Oxygenated volatile organic chemicals in the oceans: inferences and implications based on atmospheric observations and air-sea exchange models

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Oxygenated volatile organic chemicals in the oceans: inferences and implications based on atmospheric observations and air-sea exchange models

H. B. Singh1, A. Tabazadeh1, M. J. Evans2, B. D. Field2, D. J. Jacob2, G. Sachse3, J. H. Crawford3, R. Shetter4, W. H. Brune5

1 NASA Ames Research Center, Moffett Field, CA 94035

2 Harvard University, Cambridge, MA 02138

3 NASA Langley Research Center, Hampton, VA 23665

4 National Center for Atmospheric Research, Boulder, CO 80307

5 Pennsylvania State University, University Park, PA 16802

Abstract. Airborne measurements of several oxygenated volatile organic chemicals (OVOC), OH free radicals, and tracers of pollution were performed over the Pacific during Winter/Spring of 2001. We interpret atmospheric observations of acetaldehyde (CH3CHO), propionaldehyde (C2H5CHO), methanol (CH3OH), and acetone (CH3COCH3) with the help of a global 3-D model and an air-sea exchange model to assess their oceanic budgets. We infer that surface waters of the Pacific are greatly supersaturated with acetaldehyde and propionaldehyde and provide a large source. Bulk surface seawater concentration of 0.3 g l-1 and 0.1 g l-1 and net fluxes of 1.1 x 10-12 g cm-2 s-1 and 0.4 x 10-12 g cm-2 s-1 are calculated for CH3CHO and C2H5CHO, respectively. Large surface seawater concentrations are also estimated for methanol (3.3g l-1) and acetone (0.6g l-1) corresponding to a deposition flux of 0.3 x10-13 g cm-2 s-1 and 1.2 x 10-13 g cm-2 s-1 and an under-saturation of 3% and 14%, respectively. If extrapolated to global oceans, these fluxes imply a net oceanic source/sink (Tg y-1) of 125, 45, -4, and –14 for CH3CHO, C2H5CHO, CH3OH, and CH3COCH3, respectively. Assuming the classical view that surface seawater concentrations are representative of the top 50-100 meter mixed layer, an extremely large oceanic reservoir of OVOC, exceeding the atmospheric reservoir by an order of magnitude, can be inferred to be present. Available seawater data are both preliminary and extremely limited but indicate rather low bulk OVOC concentrations and provide no support for the existence of a large oceanic reservoir. We speculate on the potential causes and implications of these findings.

1. Introduction

In recent years it has become evident that significant concentrations of a large number of oxygenated organic chemicals (OVOC) are present in the global troposphere [Singh et al., 2001]. These chemicals are expected to play an important role in the chemistry of the atmosphere by sequestering nitrogen oxides (NOx) in the form of PAN and by providing HOx free radicals in critical regions of the atmosphere [Singh et al. 1995; Wennberg et al., 1998]. OVOC are believed to have large terrestrial sources (≈ 500 TgC y-1) but our quantitative knowledge about their sources is rudimentary [Fall, 1999]. Attempts to reconcile atmospheric observations with models have led to the inference that oceans must provide a sizable source of acetone [de Laat et al., 2001; Jacob et al., 2002]. Galbally et al. [2002] have suggested that an extremely large oceanic reservoir of methanol may be present. It is also known that sunlight initiated reactions can decompose organic matter in surface oceans to form a variety of oxygenated chemicals [Blough, 1997; Zhou and Mopper, 1997]. The role of oceans in the global budgets of OVOC remains largely unexplored and there is little reliable data available. Here we investigate and assess oceanic composition and budget of acetaldehyde (ethanal, CH3CHO), propionaldehyde (propanal, C2H5CHO), methanol (CH3OH), and acetone (propanone, CH3COCH3) based on airborne measurements over the Pacific and models of air-sea exchange. These OVOC are selected because they are among the most abundant and there is reason to believe that they are globally ubiquitous [Singh et al., 2001].

