conditions over the past 15 ka on the chemical composition and 14 C dating of groundwater from the Franconian Albvorland aquifer system. 14C analysis was carried out on both fulvic acid and DIC. 14C dating via fulvic acid gives groundwater ages consistent with climatic and vegetation records and variations in the groundwater composition. No correction for geochemical processes is required, because, under these geochemical conditions, fulvic acid remains stable over this time period and flow-distance. On the other hand, 14C dating via DIC requires correction of the 14C value due to perturbation by different geochemical processes. Fig. 3 shows the data from Buckau et al. (2000a). The crossing point is estimated from the empty circle which represents water that has evolved predominantly under open system conditions (see empty circles in Figs. 5 and 17). As can be seen from Fig. 3, after becoming closed, some additional carbon with 13C content close to the soil CO2, but without 14C, was added to the system (full-line arrows; symbol ‘↑ – ↓’). Most of the samples have undergone 14C decay (dashed line arrows). Buckau et al. found that in general the uncorrected 14C ages determined via DIC are greater than that determined via fulvic acid.

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3.2.3. Example 2 Buckau et al. (2000b) re-studied groundwater in the Gorleben aquifer system. Previous study had given 14C results with conflicting 3H and stable isotope data and hydrological estimates. 14C model ages of 1± 10 ka have been found for 3H containing recharge water, up to 31 ka for groundwater with Holocene stable isotope signatures and 6±10 ka for groundwater at 35 m depth. Buckau et al. show that the elevated DIC concentration in groundwater was caused by microbiologically mediated mineralization of organic components in deep sediments. For these enhanced DOC groundwaters, their study leads to groundwater ages in compliance with the age limits derived by the 2H and 18O signatures. Fig. 4 shows the data from Buckau et al. (2000b). Compared with Fig. 2, it can be seen that the DIC of the samples represented by solid circles has probably been diluted by dead carbon from fossil organic matter. The samples represented by solid triangles have higher δ 13C compared with those represented by solid circles. The higher δ 13C may indicate that these samples could contain additional carbon from non-biogenic sources (e.g. CO2 of geogenic origin, or reaction products between carbonate and non-carbonic acid e.g. H2S, see Eq. (6)). A third group of waters (empty symbols) plot below line Y in graph I and above line Y in graphs II and III. These samples have higher 14C (>50 pmc) and 3H (>20 TU), and lower pH values

0 (I) Z

-5

-10 Y -15

-20

-25 X -30 0

0.005

0.01

0.015

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1/[DIC]

120

120

80

80

(pmc)

100

(II)

60

14C

60

14C

(pmc)

X

(III)

X 100

Y

Y 40

40

Z

Z 20

20

0

0 -30

-25

-20

-15

-10

-5

0

0

0.005

0.01

0.015

0.02

1/[DIC] Fig. 3. The arrows denote the direction of chemical and isotopic evolution. Full-line arrows: effect of addition of carbon caused by oxidation of organic matter; dashed-line arrows: decay of 14C. The empty circle represents a water sample with zero age. Carbon isotopic and chemical data from Buckau et al. (2000a).

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0 (I) -5

-10

-15

-20

-25

-30 0

0.01

0.02

0.03

0.04

1/[DIC]

120

120

(II)

80

80

(pmc)

100

60

14C

14C

(pmc)

(III) 100

60

40

40

20

20

0

0 -30

-25

-20

-15

-10

-5

0

0

0.01

0.02

0.03

0.04

1/[DIC] Fig. 4. The dashed-line arrows denote the evolution of carbon chemistry and isotopes for the samples represented by triangles. The full-line arrows denote the direction of chemical evolution of the samples represented by solid circles. The samples represented by empty circles have lower dissolved organic carbon compared to the other samples. Carbon isotopic and chemical data from Buckau et al. (2000b).

(some have pH values as low as b6) indicating open system conditions (probably with organic acids). The three samples (solid squares) with 14C b 50 pmc and detectable 3H are most probably mixtures of old and young waters (see discussions in Section 4.2). Plotting of total dissolved carbon ([DIC + DOC] instead of [DIC] vs. δ 13C and 14C gives similar graphs, indicating that the amount of additional CO2 is proportional to that of the DOC. This would allow corrections for the effects of DOC based on the amount of the DIC.

et al., 2006), methylated amines (e.g. trimethylamine; Eq. (12)) (Oren, 2001), or methanol (Eq. (13)) (Oren, 2001): −

2CH3 SCH3 þ 3H2 O→3CH4 þ HCO3 þ 2H2 S þ H þ

−

3.3.1. Effect of fermentation of organic matter: methanogenesis (processes 4 and 13 in Table 1) The two most common methanogenic pathways in typical organicrich water sediments are reduction of carbon dioxide with hydrogen gas and acetoclastic methanogenesis (Conrad, 2005): CO2 þ 4H2 →2H2 O þ CH4

ð9Þ

CH3 COOH→CO2 þ CH4 :

ð10Þ

In addition to the above two reactions, there are other methanogenic pathways using substrates including methylated compounds (other than acetate) such as dimethylsulfide (Eq. (11)) (Van Leerdam

þ

ð11Þ

4ðCH3 Þ3 N þ 9H2 O þ H →9CH4 þ 3HCO3 þ 4NH4

ð12Þ

−

ð13Þ

þ

4CH3 OH→3CH4 þ HCO3 þ H2 O þ H :

3.3. Methanogenesis

þ

Due to methanogenesis the 14C reservoir will be isotopically changed. This is because 1) the reaction product HCO3− added to the system may or may not contain 14C; and 2) the acidic reaction products CO2, and H2S will react with carbonate minerals to further produce HCO3− which typically contains no 14C. Because CH4 typically is much depleted in 13C, the reaction product HCO3− will be enriched in 13C. In addition, the dissolved carbonate through reaction with CO2, H2S, and H + is also enriched in 13C, therefore, methanogenesis will increase δ 13C and change the 14C content of the DIC. The reaction lines in the case of methanogenesis will have different slopes compared with that of carbonate dissolution by soil CO2. In some cases methanogenesis may involve organic matter containing 14C (e.g. landfill). Data plotting in regions 4 and 5 in Fig. 1 may

L-F. Han et al. / Chemical Geology 318–319 (2012) 88–112

indicate methanogenesis involving organic matter containing 14C. This process would increase HCO3−, δ13C and 14C values together. 3.3.2. Example 1 In the work of Hackley et al. (1992) groundwater samples were collected from several different depths in Illinois glacial deposits at a site in east central Illinois, USA. Due to the formation of microbial CH4, some of the samples were unusually enriched in 13C, δ 13C values of − 3 to − 1‰. Isotopic and analytical results show a positive correlation between the δ 13C of the DIC and the concentration of methane in the groundwater. Fig. 5 shows the data from Hackley et al. (1992), plotting in three graphs. The samples with measureable tritium (represented with empty circles) have a pattern similar to that in Fig. 17 (see explanations there for similar samples; also see Section 4.6). As seen in Fig. 5, further dissolution of carbonate has taken place (triangles). The presence of methane in these samples indicates that the dissolution of carbonates is probably due to methanogenesis. Methane production can lead to CO2 more enriched in δ 13C compared to soil CO2. The dashed line shows the shifting of the data caused by methanogenesis from the crossing point (see footnote d in Table 1). From the vertical intersection with the zero-age line (dashed line on graph III), the age of the waters could be estimated (see age estimation for solid circles in Figs. 6 and 18).

95

3.3.3. Example 2 Aravena et al. (1995) studied a confined regional aquifer in southern Ontario, Canada, affected by methanogenesis. The 13C and 14C isotopic data for DIC, DOC, and CH4 clearly show the effects of methanogenesis occurring in the aquifer. Significant enrichment in 13C of DIC is observed along a groundwater flow path, changing from −11.6 to +4.8‰. This trend is accompanied by lowered 14C values, reflecting in part the input of older organic carbon to the DIC via methanogenesis (graph I′ in Fig. 5). 3.4. Effect of dead carbon introduced by further carbonate dissolution 3.4.1. Effect of dead carbon introduced by CO2 of magmatic origin Assuming that addition of geogenic CO2 to the system and dissolution of rock carbonate have taken place, namely (Crossey et al., 2009; James et al., 1999; Lesniak, 1985; Yamada et al., 2010): 2þ

CO2 þ H2 O þ MeCO3 →Me

−

þ 2HCO3

ð14Þ

carbon from geogenic origin CO2 is isotopically enriched in 13C relative to that from plants and contains only dead carbon. In this case the reaction product HCO3− in reaction (14) would have 14C value close to zero and increased δ 13C value (in many cases close to point B in Fig. 1), and the sample points affected by this process should 5

10

(I)

(I')

5

0 0 -5

-5

-10

-10 -15

-15 A'

-20

A -20

-25 -30

-25 0

0.005

0.01

1/[HCO3

0.015

0

0.02

0.005

-]

0.01

0.015

0.02

1/[HCO3- ] 120

120 A

100

(III)

(II)

A'

100

A

(pmc)

60

14C

(pmc)

80

14C

A' 80

60

40

40

20

20

0

0 -25

-20

-15

-10

-5

0

5

0

0.005

0.01

0.015

0.02

1/[HCO3- ] Fig. 5. The dashed line shows the shifting of the data caused by methanogenesis from the crossing point. Triangle: water with CH4; Empty circle: water with tritium. The arrows and point A and A′ are discussed in Section 4.6. In graph I′ the empty circles represent water samples with negligible CH4 concentrations. Carbon isotopic and chemical data from Hackley et al. (1992). Graph I′ shows data from Aravena et al. (1995).

