The mass balance of dissolved thallium in the oceans

Marine Chemistry 85 (2004) 125 – 139 www.elsevier.com/locate/marchem

The mass balance of dissolved thallium in the oceans Mark Rehka¨mper *, Sune G. Nielsen Institute of Isotope Geology and Mineral Resources, ETH Zu¨rich NO C61, CH-8092 Zu¨rich, Switzerland Accepted 17 September 2003

Abstract This study provides a comprehensive and quantitative discussion of the inventory and mass balance of dissolved thallium (Tl) in the oceans. The calculations are based on a critical evaluation of recently published analytical data. Thallium is estimated to have a mean seawater concentration of 65 pmol/kg and this implies a total oceanic inventory of about 8.76  104 Mmol Tl. The five primary sources of dissolved Tl in the oceans are rivers, hydrothermal fluids, eolian contributions from volcanic emanations and mineral aerosols, and benthic fluxes from continental margin sediments. Combined, these sources generate a Tl flux of between 1.4 and 12 Mmol/year, with a best estimate of 5.2 Mmol/year. Based on this, an oceanic residence time of 7 – 63 ky is calculated for Tl, with a best estimate of 17 ky. The two main oceanic sinks of Tl are scavenging from seawater by the authigenic phases of pelagic clays and uptake of Tl during low-temperature alteration of oceanic crust. These sinks remove about 1.6 – 7.7 Mmol/year of Tl, with a best estimate of 4.2 Mmol/year. This indicates an oceanic residence time of 11 – 56 ky, with a best estimate of 21 ky. The mass balance yields input and output fluxes which are identical, within the uncertainties, and this is in accord with a budget that is in steady state. D 2003 Elsevier B.V. All rights reserved. Keywords: Chemical oceanography; Fluxes; Residence time; Steady-state mass balance; Thallium; Trace elements

1. Introduction The element thallium (Tl) has two stable isotopes, Tl (29.5%) and 205Tl (70.5%), and an atomic mass of 204.38 amu. Mass-dependant variations in Tl isotope compositions of about 2x have been observed for various natural terrestrial materials, particularly seawater and sedimentary rocks (Rehka¨mper and Halliday, 1999; Rehka¨mper et al., 2002, 2003). In the silicate Earth, Tl is a trace element with a geochemical behavior that is akin to the alkalis K, Rb 203

* Corresponding author. Tel.: +41-1-632-7922; fax: +41-1-6321179. E-mail address: [email protected] (M. Rehka¨mper). 0304-4203/$ - see front matter D 2003 Elsevier B.V. All rights reserved. doi:10.1016/j.marchem.2003.09.006

and Cs (Shaw, 1952; Wedepohl, 1974; McGoldrick et al., 1979; Heinrichs et al., 1980). This resemblance is not surprising given that Tl and the alkali elements form univalent cations of very similar size. For 12fold coordination, Tl and K, Rb, Cs have ionic radii of 170 pm, and 164, 172, 188 pm, respectively (Shannon, 1976). Thallium is therefore generally classified as a highly incompatible lithophile trace element. In rocks, it is concentrated in plagioclase or, if present, in K+-minerals such as K-feldspars and biotite. In contrast to the alkali elements, Tl also displays chalcophile behavior. Thallium is observed to partition into sulfide melts and it is enriched in some sulfide minerals (Shaw, 1952; Wedepohl, 1974; McGoldrick et al., 1979; Heinrichs et al., 1980).

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In the marine environment the behavior of Tl is more enigmatic. This may reflect the redox chemistry of Tl, which can occur as either Tl+ or Tl3 + in surface environments. Some thermodynamic models predict that seawater Tl should be exclusively or predominately univalent (Vink, 1993; Kaplan and Mattigold, 1998; Lin and Nriagu, 1998), but others indicate that Tl exists primarily in the trivalent state (Batley and Florence, 1975; Turner et al., 1981). The dominant Tl(I) species in seawater are expected to be Tl+ and TlCl, whereas Tl(III) should occur either as sparingly soluble Tl(OH)3 or in the form of chloro or hydroxy complexes (Turner et al., 1981; Lin and Nriagu, 1998; Byrne, 2002). Thallium generally displays a conservative distribution in the oceans and under most circumstances it appears to be cycled through the marine system as an analogue to K (Bruland, 1983; Flegal and Patterson, 1985; Flegal et al., 1986; Whitfield and Turner, 1987; Donat and Bruland, 1995; Nozaki, 1997). These observations are in accord with the occurrence of Tl as a monovalent cation in seawater. The relatively high Tl concentrations of hydrogenous ferromanganese deposits and pelagic clays, however, indicate that Tl is much more reactive than the alkalis toward scavenging by oxyhydroxides (Matthews and Riley, 1970; McGoldrick et al., 1979; Heinrichs et al., 1980; Hein et al., 2000; Rehka¨mper et al., 2002, 2003). This has been interpreted by some as resulting from the presence of Tl3 + species in seawater (Batley and Florence, 1975; McGoldrick et al., 1979; Flegal et al., 1989). In addition, Flegal and Patterson (1985) and Flegal et al. (1989) suggested that some of the, albeit limited, perturbations of Tl concentrations and particulate fluxes in seawater may be due to dissolved Tl(III). Batley and Florence (1975) furthermore evaluated the adsorption of Tl from seawater onto ionexchange resins and concluded that up to 80% of seawater Tl may be trivalent. Based on the high reactivity of Tl toward scavenging by authigenic oxyhydroxides, it was postulated (Flegal and Patterson, 1985) that Tl should display a shorter oceanic residence time than K and Rb. These elements have residence times of 4.5 My and about 0.6– 0.9 My, respectively (Bruland, 1983). However, only two quantitative estimates for the residence time of Tl in the oceans have been published to date. A value of 10 ky was derived

