with the atmosphere, giving rise to a net air-sea gas ex-
change flux of N₂O to the atmosphere. Nitrous oxide is
produced in subsurface waters both as an intermediate
of denitrification, the reduction of nitrate ion (NO₃⁻) to
nitrogen (N₂), and as a trace by-product of nitrification,
the oxidation of NH₄⁺ to NO₃⁻. Denitrification occurs
under suboxic to anoxic conditions, and is thought to
take place mainly in restricted low-oxygen regions such
as the eastern tropical Pacific and the Arabian Sea, and
in the sediments of continental shelves. The distribu-
tion of nitrification is thought to be more widespread,
since it occurs under aerobic conditions in association
with the internal recycling of fixed nitrogen.

During the past decade there have been significant
improvements in our understanding of oceanic proc-
esses of N₂O production, of the distribution of N₂O in
the surface and subsurface ocean, and of the magnitude
of the oceanic N₂O source to the global atmosphere. The
relative roles of nitrification and denitrification proc-
esses have been addressed by measuring nitrogen sta-
ble isotopes and their fractionation between N₂O and
other dissolved nitrogen-bearing compounds. The in-
terpretation of these difficult measurements is compli-
cated by the likelihood that both nitrification and
denitrification are coupled in many oceanic systems, and
no clear picture has yet emerged. There have also been
recent advances in the study of air-sea gas exchange
processes, as indicated in Sect. 2.8.1, which will lead to
improvements in the quantification of exchange coeffi-
cients as a function of wind speed.

Finally, our understanding of the large-scale distri-
bution of N₂O in the oceans has been improved through
a number of shipboard measurement programs, such
as those associated with the World Ocean Circulation
Experiment (WOCE) and Joint Global Ocean Flux Study
(JGOFS) programs. These have generally reinforced our
view that open ocean upwelling regions along eastern
ocean boundaries and in equatorial and coastal regions,
represent major sources of atmospheric N₂O. By con-
trast, the great subtropical gyres, which represent a large
portion of the surface area of the oceans, are relatively
close to atmospheric equilibrium for N₂O. In recent
years, some extremely high N₂O concentrations have
been found in the eastern Arabian Sea, in suboxic wa-
ters over the Indian Shelf (Naqvi et al. 2000). Since an-
thropogenic impingements on the coastal ocean may
cause an increase in hypoxia, suboxia, and anoxia in
some areas, these recent observations from the Arabian
Sea are provocative. By modelling these distributions
together with the wind field (e.g. Nevison et al. 1995),
we have come to believe that the global oceans consti-
tute a net source to the atmosphere of about 4–5 Tg of
N₂O, or about one third of the global natural source
strength. This value may increase as more is learned
about the diverse distribution of N₂O in coastal waters
(e.g. Seiwinger and Kroeze 1998).

2.8.2.6 Dimethylsulphide

2.8.2.6.1 Introduction

Dimethylsulphide was discovered in ocean waters some
30 years ago by Lovelock et al. (1972). However, it re-
mained a compound of marginal scientific interest for
about a decade, until it was established that DMS is the
main volatile sulphur compound emanating from the
oceans and therefore plays a major role in the atmo-
spheric sulphur cycle (e.g. Nguyen et al. 1978; Leck and
Rodhe 1991). Interest in the biogeochemical cycle of
DMS increased sharply again in the late 1980s, when
Charlson et al. (1987) proposed a hypothesis inking bio-
genic DMS emission and global climate. In short, this
hypothesis states that DMS released by marine phyto-
plankton enters the troposphere and is oxidised there
to sulphate particles, which then act as cloud condensa-
tion nuclei (CCN) for marine clouds (see Box 4.3, how-
ever, regarding the utility of the CCN concept). Changes
in CCN concentration affect the cloud droplet number
concentration, which influences cloud albedo and con-
sequently climate. Large-scale climate change, in turn,
affects the phytoplankton number and speciation in the
oceans and thereby completes, but does not necessarily
close, a feedback loop. Recent assessments of the DMS-
climate link can be found in Watson and Liss (1998) and
Bigg and Leck (2001) (also see Chap. 4).

In the years since publication of the DMS-CCN-cli-
mate hypothesis, almost 1000 papers have been pub-
lished discussing the distribution and biogeochemistry
of DMS (and its precursors) and its link to climate. Sev-
eral IGAC-inspired studies have addressed aspects of
the DMS-aerosol-climate connection, most prominently
among them ASTEX/MAGE (e.g. Huebert et al. 1996),
ACE-1 (e.g. Bates et al. 1998), and AOE-91, 96 (e.g. Leck
et al. 1996, 2001). As a result of these projects, and the
large number of independently conducted studies re-
lated to the DMS-climate hypothesis, we now understand
many of the details of DMS production in the oceans,
its transfer to the atmosphere, and the atmospheric oxi-
dation processes (see Chap. 3) that lead to the forma-
tion of aerosols (see Chap. 4) that can act as CCN. How-
ever, in spite of this progress, fundamental gaps remain
in our understanding of key issues in this biosphere-
climate interaction, in particular with regard to the proc-
esses that regulate the concentration of DMS in seawater.
While the basic processes have been identified, and even
quantified in specific locations (e.g. Bates et al. 1994;
Simo and Pedro-Alio 1999), generally applicable mod-
els of DMS-plankton relationships are still in their in-
fancy (e.g. Gabric et al. 1993; Jodwali's et al. 2000). There-
fore, we are still not able to represent the DMS-CCN-
climate hypothesis in the form of a process-based, quan-
titative, and predictive model. Even the overall sign of
the feedback cannot be deduced with certainty, since it
is not known yet if a warming climate would result in an increase or decrease of DMS emissions. Glacial-to-interglacial changes in the amounts of DMS oxidation products in polar ice cores cannot answer this question unambiguously, as they may reflect variations in atmospheric transport patterns as much as differences in DMS production (e.g. Whunget al. 1994), as is discussed in detail in Sect. 2.3.

Early, limited data sets had suggested a possible correlation between DMS and phytoplankton concentration (e.g. Andreae and Barnard 1984). This correlation is particularly evident in vertical profiles of DMS and chlorophyll a, which in most instances show a sharp drop of both parameters around the depth corresponding to a light penetration of one percent of the surface light flux. Close correlations between DMS and phytoplankton densities were also found in situations where a single species accounted for much of the DMS production or phytoplankton biomass (e.g. Barnard et al. 1984; Matrai and Keller 1994). These findings led to the hope that global DMS distributions could be estimated from chlorophyll concentrations obtained by remote sensing, but experimental investigations of this proposal were not encouraging (e.g. Matrai et al. 1993), except in frontal regions (e.g. Holligan et al. 1993; Belviso et al. 2000). Furthermore, a statistical analysis of almost 16 000 measurements of DMS in surface seawater failed to show any useful correlations between DMS and chlorophyll or other chemical or physical parameters (Kettle et al. 1999). One reason for the absence of a general correlation between plankton biomass and DMS is that the intracellular concentration of its metabolic precursor, dimethylsulphoniopropionate (DMSP), varies between different phytoplankton species over a range of five orders of magnitude. While it is clear that some taxonomic groups typically contain higher amounts of DMSP, these relationships are by no means clear-cut (e.g. Keller et al. 1989). At least as important, however, are the complexities of DMS cycling by biological and abiotic processes in the surface ocean, which will be addressed below.