96

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0 (I) -2 -4

Z

-6 Y -8 -10 -12 X -14 0

0.005

0.01

0.015

0.02

-

1/[HCO3 ]

120

120 (III)

X

80

80

(pmc)

100

(pmc)

100

14C

60

14C

(II)

X

Y

60 Y

40

40

Z

Z 20

20

0

0 -14

-12

-10

-8

-6

-4

-2

0

0

0.005

0.01

0.015

0.02

1/[HCO3- ] Fig. 6. Empty circles represent samples that contain detectable tritium. Graphical representation of data from Bayari et al. (2008).

plot on line Z in Fig. 1. The positions of the points depend on the amount and δ 13C value of the CO2 added and the 13C of the carbonate minerals. The more CO2 added, the closer the δ 13C value to zero and the closer the points will be to point B in Fig. 1 (assuming that no other processes which modify carbon content have occurred). 3.4.2. Effect of cation and/or proton exchange Cation exchange on clay minerals and/or organic matter may take place. When this exchange takes place on clay minerals (R1), alkalineearth ions typically are removed from the solution, and thus more carbonate can dissolve (R2) until saturation with respect to carbonate is re-established: −

R1

þ ðclay−2NaÞs → þ H2 O



þ

2HCO− 3 2þ

2Na Me



þ ðclay−MeÞs R2 ← ðMeCO3 Þs − − þ HCO3 þ OH

3.4.3. Effect of reactions with non-carbonic acids Dissolution of rock carbonate can also take place if non-carbonic acids present, e.g.: −

2þ

R  COOH þ CaCO3 →HCO3 þ Ca H2 S þ

− CaCO3 →HCO3

2þ

þ Ca

−

−

þ R  COO

þ HS :

ð16Þ ð17Þ

The introduced HCO3− will have δ 13C and 14C values close to that of the rock carbonate. For a closed system, samples with additional carbon produced by reactions 16 and 17 will plot on line Z in Fig. 1.

2þ

2HCO3 þ Me

releasing both protons and sodium ions. The released protons (acid) will react directly with old carbonates, if present, further diluting the 14C reservoir. For a closed system, samples having undergone cation and/or proton exchange will plot on line Z in Fig. 1.

ð15Þ

where, the subscript “s” represents solid phase. In such a case the concentration of DIC will increase. Because HCO3− is from the dissolution of solid carbonate, δ13C will also increase and 14C value typically will decrease. In the case of proton exchange (Appelo, 1994; Plummer et al., 1994), calcium and magnesium are taken up on organic matter

3.4.4. Example Bayari et al. (2008) studied age distribution of groundwater in the Konya Closed Basin, central Anatolia, Turkey. In this area strong carbon dioxide discharges encountered in groundwater drillings seem to be the end products of Holocene volcanism. The 13C content of the CO2 gas has been determined to be ca. 0‰ VPDB. The area is also characterized by the presence of numerous gigantic collapse dolines. These

L-F. Han et al. / Chemical Geology 318–319 (2012) 88–112

landforms have been formed by hypogenic karstification, which is driven mainly by the strong geogenic carbon dioxide outflux. Spatial distribution of obruks implies that upwelling carbon dioxide flux towards the surface is routed not only by faults but also by the extent of Quaternary paleolake sediments that form a low permeability barrier on top of Neogene units. In Bayari et al. (2008), the two samples represented by empty circles contain detectable tritium (Fig. 6). The other samples (solid symbols) plot close to the “zero age” line (Z) in graphs I, II and III. The authors suggest that the age of the oldest groundwater is on the order of 40,000 years, but that seems unrealistic because this sample had a 14C activity of about 5 pmc and the high δ13C value suggests that the 14C content of this water has been further diluted from the crossing point O by dead carbon. Fig. 6 shows that some processes that add dead carbon enriched in 13C to the system have occurred (symbol ‘↑ ↑ ↓’). Most probably, the dead carbon is introduced by geogenic CO2. There seems to be a slight effect of carbon isotopic exchange with solid carbonate (because some points are above line Z in graph I). Based on the graphs the age of the oldest water could be estimated to be about 11,400 years, from an initial 14C content of ca. 20 pmc (vertical intersection with the zero-age line, Z, on graph II) (see and compare ages for solid circles in Fig. 18). 3.5. Isotopic exchange under open system conditions 3.5.1. Effect of isotopic exchange In the unsaturated zone isotopic exchange between gaseous CO2 and dissolved HCO3− will take place: 

−



−

C O2ðgÞ þ HCO3 ↔CO2ðgÞ þ HC O3

ð18Þ

where, “*” denotes carbon atoms initially in gaseous CO2. Compared with a closed system, process 18 will decrease δ13C and increase the 14C value of the DIC while [HCO3−] remains unchanged (symbol ‘– ↓ ↑’, see process 8 in Table 1). 3.5.2. Example Bredenkamp and Vogel (1970) studied the dolomitic aquifers in Transvaal, South Africa. The region is typically semi-arid and the water generally is in fissures in the solid dolomite, ranging from small cracks to substantial solution channels, sometimes measuring 10 m in diameter. The cavities are filled with chert debris or with a honeycomb residue, locally known as wad, which is formed during the process of karstification and represents the insoluble material of the dolomite. Samples from wells and springs were collected and analyzed for carbon isotopes and 3H. Bredenkamp and Vogel assume that the relatively large variations in the 14C content are due to differences in the unsaturated surface layers causing variations in the rate of infiltration and isotope exchange. Fig. 7 shows the data for samples collected from wells. The empty circles are waters containing tritium. While the bicarbonate concentration remains more or less constant, most of the waters have decreased δ 13C and increased 14C values relative to point O. The data plot approximately in the shaded areas on line X(a) in graphs I and II, and in region 1 in graph III. This distribution of data points is typical for isotopic exchange between groundwater and carbon sources which have low 13C and high 14C contents (e.g. soil CO2; symbol ‘– ↓ ↑’, process 8 in Table 1). 3.6. Isotopic exchange under closed system conditions 3.6.1. Effect of isotopic exchange Isotopic exchange between HCO3− and solid carbonate phase in the aquifer below the water table can change the isotopic composition of the DIC, e.g.: 

−

−



HC O3 þ CaCO3ðsÞ ↔HCO3 þ CaC O3ðsÞ

ð19Þ

97

where “*” denotes carbon atoms initially in the liquid phase. Because the liquid phase contains more 14C and the solid phase contains more 13C, process 19 will increase δ13C and decrease the 14C content of the DIC. For a closed system, isotopic exchange between HCO3− and solid carbonate, process 19, will cause a sample point (originally at point O) to move along line-X(b) (Fig. 1): in graph I, the points will move up; in graph II, the points will move down (symbol ‘– ↑ ↓’, see process 7 in Table 1). In Figs. 8, 15 and 17 some of the samples which had more or less constant bicarbonate concentrations, but show increases in δ13C and decreases in 14C, could have undergone process 19. Wigley et al. (1978) presented equations for calculating isotopic composition in systems undergoing geochemical reactions, and derived a relation for the isotopic compositions of DIC and a carbonate mineral as it approaches an infinite amount of isotope exchange. Plummer and Sprinkle (2001) applied the equations of Wigley et al. (1978) in correcting radiocarbon ages of DIC in the Floridan aquifer of Florida, USA. The Wigley et al. (1978) equations are incorporated in NETPATH (Plummer et al., 1994). Aeschbach-Hertig et al. (2002) showed that 14C content in a sandy aquifer, with 5–25% carbonate shell debris and calcareous-cemented sand, was affected by isotopic exchange leading to uncorrected ages that were too old, and, were difficult to correct, in part because of the low 14C content. Isotopic exchange may also happen in a system where calcite continuously dissolves and re-precipitates. Smith et al. (1975) showed that continuous dissolution and re-precipitation of calcitic minerals within the London Basin Aquifer are responsible for the enrichment of δ13C in the total dissolved inorganic carbon (TDIC) from −13 to −1‰ along a flow path of 8 km. Recently, Gonfiantini and Zuppi (2003) reviewed 14C–δ13C correlations in a number of aquifers and showed that isotope exchange caused by dissolution–precipitation process, if not accounted for, may lead to a significant overestimation of ages. As pointed out by Maloszewski and Zuber (1991), if the isotopic exchange partly takes place on the solid surface already modified by the exchange process, the δ 13C value of the DIC remains unchanged by that part of the process. As a consequence, the transport of 14C can additionally be delayed without significant changes in the δ 13C value and the 14C ages become additionally overestimated. 3.6.2. Example Fontes and Garnier (1979) investigated a carbonate aquifer composed of multi-layered karstified limestones and fractured sandstones of Paleozoic age. Most of the water samples were collected from two profiles (flowpaths). The measured δ 13C, 14C and HCO3− data for the samples are shown in Fig. 8. It can be seen from Fig. 8, that without significant change in bicarbonate concentration, 13C has relatively large variations ranging from −5.6 to − 16‰. In general an enrichment in 13C is accompanied by a decrease in 14C value (solid circles) and vice versa (empty circles). Isotopic exchange may be responsible for the changes in δ 13C and 14C (see Table 1, processes 7 and 8). The solid- and dashed-line arrows in Fig. 8 represent the effects of isotopic exchange with solid carbonate and soil CO2, respectively. All of the samples represented by empty circles contain 3H (7–140 TU). These waters could be mixtures of waters under closed system conditions with those under open system conditions. Such mixtures can be identified from the two samples which contain tritium but the 14C contents are lower than 20 pmc (see Section 4.2). Most of the samples represented by solid circles were collected from profile 1, and the samples collected from profile 2 are represented by empty circles. In profile 1, due to low permeability, groundwater circulates more slowly than that in profile 2. In profile 2 the groundwater circulation occurs through fractures and major channels of the karstic carbonates. This difference in flow rates explains the high degree of exchange with the carbonate matrix found for profile 1 (solid circles) and the high degree of exchange with soil CO2 found for profile 2 (empty circles, compare Fig. 7).

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0 (I) -2

Z

-4 Y -6

-8

-10 X -12 0

0.002

0.004

0.006

0.008

0.01

1/[HCO3-]

140

140 (III)

X

Y

(pmc)

60

80

14C

100

(pmc)

100

14C

120

80

(II)

X

120

60

Y

40

40

Z

Z 20

20

0

0 -12

-10

-8

-6

-4

-2

0

0

0.002

0.004

0.006

0.008

0.01

1/[HCO3-] Fig. 7. Typical graphs of isotopic exchange between groundwater and soil CO2, i.e., the reactions are open to soil CO2 (open system). Data are for samples collected from wells (Bredenkamp and Vogel, 1970). The empty circles are waters containing tritium.