from the Tl concentrations of seawater and pelagic sediment, but no details of the calculation were given (Flegal and Patterson, 1985). Flegal et al. (1989) computed a residence time of 30 ky, based on measurements of particulate Tl fluxes in the Northeast Pacific. A more detailed evaluation of the oceanic inventory and the major input and output fluxes of this element has not been undertaken. In part, this is probably due to the difficulty of obtaining precise Tl abundance data for geological samples. Such measurements are not straightforward because the thallium concentrations of silicate rocks are typically less than 1 Ag/g and most water samples have Tl contents of less than 100 pmol/kg. Recently, there has been a revived interest in the aqueous and marine geochemistry of Tl. This interest is based on (1) the discovery of significant variations in the stable isotope composition of Tl in seawater and ocean sediments (Rehka¨mper and Halliday, 1999; Rehka¨mper et al., 2002, 2003), and (2) the elevated Tl concentrations of many rivers, lakes and groundwaters. In many cases, anthropogenic inputs (e.g., from coal mines or coal combustion; Cheam et al., 2000) appear to be responsible for the high Tl contents, but others have been attributed to natural sources (e.g., Tl-rich sulfide mineralization; Xiao et ˚ strom, 2001). al., 2003) or a combination of factors (A The high toxicity of Tl warrants an investigation of such emissions, particularly because Tl can be accumulated by some agricultural plants and livestock (Sager, 1998; Anderson et al., 1999). The acute and chronic toxicity of Tl appears to be similar to that of Hg, Cd and Pb. For mammals, the 50% lethal dose (LD50) is between 5 and 70 mg/kg body weight (Zitko, 1975; Repetto et al., 1998; Sager, 1998). Intoxication is associated with disorders of the nerve and digestive system as well as the Na/K metabolism. Common symptoms include gastrointestinal alterations, polyneuropathy, loss of hair, and impairment of vision (Repetto et al., 1998). Both the environmental and the isotopic research would profit from an improved understanding of the budget of Tl in the marine environment. Due to advances in analytical techniques, precise and accurate concentration data have recently become available for seawater and all relevant oceanic sources and sinks of Tl. It is the intention of the present study to critically evaluate the published analytical data and to

M. Rehka¨mper, S.G. Nielsen / Marine Chemistry 85 (2004) 125–139

provide a comprehensive estimate of the oceanic inventory and mass balance of dissolved Tl.

2. Oceanic inventory The first reliable Tl concentration data for seawater were presented by Matthews and Riley (1969, 1970) and Batley and Florence (1975). These workers determined Tl contents of between 45 and 80 pmol/kg for various filtered (0.45 Am filters) seawater samples from the North Atlantic and Pacific. Later studies, which applied the precise technique of isotope dilution thermal ionization mass spectrometry (ID-TIMS), substantiated the earlier results, even though these analyses utilized unfiltered seawater samples. The agreement between data for filtered and unfiltered samples is expected, because measurements of particulate Tl fluxes indicate that solubilization of Tl from particles cannot generate significant (>0.1%) perturbations of dissolved Tl concentrations for offshore waters (Flegal et al., 1989). Flegal and Patterson (1985) reported Tl abundances of 60 –80 pmol/kg for seawater, and they noted that no significant differences existed between the Atlantic and the Pacific Oceans, between the Northern and Southern Hemispheres of the Pacific Ocean, and between surface and deep waters from both oceans. Murozima and Nakamura (1980) determined the Tl contents of seawater samples from four stations in the North Pacific and this included three depth profiles that extended below the mixed layer. The Tl concentrations were observed to vary between 55 and 100 pmol/kg, but most of the variability was localized at depths of less than 200 m. Below 200 m, two profiles showed nearly uniform Tl abundances of about 65 pmol/kg, whereas the Tl contents increased from 60 pmol/kg (at 200 m depth) to nearly 100 pmol/kg (at 6000 m) for the third. Schedlbauer and Heumann (1999, 2000) determined Tl in surface waters from the Atlantic Ocean. In addition, they analyzed one depth profile from the South Atlantic. The surface water data for 30 stations that span a N – S profile of the Atlantic Ocean (from 65jN to 70jS) appear to show that Tl abundances increase with increasing latitude. The highest Tl concentrations were obtained at >50jN (>90 pmol/kg) and at >60jS (>55 pmol/kg), whereas the abundances were lower close to the

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equator ( < 50 pmol/kg between 10jN and 50jS). The South Atlantic depth profile displayed only small changes in Tl concentration, with a maximum at 1 km depth (71 pmol/kg) and depletions both at the surface (54 pmol/kg) and at 4 km depth (64 pmol/kg). The studies of Schedlbauer and Heumann (1999, 2000) and Murozima and Nakamura (1980) provide no explanation for the variability of Tl concentrations in surface water samples. It is possible that the variations are caused by natural or anthropogenic eolian inputs or, in some cases, they may reflect fluvial or benthic fluxes from continental margins. Desorption of Tl from suspended sediments or Tl from hydrothermal sources may be responsible for the unusually high Tl content ( f 100 pmol/kg) of the deep Pacific water sample analyzed by Murozima and Nakamura (1980). An average seawater concentration of 65 pmol/kg is calculated for Tl, based on the results for deepwater samples (Table 1). An uncertainty of only F 5 pmol/ kg appears reasonable, given that most of the variability in abundance appears to be restricted to the mixed layer. The present compilation is thus in accord with earlier work, which concluded that Tl displays (nearly) conservative behavior in the marine environment (Bruland, 1983; Whitfield and Turner, 1987; Donat and Bruland, 1995; Nozaki, 1997). With a mass of 1.348  1021 kg for the oceans, the global marine inventory of dissolved Tl is calculated as 8.76 ( F 0.67)  104 Mmol (Table 1).

