Review pubs.acs.org/CR

Chemistry and Release of Gases from the Surface Ocean Lucy J. Carpenter*,† and Philip D. Nightingale‡,§ †

Wolfson Atmospheric Chemistry Laboratories, Department of Chemistry, University of York, York YO10 5DD, United Kingdom Plymouth Marine Laboratory, Prospect Place, The Hoe, Plymouth PL1 3DH, United Kingdom



Author Information Corresponding Author Notes Biographies Acknowledgments References

1. INTRODUCTION The sea-surface layer is the very upper part of the sea surface where reduced mixing leads to strong gradients in physical, chemical, and biological properties.1 This surface layer is naturally reactive, containing a complex chemistry of inorganic components and dissolved organic matter (DOM), the latter including amino acids, proteins, fatty acids, carbohydrates, and humic-type components,2 with a high proportion of functional groups such as carbonyls, carboxylic acids, and aromatic moieties.3 The different physical and chemical properties of the surface of the ocean compared with bulk seawater, and its function as a gateway for molecules to enter the atmosphere or ocean phase, make this an interesting and important region for study. A number of chemical reactions are believed to occur on and in the surface ocean; these may be important or even dominant sources or sinks of climatically active marine trace gases. However, the sea surface, especially the top 1 μm to 1 mm known as the sea-surface microlayer (ssm), is critically undersampled, so to date much of the evidence for such chemistry comes from laboratory and/or modeling studies. This review discusses the chemical and physical structure of the sea surface and mechanisms for gas transfer across it, and explains the current understanding of trace gas formation at this critical interface between the ocean and atmosphere.

CONTENTS 1. Introduction 1.1. Marine Dissolved Organic Matter (DOM) and Surface-Active DOM 1.2. Photochemical Cycling of Organic and Inorganic Material in the Surface Ocean 1.3. Physical Structure of the Sea Surface 1.4. Mixing in the Microlayer and Gas Transfer 1.5. Surfactants and Air−Sea Gas Transfer 1.6. Sampling the Sea Surface 2. Trace Gas Production in the Sea-Surface Layer 2.1. Halogenated Compounds 2.1.1. Ozonolysis of Halides at the Sea Surface 2.1.2. Photosensitized Mechanisms 2.1.3. Oxidation of Halides by •OH 2.1.4. Halogenation of Dissolved Organic Matter (DOM) by Radical and Nonradical RHS 2.2. Oxidized Nitrogen Compounds 2.2.1. Nitrogen Oxides and Alkyl Nitrates (RONO2) from Solar Photolysis of Nitrite 2.2.2. Nitryl Chloride ClNO2 2.3. Oxygenated Volatile Organic Compounds (OVOCs) 2.3.1. Acetone 2.3.2. Acetaldehyde 2.3.3. Methanol 2.3.4. Glyoxal 2.4. Inorganic Carbon 2.4.1. Carbon Monoxide (CO) 2.4.2. Carbon Dioxide (CO2) and Dissolved Inorganic Carbon (DIC) 3. Summary and Perspective Appendix 1: Uptake of Gas-Phase Molecule to Liquid Surfaces: Bulk-Phase and Surface-Phase (Langmuir−Hinshelwood) Reaction Mechanisms Appendix 2: The Haloform Reaction Appendix 3: Photochemistry of CDOM, Nitrite, and Nitrate © XXXX American Chemical Society

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1.1. Marine Dissolved Organic Matter (DOM) and Surface-Active DOM

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Marine DOM influences the production of most marine volatile trace gases. It is primarily sourced from marine biota, particularly photosynthetic algae and bacteria,4 and released during phytoplankton growth, as a consequence of grazing by predators and during viral lysis of cells.5 Transfer of terrestrial organic carbon, e.g., lignins, from plant litter, biomass, and soil organic matter appears to make a small contribution to the ocean pool.6 DOM is present at low concentrations in the oceans (0.5− 1.0 mg L−1) and is operationally distinct from particulate organic matter (POM) as the fraction passing through a filter

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Figure 1. Major known components of marine dissolved organic matter (DOM). The bottom trace shows a high-resolution FT-ICR mass spectrum (negative mode) of seawater SPE-DOM collected in the Greenland Sea (Rosie Chance, University of York).