2. Experimental

During the TRACE-P Spring 2001 experiment, NASA/DC-8 flying laboratory was used to measure a large number of OVOC in the Pacific troposphere to an altitude of 12 km (Latitude 10-45 ˚N, Longitude 100-230˚E). Complementing these were in-situ measurements of hydroxyl radicals (OH), a key constituent that controls the atmospheric removal rates of OVOC, and multiple tracers of pollution (e. g. CO, C2Cl4). Additionally, spectral radiometers were employed to determine photolytic loss rates of the OVOC of interest in the troposphere. Specific measurement details can be found elsewhere and are not repeated here [Singh et al., 2001; Tan et al., 2001; Shetter and Müller, 1999]. An overview of the mission payload, flight profiles, and prevalent meteorological conditions has been provided by Jacob et al. [2003] and Fuelberg et al. [2003].

3. Results and discussion

All of the data discussed here were collected over open oceans some 500-km or more away from the Asian continent. A “pollution filter” based on CO and C2Cl4 distribution was employed to mitigate the effects of pollution and is described in more detail in Singh et al. [2003a]. This filter eliminated all major pollution influences and resulted in mean tropospheric mixing ratios of 102(±20) ppb/CO and 3(±1) ppt/C2Cl4, and is assumed to represent near-background conditions. All analysis in this paper is based on this filtered data set. The observed vertical distribution of the four-selected OVOC under these clean conditions is shown in Figure 1.

It is evident from Figure 1 that measured mixing ratios of acetone and methanol are lower in the marine boundary layer (MBL) compared to the free troposphere, while the reverse is the case for aldehydes. In Table 1 we present mixing ratio data in the MBL (0-2 km, dMBL) and over it (2-4 km) for the OVOC of interest in this study. Figure 1 also shows the mixing ratios of select OVOC as simulated by the GEOS-CHEM global tropospheric model along the TRACE-P flight tracks for the entire period. More details about the GEOS-CHEM model can be found in Jacob et al. [2002]. An in-depth analysis of the atmospheric distribution and global budgets of the large suite of OVOC measured during TRACE-P is being presented elsewhere [Singh et al., 2003a].

3.1 Ocean flux

Oceanic fluxes are estimated assuming a steady state behavior in which the entrainment at the top of the MBL, production and losses within the MBL, and net oceanic fluxes are in balance [Singh et al., 2003b]. The various flux terms are described in equations 1-3.

Ft = Ve x C (1)

FMBL = FMBL (production) – FMBL (loss) (2)

Fo = - FMBL - Ft (3)

Where Ft is the flux at the top of the MBL, FMBL is the net flux associated with chemical production and loss of OVOC in the MBL and, Fo is the calculated net oceanic flux. Based on available literature, a mean entrainment velocity (Ve) of 0.4 cm s-1 is used in all subsequent calculations [Singh et al., 2003b]. To avoid biases due to scatter in the data, gradient across the top of the MBL (C) are determined by using median concentration differences between the 0-2 km and 2-4 km layers. The calculated value of Ft for each OVOC is provided in Table 1.

Since OVOC are both produced and destroyed in the atmosphere, FMBL is a measure of net chemical flux in the MBL. Loss rates in the MBL are calculated on the assumption that chemical loss in the MBL is due to reaction with OH and photolysis {FMBL (loss) = CMBL x dMBL x (ki [OH]MBL + ji)}. The OH reaction rate constants (ki) for all species in this study were taken from the compilations of Atkinson et al. [2002] and Sander et al. [2002]. Diurnally averaged OHMBL concentrations of 1.5 x 106 molec. cm-3 were derived from the in-situ OH measurements on board the DC-8 and were also in good agreement with GEOS-CHEM model simulations. Photolytic loss rates (ji) were directly calculated from spectral radiometric measurements. Photolytic loss rates were small in comparison with loss due to OH in all cases except acetone (55%OH; 45% hv). Combined first order loss rates are shown in Table 1.

