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Chapter 14

Po-210 in the Environment: Biogeochemical Cycling and Bioavailability Guebuem Kim, Tae-Hoon Kim, and Thomas M. Church

Abstract As the heaviest element of Group 6A, 210Po has a unique biogeochemistry in the environment that challenges our understanding. This chapter provides an overview of the research on 210Po in the atmosphere as well as in marine and other aqueous environments. Excess atmospheric 210Po has been attributed to external sources, such as volcanic emissions, resuspension of soil humus, incursion of stratospheric air, sea spray from the oceanic surface micro-layer, plant exudates including evapotranspiration, anthropogenic emissions (e.g., emission from coal combustion), and biovolatilization through the formation of dimethyl polonide. Most of these sources have been qualitatively identified, yet they remain difficult to quantify. In the aqueous environment, 210Po is efficiently accumulated in plankton and bacteria and is biomagnified through the food webs, relative to its grandparent 210 Pb, causing 210Po to be largely deficient in the euphotic zone. Globally, 210Po deficiency increases as ocean productivity decreases in the upper 1,000 m through biological transfer to the upper trophic levels. Smaller 210Po deficiencies in the productive areas of the ocean appear to be related to relatively active bacterial remineralization. Unusually high activities of 210Po are often found in the suboxic and anoxic waters in association with S, Mn, and Fe redox cycles. As many details of these processes remain elusive and under debate, we propose additional studies that should be conducted. G. Kim (*) and T.-H. Kim School of Earth and Environmental Sciences, Seoul National University, Seoul, South Korea T.M. Church College of Earth, Ocean, and Environment, University of Delaware, Newark, DE, USA

14.1 Introduction Naturally occurring 210Po (t1/2 ¼ 138 days) is the decay product of 210Pb (t1/2 ¼ 22.3 years) via 210Bi (t1/2 ¼ 5 days) in the 238U decay series and widely distributed in the earth’s crust, rivers, oceans, and the atmosphere. 210Po is highly radioactive, with alpha particle energy of 5.30 MeV. Although alpha particles are not sufficiently energetic to pass through a person’s outer skin, they easily penetrate the unprotected lining and pass through living cells when released within the lungs. Thus, 210Po and other radon daughters in the inhaled air may cause lung cancer, especially when individuals are exposed to high activities of pointsources in regions such as confined spaces associated with radioactive materials. In addition, it could pose a risk to human health when it enters the body through food consumption. The ratios of 210Po/210Pb vary in the earth’s surface, aqueous environment, and atmosphere because of their different chemical reactivity, biological enrichments, and volatility. In the atmosphere, 210Po produced from 210Pb is minor (e.g., 210Po/210Pb < 0.1), as measured in precipitation or fallout particles. However, a large excess (not produced in the troposphere) of 210Po is introduced by a variety of sources, including volcanic emissions, resuspension of top soils, incursion of stratospheric air, sea spray from the oceanic surface micro-layer, bio-volatilization of 210Po from the ocean, and anthropogenic emissions (Vilenskii 1970; Poet et al. 1972; Turekian et al. 1974, 1977; Moore et al. 1976; Bacon and Elzerman 1980; Lambert et al. 1985; Baskaran 2011). In the aquatic environment, 210Po displays a strong biogeochemical cycling, relative to its parent 210Pb, and exhibits a large deficiency in the euphotic zone because of its preferential

M. Baskaran (ed.), Handbook of Environmental Isotope Geochemistry, Advances in Isotope Geochemistry, DOI 10.1007/978-3-642-10637-8_14, # Springer-Verlag Berlin Heidelberg 2011