pore, typically 0.2−1 μm in size.7 Marine DOM composition is a heterogeneous mixture of low-molecular weight (1000 Da) biomolecules, and microgels (Figure 1).8 The low-molecular weight (LMW-DOM) fraction represents the major size fraction (65−80% of bulk DOC), and the high-molecular weight (HMW-DOM) fraction is more abundant at the surface (35−40%) compared with deep water (20−30%).9 Together this suggests that the LMW-DOM fraction is more resistant to microbial oxidation. Mass spectrometric investigations show that the majority of marine DOM is in the mass range 200−1000 Da. The major biochemical components of marine DOM are proteins (25−50%), lipids (5−25%), and carbohydrates (up to 40%) (see Figure 1).10 However, most DOM in seawater is uncharacterized at the molecular level,11 and the total contribution of these components may only reach 10−30% of bulk DOM.7,9 Polysaccharides (CHO) are the most abundant measurable component. About 7−20% of CHO has been identified as neutral sugars.12 The lipid component of marine DOM is predominantly from saturated fatty acids (FA),13 although polyunsaturated fatty acids (PUFA) can comprise a significant fraction.14 Of the proteinaceous material, only a small proportion derives from free amino acids (typical concentration in the low nanomolar range), which are highly reactive.15 Humic substances (HS), which are composed of the three operationally defined components fulvic acids, humic acids, and humins, represent a portion of the molecularly uncharacterized component of marine DOM, but comprise only a small fraction (0.7−2.4%) of DOM in the ocean.16

Oceanic humic acids are widely believed to derive predominantly from in situ degradation products of marine plankton, particularly unsaturated FA, as opposed to transfer from continents.17 Measurements of DOM composition are generally selectively defined toward a particular chemical and physical class of DOM, dependent on the sampling, isolation, and detection method. For example ultra- or molecular filtration separates according to hydrodynamic diameter and concentrates hydrophilic HMW-DOM. Mass spectrometric studies are usually preceded by separation of DOM from salt using solid phase extraction (SPE) which selectively concentrates hydrophobic to moderately polar compounds depending on the choice of solid phase. While high-resolution mass spectrometers are capable of molecular characterization at a high level of confidence, matrix effects coupled with the selectivity of SPE-DOM mean that they are most likely to capture LMW-humic (HS) DOM. Indeed, the range of H/C and O/C ratios determined by SPEDOM-MS falls outside the elemental ratios of most common proteins, carbohydrates, and lipids.7 Soluble but surface-active DOM samples (including carbohydrates, lipids, polysaccharides, amino acids, and proteins) are of particular interest in this review due to their ability to produce atmospherically relevant gases and organic particles, and to reduce air−sea transfer rates. Surface-active agents, or surfactants, are compounds that are capable of reducing the surface tension of the liquid in which they are dissolved. These molecules are typically amphiphilic, i.e., have a hydrophilic (polar) water-soluble head and a hydrophobic B

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Figure 2. Conceptual model of the sea-surface layer (modified from Hardy39). Note the streak of the surfactant oleyl alcohol (running from left to right toward the top of the ocean surface) formed during an air−sea gas transfer experiment in the Atlantic.40

seawater and is significantly enriched at the surface compared to levels in subsurface seawater,24a,25 due to many of its individual components being surface-active to varying extents.

(nonpolar) tail. The definition of surface-active agents also includes sparingly soluble substances that lower the surface tension of seawater by spreading spontaneously over its surface, henceforth referred to as hydrophobic or insoluble surfactants. The surface tension of clean (i.e., organic- and particulate-free) seawater varies with temperature and, to a lesser extent, salinity and is typically between 73 and 75 mN m−1.18 Experiments with surface films from coastal waters have found the reduction in surface tension, as compared to clean seawater, can be as high as 10−15 mN m−1, whereas reductions using films collected from the open oceans are typically only 0.5−1 mN m−1.19 Although the impact of natural surfactants on surface tension is often small, they have a much greater effect on air− sea gas transfer rates, the latter being reduced by even slight surface accumulation of organic material.20 The mechanisms responsible for the sensitivity of gas transfer to surfactants are discussed in section 1.5. The primary sources of natural surfactants are probably phytoplankton exudates,21 that are transported to the surface via turbulent mixing, diffusion, and rising bubbles coated with surface active organic matter.22 Surfactants are enriched at the sea surface, typically by up to a factor of 3, in many areas of the world’s oceans.23 Although concentrations of organics are often highest in shelf seas, and coastal and near-shore areas, enrichments at the surface compared to bulk seawater appear to be greatest in the oligotrophic oceans rather than in more productive regions.23b Further, enrichments are often greater at higher wind speeds.23b,24 Collectively, the light-absorbing fraction of DOM is referred to as chromophoric dissolved organic matter (CDOM). CDOM is the dominant absorber of UV and visible light in