3.7. Calcite precipitation under closed-system condition 3.7.1. Effect of calcite precipitation CaSO4 (gypsum, anhydrite) is typically more soluble than CaCO3 in water. On encountering CaSO4 containing sources in an aquifer calcite precipitation will occur (see reaction 23). Other processes such as dissolution of plagioclase feldspar and dissolution of other primary silicates will also cause calcite precipitation because they consume protons which increases the activity of the carbonate ion in solution, and also, in cases such as plagioclase feldspar dissolution, they release the common ion, Ca 2 + further causing calcite precipitation: þ

2þ

consumming H ; releasing Ca þ Ca CaAl2 Si2 O8ðsÞ þ 2H þ H2 O→

2þ

ð20Þ

þ Al2 Si2 O5 ðOHÞ4ðsÞ :

Calcite precipitation will cause a small decrease in δ 13C, 14C content and [HCO3−], because the heavier carbon isotopes are enriched in calcite cement due to isotope fractionation and because [HCO3−] is removed from solution in forming calcite cement. Therefore, the sample point will plot in regions 1, 2 and 3 in graphs I, II and III, respectively (assuming that the point is originally at point O).

3.7.2. Example Robertson (1992) used chemical and isotope data to evaluate the hydrological flow system of groundwater in a confined aquifer of the lower San Pedro basin in southeast Arizona, USA. Geochemical models developed from major-ion data show that reactions occurring along flow paths are weathering of plagioclase, potassium feldspar and hornblende, precipitation of calcite and formation of montmorillonite and goethite. Reactions were identified by mass-balance calculations using the program BALANCE (Parkhurst et al., 1982), which preceded NETPATH. The most significant observation is the similarity of the mass transfers of the primary aluminosilicates, although the chemistry and dissolved constituents differ significantly along flow paths. The major differences in water chemistry can be attributed to the dissolution of halite and gypsum. Chloride, sulfate and dissolved solids, for example, increase by factors of 5, 8 and 2, respectively, between the waters at the beginning of the confined aquifer system in the southern part of the basin, to that in the north. This similarity of the mass transfers of plagioclase, hornblende and potassium feldspar, and of the total mass transfer of the aluminosilicates, in view of the differing water compositions, indicates that a common geochemical process is occurring in the basin. The data from Robertson (1992) are shown in Fig. 9. All the above reactions (e.g. reaction (20)) will cause a decrease in [HCO3−], δ 13C and 14C in the water (symbol ‘↓ ↓ ↓’; see process 9 in Table 1). This

L-F. Han et al. / Chemical Geology 318–319 (2012) 88–112

99

0 (I) -5 Z -10 Y -15

-20

-25 X -30 0

0.005

0.01

0.015

0.02

-]

1/[HCO3 120

120 (III)

(II)

X 100

80

80

(pmc)

100

60

14C

14C

(pmc)

X

Y

60 Y 40

40

Z

20

Z

20

0

0 -30

-25

-20

-15

-10

-5

0

0

0.005

0.01

0.015

0.02

-]

1/[HCO3

Fig. 8. Most of the samples represented by solid circles were collected from profile 1. The samples represented by empty circles contain 3H and most of them were collected from profile 2. The full-line arrows indicate isotopic exchange with solid carbonate and the dashed-line arrows indicate isotopic exchange with soil CO2 respectively. Graphical presentation of data from Fontes and Garnier (1979).

decrease can also be seen from Fig. 9. That the sample points plot in regions 1, 2 and 3, in graphs I, II and III, respectively, indicate calcite precipitation. In Fig. 9 the symbols “+” and the oval area of graph III represent the corrected initial 14C and 13C contents (without calcite precipitation) calculated from the equation derived by Wigley et al. (1978) and used by Plummer et al. (1983). Because the soil carbonates here are not always dead carbon, in the calculations Robertson used the measured isotopic values of the soil carbonates, namely, δ 13C ~ −3.5‰ and 14C ~ 0−20 pmc. The oval areas in graphs I and II are projective images from graph III. These three areas show the isotopic and chemical data of the samples before calcite precipitation. The arrows show the chemical and isotopic evolution, with the short and long arrows representing the effects of calcite precipitation and 14C decay, respectively. 3.8. Weathering of silicates 3.8.1. Effect of weathering of silicates (process 10 in Table 1) While weathering of silicates may cause calcite precipitation (see Section 3.6.1), weathering of silicates may also cause an increase in [HCO3−]: þ

þ

Nax Kð1−xÞ AlSi3 O8ðsÞ þ 8H2 O þ CO2ðaqÞ →Nax þ Kð1−xÞ þ AlðOHÞ3 − þ 3H4 SiO4 þ HCO3

ð21Þ

with x=1 or 0. In this process CO2(aq) is transformed into HCO3− without reacting with carbonates (reaction 5). Consequently, the HCO3− produced by this process will have δ13C and 14C values close to that of the CO2(aq). 3.8.2. Example Stute and Deak (1989) studied water samples from various locations in the Great Hungarian Plain along two selected hydrogeological cross sections of about 100 km: an intermediate flow system through Quaternary (Q) layers and a regional system through underlying Pliocene (P2) layers with semi-permeable Pliocene sediments as confining beds in between. Stute and Deak found that non-radioactive reduction of the 14C/ 12 C isotope ratio plays an important role due to the addition of dead carbon species to groundwater along its subsurface pathway. This includes dead carbon from geogenic sources and oxidation of organic matter. Fig. 10 shows their carbon isotopic and HCO3− data. Based on the graphical method four groups of waters can be identified: (A) Samples in this group have not undergone any processes that could modify 14C content except for 14C decay (solid triangles; full-line arrows; symbol ‘– – ↓’). (B) Samples in this group contain additional carbon from oxidation of fossil organic matter that has δ 13C values close to the current soil CO2 (solid circles; symbol ‘↑ – ↓’) (also see Figs. 3 and 4).

100

L-F. Han et al. / Chemical Geology 318–319 (2012) 88–112

0 (I) B

Z

-5

Y -10

-15 X -20

-25 0

0.005

0.01

0.015

0.02

1/[HCO3-] 140

140

X

(III)

100

100

(pmc)

120

80

14C

14C

(pmc)

X 120

Y

60

(II)

80 60

Y

40

40 Z

Z 20

20

B

B

0

0 -25

-20

-15

-10

-5

0

0

0.005

0.01

0.015

0.02

-]

1/[HCO3

Fig. 9. The two samples represented by empty circles contain tritium of 8.4 and 25.6 TU, respectively. The symbol “+” represent corrected initial denote the chemical and isotopic evolution. Carbon isotopic and chemical data from Robertson (1992).

(C) Samples in this group contain additional carbon from oxidation of organic matter (empty circles; dashed-line arrows; symbol ‘↑ ↓ ↓’). As can be seen in Table 1, processes 3 (Cases 2 and 3) and 9 have similar effects (‘↑ ↓ ↓’). All these samples contain CO2(aq) indicating the lack of carbonates. However, the amount of CO2(aq) which is about 10% of the DIC, is unlikely to cause such a low δ 13C value. Therefore, the remarkably low δ 13C values of these samples could, at least partly, be caused by weathering of silicates (reaction 21). (D) Samples in this group contain additional carbon with δ 13C values close to zero (empty triangles; dot-dashed line arrows; symbol ‘↑ ↑ ↓’). These waters, except one, have undergone 14C decay. While group A samples were collected from the P2-flow system, group B and C samples were collected from the Q-flow system. Group D samples were collected from both systems. It can be seen therefore, while only the waters in the Q-flow system have been affected by oxidation of fossil organic matter, both systems have been affected by addition of dead carbon from geogenic sources. Most of these waters have undergone 14C decay. For waters in Group A, estimation of water age is relatively straightforward: taking initial 14C content close to the O point in graph II or III, and using

14

C and

13

C content. The arrows

Eq. (1) to calculate water age (the full-line arrows indicate the 14C decay). For waters in Groups B, C and D, the initial 14C contents are about 20 pmc, as indicated by the position of the arrow heads in graphs II and III. 3.9. Incongruent dissolution of carbonate minerals/dedolomitization 3.9.1. Effect of incongruent dissolution of carbonate minerals (process 11 in Table 1) Except for isotopic exchange, it is possible that the 14C isotope content be diluted without significant increase of dissolved carbon concentration, or, even with a decrease of [DIC]. One example is referred to as “incongruent dissolution of dolomite”, which can occur in limestone/dolostone aquifers (Wigley, 1975). Because dolomite (CaxMg(1 − x)CO3, with 0 b x b 1) dissolves more slowly than calcite, water in limestone/dolostone aquifers can attain saturation with calcite before dolomite, and as dolomite continues to dissolve, calcite precipitates due to the addition of Ca 2 +/Mg 2 + and CO32 − from dolomite:



2þ

Ca



2−

þCax Mgð1−xÞ CO3 ;−CaC O3

ðsÞ

2þ

2þ

2−

þ C O3 →Cax þ Mgð1−xÞ þ CO3 :

ð22Þ

L-F. Han et al. / Chemical Geology 318–319 (2012) 88–112

101

0 (I) -5

-10

-15

-20

-25

-30 0

0.002

0.004

0.006

0.008

0.01

-

1/[HCO3 ]

100

100

(III)

80

(pmc)

80

(pmc)

(II)

60

14C

14C

60

40

40

20

20

0

0 -30

-25

-20

-15

-10

-5

0

0

0.002

0.004

0.006

0.008

0.01

1/[HCO3-] Fig. 10. Four groups of waters can be recognized graphically (represented by solid triangles, empty triangles, solid circles and empty circles). Graphical interpretation of data from Stute and Deak (1989).

