SCOPE 21 -The Major Biogeochemical Cycles and Their Interactions  

 ABSTRACT

The seaair exchange of aerosols plays an important role in the global biogeochemical cycles of carbon, phosphorus, nitrogen, and sulphur. In this paper a brief general discussion of sea-to-air and air-to-sea transport of aerosols is given followed by a discussion of the sources, transport, and seaair exchange of compounds of each of these elements in marine aerosols. In terms of the interaction of these cycles, the clearest case for marine aerosols is for N and S, with the likelihood that most of the NH₄⁺ and SO₄^2- on the smallest particles are present as a result of the different acid/base characteristics of NH₃, SO₂/SO₃ and their condensed phases. Phosphorus transport from sea to air is clearly related to organic carbon transport, and it is possible that a significant quantity of marine aerosol nitrogen may also be associated with organic carbon. Since there is considerable recycling of trace susbtances across the airsea interface, evaluation of the net (vs gross) input rate of C, P, N, and S compounds from the continents via the atmosphere to the ocean is quite difficult but is necessary to evaluate accurately the importance of airsea exchange in the biogeochemical cycles of these elements.

16.1 INTRODUCTION 

In terms of mass, the ocean is probably the major natural source of aerosols. It is also the ultimate sink, either directly, or indirectly via rivers and other inputs, for a significant fraction of continentally derived aerosols. Aerosols over the ocean, which we will refer to as marine aerosols, are thus not only derived from the ocean. Primary and secondary aerosols from the continents can be transported thousands of miles over ocean waters before being removed by wet and dry deposition processes. Sea salt aerosols produced by the ocean do not necessarily have the same relative chemical composition as the bulk sea-water from which they are formed. Thus airsea exchange of the marine aerosol involves input of continental material to the ocean and input of modified marine source material to the atmosphere. Since there is a considerable amount of recycling of trace substances across the airsea interface, evaluation of the net input of these substances from the continents via the atmosphere to the ocean is often extremely difficult.

The purpose of this paper will be to point out what we know, what we think we know, and what we almost certainly do not know about exchange of aerosols in both directions across the airsea interface. A discussion will then follow on carbon, sulphur, phosphorus, and nitrogen in marine aerosols, possible relationships among these biogeochemical cycles relative to marine aerosols, and the airsea exchange of these elements.

16.2 SEA-TO-AIR TRANSPORT OF AEROSOLS

It has been estimated that the sea produces between 1000 and 10,000 Tg/yr of atmospheric sea salt particles with radii less than ~ 20 µm (Eriksson, 1959, 1960; Blanchard, 1963; Petrenchuk, 1980). Most of the atmospheric sea salt particles in this size range are produced by whitecap bubbles breaking at the sea surface. When a bubble bursts at the seaair interface, atmospheric particles are produced both from a central jet and from the shattering bubble film cap. When the bubbles break, they skim off a thin layer of the sea surface to form the film and jet drops. MacIntrye (1968) has investigated this `microtome' effect for jet drops and has shown that the material present in the top jet drop was originally spread over the interior of the bubble surface (both bubble cap and the portion submerged) at a thickness equal to approximately 0.05 percent of the bubble diameter. Bubbles in breaking waves range in diameter from a few tens of micrometres to a few millimetres with most being between 100 µm and 1 mm. Blanchard (1963) and Cipriano and Blanchard (1981) found that the diameter of the jet drop was about 10 percent of the diameter of the bubble producing it. Thus a 100 µm bubble will produce a 10 µm atmospheric sea salt particle which is composed of material originally present in the top 0.05 µm of the bubble (or ocean) surface. Similarly the 100 µm jet drop from a 1 mm bubble is derived from the top 0.5 µm of the airsea interface. These relationships for jet drops are summarized in Figure 16.1.

As expected, the airsea interface has been found to be highly enriched in surface active organic materials and other substances associated with them (MacIntyre,1974; Liss, 1975; Duce and Hoffman, 1976). Thus it is no surprise that these materials can be considerably enriched on the jet drops relative to their concentration in bulk sea-water a few cm below the surface. This `chemical fractionation' process is very difficult to evaluate by collecting and analysing aerosols over the ocean surface because the ocean is not the only source for particles found there. However, careful laboratory studies and carefully controlled studies of artificial bubbling and aerosol sampling over the ocean have shown clearly that such substances as iodine, phosphorus, probably organic nitrogen, many organic carbon compounds, and certain heavy metals are enriched on the sea salt particles produced by bursting bubbles. In many cases, however, this enrichment in the ambient marine atmosphere is swamped by the high concentration of these substances already present in the atmosphere from other sources.

Figure 16.1 Relationships among bubble diameter, jet drop diameter, and microtome depth (After MacIntyre, 1974). Reproduced by permission of John Wiley & Sons, Inc.

We know considerably less about film drops than jet drops. Considering the size of most of the bubbles in breaking waves, it is clear that relatively few jet drops less than 510 µm diameter will be produced, yet there is a considerable number (and mass) of sea salt particles smaller than this size in the marine atmosphere. Cipriano and Blanchard (1981) have shown that most of the smaller sea salt particles are film drops, and that these are derived primarily from the very large bubbles, perhaps 1 mm and larger, which may produce as many as 100 or more film drops per bubble. Submicrometre size film drops are commonly produced. In general, the results of Cipriano and Blanchard (1981) suggest that most of the atmospheric sea salt particles smaller than 510 µm diameter originate as film drops from bubbles larger than 1 mm, while most of the sea salt particles in the atmosphere larger than ~ 20 µm originate as jet drops derived from bubbles larger than ~ 200 µm (see Figure 16.2).

Figure 16.2 Relative jet and film drop contributions to atmospheric sea salt particles. This is a first approximation only and should not be used in an exact sense (After Cipriano and Blanchard, 1981). Reproduced by permission of American Geophysical Union

Very little is known about chemical fractionation on film drops, although it almost certainly occurs. Clearly, if the data of Cipriano and Blanchard (1981) are correct, most atmospheric sea salt particles with any significant lifetime in the atmosphere (i.e., with r20 µm) are derived from film drops. Considerable research is needed on the physics and chemistry of film drop production to understand properly sea-to-air chemical exchange of aerosols.

The size of atmospheric sea salt particles is directly related to the ambient relative humidity. For example an atmospheric sea salt particle with a radius of 10 µm when ejected from the ocean surface would have a radius of 4 µm when in equilibrium with an ambient relative humidity of 80 percent and only 2 µm when in equilibrium with a relative humidity of 70 percent, the relative humidity at which solid precipitates may begin to form in the droplet. Note that even at relative himidities of 50 percent or lower, the hygroscopic nature of salt particles assures at least a film of water on the particles. Thus at the relative humidities of 75 to 85 percent often found in the lower marine boundary layer, solution chemistry and gasliquid interaction processes apply relative to ocean derived aerosols. This must be kept in mind when considering the atmospheric chemistry of marine aerosols.

16.3 AIR-TO-SEA TRANSPORT OF AEROSOLS 

Marine aerosols are transported to the sea via wet and dry removal processes. Estimates of the wet (rain) removal of aerosols, or material present on aerosols, are often made utilizing the wash-out factor or scavenging ratio, W.