3. Input fluxes The main oceanic sources of dissolved Tl are rivers, hydrothermal fluids, atmospheric contributions derived from subaerial volcanism and the deposition of mineral aerosols, and benthic fluxes from continental margin sediments. Table 1 summarizes the annual Tl fluxes and the uncertainties of the estimates. 3.1. Riverine flux The studies of Cheam et al. (2000) and Cheam (2001) show that Canadian rivers and lakes generally have elevated Tl concentrations downstream of coal mines and coal-fired power stations. Lin and Nriagu (1999b) inferred that the high Tl contents ( f 200

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Table 1 Fluxes of dissolved Tl into and out of the oceans and the oceanic residence times calculated from these values Mass (kg) 1.348  1021

Input fluxes

Mass flux, range

Rivers Hydrothermal fluids Subaerial volcanisma Mineral aerosolsb Benthic fluxesc Total Residence time

1.5 – 15  1013 kg/year S = 3.7 – 7.7  1011 mol/year 5.3 – 9.1  1011 kg/year Mn = 0.5 – 13  1010 mol/year

Output fluxes

Mass flux, range

Pelagic claysd Altered oceanic cruste Total Residence time

Tl concentration, range

Tl concentration, best estimate

10 – 100 pmol/kg

65 F 5 pmol/kg

Tl concentration, range

Tl concentration, best estimate

3

< 5 – 100 pmol/kg

35 F 15 pmol/kg

4

0.8 – 1.9

1.3

5

7 – 92 nmol/kg

40 F 20 nmol/kg

6

0.3 – 2.7

1.2

8

0.2 – 4.3

1.6

10

0.1 – 1.0

0.4

12

0.03 – 2.0

0.7

1

Mass flux, best estimate 3.78  1016 kg/year 3 ( F 1.5)  1013 kg/year S = 5.75  1011 mol/year 7.0  1011 kg/year Mn = 6.8  1010 mol/year

Reference

Reference 2

Reference

6

7

(Tl/S)at = 8  10

9

3.7 nmol/g (750 ng/g) (Tl/Mn)at = 1  10

11

(Tl/Mn)at = 0.5 – 1.5  10

5

5

Tl mass (Mmol) 8.76 ( F 0.67)  104

Tl flux, range (Mmol/year)

1.4 – 12 7 – 63 ky

3.8 – 5.7  1012 kg/year

Tl flux, range (Mmol/year)

Tl flux, best estimate (Mmol/year)

5.2 17 ky

Mass flux, best estimate

Reference

Tl uptake, range

Tl uptake, best estimate

Reference

0.22 g/cm2/ky

13

1.0 – 2.0

1.3

15

2.0 nmol/g (0.4 Ag/g) 0.60 nmol/g (120 ng/g)

14

4.8  1012 kg/year

1.5 – 3.0 nmol/g (0.3 – 0.6 Ag/g) 0.15 – 1.0 nmol/g (30 – 200 ng/g)

16

0.6 – 5.7

2.9

1.6 – 7.7 11 – 56 ky

Tl flux, best estimate (Mmol/year)

4.2 21 ky

M. Rehka¨mper, S.G. Nielsen / Marine Chemistry 85 (2004) 125–139

Global oceans

Reference

M. Rehka¨mper, S.G. Nielsen / Marine Chemistry 85 (2004) 125–139

pmol/kg total Tl, which includes both dissolved and >0.45 Am particulate Tl) of two small rivers in Michigan (USA) reflect inputs from urban runoff. It is therefore likely that anthropogenic sources are also responsible for the high Tl concentrations of the Rhine at the Dutch/German border (350 pmol/kg total Tl; Cleven and Fokkert, 1994), and various rivers in Sweden, China and Japan (250 –6500 pmol/kg total Tl; Miyazaki and Tao, 1991; Axner et al., 1993; Luo and Hou, 1994). The present investigation intends to estimate the natural dissolved riverine flux of Tl, and it therefore needs to avoid anthropogenic contributions and discriminate between dissolved and particulate Tl budgets. For the Michigan rivers, particulate Tl accounted for about 50 – 60% of the total Tl concentration (Lin and Nriagu, 1999b), but it is unlikely that this figure is of validity for other river systems. At present, only scant high-quality data are available, which can be used to estimate the global dissolved nonanthropogenic riverine Tl flux. The most useful results were presented by Lin and Nriagu (1999a) and Cheam (2001) who analyzed a number of samples from the Great Lakes. Similar results were obtained in both studies and differences between dissolved and total Tl concentrations are negligible, because only about 5% of the Tl is associated with particulates (Lin and Nriagu, 1999a). These data may be particularly useful, because the Great Lakes collect the runoff of a large drainage basin to provide an average Tl concentration. The Tl abundances were not

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observed to increase in the direction of drainage and increasing industrial intensity from Lake Superior to Lake Ontario, but some higher concentrations were thought to reflect local sources of contamination (Lin and Nriagu, 1999a). The Tl contents were reported to vary between 6 pmol/kg for Lake Superior to 70 pmol/kg for Lake Michigan, with an overall average of about 51 pmol/kg. Numerous other (unfiltered) river and lake water samples were analyzed by Cheam et al. (2000) and Cheam (2001). Most individual data are not tabulated in the publications, however, and the text provides only a short summary of the results. High Tl concentrations generally appear to be related to industrial contamination but some natural variability is also apparent. In most instances, relatively pristine rivers and lakes were reported to have total Tl abundances that range from < 5 to about 100 pmol/kg, with many values below 50 pmol/kg. Rivers are estimated to have a mean dissolved Tl content of 35 F 15 pmol/kg. This figure is slightly lower than the Great Lakes average (to account for anthropogenic inputs) and it encompasses the lower concentrations of many other rivers. From this, the global riverine Tl flux is calculated as 0.8 –1.9 Mmol/ year, with a best estimate of 1.3 Mmol/year, for a total runoff of 3.78  1016 kg (Table 1). This calculation assumes that Tl behaves conservatively in estuaries, but this must be verified by future studies. It is possible that the riverine Tl budget is modified in estuarine environments, either due to loss of Tl by