The OVOC production in the MBL is primarily determined by the concentrations of specific precursors and OH radicals. Pathways involving halogen atom reactions are controversial but also likely small and are presently neglected. In the case of acetone and methanol the main photochemical formation pathways involve propane/i-butane and methane oxidation, respectively. The acetone yield of 0.8 from propane/OH reaction is well documented and has been used here [Atkinson et al., 2002]. The methanol yield from methane oxidation involves a series of secondary reactions {2CH3O2 ---> CH3OH + HCHO+ O2} that are dependent on available NOx. Given typical atmospheric conditions, a mean molar yield of 0.03 could be derived from the GEOS-CHEM model and is used here. Median MBL mixing ratios of propane (+i-butane) and methane were 76 ppt and 1772 ppb, respectively. Acetaldehyde sources are more complex as it is an intermediate product in the oxidation of a large number of hydrocarbons. Therefor for aldehydes we have used the chemical production rates as calculated by the GEOS-CHEM model. The net MBL flux values (FMBL) due to chemistry (equation 2) are summarized in Table 1.

The net oceanic flux estimates (Fo) calculated from equation 3 are summarized in Table 1. These fluxes are positive (source to the atmosphere) for aldehydes and negative (sink) for methanol and acetone (Table 1). It is also evident that aldehyde fluxes are extremely large in comparison to methanol and acetone. If Table 1 fluxes are extrapolated to global scales, an oceanic source/sink (Tg y-1) of 125, 45, -4, and –14 can be calculated for CH3CHO, C2H5CHO, CH3OH, and CH3COCH3, respectively. There are no seawater data available to directly verify these estimates. The finding of an oceanic sink for CH3OH and CH3COCH3 is contrary to previous suggestions of a sizable oceanic source [de Laat et al., 2001; Jacob et al., 2002; Heikes et al., 2002]. As can be seen from Figure 1 (d), the presently accepted net oceanic source of CH3COCH3 (≈13 Tg y-1) in the GEOS-CHEM model greatly over predicts MBL mixing ratios. Sources of methanol are much more uncertain and its simulation is less reliable. However, an assumed sink of the order of 15 Tg y-1 provides a reasonable fit to the observations (Figure 1c). While these results must be considered preliminary at this stage, they provide the first indications of the important role that oceans could play in the global budgets of OVOC.

Here it is appropriate to add that CH3OH, and CH3COCH3 are relatively long-lived and their global concentration fields for reasonably well established [Singh et al., 1995; 2001; de Laat et al., 2001; Wisthaler et al., 2002]. Aldehydes on the other hand, have extremely short lifetimes (<1 day) and their global ubiquity in the entire troposphere is less certain. The pervasive hydrocarbon oxidation source is much too small (model lines in Figure 1a,b). However, there are substantial CH3CHO measurements in the MBL from around the globe. Mixing ratios in the range of 70-400 ppt, indicated by the shaded area in Figure 1a, have been reported from the northern and southern Pacific [Singh et al., 1995; 2001], the Atlantic [Zhou and Mopper, 1993; Arlander et al, 1995; Tanner et al, 1996], and the Indian Ocean [Wisthaler et al., 2002]. Using a newly developed mass spectrometric technique, Wisthaler et al. [2002] report MBL mixing ratios of 212±29 ppt and 178±30 ppt from the northern (0-20 ˚N) and southern (0-15 ˚S) Indian Ocean, respectively under clean conditions, in good agreement with this study (Table 1). The ensemble of these observations supports the view that substantial CH3CHO concentrations are present in the MBL throughout the globe. Although comparable data for C2H5CHO are note available, we expect it to behave much like CH3CHO. Indeed, the mixing ratios of CH3CHO and C2H5CHO measured during TRACE-P were found to be strongly correlated (R2 = 0.9). We postulate that the observed remote MBL mixing ratios of these extremely short-lived aldehydes ( ≈ 1-day) can only be sustained by oceanic sources. The large oceanic source estimated here alone are insufficient to account for the concentrations in the free troposphere. Our present relatively poor understanding of the global sources of aldehydes is further discussed in Singh et al. [2003a].

