Ocean-Atmosphere Interactions in the Global Biogeochemical Sulfur Cycle*

Marine Chemistry, 30 (1990) 1-29 1 Elsevier Science Publishers B.V., Amsterdam

Ocean-atmosphere interactions in the global biogeochemical sulfur cycle*

Meinrat O. Andreae Biogeochemistry Department, Max Planck Institute for Chemistry, P.O. Box 3060, D-6500 Mainz (F.R.G.) (Received December 5, 1989; accepted December 15, 1989)

ABSTRACT

Andreae, M.O., 1990. Ocean-atmosphere interactions in the global biogeochemicai sulfur cycle. Mar. Chem., 30: 1-29.

Sulfate is taken up by algae and plants and then reduced and incorporated into organosulfur compounds. Marine algae produce dimethylsuifonium propionate (DMSP), which has an osmoregulating function but may also be enzymatically cleaved to yield the volatile dimethylsulfide (DMS). At- tempts to identify the variables which control the oceanic production of DMS have shown that there are no simple relationships with algal biomass or primary productivity, but suggest that the concentration of DMS in the ocean is regulated by a complicated interplay of algal speciation and trophic interactions. Part of the biogenically produced DMS diffuses into the atmosphere, where it is oxidized, mostly to aerosol sulfate. The ability of these aerosol particles to nucleate cloud droplets, and thereby influence the reflectivity and stability of clouds, forms the basis of a proposed geophysiologi- cal feedback loop involving phytoplankton, atmospheric sulfur, and climate. Carbonylsulfde (COS) is produced photochemically from dissolved organic matter in seawater. The mechanism of this reaction is still unknown. Diffusion of COS from the ocean to the atmosphere is a globally signifcant source of this gas, which participates in the stratospheric ozone cycle. Hydro- gen sulfide and carbon disulfide are produced in the surface ocean by still unidentified processes, which appear to be related to biogenic activity. For these gases, the oceans are a minor source to the troposphere.

SOURCES OF SULFUR TO THE ATMOSPHERE: AN OVERVIEW

Recent aircraft measurements of atmospheric sulfur species show that anthropogenic emissions are influencing the global atmospheric sulfur cycle even over remote ocean regions (e.g. Andreae et al., 1988 ). The human perturba- tion of the atmospheric sulfur cycle results largely from the emission of sulfur dioxide (SO2) from fossil fuel burning. A number of recent papers have reviewed these emissions and presented a detailed source allocation (e.g. Cullis

*Presented at the section on Atmospheric and Marine Chemistry of the 32nd IUPAC Congress in Stockholm, Sweden, August 2-7, 1989.

0304-4203/90/$03.50 © 1990 -- Elsevier Science Publishers B.V. 2 M.O. ANDREAE and Hirschler, 1980; M6ller, 1984). The estimates for man-made sulfur emissions fall into a relatively narrow range: about 2.5 ___0.3 Tmol yr- 1 (Tmol: 1 Teramol = 1012 mol = 32 × 1012 g). The characteristics of the natural biogeochemical sulfur cycle in the atmosphere-biosphere-ocean system are much less well known, but are currently receiving intense interest because of their potential involvement in the regulation of global climate (Charlson et al., 1987). A summary of natural sulfur emissions from all sources is given in Table 1. This table presents the best estimates of these fluxes based on current information; it must be emphasized that most of these estimates are rather uncertain. This applies especially to the emissions of particulate sulfur in the form of dust and seaspray and to the emissions from soil and plants on the continents. The main reasons for the uncertainty regarding continental emissions ofbiogenic sulfur compounds are ( 1 ) the difficulty of accurately determining the various biogenic sulfur species, particularly hydrogen sulfide (H2S), at the low levels found in unpol- luted environments, (2) the technical problems of measuring emission fluxes from forest and brush ecosystems, and (3) the inadequate geographical coverage of existing data. Recent measurements of biogenic sulfur fluxes from terrestrial ecosystems have shown much lower emission rates than had been assumed just a few years ago, leading to lower estimates of their contribution to the atmospheric sulfur cycle and consequently making the oceans and fossil fuel burning by

TABLE 1

Estimates of natural sulfur emissions (in Tmol S year-1 )

SO2 H2S COS DMS CS2 Sulfate Other Total

Seaspray 1.2-10 1.2-10 Dust 0.1-1 0.1-1 Total 1.3-11 1.3-11 particulates

Volcanoes 0.23-0.29 0.03 0.0003 - 0.0003 <0.1 ? 0.3-0.4 Soils and - 0.1-0.3 +0.02 0.006-0.12 0.02-0.025 - 0.03 0.15-0.4 plants Coastal - 0.03 0.004 0.02 0.002 - 0.004 0.06 wetlands Biomass 0.08 ? 0.003 - ? ? ? t> 0.08 burning Oceans - 0.05-0.2 0.011 0.6-1.6 0.01 - ? 1.1-1.8 (gases)

Totalgases 0.3-0.4 0.2-0.6 0.00-0.04 0.6-1.7 0.03-0.04 <0.1 0.03 1.2-2.8"

"Equivalent to 38-89 Tg S year-~. GLOBAL B1OGEOCHEMICAL SULFUR CYCLE 3 far the most important sources of atmospheric sulfur. Among continental sources of sulfur gases, emissions from plants are now recognized as being at least as important as soil emissions. The results from recent work on the biogenic sulfur cycle over the continents have been reviewed by Andreae (1990a); a more detailed discussion of sulfur fluxes over the tropical continents can be found in Andreae and Andreae (1988), Andreae et al. (1990), and Bingemer et al. (1990). On a global scale, biomass burning appears to be a minor source of atmospheric sulfur, with an annual sulfur release rate of ~ 0.08 Tmol year- 1. It is however, a regionally important source in the tropics, where other sulfur emissions are sparse (Andreae, 1990b). In the following sections, I will discuss the principles of biogenic sulfate reduction and synthesis of volatile species, the oceanic emission of dimethylsulfide (DMS), carbonyl sulfide (COS) and other volatile sulfur species, and the fate of these compounds in the atmosphere. Additional information on other aspects of the sulfur cycle can be found in recent reviews (Andreae, 1985a; Andreae, 1986, and references therein) and in the proceedings volume from the Symposium on Biogenic Sulfur in the Environment (Saltzman and Cooper, 1989).