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removal by biota and rapid regeneration in the subsurface layer because of preferential remineralization from sinking biogenic debris. Concentration factors for 210Po in the pelagic ecosystem increase through the food web from phytoplankton and finally to fish (15  104, shrimp), although they are only about 12–23  102 for 210Pb (e.g., Fisher et al. 1983; Stewart and Fisher 2003; Stewart et al. 2005; Waska et al. 2008). Uranium mining operations often result in elevated levels of 210Po in freshwater fish (average 208 mBq g1), especially in the viscera, while 210Pb concentrates in the bone (Carvalho 1988; Carvalho et al. 2007). Po is effectively taken up by bacteria and dispersed between the cell walls, cytoplasm, and high-molecular-weight proteins in a manner similar to sulfur (Fisher et al. 1983; Cherrier et al. 1995; LaRock et al. 1996; Momoshima et al. 2001; Kim et al. 2005a). Thus, the volatility of Po could be due to the formation of biovolatile species such as dimethyl polonide (Hussain et al. 1995). Given its high affinity for bacteria, it is not surprising that extremely high and unsupported 210Po activity was found in some sulfide-bearing shallow groundwater in central Florida (Harada et al. 1989). Thus, studying 210Po in the environment is important in many aspects: (1) it can serve as a tracer for sulfur group elements such as Se and Te in aqueous environments, (2) its bio-accumulation and transfer through terrestrial and marine food chains are of great health concerns for human beings, and (3) the occurrence of excess 210Po in the atmosphere is also important for tracing atmospheric emissions and in enclosed areas for human health. In this chapter, we review recent advances in our understanding of the fate and cycling of 210Po in the atmospheric and aqueous environments.

14.2 Analytical Methods The analytical methods for measuring 210Po and 210Pb in the same sample are relatively straightforward (e.g., Sarin et al. 1992; Kim et al. 1999; Masque et al. 2002; Stewart et al. 2007; IAEA 2009; Baskaran et al. 2009; GEOTRACES User Manual, Church and Baskaran, http://www.ldeo.columbia.edu/res/pi/geotraces/ documents/GEOTRACESIPYProtocols-Final.pdf). The source of 210Po is prepared by spontaneous

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self-deposition onto a silver disc and following simple pre-concentration procedures to separate 210Pb. Then, 210 Pb is measured, via ingrown 210Po, after storing the sample for more than 3 months. Some researchers determine 210Pb activities by beta counting of its daughter, 210Bi, following a specific separation and source preparation in order to obtain 210Pb data without delay (e.g., Nozaki and Tsunogai 1973). Also, some researchers conduct the separation of 210Po from 210 Pb via specialized resins before auto-plating in order to eliminate any interference in the alpha counting (e.g., IAEA 2009). However, in this chapter we describe the simplest and most common method for 210 Po and 210Pb analyses in water samples. Although we do not include the method for solid samples, the methods after sample dissolution and subsequent Fe(OH)3 precipitation are the same. In general, approximately 10–20 L of water is required for the analysis of dissolved or total 210Po and 210Pb in seawater, river water, or lake water. Because the activity of 210Pb in particulate matter is approximately one order of magnitude lower than that in solution, the volume of water required for the analysis of particulate 210Po and 210Pb is approximately 3–5 times higher than that required for the analyses of 210Po and 210Pb in the solution form. Although the uncertainties in measuring the sample volume can easily be reduced in the laboratory, special tools, such as an electronic balance, may be required to weigh the samples in order to reduce the uncertainties onboard ships. The separation between the dissolved and particulate phases can be performed using the membrane or cartridge filters with a pore size of 0.4–0.8 mm immediately after sample collection. From the GEOTRACE intercalibration results (GEOTRACES User Manual), the difference between the particulate 210Po and 210Pb concentrations using 0.4- or 0.8-mm filters was found to be insignificant, but the choice of the filter materials is important – Supor 0.4–0.8 mm filter cartridges (e.g. Acropak 500) have been found to be reliable. The filtered water samples are stored in acid-cleaned plastic containers with acidification to a pH lower than 2. Because the intercalibration showed higher than 20% uncertainties for 10–20 L water volumes for both nuclides, water volumes of at least 50 L are recommended for particulate 210Po and 210Pb measurements. The water samples are transferred into plastic buckets or cubitainers, and then spiked with NIST-traceable