1.2. Photochemical Cycling of Organic and Inorganic Material in the Surface Ocean

While consumption by heterotrophic bacteria is the main removal pathway of DOM, photochemical oxidation to carbon dioxide,26 carbon monoxide,27 and oxygenated volatile organic compounds (OVOCs)28 also occurs in or on the surface ocean. Photochemical oxidation of DOM thus represents an important route for release of trace gases,29 and is discussed in more detail in later sections. It can also represent a significant component of the ocean carbon cycle.30 For example, methanol and acetaldehyde are used as sources of energy and carbon by marine microbes; thus, their production has important implications for marine productivity and hence carbon cycling.28c,31 Components of DOM with carbon−carbon double bonds such as PUFA can react rapidly with ozone in heterogeneous reactions at the ocean surface to produce glyoxal (CHOCHO) and other OVOCs.28d,32 Pure water is relatively transparent to light between the nearultraviolet and near-infrared components of the spectrum. Dissolved salts in seawater have little effect on light absorption at visible wavelengths but increase absorption slightly in the ultraviolet.33 The downward attenuation of irradiance is primarily controlled by concentration of light-absorbing DOM in the UV and by a combination of DOM and chlorophyll (and other pigments) contained within phytoplankton cells in visible light. In very clear oligotrophic waters, absorption of light by viruses, colloids, and bacteria may C

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transfer through the air-side film. Although conceptually useful, this model has been shown to be too simple to represent gas transfer in laboratory and field experiments.41c Rigid boundary or solid wall models apply a velocity profile to either side of the interface to describe a smooth transition between diffusive and turbulent regimes.42 In such models, turbulence plays a modest role in controlling the rate of transfer. The predicted transfer velocity scales with D2/3, demonstrating that diffusivity becomes less important in determining gas transfer at higher winds. The rigid boundary models are found to work well in tank and wind tunnel experiments when the surface is smooth but underestimate gas transfer rates in the presence of even small waves.43 Surface renewal models of the water-side of the interface improve on this by periodically and instantaneously replacing the stagnant film with material from the bulk.44 These models predict that gas transfer scales with D1/2 as has been observed in laboratory experiments with heat and with gases and at sea.45 Similarly, the so-called large and small eddy models,46 in which transfer in the near-surface water is described in terms of a series of cells of rotating fluid, also lead to the conclusion that gas transfer is proportional to D1/2. Finally, a surface penetration model, which considers incomplete replacement of the surface thin layer by eddy transport from the bulk,47 has been applied to explain the apparent discrepancy between transfer velocities of heat and mass from observations compared with predictions from the renewal models.48 Their findings imply that the diffusivity dependence of transfer velocity is not independent of turbulent forcing, but molecular diffusion still plays a controlling role in air−sea gas transfer.

become significant.33 The maximum depth of the 10% irradiance level for ultraviolet light is predicted to be in the oligotrophic ocean gyres (18, 32, 44, and 70 m for wavelengths at 305, 325, 340, and 380 nm, respectively) and at a minimum (0−5m) for all wavelengths in waters on the continental shelves, upwellings, and coastal regions.34 The wavelength of maximum transmittance shifts from blue in the oligotrophic oceans, to green in more productive waters, and to yellow in highly turbid near-shore waters.33 Photolysis and chemical oxidation of inorganic compounds, including halides, nitrates, and nitrites, also play important roles in ocean−atmospheric exchange of key constituents of atmospheric composition. For example, ozonolysis of surface iodide produces photochemically active iodine compounds, which may then participate in catalytic ozone destruction cycles in the troposphere.35 Given the reactivity of many chemicals in the surface ocean with ozone, it is not surprising that models of tropospheric ozone deposition that ignore these reactions in the interface region underestimate observed O3 deposition velocities.36 Inorganic and organic ocean chemistry can be strongly coupled; for example, CDOM has also been shown to facilitate photosensitized chemical oxidation of halides to volatile halogen compounds,37 and reactions of the photochemical products of CDOM and nitrite are believed to be a major routes to oceanic alkyl nitrate production.38 These reactions are discussed in more detail in section 2. 1.3. Physical Structure of the Sea Surface

The sea surface can be thought of as a series of sublayers varying in thickness from nanometers (for insoluble surfactants) to micrometers (for gases, bacterioneuston and phytoneuston) and millimeters for zoonueston.1 Note that the term “neuston” simply refers to organisms that inhabit the sea-surface layer. Of particular relevance to this study is the sea surface microlayer (ssm), usually defined operationally and typically covering the top 1 μm to 1 mm of the sea surface and the very surface of the microlayer (termed the nanolayer in Figure 2). Both theoretical and experimental studies have shown that the upper part of the ssm is a region where turbulence is dramatically reduced and molecular diffusion dominates resulting in strong gradients in heat, pH, and gas concentrations.41 Indeed, much of our information on microlayers has been derived from modeling, laboratory, and field investigations into the transfer of gases (and heat) across the air−sea interface.