SULFATE REDUCTION BY BIOLOGICAL PROCESSES

In the + 6 oxidation state, the chemistry of sulfur is dominated by sulfuric acid and sulfate, which are rather involatile chemical species. As only this oxidation state is stable in the presence of oxygen, sulfate is the predominant form of sulfur in seawater, fresh waters and soils. Therefore, the reduction of sulfate to a more reduced sulfur species is a necessary prerequisite for the formation of volatile sulfur compounds and their emission to the atmosphere. In the global geochemical cycle, there are two types of biochemical pathways which lead to sulfate reduction: assimilatory and dissimilatory sulfate reduction. Table 2 shows estimates of the rates of sulfate reduction by these processes and compares these rates with the flux of sulfur through the atmosphere. Biological sulfate reduction has two major objectives: ( 1 ) the biosynthesis of organic sulfur compounds which are used for various purposes by the cell, e.g. in amino acids, and (2) the use of sulfate as a terminal electron acceptor to support respiratory metabolism in the absence of molecular oxygen. The former process is called assimilatory sulfate reduction (sulfur is being 'assimilated' ), the latter dissimilatory sulfate reduction. It is important to under- stand the ecological and biogeochemical differences between these two mechanisms: inadequate awareness of these differences between the two pathways of sulfate reduction has led to many of the misinterpretations and false as- sumptions found in the literature on the atmospheric sulfur cycle, e.g. the assumption that H2S is the major reduced sulfur compound emitted from the oceans. 4 M.O.ANDREAE

TABLE 2

Rates of sulfate reduction by major biogeochemical processes compared with anthropogenic and biogenic sulfur emissions to the atmosphere

Process Tmol year- Bacterial, dissimilatory sulfate reduction Coastal zone 2.2 Shelf sediments 6 Slope sediments 9 Total 12-20 a

Assimilatory sulfate reduction Land plants 3-6 Marine algae 10-20 Total 12-25 b

Anthropogenic emission of SO 2 ~ 3 Total biogenic sulfur gas emissions ~ 1.5 Total natural sulfur emission ~ 2 alvanov and Freney (1983). bEhrlich et al. (1977).

3 Tmol SO2 yr -I

. A~roposphe~

Assimtlatory sulfate reduction in the presence of 02 COS4 ~ (Plants and algae) (Land plants 3-6 Tmol yr -I) @/ o DMSj~ (Marine algae 10-20 Tmol yr "1)

_~_~__oxic___mixing barrier (redoxcline) ~ HZs ~ onoxic FeS/ Dissimilatorysulfate reduction in the absence of 0 2 (Anaerobic bacteria) (12-20 Tmol yr -I)

Fig. 1. Interactions in the global biogeochemical sulfur cycle.

Figure 1 gives a simplified, conceptual overview of the biogeochemical sulfur cycle. The global environment is subdivided into four compartments: atmosphere, biosphere, hydrosphere and lithosphere (the last standing for the sediments and rocks of the Earth's crust). The major pathway for the production of H2S is dissimilatory sulfate reduction, which is used by microbes to obtain thermodynamic energy in an oxygen-depleted environment. The oxidation of organic matter by available electron acceptors is the energetic basis GLOBAL BIOGEOCHEMICALSULFUR CYCLE 5 for essentially all life processes. Molecular oxygen is the thermodynamically most favorable electron acceptor which, if available, will be used preferen- tially in any ecosystem. However, if the supply of organic compounds exceeds that of oxygen, other electron acceptors (e.g. nitrate or sulfate ) are used when oxygen has been depleted. Dissimilatory sulfate reduction is therefore most commonly observed in marine environments where water circulation, and consequently oxygen availability, is limited (e.g. in stratified basins or in sedimentary pore waters) but where sulfate is easily available because of its relatively high concentration in seawater (28 mmol kg -~ ). The interface between the oxic and anoxic regimes (the 'redoxcline' ) is indicated in Fig. 1 by a dashed line through the biosphere and hydrosphere compartments. Under favorable conditions, the rate of sulfate reduction to HES in anoxic environments can be high, of the order of hundreds of mmol m -2 day -1. However, as the occurrence of this process is dependent on the existence of a mixing barrier which prevents oxygen from entering the system, the escape of H2S from the system will be limited by the same barrier. Furthermore, in the presence of oxygen, H2S provides an excellent substrate for microbial oxidation from which certain bacteria can obtain a substantial amount of energy. Such microorganisms tend therefore to be present in high numbers at the oxic- anoxic interface. They are very efficient in removing H2S and can completely oxidize this compound in a sediment layer only a fraction of a millimeter thick. Consequently, the very large amounts of H2S which are produced in the coastal and marine environment (Table 2) cannot usually be transferred to the atmosphere (Andreae, 1984, and references therein), but are either re- oxidized at the oxic-anoxic interface, or precipitated in the form of iron sul- fides and locked up in sediments and sedimentary rocks. Only under excep- tional conditions in shallow-water environments, can a fraction of the H2S escape: through temperature- or wind-driven turnover in estuaries, through scouring of muds in tidal channels, through bubbling of gas from anoxic environments, etc. Significant H2S emissions from the marine environment are therefore limited to nearshore environments such as estuaries and salt marshes.

Assimilatory sulfate reduction

In the form of a large variety of organosulfur compounds, sulfur is an essential element for biological organisms. Animals and protozoans are dependent on organosulfur compounds in their food to supply their sulfur requirement. All other biota - bacteria, blue-green algae, fungi, eucaryotic algae and plants - are able to carry out assimilatory sulfate reduction, i.e. they can synthesize organosulfur compounds from sulfate (Anderson, 1980). The biochemistry of assimilatory sulfate reduction has been studied mostly using the green alga Chlorella, therefore most of the following discussion refers specifically to this 6 M.O. ANDREAE organism and it is not altogether clear at this time how far these conclusions can be generalized to other organisms. The assimilation of sulfate to cysteine, the first organosulfur metabolite produced, is a complex, multi-step process (Fig. 2 ). Sulfate is taken up into the cell by an active transport mechanism, and inserted into an energetically activated molecule, APS (adenosine-5'-phosphosulfate), which can be further activated at the expense of one more ATP molecule to PAPS (3'-phosphoadenosine-5'-phosphosulfate). It is then transferred to a thiol carrier (RSH) and reduced to the -2 oxidation state. In contrast to nitrate assimilation, where the various intermediates are present free in the cytoplasm, sulfur remains attached to a carrier during the reduction sequence. In a final step, the carrier-bound sulfide reacts with O-acetyl-serine to form cysteine. Wilson et al. ( 1978 ) have suggested that under conditions when the availability of this or other endogenous sulfide acceptors is limiting the rate of cysteine synthesis, the volatilization of H2S could serve as a mechanism for removing excess reduced sulfur. Such volatilization has been observed from plants (Winner et al., 1981; Rennenberg, 1989), but its possible occurrence in marine algae has yet to be investigated. Cysteine serves as the starting compound for the biosynthesis of all other sulfur metabolites, especially the sulfur-containing amino acids homocysteine and methionine (Fig. 2). Cysteine and methionine are the major sulfur amino acids in plants and represent usually a very large fraction of the sulfur content of biological materials (Giovanelli et al., 1980). Glutathione (L-glu- tamyl-L-cysteyl-L-glycine) plays a variety of biochemical roles, including re- dox transfer reactions and the removal of H202 in chloroplasts. Methionine reacts with ATP to form S-adenosyl-methionine (SAM), the most important methyl group donor in methyl group transfer reactions in plants and algae. Transfer of a methyl group from SAM to methionine yields S-methyl-methionine, the precursor of dimethylsulfide in terrestrial plants. In marine algae, dimethylsulfonium propionate (DMSP) is formed in a multi-step process from methionine.