		CRD
	W =
		CA

and C_R is the concentration of material in rain (e.g., µg kg^-1), C_A is the concentration of material in air (e.g., µg m^-3), D is the density of air (e.g., ~ 1.2 kg m^-3 at 20°C, 1 atm), and W is dimensionless. Values for W generally range from a few hundred to a few thousand, which roughly means that one gram (or one cm³) of rain scavenges about one m³ of air. W is, of course, dependent upon a number of factors, including particle size, chemical composition, vertical concentration distribution of the aerosol, vertical extent of the precipitating cloud, etc. Gatz (1977) and Duce et al. (1979) have shown that W decreases with decreasing particle size for certain trace metals at urban and continental coastal sites. Recent studies by Ng and Patterson (1981) indicate that W for Pb, a small particle element, is greater than that for continental dust at a mid-Pacific Ocean site. Thus effects of size, particle chemistry, and other factors make it difficult to predict W more accurately than in the range between a few hundred and a few thousand.

v_d = F/M

and M is the mass of aerosol (or material on aerosol) (e.g., µg cm^-3), F is the flux of particles to the surface (e.g., µg cm^-2 s^-1), and v_d is the deposition velocity (e.g., cm s^-1).

v_d can be estimated from theoretical considerations or it can be measured in laboratory or field situations. Considerable controversy exists about the validity of dry deposition velocity values determined from measurements to buckets or dry plates (see for example Sehmel, 1980). Effective deposition velocities for particles in the stable aerosol size range near ground are often found to be near 1 cm s^-1, but this varies considerably with particle size, wind speed, and surface roughness.

In a laboratory wind tunnel experiment Sehmel and Sutter (1974) investigated aerosol deposition velocity over a water surface as a function of wind speed and particle size. Their results are presented in Figure 16.3. Slinn and Slinn (1980) developed a theoretical two-layer model for the prediction of particle dry deposition to natural waters. They pointed out that if particle growth by water vapour condensation occurs in the humid regions near an air-water interface, the deposition velocity of certain types of particles (they used (NH₄)₂SO₄ as an example) with dry radii near 1 µm is nearly independent of particle size and can be approximated by v_d = 1.3 x 10^-3 , where is the mean wind speed in cm s^-1. At winds of 510 ms^-1, this results in an aerosol v_d of ~ 1 cm s^-1 (see Figure 16.4).

Figure 16.3 Atmospheric particle dry deposition velocity, v_d, to a water surface as a function of particle size and wind speed from wind tunnel studies (After Sehmel and Sutter, 1974). Reproduced by permission of J. Rech. Atmos.

Aside from the size range ~ 1 µm described above, both Sehmel and Sutter (1974) and Slinn and Slinn (1980) predict a large difference in v_d for very small (e.g., 0.1 µm) vs very large (e.g., 20 µm) particles. For example, for a 20 µm and 0.1 µm radius particle at a wind speed of 10 m s^-1, Sehmel and Sutter (1974) predict 40 cm s^-1 and 0.01 cm s^-1 respectively and Slinn and Slinn (1980) predict 10 cm^-1 and 0.01 cm s^-1 respectively. Thus, if even a very small percentage of a substance is present over the ocean on the larger particles, these larger particles may dominate the dry deposition to the ocean.

As an example, recent studies of lead in the atmosphere at Enewetak Atoll in the mid-Pacific Ocean and its deposition to the ocean surface during the SEAREX (Sea/Air Exchange) Program have shown that the mass median radius of the Pb containing particles is ~ 0.2 µm) (R. Duce, unpublished data). However, the measured dry deposition velocity of the Pb to a flat plate several metres above the ocean surface was ~ 0.7 cm s^-1 (C. C. Patterson, unpublished data), much greater than that expected from the mass median radius. However, if the measured Pb particle size distribution was applied to the theoretical deposition velocity relationships derived by Slinn and Slinn (1980), the predicted total Pb dry deposition agreed well with that observed. The small percentage of Pb on the largest particles was controlling the Pb deposition. Thus the size distribution of any substance of interest on aerosols and the deposition velocity as a function of particle size must both be known to adequately approximate the dry deposition of any substance to the ocean.

Figure 16.4 Theoretically derived particle dry deposition velocity, v_d, to a water surface as a function of particle size and wind speed (After Slinn and Slinn, 1980). Reproduced by permission of Pergamon Press Ltd.

Perhaps the most serious problem in interpreting wet and dry deposition calculations for any substance entering the ocean is evaluating gross vs net deposition. In evaluating the airsea exchange portion of biogeochemical cycles, we are most often interested in the net input of material into (or in some cases out of) the ocean. Measurements of rain concentrations and measurements or calculations of dry deposition from concentration and dry deposition velocities only give us gross input values. From this we must remove the recycled component coming from the ocean itself. If atmospheric sea salt particles had the same chemical composition as bulk sea-water, this would be simple, as we could simply use the X:Na ratio in the atmospheric sample and in the ocean and factor out the marine component. If, however, the marine component does not have the composition of sea-water, this approach will not workit will over-estimate the net flux into the sea. There are several approaches that give some promise of being able to evaluate this problem, and these will be discussed in the following sections on C, S, P, and N exchange across the seaair interface.

16.4 CARBON

Carbon can be present on marine aerosols in essentially three forms-carbonate-bicarbonate, organic, and elemental carbon or soot. Except in areas downwind from continental or island limestone or coral deposits, carbonate is a small fraction of marine aerosol carbon and will not be considered further here.

Hundreds of different organic compounds are present on atmospheric particles and each has its own characteristic physical and chemical properties and associated atmospheric sources, residence times, and sinks. Data are very sparse on individual compounds in the marine aerosol, however, so we will restrict our discussion to total organic carbon.

Organic carbon in the marine aerosol has been measured by several investigators in the past few years. In many cases the concentrations reported for organic carbon probably include elemental carbon as well. Reported values in the near surface marine atmosphere (1090 metres) are presented in Table 16.1. Note that the mean concentrations observed are very similar at the various marine sites in the northern and southern hemispheres, with a range of mean concentrations of only 0.2 to 0.9 µg m^-3. The overall mean concentration at all eleven sites is 0.48 ± 0.20 µg m^-3. Thus, a concentration of about 0.5 µg m^-3 is representative of the organic (or total) carbon content of the marine aerosol over much of the world ocean. Hoffman and Duce (1977) found that about 80 percent of the carbon was on particles with radii < 0.5 µm at Bermuda, Hawaii, and Samoa sites (see Figure 16.5). They suggested this was the result of gasparticle conversion reactions in the atmosphere. The distribution of organic carbon and Na as a function of particle size for laboratory generated sea salt particles and for a typical atmospheric sample from Bermuda is represented in Figure 16.6. While the Na size distribution is quite similar in the ambient and laboratory aerosol, the organic carbon distribution is quite different. The organic carbon in the laboratory generated sea salt particles is enriched several hundred-fold over the sea-water concentration (relative to Na) and is present on generally the same size particles as the Na. This suggests that the large-particle organic carbon is probably present on the sea salt aerosols when they are produced by the ocean, while the small particle, or dominant, organic carbon comes from some other source.

Table 16.1 Organic carbon concentration in the marine aerosol

Location	Mean		Range	Reference
	(µg Cm-3 STP)		(µg Cm-3 STP)
Northern Hemisphere
Bermuda	0.29 ±0.09		0.150.47	Hoffman and Duce (1974)
North Atlantic	0.76 ±0.42		0.331.6	Ketseridis et al. (1976)
West Ireland	0.57 ±0.29		0.200.86	Eichmann et al. (1979)
Bermuda	0.37 ±0.23		0.150.78	Hoffman and Duce (1977)
Sargasso Sea	0.44 ±0.04		0.380.48	Chesselet et al. (1981)
Hawaii	0.39 ±0.03		0.360.43	Hoffman and Duce (1977)
Enewetak Atoll	0.89 ±0.27		0.731.2	Chesselet et al. (1981)
Eastern Tropical Pacific	0.49 ±0.26		0.220.74	Hidy et al. (1974)
Southern Hemisphere
Samoa	0.22 ±0.09		0.130.41	Hoffman and Duce (1977)
Eastern Tropical Pacific	0.21 ±0.18		0.070.53	Barger and Garrett (1976)
Tasmania	0.53			Eichmann et al. (1980)

Figure 16.5 Organic carbon concentration as a function of marine aerosol particle size (After Hoffman and Duce, 1977). Each symbol indicates one sample. Reproduced by permission of American Geophysical Union

Recent SEAREX studies by Chesselet et al. (1981) utilizing stable carbon isotopes lend support to the suggestion that the small particle organic carbon is not of marine origin. Figure 16.7 shows the Na and organic carbon concentration and the ¹³C value for a size-separated aerosol sample collected at Enewetak Atoll. ¹³C values for the smallest particle are 26‰ to -28‰. Chesselet et al. (1981) point out that this range is similar to ¹³C values for continental vegetation, coal, and the products of petroleum combustion, 26 ± 2 ‰, suggesting that the small-particle carbon is of continental origin. The ¹³C values for the large-particle carbon are 18 to -22‰. This is similar to 8¹³C of marine organic carbon, which is generally 21 ± 2‰ in low latitude regions (40°S50°N), suggesting that the large-particle organic carbon in the atmosphere is of marine origin.