Notes to Table 1: References. 1: Baumgartner and Reichel (1975). 2: Murozima and Nakamura (1980); Flegal and Patterson (1985); Schedlbauer and Heumann (1999, 2000). 3: Baumgartner and Reichel (1975). 4: Lin and Nriagu (1999a); Cheam et al. (2000); Cheam (2001). 5: Elderfield and Schultz (1996). 6: Metz and Trefry (2000). 7: Berresheim and Jaeschke (1983); Lambert et al. (1988). 8: Patterson and Settle (1987); Lambert et al. (1988); Nriagu (1989); Hinkley et al. (1994); Gauthier and Le Cloarec (1998). 9: Prospero (1981); Duce et al. (1991); Rea et al. (1994). 10: Taylor and McLennan (1985). 11: Sawlan and Murray (1983); Heggie et al. (1987); Klinkhammer and Palmer (1991); Johnson et al. (1992); Reimers et al. (1992). 12: estimated from data for Baltic Fe – Mn deposits from Ingri and Ponte´r (1986a,b); Rehka¨mper et al. (2002). 13: Osmund (1981); Hay et al. (1988). 14: Matthews and Riley (1969); Heinrichs et al. (1980); Rehka¨mper et al. (2003). 15: Hart and Staudigel (1982); Smith et al. (1995); Staudigel et al. (1995); Rowley (2002). 16: Hart and Staudigel (1982); Staudigel et al. (1995); Jochum and Verma (1996); Teagle et al. (1996); Bach et al. (2003). a The calculations assume that 70% of the volcanic aerosols are deposited over the oceans and that between 10% and 100% of the Tl dissolves in seawater. b Between 5% and 30% of the Tl in mineral aerosols is assumed to dissolve in seawater. c Benthic fluxes are produced by C-rich shelf sediments, which cover 8 – 13% of the ocean floor. d The area of pelagic sedimentation is assumed to be 300  106 km2. e An oceanic plate production rate of 3.4 km2/year is used and the upper 0.4 – 0.6 km of oceanic crust are assumed to be affected by lowtemperature alteration. For simplicity, all calculations assume the following relationship for aqueous samples: 1 l = 1 dm3 = 1 kg. The uncertainties in the calculated residence times do not include the uncertainty in the Tl inventory of the oceans, which is assumed to be well defined.

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adsorption or Tl additions from benthic fluxes or the desorption from suspended sediment. 3.2. Hydrothermal flux At present, Tl concentration data for high-temperature endmember hydrothermal fluids are only available from a single source (Metz and Trefry, 2000). These results indicate that Tl is enriched in hydrothermal fluids by a factor of about 0.5 – 1  103, compared to seawater. Data are reported for endmember fluids from five different vents at the Southern Juan de Fuca Ridge (SJFR) and two vents of the TAG Hydrothermal Field on the Mid-Atlantic Ridge, and they yield an average Tl concentration of about 55 nmol/kg for hydrothermal fluids. This mean value, however, may be strongly biased by the SJFR fluids, which have higher Tl concentrations (40 – 92 nmol/kg) than the two TAG samples ( f 10 nmol/kg). An average Tl concentration for hydrothermal fluids can also be calculated from the correlation of Tl with the K and Rb abundances of endmember fluids. To this end, the Tl data of Metz and Trefry (2000) were combined with compatible K and Rb abundances from various sources (Butterfield and Massoth, 1994; Von Damm, 1995; Edmonds et al., 1996). Plots of hydrothermal Tl versus K and Rb display correlations with r2 values of about 0.85 – 0.90, as was first noted by Metz and Trefry (2000). A global mean Tl concentration for hydrothermal fluids was then obtained by assuming (1) hydrothermal fluxes of 6.9  105 and 9.5  102 Mmol/year for K and Rb, respectively (Elderfield and Schultz, 1996), and (2) a hydrothermal water flux of 3  1013 kg/year (Elderfield and Schultz, 1996). These calculations indicate average Tl abundances of about 20 nmol/kg (based on K) and 60 nmol/kg (based on Rb), for endmember hydrothermal fluids, which is in reasonable agreement with mean Tl content of the published dataset (Metz and Trefry, 2000). Based on these results, hydrothermal fluids are estimated to have a Tl concentration of 40 F 20 nmol/kg. Combined with a water flux of 3 F 1.5  1013 kg/year, this yields a hydrothermal Tl flux of 0.3 – 2.7 Mmol/ year (Table 1). A best estimate of 1.2 Mmol/year is based on a Tl abundance of 40 nmol/g and a water flux of 3  1013 kg/year. These calculations assume that Tl displays conservative behavior upon mixing of the hot

vent fluids with ambient seawater, in analogy to the alkali elements. It is also possible, however, that Tl is lost from the hydrothermal plume, and this may occur if Tl is incorporated into precipitating sulfide phases or scavenged by Fe- and Mn-oxyhydroxides. No data are presently available to evaluate whether and how such processes alter the Tl flux provided by hydrothermal fluids. Further investigations are therefore required, that study the behavior of Tl in buoyant and neutrally buoyant hydrothermal plumes. The estimated Tl flux from vent fluids is compared in the following with the total flux of Tl that can be provided by leaching of pristine oceanic crust. The calculations utilize an ocean crust production rate of 3.4 km2/year (Rowley, 2002), an average crustal thickness of 6.4 km (Staudigel et al., 1995), and a density of 2.8 g/cm3 (Smith et al., 1995). A maximum estimate of how much Tl is available for leaching can be obtained by assuming that the Tl abundance of the total ocean crust is identical to the concentration in normal mid-ocean ridge basalts (N-MORB). This yields a maximum estimate because the cumulate rocks of the lower oceanic crust typically have much lower contents of incompatible trace elements than MORB (Hart and Staudigel, 1982). Fresh N-MORB is estimated to have a Tl concentration of 3 ng/g (15 pmol/g), based on a Cs abundance of 14.1 ng/g (Hofmann, 1988) and a Tl/Cs weight ratio of 0.2 (Sun and McDonough, 1989; Jochum and Verma, 1996). It is furthermore assumed (1) that only the lowermost 5.6 km of oceanic crust are leached by hydrothermal fluids and (2) that Tl is extracted with an efficiency of 100%. This yields a maximum flux of 0.8 Mmol/year of Tl, which can be provided by leaching of oceanic crust. A more realistic estimate assumes, in analogy to the calculation for Cs (Hart and Staudigel, 1982), that the Tl concentration of the lower 5.6 km of oceanic crust is only about 50% of the N-MORB abundance, and that hydrothermal alteration only leaches 85% of the Tl inventory from the rocks. With these values, the pristine oceanic crust can provide a hydrothermal fluid Tl flux of only about 0.3 Mmol/year and this is significantly lower than the preferred output flux of 1.2 Mmol/year (Table 1). A similar discrepancy was previously noted for the alkali elements Rb and Cs. Edmond et al. (1979) suggested that this could be due to a non-steady state of the hydrothermal systems during sampling. Alter-

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natively, it is conceivable that analytical data are only available for fluids that are atypically enriched in Tl. The most likely explanation, which was originally proposed by Palmer and Edmond (1989) for Rb and Cs, is that the high Tl contents of hydrothermal fluids are due to leaching of ocean crust that was previously enriched in Tl by low-temperature alteration processes.