### 2.8.2.6.2 Physiological and Ecological Controls of DMS Production

The pathways of DMS biosynthesis in phytoplankton have been studied and have shed light on potential regulating mechanisms such as nitrogen nutrition (e.g. Gröne and Kirst 1992; Keller et al. 1999a,b), temperature (Baumann et al. 1994), and light (Vetter and Sharp 1993; Matrai et al. 1995). While DMSP, and sometimes DMS, is directly released by phytoplankton, zooplankton also play a role by grazing, or avoiding, DMSP-rich cells (e.g. Dacey and Wakeham 1986; Wolfe et al. 1997; Tang 2000).

Very high concentrations of DMS and dissolved DMSP have been reported from several coastal and/or

high latitude areas, especially where blooms of DMSP-producing phytoplankton such as the coccolithophore *Emiliania huxleyi* and the prymnesiophyte *Phaeocystis pouchetii* occur (e.g. Malin et al. 1993; Barnard et al. 1984). In this context, it is interesting to note that Kettle et al.'s (1999) DMS database revealed that high DMS regions corresponded roughly to the coccolithophorid bloom areas derived by Brown and Yoder (1994) from remotely sensed ocean colour data. Prymnesiophytes (including coccolithophores) and dinoflagellates are phytoplankton groups that tend to have high DMSP cell quotas (Keller et al. 1989) and, not surprisingly, DMS is often relatively high when these groups dominate the phytoplankton assemblage. Diatoms, on the other hand, tend to have low intracellular DMSP concentrations and it is generally observed that diatoms are less important DMSP producers in the field (e.g. Keller et al. 1989). Predicting DMS concentrations from the algal assemblage is not straightforward, however. For example, Matrai and Vernet (1997) reported that DMS concentrations were as high in diatom-dominated, Arctic waters as they were in those dominated by *Phaeocystis* sp. It is now recognised that some phytoplankton species not only produce high intracellular concentrations of DMSP, but they also have cell-surface (Stefels and Dijkhuizen 1996) and intracellular (Steinke et al. 1996) DMSP lyase enzymes that may be involved actively in DMS production, thereby contributing further to the elevated DMS concentrations associated with these organisms. The ecological roles of these lyase enzymes are not well understood but several recent studies have pointed to very interesting functions such as in grazing deterrence, carbon acquisition, and bacterial inhibition (Noordkamp et al. 1998; Wolfe and Steinke 1996; Wolfe et al. 1997).

Blooms of marine phytoplankton provide convenient natural "laboratories" for investigating the production of DMS in relation to phytoplankton community dynamics and species succession and associated processes, including grazing and bacterial turnover. However, this apparent focus on "hotspots" of DMS production in relatively nutrient rich areas can be criticised in that oligotrophic areas of the oceans, which generally have relatively low levels of DMS and DMSP throughout the year, make up a large fraction of the total ocean area and so must contribute significantly to the total global flux of DMS (Bates et al. 1992). These pioneering studies established the link between phytoplankton and DMS levels, but failed to account for a large part of the natural variability in DMS concentrations. There have been rather few actual DMS time-series studies (Leck et al. 1990; Turner et al. 1996a; Dacey et al. 1998), all of which noted seasonal periods of elevated DMS concentrations.

We now realise that bacterial processes are also very important in the overall DMS cycle. More isolates of bacteria are available with which to study biochemical pathways and physiology of DMSP and DMS metabolism
(e.g. Ledyard and Dacey 1994; Yoch et al. 1997). New methods, including use of ³⁵S tracers, improved inhibitors, and molecular genetics techniques have allowed ever more sensitive analyses of DMSP-DMS cycling rates and fates, and have permitted more detailed examination of the complex microbial communities involved (e.g. González et al. 1999; Wolfe and Kiene 1993). The potential for DMS production from dissolved DMSP is quite large (e.g. Kiene 1996b; van Duyl et al. 1998), but recent studies indicate that most of the DMSP in the sea is not converted to DMS. A demethylation-demethiolation pathway leading to production of methanethiol (MeSH) can account for 70-95% of DMSP metabolism *in situ* thereby diverting sulphur away from DMS (Kiene 1996a). The predominance of this non-DMS producing demethylation-demethiolation pathway is explained by the fact that bacteria use it to assimilate the sulphur from DMSP into protein amino acids (Kiene et al. 1999). Further understanding of this DMSP-DMS-MeSH-bacteria interaction is critical because a relatively small change in the yield of DMS from DMSP could have a major impact on the gross production of DMS, which would then be available for sea-air exchange.

Removal of DMS from the water column by biological and photochemical mechanisms also exerts a great influence on the net accumulations of DMS in surface waters. Slow biological degradation of DMS may partially explain the rise in DMS concentrations observed at the peak and initial decline phases of phytoplankton blooms (e.g. Matrai and Keller 1993; Nguyen et al. 1988). Net consumption of DMS appears to occur in the later stages of blooms after DMS-consuming bacteria have had time to develop (Kwint et al. 1996; van Duyl et al. 1998). The photochemistry of DMS in seawater remains poorly understood, despite the fact that it has been identified as a major removal mechanism under some circumstances (e.g. Kieber et al. 1996; Sakka et al. 1997; Brugger et al. 1998). DMS photooxidation appears to depend on photosensitisers in seawater, which are most likely part of the coloured dissolved organic matter (CDOM) (Dacey et al. 1998). In the open ocean CDOM originates from autochthonous primary productivity and food web processes so the interaction with DMS is probably complex. Add to this the fact that DMS producing and consuming bacterial populations are likely to be strongly influenced by UV-B in surface waters, and one can easily see the importance of understanding photophysical effects on the DMS cycle. Recently, it has been shown that viruses are significant agents in the control of bacteria and phytoplankton. Viral infections can cause a total release of intracellular DMSP (Hill et al. 1998) and viruses are known to infect DMSP-containing bloom organisms such as *Emiliania huxleyi* (Brussaard et al. 1996) and *Phaeocystis* sp. (Malin et al. 1998). It seems clear from studies such as these that the overall food web dynamics, including macro- and microzooplankton graz-

ing, bacterial, and viral activities, as well as the physicochemical dynamics of the upper ocean (e.g. incoming solar radiation, mixing, temperature, air-sea exchange) are important factors governing DMS accumulation.