Due to precipitation of calcite some 14C will be removed from the solution. The removed carbon (C*) is then replaced by 14C-free carbon (C without asterisk) from dolomite. This combination of dissolution and precipitation will cause 14C-dilution with slightly increased DIC concentration. Incongruent dissolution (22) in a closed system will causes a sample point (assuming that the point is originally at point O) to move along line X(b) in graphs I and II toward higher δ 13C and lower 14C values. For a sample point at any position, the movement will be parallel to Line X toward higher δ 13C and lower 14C values. Some of the samples that could have undergone incongruent dissolution or isotopic exchange are plotted in Figs. 8, 15 and 17. Other incongruent reactions are driven by an irreversible reaction, like gypsum (CaSO4·2H2O) or anhydrite (CaSO4) dissolution that provides a common ion (Ca 2 + typically) and causes precipitation of calcite (through reactions R1 and R2): Ca

2þ





"

2− þCaSO4 ;−CaC O3 ðR1Þ

þ C O3 →

#þCaSO ;−CaCO ðR2Þ 4 3 2þ 2− Ca2þ þ SO2− →Ca þ SO4 4 2þ 2− dolomite dissolutionðR3Þ Ca2þ þ Mg þ CO Cax Mgð1−xÞ CO3 : x ð2−xÞ 3 ←

ð23Þ Due to reactions R1 and R2 some carbon will be removed. The difference between dedolomitization and calcite precipitation (processes 9 and 12 in Table 1) is that in the process of dedolomitization the

removed carbon is compensated through dolomite dissolution (R3) (i.e. calcite precipitation and dolomite dissolution driven by anhydrite dissolution) (Plummer, 1985; Plummer et al., 1990). The net effect is a decrease in [DIC] and 14C, accompanied by an increase in δ 13C. Wigley (1975) and Wigley et al. (1978, 1979) presented a quantitative discussion on the evolution of δ 13C (DIC) in groundwater with special emphasis on incongruent dissolution. 3.9.2. Example 1 Chen et al. (2003) conducted isotopic studies of four aquifers of the North China Plain. Several chemical processes in addition to carbonate mineral dissolution have been identified as being important in controlling the major ion chemistry in the samples. These processes include: ion exchange of Na + on the solids for Ca 2 + and Mg 2 + in solution; dissolution of carbonate minerals and gypsum in response to cation exchange; and mixing of saline water from the overlying aquifer through leakage. In the recharge area, the 3H contents are more than 6 TU. Fig. 11 shows the data from Chen et al. (2003). The samples denoted by empty triangles in the shaded area have distinctly lower [HCO3−] compared with other samples. The lower [HCO3−] could be caused by calcite precipitation, or due to lower soil CO2 content in the recharge area during the time of recharge, or, due to different evolution paths concerning open and closed system conditions (see

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1/[HCO3

Fig. 11. The empty circles are samples containing detectable tritium. The arrows and point A, A′ and A″ are discussed in Section 4.6. Graphs of data from Chen et al. (2003).

Section 4.6). These samples have undergone 14C decay. The samples denoted by solid circles in the shaded area are typical old waters (compare Fig. 14). The solid triangles represent waters in which HCO3− concentration is decreasing as 14C decreases. Because gypsum and dolomite are present, this represents calcite precipitation and dedolomitization (dolomite dissolution and calcite precipitation driven by gypsum dissolution). For a single process involving carbonate precipitation, with decreasing HCO3− concentration, both δ 13C and 14 C should decrease. But it can be seen from Fig. 11(I) that δ 13C has slightly increased with decreasing HCO3− concentration. This can be explained by the addition of 13C enriched HCO3− from dolomite dissolution (dedolomitization). Dolomite dissolution would cause an increase in δ13C and further decrease in 14C. The shifting of the solid circles toward slightly higher δ13C may be caused by isotopic exchange. In such a case combined use of geochemical modeling and the graphical method would provide a more detailed explanation. 3.9.3. Example 2 Plummer et al. (1990) reported results of a geochemical modeling study of the Madison Aquifer in parts of Montana, Wyoming, and South Dakota, a geochemically complex system undergoing dedolomitization, with varying amounts of cation exchange, oxidation of organic carbon, sulfate and iron reduction, and halite dissolution. It can be seen from Fig. 12 that most of the water compositions from the Madison aquifer are consistent with dedolomitizarion, showing an

increase in δ 13C with decrease in [DIC] (Fig. 12(I), points plotting in region 2). The increase in δ 13C results from dissolution of dolomites enriched in 13C (range of + 1 to + 5‰) and the decrease in [DIC] is due to calcite precipitation (dissolution of dolomites compensates the decrease in [DIC], because anhydrite is present, the net effect is a decrease in [DIC]). The enrichment in 13C from dolomite dissolution overwhelms the small effect to lower δ 13C values due to isotope fractionation in calcite precipitation. Three samples in the shaded area along line X(b) (Fig. 12(I)) were modeled by Plummer et al. (1990) as a combination of dedolomitization, cation exchange, and oxidation of organic carbon, which produces a close match in the measured and calculated δ 13C, although they might also be interpreted as affected by isotopic exchange. If dedolomitization were the only reaction and there were no decay of 14C, the sample points would shift in the directions indicated with the arrow in Fig. 12(III). Fig. 12(II) shows the effect of 14C decay on samples affected by (predominantly) dedolomitization, shifting from the shaded zone on line X(b) as calcite precipitates and 14 C decays. In Fig. 12(III), young samples plot in the shaded zone along the zero-age line (line Z), and those plotting below the Line Z have radiocarbon age. These observations are consistent with the geochemical modeling results reported by Plummer et al. (1990). The common characteristic of these two data sets (Figs. 11 and 12) is that most of the data plot on the right side of line X in graphs I and II, above line Y in graph I and below line Y in graph II. This is typical of samples affected, at least in part, by dedolomitization, but also is a

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1/[DIC] Fig. 12. Geochemical reaction models indicate that the dominant groundwater reaction is dedolomitization (arrows). The empty circles represent waters with detectable tritium. Data from Plummer et al. (1990).

property of samples from widely distributed sources. Indeed, unlike data discussed elsewhere in this paper, which were collected in relatively small areas, the samples of Chen et al. and Plummer et al. were collected from areas covering distances of several hundred km. However, while the samples represented by empty triangles in Fig. 10 can be interpreted as samples collected from an area with relatively low soil CO2 content, the spread of these data points over the [DIC] can also be interpreted as the effect of carbonate precipitation. 3.10. CO2 gas loss 3.10.1. Effect of CO2 gas loss (process 15 in Table 1) Loss of gaseous CO2 by degassing from water surfaces is the reverse process of introduction of carbon into water: þMeCO3ðsÞ

→ H CO → HCO− þ Hþ CO2ðgÞ þ H2 O ← 2 3← 3 R1

−MeCO3ðsÞ

→ Me2þ þ 2HCO− þ MeCO3ðsÞ ← 3 : R2

ð24Þ

Because the enrichment of heavier carbon isotopes in the water phase caused by reaction R1 overwhelms the depletion effect caused by reaction R2, process (24) will cause an increase in δ 13C and 14C values while a decrease in [HCO3−] (symbol ‘↓ ↑ ↑’).

3.10.2. Example Land and Huff (2010) conducted a multi-tracer investigation of groundwater residence time in a karstic aquifer in New Mexico, USA. The investigated system is a carbonate aquifer, and dissolved HCO3− in the waters could have undergone isotopic exchange with solid carbonate. Three samples have tritium content between 0.58 and 2.48 TU. All of the samples contain CFCs. As pointed out by the authors, these samples are mixtures of young and old waters. The low tritium and high CFCs indicate that the younger CFC apparent ages could be caused by “re-aeration” of the groundwater. Fig. 13 shows the data from Land and Huff (2010). The plot pattern is similar to that of Fontes and Garnier (see Fig. 8). The three samples with detectable tritium are represented by empty circles. There are no indications that additional carbon has been added to the system because no sample points plot in the regions left of line X in graphs I and II. Isotopic exchange between carbonate and water has caused an increase in δ 13C and a decrease in 14C value (Fig. 13, full-line arrows). On the other hand, “re-aeration” of the groundwater has caused a loss of CO2 from the water. Re-aeration causes an increase in 13C and 14C content while the HCO3− concentration decreases (Fig. 13, dashed-line arrows; symbol ‘↓ ↑ ↑’. Data plot right of line X in graphs I and II). Similar changes in δ 13C and HCO3− concentration were observed by Deines et al. (1974) for a spring sample that has lost CO2 to a cave atmosphere (see Fig. 19, below).

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1/[HCO3-] Fig. 13. The full arrows represent isotopic exchange between water and carbonate. The dashed arrows represent loss of CO2 from water. The symbols ‘+’ show the initial 14C and δ13C values modeled using NETPATH (Plummer et al., 1994) for selected sampling sites. The calculated 14C contents are corrected for predicted mass transfer in the absence of radioactive decay. Graphical interpretation of data from Land and Huff (2010).

The symbols ‘+’ in Fig. 13 show the initial 14C ( 14C0) and δ 13C values calculated with mass balance modeling (NETPATH) of carbon isotopic data between the starting point and downgradient sampling sites. It can be seen that the initial carbon isotopic data after correction plot close to the zero-age line. 3.11.

14

C decay

3.11.1. Effect of 14C decay (process 16 in Table 1) Unlike other processes, 14C decay will only affect 14C while the values of [DIC] and δ 13C are not affected (symbol ‘– – ↓’). 3.11.2. Example Evin and Vuillaume (1970) studied the confined groundwater of the Albian in the Paris Basin. The selected data from Evin and Vuillaume are shown in Fig. 14. From the three graphs in Fig. 14 it can be seen that 14C decay seems to be the dominant process that causes a decrease in 14C. This is because if there were dead carbon added to the system after the system became closed the 1/[HCO3−] value would decrease. And any isotopic exchange would change δ 13C. Therefore, the 14C0 can be assumed to be about 50 pmc and these waters seem to be very old.