14 Po-210 in the Environment: Biogeochemical Cycling and Bioavailability 209

Po, 25 mg of ancient Pb carrier (with a negligible Pb activity), and 70 mg of Fe3+ carrier for 10–20 L volume. Following stirring with rods or bubbling with N2 gas for 10–30 min for equilibration of the spikes and carriers, the samples are allowed to stand for an additional hours to ensure complete equilibration (Kim et al. 1999). The Po and Pb are co-precipitated with Fe(OH)3 at pH 7–8 using ammonia solution, with vigorous stirring during the first 10 min, and the precipitate is then allowed to settle for hours. The supernatant is siphoned off, and the precipitates are collected using a centrifuge or Whatman 54 quantitative-grade paper. The precipitates on the filter paper are dissolved by adding approximately 3 mL of 6 M HCl and are then rinsed thoroughly using about 40 mL of deionized water to bring the samples for plating to approximately 0.5 M HCl. About 200 mg of ascorbic acid is added to reach a colorless solution (Fe is fully reduced). The Po in the solution is spontaneously plated onto a silver disc (the reverse side of which is covered by a neutral cement or plastic film) by swirling the disc in water at temperature at approximately 90 C. After the self-plating of Po onto the disc, the solution is dried and then dissolved in 5 mL of 9 M HCl. The Pb is separated from Po using a preconditioned 9 M HCl anion-exchange column. The remaining Po is adsorbed onto the column, while Pb passes through the column. To the collected Pb solution, a known amount of 209Po spike is added, and the sample is stored in a clean plastic bottle for more than 3 months. The 210Pb activity is measured by determining the ingrown activity of 210Po. The chemical yield of Pb from all the procedures is determined by measuring the stable Pb recovery for an aliquot of the 210 Pb solution. The 210Po activity is determined using alpha spectroscopy. The time of sampling, anion column separation, and alpha counting should be recorded to conduct the appropriate corrections for the ingrowth and decay of 210Po and 210Pb from sampling to counting (IAEA 2009). The activity of 210Po is calculated by the 210Po/209Po counting ratios multiplied by the activity of the added 209Po spike. The counts of each Po peak should be carefully corrected for background because Po background builds up easily owing to its volatile nature. The blank activities should be further corrected for the 210Po and 210Pb activities. 210

14.3

210

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Po in the Atmosphere

In the atmosphere, 222Rn daughters, 210Pb, 210Bi, and 210Po, grow in during the residence of aerosols because they are adsorbed readily onto particles. Thus, their disequilibria, 210Pb/222Rn, 210Bi/210Pb, and 210 Po/210Pb, have been used to determine the residence times of aerosols (e.g., Turekian et al. 1977; Carvalho 1995a; Church and Sarin 2008). Amongst these pairs, the residence times (10–300 days) of aerosols based on 210Po/210Pb disequilibria are generally much longer than those based on the other pairs (2–20 days) and aerosol deposition models (e.g., Kim et al. 2005a). Thus, the various possible sources for the excess 210Po (the observed 210Po activity minus the 210 Po activity produced from 210Pb in the troposphere) have been reported in many studies. The suggested primary sources include the input of aged aerosols from the stratosphere (residence time of about 1 year) or resuspended dust from topsoil in which the 210 Po/210Pb activity ratios are close to 1. Poet et al. (1972) and Moore et al. (1976) suggested that 210Po mixed down from the stratosphere was 0.2–7% of that in the troposphere. At Lamto, on the west coast of the African continent, soil dusts from Sahara desert contributed to very high 210Po/210Pb ratios, ranging from 0.30 to 0.54, in the winter atmosphere (Nho et al. 1996). In this chapter, we introduce recent progress in finding the sources of preferential 210Po inputs into the atmosphere via geophysical or biogeochemical processes. First, high 210Po can be introduced into the atmosphere by volcanic activities because of its high volatility relative to 210Pb. The activity ratios of 210 Po/210Pb were found to be between 5 and 40 in volcanic plumes from Mt. Etna (Lambert et al. 1985) and up to ~600 at the Stromboli volcano (Gauthier et al. 2000). For this reason, the 210Po/210Pb ratios in freshly erupted pyroclastic or lava flows are close to zero (Bennett et al. 1982; Gill et al. 1985; Rubin et al. 1994). Although quite patchy in space and time, the input of 210Po from volcanic emissions can account for more than one-half of the global budget of 210Po in the atmosphere (Lambert et al. 1979, 1982). In northern Taiwan, Su and Huh (2002) observed higher 210 Po/210Pb, derived from pre-eruption gas emission from the Philippines’ Mayon volcano. Based on these observational results, they suggested that 210Po might