1.5. Surfactants and Air−Sea Gas Transfer

Films occur when there is an excess of surface-active material at the air−sea interface. Such films may be visible or invisible. Hydrophobic surfactants, either anthropogenic or biogenic in origin, are likely to be present as either a condensed film or slick on the sea surface and are represented in Figure 2 by the nanolayer. These monolayers are typically 2−3 nm thick49 and can inhibit air−sea gas transfer by acting as a physical barrier to exchange,20,50 or by providing an additional liquid phase that may enhance resistance to transfer.51 These impacts are known as “static” effects and are thought to be important only at low wind speeds as slicks are easily dispersed by wind and waves.41c The main effect of surface-active material on air−sea gas transfer is believed to be due to the presence of hydrophilic surfactants. Unlike for hydrophobic surfactants, static effects are relatively unimportant as soluble surfactants films are too permeable to offer significant liquid phase resistance.50 Instead, hydrophilic surfactants are thought to impact gas transfer via hydrodynamic effects.41c In particular, soluble surfactants alter the viscoelastic modulus of the sea surface (i.e., the resistance of the surface to a change in surface area). Impacts of changes in the viscoelastic properties of the sea surface include a reduction in near-surface turbulence length and velocity scales, the inhibition of wave growth, and an enhancement to wave energy dissipation (wave damping).20,52 The last effect has been the most studied and is thought to be due to gravity and capillary waves creating transient surface tension gradients during their compression and dilation of the sea surface.19c,20,53,54 Even at the low surface film pressures typical of the open ocean, surface viscoelastic effects are thought to be sufficient to significantly reduce gas transfer rates.20,55

1.4. Mixing in the Microlayer and Gas Transfer

The simplest model of the role of the sea surface in air−sea gas transfer is the stagnant film (or thin film) model.41f This model assumes the sea surface is flat with stagnant mass boundary layers on either side of the interface, through which molecular diffusion is the sole transport process, while turbulent mixing dominates in the bulk fluids. The aqueous thin film is part of, not separate to, the ssm. In this model, gas transfer scales with molecular diffusion (D). The flux of gas across the interface is also driven by the concentration gradient across the two thin films. The depth of the water-side thin film varies with wind speed, or turbulent forcing, but is typically around 40 μm41f and hence represents the upper part of the ssm as depicted in Figure 2. The model can allow for chemical reactivity of the gas of interest within the film. Gases that are relatively insoluble (e.g., CO2) are limited by transfer through the aqueous thin film, while highly soluble gases (e.g., OVOCs) are limited by D

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previously for estimating ssm concentrations of DMS,63a,69 and OVOCs,28b,70 and significant enrichments were observed.

Surfactants have long been known to reduce gas transfer in laboratory experiments,45a,56 but it has proven very difficult to test their impact in the field.57 An in situ study in the Pacific Ocean found that the flux of heat was overpredicted by wind speed based relationships when winds were below 6 m s−1 and when levels of CDOM were enhanced in the ssm relative to bulk water.25a Although surfactants are commonly thought to only inhibit surface turbulence and hence gas transfer at relatively low winds, laboratory experiments have shown that soluble surfactants can inhibit gas transfer even under breakingwave conditions.58 Further evidence comes from a large-scale deliberate release of oleyl alcohol (an artificial surfactant) in the Atlantic Ocean that reduced gas transfer rates, determined via two independent techniques (dual tracers and direct covariance) by about 50% at 7 m s−1 and by about 25% at 11 m s−1.40 While this study has proved that surfactants can significantly modify air−sea gas transfer in the ocean, it is not yet clear that they do so at ambient levels. However, many of the world’s oceans have been found to be covered to a significant extent by surfactants, with ssm enrichments persisting at wind speeds of at least up to 10 m s−1.23b,24

Table 1. Comparison of DMS and SF6 Concentrations in Bulk Seawater with Sea-Surface Microlayer Samples Collected Using Various Techniquesa DMS (nM) SF6 (pM) a

bulk seawater

glass plate

Garrett screen

rotating drum

7.1 22

1.6 ± 0.4 3.5 ± 0.6

2.0 ± 0.3 2.5 ± 0.6

4.6 ± 0.5 14 ± 1.4

N = 5, ±1 std dev. Data from refs 40, 64, 67.