Modelling efforts have expanded our understanding of DMS production, both for field situations (e.g. Gabric et al. 1999) and laboratory systems (Laroche et al. 1999). However, our current knowledge base is not sufficient to develop and constrain predictive DMS production models for diverse biogeographic regions, in order to allow interpretation of the role of DMS in climate change, for example. Future research will need to focus on (1) gaining a full understanding of the processes that control DMS production and allow the prediction of DMS emissions, and (2) obtaining much more data concerning spatial, temporal, and interannual variation in the concentration of DMS and related compounds. Emphasis on undersampled areas and seasons would be valuable. For process studies, there is an increasing need to cross disciplinary and international boundaries to bring together experts on different aspects of DMS and related compounds for integrated field campaigns. For analysis of variability, "remote" sampling systems could be considered (such as attempted in ACE-1). It might be possible to develop a buoy-mounted monitoring system whereby samples were stored on a camel for later analysis. Alternatively, we might follow the example of the pCO₂ measuring community, who have demonstrated that it is feasible to employ unmanned instruments on merchant ships (Cooper et al. 1998a). This would enable the collection of large data sets during long passage routes, covering diverse biogeographic areas, and different seasons, and the chance to investigate interannual variability at relatively low cost. New techniques will be needed to circumvent the present lack of a reliable storage method for DMS samples. In the first instance, it might be more realistic to concentrate on DMSP analyses.

### 2.8.2.7 Carbonyl Sulphide

The oceans represent approximately 30% of the total atmospheric source of COS, and much of the oceano-graphic work on COS over the last decade has focussed on assessing the spatial and temporal distributions of COS concentration and understanding the processes that control its temporal and spatial distribution. The photochemical source of COS was first recognised by Ferek and Andreae (1984), who demonstrated a clear diurnal cycle in the sea surface concentration of the compound. A mechanism of formation of COS was proposed by Pos et al. (1998) who suggested that the photochemical production of COS and carbon monoxide proceeds along a coupled pathway which first involves the photochemical formation of an acyl radical from coloured dissolved organic matter (CDOM). Flöck et al.
(1997) and Ulshöfer et al. (1996) suggested that cysteine is probably implicated in the reaction mechanism of COS formation as the result of its reactivity and abundance in the oceans. The photochemical COS production in natural seawater is probably not limited by the concentration of a precursor sulphur compound but rather by the concentration of CDOM represented by its ultraviolet attenuation coefficient (Ulshöfer et al. 1996; Uher and Andreae 1997). Zepp and Andreae (1994) and Weiss et al. (1995) quantified the wavelength dependence of COS photoproduction from CDOM and found that quantum efficiency of photoproduction decreases monotonically with increasing wavelength. The dark (or non-photochemical) production of COS has been proposed on the basis of the non-zero COS concentration observed at ocean depths where there is no photochemical production and where there is no mixing from the surface (Radford-Knoery and Cutter 1994; Flöck and Andreae 1996) and also on the basis of careful interpretation of sea surface COS concentration measurements using inverse models (Ulshöfer 1995). COS hydrolysis varies as a function of temperature and pH and has been evaluated several times over the last decade (Elliott et al. 1989; Radford-Knoery and Cutter 1994; Uher and Andreae 1997).

Recent models have used laboratory results for the photoproduction and hydrolysis rate constants to explain COS sea surface measurements obtained during expeditions made in the 1980s and 1990s (see Ulshöfer (1995) for a review of recent sea surface COS concentration measurements). Andreae and Ferek (1992) developed the first chemical box model to explain the diurnal variation of COS in terms of photochemical formation and hydrolysis destruction. Ulshöfer (1995) adopted an optimisation scheme based on the coupled photochemical-mixed layer used by Kettle (1994) to calculate the photoproduction and dark production constants for COS from a series of sea surface measurements made between 1992 and 1994 in the North Atlantic Ocean. von Hobe (1999) extended this work for other models and expedition measurements. Najjar et al. (1995) generalised a simplified coupled physical-chemical model on a global scale to investigate the sensitivity of COS sea surface concentration on ozone reduction and tropospheric increases of carbon dioxide. Kettle and Andreae (1998) and Preiswerk and Najjar (1998) have used existing measurements of the CDOM absorption coefficient of seawater to predict a seasonal variation in the absolute COS concentration, with maximum values at high latitudes in the summer of either hemisphere.

Future work on COS should aim to quantify more accurately the role of the oceans as a source or sink of the gas to the atmosphere. The global application of the photochemical production model for COS is currently limited by the absence of an algorithm to predict the global CDOM absorption coefficient and by the sugges-

tion that the apparent quantum yield of COS formation may vary by more than an order of magnitude in different regions of the ocean. The scarcity of profile measurements of COS concentration has been problematic for modelling efforts which have so far been developed to explain only the surface COS concentration distributions. Finally, the precise quantification of the sea-air flux of all gases produced in the upper ocean (including COS) is currently limited by the absence of an effective gas exchange parameterisation based on wind speed, average wave slope, or other measure of upper ocean turbulence, as already indicated.

### 2.8.2.8 The Ocean's Role as Source and Sink of Atmospheric Methyl Bromide and other Methyl Halides

Methyl halides are produced and consumed biologically ($CH_3Br$) (Moore and Webb 1996; Baker et al. 1999); ($CH_3I$) (Moore and Groszko 1999); photochemically ($CH_3I$) (Happell and Wallace 1996); and in surface ocean waters ($CH_3Cl$) (Moore et al. 1996). Recent measurements have shown that the flux of $CH_3Cl$ is significantly less than early estimates (Moore et al. 1996) and that the open ocean is a net sink, rather than a source, for $CH_3Br$ (see below).

