Marine Chemistry, 30 (1990) 1-29
Elsevier Science Publishers B.V., Amsterdam
1
Ocean-atmosphere interactions in the global
biogeochemical sulfur cycle*
M e i n r a t O. A n d r e a e
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). Attempts 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 geophysiological 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. Hydrogen 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
o v e r r e m o t e o c e a n r e g i o n s ( e . g . A n d r e a e e t al., 1 9 8 8 ). T h e h u m a n p e r t u r b a 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 y r - 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 unpolluted 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
Seaspray
Dust
Total
particulates
Sulfate
Other Total
1.2-10
0.1-1
1.3-11
Volcanoes
Soils and
plants
Coastal
wetlands
Biomass
burning
Oceans
(gases)
0.23-0.29 0.03
0.1-0.3
Totalgases
-
0.03
0.08
?
-
0.3-0.4
0.0003
+0.02
<0.1
-
?
0.03
0.3-0.4
0.15-0.4
0.06
0.02
0.002
-
0.004
-
?
?
?
0.05-0.2 0.011
0.6-1.6
0.01
-
?
1.1-1.8
0.2-0.6
0.6-1.7
0.03-0.04
0.03
1.2-2.8"
"Equivalent to 38-89 Tg S year-~.
0.004
0.0003
0.006-0.12 0.02-0.025
1.2-10
0.1-1
1.3-11
0.003
0.00-0.04
<0.1
t> 0.08
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 understand 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 assumptions 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
Shelf sediments
Slope sediments
Total
2.2
6
9
12-20a
Assimilatory sulfate reduction
Land plants
Marine algae
Total
3-6
10-20
12-25 b
Anthropogenic emission of S O 2
Total biogenic sulfur gas emissions
Total natural sulfur emission
~ 3
~ 1.5
~2
alvanov and Freney (1983).
bEhrlich et al. (1977).
3 Tmol SO2 yr -I
. A~roposphe~
COS4
~
@/
DMSj~
~
o
_~_~__oxic___mixing
HZs ~
onoxic
FeS/
Assimtlatory sulfate reduction
in the presence of 02
(Plants and algae)
(Land plants 3 - 6 Tmol yr -I)
(Marine algae 10-20 Tmol yr "1)
barrier (redoxcline)
Dissimilatory
sulfate 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 o f the biogeochemical sulfur cycle. The global e n v i r o n m e n t is subdivided into four compartments: atmosphere, biosphere, hydrosphere and lithosphere (the last standing for the
sediments and rocks o f the Earth's crust). The major pathway for the production o f H2S is dissimilatory sulfate reduction, which is used by microbes to
obtain t h e r m o d y n a m i c energy in an oxygen-depleted environment. The oxidation o f 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 preferentially 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 oxicanoxic 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 reoxidized at the oxic-anoxic interface, or precipitated in the form of iron sulfides and locked up in sediments and sedimentary rocks. Only under exceptional 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-glutamyl-L-cysteyl-L-glycine) plays a variety of biochemical roles, including redox 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 )
--
ADP
i "~ATP
O O
carrier) AP
RSH
(Thiol
Ferredoxin
1
(APS)
Adenosine-5'- phosphosulfate
OH OH
(t%)
(SAM)
\
J
~%)
/
(38%]
Protein
cyeteine
HS-CH2CH2CHCOOH
NH2
Homocysteine
= HS-CH2CHCOOH
NH2
Cysteine
O-ocetyl- Acetate
$erine
Glutathlone
(glu-cys-gly)
CH2=CHCOOH
Acrylic acid
CHsSCH3
proplonote (DMSP)
HsC r Dimethyleulfonium
H3C-S+_CH2CH2CO0 -
Methionine
Protein
methlonine
(58%)
,3C_S_CH2CH2CHCOOH
S-adenosyl-methlonln~Plonts
e
NH2
Su,fide
|
I
H2S
I
Dimethyleulflde
(DMS)
OH
NHo
1
CH2CH2CHCOOH
Homoeerine
S-methyl-methionine
H.C
NH2
i
H;~S'-CH2CH2CHCO0- ~
,, , ,.o.~..~-o-P-o-s-o.~.-s-so;-~J--,-'-s'-'.~--"-~
~,. ~.
su,f..