Both gases were highly supersaturated in bulk seawater compared to the marine atmosphere, but concentrations immediately below the air−sea interface (i.e., in the nanolayer) will have been close to equilibrium with the atmosphere (i.e., close to zero for DMS and ∼2 fM for SF6). Hence, there will have been strong gradients in concentrations of both SF6 and DMS within the ssm. The stagnant film model (see section 1.3) would predict that mean concentrations of both gases in the upper ssm, under the conditions outlined above, would be about 50% of bulk water values. However, given the likelihood of some turbulent mixing occurring within the ssm, and the possible entrainment of bulk water during sampling, the expected mean concentrations of both gases are likely to be higher. Samples obtained using the rotating drum were 65% and 64% of the concentrations of DMS and SF6 in bulk water and are in line with the predictions above. There was no evidence of any enrichment in DMS concentrations compared to the inert tracer. Samples collected using the glass slide and Garrett screen were just 23% and 28%, respectively, of the bulk concentration of DMS, while SF6 was reduced by an even greater extent with values only 16% and 11% of those in bulk water. Larger losses with the glass plate compared to the Garrett screen have previously been observed for DMS,63a but our differences between these two techniques are not significant. Both the Garrett screen and glass plate show strong evidence of losses to the atmosphere during sample collection and appear to be of limited use for ssm measurement of volatile species. Note that for a gas strongly undersaturated in bulk water compared to the atmosphere, there would have been a similarly large transfer of gas from the atmosphere into the sample during collection.

1.6. Sampling the Sea Surface

Common ssm sampling devices include the mesh screen,59 the glass plate,60 and the rotating drum sampler or surface skimmer.61 These approaches have sampling depths of typically 200−440 μm,25b,62 20−150 μm,23b,63 and 50 μm,64 respectively. Other techniques include the use of polycarbonate and PTFE membrane filters to examine the diversity and activity of the bacterioneuston and these are thought to sample just the top 40 and 6 μm, respectively.65 The sampling depths of both the Garret Screen and glass plate are probably deeper than the depth at which gradients in gas concentrations due to molecular diffusion in the upper ssm would be expected (10−100 μm) indicating that bulk water is likely to be entrained into microlayer samples. Sampling for dissolved gases in the ssm is particularly challenging given the high potential for enhanced gas exchange between the microlayer and atmosphere during sample collection, which would act to reduce any concentration difference between the aqueous and air phases. For example, for a gas transfer velocity of 5 cm h−1, i.e., equivalent to that predicted for dimethyl sulfide (DMS) at a windspeed of 5 m s−1,57 the residence time of DMS within a 50 μm thin film would be less than 4 s. One approach to the sampling of trace gases in the ssm has been to use a cryogenic technique that freezes the sea surface and collects it as an ice layer,66 although the impact of liquid nitrogen on microbial communities and hence cycling of biogenic gases might be an issue. The effectiveness of three commonly used sampling devices was evaluated during a gas exchange experiment in the Atlantic.40,64,67 Surface seawater was pumped into a large ondeck container, spiked with the inert tracer sulfur hexafluoride (SF6), and then left for approximately 30 min for a microlayer to form. Marine microlayers in the laboratory and in the field have been previously shown to re-form within seconds.23a,68 Replicate microlayer samples were then collected using a Garrett screen, glass plate, and rotating drum, analyzed for SF6 and DMS, and compared to near surface bulk water values collected from a few centimeters below the interface (see Table 1). The Garrett screen and glass plate have been used

2. TRACE GAS PRODUCTION IN THE SEA-SURFACE LAYER 2.1. Halogenated Compounds

The surface ocean offers a very rich potential for chemical production of volatile halogenated species. Halide anions (X− where X = Cl, Br, or I) are highly abundant in surface seawater ([Cl−] 5.6 M; [Br−] 8.6 × 10−4 M; [I−] 1−20 × 10−8 M with a maximum in the tropics), and may be further enhanced in the nanolayer (Figure 2) close to the air−water interface.71 A number of halide oxidation mechanisms have been identified that could operate in or on the ocean surface and produce radical and nonradical reactive halogen species (RHS) including X•, X2•−, XY•−, X2, XY, and HOX (Figure 3). RHS may be released directly to the overlying atmosphere or react first with dissolved organic material in the ssm to produce organic halogens that may themselves be transferred across the air−water interface. In the atmosphere, these compounds E

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Figure 3. Chemical mechanisms for production of halogenated species in surface seawater.

Figure 4. Kinetics of surface and bulk liquid reactions. Definition of terms is given in Appendix 1.