3.2 Sea water concentrations

For oceans to be a net source or a sink, active in-situ production and/or destruction of OVOC in seawater is necessary. To explore oceanic budgets of these OVOC we use a two-film model of air-sea exchange described by Liss and Slater [1974] and further discussed by Donelan and Wanninkhof [2001]. Based on this model, equations 4-5 were used to derive surface seawater concentration and saturation (S+1) necessary to maintain the oceanic flux, Fo.

Fo = Kl (Cl - Cg/H) = (1/kl +1/Hkg)-1(Cl - Cg/H) (4)

S = HCl/Cg - 1 = HFo/K1Cg (5)

where Cl and Cg are concentrations in the bulk liquid and gas phases, H (Csg/Csl) is the dimensionless Henry’s law constant, kl and kg are exchange constants for liquid and gas phases, and  is an enhancement factor due to species reaction in solution. The temperature dependent partition coefficients (KH) for pure water published by Sander, [1999] were adopted and are listed in Table 2. The corresponding dimensionless Henry’s Law constants (H) were increased by 20% to correct for the reduced solubility due to the salting out effect in seawater [Donelan and Wanninkhof, 2001]. For aldehydes, these were further multiplied by the hydration factor which has been determined to be ≈ 2.4 for acetaldehyde and 1.0 for methanol [Zhou and Mopper, 1997 and references there in]. A liquid phase exchange constant (kl) of 11 cm h-1 and a gas phase exchange constant (kg) of 6912 (18/MW)1/2 cm h-1 was used based on recommended values in the literature for a mean surface wind speed of 6 m s-1 [Donelan and Wanninkhof, 2001, Asher, 1997]. These calculations are relatively insensitive to , which is taken to be 2.0 for aldehydes and 1.0 for other OVOC [Zhou and Mopper, 1997]. Results derived from the use of equations 4-5 are summarized in Table 2.

It is evident from Table 2, that for chemicals of moderate solubility, such as OVOC, resistance to air-sea exchange both in the gas phase (rg=1/Hkg) and in the liquid phase (rl=1/kl) can be important. For acetaldehyde these resistances are nearly equal while for methanol, nearly all of the resistance (91%) is in the gas phase. Air-sea exchange models have been predominantly validated for sparingly soluble and relatively inert species (such as O2, N2, N2O, SF6, and CO2) where the liquid phase resistance dominates. Reactive organic chemicals such as OVOC as well as chemicals for which resistance to transfer resides principally in the gas phase have not been extensively studied. For OVOC, surface organic films typically present on seawater may greatly interfere with the air-sea exchange processes [Frew, 1997]. Therefore, the possibility can not be ruled out that the air-sea exchange models being applied here may suffer from large unknown errors. Given these caveats, bulk surface seawater concentrations (g l-1) of 0.3, 0.1, 3.3, and 0.6 are calculated for CH3CHO, C2H5CHO, CH3OH, and CH3COCH3, respectively (Table 2). This also implies a very large seawater super-saturation of aldehydes. In comparison, Pacific waters appear to be negligibly under-saturated in methanol (3%) and moderately in acetone (14%) and provide a net sink. This sink corresponds to an estimated deposition velocity (SKl/H) of 0.04 cm s-1 methanol and 0.12 cm s-1 acetone. These results also imply that there are active OVOC production and loss processes in the oceans.