EMISSION OF DIMETHYLSULFIDE FROM THE OCEANS

Biosynthesis of dimethylsulfide

Dimethylsulfide was first identified in the gaseous emissions of the marine red macroalga Polysiphonia lanosa by Haas ( 1935 ). Challenger and Simpson (1948) showed that DMS was evolved from DMSP, which was present in substantial concentrations in the algal tissue. Later investigators found DMSP to be present in most algal species studied (Ackman et al., 1966; Tocher et al., 1966; Craigie et al., 1967; Granroth and Hattula, 1976; White, 1982 ). In a recent survey ofphytoplankton species in pure cultures, Keller et al. ( 1989 ) NH2 r-C~ ~N 0 0 © N~'~r-" ~ ,, , ~o I II I H.o.CH2"O'P'O'S'OH ~ Sulfated > L"~--'L--'--N--'~ "1 ()H ~)H polysocchorldes N ~/ (algae) =0 OH 0 C~ 5"-phosphoadenosine- rn 5'-phoephoeulfote HO-P-O O (PAPS) OH Glutathlone Protein K ~1~ H2S (glu-cys-gly) cyeteine N ADP RSH I ATP PPi NH2 i "~ATP (Thiol I O-ocetyl- Acetate ~%) / (38%] i 4 .L - N -- O O carrier) AP Ferredoxin | $erine so.--~----, ,, , ,.o.~..~-o-P-o-s-o.~.-s-so;-~J--,-'-s- \ J = HS-CH2CHCOOH " '-'.~--"-~ ~,. ~. su,f.. Su,fide Sulfate NH2 OH OH Cysteine M Adenosine-5'- phosphosulfate (APS) HS-CH2CH2CHCOOH NH2 1 Homocysteine Sulfolipid ( Diacylsulfoquinovosyl glycerol) H.C NH2i (All plonts, algae, cyanobacteria) H;~S'-CH2CH2CHCO0- ~ S-adenosyl-methlonlne~Plonts ,3C_S_CH2CH2CHCOOH (t%) NH2 S-methyl-methionine (SAM) Methionine

Protein OHCH2CH2CHCOOH NHo 1 H3C -S+_ CH2CH2CO0 - methlonine Homoeerine HsC r Dimethyleulfonium (58%) proplonote (DMSP)

CHsSCH3 CH2=CHCOOH Dimethyleulflde Acrylicacid (DMS) Fig. 2. Major metabolic pathways of sulfur in algae and plants. The percentages represent the approximate distribution of the major organosulfur compounds in Chlorella. 8 M.O. ANDREAE found that species of dinoflagellates, prymnesiophytes (in particular coccol- ithophores) and chrysophytes contained the highest DMSP concentrations. Maximum reported concentrations are generally in the range 0.2-0.4 mol DMSP 1- ~cell volume (Dacey and Wakeham, 1986; Dickson and Kirst, 1986, 1987; Keller et al., 1989 ). Groups of marine phytoplankton that usually contain only small amounts of DMSP include the chlorophytes, cryptomonads and cyanobacteria. There is compelling evidence that DMSP has an osmostatic and osmore- gulatory function in marine algae (Dickson et al., 1980, 1982; Vairavamurthy et al., 1985). The similarity in structure and chemical behavior between DMSP and other plant osmolytes, e.g. glycine betaine and proline, suggests that DMSP has similar enzyme-protective properties to these other 'compat- ible' solutes (Brown and Simpson, 1972). DMSP is produced from methionine by successive S-methylation, deamination, and decarboxylation. Its enzymatic cleavage produces DMS and acrylic acid on a one-to-one basis. Cantoni and Anderson (1956 ) have shown that the enzyme responsible for cleaving DMSP contains sulfhydryl groups and is bound to the membrane system. The release of DMS from the DMSP in algae occurs continuously at a relatively slow rate, but increases greatly when the organism is subjected to external stress, e.g. salinity changes, physical disturbance (e.g. stirring), or exposure to the atmosphere. This effect leads to pronounced DMS emissions from intertidal macro-algae during exposure at low tide. The physiological state of phytoplankton also appears to influence the rate of DMS emission, with the highest amounts being emitted during senescence (Nguyen et al., 1988). DMSP is also released by algae, and is cleaved in seawater to produce DMS (Turner et al., 1988, 1989). Although this reaction is extremely slow under abiotic conditions (Dacey and Blough, 1987 ), it is enhanced by the presence of microorganisms (Kiene, 1988). The relative contributions of the direct emission of DMS into seawater by algae and the breakdown of dissolved DMSP to DMS in the water column have not yet been determined. The biological or ecological function of DMS and DMSP excretion by algae also remains unknown at this time.

Marine chemistry and distribution of dimethylsulfide

In open ocean waters, DMS is the predominant volatile sulfur compound (Barnard et al., 1982; Andreae et al., 1983; Cline and Bates, 1983; Andreae and Barnard, 1984; Nguyen et al., 1984; Bates et al., 1987; Turner et al., 1988, 1989 ). Figure 3 shows a typical vertical distribution of particulate (intracellular) DMSP, dissolved DMSP and DMS, and chlorophyll (an indicator of phytoplankton biomass ) in the marine water column for the example of data from the northwestern Atlantic. The vertical distribution of DMS and DMSP GLOBAL BIOGEOCHEMICAL SULFUR CYCLE 9

0 0.1 0.2 0.3 0.4 0.5 0.6 pg L -~ (X) 0 0 2 4 6 8 I0 12 14 nrnol L'=(e,I:],ZI)

,oo

E 3: I-- G. I,i 13 28"05' N, 73"30'W 200 I MAY 1986

300

Fig. 3. Typical vertical distribution of particulate DMSP, dissolved DMSP and DMS, and chlorophyll a during the April-May 1986 cruise of R/V "Columbus Iselin" in the northwestern Atlantic Ocean.