It should be pointed out, however, that such a case may not be simple. If the groundwaters are old, the initial isotopic signature might be lost by water–rock reactions. And there might be redox reactions involving organic carbon leading to light 13C, and isotopic exchange with carbonate causing enriched values (both of the effects dilute 14C content, however, they have opposite effect on 13C). For example, the spread of δ 13C values (from −16 to −9) could be caused by isotopic exchange between water and carbonate (data above line Y in graph I), and precipitation of calcite (see Fig. 9). Therefore one would want to look at more chemistry (sulfate, iron, Ca, Na, etc.), and probably turn to NETPATH where there are both chemical and electron balance equations. Finally, as pointed out by Maloszewski and Zuber (1991), if there is an isotopic exchange process that partly takes place on the solid surface already modified by the exchange process, the δ13C value of the DIC may remain unchanged while the 14C content is diluted. Consequently, the 14C ages may be overestimated.

4. Other utilities of the graphs In addition to identifying processes that may be affecting 14C content of the water sample, the graphical method can provide a way of analyzing groundwater systems and comparing the data set.

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1/[HCO3 Fig. 14. Selected data from Evin and Vuillaume (1970) indicating old groundwater.

4.1. Identify different water systems Zuber et al. (2004) combined isotope, noble gas and chemical data to investigate age and flow patterns of groundwater in a Jurassic limestone aquifer and related Tertiary sands in the northwestern part of Cracow, Poland. Zuber et al. identified several processes, in addition to carbonate mineral dissolution and 14C decay, as being important in accounting for changes of water chemistry, δ 13C and 14C content in the samples. These processes include: isotope exchange between dissolved and solid carbonates for the confined part of the aquifer, admixture of older water ascending from underlying formations, and admixture of modern water as revealed by the presence of 3H. Fig. 15 shows the data from Zuber et al. (2004). As identified by Zuber et al. based on other evidence, the samples represented by an empty square and a triangle represent samples containing dominantly modern water. They were collected from the main recharge area. Because they are in closed systems (δ 13C ~ ½ δ 13C0), the empty triangle represents a water containing bomb- 14C (compare this sample with the empty circle, see below). The empty circle above line Y in graphs II and III is a sample collected from a spring. This sample contains some fractions that could have bomb- 14C. The two samples represented by ‘x’ were collected from production wells for mineral water. The lower 14C contents of these two waters can be attributed to 14C decay. The ages of these two waters could be estimated by using an initial 14C content of about 50 pmc.

The samples represented by solid circles have higher 13C contents than the other samples. Because all of these waters have similar [HCO3−] values, the higher 13C contents could be attributed to isotopic exchange as pointed out by Zuber et al. These waters should have had some radioactive decay of 14C, otherwise they would plot on line Z in graph III (see process 7 in Table 1). As identified by Zuber et al. based on other tracers, these waters have signatures of cooler climate. The sample represented by a solid square is identified by Zuber et al. to be a mixture of water from cooler period with modern water. Mixing with modern water has caused the sample point to move along line X in graphs I and II and along line Z in graph III toward the crossing point (see discussions in Section 4.2.). 4.2. Identify mixing of old water with young water Samples containing detectable 3H, below modern background and, plotting below line Y in graphs II and III can be identified as mixtures of old water with young water. One example is Fig. 4. In Fig. 4 the samples represented by solid squares plotting below line Y in graphs II and III, and they contain detectable 3H (0.28, 1.1 and 8 TU, respectively). These samples are most probably mixtures of old and young waters. A second example is Fig. 8 (see Section 3.6.2). A third example is Fig. 15. In Fig. 15 the sample represented by a solid square contains 3H below modern background (5.4 TU). This sample is identified by Zuber et al. to be a mixture of water from cooler period with modern water.

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1/[HCO3-] Fig. 15. Waters in different systems are represented with different symbols. The empty symbols represent samples with tritium contents between 5.4 and 26.4 TU. The solid triangle is apparently an outlier (see Fig. 12). Data from Zuber et al. (2004).

4.3. Comparison of different models in a graphical way In their investigation of a carbonate aquifer Fontes and Garnier (1979) compared initial 14C content of DIC calculated from different models and listed their results in a table. Fig. 16 shows the graphical representation of the data. It can be seen from Fig. 16 that for some of the data points below line Y the model of Mook can lead to negative ages (estimated initial 14C lower than the measured values). On the other hand, the model of Tamers can lead to older ages (estimated initial 14C too high). For these samples the other two models (Ingerson and Pearson, Fontes and Garnier) reconstructed the initial 14 C values close to the zero-age line. It can be seen that graphical presentations are superior to tables because they show the trend lines, data clusters, or curves that reveal a certain pattern of change easily. For example, it can be seen clearly that Tamers' model cannot describe the 14C isotopic dilution correctly. The reason why Tamers' model has failed in describing isotopic dilution in such cases is that in the Tamers' equation (Eq. (25)) only chemical reaction between soil CO2 and carbonate is considered:

A0 ¼

−

mCO2 þ 0:5mHCO3 Ag mCO2 þ mHCO− 3

ð25Þ

where m stands for molalities; A0 is the calculated initial 14C content and Ag the 14C content of the soil CO2 (generally taken as 100 pmc). Because this model only considers chemical reactions involving carbon and it does not consider other processes, this model would fail if the changes in δ 13C and 14C of DIC are attributed to other processes, e.g. isotopic exchange of water, either with soil CO2 or with carbonate. This is because these processes do not change chemical composition of the system, that is, the amount of the substances in Eq. (25) remains unchanged but the isotopic composition is changed. 4.4. Divide waters into different groups With the help of the graphical method the samples can be classified more easily based on the isotopic and dissolved inorganic carbon data. Jirakova et al. (2009) studied the paleorecharge conditions of the deep aquifers of the Northern Aquitaine region (France). Their carbon isotopic and HCO3− data are shown in Fig. 17. With the help of the graphical method, three groups of samples can be identified: (A) Samples represented by empty circles. For this group of samples one possible interpretation is that the dissolution of carbonates

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Fig. 16. The models are: Tamers (A); Mook (B); Ingerson and Pearson (C); Fontes and Garnier (D). Initial 14C values of total dissolved inorganic carbon from different models calculated by Fontes and Garnier (1979).

occurred predominantly under open system conditions (see Section 3.1.1). (B) Samples represented by triangles. These samples have been affected by exchange with carbonate and/or incongruent dissolution of carbonate. These processes will decrease 14C content and increase 13C content while [HCO3−] remains more or less constant (symbol ‘– ↑ ↓’). These waters are old because isotopic exchange without 14C decay would cause the sample points to be located in the shaded areas in graphs I and III. In the case of isotopic exchange, there would be a linear relationship between 14C content and δ 13C (see Appendix B, Eq. (B.4)). Compared with Group A samples, the bicarbonate concentrations of this group of waters are in a broader range. This indicates that some other processes that could change [HCO3−] have occurred. These processes also could affect the 14C content in the system to some extent. (C) Samples represented by solid circles. These samples have not been affected by any process that can modify initial 14C content in the water under closed system condition. This is because samples in this group have similar δ13C values. However, these samples have different 14C activities indicating that these waters have different ages due to 14C decay (symbol ‘– – ↓’). Estimation of water age of this group of waters is relatively straightforward: taking the initial 14C content of ca. 50 pmc (not necessarily

50 pmc, depends on the position of line Y), using Eq. (1) to calculate water age. As pointed out by Maloszewski and Zuber (1991), however, isotopic exchange may have taken place on the solid surface, but the δ13C value of the DIC remains unchanged. As a consequence, the initial 14C content may be lower than line Y in graphs II and III, and the 14C ages may be overestimated. Besides, similar to the samples in Group B, in this group of waters some processes that could change [HCO3−] could have occurred (compare Figs. 9 and 12). This would also lead to overestimation of the water ages. 4.5. Comparison of different data sets Similar graph patterns often indicate similar effects of processes on the initial 14C content (e.g. Figs. 2 and 4). Therefore, identifying similarity between graphs can help in data interpretation. Using 14C techniques, Muller and Mayo (1983) studied groundwater circulation in the Meade thrust allochthon, southeastern Idaho, U.S.A. Data from the analyses of major ions, 2H/ 18O, and 14C/ 13C, and from field geologic evidence have identified two distinct flow regimes within the Meade thrust allochthon. Shallow flow systems lie above the impermeable Phosphoria Formation, usually within a few hundred meters of the surface. Most of the spring waters from this system are recent and cool. In all cases, they have mean

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1/[HCO3-] Fig. 17. Three groups of waters can be recognized (Represented by empty triangles, solid circles and empty circles). Graphical interpretation of data from Jirakova et al. (2009).

subsurface residence times of less than a few hundred years (empty circles in Fig. 18). The deeper flow systems which lie below the Phosphoria formation are hydraulically isolated from the shallow system. The authors suggest mean groundwater residence times on the order of 15,000 years for the warm waters from these springs (solid circles in Fig. 18). The data set from Muller and Mayo (1983) have similar graphical patterns to that of Bayari et al. (2008) (Fig. 6) except for the ½ δ 13C0 values (in Fig. 6 the ½ δ 13C0 value is about − 6.5‰ for semi-arid land). Therefore, there is clear evidence that 14C content of the DIC in the warm waters has been diluted. Based on the vertical offset from the zero-age line on graph II the age of the water with the lowest 14C content (0.8 pmc) could be about 15,000 years if an initial 14C content of 5 pmc is used for calculation, though the other three samples are considerably younger than 15,000 years. 4.6. Identify evolution path regarding open and closed system conditions Different environmental conditions may affect the initial 14C content in different ways. Therefore, identifying the evolutionary path with respect to open and closed system conditions may help data interpretations. In Fig. 5 the position of point A can be estimated based on the soil PCO2. As can be seen from graphs I, II and III, the system is not a ‘pure’ closed system. Rather, for the samples represented by triangles, the

initial CO2(aq) had reacted with carbonate minerals at least to point A′ before the system became closed. On the other hand, the samples represented by empty circles were probably in systems open to soil CO2. In Fig. 11, the samples represented by empty triangles contain distinctly lower HCO3− than the other samples. This could be caused when the closed system conditions are predominant (long full-line arrows). In this case the crossing point would be at a location with δ 13C and 14C about ½ of the isotopic composition and, DIC about twice of the amount, of the initial CO2(aq). For these waters an initial 14 C value can be estimated to be about 50 pmc from the position of the crossing point represented by dashed lines. For the other samples (empty and solid circles, solid triangles), closed-system conditions were encountered probably after point A′ (dot-dashed and short full-line arrows). The initial 14C value of these samples can be estimated to be about 60 pmc. In the latter case the higher initial 14C value is due to the fact that more DIC has been exposed to the soil CO2 compared to the samples represented by empty triangles. Finally, as pointed out by Deines et al. (1974), the groundwater systems in Nittany Valley, Pennsylvania, are closed to the gas phase. Fig. 19 shows that the waters become isolated from soil CO2 with different initial [DIC]. This pattern is in agreement with the recharge that prevails in the study area. Much or most of the groundwater recharge which feeds conduit-type springs is surface runoff from noncarbonate ridges that enters the subsurface via sinkholes or fractured