In the classical air-sea exchange view, it is common to assume that the bulk surface concentration (Cl) is representative of the top 50-100 m well-mixed region of the oceans. Assuming a 50 m mixed layer, the oceanic loading of these species over the Pacific can be calculated to be 2-25 times their atmospheric loading (R in Table 2). Adjusting for hemispheric differences and extrapolating to global scales, a very large oceanic reservoir of these species (≈ 65 Tg, Table 2) can be calculated to be present. A deeper mixed layer would only increase the size of this reservoir. In their assessment of the global budget of methanol, Galbally et al. [2002] assumed saturation in the top 80 m to calculate an oceanic methanol reservoir some 66 times larger than the atmospheric reservoir. We note that based on our results (Table 2), the assumption of surface saturation is very nearly correct. Heikes et al. [2002], in an independent review of the methanol budget, estimate a methanol oceanic source of 0-80 Tg y-1 but do not imply a reservoir.

Is there evidence for the large oceanic concentrations and the implied reservoir of OVOC estimated in Table 2? Two studies provide some preliminary indications. In the first study, Zhou and Mopper [1997] use a wet chemical technique to measure CH3CHO and CH3COCH3 in coastal and open ocean areas east of Florida. They report a bulk seawater concentration range of 0.04-0.66 g l-1 for CH3CHO and 0.17-2.61 g l-1 for CH3COCH3. Bulk water concentrations in two open ocean samples collected from a location some 100-km east of Bahamas were 0.06 g l-1 for CH3CHO and 0.17 g l-1 for CH3COCH3. In a more recent investigation Williams et al. [2003], using a sensitive mass spectrometric technique over the open tropical Atlantic ocean, report bulk water concentrations of <0.05 g l-1 for CH3OH and CH3COCH3. They also find that these low concentrations are nearly uniform in the top 50 m of the ocean. The atmospheric concentrations measured by them are comparable to those in Table 1. The available oceanic data at the moment are preliminary and too limited to draw any firm conclusions. However, they do indicate that bulk seawater concentrations in remote oceans are much lower than predicted by current air-sea exchange models (Table 2). These measurements provide no support for the presence of a large OVOC oceanic reservoir.

A possible reason for the low observed concentrations in bulk seawater could be that OVOC are rapidly consumed by bacteria typically present in the top few meters. No biodegradability data from the oceans are available, but in other aquatic media these degradation rates appear to be comparable to the mixing times (1-2 days) in the oceanic mixed layer [Verschueren, 1996]. It is also possible that high OVOC concentrations inferred in Table 2 are present only in surface microlayers (≈102 M) over the ocean [Frew, 1997; Zhou and Mopper, 1997]. Open oceans are known to have broken surface organic films originating from lipids, organic deposition, and biological degradation. These films are expected to have little effect on the gas phase resistance (1/kg) but they can provide additional resistance in the liquid phase to impede OVOC exchange of the microlayer with the underlying mixed layer. [Asher, 1997]. We also compared the steady state surface coverage of OVOC (ss) based on the instantaneous flux (Fo in molec. cm-2 s-1) with the equilibrium coverage (eq) as defined in equations 6 and 7, respectively [Donaldson and Anderson, 1999]. A “standard” surface site density of 1015 cm-2 was used (Adamson, 1990).

ss = 10-15 [6.03.1023/MW] Fo (6)

eq = 10-15 satCl/[b+Cl] (7)

where gives the surface coverage at saturation, b is the adsorption constant in pure water in molar units and Cl is the concentration of OVOC in bulk water in molar units (Table 2). Using published values of sat and b, we find that the instantaneous surface coverage (ss) is some three orders of magnitude larger than the equilibrium coverage (eq). This implies that the sea surface layer is highly enriched in OVOC relative to that expected for a pure water sample. This enrichment can occur for multiple reasons. Surface organic films can increase the partitioning of OVOC at the air-water interface by increasing adsorption (lower b) and the adsorbed OVOC in the organic film layer can participate in fast chemistry near or at the interface layer.