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serve as a precursor of volcano eruptions worldwide if a global network can be properly designed. Second, the source of excess 210Po activities in the atmosphere can be from sea spray. The sea-surface microlayer has high activity of 210Po relative to 210 Pb. The secretion from phytoplankton concentrated in the surface microlayer, where the 210Po/210Pb activity ratios are as much as 3.8 (Bacon and Elzerman 1980), could increase the 210Po/210Pb activity ratios in the atmosphere by sea-sprays. Heyraud and Cherry (1983) showed that the 210Po/210Pb ratios were in the ranges of 1–3 and 2–80 for the surface microlayer and neuston, respectively. Kim et al. (2005a) showed that the seawater fraction of 210Po in aerosol samples from Busan, a harbor city in Korea, explained the unusually high excess of 210Po activities. This suggests that a small fraction of the sea spray can result in a large excess of 210Po in the coastal atmosphere. Third, bio-volatile 210Po from the eutrophic coastal region could result in high excess 210Po in the atmosphere. Hussain et al. (1995) documented that biovolatile Po (i.e., dimethypolonide) can be formed in natural water together with other species, such as DMS, DMSe, and MMHg, through theoretical calculations and laboratory experiments. Similarly, Momoshima et al. (2001) reported that the formation and emission of volatile Po compounds occurred associated with the biological activity of microorganisms in culture medium as well as in natural seawater. Kim et al. (2000) observed that the increase in the excess 210 Po in the Chesapeake Bay air was dependent on the wind speed over a threshold of 3 m s1 (mean), similar to other gases (i.e., CO2, SF6, and DMS) (Fig. 14.1). They showed that the simultaneously measured activity ratios of 7Be/210Pb and 210Pb/222Rn argued against either higher-altitude air or continental soils as the source of this excess. Thus, Kim et al. (2000) suggested that the source of excess 210Po could be from surrounding coastal seawater through the gaseous air-sea exchange of volatile biogenic species (e.g., dimethyl polonide). They found that less than 10% of the total 210Po in the Chesapeake or Delaware Bay was required to be volatile to account for the excess 210Po in the observed air. Hussain et al. (1993) showed nearly threefold higher activity ratios of 210Po/210Pb in aerosols in and around the mid-Atlantic region in the summer than that in the fall and winter, even through the winter winds are stronger, as an indication

G. Kim et al.

Fig. 14.1 Plots of maximum wind speed vs. 210Po/210Pb activity ratio for aerosol samples collected in Chesapeake Bay, USA, during 10–24 August 1995 (adapted from Kim et al. 2000). Wind speeds were measured on board hourly

of an additional bio-volatile 210Po source in the marine environment during summer. Forth, plant exudates, which are submicron particles, can be important sources of excess 210Po in the atmosphere (Moore et al. 1976). Plant exudates are formed by the condensation of volatile organic materials released from the plant surfaces. Stress conditions such as leaf expansion during active growth or by normal weathering due to wind action and abrasion may lead to loss of these exudates (Moore et al. 1976). Moore et al. (1976) predicted that plant exudates contribute to almost 40% of the natural and anthropogenic 210 Po fluxes in the atmosphere over the continental USA, on the basis of a weak correlation between submicron aerosols and vegetation density, without actual 210Po data. Since this process seems to be quantitatively important, more direct measurements are necessary in the future. Fifth, anthropogenic 210Po input can result from its high volatility. Moore et al. (1976) found that up to 7% of the total 210Po flux to the atmosphere in Boulder, Colorado, USA could be from anthropogenic sources. These can be produced from phosphate fertilizer dispersion, a by-product of gypsum or lead refinement, cement and other metal production, and fossil-fuel burning. As such, Carvalho (1995b) reported increased 210 Po/210Pb ratios in precipitation from the industrial emission of 210Po in Lisbon, Portugal. Kim et al. (2005a) documented that excess 210Po in the