in seawater as shown in Fig. 3 is typical for these compounds as well as for a number of other phytoplankton metabolites, e.g. dimethylsulfoxide (DMSO) and the methylarsenates (Andreae, 1979, 1980). The characteristic features of this distribution are the existence of a maximum at, or a few meters below, the sea surface, and a sharp decrease in DMS concentration near the level of 1% light transmission ( ~ 100 m in the example shown in Fig. 3 ). This depth represents the base of the euphotic zone, defined as the depth range in which enough light is present to permit the growth of phytoplankton. In deep water, DMS is present only at relatively low levels: ~ 0.03-0.15 nmol 1- t. In contrast to the distribution of DMS, the vertical profile of chlorophyll a shows a pronounced maximum at ~ 100 m. This deep chlorophyll maximum is a characteristic feature of the low-productivity regions of the central ocean basins and represents populations with very high intracellular chlorophyll levels. High levels of DMS are not to be expected here, as the shade flora characteristic of the deep chlorophyll maximum is usually dominated by the dinoflagellate genera Ceratium or Pyrocystis, neither of which is a major DMS producer (Keller et al., 1989). Furthermore, these phytoplankters are growing very slowly in the low-light conditions prevailing at the base of the euphoric zone. The sharp decrease in DMS concentration at the base of the euphoric zone suggests that there is consumption of DMS in the upper ocean, presumably by bacteria. The steep gradient in DMS concentration at the level of 1% light 10 M.O. ANDREAE penetration would then be explained by the relative dominance of bacterial consumption over the production of DMS by phytoplankton in this region of light-limited growth. The ability of bacteria to grow on DMS has been demonstrated both for anaerobic conditions (Zinder and Brock, 1978; Kiene, 1988 ) and for aerobic environments (Sivel~i and Sundman, 1975; Kanagawa and Kelly, 1986; Suylen and Kuenen, 1986 ). That such bacterial consumption of DMS actually takes place in the marine environment is also suggested by its behavior in anoxic basins (Wakeham et al., 1984) and in sedimentary porewaters (Andreae, 1985b). Studies on the anaerobic decomposition of DMS in sediment slurries (Kiene, 1988) showed that both sulfate-reducing and methanogenic bacteria are responsible for the removal of DMS and DMSP in marine sediments. The photochemical decomposition of DMS in surface seawater has also been demonstrated (Brimblecombe and Shooter, 1986 ). The concentration of DMS in the water column at any given place and time is thus the result of the interplay of DMS production by phytoplankton excretion and DMSP hydrolysis, DMS consumption by bacterioplankton and by photo-oxidation, volatilization of DMS across the air-sea interface, and downward mixing of DMS into the deep ocean by eddy diffusion (Andreae and Barnard, 1984; Wakeham and Dacey, 1989). The presence of DMS in the deep ocean at relatively constant levels suggests that the abiotic chemical breakdown of DMS under seawater conditions is a very slow process and does not contribute significantly to the removal of DMS from surface waters (Shooter and Brimblecombe, 1989). Based on data on the uptake of sulfate and the concentration of DMS in the water column of the Peru shelf upwelling region, I have estimated the relative rates of production, consumption, and ventilation loss of DMS. The results suggest that on the order of 1% of the sulfur assimilated by phytoplankton in this region is converted to DMS, and that roughly comparable amounts are lost by ventilation and by bacterial consumption (Andreae, 1985b). In a study of the cycle of methylated sulfur species in a coastal saline pond, Wakeham et al. (1987) concluded that, in this system, microbial consumption was the major sink for DMS, exceeding emission to the atmosphere by a factor of seven. These observations are consistent with the requirement that the release of DMS to the atmosphere should be only a relatively small fraction of the total sulfur assimilated by plankton, as most of the sulfur is required for other biochemical functions. For the assessment of the sea-to-air flux of DMS, knowledge of the ocean- wide distribution of DMS in the upper meter of the ocean is required. As it is not realistic to try to measure DMS everywhere, we have attempted to find relationships between DMS and other observable parameters which could be used for the prediction of DMS levels in regions for which no direct measurements of its concentration exist. A measure of phytoplankton biomass, e.g. GLOBAL BIOGEOCHEMICAL SULFUR CYCLE | 1 chlorophyll a concentration, or of phytoplankton productivity, e.g. 14C uptake, would be an obvious candidate for such a predictor variable. Chloro- phyll would be especially attractive as it can be estimated by remote sensing either from aircraft or from satellites. Our attempts to find consistent relationships between chlorophyll and DMS have met with mixed success however. When we subject our entire data set on DMS and chlorophyll concentrations to regression analysis, we obtain values of r 2 near 0.3, which, because of the large number of data (over 1000), are highly significant. As the value of r 2 suggests, however, this correlation explains only about 30% of the variability. Although such analysis of large data sets (as well as the vertical distribution of DMS in the marine water column) demonstrates a significant overall re- lationship between the distributions of DMS and phytoplankton in the surface ocean, it is difficult to find a clear correlation between total plankton abundance and DMS concentration within a given region. This is most probably due to the substantial differences in the DMS output rate between different plankton species (Andreae et al., 1983; Barnard et al., 1984; Turner et al., 1988; Keller et al., 1989). In some cases, a single phytoplankton species can be responsible for most of the DMS production in a given oceanic region, e.g. Phaeocystis poucheti in the Bering Sea shelf region (Barnard et al., 1984 ) and on the shelf west of the English Channel (Holligan et al., 1987). Our data also show that the DMS concentrations in the low-productivity regions of the oceans, especially the subtropical gyres, are substantially higher than expected on the basis of the abundance of phytoplankton in these areas. An example of this behavior is shown in Fig. 4, where the distributions of chlorophyll, DMSP and DMS along a cruise track in the northwestern Atlan- tic are compared. The surface water temperature measured along the cruise track is shown in Fig. 4 as a water-mass indicator for the warm waters of the Gulf Stream and the Sargasso Sea and for the cold waters of the Mid-Atlantic Bight. We see that the consistently highest DMS levels are found in the oligotrophic waters of the Sargasso Sea, whereas the very high phytoplankton densities in the frontal areas off Cape Hatteras are not reflected in significantly elevated DMS levels. This is most probably due to species-related effects, as the blooms of Cape Hatteras are dominated by diatoms, which tend to produce little DMS, whereas the coccolithophorid species common in the tropical gyres are prolific emitters of DMS. The underlying reason for the relatively high abundances of DMS and DMSP in oligotrophic waters may be related to the scarcity of nitrate in these environments: to achieve the required high internal osmotic pressure to balance that of the seawater surrounding the cell (osmolarity ~ 1.1 mol 1-~), marine microorganisms must produce a substantial amount of osmoregula- tory substances. Many of the preferred osmolytes, however, contain nitrogen (e.g. proline, betaine). This does not present a serious problem in the pro- 12 M.O. ANDREAE

4 , , , , , , , , , i ' New Yo~'k~ ' 40 oN • Chl a I 30 ~ "" i UNITED 2o ? STATES ~L"~\ J "-,,..-,i"i Jl/=,]"---;;-;e; '.L. ~i~A fI' ~°~ temperature 55 ~.~ H-atteras / ,u io