.L - N
NH2
Sulfolipid
( Diacylsulfoquinovosyl glycerol)
(All plonts, algae, cyanobacteria)
4
PPi
~1~
OH
polysocchorldes
(algae)
~ Sulfated
Fig. 2. Major metabolic pathways of sulfur in algae and plants. The percentages represent the approximate distribution of the major organosulfur compounds in Chlorella.
Sulfate
so.--~----,
"
i
ATP
(PAPS)
NH2
N ~
0,, 0,
N~~'~r-"
I II
I H.o.CH2"O'P'O'S'OH
L"~--'L--'--N--'~ "1
()H ~)H
N
~ /
OH 0
5"-phosphoadenosine5'-phoephoeulfote
HO-P-O
M
K
N
C~
r©
~o
>
=0
C~
rn
O
8
M.O. ANDREAE
found that species of dinoflagellates, prymnesiophytes (in particular coccolithophores) 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 osmoregulatory 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 'compatible' 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
0
00
0.1
2
0.2
4
0.3
6
0.4
8
9
0.5
I 0
0.6
12
pg L-~ (X)
14 nrnol L'=(e,I:],ZI)
,oo
E
3:
I-G.
I,i
13
2 8 " 0 5 ' N, 7 3 " 3 0 ' 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 m a x i m u m 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 m a x i m u m at ~ 100 m. This deep chlorophyll m a x i m u m 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 m a x i m u m 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 oceanwide 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. Chlorophyll 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 relationship 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.
Phaeocystispoucheti 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 Atlantic 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 c o m m o n 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 a m o u n t of osmoregulatory 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
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New Yo~'k~
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75
10
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6
DMS
o
E
c
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 D M S P and D M S 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
Synecchococcussp., 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 combination 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 numerous 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 KoblentzMishke 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 Smethie 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
( 106 km 2)
Mean concentration
(nmol 1- l )
Total flux
(Tmol S year- ~)
Oligotrophic (tropical/low
productivity )
Temperate
Upwelling (coastal
and equatorial)
Coastal/shelf
148
2.4
0.2-0.6
83
86
2.1
4.9
0.1-0.3
0.2-0.7
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 s u m m e r 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 n u m b e r 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 R a e m d o n c k (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 eddycorrelation 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 Bingemer (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 schematic description of the major atmospheric oxidation reactions of DMS. Currently 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 communication, 1989), which would make this reaction negligible compared with
the OH oxidation.
16
M.O.ANDREAE
03
02
OH• H20
NO3
HNO3
CH3S.
o.y
v
CH3SCH3
HO~.
CH3SO2"
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 a m o u n t of work has been done to determine the rate of the
reaction between DMS and these radicals; however, the actual reaction sequences and products are still uncertain. Observations on the relative abundances of SO2 and the other possible DMS oxidation products (DMSO,
methanesulfonic acid ( M S A ) ) 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 c o m p o u n d should be conducted. Uncertainty 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 seasurface 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 predominantly in the form of aerosol particles, even though in the case of MSA a
significant a m o u n t (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; Berresheim, 1987; Berresheim et al., 1990) or time-dependent numerical models
(Ferek et al., 1986 ), and assuming DMS fluxes of the order estimated in Table 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 (Australia). 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 1.01.2 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)
nss-S042-
MSA
DMS
MSA/nss- S04z-
4
NE PACIFIC
3
km
2
2o
~o '
o ~
,o,~
o
I00
150
0 0.04 0.08
Ii5O ,60
pptv
,~o
o
50
i
i
i ~
i
(b)
"r
~ I SOUTHERN OCEAN
krn
CO
20
pptv
40 J
0 5 I0 15
pptv
0
0.4
d.8'
tool/tool
Fig. 6. Verticaldistributions of DMS, MSA, non-seasaltsulfate (nss-SO2- ) and the MSA/nssSO42- ratio over (a) the northeasternPacificOcean (May 1985) and (b) the SouthernOcean
near Tasmania (December1986). (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-
GLOBALBIOGEOCHEMICALSULFURCYCLE
19
Radiation
budget_
Cloud condensation
nuclei
Global temperature
Sulfate aerosol
,+
~'~ ~---lv
Climate feedbacks
D~M:
-I-
+or-?