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120 (III)

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1/[HCO3-] Fig. 18. Two distinct flow regimes can be identified graphically. Graphical representation of data from Muller and Mayo (1983). The graph pattern is similar to that of Bayari et al. (2008) (Fig. 6).

bed rock. As illustrated by Fig. 19, with increasing initial [DIC] the slope of the full-line arrows becomes smaller because of the smaller CO2(aq)/[HCO3−] ratio. In extreme cases where CO2(aq) ≪ [HCO3−], after the system becomes closed, the change in isotopic composition of the DIC will be very small because the contribution of Reaction (5) to the DIC is negligible, provided no other processes including 14C decay have taken place.

0 (I) -5

-10

5. Discussion

-15

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1/[HCO3

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Fig. 19. The dot-dashed- and full-line arrows indicate chemical and isotopic evolution under open and closed system conditions, respectively. The sample represented by an empty circle is influenced by CO2 exsolution and CaCO3 precipitation, as pointed out by Deines et al. (see Section 3.9). The samples represented by empty triangles are probably in open system environments. Data from Deines et al. (1974).

Radiocarbon dating of DIC in groundwater is the process of dating the 14C that is in the fraction of the dissolved inorganic carbon that was initially in isotopic exchange equilibrium with the modern CO2 reservoir, that is, the part of the DIC that was CO2 derived from plant root respiration in the recharge water. Because there can be many sources and sinks for 14C of DIC in the aquifer, it can be a formidable task to determine the amount of DIC and its 14C content that is of atmospheric origin from the recharge process. Dissolved inorganic carbon in groundwater systems can have many sources, including soil gas CO2, dissolved or particulate organic carbon, carbon from dissolution of carbonate minerals, or carbon of geogenic origin. The DIC in groundwater is further modified by chemical and isotopic exchange with solid and gaseous carbon phases during recharge, and within the aquifer by water–rock reactions, oxidation–reduction reactions, isotopic exchange, mixing with fluids from, for example, aquitards,

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isotopic fractionation in carbonate mineral precipitation, gas exchange, changes in temperature and pressure, mixing, and other physical processes. Radiocarbon dating of DIC in groundwater can be complicated by difficulties in recognizing which reaction processes are predominant, and often can be limited by incomplete and/or inadequate chemical and isotopic data. Ultimately, the process of radiocarbon dating in groundwater systems can depend on geochemical modeling approaches that calculate the amounts of the various fractions of the DIC and their isotopic composition. But before one turns to geochemical adjustment models (either the traditional adjustment models, e.g. Tamers, Ingerson and Pearson, Mook, Fontes and Garnier, etc., or more advanced geochemical modeling software, e.g. NETPATH, PHREEQC, etc.), there needs to be a conceptualization of the various reactions and/or processes that may be occurring in the system. This conceptualization usually is developed by examination of the chemical and isotopic data for the aqueous and solid phases in the system. The graphical approach presented here aids in this conceptualization process because (1) it provides a systematic way of examining the critical parameters, δ 13C, 14C content of DIC, and [DIC] or [HCO3−]; (2) it helps identify those samples that may be consistent with simple geochemical processes, and those samples that may be influenced by more complex geochemical and physical processes; (3) it permits comparison of data from different parts of the groundwater system, or from different aquifers; and (4) in many cases, the approach can lead to estimation of radiocarbon age. In most cases samples that plot close to the zero-age line (line Z in Fig. 1) do not contain detectable tritium. This means that the closed system development of the groundwater is a ‘slow’ process in the tritium age scale. Therefore, samples that plot close to line Z are only regarded as ‘young’ waters in the 14C age scale (up to a few thousands of years). In constructing the graphs, representative data are needed to define the critical points in the graphs, and depend on measurement or assumptions, such as δ13C of soil gas CO2, carbonates, and magmatic CO2 etc., that can have a certain range in δ13C and 14C values. Also, the water equilibrated with soil CO2 can have a wide range of DIC concentration and 14C value (for example, a silicate aquifer can have very low HCO3− concentration and low 14C content, although in general a low HCO3− concentration is associated with relatively high 14C). If no measured data are available, geochemical modeling techniques, such as those of PHREEQC, can be used to reconstruct hypothetical solutions according to general assumptions about the system under consideration. Many processes cause changes in δ 13C, 14C content and [DIC] differently. These processes are often readily identified. For example, as can be seen from Table 1, the process represented by different symbols ‘– ↓ ↑’, ‘↓ ↓ ↓’, ‘↑ ↑ ↑’, ‘↑ ↓ ↓’, ‘– ↑ ↑’, and ‘– – ↓’, can be identified with the help of the graphical method. However, there are processes which can cause similar changes in δ13C, 14C content and [DIC]. For example, dissolution of marine carbonates caused by addition of geogenic CO2, ion exchange on clay mineral and methanogenesis involving old organic matter will all cause an increase in [DIC] and δ 13C, and a decrease in 14C content (symbol ‘↑ ↑ ↓’). In such cases, chemical analyses such as analysis of CH4, Na+, Ca 2 +, etc., can help to resolve the process possibilities. In some cases mixing of waters of different age can be considered. For example, the sample points represented by empty circles in Figs. 4 and 8 could be caused by mixing processes. In general, if a sample point in the graph cannot be interpreted by simple processes, either multiple processes, or, mixing of waters should be considered. Other possible processes that can lead to changes in δ 13C and 14C content include leakage of saline waters; increase in hydrostatic pressure with depth; warming along a geothermal gradient, recrystallization of calcite, aragonite, or Mg-calcites (most of these processes would cause increase in [DIC] due to calcite dissolution in an closed system); leakage of aquitards or exchange with aquitards; diffusion; matrix diffusion; etc.

There are many more reaction possibilities that are not considered in the graphical method. Often these occur in multiple combinations and to varying extents. Even a seemingly simple process could be a complex one. Therefore, this method would only aid in the initial evaluation of the process of finding conceptually the geochemical (and other) processes that may be affecting the water sample. It should be used with the results obtained by the more advanced dating models, e.g. geochemical adjustment models, such as NETPATH, in order to provide a better understanding of the systems under investigation. 6. Conclusions The combined use of 14C content, δ 13C and concentration of dissolved inorganic carbon data can aid in the initial evaluation of geochemical data from groundwater systems by identifying processes that may be affecting the initial 14C content of the water sample. This evaluation leads to improved understanding of the overall water–rock reactions occurring in the system and ultimately, in conjunction with geochemical modeling efforts, improved estimation of radiocarbon age. The graphical method presented here can provide an overview of the entire data set, compare different data sets, help in identifying different water systems, and divide waters into different groups. The graphs provide an alternative method to compare different correction models for initial 14C content in a graphical way. Furthermore, the graphs can be used to provide information about hydrochemical complexity of a groundwater system, and provide a qualitative estimate of water age. Due to the complexity of geochemical and physical processes affecting the 14C content of DIC, this method cannot replace the more advanced approaches to radiocarbon dating in groundwater systems which are treated within the context of geochemical modeling; that is, the study of geochemical evolution of water–rock systems. However, these graphical techniques can help improve conceptualization of the predominant processes in groundwater systems that affect the radiocarbon content. Such an analysis aids in the construction of more advanced adjustment models that ultimately lead to more reliable 14 C ages (e.g. NETPATH Plummer et al., 1994). Acknowledgment The authors wish to acknowledge D. Parkhurst and two anonymous reviewers. The manuscript has been greatly improved because of their comments. Appendix A. Constructing the graphs Graph I is the relation between δ 13C and reciprocal of dissolved inorganic carbon concentration ([DIC]), graph II is the relation between 14C and 1/[DIC], and graph III is the relation between 14C and δ 13C. When constructing the graphs, follow the steps below: 1. Estimate initial 13C and 14C value of soil CO2 in the recharge area (δ 13C0 and 14C0). To a first approximation δ 13C0 can be estimated to be −30 to −20‰ (C3 type of land plants) or −20 to −15‰ (C4 type of land plants); and 14C0 to be 100 pmc. 2. Calculate ½ δ 13C0 and ½ 14C0. 3. Draw horizontal lines (Y) with the ordinate values of ½ δ 13C0 and ½ 14C0 in graphs I, II and III, respectively. 4. Draw vertical line (X) in graph III with the abscissa value of ½ δ 13C0. 5. Draw line Z connecting the crossing point (point O) and the origin in graph III. 6. Plot the measured data in graphs I, II and III.