01 , / /'Flor!da Sargos ~ /coost s;o , , , , , , , , , , 3C

10 800W 75 70

C) i i i i i i i i i i

, , , , , , , , , , DMS 6

o E c

i i i i i i i i I 22 23 24 25 26 27 28 29 :30 I 2 APRIL 1986 MAY

Fig. 4. Cruise track of R/V "Columbus Iselin", April 22-May 3, 1986, with surface water temperatures and concentrations of chlorophyll a, particulate DMSP and DMS measured during this cruise. ductive regions, where nitrate is present in the water column in relatively high concentrations. On the other hand, in the nutrient-depleted regions of low productivity, e.g. the oceanic gyres, the use of a sulfur osmolyte (DMSP) instead of a nitrogen osmolyte would make all bound nitrogen available for essential uses in amino acids, etc. Although the thermodynamic energy required to assimilate sulfate (involving the reduction from the oxidation state + 6 to -2) is higher than that needed to assimilate nitrate (reduction from oxidation state -t-5 to -3), it is comparable to the energy requirement for nitrogen fixation (reduction from oxidation state 0 to - 3 ). Marine blue-green algae solve the problem of nitrogen limitation by fixing (i.e. assimilating) molecular nitrogen, and therefore would not benefit from the synthesis of a sulfur-containing osmolyte. Consistent with this argument, we have found that Synecchococcus sp., a common blue-green alga of oceanic gyres, produces nei- GLOBAL BIOGEOCHEMICAL SULFUR CYCLE 13 ther DMSP nor DMS. Nitrogen fixation is, however, not available to other algal taxa. These organisms could, therefore, benefit from replacing some of their nitrogen requirement with a molecule which contains sulfur in lieu of nitrogen (Andreae, 1986). Some experimental support for this hypothesis has been provided by laboratory experiments in which planktonic algae grown at high nitrate levels showed lower intracellular DMSP concentrations than algae grown under nitrate-limited conditions (Turner et al., 1988 ). However, although this hypothesis could explain the increased levels of DMSP in species living in nutrient-depleted regions, it does not explain why some of this DMSP is broken down to DMS and excreted.

Estimating the air-sea flux of dimethylsulfide

Volatile substances are transferred across the air-sea interface by a combi- nation of molecular and turbulent diffusion processes, which are still poorly understood and for which no entirely satisfactory physical and mathematical models are available. A discussion of the state of the art in this field is given in the review by Liss and Medivat (1986). The sea-to-air flux is proportional to the air-sea concentration gradient and the gas transfer velocity across the air-sea interface. The atmospheric concentration of DMS is several orders of magnitude below the value in equilibrium with seawater. It can therefore be ignored for the purpose of estimating the sea-to-air concentration gradient and only the concentration of DMS in seawater is required to estimate the emission flux of DMS. As a result of numer- ous cruises conducted by several groups (Barnard et al., 1982; Andreae et al., 1983; Cline and Bates, 1983; Bingemer, 1984; Andreae and Barnard, 1984; Nguyen et al., 1984; Bates et al., 1987 ), we now have a relatively good picture of the distribution of DMS in the World Oceans. These data are summarized in Table 3, which is based on the compilation of DMS data in Andreae ( 1986 ). The data are organized by biogeographical regions as defined by Koblentz- Mishke et al. ( 1970); averages for each of these regions are used together with an estimate of their areal extent for the prediction of the flux of DMS from each region. The data base used for Table 3 does not contain any measurements from the Southern Ocean; recent work by Berresheim (1987) has shown, however, that oceanic emissions of DMS in this region are similar to those found in temperate regions. To obtain the DMS transfer velocities used in the flux calculations in Table 3, we adjusted the radon transfer velocities of Peng et al. ( 1979 ) and of Sme- thie et al. ( 1985 ) by assuming that the transfer velocity is proportional to the square root of the diffusivity. If we use the global average ~4CO2 transfer velocity ( ~ 21 cm h- t: Liss and Merlivat, 1986 ) to estimate the DMS flux (after adjusting for diffusivity and dissociation effects), we obtain a significantly higher flux: 1.6 instead of 1.2 Tmol DMS year- ~. This is probably due to the 14 M.O. ANDREAE

TABLE 3

DMS concentrations and fluxes for the world oceans

Biogeographic region Area Mean concentration Total flux ( 10 6 km 2 ) (nmol 1- l ) (Tmol S year- ~)

Oligotrophic (tropical/low 148 2.4 0.2-0.6 productivity ) Temperate 83 2.1 0.1-0.3 Upwelling (coastal 86 4.9 0.2-0.7 and equatorial) Coastal/shelf 49 2.8 0.1-0.2

Mean: 3.0 Total: 0.6-1.7 fact that the 14CO2transfer velocity integrates over the whole year, whereas the radon transfer velocity is based almost entirely on summer data when wind speeds are lower. In view of the extensive data on DMS concentrations in the surface ocean as presented in Table 3, I feel that the major uncertainty about the sea-to-air flux of DMS now rests in the uncertainties associated with the use of the 'stagnant-film' model, and in particular with the estimation of the transfer velocities. This uncertainty may be as large as a factor of two. From Table 3 we can reach some interesting conclusions. First, there is a surprisingly small difference in the average DMS concentrations for the different regions. The average for the oligotrophic areas is essentially the same as for the transitional areas of the temperate oceans, and both types of openocean regimes have DMS concentrations similar to the coastal waters. Only in upwelling areas do we observe a substantially higher average concentration, but even here the difference is only a factor of two. One reason for these relatively small differences is that the tropical regions have relatively high DMS concentrations year-round, whereas in temperate regions, especially the coastal temperate areas, there is a pronounced seasonality with low values during the cold season. Second, the large areas of low and moderate biological productivity contribute amounts of DMS to the atmosphere comparable to those from the relatively small regions of high productivity in the upwelling regions and the coastal areas. This is in contrast to earlier views which had assumed that the biogenic sulfur flux from the oceans would be dominated by localized 'hot-spots' of biological productivity. Finally, we find that the estimate for the global flux has by now become very robust relative to the addition of new data (even including data from a number of different groups). Although the number of data points in Table 3 is ~ 2.5 times greater than in the comparable table in Andreae and Raemdonck (1983 ), the estimate for the global mean DMS concentration has only changed from 3.2 to 3.1 nmol GLOBAL BIOGEOCHEMICAL SULFUR CYCLE 15

1-~, and that for the global flux remains unchanged at ~ 1.2 Tmol year -1. Using a data set from the Pacific Ocean only, Bates et al. ( 1987 ) obtained a lower mean DMS concentration (~ 1.8 nmol 1-1), and a correspondingly lower flux of 0.5 Tmol year- t with an estimated uncertainty of a factor of two. In view of the large uncertainties associated with the 'stagnant-film' model, it seems very important that independent methods be developed to test the predictions based on this model. However, alternative methods to determine the flux, e.g. the eddy-correlation or gradient techniques, still face large experimental difficulties. No rapid-response sensor which would make the eddy- correlation technique possible is available for DMS, or in fact any of the reduced sulfur gases. The gradient method has been used on board ship by Bin- gemer (1984) and by Nguyen et al. (1984) by sampling at different levels above the waterline. Although the results compare well with predictions from gas transfer calculations, they may contain substantial error because of the influence of the ship on the air flow characteristics. Because of the difficulty of simulating a realistic wave climate inside a flux chamber, direct measurements of sulfur gas fluxes across the air-sea interface by the chamber technique have not been attempted.