Figure 16.6 Comparison of organic carbon and sodium concentration as a function of particle size on ambient marine aerosols and laboratory generated sea salt particles (After Hoffman and Duce, 1977). Reproduced by permission of American Geophysical Union

Figure 16.7 Sodium, organic carbon, and ¹³C as a function of marine aerosol particle size. ¹³C is defined in the usual notation:

(From data reported in Chesselet et al., 1981). Arrows indicate less than or equal to.

The global production of aerosol organic carbon is rather poorly known. Duce (1978) estimated a global source strength of organic carbon on primary aerosols of ~ 56 Tg yr^-1, of which about half was from natural sources (the ocean, crustal weathering, forest fires) and about half from pollution sources. Duce (1978) estimates that about 80 to 160 Tg yr^-1 of particulate organic carbon results from gas to particle conversion processes, resulting in a total global production rate of ~ 140 220 Tg yr^-1. This agrees well with the estimate of Jaenicke (1978) of ~ 200 Tg yr^-1, ~ 150 Tg yr^-1 from gas-particle conversion and 50 Tg yr^-1 from direct production.

A complication here is elemental carbon. Elemental carbon has been found in marine sediments in the mid-Atlantic and Pacific Oceans (~ 0.02 0.1 percent dry weight; Smith et al., 1973; Griffin and Goldberg, 1975). This material originates from the combustion of coal, wood, and petroleum products (Griffin and Goldberg, 1979). It has been measured extensively in urban areas, but few data are available in remote regions. Rosen et al. (1981) find elemental carbon concentrations ranging from ~ 0.1 to 0.9 µg m^-3 at Barrow, Alaska, with the higher concentration being observed in the winter months. The high concentrations coincide with apparent long range transport of pollution across the Arctic during the winter (Rahn and McCaffrey, 1980) and indeed are near typical concentrations in many urban areas (Rosen et al., 1981). No data for elemental carbon are available from remote marine regions, but the concentrations are probably considerably less than at Barrow in the winter since filters are typically still white in mid-Pacific locations even after the passage of 10⁴ m³ of air.

The global production of elemental carbon by fires may be very high. Ryan and McMahon (1976) indicate that as much as 40 percent of the particles produced during burning of temperate forests and agricultural wastes may be elemental carbon. Seiler and Crutzen (1980) estimated that as much as 90180 Tg yr^-1 of aerosol carbon may be produced in this way. Note this is in the range of the total production of aerosol organic carbon estimated by Duce (1978) and Jaenicke (1978). During discussions at this SCOPE meeting it was suggested that perhaps 10 percent of the 90180 Tg yr^-1 of aerosol carbon may be emitted as fine particles that can be transported great distances. Elemental C may thus play a very important role in the global aerosol carbon budget. The historical record of burning is maintained in coastal marine and in lake sediments. As emphasized by UNESCO (1980), the fate of this elemental carbon may be important for the global carbon budget, and its production and distribution should be investigated in more detail.

Information on the net input of organic carbon to the ocean from the atmosphere is essentially nil. We have virtually no data on organic carbon in marine rains. Neumann et al. (1959) reported 1.73.4 mg kg^-1 organic carbon over Sweden, while Williams (1967) reported 0.7 mg kg^-1 just north of Samoa. Gagosian (1981) found a mean organic carbon concentration of 0.64 ± 0.48 mg kg^-1 in 10 samples collected during the SEAREX Program at Enewetak Atoll. If we assume a global value of 0.51 mg kg^-1 organic carbon (i.e., a W of 1000 to 2000) and a rainfall over the ocean of 3.9 x10¹⁷ kg yr^-1 (Baumgartner and Reichel, 1975), there is a total annual input of ~ 200 400 Tg yr^-1. This is very large, about 24 times the estimates of the riverine input of organic carbon to the ocean (Duce and Duursma, 1977) and higher than the estimates of global organic carbon aerosol production described above (which may well indicate that the production estimates are too low). This calculated input is, of course, a gross input to the ocean and is likely to be considerably higher than the net input because much of it is probably recycled carbon from the ocean surface, present on the largest and most scavengeable sea salt particles.

We have a similar problem distinguishing net from gross dry deposition, as it is very possible that the gross deposition will be dominated by the relatively small concentration of organic carbon on the large particles, which have high dry deposition velocities. For example, if we assume we have 0.4 µg m^-3 of organic carbon on particles with radii ~ 0.20.3 µm with a v_d of ~ 0.05 to 0.5 cm s^-1 (see Figure 16.4), the global dry deposition of this fraction would be ~ 2 to 20 Tg yr^-1. If we assume the other 0.1 µg m^-3 of organic carbon is on 15 µm particles with a v_d of 13 cm s^-1, this size fraction would contribute 1030 Tg yr^-1 to the ocean. There are no field data on dry deposition of organic carbon to the ocean at this time.

The use of carbon isotopes offers an excellent opportunity for untangling the relative role of small (net input) particles and large (recycled material) particles and may allow us to evaluate the net input of organic carbon to the ocean. Experiments to do this are planned by R. Chesselet and his group at the SEAREX Program study in Samoa during the summer of 1981. This group will measure organic carbon concentrations and ¹³C values in rain, in dry deposition, and in the ambient aerosol as a function of particle size. Results of the analyses of these samples should indicate clearly which size particles are primarily involved in the wet and dry deposition processes and thus what the net input of organic carbon is to the ocean.

16.5 SULPHUR

Sulphate in the marine aerosol is generally present in concentrations higher than expected for particles generated from sea-water, i.e., the SO₄^2-:Na ratio in marine aerosols is almost always higher than the value of 0.25 found in sea-water (Buat-Menard et al., 1974; Gravenhorst, 1978; Bonsang et al., 1980). Several investigators in the 1950s and 60s suggested this `excess' sulphate was the result of chemical fractionation processes during bubble bursting and sea salt particle production, but it has generally been accepted that the excess sulphate has a different source. In a recent paper, however, Garland (1981) found small enrichments (1030 percent) of sulphate during bubble bursting using radiotracers in a laboratory study. Thus this process cannot be ruled out completely as a potential source for some of the excess sulphate observed.