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Mmol/year (Table 1). It is likely that the Tl flux from volcanoes is not at the lower end of this estimate, because Tl appears to be adsorbed onto the S aerosols as a chloride (Patterson and Settle, 1987; Hinkley, 1991; Churakov et al., 2000) and such species should be readily mobilized in seawater. The best estimate of the input flux (1.6 Mmol/year, Table 1) therefore assumes that 50% of the Tl from volcanic aerosols dissolves in the oceans.

3.3. Flux from subaerial volcanism 3.4. Flux from mineral aerosols Patterson and Settle (1987) noted that volcanic gasses have high Tl concentrations, such that volcanic emissions should generate a sizeable atmospheric flux of Tl. The present estimate of the global atmospheric input follows the approach outlined by the former study. We produced a new compilation of trace element contents for volcanic emissions, based on a survey of the published literature. The compilation includes data for samples from White Island, New Zealand (Patterson and Settle, 1987), Kilauea (Hinkley et al., 1994), and Mt. Etna (Gauthier and Le Cloarec, 1998) which yield an average (Tl/Pb)wt weight ratio of f 0.65 for volcanic gasses. Various studies furthermore indicate that volcanic gasses have an average (Pb/S)wt of about 8  10 5 (Patterson and Settle, 1987; Lambert et al., 1988; Nriagu, 1989). Taken together, this suggests a mean (Tl/S)wt ratio of circa 5  10 5 for subaerial volcanic emanations, equivalent to an atomic (Tl/S)at of about 8  10 6 (Table 1). The global volcanic flux of S from both eruptive and fumarolic sources has been estimated at between 3.7 and 7.7  1011 mol/year (Berresheim and Jaeschke, 1983; Lambert et al., 1988). These values yield a total atmospheric flux of Tl from volcanoes of 3.0 –6.2 Mmol/year, with a best estimate of 4.6 Mmol/ year. The real uncertainty of this figure is even larger and difficult to estimate, primarily because the Tl/S ratio of volcanic gases is based on a limited dataset. Even larger uncertainties, however, are encountered in calculating an oceanic flux of dissolved Tl from this atmospheric input. The oceanic flux is estimated from the atmospheric input by assuming (1) that 70% of the volcanic aerosol is transported to the oceans, and (2) between 10% and 100% of the Tl is then released into the dissolved form. From this, it can be inferred that the global volcanic Tl flux to the oceans is 0.2 – 4.3

The mass flux of mineral dust delivered to the oceans has been estimated at 5.3– 9.1  1011 kg/year (Prospero, 1981; Duce et al., 1991; Rea et al., 1994). On average, this eolian flux can be assumed to have a composition akin to upper continental crust with 750 ng/g (3.7 nmol/g) Tl (Taylor and McLennan, 1985). The calculation of the oceanic flux of dissolved Tl from mineral aerosols also requires knowledge about the dissolution behavior of particulate Tl. Such data are presently not available for Tl, and only a few other elements have been investigated. Between 2% and 20% of the particulate Nd from mineral dust is thought to contribute to the dissolved oceanic budget (Greaves et al., 1994; Tachikawa et al., 1999). The elements Sr, Ba and Mn furthermore appear to dissolve preferentially relative to Nd (Arraes-Mescoff et al., 2001). In the present study, it is assumed that between 5% and 30% of the Tl in mineral aerosols dissolves in seawater, with a preferred value of 15%. These data yield an eolian Tl flux from mineral dust of 0.1 – 1.0 Mmol/year, with a best estimate of 0.4 Mmol/year (Table 1). The high affinity of Tl towards adsorption onto clay minerals (Matthews and Riley, 1969) suggests that it is unlikely that a larger proportion (>50%) of this particulate Tl will dissolve in seawater. 3.5. Diagenetic fluxes from continental margin sediments Thallium is highly enriched in ferromanganese nodules and crusts where it appears to be primarily associated with Mn-oxide phases (Koschinsky and Halbach, 1995; Koschinsky and Hein, 2003). It is also relatively abundant in pelagic sediments, due to adsorption onto Fe –Mn oxyhydroxides (Matthews

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and Riley, 1970; McGoldrick et al., 1979; Heinrichs et al., 1980; Rehka¨mper et al., 2003). In addition, Tl is redox-sensitive, and sparingly soluble Tl(OH)3 can form in oxidizing environments. These observations indicate that Tl should be mobilized in the pore fluids of marine sediments during recycling of Mn. A number of studies have shown that significant benthic fluxes of Mn are generated only by continental margin sediments rich in organic matter, but not by oxic pelagic clays (Elderfield, 1976; Callender and Bowser, 1980; Sawlan and Murray, 1983). Benthic fluxes from shelf sediments are therefore inferred to be an additional source of dissolved oceanic Tl. Support for this conclusion is provided by the relatively high Tl concentrations (75 – 90 pmol/kg) measured for the coastal waters of the Irish Sea (Matthews and Riley, 1969; Riley and Siddiqui, 1986). The pore water data of Sawlan and Murray (1983) indicate that sediments from the shelf of Baja California and Guatemala generate a benthic Mn flux of about 2.8  10 4 mol/cm2/ky. Johnson et al. (1992) measured an average benthic Mn flux of 1.8  10 4 mol/cm2/ky for a transect across the continental shelf and upper slope of the California margin. The benthic flux produced by the shelf sediments of the eastern Bering Sea is lower at about 1.8  10 5 mol/ cm2/ky (Heggie et al., 1987). These estimates are combined with the total area covered by organic rich shelf sediments, which is about 8 – 13% of the seafloor (2.9 – 4.7  1018 cm2) (Klinkhammer and Palmer, 1991; Reimers et al., 1992). Together, this yields a global benthic Mn flux from continental margin sediments of 0.5– 13  1010 mol/year (Table 1). Calculation of a global Tl flux then requires knowledge of the Tl/Mn ratio for pore fluids emanating from continental margin sediments, but no such data are available. A rough estimate of the Tl/ Mn ratio is therefore obtained from Fe –Mn deposits of the Baltic Sea (Rehka¨mper et al., 2002). The validity of this approach is indicated by the similarity of Ni/Mn and Cu/Mn ratios for pore fluids and shelf sediments enriched in recycled Mn (Sawlan and Murray, 1983). Ferromanganese micronodules from the Gulf of Bothnia, which precipitated from pore fluids that remobilized Mn and Tl from anoxic sediment layers, have Tl concentrations of about 17 Ag/g Tl, whereas a Fe –Mn encrustation from the