H+(aq) reaction releases volatile halogens77 and occurs readily on aerosol containing sufficient acidity and halide concentrations such as sea spray aerosol.72 However, at the air− seawater interface, ozonolysis appears to be the primary chemical (dark) mechanism for brominated and iodinated RHS release. Hunt and co-workers78 proposed an interface reaction between Br− and O3 to explain measured Br2 production from the reaction of gas-phase O3 with aqueous sodium bromide particles. The measured rates were an order of magnitude higher than could be explained from known gasphase and aqueous chemistry alone. Clifford and Donaldson79 measured a pH increase at the air−water interface during the

initiate the destruction of tropospheric ozone (Br and I compounds),72 oxidize hydrocarbons (as observed for Cl),73 and influence aerosol formation.74 Below, we discuss the different halide oxidation mechanisms, and then evidence for reactions of RHS with organic material producing organic halogens (org-X). 2.1.1. Ozonolysis of Halides at the Sea Surface. Heterogeneous reactions on salty ice and snow and on aerosol surfaces release a wide range of halogenated species and are believed to be the dominant sources of bromine and chlorine to the polar and marine boundary layer.72,75 On ice, surface-phase ionic processes (such as the dissociation of HCl) and solvation processes are crucial for such chemistry to occur.75b,76 On liquid surfaces, the well-established HOX(g) + X−(aq) + F

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iodide activity being important under the experimental conditions of these studies. A coupled atmosphere−surface time dependent box model, which assumed bulk kinetics with reaction 6 occurring over the reactodiffusive length of the aqueous surface, was used by Carpenter and co-workers35 to extrapolate laboratory measurements to the ocean−atmosphere interface. This study concluded that the sea surface iodide−O3 reaction is the dominant contributor to global iodine emissions. Emissions of HOI as well as I2 were found to be important, despite the low Henry’s Law coefficient of HOI, with calculated typical fluxes of 100−250 nmol m−2 d−1 HOI and 2−10 nmol m−2 d−1 I2 in the clean marine boundary layer.35 Globally, these emissions together strongly outweigh those from organic iodine compounds.82 A parametrization of oceanic HOI and I2 emissions developed from this work yields a good match between modeled and measured atmospheric iodine oxide (IO) levels except at low wind speeds.35,83 This could be due to the model underestimating mixing of interfacial iodine through the ssm to bulk waters, which reduces the quantity volatilized from the surface, under low wind speed conditions. The presence of organic films on the surface of the ocean has also been shown to inhibit the transfer of I2 to the air.81d,84 Potentially relevant to these surface reactions is that both experimental and theoretical studies have established that heavy halides are enhanced at the water−air interface with enhancements in the order I− > Br− > Cl−.71a,c,85 The surface preference of these large anions has been explained by anion polarizability,86 and by more favorable water−water interaction energies when the ions are partially desolvated compared to a fully solvated ion.85b These enhancements are restricted to the outermost interfacial layer, i.e., the nanolayer in Figure 2, with depletions in the layer immediately below. Thus, in terms of I− interaction with O3, which appears to occur over the reactodiffusive length of the first few micrometers of the sea surface, it is likely that the bulk I− concentration is most relevant. In the case of a Langmiur−Hinshelwood mechanism, e.g., for Br−, then surface enhancement is likely to be important. The second-order rate coefficient between chloride and ozone is very low (kbulk,Cl‑ ∼ 3 × 10−3 M−1 s−1),87 and this reaction is not expected to be important in seawater. 2.1.2. Photosensitized Mechanisms. Another potential route to direct ocean surface production of small halogen molecules is via oxidation of halogen anions to their radical forms by photosensitizers, a type 1 reaction (see Appendix 3). Reeser and co-workers88 found that visible illumination (up to 700 nm) of chlorophyll produces an excited triplet state (see Appendix 3) that can decay to chlorophyll cations; these can subsequently oxidize halide anions to halogen atoms. The oxidation is enhanced in the presence of atmospheric O3, which acts as an electron acceptor, thus promoting the cationic form of the photosensitizer. Although chlorophyll itself is unlikely to be freely present in seawater to act as a specific photosensitizer, other photosensitizers (e.g., carbonyl containing compounds) are abundant at the sea surface. Jammoul and co-workers37b used benzophenone (Ph2CO) as a proxy for CDOM and suggested that photoexcited Ph2CO* forms a contact complex with a halide ion, then an ion pair via intramolecular electron transfer:

reaction on aqueous bromide, consistent with the proposed reaction mechanism:80 O3 + Br − → OBr − + O2

(R1)

OBr − + H+ ↔ HOBr

(R2)

HOBr + Br − → Br2 + OH−

(R3)

overall: O3 + 2Br − + H+ → Br2 + O2 + OH−

(R4)