Several studies have shown that the sulphate size distribution in the marine aerosol is bimodal, with one fraction present on sea salt size particles and the other on submicrometre size particles. Gravenhorst (1978) found that the mass median radius of the fine-particle sulphate on marine aerosols over the North Atlantic was ~ 0.4 µm while that for sea salt sulphate was ~ 2.5 µm. Lawson and Winchester (1978, 1979) found similar results in the southern hemisphere. In a review of recent studies of SO₄^2- over the Atlantic, Georgii (1978) points out that the contribution of non-sea salt sulphate to total sulphate in marine aerosols may be 60 to 80 percent, with as much as 1 µg m^-3 of excess sulphate over the North Atlantic. Huebert and Lazrus (1980b) found many of their samples over the North and South Pacific had SO₄^2-:Cl ratios similar to sea-water, suggesting the excess SO₄^2- was rather small in this area. Bonsang et al. (1980) have shown that the SO₄^2-:Na ratio on particles with radii > 1 µm is the same as in sea-water, and these larger sulphate particles are clearly derived from the ocean, similar to organic carbon, but in this case with no chemical fractionation. Thus it is possible that `excess sulphate' and 'fine-particle sulphate' are essentially synonymous in the marine aerosol. Some recent data on fine-particle or excess sulphate are presented in Table 16.2. Note that excess SO₄^2- is generally 0.5 to 1 µg m^-3 over the North Atlantic and apparently somewhat lower over the North Pacific. It is considerably lower in the southern hemisphere, generally 0.050.1 µg m^-3. The data of Bonsang et al. (1980) in the southern Indian Ocean are an exception to this. These data were obtained from 20°S. to 60°S., and there is no explanation at this time for the high excess sulphate concentration observed in that region. Lawson and Winchester (1979) point out that fine-particle sulphate in marine air at Punta Arenas, Chile is much lower in the winter (0.01 µg m^-3) than the summer (0.05 µg m^-3), suggesting possible seasonal (and perhaps biological) effects. Meszaros (1978) suggested that there may be a SO₄^2-maximum over the North Atlantic at ~ 40^°N. and that this may be a result of conversion of high concentrations of SO₄^2-_, from pollution origin in the westerlies.

Table 16.2 Fine particle or excess sulphate in marine aerosols-some recent data

Location		Mean	Reference
		(µg m-3 STP)
Northern Hemisphere
Central East Pacific,
0°10°N.		~0.12	Maenhaut et al. (1981)
North Atlantic		0.81 ± 0.49	Darzi (1977)
Bermuda		~0.53	Meinert and Winchester (1977)
North Atlantic		0.9 ± 0.5	Gravenhorst (1978)
North Atlantic, 15°N.		0.7 ± 0.2	Bonsang et al. (1980)
Faroe Islands		0.41 ± 0.28	Prahm et al. (1976)
Southern Hemisphere
Central East Pacific,
0°15°S.		~0.05	Maenhaut et al. (1981)
Indian Ocean		0.050.1	Darzi and Winchester (1981)
Samoa		~0.06	Maenhaut et al. (1979)
Salvador, Brazil		~0.10	Lawson and Winchester (1979)
	Punta Arenas, Chile	~0.05	Lawson and Winchester (1979)
Southern Indian Ocean		~ 2	Bonsang et al. (1980)

The source of the small-particle SO₂ is a matter of considerable interest. That much if not most of it is derived from gasparticle conversion involving SO₂, is almost certain, but the source of the SO₂, is not. The sulphur cycle, and gasparticle sulphur exchange, has been discussed at great length by others and will not be developed here. Bonsang et al. (1980) contend that in marine regions far from continents the SO_2, is not continentally derived but is probably from the oxidation of marine derived reduced forms of sulphur such as dimethylsulphide. Maroulis et al. (1980) suggest that much of the background SO₂, results from oxidation of COS. The higher concentrations of fine-particle SO₄^2- over the North Atlantic support a continental, and probably anthropogenic, source for the SO₂, that is converted to SO₄^2- in this region.

Taylor et al. (Chapter 4, this volume) have also discussed the bimodal size distribution of aerosol sulphate and their conclusions are similar to those reached in this paper. Of particular interest in the paper by Taylor et al. (Chapter 4, this volume) is their discussion of the relationships among NH₄⁺ , SO₄^2-, and rain-water acidity. Even in remote marine areas there do not appear to be sufficient basic substances to neutralize the fine-particle H₂SO₄; the only basic material evaluated to date is NH₃. No data are available on the presence, for example, of basic amines or other basic gases in the marine atmosphere.

In terms of seaair exchange of SO₄^2- consideration of the primary sea salt sulphate would appear to be simple and straightforward. That may not be the case, however. For example, Lawson and Winchester (1979) point out that in the southern hemisphere marine atmosphere, the highest fine particle SO₄^2- concentrations were observed along the coast during the most intense periods of large-particle sea salt sulphate. This suggests there may be some direct contribution to fine-particle SO₄^2- from the ocean, and agrees with Cipriano and Blanchard's (1981) recent findings that the ocean can apparently produce reasonably large concentrations of sub-micrometre size aerosols.

With no intention of summarizing the literature, it is interesting to comment briefly on sulphate in marine rains. Numerous investigators have found excess sulphate in marine rains (e.g., Eriksson, 1957; Junge, 1963; Nyberg, 1977; Asman and Ridder, 1981, among others). Perhaps the most geographically extensive set of data to date were reported by Kroopnick (1977). Rain from 19 stations in the North and South Atlantic and Pacific Oceans was analysed for SO₄^2- and Cl^-. The mean SO₄^2-:Cl^- ratio observed for precipitation at each sea level marine site at which eight or more samples were collected is presented in Table 16.3. Data are also given for Mauna Kea, Hawaii, 3500 m above sea level. The sea-water SO₄^2-:Cl^- ratio is 0.14 and the mean rain-water SO₄^2-:Cl ratio at all the sites (except Mauna Kea) is 0.20 ± 0.09. While excess SO₄^2- is certainly observed, the data from this study suggest it is not observed at all sites all of the time. Note also that there is independent evidence for CI loss from marine aerosols, probably partly as HCl due to acidification of the marine aerosols by, among other things, H₂SO₄ formation. Thus Cl itself is not always a conservative indicator of the sea salt contribution for any substance.

Table 16.3 The SO₄^2-:Cl ratio in marine rains (From data in Kroopnick, 1977). Reproduced by permission of University of Hawaii Press

Excess sulphate in rain should be a maximum in areas where the excess sulphate in the marine aerosol is highest, e.g., over the North Atlantic. This is not apparent in the data of Table 16.3. However, Asman and Ridder (1981) found a mean excess SO₄^2- concentration of 2.6 mg litre^-1 for a series of samples collected over the North Atlantic from a Dutch weathership at 60°N., 2°E. In one low-pH sample (3.8), the excess sulphate was 16 mg litre^-1. Nyberg (1977) found a mean excess SO₄^2- concentration of 1.2 mg litre^-1 for 26 rain samples collected from ships in the North Atlantic from 45° to 65°N. and 33°W. to 4°E. In contrast, Pszenny and Duce (1981) collected 33 rain samples under onshore wind conditions from a tower about 50 metres above sea level on the windward shore of American Samoa in the South Pacific. They found `a hint of an excess SO₄^2- component' of ~ 0.04 mg litre^-1. With total SO₄^2- concentrations averaging 1.3 mg litre^-1, this was essentially a statistically insignificant enrichment. Thus, there is some evidence that excess sulphate in rain is considerably higher in some marine regions that others, as expected. A number of detailed studies of rain chemistry over the world ocean are presently underway, and considerably more data should be available soon.