Methyl bromide ($CH_3Br$) in the environment began to receive considerable attention in the early 1990s when it was being evaluated as an ozone-depleting gas, along with chlorofluorocarbons, chlorocarbons, and halons. First-order calculations indicated that methyl bromide was likely to be a significant contributor to stratospheric ozone depletion. Before then, only a few studies of $CH_3Br$ in the ocean and atmosphere had been conducted. Lovelock (1975) detected $CH_3Br$ in coastal waters of England and suggested that this gas could have a large natural source. Singh et al. (1983) later reported widespread supersaturations greater than 200% off the west coast of North America, lending support to the ocean as a large natural source of $CH_3Br$. Khalil et al. (1993) suggested that the open ocean was supersaturated in methyl bromide by 40-80%. However, prompted in part by calculations showing that the ocean simultaneously had to be a large sink for $CH_3Br$ because of reaction with $Cl^-$ in seawater (Elliott and Rowland 1995; Jeffers and Wolfe 1996; King and Saltzman 1997), numerous investigations, using *in situ* mass spectrometry-gas chromatography, demonstrated that the ocean on average was a net sink for atmospheric $CH_3Br$, with tropical and subtropical waters of the open ocean highly undersaturated and coastal waters often supersaturated in this gas (Lobert et al. 1995, 1996, 1997; Moore and Webb 1996; Groszko and Moore 1998). Certain species of phytoplankton produce $CH_3Br$, but apparently not at rates sufficient to explain the observed saturation levels (Saemundsdottir and Matrai 1998; Moore et al. 1995;
Scarratt and Moore 1996). Most recently, there have been suggestions that CH₃Br in temperate and coastal waters might undergo a seasonal cycle, with higher concentrations or supersaturations in the spring and early summer and undersaturations the rest of the year (Balker et al. 1999; King et al. 2000). About the same time, it also became clear that chemical and biological removal of CH₃Br in seawater constituted such a large sink for this gas that it would have a profound effect on the lifetime of CH₃Br in the atmosphere, even if the ocean were everywhere a net source (Butler 1994; Yvon et al. 1996b; Yvon-Lewis and Butler 1997). In the latest budget calculations, irreversible loss of atmospheric CH₃Br to the ocean accounts for one-quarter to one-third of the total removal (Kurylo et al. 1999).

These two findings – that the oceanic source was outweighed by its sinks and that the lifetime of atmospheric CH₃Br depended strongly upon its reaction in seawater – necessitated a re-evaluation of the global budget of this gas in the atmosphere. Once the apparently large soil sink was discovered and confirmed (Serça et al. 1998; Shorter et al. 1995; Varner et al. 1999), the calculated budget of atmospheric CH₃Br was no longer in balance. The latest calculations have sinks outweighing sources by 80 Gg yr⁻¹, out of a budget of 205 Gg yr⁻¹ (Kurylo et al. 1999). It is unlikely that this

additional source will come from the ocean, as the current global coverage of surface measurements, although not complete, is representative of the various oceanic regimes, although with reduced coverage of coastal waters. Currently, a small net sink is calculated for the ocean (3–30 Gg yr⁻¹) which is unlikely to change much, unless, of course, there is some significant global change driving it. Furthermore, recent studies are identifying terrestrial sources from plants and salt marshes that are making the budget gap smaller (Gan et al. 1998; Rhew et al. 2000; Dimmer et al. 1999).

Perhaps one of the most significant things to come out of these intensified studies of methyl bromide in the ocean is that other halogen gases may behave in similar, quantifiable ways. Many of these gases, which may include CH₃I, CHBr₃, CH₂Br₂, CH₂BrCl, and C₂H₅Br, among others, also have climatic implications through their chemistry or radiative effects; however, specific studies of them in the past have been limited (e.g. Sturges et al. 1992, 1993; Nightingale et al. 1995). When gases are produced and destroyed in seawater and exchanged with the atmosphere on similar time scales, their exchange with the atmosphere can be controlled in good part by their biogeochemical cycling in seawater. Recent analyses of polar firn air have provided global temporal trends for CH₃Br, while also showing in situ,

<figure><img src="image_21.png" /><figcaption>Arctic 20.04.1997</figcaption></figure>

<figure><img src="image_22.png" /><figcaption>Antarctic 18.09.1997</figcaption></figure>

<figure><img src="image_23.png" /><figcaption>Fig. 2.19. Satellite (Global Ozone Monitoring Experiment, GOME instrument on ERS-2) observations of tropospheric BrO “clouds” in the Arctic and over Antarctica (Wagner et al. 2001). Total columns in the centre of the clouds exceed 10¹⁴ BrO molecules cm⁻². The clouds are visible only in springtime and have a typical lifetime of one to a few days</figcaption></figure>
seasonal production for other organohalogens (Butler
et al. 1999; Sturges et al. 2001).

Large amounts of reactive bromine (and smaller
amounts of chlorine) are also found in polar regions and
near salt pans likely due to oxidation of halides by inor-
ganic reactions (see also Table 3.4). However, the source
of iodine compounds in the coastal marine air and of
bromine in the free troposphere is much more likely to
be the photochemical degradation of organohalogen
compounds (like CH<sub>2</sub>I<sub>2</sub>, or CH<sub>3</sub>Br, respectively) of ma-
rine biogenic origin, as indicated above.

In field campaigns of the IGAC's Polar Air and Snow Chemistry (PASC) activity, it was discovered in the late 1980s and early 1990s (Barrie et al. 1988, 1992; Barrie and Platt 1997) that surface ozone depletion chemistry occurring in spring over the Arctic Ocean is a tropospheric analogue of stratospheric ozone chemistry, with a difference. It is driven by sea salt halogens from heterogeneous reactions occurring in sunlight on surface snow and ice, rather than by halogens from photolysis of spray can propellants. The existence of BrO and ClO as well as Cl and Br reactions with hydrocarbons was well documented in numerous measurements in air just above the surface of frozen marine areas. In spring, BrO was found in the Arctic (Hausmann and Platt 1994; Tuckermann et al. 1997) and Antarctic marine boundary layer (Kreher et al. 1997) by ground based and satellite observations (Fig. 2.19) (Wagner and Platt 1998; Richter et al. 1998; Hegels et al. 1998; Wagner et al. 2001). In addition, measurements made by chemical amplification (Perner et al. 1999), DOAS (Tuckermann et al. 1997), and by the "hydrocarbon clock" technique (Jobson et al. 1994, Solberg et al. 1996; Ramacher et al. 1999) suggest ClO levels in the pmol mol⁻¹ range in the Arctic marine boundary layer. An unexpected link to the mercury cycle was discovered to result in enhanced inputs of mercury to the biosphere in these regions, when long lived elemental mercury is converted to shorter lived particulate and reactive gaseous forms of mercury (Schroeder et al. 1998). Halogen reactions are suspected to be the cause of this conversion of mercury.

As a result of these polar discoveries as well as mod-
elling studies (e.g. Vogt et al. 1996; Sander and Crutzen
1996) (see also Chap. 3), researchers have begun to seek
and confirm the occurrence of reactive halogen com-
pounds (IO, BrO, ClO) from air-surface exchange proc-
esses in other regions (e.g. remote mid- and lowlatitude
marine sites, midlatitudes coastal sites, Dead Sea basin,
and the free troposphere).