Chemical reactions and transformations of dimethylsulfide in the marine atmosphere

After its transition from the ocean into the atmosphere, DMS can react with a variety of oxidizing atmospheric trace species. The rates and pathways of DMS oxidation in the atmosphere have been reviewed recently (Andreae, 1986; Yin et al., 1986; Toon et al., 1987; Plane, 1989). Figure 5 gives a sche- matic description of the major atmospheric oxidation reactions of DMS. Cur- rently available information suggests that the reaction with hydroxyl radical (OH) is the predominant oxidation process, with a potentially significant contribution from the reaction of DMS with the nitrate (NO3) radical. The latter reaction is relevant only in moderately to highly polluted airmasses, where the concentrations of NOx and ozone are high enough to lead to significant night-time production of NO3. Consequently, NO3 may be the most important oxidant for DMS in polluted ocean regions, e.g. over the western North Atlantic, whereas over the remote oceans it probably does not contribute significantly to DMS oxidation (Andreae et al., 1985). The reaction of DMS with the iodine oxide (IO) radical to form DMSO has been proposed as a major sink for DMS (Barnes et al., 1987), but recent work suggests that the reaction rate constant between IO and DMS may have been overestimated by a factor of 1000 (P.H. Wine, Georgia Institute of Technology, personal com- munication, 1989), which would make this reaction negligible compared with the OH oxidation. 16 M.O.ANDREAE

03 02 OH • H20

NO3 HNO3 CH3S. o.yCH3SCH 3 v CH3SO2" HO~. CH3SOH CH~. i DMSO Met~c acid Sulfur dioxide

Fig. 5. Reaction pathways for the oxidation of DMS by OH, NO3 and IO radicals.

A considerable amount of work has been done to determine the rate of the reaction between DMS and these radicals; however, the actual reaction se- quences and products are still uncertain. Observations on the relative abundances of SO2 and the other possible DMS oxidation products (DMSO, methanesulfonic acid (MSA)) in the marine atmosphere suggest that SO2 is the dominant product (Saltzman et al., 1983; Andreae, unpublished data, 1988 ). However, under specific circumstances, e.g. over the Southern Ocean and in the subantarctic region, MSA appears to be a major product of DMS oxidation (Ayers et al., 1986; Berresheim, 1987; Berresheim et al., 1990). The information on the atmospheric abundance of DMSO, produced by the minor OH addition reaction sequence and possibly by the DMS + IO reaction (Fig. 5), is currently limited to a few measurements in marine rain (Andreae, 1980, and unpublished data, 1988) and some recent measurements of its gas-phase concentration (Harvey and Lang, 1986; Andreae, unpublished data, 1988 ). These measurements are not sufficient to assess the role of DMSO as a product of DMS oxidation in the marine atmosphere, and further studies on the abundance of this compound should be conducted. Un- certainty also exists about the fate of DMSO in the marine atmosphere. It reacts rapidly with OH, probably resulting in the formation of SO2 and MSA. However, it is also highly water-soluble, so that dry deposition to the sea- surface may also play an important role as a sink for DMSO. SO2 is rapidly oxidized to sulfate in the marine boundary layer, both by gas-phase and liquid-phase processes (Calvert et al., 1985; Bonsang et al., 1987 ). Because of their low volatility, sulfate and MSA are present predomi- nantly in the form of aerosol particles, even though in the case of MSA a significant amount (up to 30%) may be present in vapor form (Andreae, GLOBAL BIOGEOCHEMICAL SULFUR CYCLE 17

unpublished data, 1988 ). MSA is very highly soluble and will be efficiently scavenged by cloud droplets and precipitation (Clegg and Brimblecombe, 1985). Based on intensive field studies during the last few years, we now have a reasonably good idea of the concentrations and vertical distribution of DMS in the lower troposphere over most of the major ocean regions, both from shipboard measurements (Andreae and Raemdonck, 1983; Nguyen et al., 1984; Andreae et al., 1985; Berresheim, 1987; Saltzman and Cooper, 1988; Church et al., 1990) and from aircraft data (Ferek et al., 1986; Van Valin et al., 1987; Andreae et al., 1988; Berresheim et al., 1990). These data sets show a rather consistent distribution pattern of DMS in the marine atmosphere: at ground level, DMS concentrations are typically on the order of 20-200 pptv (parts per trillion by volume), depending on ocean area, season, etc. This concentration remains nearly constant with altitude through the subcloud mixed layer (typically ~ 1 km), and then decreases rapidly with altitude in the free troposphere. We can use this information, combined with available data on the concentrations of the products of DMS oxidation in the marine atmosphere and the estimates of DMS emission from the oceans and the deposition fluxes of the oxidation products, to assess the validity of our knowledge of the major features of the marine biogenic sulfur cycle. This is done by comparing the measured concentrations of atmospheric sulfur species with predictions from model calculations. Using either simple box models (Andreae, 1986; Berresh- eim, 1987; Berresheim et al., 1990) or time-dependent numerical models (Ferek et al., 1986 ), and assuming DMS fluxes of the order estimated in Ta- ble 3 as the only input of gaseous sulfur, we find that we can construct a reasonably consistent picture of the cycle of biogenic sulfur for the marine boundary layer. This is demonstrated in Fig. 6, where the vertical distributions of DMS, MSA, non-seasalt sulfate and the MSA/non-seasalt sulfate ratio over two temperate ocean areas are shown: the northeastern Pacific off the state of Washington (U.S.A.), and the Southern Ocean off Tasmania (Aus- tralia). In the subcloud mixed layer, comparable concentrations of all sulfur species are present, the amounts of which can largely be explained on the basis of the oxidation of DMS (Andreae et al., 1988; Berresheim et al., 1989, 1990). This suggests that DMS is indeed the major source of non-seasalt sulfur in the remote marine atmosphere, and that its flux is of the order of 1.0- 1.2 Tmol year-~ as estimated on the basis of its concentration in surface seawater. Models which include intermittent, rapid transport of boundary layer air into the upper troposphere by large convective cloud systems predict that DMS can make a significant contribution to free tropospheric SO2 as well (Gidel, 1984; Chatfield and Crutzen, 1984). These predictions were substantiated by our measurements in the marine troposphere near Barbados, which showed 18 M.O. ANDREAE