Interestingly in eleven global cycles of atmospheric sulphur published since 1963, all use the same value for the sea spray sulphate-sulphur source strength, 44 Tg yr^-1. This number is derived from the excellent pioneering papers of Eriksson (1959, 1960) in which he derived the global sea salt flux (~1200 Tg yr^-1). A fundamental assumption in the derivation of the 44 Tg yr^-1 figure was that the amount of sea salt SO₄^2- deposited on continents was 10 percent of that redeposited directly in the ocean from the atmosphere. The 10 percent figure was derived from the total chloride entering the ocean each year via river run-off (it was all assumed to be of sea salt origin deposited on the continents via rain and dry deposition) and Eriksson's calculation of the annual deposition of atmospheric sea salt (and thus chloride) directly to the ocean. Using the Cl^-:SO₄^2- ratio in sea-water and the riverine chloride input to the sea, it was calcualted that 4 Tg yr^-1 of sea salt SO₄^2--S entered the ocean via rivers. This results, then, in a global production of 44 Tg yr^-1 of atmospheric sea salt SO₄^2-S. Clearly the 44 Tg yr^-1 is directly related to the assumption that 10 percent of the atmospheric sea salt is deposited on land. If the percentage is less than 10 percent, the 44 Tg yr^-1 figure is too low. Similarly, if the 1200 Tg yr^-1 figure for global atmospheric sea salt production is too low, the estimate of 44 Tg yr^-1 SO₄^2-S from sea salt is also too low.

Blanchard (1963) re-evaluated the global sea salt production figures using the general approach of Eriksson, but utilizing actual measurements of the temporal and geographical variation of wind speeds over 5° to 20° latitude/longitude squares over the global coean (Eriksson assumed a 12 kt wind over the entire ocean). Blanchard estimated a global sea salt production of 10,000 Tg yr^-1, about 8 times that of Eriksson. Petrenchuk (1980) calculated a value of 1300 Tg yr^-1 salt removal by rainfall over the ocean each year. Relatively little consideration was given to dry deposition of sea salt, which may be equal to or greater than rain removal for sea salt. Thus it is quite possible that the commonly used estimate of 44 Tg yr^-1 SO₄^2--S introduced into the atmosphere on sea salt particles may be low by a factor of 2 to 8.

16.6 PHOSPHORUS

Much less is known about phosphorus in the atmosphere than the other three elements we are considering. In his discussion of the global phosphorus cycle, Pierrou (1976) devotes one paragraph to atmospheric phosphorus, and concludes `Unfortunately no measurements or estimates seem to have been published on this subject.' There are at least a few more data available now as a result of several recent papers on the global atmospheric phosphorus cycle (Graham and Duce, 1979) and on the sea as a source for atmospheric P (Graham et al., 1979).

There is no information on the presence of any significant concentration of vapour phase phosphorus compounds in the atmosphere, although certainly some do exist, for example certain pesticides. It is assumed, however, that the atmospheric phosphorus cycle is largely controlled by aerosol P. Graham and Duce (1979) calculated a global atmospheric P cycle, which is shown in Figure 16.8. The primary sources of atmospheric P are crustal weathering, the ocean, and pollution. The major pollution sources are from phosphate manufacturing and processing, from the soil as a result of man's activities, and from coal combustion. A detailed breakdown of sources, sinks, and the entire cycle is given in the original paper.

Phosphorus concentrations in the marine aerosol have been measured by Graham and Duce (1979) and by Graham et al. (1979). These data are summarized in Table 16.4. The effect of the continents, particularly the Sahara dust plume, is clearly evident in the data. Concentrations over the mid-Pacific appear to be ~ 0.5 ng m^-3. Concentrations north of'- 45°N. over the Atlantic are similar to this, but concentrations closer to North and South America are typically 5-10 ng m^-3, with 50 ng m^-3 common in the Sahara dust plume. Graham and Duce (1981) found the major mass of total P in the marine aerosol to be on particles with radii of 1 3µm, with perhaps 1020 percent of the mass on submicrometre particles. The calculated gross input of atmospheric P to the ocean is ~ 1.4 Tg yr^-1 with a net input of ~ 1.1 Tg yr^-1. The global annual mobilization of aerosol phosphorus is ~6 Tg yr^-1 (see Figure 16.8).

Figure 16.8 The atmospheric phosphorus cycle. The numbers in parentheses are the estimated inputs of sea-water soluble phosphorus through rivers and through the atmosphere (After Graham and Duce, 1979). Reproduced by permission of Pergamon Press Ltd.

Graham and Duce (1982) have considered the relative importance of the atmospheric transport of phosphorus to the phosphorus content of the waters in the western North Atlantic. They found that about 35 ± 15 percent of the total phosphorus present in marine aerosols over the western North Atlantic was soluble in sea-water within 12 hours. A somewhat lower percentage was found for Sahara dust aerosols. Since the total phosphorus input into an area of the ocean bounded by the North American coast, 25°N., and 65°W. was calculated to be 0.5 to 1.0 x 10¹⁰g yr^-1, at least 0.2 to 0.4 x 10¹⁰ g yr^-1 of phosphorus available for rapid nutrient use was deposited. This compares with ~ 3 x 10¹⁰ g P yr^-1 which is estimated to enter estuarine areas through east coast North American rivers, i.e., atmospheric input is at least 10 percent of riverine input to this region. However, the riverine input is added to estuaries and coastal waters, and it is uncertain how much of it may be deposited there. Atmospheric deposition occurs over a wide area, including much of the nutrient-poor Sargasso Sea. Thus the atmosphere may be a more important source of nutrient P to open ocean regions than the gross figure would indicate.

Table 16.4 Phosphorus concentrations in the marine aerosol (From data in Graham and Duce, 1979)

Location	Concentration
	Range
	(ng P m-3 STP)
Western North Atlantic (30°45 °N.)	1.0 15
Bermuda	0.6 9.2
Near Shore North America Atlantic	15 60
North Atlantic trades	3 80
Central North Atlantic (45°60 °N.)	0.6 1.0
West coast of South America	6 20
Hawaii	0.4 0.9
Samoa	0.4 0.8
Central Pacific (20 °N15 °S.)	0.04 0.9

Sutcliffe et al. (1963), Bruyevich and Kulik (1967), and MacIntyre and Winchester (1969) all found that in laboratory bubbling studies, phosphate was significantly enriched on sea salt aerosols produced by bursting bubbles compared with sea-water. Using a bubble generator in coastal sea-water, Graham et al. (1979) found that total phosphorus enrichment ranged from 4 to 170, i.e., P:Na on the sea salt aerosols was 4 to 170 times that of the sea-water. In their studies, Graham et al. (1979) measured total phosphorus (using persulphate oxidation), and reactive phosphorus (that soluble in distilled water, primarily PO₄^3- , HPO₄^2-, etc.). The difference between these two was termed `organic' phosphorus. It was found that the enrichment of reactive phosphorus was only a factor of 28 above sea-water whereas that of organic phosphorus was often 100200, suggesting most of the enriched phosphorus is associated with surface-active organic substances. Factor analysis and multi-variate regression analysis of sodium, aluminium (an indicator of crustal sources), excess vanadium (an indicator of anthropogenic sources), reactive P, and organic P in a number of samples collected over the eastern equatorial Pacific strongly suggested that crustal weathering was the primary source for the reactive P, the ocean was the primary source for organic P, and pollution apparently represented a minor source for both forms in this area. The mean atmospheric phosphorus concentrations observed in this region were 4.1 ± 1.2 ng m^-3 for reactive P and 2.8 ± 1.6 ng m^-3 for organic P, i.e., ~ 60 percent reactive P. In Hawaii and Samoa the reactive P, the form clearly from continental areas, was ~ 35 percent and 20 percent respectively, reflecting the increasing distances from continental regions (Graham and Duce, 1981). On the basis of these studies it would appear that organic P over the ocean is largely recycled from the ocean whereas reactive P is from continental areas. The input of reactive P to the ocean is thus of primary interest in the P biogeochemical cycle. More detailed measurements of the wet and dry input to the ocean of atmospheric reactive phosphorus must be made before an accurate assessment can be made of the importance of atmospheric P transport to the marine chemistry of surface waters. At present it would appear from Figure 16.8 that perhaps ~10 percent of the dissolved form of phosphorus entering the oceans each year is via atmospheric transport.