2.8.2.9 Primary Marine Aerosols

Primary aerosols are also emitted directly from the oceans.
The work of Blanchard and colleagues (Blanchard 1983)
has shown that bubble bursting at the air-water inter-
face injects aerosols into the atmosphere from two

sources. One is from fragments of the bubble film (film
drops), the other from a jet of water that follows the
bubble burst. Bubbles selectively scavenge high molecu-
lar weight surface-active compounds (Gershey 1983) and
viable particulate material from the water such as bac-
teria and viruses (Blanchard 1983), leading to a consid-
erable enrichment of these organic components in the
aerosol relative to the water. As a result, primary parti-
cles in the marine environment will usually contain a
wide range of biogenic compounds. Long-chain fatty
acids, alcohols, esters, and soluble proteins have all been
found in marine aerosols. Proteinaceous material and
free amino acids are present in marine rain (Mopper
and Zika 1987). In the atmosphere some of these com-
pounds are degraded to form secondary aerosols, such
as the fatty acids which may break down to short chain
forms such as oxalic acid. Others, like the amino acid L-
methionine, are oxidised. Bacteria and remains of or-
ganisms have been observed to become separated from
the other aerosol components in Arctic conditions (Bigg
and Leck 2001). Estimates of the organic components
of marine aerosols in relatively unpolluted environ-
ments vary widely, the order of magnitude being around
10-20% by number, but this may include secondary
aerosols as well as those transported from continents
(Mathias-Maser 1998) (see Chap. 4).

2.8.3 Biological and Chemical Impacts of Atmospheric Deposition on Marine and Estuarine Systems

2.8.3.1 Atmospheric Iron Input to the Ocean and its Role in Marine Biogeochemistry

2.8.3.1.1 Introduction

It is now recognised that a primary transport path for iron found in the ocean is through the atmosphere. Among the first papers to address the importance of atmospherically derived iron were those of Moore et al. (1984) and Duce (1986). These authors calculated the aeolian transport of mineral matter into many areas of the ocean, and pointed out that some fraction of the iron from the mineral matter dissolved into seawater after the dust was deposited to the ocean surface. Duce and Tindale (1991) and, more recently, Jickells and Spokes (2001) have reviewed this topic.

The major reason why atmospheric dust transport
has received considerable research effort over the last
decade is because of the role iron has been hypothesised
to play in controlling marine primary productivity over
large areas of the oceans remote from land. Because of
their distance from riverine and shelf inputs in these
regions (e.g. Southern Oceans, North and Equatorial
Pacific) one of the primary ways in which "new" iron
gets into the system is via deposition from the atmo-
sphere of terrestrially derived material. The idea of iron
being a major control on ocean production is not new.
In the early decades of the 20th century it was hypoth-
esised that the reason why large areas of the Southern
Oceans contained significant amounts of residual con-
ventional plant nutrients (nitrate and phosphate), when
light and other conditions for plant growth were favour-
able (the HNLC, high nutrient-low chlorophyll, regions),
was because of iron deficiency in the water (see, for ex-
ample, Gran 1931; Harvey 1933; Hart 1934; and the recent
review by DeBaar and Boyd 2000). However, it is only
in the last decade that analytical techniques for iron and
field-going experimental approaches have been good
enough to begin to test the hypothesis critically.

2.8.3.1.2 Sources and Transport of Mineral Aerosol to the Oceans

The primary sources of mineral aerosol are arid and semi-arid continental regions (e.g. Tegen and Fung 1994; Duce 1995; IPCC 1996) (see also Chap. 4). The atmospheric concentrations of dust and the deposition of dust to the ocean surface are both very episodic and are primarily associated with the transport of aerosol from dust storms or major dust outbreaks. The typical duration of such dust pulses over the ocean may range from one to four days, and the transport and deposition may also vary seasonally. Due to the episodic character of both the atmospheric dust concentrations and local rainfall, the primary removal process for dust (Duce 1995) input to the ocean in a particular region can often occur during a few events covering a relatively short period of time. For example, the results of one multi-year study showed that half of the annual deposition of dust to the ocean at Midway Island in the central Pacific occurred during only two weeks (Prospero et al. 1989). In Bermuda, Arimoto et al. (1992) found that mineral aerosol concentrations ranged over four orders of magnitude, from 0.001 to 11 µg m⁻³.

We have very few data sets of marine surface dust concentrations collected over long periods of time. In general, the highest atmospheric concentrations of dust in marine areas are found over the North Pacific and the tropical Atlantic. Other high concentration areas are found in the Arabian Sea and the northern Indian Ocean, but there are very limited data in these regions. Accurate estimation or calculation of dust deposition is still quite difficult. An estimate of the geographical distribution of the flux of mineral matter to the global ocean is presented in Fig. 2.20 (Duce et al. 1991). Note that by far the major fraction of mineral dust is deposited in the Northern Hemisphere. The atmospheric deposition has clearly fluctuated significantly in the past, as seen in ice core and deep sea sediment samples (see, for example, Rea 1994; Andersen et al. 1998; Maher and Hounslow 1999). Numerical simulations of the mineral dust cycle are attempting to improve global data sets by linking soil types, particle emissions, gas-particle heterogeneous chemistry, and wind transport in the troposphere with aerosol satellite measurements (e.g. Marticorena and Bergametti 1995; Phadnis and Carmichael 2000) (also see Chap. 6).

2.8.3.1.3 Iron in Mineral Aerosol over the Oceans

The atmospheric deposition of iron is associated with
the eroded mineral aerosol particles, and the iron is
primarily bound in their aluminosilicate matrices. It is
thus possible to convert mineral aerosol concentrations
or fluxes to an iron concentration or flux by knowing
the abundance of iron in the earth's crust. This ranges
from ~3 to 5% (Taylor and McClennan 1985). Typically
a value of 3.5% is used. With a mineral aerosol flux
of 500-2000 Tg yr⁻¹, the input of iron would be ca.
15-100 Tg yr⁻¹. However, before the iron deposited from
the atmosphere can be utilised by phytoplankton, it must
be in a form that is available to these organisms. Proc-
esses that change the solubility or lability of the iron in
the atmosphere will then have potential for influencing

<figure><img src="image_24.png" /><figcaption>Fig. 2.20.
Calculated global fluxes of
atmospheric mineral matter
to the ocean (Duce et al. 1991)</figcaption></figure>
the availability of the iron when the atmospheric mate-
rial enters the ocean. Jickells and Spokes (2001) have
carefully reviewed the information to date on the mecha-
nisms that may control the distribution of dissolved and/
or particulate iron in the material entering the ocean
from the atmosphere. Some studies have observed Fe(II)
in aerosol iron and its formation is postulated to occur
via photochemical reduction of Fe(III) hydroxides. Cer-
tain organic compounds such as oxalate, acetate, and
formate can facilitate this photoreduction. It has been
suggested by several authors that the low pH (0-5) char-
acteristic of the cloud cycling process produces acidic,
hygroscopic aerosols. This combined with possible pho-
tochemical reactions results in an increase in the lability
of crustally derived metals, such as iron, in the atmo-
sphere over that seen in the parent material. In turn, this
will play a role in the availability of the iron when the
aerosol enters the ocean (e.g. Jickells and Spokes 2001;
Zhu et al. 1992). In addition, high ionic strength solutions
and alternating wet and dry cycles during cloud forma-
tion and evaporation would be common. There are likely
to be many such cycles before the particles ultimately
enter the ocean by dry deposition or precipitation.