DMS
?. ~
+ or-.
Atmosphere
Ocean
Phytoplankton
?
abundance and ~ '
speciation
Marine
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 predicted 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 C O 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. Because 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 violent volcanic eruptions ). The input of COS is considered to be responsible for
the maintenance of the sulfate aerosol layer in the stratosphere during volcanically 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 ~ 0.031.0 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 solutions 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
GLOBALBIOGEOCHEMICALSULFURCYCLE
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 y e a r - ' ; Ferek and Andreae ( 1983): ~ 16 Gmol year-= ).
i
i
I
i
I
I
I
i
I
i
i
FLORIDA BAY - B A H A M A S
8-21 NOV 1983
I T
15
l
0
I-<t
n,,
m
z
0
I-<[
r.,'
l
NEARSHORE
l
I0
I.<[
(n
u~
o
u
J
. OPENOCEAN "',,
Oi
i
i
i
3 05 07 09
I'I
= '
I
i
15 15 17 19
I
I
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
( 106 km 2)
Mean
concentration
(pmol 1- l )
Oligotrophic (tropical/low
productivity )
Temperate
Upwelling (coastal and
equatorial)
Coastal/shelf
148
11.3
14.0
0.8
83
86
20.3
24.1
45.0
64.0
1.4
2.0
49
95.0
373.0
6.7
Mean: 27.6
Flux/area
(nmol m - 2 day- ~)
Total flux
(Gmol year- ~)
Total: 10.9
F O R M A T I O N A N D EMISSION O F H Y D R O G E N S U L F I D E A N D C A R B O N D I S U L F I D E
Hydrogen sulfide
There are few data on the concentration of dissolved H 2 S in surface seawater, and only a few reliable measurements of H2S in the marine atmosphere; therefore the air-sea exchange flux of this c o m p o u n d 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 workers 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 m a x i m u m 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. However, 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
( 106 km2)
Open oceans
Coastal/shelf
310
50
Mean
concentration
(pmol S year -~ )
16
33
Mean: 18
Flux/area
( nmol m - 2 day- ~)
45
90
Total flux
( Gmol S year- t )
5.1
1.6
Total: 6.7
We h a v e r e c e n t l y d e t e r m i n e d C S 2 in o p e n o c e a n a n d coastal s e a w a t e r f r o m
the N o r t h Atlantic, a n d h a v e o b s e r v e d m e a n c o n c e n t r a t i o n s o f 16__ 8 a n d
33_+ 19 p m o l S ( C S 2 ) 1- l , r e s p e c t i v e l y ( T a b l e 5; K i m a n d A n d r e a e , 1 9 8 7 ) ,
s o m e w h a t h i g h e r t h a n L o v e l o c k ' s results. F r o m these data, we e s t i m a t e a flux
o f ~ 7 G m o l S y e a r - i in the f o r m o f CS2 f r o m the W o r l d O c e a n surface, a b o u t
0.6% o f the D M S flux. T h e p h o t o c h e m i c a l o x i d a t i o n o f CS2 p r o d u c e s o n e
m o l e c u l e each o f SO2 a n d C O S per m o l e c u l e o f CS2 oxidized. T h u s , the m a rine e m i s s i o n o f CS2 p r o v i d e s a significant i n d i r e c t s o u r c e o f COS, w h e r e a s it
is clearly i n c o n s e q u e n t i a l as a s o u r c e o f t r o p o s p h e r i c SO2.
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