16.7 NITROGEN

Nitrogen in marine aerosols is somewhat similar to sulphur, i.e., gas-particle conversion plays a significant, in this case major, role in determining the NH₄⁺ and NO₃ content of marine aerosols. Evidence suggests that the ocean plays a relatively minor role in the direct production of marine aerosol NH₄⁺ and NO₃, although data from mid-ocean regions are sparse.

In the case of NO₃, much of the early aerosol NO₃ data are suspect due to artifact NO₃ formation from NO₂ on the filters utilized to collect particulate NO₃ (Appel et al., 1979). However the early data of Junge (1957) from Hawaii appear to show little evidence of this problem. Huebert and Lazrus (1980a) showed there was no correlation between aerosol NO₃ and Cl collected from an aircraft in the marine boundary layer over the North and South Pacific, indicating that sea salt aerosols were not a significant source for NO₃. Similarly Huebert (1980) found no correlation between aersosol NO₃ concentration and sea state in the central Pacific. Pszenny and Duce (1981) found no correlation between NO₃ and Na⁺ in marine rains collected in American Samoa (see Figure 16.9). Huebert and Lazrus (1980a) point out that, assuming there was no chemical fractionation of sea-water NO₃ during sea salt particle production, sea spray alone could be expected to contribute no more than 0.10.2 ng m^-3 of particulate NO₃, whereas measured marine aerosol NO₃ concentrations are typically 0.1 to 0.3 µg m^-3 (see Table 16.5). NO₃ showed no significant correlation with HNO₃, SO₄²-, NH₄⁺ or Rn in the studies of Huebert and Lazrus (1980b).

Figure 16.9 Nitrate versus sodium concentration in rain samples collected in Samoa (After Pszenny and Duce, 1981). Arrows indicate less than or equal to.

On the other hand Sequeira (1981) found a significant correlation (r = 0.81, 68 data pairs) between NH₄⁺ and NO₃ in precipitation collected at Valentia on the Irish coast. He also found a stoichiometric ratio for NH₄⁺ : NO₃ of near unity and suggested NH₄⁺ and NO₃ ions are tied up as NH₄⁺NO₃. Other evidence does not support this, however. Parungo et al. (1981) and Gravenhorst et al. (1979, 1981) found that most of the aerosol NO₃ was concentrated in the sea salt size range (110 µm). Gravenhorst et al. (1979, 1981) found most of the NH₄⁺ on particles with radii < 0.5 µm. the same size which contains the excess sulphate. They suggest that this indicates that conversion of gaseous NH₃ and HNO₃ over the ocean does not result in NH₄NO₃.

Table 16.5 Nitrate concentrations in marine aerosols (µg m^-3 STP)

Location	Range	Median	Mean	Reference
Cornwall, England		Minimum of 0.21		Brice et al. (1981)
(NO3 +HNO3)
North Atlantic			~0.06	Gravenhorst (1978)
North Pacific	<0.09 1.4	0.25	0.38	Huebert and Lazrus (1980a)
Hawaii			0.09	Junge (1957)
Pacific, 7°N.9°S.	0.080.3	0.20	0.21	Huebert (1980)
South Pacific	< 0.10 0.5	0.12	0.16	Huebert and Lazrus (1980a)
Table 16.6 Ammonium concentrations in marine aerosols (µg m-3 STP)
Location	Range	Median	Mean	Reference
North Atlantic	0.010.15			Gravenhorst et al. (1979)
North Atlantic			~ 0.06	Gravenhorst (1978)
Cornwall, England		Minimum of 0.28		Brice et al. (1981)
Hawaii			~ 0.03	Junge (1957)
North and South Pacific	< 0.01 0.2	< 0.05		Huebert and Lazrus (1980b)
North and South Pacific			~0.5	Tsunogai (1971)

Data on aerosol NH₄⁺ concentrations are also sparse (See Table 16.6). Huebert and Lazrus (1980b) found that the ammonium concentration ranged from < 0.01 to 1 µg m^-3 over the North and South Pacific and Gulf of Mexico. Only two samples were above 0.2 µg m^-3 and they were from the Gulf of Mexico. Remote marine samples were generally well below 0.1 µg m^-3. Again, no correlation was found between NH₄⁺ and NO₃, SO₄²-, or Cl. This, along with the NH₄⁺ aerosol size distribution, suggests the NH₄⁺ does not come directly from the ocean on sea salt particles. Georgii and Lenhard (1978) point out that NH₃ and NH₄⁺ concentrations decrease over the ocean as one moves away from continental areas, with continental areas as the primary NH₃ source. The data are too sparse to date to suggest that the concentrations are higher over the Atlantic than the Pacific. A background marine aerosol NH₄⁺ concentration appears to be ~ 0.05 µg m^-3. While this almost certainly varies geographically, the details of the variability cannot be detected from the available data.

Concentrations of NO₃ and NH₄⁺ in precipitation collected in marine areas are presented in Tables 16.7 and 16.8. The NO₃ concentration in rain varies considerably, according to the available data, with means ranging from 10 to 1000 µg kg^-1. It is difficult to know at present whether these differences are real. Some samples include wet and dry deposition, others wet only, samples were stored varying lengths of time, etc. However, there are a number of studies of NO₃ (and NH₄⁺) in remote marine rains currently getting under way, and reliable numbers should be available soon. Note that the available data for NH₄⁺ in rain show much less scatter, with most values falling between 10 and 100 µg kg^-1. Also note that with a mean aerosol NH₄⁺ concentration of ~ 0.05 µg m^-3, this results in an aerosol NH₄⁺ wash-out factor W ranging from 240 to 2400, which is reasonable. For NO₃ with rain concentrations from 10 to 1000 µg kg^-1 and assuming a mean aerosol NO₃ concentration ~ 0.1 µg m^-3, the aerosol wash-out factor W ranges from 120 to 12,000. The higher values seem unrealistic and may be a result of a significant amount of dry deposition being collected with the rain (since NO₃ is on the larger particles) or effective scavenging of vapour-phase HNO₃ by the rain.

As mentioned previously, Gravenhorst et al. (1979) point out that for the marine aerosol over the North Atlantic, the maximum nitrate concentration is found on particles with radii ~ 1 µm. This coincides with the lower end of the sea salt size distribution. On the other hand, 90 percent of the mass of NH₄⁺ is found on particles with radii < 0.5 µm, the same size in which most of the excess SO₄^2- is found. Gravenhorst et al. (1979) suggest this distribution results from the heterogeneous reaction of gaseous NH₃ with acidic sulphur containing particles with radii < 0.5 µm, as suggested also by Georgii (1978). Gravenhorst et al. (1979) point out that the different size distributions argue against a significant gas phase reaction of NH₃ with HNO₃. He does suggest the possibility, however, of attachment of HNO₃ or NO₂ to the alkaline sea salt particles and not to the acidic sulphate particles. In this case, the nitrate mass distribution should be proportional to the surface area of the sea salt particles, not their mass, and this may explain why the NO₃ maximum is on the size of the smaller sea salt particles. Thus Gravenhorst et al. (1979) suggest that NH₃ and NO₂/HNO₃ interact with the aerosol population in different ways due to the different acidic and basic nature of these gases and the various size fractions of the marine aerosol.