79

Clifford and Donaldson found that the kinetics of the reaction, as determined by the dependence of the rate of pH change upon bulk Br− concentration, exhibited behavior quantitatively consistent with adsorption of O3(g) at the surface interface, i.e., following a Langmuir−Hinshelwood reaction mechanism (Figure 4 and Appendix 1). Oldridge and Abbatt80b studied the Br2 formed from interaction of O3(g) with both frozen and liquid solutions of Br− and found, for both types of solution, behavior consistent with a bulk-phase reaction (Appendix 1) and a surface-phase reaction operating simultaneously. The dependence of the reactive uptake coefficient for ozone on ozone concentration indicated that surface-phase kinetics are relatively more important at low ozone levels, but that bulk-phase kinetics dominate at high ozone concentrations. At low ozone concentrations, the reactive uptake coefficient for ozone increases with decreasing ozone concentrations, as in the Langmuir−Hinshelwood mechanism, since the surface is unsaturated with adsorbing ozone molecules. At high ozone concentrations the surface becomes saturated; thus, surface-phase kinetics are less important than bulk phase kinetics. These studies confirm that for liquid bromide-containing particles in the atmosphere under ambient conditions, surfacephase kinetics are significantly faster than bulk phase. Hunt and co-workers78 estimate that several ppt of bromine could potentially be produced in the night-time marine boundary layer from this aerosol surface chemistry. Further work is required to estimate the impact of the analogous reaction at the air−seawater interface. Clifford and Donaldson79 suggest that the presence of organic material in the nanolayer will enhance the Br− and O3 surface reaction by increasing the ozone residence time at the interface. The corresponding reaction between aqueous iodide and gaseous ozone is more than 6 orders of magnitude faster than the ozone−bromide bulk reaction at around 1−2 × 109 M−1 s−1. The reaction is well-studied due to its relevance to iodine biogeochemical cycling and tropospheric O3.35,81 Various laboratory studies have established the following basic mechanism from measurements of ozone uptake and I2 emission: H+ + I− + O3 → HOI + O2

(R5)

H+ + HOI + I− ⇋ I 2 + H 2O

(R6)

Both reactions 5 and 6 occur in several steps.18,61 Unusually for ozone reactions on aqueous surfaces (Appendix 1), the reaction kinetics of I− observed by many investigators are consistent with bulk phase kinetics.35,81a,c,e This could be due to the strong reactivity of iodide to ozone, such that interfacial reactivity of iodide is unimportant and chemistry occurring in the first few microns of the ssm dominates.35,80b,81e Conversely, some studies have observed Langmuir−Hinshelwood kinetics,81d,f possibly due to surface

Ph 2CO* + X− ↔ Ph 2CO*··· X− ↔ [Ph 2C• − O− ···X•]* (R7) G

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Chemical Reviews

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competes over disproportionation of Br2•− into bromide and bromine, as well as with its reaction with nitrite radicals. Nonradical RHS (e.g., X2, HOX) species are widely known to halogenate a variety of organic compounds including aromatics via electrophilic substitution (which occurs in the biological synthesis of naturally occurring compounds by marine organisms)96 and alkenes and alkynes via addition. In the marine environment, major nonradical processes for production of small volatile organohalogens are halogenation of phenolic compounds,97 and of methyl ketones via the wellknown haloform reaction (Appndix 2). An early demonstration of the importance of the haloform reaction was by Theiler and co-workers98 who showed that simple brominated hydrocarbons (e.g., CH2Br2) are formed by marine algae via enzymecatalyzed bromination of ketones. The likely reaction steps are the production of HOBr, released from extracellular bromoperoxidases (BrPO), followed by reaction with organic nucleophiles in seawater such as ketones.99 The same reactions are possible for HOI and perhaps HOCl if the peroxidase enzyme is strong enough to oxidize Cl−. The haloform mechanism was postulated to be responsible for the formation of a range of volatile organoiodine compounds including diiodomethane (CH2I2), dichloroiodomethane (CHICl2), and iodoform (CHI3) during ozonolysis of natural seawater, which produces HOI and/or I2.100 Bichsel and von Gunten97 studied the reactions of HOI with a range of substituted phenols as well as with carbonyls. At pH 8, rates of HOI loss were much greater for phenols than for carbonyls (per C atom: phenols 30−6 × 105 M−1 s−1, carbonyl compounds 6 × 10−9 to 5 × 10−7 M−1 s−1). Yields of CHI3 were at least an order of magnitude greater for phenols (e.g., resorcinol, benzene-1,3-diol) than for α-carbonyls. Concentrations of these compounds vary widely in the marine environment, but it seems likely that phenolic groups in the DOM pool are as, or more, important as carbonyls in halogenation. Several laboratory investigations have utilized organic probe compounds (e.g., bisphenol A, phenol, salicylic acid, allyl alcohol, and dimethyl-1H-pyrazole (DMPZ)) to confirm the presence of halogenating activity in illuminated natural or artificial seawater.101 These studies provide evidence for sunlight-driven DOM halogenation in seawater. However, as yet it is not clear which mechanism/s are predominantly responsible for photoinduced production of RHS in ambient seawater, nor the contribution of oceanic photoproduced RHS to the reactive halogen pool in the atmosphere. A large fraction of org-X produced by such mechanisms is indeed likely to be nonvolatile (e.g., halogenated phenols), although, from the discussion above, it is clear that volatile forms are also produced and are released to the atmosphere. Since photoproduced RHS are not only produced by DOM (3CDOM*) but also scavenged rapidly by DOM,101e it will be important to carry out studies using natural seawater to ascertain the balance of their production and destruction.