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7. Observe the measured data in graph III. If there are data points plotted close to line X, then draw the vertical line (X) in graphs I and II, crossing the abscissa at the 1/[DIC] value of one of these samples. (It is assumed that in a closed system if δ13C is about the value of ½ δ13C0, no additional processes should have taken place except for reactions 4 and 5 involving soil CO2. However, this assumption is not necessarily true. Therefore, the graphs need to be refined.) 8. Draw line Z connecting point B and crossing point (O) in graphs I and II. 9. Refine the graphs based on other information or interpretation (soil CO2 partial pressure, pH value, temperature, oxidation of old organic matter, data cluster pattern, etc.).

where A and B are constants. If 14C decay is negligible during the change in 13δm (i.e. 14C decay rate bb 13δm change rate), a plot of measured 14C contents vs. 13C contents would be a straight line with a slope of A. (Fig. 1(III) line Z). In this case, Eq. (B.3) is identical with Pearson's model (Ingerson and Pearson, 1964), with 14δm equals corrected initial 14C content. In most cases, however, measurable 14 C decay occurs during the change in 13δm. In such cases, the data points will plot below line Z. Assuming that the measured 13C content ( 13δm) changes, and the change is proportional to time: (for example, thermal waters in a carbonate aquifer), 13

13

Appendix B. Using δ C as a qualitative indicator of groundwater age Fig. B.1 shows a rough exponential relationship between 14C and 13C contents. Such a rough exponential relationship has been observed in many studies, e.g. Deak (1978) (thermal hot spring); Deak et al. (1987) (deep Pliocene aquifers [2000 to 3000 m below land surface]); Dennis et al. (1997) (carbonate aquifer); Qin et al. (2005) (thermal hot water); Fontes and Garnier (1979) (carbonate aquifer); and Gonfiantini and Zuppi (2003). Fig. B.1. A curved 14C–δ 13C relationship (data from Dennis et al., 1997). Points: measured data; curve: exponential relationship simulated based on Eq. (B.6) with constant a = 4 × 10 − 4. If Eq. (3) can be used to describe a single binary mixing process of isotopes for both 13C and 14C, similarly, we obtain 13

14

 ½HCO− 13 13 13 3 0 δm − δx ¼  δ0 − δx Þ − HCO3 m

ðB:1Þ

 ½HCO− 14 14 14 3 0 δm − δx ¼  δ0 − δx Þ; − HCO3 m

ðB:2Þ

where subscripts 0, x, m denote initial carbon, added carbon (from sources other than soil CO2), and mixing product m (=measured value), respectively. Combining Eqs. (B.1) and (B.2), we obtain 14

13

δm ¼

δm −13 δx 14 14 14 δ0 − δx þ δx : 13 δ0 − δx

ðB:3Þ

13

namely, 14

13

δm ¼ A δm þ B

ðB:4Þ

100

111

13

δm ¼ atþ δ0 ;

ðB:5Þ

where t is the water age and a is a constant. At the time of t=0, 13δm = 13 δ0 (from Eq. (B.4) 14δm = 14δ0). As t increases, 13δ also increases due to isotopic exchange, incongruent dissolution of carbonate minerals, addition of CO2 of geogenic origin, etc. Combining Eq. (B.5) with the equation of radioactivity decay gives: −λ 13 δ −13 δ0 Þ C ¼ C0 e að m

ðB:6Þ

where C and C0 are the measured and initial 14C content, respectively, λ is the decay constant of 14C, 13δm and 13δ0 are the measured and initial 13 C content. According to Eq. (B.6), a plot of 14C content vs. 13δ would show an exponential relationship between the two carbon isotopes (Fig. B.1). Eq. (B.6) is similar to the equation derived by Gonfiantini and Zuppi (2003) which can also describe a curved 13δ vs. 14C relationship. Gonfiatini and Zuppi derived their equation by using a kinetic isotopic exchange model. In their model the isotopic exchange between water and carbonate is considered. In the derivation of Eq. (B.6), we use a binary mixing model, assuming that many other processes in addition to isotopic exchange between water and carbonate could also cause a curved 13δ vs. 14 C relationship (e.g. steady addition of geogenic CO2 to the system). Although using 13δ as a qualitative indicator of groundwater age is theoretically possible, age estimation of groundwater by using Eq. (B.6) could lead to a large bias. This is because the aquifer mineralogy and reactive surface area probably are not homogeneous enough to warrant a linear rate of reaction with time, and because some reactions may approach equilibrium or steady states along a flow path. It is often observed that along reaction fronts, some reactions take place faster than throughout the aquifer. Even though the curved δ 13C– 14C relationship cannot give quantitative age information of the groundwater system, the curved relationship may indicate decay of 14C (because the points do not plot on line Z, i.e. zero-age line). The occurrence of curved δ 13C– 14C relationships in carbonate, fractured, and/or geothermal aquifers, may support the hypothesis that the δ 13C value of the groundwater has more or less continually increased with time.

80

14

C (pmc)

References 60

40

20

0 -15

-10

-5

0

δ13C (per mil) Fig. B.1. A curved 14C–δ13C relationship (data from Dennis et al., 1997). Points measured data; curve: exponential relationship based on Eq. (B6) with constant a = 4 × 10− 4

Aeschbach-Hertig, W., Stute, M., Clark, J., Reuter, R., Schlosser, P., 2002. A paleotemperature record derived from noble gases in groundwater of the Aquia Aquifer (Maryland, USA). Geochimica et Cosmochimica Acta 66, 797–817. Appelo, C.A.J., 1994. Cation and proton exchange, pH variations, and carbonate reactions in a freshening aquifer. Water Resources Research 30 (10), 2793–2805. Aravena, R., Wassenaar, L.I., Plummer, L.N., 1995. Estimating 14C groundwater ages in a methanogenic aquifer. Water Resources Research 31, 2307–2317. Bayari, C., Serdar, Ozyurt, Nur, N., Kilani, Susan, 2008. Radiocarbon age distribution of groundwater in the Konya Closed Basin, central Anatolia, Turkey. Hydrogeology Journal 17, 347–365. Blaser, P.C., Coetsiers, M., Aeschbach-Hertig, W., Kipfer, R., Van Camp, M., Loosli, H.H., Walraevens, K., 2010. A new groundwater radiocarbon correction approach accounting for palaeoclimate conditions during recharge and hydrochemical evolution: the Ledo-Paniselian Aquifer, Belgium. Applied Geochemistry 25, 437–455. Bredenkamp, D.B., Vogel, J.C., 1970. Study of a dolomitic aquifer with carbon-14 and tritium. Isotope Hydrology 1970. IAEA, Vienna, pp. 349–372. Brinkmann, R., Münnich, K., Vogel, J.C., 1960. Anwendung der 14C-Methode auf Bodenbildung und Grundwasserkreislauf. Geologischen Rundschau 49, 244–253.

112

L-F. Han et al. / Chemical Geology 318–319 (2012) 88–112

Brook, G.A., Folkoff, M.E., Box, E.O., 1983. A world model of soil carbon dioxide. Earth Surface Processes 8, 79–88. Buckau, G., Artinger, R., Kim, J.I., Geyer, S., Fritz, P., Wolf, M., Frenzel, B., 2000a. Development of climatic and vegetation conditions and the geochemical and isotopic composition in the Franconian Albvorland aquifer system. Applied Geochemistry 15, 1191–1201. Buckau, G., Artinger, R., Geyer, S., Wolf, M., Fritz, P., Kim, J.I., 2000b. 14C dating of Gorleben groundwater. Applied Geochemistry 15, 583–597. Chen, Z.Y., Qi, J.X., Xu, J.M., Xu, J.M., Ye, H., Nan, Y.J., 2003. Paleoclimatic interpretation of the past 30 ka from isotopic studies of the deep confined aquifer of the North China plain. Applied Geochemistry 18, 997–1009. Coetsiers, M., Walraevens, K., 2009. A new correction model for 14C ages in aquifers with complex geochemistry — application to the Neogene Aquifer, Belgium. Applied Geochemistry 24, 768–776. Conrad, R., 2005. Quantification of methanogenic pathways using stable carbon isotopic signatures: a review and a proposal. Organic Geochemistry 36, 739–752. Crossey, L.J., Karlstrom, K.E., Springer, A., Newell, D., Hilton, D., Fischer, T., 2009. Degassing of mantle-derived CO2 and 3He from springs in the southern Colorado Plateau region — neotectonic connections and implications for groundwater systems. Geological Society of America Bulletin 121, 1034–1053. Deak, J., 1978. Environmental isotopes and water chemical studies for groundwater research in Hungary. Isotope Hydrology, 1978. IAEA, Vienna, pp. 221–249. Deak, J., Stute, M., Rudolph, J., Sonntag, C., 1987. Determination of the flow regime of Quaternary and Pliocene layers in the Great Hungarian Plain (Hungary) by D, 18 O, 14C and noble gas measurements. Isotope Techniques in Water Resources Development 1987. IAEA, Vienna, pp. 335–350. Deines, P., Langmuir, D., Harmon, R.S., 1974. Stable carbon isotope ratios and the existence of a gas phase in the evolution of carbonate ground waters. Geochimica et Cosmochimica Acta 38, 1147–1164. Dennis, F., Andrews, J.N., Parker, A., Poole, J., Wolf, M., 1997. Isotopic and noble gas study of chalk groundwater in the London Basin, England. Applied Geochemistry 12, 763–773. Eichinger, L., 1983. A contribution to the interpretation of 14C groundwater ages considering the example of a partially confined sandstone aquifer. Radiocarbon 25 (2), 347–356. El-Kadi, Aly I., Plummer, L. Niel, Aggarwal, Pradeep, 2010. NETPATH-WIN: an interactive user version of the mass-balance model, NETPATH (Article first published online: 6 DEC 2010) Ground Water, http://dx.doi.org/10.1111/j.1745-6584.2010.00779.x (http://onlinelibrary.wiley.com/doi/10.1111/j.1745-6584.2010.00779.x/pdf.). Evans, G.V., Otlet, R.L., Downing, A., Monkhouse, R.A., Rae, G., 1979. Some problems in the interpretation of isotope measurements in United Kingdom aquifers. Isotope Hydrology 1970. IAEA, Vienna, pp. 679–708. Evin, J., Vuillaume, Y., 1970. Etude par le radiocarbone de la nappe captive de l'albien du bassin de paris. Isotope Hydrology 1970. IAEA, Vienna, pp. 315–332. Fontes, J.-Ch., Garnier, J.M., 1979. Determination of the initial activity of the total dissolved carbon. A review of existing models and a new approach. Water Resources Research 12, 399–413. Fontes, J.-Ch., 1983. Dating of groundwater. Guidebook on Nuclear Techniques in Hydrology, 1983. IAEA, Vienna. Fontes, J.-C.h., 1992. Chemical and isotopic constraints on 14C dating of groundwater. In: Taylor, R.E., Long, A., Kra, R.S. (Eds.), Radiocarbon Dating after Four Decades: An Interdisciplinary Perspective. Springer-Verlag, New York, pp. 242–261. Geyh, M.A., 2000. An overview of 14C analysis in the study of groundwater. Radiocarbon 42 (1), 99–114. Geyh, M.A., 2005. Dating of old groundwater — history, potential, limits and future. In: Aggarwal, P.K., Gat, J.R., Froehlich, K.F.O. (Eds.), Isotopes in the Water Cycle, Past, Present and Future of a Developing Science. Springer, The Netherlands, pp. 221–241. Glynn, P.D., Plummer, L.N., 2005. Geochemistry and the understanding of groundwater systems. Hydrogeology Journal 13, 263–287. Gonfiantini, R., Zuppi, G.M., 2003. Carbon exchange rate of DIC in karst groundwater. Chemical Geology 197, 319–336. Hackley, K.C., Liu, C.L., Coleman, D.D., 1992. Dating of groundwater containing microbial CH4. Radiocarbon 34 (3), 686–695. Ingerson, E., Pearson Jr., F.J., 1964. Estimation of age and rate of motion of ground water by the 14C method. Recent Researches in the Fields of Hydrosphere, Atmosphere and Nuclear Chemistry, pp. 263–283. James, E.R., Manga, M., Rose, T.P., 1999. CO2 degassing in the Oregon Cascades. Geology 27 (9), 823–826. Jirakova, H., Huneau, F., Celle-Jeanton, H., Hrkal, Z., Coustumer, P., 2009. Palaeorecharge conditions of the deep aquifers of the Northern Aquitaine region (France). Journal of Hydrology 368, 1–16. Kalin, R.M., 1999. Radiocarbon dating of groundwater systems. In: Cook, P.G., Herczeg, A.L. (Eds.), Environmental Tracers in Subsurface Hydrology 1999. Kluwer Academic Publishers, New York, pp. 111–144. Land, L., Huff, G.F., 2010. Multi-tracer investigation of groundwater residence time in a karstic aquifer: Bitter Lakes National Wildlife Refuge, New Mexico, USA. Hydrogeology Journal 18, 455–472. Lesniak, P.M., 1985. Open CO2 – underground water system in the West Carpathians (South Poland) – chemical and isotopic evidence. Chemical Geology 49, 275–286. Maloszewski, P., Zuber, A., 1991. Influence of matrix diffusion and exchange reactions on radiocarbon ages in fissured carbonate rocks. Water Resources Research 27, 1937–1945. Mook, W.G., 1972. On the reconstruction of the initial 14C content of groundwater from the chemical and isotopic composition. Proceedings of Eighth International Conference on Radiocarbon Dating, vol. 1. Royal Society of New Zealand, Wellington, pp. 342–352.