Baltic Sea has f 5.5 Ag/g Tl (Rehka¨mper et al., 2002). If these abundances are combined with a Mn content of 20 –40% (Ingri and Ponte´r, 1986a,b), this yields an estimated (Tl/Mn)wt weight ratio for pore fluids of about of 2 – 6  10 5, equivalent to a (Tl/ Mn)at of 0.5– 1.5  10 5. Combination of the benthic Mn flux with this range of Tl/Mn ratios indicates a global Tl flux from shelf sediments of 0.03– 2.0 Mmol/year, with a best estimate of 0.7 Mmol/ year (Table 1). 3.6. Other input fluxes Cosmic dust is an additional potential source of dissolved oceanic Tl. The mass flux of cosmic dust has been estimated at 3.5 – 7.2  107 kg/year (PeuckerEhrenbrink, 1996; Levasseur et al., 1999). The maximum Tl concentration of this material is probably 150 ng/g ( f 0.75 nmol/g), which is given by the Tl content of volatile-rich CI chondritic meteorites. These figures imply a global extraterrestrial Tl flux of less than 75 mol/year, which is insignificant compared to the fluxes provided by other sources (Table 1).

4. Output fluxes The two mechanisms, which appear to dominate the removal of dissolved Tl from the oceans, are uptake by the authigenic phases of pelagic clays and low-temperature alteration of oceanic crust. Estimates of these fluxes and the associated uncertainties are summarized in Table 1. 4.1. Authigenic flux into pelagic clays The seawater-derived Tl component of pelagic clays originates from adsorption either onto Fe– Mn oxyhydroxides or the surfaces of clay minerals (Matthews and Riley, 1969; McGoldrick et al., 1979; Rehka¨mper et al., 2003). The difference in the Tl concentration between pelagic clays and shelf sediments can be used to estimate the authigenic, seawater-derived fraction of Tl, because the latter sediments are unlikely to have incorporated a significant hydrogenetic component (Wedepohl, 1960; Krishnaswami, 1976; Thomson et al., 1984). It is emphasized that the calculations are only approximate, because the avail-

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able database does not permit a comparison that utilizes continental margin and pelagic sediments with similar mineralogy and grain size. Heinrichs et al. (1980) determined Tl contents of between 0.8 and 1.3 Ag/g (3.9 – 6.4 nmol/g) for four composite pelagic clay samples from the Atlantic and Pacific, and this yields an overall average of 1.0 Ag/g (4.9 nmol/g) Tl. Similar results were obtained by Rehka¨mper et al. (2003), who measured Tl concentrations of 0.8– 1.2 Ag/g (3.9 – 5.9 nmol/g) for three pelagic clays from the Pacific. A Tl abundance of 1.8 Ag/g (8.8 nmol/g) was furthermore determined by Matthews and Riley (1969) for a red clay from the North Pacific Ocean. In comparison, sediments from the continental shelf have significantly lower Tl contents. Rehka¨mper et al. (2003) determined 0.2 Ag/g (1.0 nmol/g) Tl for a clay from Yellow Sea and Matthews and Riley (1969) measured a Tl abundance of 0.1 Ag/g (0.5 nmol/g) for a glauconitic mud from the continental shelf. In addition, Heinrichs et al. (1980) obtained mean Tl contents of 0.2 –0.7 Ag/g (1.0 – 3.4 nmol/g) for composite samples of sandstones, siltstones, graywackes and shales with < 0.5% organic carbon. These data indicate that pelagic clays have an authigenic Tl concentration of about 0.3 – 0.6 Ag/g, equivalent to 1.5– 3.0 nmol/g (Table 1). The authigenic Tl concentration is combined with a global average accumulation rate for pelagic clays to obtain the output flux of Tl. Hay et al. (1988) estimated the mass of pelagic clays that were deposited during the last 5 My. This value and the area of pelagic sedimentation, which is about 300  106 km2 (Hay et al., 1988; Milliman, 1993), yields an accumulation rate of 0.23 g/cm2/ky. A similar result follows from the global average sediment accumulation rate of 2 g/cm2/ky (Osmund, 1981). Pelagic clays constitute about 10% of the total oceanic sediment mass (Hay et al., 1988), which indicates that they accumulate at a rate of about 0.2 g/cm2/ky. A Tl output flux of about 1.0– 2.0 Mmol/year is obtained by combining the mean accumulation rate of 0.22 g/ cm2/ky with an authigenic Tl concentration of 1.5– 3.0 nmol/g. The best estimate for this flux is 1.3 Mmol/year, and this assumes an authigenic Tl content for pelagic clays of 2.0 nmol/g (Table 1). This estimate of the Tl output flux was confirmed by an independent calculation. Previous studies de-