Jickells and Spokes (2001, and references therein) state, in summary, that it is likely that the overall iron solubility of dry deposited mineral aerosol is <1% at a seawater pH of 8, and that a significant proportion of this iron is photoreduced to Fe(II), which is bioavailable. The solubility of iron in marine rains with a pH of 4-7 is generally 14%. Thus the input of soluble atmospheric iron to the oceans is apparently dominated by wet deposition. These estimates, based on laboratory studies, are somewhat lower than those made earlier by other authors. However, Jickells and Spokes (2001) made other oceanographic approaches to estimate the solubility of atmospheric iron. All of these approaches result in low overall iron solubility, probably less than 2%. Their final conclusion is that approximately 0.8 to 2.1% of the total iron deposited in the ocean is soluble. With a total input of 15 to 100 Tg yr⁻¹, this would result in a total soluble iron atmospheric input of from 0.12 to 2.1 Tg yr⁻¹.

2.8.3.1.4 Iron and Marine Biogeochemistry

Once the atmospheric iron has entered the oceans by either wet or dry deposition, it is hypothesised to play potentially important roles in the primary productivity of surface waters in substantial areas remote from land. These HNLC regions are estimated to cover ~20-25% of the area of the oceans. An up-to-date and detailed assessment of the chemical form of iron in seawater and how this relates to its uptake by marine organisms is to be found in several chapters in the book edited by Turner and Hunter (2001).

John Martin and his colleagues made some of the first
reliable measurements of iron in the oceans and con-

ducted shipboard incubation studies in flasks and car-
boys of HNLC seawater that had been amended with
soluble iron (Martin et al. 1994). The results were prom-
ising (e.g. Martin and Fitzwater 1988) and clearly showed
that addition of iron (normally added as ferrous sul-
phate or other simple inorganic salts) could lead to sub-
stantial increases in plankton growth, as indicated by
increasing chlorophyll concentrations with time in the
experimental flasks. An interesting variant on this ba-
sic experiment, which is particularly relevant in the
present context, was a study conducted in the equato-
rial Pacific by Johnson et al. (1994). They added the iron
in a variety of inorganic and organically complexed
forms, but they also used natural Asian dust aerosol
particles (collected in Hawaii) and added them to one
of the carboys of seawater. In this carboy the rate of
plankton growth was found to be the most rapid and
attained the highest chlorophyll levels, indicating that
the aerosol particles were more effective at promoting
growth than artificial iron supplements.

Other avenues have been explored to attack the problem in a more direct way. Young et al. (1991) monitored natural dust inputs to the North Pacific and examined any resulting change in productivity in the receiving water. Several dust deposition events appeared to be correlated with increases in primary productivity measured in on-deck incubators, but with a four day lag between the dust input and the peak in productivity. Although suggestive of a relationship, the results were too few and insufficiently clear-cut to be totally convincing. In addition, interpretation was complicated because productivity change was measured in a deck incubator, not in the ocean itself. Also, when deposition occurred, meteorological conditions changed, with greater stirring of near-surface water, which itself may have changed the productivity. However, this experiment represents a novel and potentially powerful tool since it uses the natural atmospheric input and examines the response of the real oceanic system.

A different approach to testing the iron hypothesis is that of adding inorganic iron (FeSO₄) directly to a small patch (of the order of 100 km²) of the oceans. In order to be able to track the iron enriched patch as it moves in the ocean, the gas sulphur hexafluoride is added along with the iron. Sulphur hexafluoride can be easily measured at tracer (femtomolar) concentrations in almost real time from the research vessel, enabling the enriched patch to be tracked. The principles underlying this approach are outlined in Watson et al. (1991a). It has been utilised three times to date; twice in the equatorial Pacific (IronEx I: Martin et al. (1994) and IronEx II: Coale et al. (1996)) and very recently in the southern oceans (SOIREE: Boyd et al. (2000)). On all three occasions, raising the iron level in the water by a few nanomoles per litre produced a significant enhancement in phytoplankton activity, as measured by chlorophyll concen-
<figure><img src="image_25.png" /><figcaption>**Fig. 2.21.**
Carbon dioxide changes inside and outside the enriched patch during the course of SOIREE (Boyd et al. 2000)</figcaption></figure>

<figure><img src="image_26.png" /><figcaption>**Fig. 2.22.** Depth-integrated time evolution of DMS and DMSP inside (open circles) and outside (closed circles) the enriched patch during the course of SOIREE (Boyd et al. 2000)</figcaption></figure>

<figure><img src="image_27.png" /><figcaption>**Fig. 2.23.** Compendium of results from ice cores for iron, CO₂, MSA, and several other parameters (Turner et al. 1996b)</figcaption></figure>

tration increase, consistent with the iron fertilisation hypothesis. In the case of IronEx II, the increase was at least an order of magnitude. Smaller organisms were the first to utilise the iron supplement, with the larger plankton (mainly diatoms) benefitting later.

Trace gases measured in these experiments were CO₂ and DMS. The former was drawn down due to the enhanced primary production. The extent of CO₂ removal roughly mirrors the increase in chlorophyll, except for IronEx I where it was very small, probably due to rapid recycling of the fixed carbon by grazers. For DMS, three-to five-fold increases occurred in all three studies, with much less variation than for CO₂. Carbon dioxide changes between inside and outside the enriched patch during the course of SOIREE are shown in Fig. 2.21, and the time evolution of DMS and its precursor DMSP integrated over a vertical column are shown in Fig. 2.22. Such a fertilisation experiment is akin to a batch culture perturbation and it is not clear whether long-term Fe enrichment and sustained higher productivity would lead to higher steady-state DMS concentrations. Given

enough time, shifts in the population structure, or simply growth of the microbial populations, might re establish low steady-state DMS concentrations with perhaps higher turnover of both DMSP and DMS. Our present understanding of the response of microbial populations to changes in DMSP and DMS supply is insufficient to make confident predictions in this regard. To put these results in a broader time context, a compendium of results from ice cores for iron, CO₂, MSA (an atmospheric oxidation product of DMS), and several other parameters is provided in Fig. 2.23. It is notable
worthy that the elevated iron and MSA and lowered CO₂ levels during the last glacial period are consistent with a scenario wherein ocean productivity was higher then due to enhanced atmospheric inputs of iron. For further discussion of the use of ice core records to examine the overall sulphur cycle see Sect. 2.3.