Table 16.7 Nitrate concentration in precipitation from marine areas (µg kg^-1)

Location	Range	Median	Mean	Reference
Cloud water, Hawaii			~ 600	Parungo et al. (1981)
Rain, North Atlantic			~1100	Asman and Ridder (1981)
Rain, Hawaii,
surface to 2000 m	4260	< 70	~ 70	Eriksson (1957)
Rain, Hawaii,
surface to 3400 m	<20 300			J. Miller (personal communication)
Rain, Samoa	< 1 120	10	12	Pszenny and Duce (1981)
Rain, Samoa	< 20 150		~ 100	D. Nelson (personal communication)
Table 16.8 Ammonium concentration in precipitation from marine areas ( µg kg-1)
Location	Range	Median	Mean	Reference
Rain, North and
South Pacific	10220	50	80	Tsunogai (1971)
Cloud water, Hawaii			~170	Parungo et al. (1981)
Rain, Hawaii	20 90	40	50	Eriksson (1957)
Rain, North Atlantic		Often less than 10 µg/kg		Asman and Ridder (1981)
Rain, Samoa			< 20	D. Nelson (personal communication)
Rain, 500 km north
of Samoa			5*	Williams (1967)
Rain, Samoa			< 20	Pszenny and Duce (1981)
*1 sample.

One must also mention the possibility of organic N compounds in the marine aerosol. Wilson (1959) investigated snow on mountain-sides in remote areas of New Zealand and found organic N concentrations from 20 to 200 µg kg^-1, similar to the NH₄⁺ concentration. He detected very little NO₃. Similarly Dean (1963) found an organic N concentration of 170 µg kg^-1 in a coastal New Zealand rain while Williams (1967) found 23 µg kg^-1 organic N in a rain sample collected north of Samoa. No data on organic N in aerosols are available, but one might suspect that much of the organic N may come directly from the sea, similar to organic P. Note that for an organic N concentration in rain of 50 µg kg^-1 and an aerosol wash-out factor W of 1000, the aerosol organic N concentration would have to be ~ 0.06 µg m^-3 as N, a substantial concentrations compared with background aerosol concentrations of NO₃ and NH₄⁺. Clearly this requires further study.

16.8 CONCLUSIONS

In the discussion of the marine aerosol chemistry and seaair exchange of carbon, sulphur, phosphorus, and nitrogen, we have seen a number of instances where these element cycles interact. The clearest case is for N and S, with the likelihood that most of the NH₄⁺ and SO₄^2- on the small size marine aerosols are present as a result of the different acid/base characteristics of NH₃ and SO₂/SO₃ and their condensed phases. Phosphorus transport from sea to air is clearly related to organic carbon transport, and a significant quantity of marine aerosol nitrogen may also be associated with organic carbon.

We still have a discrepancy of almost an order of magnitude in the global production estimates for atmospheric sea salt. More accurate estimates should now be able to be made. The chemical composition of film drops, their size distribution, and their total production must be evaluated. Actual measurements of wet and dry deposition in marine areas are few. Remote rain data are sparse and dry deposition data are virtually non-existent and difficult to interpret in any event. Application of dry deposition velocities as a function of particle size to measured size distributions of the substance of interest is probably the best approach to evaluate dry deposition at present. Obviously reliable atmospheric concentration data are required for all these substances from all areas of the global ocean, particularly the South Atlantic, North and South Pacific, and Indian Ocean. Analysis as a function of particle size is of particular importance. In truly remote marine areas, however, concentrations of many of these substances are quite low, and state-of-the-art analytical techniques coupled with stringent precautions against local contamination are required to obtain accurate values.

In terms of the individual elements and their cycles, an accurate evaluation of the net input of carbon to the ocean is required. Stable isotope studies and investigation of elemental carbon in lake sediments on mid-ocean islands and in coastal sediments offer a good possibility of yielding valuable information. Detailed measurement of the classes of organic compounds as well as specific compounds present in the marine aerosol will aid significantly in evaluating source regions which may have `marker' compounds.

For the sulphur cycle, we must accurately evaluate the possibility of fine particle sulphate production directly by the ocean. In addition, in concert with the more accurate estimate of atmospheric sea salt particle production, we must be able to evaluate the total SO₄^2- production by the sea more accurately. Finally and obviously, we must understand the SO₂, SO₄^2- NH₃, NH₄⁺ interactions in the fine-particle marine aerosol.

In the case of phosphorus, to obtain net input values of phosphorus to the ocean we need to obtain information on phosphorus speciation in rain and dry deposition samples from marine areas. More detailed information on the concentration and speciation of phosphorus as a function of particle size is also required.

For the nitrogen cycle we require accurate measurements of NH₄⁺ and NO₃ in precipitation and dry deposition, and the role of organic nitrogen compounds, if any, must be determined. Of particular importance is evaluating the proposed mechanisms of Gravenhorst et al. (1979) for explaining the sources and size distribution of NH₄⁺ and NO₃ on marine aerosols.

ACKNOWLEDGMENTS

16.9 REFERENCES

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Bonsang, B., Nguyen, B. C., Gaudry, A., and Lambert, G. (1980) Sulfate enrichment in marine aerosols owing to biogenic gaseous sulfur compounds, J. Geophys. Res., 85, 7410-7416.

Bruyevich, S. V., and Kulik, Ye. Z. (1967) Chemical interaction between the ocean and the atmosphere (Salt exchange), Oceanology, 7, 279-293.

Buat-Menard, P., Morelli, J., and Chesselet, R. (1974) Water soluble elements in atmospheric particulate matter over the tropical and equatorial Atlantic, J. de Rech. Atmos., 8, 661-673.

Cipriano, R. J., and Blanchard, D. C. (1981) Bubble and aerosol spectra produced by a laboratory `breaking wave', J. Geophys. Res., 86, 8085-8092.

Darzi, M., and Wincheter, J. W. (1981) Marine aerosol composition in the Indian Ocean, Abstract for the Symposium on the Role of the Oceans in Atmospheric Chemistry, IAMAP Third Scientific Assembly, Hamburg, FRG.

Duce, R. A. (1978) Speculation on the budget of particulate and vapor phase non-methane organic carbon in the global troposphere, Pure Appl. Geophys., 116, 244-273.

Duce, R. A., and Hoffman, E. J. (1976) Chemical fractionation at the air/sea interface, Ann. Rev. Earth Planetary Sci., 4, 187-228.

Eichmann, R., Ketseridis, G., Schebeske, G., Jaenicke, R., Hahn, J., Warneck, P., and Junge, C. (1980) n-Alkane studies in the troposphereII: gas and particulate concentrations in Indian Ocean air, Atmos. Environ., 14, 695-703.

Eriksson, E. (1957) The chemical composition of Hawaiian rainfall, Tellus, 9, 509-520.

Eriksson, E. (1960) The yearly circulation of chloride and sulfate in nature; meteorological, geochemical, and pedological implications, II, Tellus, 12, 63-109.

Garland, J. A. (1981) Enrichment of sulfate in maritime aerosols, Atmos. Environ., 15, 787-791.

Georgii, H.-W. (1978) Large scale spatial and temporal distribution of sulfur compounds, Atmos. Environ., 12, 681-690.

Graham, W. F., and Duce, R. A. (1979) Atmospheric pathways of the phosphorus cycle, Geochim. Cosmochim. Acta, 43, 1195-1208.

Graham, W. F., and Duce, R. A. (1981) The atmospheric input of phosphorus to remote tropical islands, Pacific Science, 35, 241-255.

Gravenhorst, G. (1978) Maritime sulfate over the North Atlantic, Atmos. Environ., 12, 707-713.

Gravenhorst, G., Muller, K. P., and Schutz, L. (1981) Inorganic nitrogen compounds over the Atlantic, Abstrct for the Symposium on the Role of the Oceans in Atmospheric Chemistry, IAMAP Third Scientific Assembly, Hamburg, FRG.

Griffin, J. J., and Goldberg, E. D. (1979) Morphologies and origin of elemental carbon in the environment, Science, 206, 563-565.