The formation of the charge transfer complex is calculated to be thermodynamically favorable for I− and Br− and unfavorable (with free energy of the electron transfer ∼0) for Cl−.37b Once formed, the charge transfer complex may undergo charge separation, leading to formation of a halogen radical (reaction 2), or return to the initial state. [Ph 2C• − O− ···X•]* ↔ Ph 2C• − O− + X•

(R8)

Similar reactions are expected to occur between halides and the triplet states of naturally occurring CDOM (3CDOM*). Further products of the reaction include X2•−, formed from reaction of X• with X−. De Laurentiis and co-workers89 utilized the excited triplet state of anthraquinone-2-sulfonate (AQ2S) as a proxy for 3CDOM* to oxidize Br− to Br•, leading to the formation of Br2•−. The authors suggested that, in seawater, the oxidation of bromide by 3CDOM* to Br2•− would be faster than oxidation by •OH. 2.1.3. Oxidation of Halides by •OH. The OH radical is the most reactive transient species in seawater.90 Solar photolysis of CDOM, dissolved nitrate (NO3−), and nitrite (NO2−) are believed to be the primary sources (see Appendix 3), and Br− the primary sink (∼90%), the latter forming Br with a second-order rate constant of 3 × 109 M−1 s−1.91 Flash photolysis studies of Zafiriou91a confirmed that •OH in seawater is almost quantitatively converted to Br•, which then is quenched by reaction with Br− to Br2•−. OH + X− → HOX−

(R9)

HOBr − + H+X + H 2O

(R10)

X− + X − ↔ X 2− −



(R11) −

HOX + X ↔ X 2 + OH



Using a chemical probe technique, Mopper and Zhou90 determined levels of •OH in seawater of 1−12 × 10−18 M. Thus, the lifetime of Br− against •OH oxidation is around 1−10 years. De Laurentiis and co-workers89 argue that the formation rate of 3CDOM* is much higher compared to that of •OH in most surface waters and would provide a large 3CDOM* reservoir for bromide to react with. 2.1.4. Halogenation of Dissolved Organic Matter (DOM) by Radical and Nonradical RHS. As discussed above, oxidation of halides by •OH or 3CDOM* can lead to radical RHS, predominantly X• and X2•−. These radicals can add to unsaturated carbon−carbon bonds and/or recombine with carbon-centered radicals.88,92 The formation of CH3I at the ocean surface is believed to proceed in part through recombination of I• with methyl radicals produced from photodecomposition of DOM,93 and occurs at much lower rates in oxygenated waters compared to deoxygenated because of the very fast reaction of CH3 with O2.92b Measurements in the north and tropical Atlantic Ocean have confirmed that surface saturation anomalies of CH3I are correlated with light intensity, and that the photochemical source of CH3I is abiotic, and suggest that this route can support at least half of the average sea-to-air flux of 23 nmol m−2 day−1.94 Similar photochemical aqueous reactions involving CDOM and Cl− have been shown in the laboratory to produce methyl chloride (CH3Cl), and are suggested to play a role in CH3Cl production in estuarine and coastal waters and to a lesser extent in open ocean waters.95 De Laurentiis and co-workers89 suggest that the reaction of Br2•− with DOM is an important radical bromination mechanism for production of org-Br, and

2.2. Oxidized Nitrogen Compounds

2.2.1. Nitrogen Oxides and Alkyl Nitrates (RONO2) from Solar Photolysis of Nitrite. Nitrate is the most abundant nitrogenous compound in the ocean and is a key limiting nutrient in many oceanic regions. Concentrations at the surface are typically 2−30 μg N L−1, and increase with depth.102 Nitrite is an intermediate in the biological interconversion between ammonium and nitrate. ConcentraH

DOI: 10.1021/cr5007123 Chem. Rev. XXXX, XXX, XXX−XXX

Chemical Reviews

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tions in the open ocean are generally low in surface waters (

Chemistry and release of gases from the surface ocean.

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