Air-sea exchange models have been largely untested for moderately soluble reactive organic species such as OVOC. For species where oceans are a source, a steady flux can be maintained from deep reservoirs but thinner layers must be continuously replenished if they are not to be quickly depleted. This almost certainly requires a mechanism for near-continuous OVOC formation in seawater. If these processes are driven by the availability of sunlight, then strong diurnal variations should be expected. The super-saturation of aldehydes and under-saturation of acetone are indicative of the fact that the OVOC in the oceans are biologically and photochemical active and are not present in a dormant state.

4. Conclusions

In this study we have used atmospheric measurements of OVOC and a standard air-sea exchange model to explore their oceanic budgets. Current models suggest that oceans should contain a very large reservoir of OVOC and may provide a net source in some cases and a sink in others. However, the reliability of these models for reactive organic species remains untested. At the moment, the possibility that air-sea exchange processes are greatly impeded for chemicals such as OVOC or that the oceanic bacteria prevent their accumulation can not be ruled out. There is clearly an urgent need for more research on the abundance and fate of OVOC in the oceans. The existence of a large OVOC reservoir or its depletion by biological processes would have important consequences for ocean biogeochemistry. The development of reliable air-sea exchange models for reactive and moderately soluble organic species is also a worthwhile long-term objective. A newly initiated international program called SOLAS (Surface Ocean/Lower Atmosphere Study) is in place to lead and prioritize future research in this area [Duce and Liss, 2002].

Acknowledgments: This research was funded by the NASA Tropospheric Chemistry Program. Harvard investigators were also supported by the NSF Atmospheric Chemistry Program. Constructive discussions with R. Wanninkhof of NOAA are appreciated.


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Table 1: OVOC mixing ratios in and above the Pacific MBL under pristine conditions







Median Mixing ratio (0-2 km)

205 (204±40)#

60 (68±24)

563 (575±211)

437 (466±97)

Oceanic source (Tg y-1)***





*An entrainment velocity of 0.4 cm s-1 across MBL top and a mean MBL OH concentrations of

1.5 x 106 molec. cm-3 is used. Loss rate (e-fold) in MBL is due to OH reaction and photolysis (ki [OH]MBL + ji).

Positive and negative signs indicate MBL source and sinks, respectively.

** Fmbl = Fmbl (production) – Fmbl (loss). See text for details.

***Assumes that the calculated oceanic flux is globally applicable.

#Median mixing ratio (Mean ± 1)

Table 2: Seawater OVOC concentrations, saturation, and potential reservoir estimated by using an air-sea exchange model



(M atm-1)



(cm h-1)




(g cm-2 s-1)


(g l-1)




Oceanic reservoir







+1.1 x 10-12










-1.2 x 10-13





aPartition coefficient in pure water at 298 K

b Dimensionless Henry’s Law constants. These are increased for seawater conditions (x1.2) and to account for hydration effects ( x 2.4

for aldehydes and x 1.0 for others). See text.

c Total transfer velocity in gas and liquid phases. See text.

d Gas phase resistance to exchange (rg) as a percent of total resistance (rg+rl). See text.

e Calculated bulk seawater concentration

f Ratio of the oceanic burden (g cm-2) and the atmospheric burden (g cm-2) over the Pacific. A 50 meters mixed layer depth is assumed.

g Extrapolated global oceanic reservoir assuming R is maintained. Concentrations in the SH are assumed to be 80% of NH.

Figure 1: Measured and modeled mixing ratios of selected OVOC. Shaded area in (a) represents the range of published MBL acetaldehyde mixing ratios from many regions of the globe (see text). Model results are based on the GEOS-CHEM 3-D model which assumes a 13 Tg y-1 net oceanic source for acetone and none for aldehydes. The divergence between measured and modeled acetone in the MBL in (d) is largely due to assumption. For methanol a 15 Tg y-1 oceanic sink is shown to provide a reasonable fit to data (c). Altitudes of measured data are shifted by -0.25 km for clarity.


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