### 2.8.3.2 The Input of Atmospheric Nitrogen to the Ocean

#### 2.8.3.2.1 Coastal Regions

There is now widespread evidence that atmospheric fixed nitrogen compounds contribute to enrichment and in some areas probably to coastal and estuarine eutrophication (Jaworski et al. 1997; Howarth et al. 1996). Current estimates of the percentage of total (natural + anthropogenic) new N loading attributed to direct atmospheric deposition at a number of North American and European locations range from 5% to over 50% (Duce 1991; Fisher and Oppenheimer 1991; Valigura et al. 1996; Dennis 1997; Holland et al. 1999). Inputs of N to estuarine systems that result from direct atmospheric deposition by-pass much of the estuarine N "filter" (Kennedy 1983; Paerl 1995, 1997). Thus, atmospheric deposition assumes an increasingly important role as a new N source in lower estuarine and coastal waters below the biological N filtering zone (Fig. 2.24).

Dry and wet atmospheric deposition introduces into estuaries a variety of biologically available inorganic ($NO_3^-$, $NH_4^+$, DON) compounds, most of which result from human activities (Likens et al. 1974; Galloway et al. 1994). In addition, organic nitrogen (ON) comprises a significant fraction (from 15 to over 30%) of wet and dry atmospheric deposition in coastal watersheds (Correll and Ford 1982; Scudlark and Church 1993; Peierls and

Paerl 1997). Although the composition of atmospheric ON is poorly known, recent work (Peierls and Paerl 1997; Seitzinger and Sanders 1999) indicates that constituents of this pool are biologically utilised and, hence, should be included in eutrophication assessments.

In situ bioassays and field surveys show that enrichment with the major deposition constituents $NH_4^+$ and $NO_3^-$ at natural dilutions and atmospherically derived dissolved organic nitrogen (DON) results in enhanced phytoplankton primary production and increased biomass (Paerl 1985; Willey and Paerl 1993; Paerl and Fogel 1994; Peierls and Paerl 1997). Atmospheric DON may selectively stimulate growth of specific types of marine phytoplankton (Neilson and Lewin 1974; Antia et al. 1991). These selective phytoplankton responses to specific nitrogen inputs, and changes in stoichiometric C:N ratios resulting from these inputs may induce changes at the zooplankton, invertebrate, herbivorous fish, and higher trophic levels.

<figure><img src="image_28.png" /><figcaption>**Fig. 2.24.** Schematic representation of airshed, upper estuarine, and lower estuarine processes</figcaption></figure>

<figure><img src="image_29.png" /><figcaption>**Fig. 2.25.** Trends in annual atmospheric deposition (wet deposition as $NH_4^+$ and $NO_3^-$ expressed in $kg N ha^{-1} yr^{-1}$) collected during a 20 year period at the National Atmospheric Deposition Program (NADP) site NC-35 in Sampson County, North Carolina (data re-plotted from NADP information)</figcaption></figure>
Atmospheric nitrogen inputs have been examined recently in North Carolina's Albemarle-Pamlico Sound System (APSS) (Paerl 1995, 1997). Within the APSS, the Neuse River Estuary receives N inputs from a mosaic of upstream and upwind agricultural, urban, and industrial sources. Fossil fuel combustion and agricultural and industrial N emissions represent a significant and growing source of new N to this system (Paerl and Fogel 1994), reflecting a national and world-wide trend (Duce 1986; Luke and Dickerson 1987; Asman 1994; Paerl 1995; Holland et al. 1999). Depending on the relationship between watershed-estuary surface areas, degree of watershed N retention, seasonal rainfall, discharge and flow patterns, and proximity of atmospheric sources, an important fraction of nitrogen from atmospheric deposition is directly deposited on the estuary. In the case of the APSS, recent estimates are on the order of 20% (for its estuarine tributaries) to 40% (for the downstream waters of Pamlico Sound) (Paerl and Fogel 1994; Paerl 1995, 1997) of the N being directly deposited.

Atmospheric N generated from expanding intensive animal farming is of particular concern. Examination of the long-term record of atmospheric $NH_4^+$ and $NO_3^-$ deposition in Sampson County, eastern North Carolina, shows a nearly three-fold increase in annual $NH_4^+$ deposition (also relative to $NO_3^-$) since 1977, with a particularly precipitous rise since the late 1980s (Fig. 2.25). The reason for this may be that unlike human waste, swine waste is stored in open lagoons and remains largely untreated, and substantial amounts (30 to >80%) of N are lost via $NH_3$ volatilisation alone (O'Halloran 1993).

### 2.8.3.2.2 The Open Ocean

There is also growing concern about the increasing input of human-derived nitrogen compounds to the open ocean. This is especially important in parts of the open ocean where nitrogen is the nutrient that limits biological growth. This is the case in the nutrient-poor waters of the large central oceanic gyres in the North and South Pacific and Atlantic Oceans and the southern Indian Ocean. Current estimates suggest that, at present, atmospheric nitrogen accounts for only a few percent of the total new nitrogen delivered to surface waters in these regions, with upwelling from deep waters being the primary source of new surface nitrogen. It is recognised, however, that the atmospheric input to the ocean is highly episodic, often coming in large pulses extending over a few days. At such times, atmospheric input plays a much more important role as a source for nitrogen in surface waters. A recent estimate of the current input of fixed nitrogen to the global ocean from rivers, the atmosphere, and nitrogen fixation indicates that all three sources are important (Cornell et al. 1995). Paerl and Whitall (1999) estimate that 46-57% of the total human-mobilised nitrogen entering the North Atlantic Ocean is coming via the atmosphere.

In addition, the atmospheric organic nitrogen flux may be equal to or perhaps greater than the inorganic (i.e. ammonium and nitrate) nitrogen flux in open ocean regions. The source of the organic nitrogen is not known, but a large fraction of it is likely to be human-derived. This form of atmospheric nitrogen input to the open ocean had not been considered in detail until very recently.

Not only will the input of atmospheric fixed nitrogen to the open ocean increase significantly in the future as a result of increasing human activities, but the geographical locations of much of this input will probably change as well. Galloway et al. (1994, 1995) have evaluated pre-industrial nitrogen fixation (formation of the so-called reactive nitrogen) on the continents; the near-current (1990) reactive nitrogen generated from human activities such as energy production (primarily as nitrogen oxides), fertiliser use, and legume growth; and the estimated reactive nitrogen that will be produced in 2020 as a result of human activities.