Muller, A.B., Mayo, A.L., 1983. Ground-water circulation in the Meade thrust allochthon evaluated by radiocarbon techniques. Radiocarbon 25 (2), 357–372. Münnich, K.O., Vogel, J.C., 1962. Untersuchungen an pluvaialen Wässern der OstSahara. Geologischen Rundschau 52, 611–624. Oren, A., 2001. The biogenic basis for the decrease in metabolic diversity at increasing salt concentration: implications for the functioning of salt lake ecosystems. Hydrobiologia 466, 61–72. Parkhurst, D.L., 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 calculations. U.S. Geological Survey Water—Resources Investigations Report 99-4259 (312 pp. ( ftp://brrftp.cr.usgs.gov/pub/dlpark/geochem/unix/phreeqc/ manual.pdf)). Parkhurst, D.L., Charlton, S.R., 2008. NetpathXL — an Excel® interface to the program NETPATH. U.S. Geological Survey Techniques and Methods, 6-A26 (11 pp.). Parkhurst, D.L., Plummer, L.N., Thorstenson, D.C., 1982. BALANCE — a computer program for calculating mass transfer for geochemical reactions in ground water. U.S. Geological Survey Water—Resources Investigations, 82-14. 29 pp. Pearson, F.J., 1992. Effects of parameter uncertainty in modelling 14C in groundwater. In: Taylor, R.E., Long, A., Kra, R.S. (Eds.), Radiocarbon Dating After Four Decades: An Interdisciplinary Perspective. Springer-Verlag, New York, pp. 262–275. Pearson, F.J., Hanshaw Jr., B.B., 1970. Sources of dissolved carbonate species in groundwater and their effects on carbon-14 dating. Isotope Hydrology 1970. IAEA, Vienna, pp. 271–286. Pearson Jr., F.J., Swarzenki, W.W., 1974. 14C evidence for the origin of arid region groundwater, Northeastern Province, Kenya. Isotope Techniques in Groundwater Hydrology 1974. IAEA, Vienna, pp. 95–108. Plummer, L.N., 1977. Defining reactions and mass transfer in part of the Floridan Aquifer. Water Resources Research 13 (5), 801–812. Plummer, L.N., 1985. Geochemical modeling: a comparison of forward and inverse methods. Published by the In: Hitchon, B., Wallick, E.I. (Eds.), Proceedings First Canadian/American Conference on Hydrogeology, Banff, Alberta, June 1984. National Water Well Assoc, pp. 149–177. Plummer, L.N., Busby, J.F., Lee, R.W., Hanshaw, B.B., 1990. Geochemical modeling of the Madison aquifer in parts of Montana, Wyoming, and South Dakota. Water Resources Research 26 (9), 1981–2014. Plummer, L.N., Glynn, P.D., in press. Radiocarbon dating in groundwater systems. Dating Old Groundwater: A Guidebook. IAEA, Vienna. Plummer, L.N., Parkhurst, D.L., Thorstenson, D.C., 1983. Development of reaction models for groundwater systems. Geochimica et Cosmochimica Acta 47, 665–686. Plummer, L.N., Prestemon, E.C., Parkhurst, D.L., 1994. An interactive code (NETPATH) for modeling net geochemical reactions along a flow path, version 2.0. U.S. Geol. Surv., Water Resour. Invest. Rep, pp. 94–4169. Plummer, L.N., Sprinkle, C.L., 2001. Radiocarbon dating of dissolved inorganic carbon in groundwater from confined parts of the Upper Floridan aquifer, Florida, USA. Hydrogeology Journal 9, 127–150. Qin, D.J., Turner, J.V., Pang, Z.H., 2005. Hydrogeochemistry and groundwater circulation in the Xi'an geothermal field, China. Geothermics 34, 471–494. Robertson, F.N., 1992. Radiocarbon dating of groundwater in a confined aquifer in Southeast Arizona. Radiocarbon 34 (3), 664–676. Salem, O., Visser, J.H., Dray, M., Gonfiantini, R., 1980. Groundwater flow patterns in the western Libyan Arab Jamahiriaya. Arid-zone Hydrology: Investigations with Isotope Techniques 1980. IAEA, Vienna, pp. 165–179. Smith, D.B., Downing, R.A., Monkhouse, R.A., Pearson, F.J., 1975. Stable carbon and oxygen isotope ratios of groundwater from the Chalk and Lincolnshire Limestone. Nature 257, 783–784. Stute, M., Deak, J., 1989. Environmental isotope study (14C, 13C, 18O, D, noble gases) on deep groundwater circulation systems in Hungary with reference to paleoclimate. Radiocarbon 31 (3), 902–918. Tamers, M.A., 1967. Surface water infiltration and groundwater movement in arid zones of Venezuela. Isotopes in Hydrology 1967. IAEA, Vienna, pp. 339–351. Tamers, M.A., 1975. Validity of radiocarbon dates on groundwater. Geophysical Surveys 2, 217–239. Taylor, C.B., 1997. On the isotopic composition of dissolved inorganic carbon in rivers and shallow groundwater: a diagrammatic approach to process identification and a more realistic model of the open system. Radiocarbon 39 (3), 251–268. Van Leerdam, R.C., de Bok, F.A.M., Lomans, B.P., Stams, A.J.M., Lens, P.N.L., Janssen, A.J.H., 2006. Volatile organic sulfur compounds in anaerobic sludge and sediments: biodegradation and toxicity. Environmental Toxicology and Chemistry 25, 3101–3109. Vogel, J.C., Ehhalt, D., 1963. The use of carbon isotopes in groundwater studies. Radioisotopes in Hydrology 1963. IAEA, Vienna, pp. 383–395. Wigley, T.M.L., 1975. Carbon-14 dating of groundwater from closed and open systems. Water Resources Research 11, 324–328. Wigley, T.M.L., Plummer, L.N., Pearson Jr., F.J., 1978. Mass transfer and carbon isotope evolution in natural water systems. Geochimica et Cosmochimica Acta 42, 1117–1139. Wigley, T.M.L., Plummer, L.N., Pearson Jr., F.J., 1979. Errata: Geochimica et Cosmochimica Acta 43, 1393. Yamada, M., Ohsawa, S., Kazahaya, K., Yasuhara, M., Takahashi, H., Amita, K., Mawatari, H., Yoshikawa, S., 2010. Mixing of magmatic CO2 into volcano groundwater flow at Aso volcano assessed combining carbon and water stable isotopes. Journal of Geochemical Exploration 26, http://dx.doi.org/10.1016/j.gexplo.2010.10.007 (Key: citeulike: 8135422.). Zuber, A., Weise, S.M., Motyka, J., Osenbrück, K., Rozanski, K., 2004. Age and flow pattern of groundwater in a Jurassic limestone aquifer and related Tertiary sands derived from combined isotope, noble gas and chemical data. Journal of Hydrology 286, 87–112.