14 Po-210 in the Environment: Biogeochemical Cycling and Bioavailability

metropolitan areas of Seoul, Korea, originates mostly from anthropogenic sources. They observed a strong correlation between non-sea-salt (nss) SO42 and excess 210Po, although there was no correlation between nss-SO42 and its parent 210Pb, suggesting that both anthropogenic SO42 and excess 210Po are controlled mainly by the same factor (Fig. 14.2). Based on this correlation and the d34S values, they concluded that the major source for excess 210Po in Seoul precipi-

tation is anthropogenic, likely from the burning of fossil fuels such as coals and petroleum oils, biomass burning, and/or high-temperature incineration. Gaffney et al. (2004) showed that the contribution of soils or coal-fired power plants to the 210Po/210Pb activity ratio in aerosols could be successfully evaluated by measuring the activity ratios for both fine (1 mm) fractions. Nevertheless, 210Po/210Pb disequilibria have been successfully used for estimating the aerosol residence times in some isolated areas (i.e., Arctic haze). Baskaran and Shaw (2001) observed that the 210 Po/210Pb ratios in Arctic haze aerosols varied between 0 and 0.177; the residence times of the aerosols calculated from these ratios were between 0 and 39 days. In addition, the aerosol residence times calculated from the 210Po/210Pb ratios at Centerton, New Jersey, USA, were equal to or slightly shorter than those calculated using 210Bi/210Pb ratios for the same aerosol samples; this indicated the negligible input of excess 210Po in this area. The mean deposition velocity of the aerosols using 210 Po was found to be 2.2 cm s1; this velocity was higher than that reported for 210Pb at the same site (McNeary and Baskaran 2007). The aforementioned difference could be attributed to the difference in the scavenging behaviors of these two nuclides. Thus, to understand the scavenging mechanisms for 210Poand 210Pb-laden aerosols and to validate the residence time calculation methods using Rn daughters, more extensive measurements must be performed on all Rn daughter pairs under different environmental conditions and for different particle sizes.

14.4

Fig. 14.2 Plots of excess 210Po activities vs. (a) non-sea-salt SO42 and (b) d34S in precipitation from Seoul (SNU, Korea) and Busan (PKNU, Korea) stations (adapted from Kim et al. 2005a). A 95% confidence interval is shown for the regression (a). In general, the d34S value is approximately 21‰ for sea spray, 5‰ on average for volcanic emissions, and lower than 4‰ for most of the biogenic emissions (Pichlmayer et al. 1998; Rees 1970). Anthropogenic sulfates exhibit a wide range of values from 0 to 10‰

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Po in the Ocean

In the surface ocean, 210Po is generally deficient relative to its parent 210Pb due to preferential removal by biota, where it is in near equilibrium or excess below the surface mixed layer due to rapid regeneration from sinking organic matter (Shannon et al. 1970; Bacon et al. 1976; Thomson and Turekian 1976; Cochran et al. 1983; Chung and Finkel 1988; Bacon et al. 1988). Sarin et al. (1999) showed that the dissolved 210Po activities in the surface waters of the equatorial and South Atlantic Ocean are about a third of the equilibrium concentrations. Based on these