The most highly developed regions in the world are predicted to show relatively little increase in the formation of reactive nitrogen, with none of these areas contributing more than a few per cent to the overall global increase. However, other areas will contribute very significantly to increased human-derived reactive nitrogen formation in 2020. For example, it is predicted that Asia will account for ~40% of the global increase in energy-derived reactive nitrogen, while Africa will have a six-fold increase accounting for 15% of the total global increase. It is predicted that production of reactive nitrogen from the use of fertilisers in Asia will account for ~87% of the global increase from this source. Both energy sources (nitrogen oxides and ultimately nitrate) and fertiliser (ammonia, urea) result in the extensive release of reactive nitrogen to the atmosphere. Thus, these predictions indicate very significant potential increases in the atmospheric deposition of nutrient nitrogen compounds to the ocean downwind of such regions as Asia, Central and South America, Africa, and the former Soviet Union (see Fig. 2.26).

The potential problem outlined above was highlighted by a computer modelling study undertaken by Galloway et al. (1994), who generated maps of the recent (1980) and expected (2020) annual deposition of reactive nitrogen compounds from the atmosphere to the global ocean. Figure 2.26 is a map of the projected ratio of the estimated deposition of oxidised forms of nitrogen in 2020 to the values for 1980. It appears that from one and a half to three, and in some limited areas up to four, times the 1980 rate will occur over large areas of the oceans. This increased nitrogen deposition will provide new sources of nutrient nitrogen to some regions of the ocean where biological production is currently limited by nitrogen. There is thus the possibility of important impacts on regional biological production and the marine carbon cycle in these regions of the open ocean.
Increase in reactive nitrogen deposition, 1980–2020

<figure><img src="image_30.png" /><figcaption>Fig. 2.26. Ratio of the estimated deposition of oxidised forms of nitrogen to ocean and land surfaces in 2020 relative to 1980 (adapted from Galloway et al. 1994, and Watson 1997)</figcaption></figure>

Debate continues about the relative importance of iron, nitrogen, or other compounds as prime determinants of oceanic phytoplankton productivity and, consequently, potential controls of marine gas emissions. On geological time scales, phosphorus is accepted to be the ultimate control. Silica, another component of the same Fe and N-containing dust but with a significantly longer residence time, has also been examined as an altering agent of the species composition of marine phytoplankton in oceanic regions, favouring siliceous organisms (e.g. diatoms) (e.g. Harrison 2000; Treguer and Pondaven 2000). Such organisms would differently affect total gas emissions but not total primary production. Such a silica hypothesis reinforces the link between marine biogeochemistry and resulting sea-air gas emissions.

## 2.9 Summary of Achievements and Remaining Research Challenges

Much progress has been achieved over the last decade through technological advances and appropriate scientific approaches. The advances include: remote sensing instrumentation to provide detailed spatial and temporal data; micrometeorological and isotopic techniques for estimating the flux of matter and energy within ecosystems and between ecosystems and the atmosphere, geosphere, and biosphere; techniques for manipulations

of local scale selected environmental factors; and, statistical and numerical modelling techniques capable of analysing multi-variate, nonlinear problems.

A developing system-based approach, including Lagrangian studies of air and water masses, comprising all components, e.g. soils, vegetation, and atmosphere, has led to an understanding of biogeochemical cycles of individual chemical compounds and interactions among chemical compounds. The campaign mode of carrying out field measurements has enhanced the understanding of the interconnectedness of systems and the importance of scaling issues.

Highlights of the research include:

* Reduced uncertainties in N₂O, NO, CH₄, DMS, and certain organohalogen emissions, and a better characterisation of local and regional distribution patterns of fluxes together with a mechanistic, but not necessarily integrated, understanding of the surface factors which control these emissions and exchange.

* Effective mitigation strategies have been developed for some CH₄ emissions and a better understanding of how land management practices influence N₂O and CH₄ emissions has been gained.

* Mechanisms and pathways of production and environmental controls have been identified for a large number of VOC compounds emitted from vegetation, including canopy transfer processes. Emission models estimating global emissions have been developed.
* VOC emissions can account for a loss of two to four percent of C taken up by photosynthesis, which has implications for understanding and quantifying the C cycle.

* Improved understanding of atmospheric input of inorganic N and Fe, mainly of anthropogenic and soil origin, into coastal and open ocean environments representing 5-70% of total input in the case of nitrogen and the subsequent impacts on the C uptake of the oceans and the C and S cycles.

* The acidic nature of wet deposition, which differs by source and region, has been characterised.

* Improved understanding of the partitioning of dry deposition (particularly O₃ and NO₂) on leaves and soil surfaces and related physiological mechanisms has been developed.

* Emission ratios for biomass burning are well described for savannas, but less well described for humid forests and biofuels. Broad databases are available of emission factors for a large number of substances.

To achieve more plausible and quantitatively reliable answers, several key issues remain:

* To investigate mechanisms (chemical and biological) responsible for trace gas cycling (emission and deposition) in oceans, soils, and plants, and to establish long-term sites/studies to provide that information, undertaking field experiments to determine, quantify, and discriminate among driving variables.

* To study the exchange of VOCs between vegetation, oceans, and the atmosphere along with the exchange of other trace gases.

* To understand and quantify the effect of soil-released NO and its oxidation product NO₂, under different management practices, to the atmosphere including interaction of these gases both within and above the canopy.

* To determine whether changes in the marine emissions of trace gases and particles are likely to have a significant influence on atmospheric chemistry and vice-versa, resulting from climatic (e.g. rainfall, temperature; perhaps small), elevated CO₂ (perhaps large), and/or land use changes. Key areas may include the greenhouse effect (tropospheric O₃), stratospheric ozone (CH₃Br), radiation and clouds (DMS), VOCs, tropospheric chemistry (dust-FE-DMS-CO₂), and other unpredicted impacts (e.g. the change in marine phytoplankton communities coupled with changes in N and Fe deposition), especially in the Tropics and high latitudes.

* To understand how the hydrological cycle will be affected in various regions with climate change and the subsequent impacts on emissions.

* To improve the parameterisation of air-sea exchange and its links to biogeochemical cycling in surface waters as well as improve Lagrangian studies in water, air, and the combined ocean-atmospheric front, including international participation in order to overcome the intrinsic organisational and logistical difficulties.

* To promote the establishment, wherever possible, of long-term sites for flux measurements, to investigate the magnitude of interannual variation and thus achieve more robust estimates of mean annual fluxes and global budgets.

* To design experiments that will bring synthesis from emission-type studies, regional means of fire detection and prediction, spatially and temporally resolved, and chemical transport models (see Chap. 6) in order to determine the impact of burning on atmospheric chemistry.

* To develop more realistic biological and deposition process-oriented models with interaction and feedback among process-oriented, regional models and global models in order to provide improved estimates of emission and deposition fluxes.