(0) DMS MSA nss-S042- MSA/nss- S04z- 4 NE PACIFIC 3 km 2

i i i ~ i 2o ~o ' o ~ ,o,~ o 50 I00 150 0 0.04 0.08 (b) "r

~I SOUTHERN OCEAN

krn

Ii CO 20 40 J 0 5 I0 15 0 5O ,60 ,~o o 0.4 d.8' pptv pptv pptv tool/tool Fig. 6. Vertical distributions of DMS, MSA, non-seasalt sulfate (nss-SO2- ) and the MSA/nss- SO42- ratio over (a) the northeasternPacific Ocean (May 1985 ) and (b) the Southern Ocean near Tasmania (December 1986). (Note differentMSA/nss-SO42- scales.) that convective transport increased the free tropospheric DMS levels by a factor of 10 over the values found during non-convective conditions. The resulting rate of SO2 production can account for much, if not all, of the SO2 and consequently the sulfate aerosol, in the free troposphere at least in tropical regions (Ferek et al., 1986). However, during aircraft experiments over the eastern North Pacific we observed elevated concentrations of aerosol sulfate in the free troposphere which could not be explained on the basis of DMS oxidation (Andreae et al., 1988) (Fig. 6a). Airmass trajectories and radon measurements both pointed towards long-range transport from Asia as the most likely source of these elevated sulfate levels. The absence of similar concentrations in the free troposphere over the Southern Hemisphere oceans, where continental sources are much less important, is consistent with this finding (Berresheim et al., 1990) (Fig. 6b ).

The climatic significance of marine dimethylsulfide emission

In a recent paper, Charlson et al. ( 1987 ) proposed the existence of a climatic feedback loop involving marine biogenic sulfur (Fig. 7 ). The emission of DMS by marine phytoplankton leads to the presence of the gas in the ma- GLOBALBIOGEOCHEMICAL SULFUR CYCLE 19

Radiation budget_

Cloud condensation Global temperature nuclei

Sulfate aerosol ~'~ ~---lv ,+ Climate feedbacks

D~M: +or-? -I- Atmosphere DMS Ocean .?~ Phytoplankton ? Marine + or-. abundance and ~ ' speciation ecology

Fig. 7. Proposed feedback cycle between climate and marine DMS production. The pluses and minuses indicate if an increase in the value of the preceding parameter in the cycle is expected to lead to an increase ( + ) or decrease ( - ) in the value of the subsequent parameter.

rine atmosphere where it is oxidized, forming sulfate (and methanesulfonate) aerosol. This aerosol provides the majority of cloud condensation nuclei (CCN) over the remote oceans. Model calculations show that the albedo (reflectivity) of clouds over the remote oceans increases with increasing CCN concentration. As the global radiation balance, and thus the global mean temperature, is sensitively dependent on the albedo of marine clouds, changes in global mean temperatures of the order of a few degrees centigrade are pre- dicted to result from a change in DMS flux by a factor of two. Currently, we have little information on which climatic, environmental and ecological factors control the global rate of DMS production and its flux to the atmosphere. As only a modest fraction of the marine primary producers (i.e. the dinoflagellates and prymnesiophytes) is responsible for the production of most of the DMS, global DMS production is not closely tied to global primary production. This decoupling between DMS emission and primary production, which is tightly constrained by the global carbon cycle, makes it plausible that the abundance of DMS-producing phytoplankton may have varied over a factor of two over glacial/interglacial time periods. Such variations in the marine DMS source may have caused 1 °C variations in global 20 M.O. ANDREAE temperature, reinforcing the effect of changes in atmospheric CO 2 levels ( ~ 0.6 ° C ) and solar radiation intensity (0.2 ° C ) (Legrand et al., 1988 ).

CARBONYL SULFIDE

Carbonyl sulfide (COS) is the most abundant atmospheric sulfur species in the remote troposphere, with an average concentration near 500 pptv. Be- cause of its low reactivity in the troposphere and its correspondingly long residence time (of the order of 1 year), it is the only sulfur compound which can enter the stratosphere (with the exception of SO2 injections during vio- lent volcanic eruptions ). The input of COS is considered to be responsible for the maintenance of the sulfate aerosol layer in the stratosphere during volcan- ically quiescent periods (Servant, 1986). Therefore, even a relatively small COS source flux can be of considerable importance in atmospheric chemistry. Carbonyl sulfide is present in surface seawater at concentrations of ~ 0.03- 1.0 nmol l- ~ (Rasmussen et al., 1982; Ferek and Andreae, 1983, 1984; Turner and Liss, 1985 ). The observed concentrations are almost always higher than the equilibrium concentration relative to the overlying atmosphere, so that a net sea-to-air flux exists essentially from the entire ocean surface. Johnson ( 1981 ) has speculated that the ocean should be a sink for COS because of its hydrolysis at the slightly alkaline pH of seawater. This suggestion is clearly not supported by the measured COS supersaturation ratios across the air-sea interface. Pronounced diel variations of the COS concentration in surface seawater ( Fig. 8 ) suggest that COS is produced there by photochemical reactions (Ferek and Andreae, 1984). Laboratory experiments with seawater and with solu- tions of organosulfur compounds in distilled water showed that seawater sulfate did not participate in the reaction, and that only the presence of dissolved organic sulfur compounds, dissolved 02 and light were necessary to produce COS. Carbonyl sulfide was formed by irradiation of a variety of organic sulfur compounds commonly found in biological materials, e.g. cysteine, methionine, glutathione and dimethylsulfonium proprionate. The mechanism of this reaction is not yet known, but it is likely that short-lived, photochemically produced radicals (e.g. OH) are involved. The photochemical production of COS in seawater is the result largely of the UV-B part of the solar spectrum, and is strongly enhanced by the presence of photosensitizing compounds, e.g. humic and fulvic acids (Zepp and Andreae, 1989). The presence or absence of living micro-organisms - planktonic algae or bacteria - has no influence on the rate of formation of COS in seawater. It appears that the role of organisms in the production of COS in seawater is limited to the synthesis of dissolved organic sulfur compounds which are then abiotically photolyzed to COS. The dependence of the rate of COS formation on the concentration of dissolved organic sulfur in seawater is reflected by the GLOBALBIOGEOCHEMICAL SULFUR CYCLE 21

difference between the COS supersaturation measured in coastal and open ocean waters (Fig. 8 ). An attempt to obtain a representative estimate of the sea-to-air flux of COS is presented in Table 4, where I have divided up the ocean surface into the same biogeographic regions as used in Table 3. Then, based on our (diurnally averaged) data on the supersaturation of COS in surface seawater relative to the overlying atmosphere and the average temperature of the surface ocean in these regions, I have calculated the flux of COS across the air-sea interface for these regions (the piston velocities for COS are a factor of 1.3 higher than for DMS, because of the higher diffusivity of COS ). We see that, in contrast to DMS, the flux of COS is dominated by the high-productivity regions, especially the coastal and shelf areas. As a result of the low levels of COS in oligotrophic areas, they contribute little to the global flux, which I estimate to be ~ 11 Gmol year- ', similar to previous flux estimates (Rasmussen et al., 1982: ~ 10 Gmol year-'; Ferek and Andreae ( 1983): ~ 16 Gmol year-= ).