Hoffman, E. J., and Duce, R. A., (1974) The organic carbon content of marine aerosols collected on Bermuda, J. Geophys. Res., 79, 4474-4477.

Huebert, B. J. (1980) Nitric acid and aerosol nitrate measurements in the equatorial Pacific region, Geophys. Res. Lett., 7, 325-328.

Huebert, B. J., and Lazrus, A. L. (1980a) Tropospheric gas-phase and particulate nitrate measurements, J. Geophys. Res., 85, 7322-7328.

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Junge, C. E. (1963) Air Chemistry and Radioactivity, New York, Academic Press. 

Ketseridis, G., Hahn, J., Jaenicke, R., and Junge, C. (1976) The organic constituents of atmospheric particulate matter, Atmos. Environ., 10, 603-610.

Lawson, D. R., and Winchester, J. W. (1978) Sulfur and trace element concentration relationships in aerosols from the South American continent, Geophys. Res. Lett., 5, 195-198.

Liss, P. S. (1975) Chemistry of the sea surface microlayer, in Riley, J. P., and Chester, R. (eds) Chemical Oceanography, New York, Academic Press, Vol. 2, 193-244.

MacIntyre, F. (1968) Bubbles, a boundary layer `microtome' for micron-thick samples of a liquid surface, J. Phys. Chem., 72, 489-592.

MacIntyre, F., and Winchester, J. W. (1969) Phosphate ion enrichment in drops from breaking bubbles, J. Phys. Chem., 73, 2163-2169.

Maenhaut, W., Raemdonck, H., Selen, A., Van Grieken, R., and Winchester, J. W. (1981) Characterization of the atmospheric aerosol over the central east Pacific, Abstract for the Symposium on the Role of the Oceans in Atmospheric Chemistry, IAMAP Third Scientific Assembly, Hamburg, FRG.

Meinert, D. L., and Wincheter, J. W. (1977) Chemical relationships in the North Atlantic marine aerosol, J. Geophys. Res., 82, 1778-1782.

Neumann, G. H., Fonselius, S., and Wohlman, L. (1959) Measurements on the content of non-volatile organic material in atmospheric precipitation, Int. J. Air Pollut., 2, 132-141.

Nyberg, A. (1977) On airborne transport of sulphur over the North Atlantic, Q. J. Roy. Met. Soc., 103, 607-615.

Parungo, F. P., Nagamoto, C., Nolt, I., and Nickerson, E. (1981) Chemical analysis of cloud water and aerosols over Hawaii, Abstract for the Symposium on the Role of the Oceans in Atmospheric Chemistry, IAMAP Third Scientific Assembly, Hamburg, FRG.

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Pszenny, A., and Duce, R. (1981) Major anions and cations in Samoa rains, SEAREX Newsletter, Vol. 4, No. 2, May.

Rosen, H., Novakov, T., and Bodhaine, B. A. (1981) Soot in the Arctic, Atmos. Environ., 15, 1371-1374.

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Seiler, W., and Crutzen, P. J. (1980) Estimates of gross and net fluxes of carbon between the biosphere and the atmosphere from biomass burning, Climatic Change, 2, 207-247.

Slinn, S. A., and Slinn, W. G. N. (1980) Predictions for particle deposition on natural waters, Atmos. Environ., 14, 1013-1016.

Sutcliffe, W. H., Baylor, E. R., and Menzel, D. W. (1963) Sea surface chemistry and langmuir circulation, Deep-Sea Res., 10, 232-243.

Tsunogai, S. (1971) Ammonia in the oceanic atmosphere and the cycle of nitrogen compounds through the atmosphere and hydrosphere, Geochem. J., 5, 57-67.

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Wilson, A. T. (1959) Surface of the ocean as a source of air-borne nitrogenous material and other plant nutrients, Nature, 184, 99-101.

COMMENT TO CHAPTER 16

R. J. CHARLSON

Besides being a dominant type of aerosol in terms of mass, marine aerosols serve several important functions that are closely connected to the meteorological process involved in element cycles. Perhaps the most important of these functions is the nucleation of clouds in maritime air masses, with the consequent possibility of precipitation and a flux of dissolved or scavenged material from the atmosphere to the earth's surface. Marine aerosol is thus intimately involved in the chemistry of rain-water, and the cycling of SO₄^2-, NH₄⁺, Na⁺, Cl^- and other species. The understanding of its composition is therefore a necessary part of understanding these aspects of cycles.

With a composition far from that of sea-water, the aerosol in marine settings is the result of several simultaneous processes. Besides direct production of rather coarse (several µm) salt particles by mechanical processes at the sea surface, the shattering of bubble films gives rise to smaller particles enriched in organic species. In addition, production of sub-µm sulphate particles by the oxidation of SO₂ adds to the complexity of marine aerosol chemistry.

The data presented by Duce generally corroborate the qualitative, process-based model of aerosol mass distribution presented by Taylor et al. (Chapter 4, this volume). It is important to note that, even in marine settings, a fine particle (sub-µm) mass concentration of sulphate can be identified. As a result of up to ca. 0.5 µg m^-3 in mass concentration of fine-particle or excess sulphate and co-existing cations, the non-sea-salt fraction of ions in cloud and precipitation water may be significantly influenced. This is consistent with recent observations that precipitation water in some remote maritime areas has pH values below 5 (Kerr, 1981).

In this connection, the stoichiometric relationship of H⁺, NH₄⁺ and excess sulphate in fine particles is of some interest. Comparing Tables 16.2 and 16.5 from Duce, sulphate in the northern hemisphere around 0.7 µg m^-3 and ammonium around 0.1 µg m^-3 yields an empirical molar ratio of [NH₄⁺]/[SO₄^2-] 1. Although the data were not gathered simultaneously, it seems clear that excess sulphate is present in amounts that are larger than would be given by (NH₄)₂SO₄. If the other cation primarily associated with excess SO₄^2- is H⁺, then [NH₄⁺] [H⁺], and the empirical formula of the aerosol would be NH₄HSO₄. Since many of the other inorganic materials may be concentrated in larger size classes, the excess sulphate in the small size classes implies that they could be quite acidic. The data indicate that this acidity is more common in the northern than in the southern hemisphere. These data also indicate that the source strength of NH₃ is low and smaller than the production rate of H₂SO₄. These results are consistent with the low NH₃ values of Ayers and Gras (1980) in remote maritime air.

It seems clear from Duce's summary that the global flux of sea salt and its associated sea-water [SO₄^2-] should be re-evaluated, and it is likely that total fluxes will be found to be larger than those in the older literature (cf. Freney et al. Chapter 2, this volume). However, judging the overall importance of this cycle solely on the basis of mass may be misleading. A large fraction (perhaps most) of the coarse sea salt particles have very short lifetimes (hours or less). They are thus confined to a layer close to the sea surface and do not effectively interact with the chemical processes of the rest of the atmosphere or with clouds. In any case, the number concentration of salt particles, which would be important for cloud nucleation, is likely to be small.

Finally, the natural, marine fine-particle sulphate aerosol seldom, if ever, approaches the concentrations found in industrialized regions. Hence, even though the global burden of naturally occurring, marine excess sulphate aerosol may be appreciable, the much higher concentrations of sulphate aerosol in industrialized regions must derive mainly from land-based sources.

Ayers, G. P., and J. L. Gras, (1980) Ammonia gas over the southern ocean, Nature, 284, 539-540.

Freney, J. R., Ivanov, M. V. and Rodhe, The Sulphur Cycle, Chapter 2, this volume. 

Kerr, R. A. (1981) Is all acid rain polluted?, Science, 212, 10-14.

Taylor, G. S., Baker, M. B. and Charlson, R. J., Atmospheric heterogeneous interactions of C, N and S cycles: the role of aerosols and clouds, Chapter 4, this volume.