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termined authigenic accumulation rates of between 6 and 23 Amol/cm2/ky (mean c 13 Ag/cm2/ky) for Mn in pelagic environments (Bender et al., 1970; Krishnaswami, 1976; Brewer et al., 1980; Martin and Knauer, 1980; Bacon and Rosholt, 1982; Thomson et al., 1984). Various analytical results indicate that the authigenic phases of pelagic clays have (Tl/ Mn)wt c 1.0– 2.5  10 4, which corresponds to (Tl/ Mn) at c 2.5 – 7  10 5 (Turekian and Wedepohl, 1961; Matthews and Riley, 1969; Krishnaswami, 1976; Heinrichs et al., 1980; Thomson et al., 1984; Taylor and McLennan, 1985; Rehka¨mper et al., 2003). Taken together, these data yield a depositional flux of about 1.0– 2.7 Mmol/year for Tl, and this agrees very well with the estimate obtained above. 4.2. Flux into altered oceanic crust The flux of dissolved Tl that is lost from seawater by low-temperature alteration of oceanic crust is estimated by comparing the Tl concentrations of fresh and altered MORB. Fresh N-MORB have Tl abundances of about 3 ng/g (15 pmol/g; see Section 3.2). Jochum and Verma (1996) analyzed altered MORB and obtained Tl concentrations of between 12 and 610 ng/g (0.06 – 3.0 nmol/g). The majority of the samples displayed Tl contents of about 100 –300 ng/g (0.5 – 1.5 nmol/g). McGoldrick et al. (1979) found that palagonites associated with MORB were characterized by Tl abundances of about 0.3 –3 Ag/g (1.5 –15 nmol/g). Teagle et al. (1996) analyzed 18 basalts from ODP Hole 896A, which covers the top 250 m of ocean crust and the data indicate an average Tl concentration of 110 ng/g (0.54 nmol/g). More representative results were obtained as follows. The data of Jochum and Verma (1996) and Teagle et al. (1996) indicate that the element flux into the ocean crust has Tl/Rb and Tl/Cs weight ratios of 0.07– 0.13 and 0.95– 0.99, respectively. These element ratios were then combined with the extensive Rb– Cs datasets available for the altered ocean crust of ODP Hole 418A (Hart and Staudigel, 1982) and Hole 504B (Bach et al., 2003). Hart and Staudigel (1982) infer that an average of about 8.3 Ag/g Rb and 0.17 Ag/g Cs were added to the upper oceanic crust of Hole 418A. The results of Bach et al. (2003) for Hole 504B suggest that Rb and Cs were enriched by 1.2 and 0.04 Ag/g, respectively. Combination of the Rb enrich-

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ments with the Tl/Rb ratios calculated above suggests that between 80 and 1100 ng/g (0.4 – 5.5 nmol/g) of Tl are added to the ocean crust during low-temperature alteration. The same calculations for Cs indicate addition of 30 – 170 ng/g (0.15 – 0.85 nmol/g) Tl. The latter result is considered to be more reliable because the element addition fluxes of Holes 418A and 504B have very similar Tl/Cs ratios, and this indicates that Tl and Cs display nearly the same behavior during low-temperature basalt alteration. Based on these data, the average uptake of dissolved Tl during alteration of oceanic crust is estimated to be 30– 200 ng/g (0.15 – 1.0 nmol/g), with a preferred value of 120 ng/g (0.60 nmol/g; Table 1). Studies of ODP drillcores indicate that low-temperature alteration of ocean crust extends down to maximum depths of approximately 400 – 600 m (Hart and Staudigel, 1982; Staudigel et al., 1995; Bach et al., 2003). If this range of depths is combined with an ocean crust production rate of 3.4 km2/year (Rowley, 2002), and a density of 2.8 g/cm3 (Smith et al., 1995), this yields a mass flux of 3.8 – 5.7  1012 kg/year for altered oceanic crust. Assuming a net addition of 0.15 – 1.0 nmol/g Tl, this implies an output flux for Tl of 0.6 –5.7 Mmol/year (Table 1). The best estimate of 2.9 Mmol/year assumes addition of 0.60 nmol/g Tl down to a depth of 500 m in the ocean crust. 4.3. Other fluxes Hein et al. (2003) recently noted that ferromanganese (Fe – Mn) crusts and nodules are a major sink for dissolved marine Te. Thallium is highly enriched in such Fe –Mn deposits (Hein et al., 2000), such that it is conceivable that they are also a major sink for Tl. The mass of Fe – Mn crusts and nodules that are currently deposited on the ocean floor has been estimated at 2  1014 kg (Hein et al., 2003) and 0.9– 15  1014 kg (Mero, 1965), respectively. Both types of deposits have Tl concentrations of about 100 Ag/g, equivalent to 500 nmol/g (Hein et al., 2000; Rehka¨mper et al., 2002). If a mean age of 2 My is assumed for the Fe – Mn crusts (which is probably on the low side), this translates into a depositional flux of 0.05 Mmol/year for Tl, which is insignificant compared to other output fluxes. Hydrogenetic Fe – Mn crusts are, therefore, recorders of the Tl isotope

composition of seawater (Rehka¨mper et al., 2002, 2003) that have no significant impact on the oceanic budget of this element. Assuming that 7.5  1014 kg of Fe – Mn nodules have accumulated on the seafloor with a mean age of 2 My, this implies a Tl output flux of only 0.2 Mmol/ year. A number of studies have furthermore shown that a significant fraction of the elemental budget of many Fe – Mn nodules is acquired during diagenesis from the pore fluids of substrate sediments (Price and Calvert, 1970; Dymond et al., 1984; Piper, 1988; Rehka¨mper et al., 2002). If sediments are indeed an important source of Tl for Fe – Mn nodules, this implies that the hydrogenetic Tl fluxes into Fe– Mn nodules are unlikely to have a significant impact on the dissolved seawater budget of this element. It is furthermore conceivable that biogenic carbonates or silica could represent an important sink for oceanic Tl. No Tl data are presently available for pure marine opal or carbonate to evaluate this possibility in detail. Several considerations, however, indicate that such fluxes will not be of primary importance. First, previously published analytical data suggest that Tl is not primarily associated with the biogenic carbonate or opal of marine sediments (Matthews and Riley, 1969; Heinrichs et al., 1980; Rehka¨mper et al., 2003). Second, the distribution coefficient of Tl for partitioning between the tests of foraminifera and seawater can be approximated using the data for the alkali elements Na and Li. Weight distribution coefficients of 0.1 and 8 have been estimated for these two elements, respectively (Lea, 1999). If these values are combined with a seawater Tl concentration of 65 pmol/kg (Table 1) and a calcite accumulation rate of 3  1012 kg/year (Milliman, 1993), this yields a Tl output flux of < 2000 mol/ year into carbonates, which is insignificant. Co-precipitation of Tl with biogenic opal is also unlikely to be a major Tl sink, given that the global silica removal rate from the oceans (about 4  1011 kg/year; LedfordHoffmann et al., 1986) is about an order of magnitude lower than the rate of carbonate accumulation.