i i I i I I I i I i i FLORIDA BAY - BAHAMAS 8-21 NOV 1983 IT 15

0 m l l NEARSHORE I--

I.- <[ (n u~ o u l J . OPENOCEAN "',,

Oi i i i I'I = ' I i I I 3 05 07 09 15 15 17 19 21 25 LOCAL TIME (h}

Fig. 8. Mean diurnal variation of COS in surface seawater during a cruise of R/V "Bellows" in November 1983. The concentration of COS is indicated as a saturation ratio, i.e. the ratio between the measured concentration and the concentration in equilibrium with ambient air with 500 pptv COS. 22 M.O. ANDREAE

TABLE 4

COS concentrations and fluxes for the world oceans

Biogeographic region Area Mean Flux/area Total flux ( 10 6 km 2 ) concentration (nmol m- 2 day- ~) (Gmol year- ~) (pmol 1- l )

Oligotrophic (tropical/low 148 11.3 14.0 0.8 productivity ) Temperate 83 20.3 45.0 1.4 Upwelling (coastal and 86 24.1 64.0 2.0 equatorial) Coastal/shelf 49 95.0 373.0 6.7

Mean: 27.6 Total: 10.9

FORMATION AND EMISSION OF HYDROGEN SULFIDE AND CARBON DISULFIDE

Hydrogen sulfide

There are few data on the concentration of dissolved H2S in surface seawater, and only a few reliable measurements of H2S in the marine atmosphere; therefore the air-sea exchange flux of this compound is difficult to estimate. H2S is oxidized rapidly in oxygenated seawater: half-lives of the order of a few hours are reported (Almgren and Hagstr/Sm, 1974); other work- ers have found values as high as 50 h, however (Chen and Morris, 1972 ). The most reliable measurements appear to be those of Millero et al. ( 1987 ), who found a half-life of 26 h at 25 °C. Cutter and Krahforst ( 1988 ) have recently developed a technique for the determination of HaS in seawater and have observed concentrations of < 0.1-1.1 nmol 1-1 in surface seawater from the western Atlantic Ocean. The concentrations show a pronounced diel variation, with a maximum just before sunrise. The production mechanism of this H2S remains unclear, but its vertical distribution in the ocean suggests that bacterial reduction in microbial microenvironments may play an important role. It must, however, be remembered that biological processes, e.g. in plants, can result in the production and release of substantial amounts of H2S even in the presence of oxygen. This is especially true in the presence of high ambient sulfate concentrations, as is the case in seawater. H2S has been observed in the marine atmosphere at levels of a few pptv to a few tens of pptv (Slatt et al., 1978; Delmas and Servant, 1982; Herrmann and Jaeschke, 1984; Cooper and Saltzman, 1987). Cooper and Saltzman (1987) found a positive interference in the determination of H2S by the method used by the previous authors (trapping on AgNO3-impregnated filters and determination by the quenching of the fluorescence of fluorescein GLOBAL BIOGEOCHEMICAL SULFUR CYCLE 23 mercuric acetate), and suggested that the mean concentration of H2S in the marine boundary layer does not exceed 10 pptv (Saltzman and Cooper, 1988 ). At these levels, H2S in the atmosphere is near thermodynamical equilibrium for the concentrations in surface seawater observed by Cutter. To obtain an estimate of the rate of H2S oxidation in the marine atmosphere, we can simply use an average concentration of 10 pptv with a scale height of 2 km, a diurnally averaged OH concentration of 2 × 106 molecules cm -3, and the measured reaction rate for the oxidation of H2S by OH ( 5 × 10-12 cm 3 molecules- l s- 1: Cox and Sheppard, 1980 ). The resulting estimate, 0.09 Tmol year- 1, is an upper limit for the sea-to-air flux of H2S, and is much smaller than the DMS flux of ~ 1.2 Tmol year-1. Based on their measurements in the Caribbean and the Gulf of Mexico, Saltzman and Cooper ( 1988 ) suggest that the oxidation of H2S accounts for only 11% of the production of biogenic non-seasalt sulfate in the remote marine boundary layer, the rest being produced by the oxidation of DMS. It is not clear, however, if the source of the HRS found in the marine troposphere is necessarily the ocean surface or if other processes could be responsible for its presence. For example, advection from coastal regions, where H2S is emitted from salt marshes, may supply some of this H2S. This hypothesis is supported by recent measurements over the western North Atlantic, which show a clear correlation between atmospheric concentrations of H2S and radon, an indicator of continental airmass origin (Andreae, unpublished data, 1989 ). On the other hand, McElroy et al. (1980) have speculated that atmospheric reactions of COS and CS2 with OH radical could produce the necessary amounts of HaS. How- ever, this suggestion has not yet been verified experimentally.

Carbon disulfide

The presence of CS2 in seawater was first observed by Lovelock (1974), who measured an average concentration of 14 pmol S ( CS2 ) 1-1 in 35 samples taken in the open Atlantic Ocean. Inshore values were about an order of magnitude higher. Turner and Liss ( 1985 ) also report the presence of high levels of CS2 in coastal waters off England, but give quantitative information for only a few samples with values near 300 pmol S(CS2) 1-i. They found substantially higher concentrations in the low-salinity region of an estuary (up to ~ 2 nmol S 1-1 ). It is possible that much of the CS2 found in coastal waters is the result of the diffusion of this substance from the porewaters of the underlying sediments. This would be consistent with the relatively high concentrations and fluxes of CS2 observed in coastal marsh environments (Adams et al., 1981, Steudler and Peterson, 1984 ). CS2 could be formed there either by fermentation reactions of organosulfur compounds or by 'pulp-mill'-type reactions of terrigenic plant matter with dissolved polysulfides originating from bacterial dissimilatory sulfate reduction. 24 M.O. ANDREAE

TABLE 5

CS2 concentrations and fluxes for the world oceans

Region Area Mean Flux/area Total flux ( 106 km 2) concentration ( nmol m - 2 day- ~) ( Gmol S year- t ) (pmol S year -~ ) Open oceans 310 16 45 5.1 Coastal/shelf 50 33 90 1.6

Mean: 18 Total: 6.7

We have recently determined CS 2 in open ocean and coastal seawater from the North Atlantic, and have observed mean concentrations of 16__ 8 and 33_+ 19 pmol S(CS2) 1-l, respectively (Table 5; Kim and Andreae, 1987), somewhat higher than Lovelock's results. From these data, we estimate a flux of ~ 7 Gmol S year- i in the form of CS2 from the World Ocean surface, about 0.6% of the DMS flux. The photochemical oxidation of CS2 produces one molecule each of SO2 and COS per molecule of CS2 oxidized. Thus, the marine emission of CS2 provides a significant indirect source of COS, whereas it is clearly inconsequential as a source of tropospheric SO2.

REFERENCES

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