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disequilibria, they calculated that the residence times of dissolved 210Po in the surface water at the equatorial and southern sites were about 73 and 130 days, respectively, with longer residence times for the intermediate depths (100–500 m). The observed 210Po-210Pb disequilibria showed a significant positive correlation (r2 ¼ 0.61) with the POC concentrations, suggesting that the 210Po deficiency in surface water is associated mainly with biological removal from surface waters (Sarin et al. 1999). Therefore, a larger deficiency of 210 Po would be expected in the more productive areas of the ocean, where the population of sinking particles is larger. However, Kim (2001) showed that 210Po deficiencies decrease as ocean productivity increases for the globally available 210Po data in the upper 1,000 m (Fig. 14.3). The removal of 210Po through the 500 m depth was about an order of magnitude higher in the oligotrophic ocean than in the productive areas of the ocean. Nozaki et al. (1990), for the first time, pointed out an unusually large deficiency of Po in the oligotrophic Philippine Sea. They explained this phenomenon with a 2–3-fold larger focusing of atmospheric 210Pb, relative to the model-based estimate, and subsequent ventilation into the deeper ocean. However, Kim (2001) posited that episodic atmospheric focusing could not significantly affect on the large deficiency of 210Po in the upper oligotrophic Sargasso Sea or Philippine Sea due to the following reasons: (1) the effect of a sudden increase in the atmospheric input of 210Pb (with negligible 210Po) cannot be significant because the water

residence times in the upper 500 m are longer than 10 years in the major oceans, allowing almost 100% 210 Po ingrowth; (2) the directly measured annual atmospheric deposition of 210Pb during the study period was less than 10% of the deficiency of the 210 Po inventory in the 0–700 m layer in the SargassoSea; (3) the largest deficiency of 210Po was found in the subsurface layer (Fig. 14.3) rather than in the surface layer; and (4) the large deficiencies are perennial in the Sargasso Sea based on bimonthly measurements (Kim 2001). Kim (2001) showed that the horizontal transport of waters from the ocean margins also could not be the source of this large 210 Po deficiency water because 210Po deficiencies are generally smaller in coastal waters due to rapid remineralization. Alternatively, Kim (2001) suggested that the large deficiencies of 210Po are likely due to biological removal from the total (dissolved plus particulate) pool by cyanobacteria and subsequent transfer to nekton (via grazing), which is unavailable to normal oceanographic water sampling. He showed the following evidence supporting this hypothesis: (1) the sediment trap based fluxes of 210Po (Bacon et al. 1985) through 3,000 m in the Sargasso Sea were an order of magnitude lower than the calculated 210Po removal flux through 3,000 m; (2) the calculated 210Po/210Pb export flux ratio in the upper Sargasso Sea (0–500 m) is about 20, which is similar to that in zooplankton (about 20), but much higher than that in sediment trap samples at 500 m in the Sargasso Sea (about 2–3);

Fig. 14.3 Total (dissolved + particulate) deficiency of 210Po relative to 210Pb in the upper ocean and removal fluxes of 210Po at 500 m (adapted from Kim 2001). The oceanic locations plotted are divided into three categories depending on biological productivity, (a) oligotrophic-very low productive ocean,

(b) mesotrophic-intermediate productive ocean, and (c) eutrophic-relatively high productive ocean. The abbreviations Eq., SW, NW, SE represent equatorial, south western, north western, and south eastern, respectively

14 Po-210 in the Environment: Biogeochemical Cycling and Bioavailability

(3) in the oligotrophic oceans, such as the Sargasso Sea, bacteria constitute the dominant biomass of the microflora and are partially predated upon by protozoa and higher trophic levels; (4) the particulate fraction (15–75%) in the oligotrophic Sargasso Sea is much higher than that in the productive areas of the ocean (e.g. the Bay of Bengal, 95% are free-living bacteria) are about fourfold faster, and the amount of free-living bacteria is an order of magnitude larger, relative to the oligotrophic ocean. This suggests that the Po could reside in the non-sinking organic pool for a much longer time in the eutrophic ocean, potentially with other sulfur group elements. Recent studies have suggested that 210Po is a potentially good tracer for particulate organic carbon (POC) export (Shimmield et al. 1995; Friedrich and Rutgers van der Loeff 2002; Murray et al. 2005; Stewart et al. 2010), similar to 234Th. Friedrich and Rutgers van der Loeff (2002) reported that the observed fractionation of 234Th and 210Po on particles, dependent on particle