5. Discussion and conclusions Fig. 1 provides an overview of the proposed oceanic cycle of dissolved Tl. The estimates for the global input fluxes of Tl to the oceans imply a residence time

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Fig. 1. Oceanic cycle of dissolved thallium. The oceanic inventory of Tl and the best estimates for the input and output fluxes (Table 1) are shown. The Tl fluxes and the mass of Tl in the oceans are given in units of Mmol/year and Mmol, respectively.

of 7 –63 ky, with a best estimate of about 17 ky. Based on the output fluxes, a residence time of 11 – 56 ky is calculated, with a best estimate of 21 ky (Table 1). These results are in accord with a mass balance that is in steady state and they are in good agreement with the published residence time estimates of 10 –30 ky, which were based on significantly more limited datasets (Flegal and Patterson, 1985; Flegal et al., 1989). Many of the flux estimates have large uncertainties, however. It is therefore desirable that the present budget is further evaluated in the future, based on a mass balance of Tl stable isotope compositions for the primary oceanic sources and sinks. Despite of the uncertainties, it is reasonably clear that the residence time of Tl is longer than the mixing time of the oceans (about 1 – 1.5 ky) but shorter than the oceanic residence time of Rb, which is about 600– 900 ky (Bruland, 1983). The present-day marine inventory of Tl is well-defined. A Tl residence time that is similar to the mixing time of the oceans would therefore require a total input flux that is at least a factor of three larger than the maximum estimate of 12 Mmol/year. It is unlikely that the present oceanic input fluxes are this large because (1) our maximum estimates for the riverine, hydrothermal, eolian, and benthic input fluxes are already quite optimistic, and (2) we are unaware of additional potential sources of oceanic Tl that could provide a flux of at least 25 Mmol Tl/year. With a marine residence time of about 10– 60 ky and a nearly conservative distribution, the ocean basins should be relatively homogeneous with respect

to their Tl isotope compositions. This conclusion is in accord with published data. Rehka¨mper et al. (2002) demonstrated that the surface layers of ferromanganese crusts are characterized by a global signature of e205Tl= + 12.8 F 1.2 (1r), whereas three seawater measurements yielded e205Tl c 6 (e205Tl is the deviation of the 205Tl/203Tl ratio of a sample from the Tl isotope standard in parts per 104). The difference in e205Tl between Fe – Mn crusts and seawater is thought to result from the nearly constant isotope fractionation that accompanies the adsorption of Tl onto the surfaces of Fe– Mn particles (Rehka¨mper et al., 2002). This interpretation implies that the Tl isotope compositions of the present oceans should also be constant to within about F 1.5 e205Tl. It is possible, however, that the Tl input or output fluxes were considerably larger in the geological past, for example, due to increased rates of volcanic activity or enhanced low-temperature alteration of oceanic crust. If the oceanic residence time of Tl was similar to the mixing time of the oceans due to such processes, different ocean basins may have exhibited significant differences in Tl isotope composition. At f 20 ky, the present residence time of Tl is comparable to the 18 ky time span that has elapsed since the last glacial maximum (LGM). If the assumption of steady state is correct, this implies that the marine Tl budget was only marginally affected by the changes in environmental conditions that were associated with deglaciation. The uncertainties of the mass balance calculations are, however, too large to infer a steady state Tl mass balance with certainty. Our

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results are, therefore, also in accord with a non-steady state marine Tl budget, which is still in the process of readjusting to changes in fluxes that may have occurred following the LGM. The new results confirm that Tl is indeed significantly more reactive in the marine environment than Rb, which is its closest geochemical analogue in the silicate Earth. It is conceivable that the higher reactivity is a consequence of the presence of at least some trivalent Tl species in seawater, which are expected to be much more particle reactive than dissolved Tl(I) (Whitfield and Turner, 1987; Flegal et al., 1989). If the calculated oceanic residence time for Tl is assumed to be the product of a binary mixture of Tl+ and Tl3 +, the data can be used to estimate the fraction of Tl(III) present in seawater. The calculations assume that Tl+ has a residence time similar to Rb (s c 750 ky), whereas Al3 + (s c 150 years; Whitfield and Turner, 1987) and Nd3 + (s c 1 ky; Tachikawa et al., 1999) are used as analogues for Tl3 +. These values imply that only about 1 –5% of the total seawater Tl must be present as trivalent species to account for the higher particle reactivity of Tl compared Rb. There is, however, also an alternative explanation for the comparatively high reactivity of Tl. It has been inferred from leaching experiments that Tl is preferentially associated with the Mn-oxide phases of hydrogenetic Fe – Mn crusts, which indicates initial adsorption as Tl+ (Koschinsky and Halbach, 1995; Koschinsky and Hein, 2003). Following adsorption, Tl is inferred to be rapidly oxidized to Tl3 +, whereby it is precipitated as Tl2O3 (Bidoglio et al., 1993). It is possible that such a surface oxidation mechanism, rather than the presence of dissolved trivalent Tl in seawater, may be responsible for higher particle reactivity of Tl compared to the alkali elements (Li, 1981; Whitfield and Turner, 1987). The present data therefore do not provide any direct evidence for the occurrence of dissolved Tl3 + in the oceans.

Acknowledgements Preparation of this paper was aided by discussions with Martin Frank, Alex Halliday, Jim Hein, Ralf Schiebel and Tina van de Flierdt. Sylvain Levasseur and Ben Reynolds read various versions of the script and their inputs were particularly valuable. The paper

also profited from the constructive formal reviews of two anonymous referees and, in particular, Bernhard Peucker-Ehrenbrink. Financial support by the Schweizerische Nationalfond (SNF) and the ETH Forschungskommission is gratefully acknowledged.

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