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Fig. 14.4 Enrichment of polonium into cyanobacteria (adapted from Kim 2001). (a) A plot of dissolved vs. particulate 210Po in the surface (0–100 m) Sargasso Sea between 1996 and 1997, showing higher enrichment of 210Po to particulate matter during summer (the numbered percentages represent the proportion of particulate 210Po), (b) a correlation between Trichodesmium N2 fixation rates in the surface ocean and relative activities of particulate to dissolved 210Po ratios in surface ocean (upper 20 m). Errors for Trichodesmium N2 fixation rates are from the standard deviation of the average N2 fixation in Puffs and Tufts for each month, and errors for 210Po are based on 1s counting error propagation. The dotted lines indicate 95% confidence intervals

composition (POC/biogenic silica ratio), is in accordance with the known preference of 210Po for cytoplasm in the Antarctic Circumpolar Current. On this basis, they suggested that the utilization of the two tracers together would enable a more detailed

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interpretation of POC fluxes than would be possible by using 234Th alone. Murray et al. (2005) found that the POC export fluxes calculated from 210Po were much more variable than those calculated from 234Th because of the more variable correction factors for advection in the central equatorial Pacific. Stewart et al. (2010) found a very good correlation (>80%) between the relative fraction of 210Po and POC in sizefractionated particles, with a better relationship during non-bloom conditions at the Bermuda Atlantic Timeseries Study (BATS) site. They suggested that 210Po traces POC export fluxes more accurately in lowexport seasons than during high-export seasons, such as during a spring bloom. In the Bellingshausen Sea in Antarctica, the export production measured using 210 Po was considerably lower than that measured using 234Th, suggesting that a better understanding of radionuclide uptake and recycling in conjunction with POC is necessary to trace export production (Shimmield et al. 1995). Careful consideration is thus needed with regard to the use of 210Po as a POC tracer because 210Po uptake rates appear to be largely dependent on ecosystem structures and microbial roles.

14.5

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Po in Suboxic and Anoxic Waters

The activity of 210Po in oxic waters of major oceans ranges from 1 to 5 mBq L1. Similarly, in the oligotrophic Crystal Lake in Wisconsin, USA, the mean annual total concentration of 210Po is 1.6  0.7 mBq L1. However, 210Po activities are higher, up to 17 mBq L1, in seasonally anoxic ponds such as Pond B in South Carolina, USA (Kim et al. 2005b), and Bickford Pond in Massachusetts, USA (Benoit and Hemond 1990), as well as in permanently anoxic seawater

such as Framvaren Fjord in Norway (Swarzenski et al. 1999) (Table 14.1). In the permanently anoxic Jellyfish Lake in Palau, the maximum activity of 210Po was 133 mBq L1, among the highest found in natural waters, except for some sulfide-bearing shallow groundwater (Harada et al. 1989). Kim et al. (2005b) showed that, in the seasonally oxic environment, Pond B, the activity of 210Po increases sharply from the surface to the bottom layer, with a maximum activity of 14 mBq L1, in the summer, while it is vertically uniform and low in winter (Fig. 14.5). In Pond B, the bottom layer becomes anoxic from May to October, and the concentrations of Fe and Mn increase sharply from the surface to the bottom layer during this period, similar to the 210Po pattern, due to the dissolution of metal oxides and oxyhydroxides of Fe and Mn in the reducing bottom sediments (Kim et al. 2005b). It has been reported that Po (IV) is insoluble and is reduced to Po (II) near the potential at which Mn (IV) is reduced to Mn (II) (Benoit and Hemond 1990; Balistrieri et al. 1995). Thus, the summer profile of 210Po resulted from its diffusion from bottom sediments under reducing conditions, together with Fe and Mn, in contrast to its much lower winter patterns as a consequence of efficient co-precipitation with Fe and Mn oxides. However, 210Pb did not show such seasonal variations and was vertically uniform in Pond B, similar to Cs, Al, Na, and Cu, those are not sensitive to redox conditions (Kim et al. 2005b). Based on the seasonal changes in the 210Po budget, Kim et al. (2005b) showed that the actual amount of labile 210Po inputs to the water column is a small fraction (