On the flux of oxygenated volatile organic compounds from organic aerosol oxidation

the

flux

oxygenated

volatile

organic

compounds

from organic

aerosol oxidation

Alan J. Kwan,

John D. Crounse,

Antony D. Clarke,

Yohei Shinozuka,

Bruce E. Anderson,

James H. Crawford,

Melody A. Avery,

Cameron S. McNaughton,

William H.

Brune,

Hanwant B. Singh,

and Paul O. Wennberg

1,7

Received 24 February 2006; revised 17 April 2006; accepted 12 June 2006; published 9 August 2006.

[

] Previous laboratory and field studies suggest that

oxidation of organic aerosols can be a source of oxygenated

volatile organic compounds (OVOC). Using measurements

of atmospheric oxidants and aerosol size distributions

performed on the NASA DC-8 during the INTEX-NA

campaign, we estimate the potential magnitude of the

continental summertime OVOC flux from organic aerosol

oxidation by OH to be as large as

70 pptv C/day in the

free troposphere. Contributions from O

, photolysis,

and other oxidants may increase this estimate. These

processes may provide a large, diffuse source of OVOC

that has not been included in current atmospheric models,

and thus have a significant impact on our understanding of

organic aerosol, OVOC, PAN, and HO

chemistry. The

potential importance and highly uncertain nature of our

estimate highlights the need for more field and laboratory

studies on organic aerosol composition and aging.

Citation:

Kwan, A. J., et al. (2006), On the flux of oxygenated

volatile organic compounds from organic aerosol oxidation,

Geophys. Res. Lett.

, L15815, doi:10.1029/2006GL026144.

1. Introduction

[

] Oxygenated volatile organic compounds (OVOC)

comprise a large number of the species whose transport to

the remote troposphere can impact radical budgets and

sequester NO

in the form of nitrates [

Singh et al.

, 1995;

Wennberg et al.

, 1998;

Muller and Brasseur

, 1999]. In

addition, they play a role in the formation of organic

aerosols (OA) [

Kanakidou et al.

, 2005, and references

therein]. Field campaigns have noted large concentrations

of OVOC throughout the troposphere, but their budgets are

poorly understood [

Singh et al.

, 2001, 2004;

Wisthaler et

al.

, 2002].

[

]

Ellison et al.

[1999] suggest that oxidation of OA

may provide an OVOC source to the remote troposphere.

Field campaigns have established the ubiquity of OA

throughout the troposphere [

Murphy et al.

, 1998;

Kanakidou

et al.

, 2005]. Most aerosols contain both organic and

inorganic components, though significantly, the organic

fraction tends to be found on aerosol surfaces [

Tervahattu

et al.

, 2002a, 2002b, 2005, and references therein].

[

] Laboratory studies have demonstrated that organic

surfaces can be oxidized by OH and O

[e.g.,

Rudich

, 2003,

and references therein;

Thornberry and Abbatt

, 2004;

Molina et al.

, 2004]. Several of these studies have shown

volatilization of OVOC resulting from organic surface

oxidation.

Molina et al.

[2004], for example, report the full

volatilization of a C

alkane monolayer following hetero-

geneous loss of 2–3 OH radicals to the surface, with many

(though not exclusively) OVOC products. Evidence for

such chemistry in the ambient environment include the

demonstration by

Grannas et al.

[2004] that photooxidation

of snow phase organic matter may explain the daytime flux

of lightweight OVOC from snowpack to the boundary layer,

and field observations that atmospheric OA becomes more

oxidized with greater oxidant exposure and/or age [

Gogou

et al.

, 1996;

de Gouw et al.

, 2005;

Quinn et al.

, 2005;

McFiggans et al.

, 2005;

Robinson et al.

, 2006].

[

] Here, we use data collected on the NASA DC-8 air-

craft during the INTEX-NA campaign (H. Singh et al.,

manuscript in preparation, 2006) to place constraints on

OA oxidation’s potential contribution to continental sum-

mertime OVOC budgets. This campaign was designed to

examine the transport and transformation of airmasses over

continental scales. Most of the flights took place over the

North American continent and the North Atlantic in the

summer of 2004.

2. Method

[

] Aerosol size distributions, surface area, and volume

for particles 10 nm to 3

m in diameter were measured on

the DC-8 using a differential mobility analyzer (DMA) and

an optical particle counter (OPC) [

Clarke et al.

, 2004].

Comparison of the DMA and OPC data to measurements

of larger and ultrafine particles indicates that these instru-

ments generally capture >90% of the total aerosol surface

area except in a few select plumes. The aerosol size is

quantified in conditions that are often dryer than the ambient

atmosphere, so the aerosols may lose water (and thus mass)

prior to measurement. Correcting for this effect is non-trivial,

particularly for submicron particles, so we neglect it in our

calculations. Based on the ratio of ambient to dry aerosol

scattering coefficients, we believe the resulting underesti-

mate of ambient surface area is significantly less than a factor

GEOPHYSICAL RESEARCH LETTERS, VOL. 33, L15815, doi:10.1029/2006GL026144, 2006

Division of Engineering and Applied Science, California Institute of

Technology, Pasadena, California, USA.

Division of Chemistry and Chemical Engineering, California Institute

of Technology, Pasadena, California, USA.

Department of Oceanography, University of Hawai’i at Manoa,

Honolulu, Hawaii, USA.

NASA Langley Research Center, Hampton, Virginia, USA.

Department of Meteorology, Pennsylvania State University, University

Park, Pennsylvania, USA.

NASA Ames Research Center, Moffett Field, California, USA.

Also at Division of Geological and Planetary Sciences, California

Institute of Technology, Pasadena, California, USA.

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of 2. OH measurements were made by laser induced fluores-

cence [

Faloona et al.

, 2004], O

by chemiluminescence (M.

Avery et al., FASTOZ: An accurate, fast-response in situ

ozone measurement system for aircraft campaigns, submitted

Journal of Oceanic and Atmospheric Technology

, 2006),

and H

by chemical ionization mass spectrometry (J. D.

Crounse et al., Measurements of gas-phase hydroperoxides

by chemical ionization mass spectrometry, submitted to

Analytical Chemistry

, 2006).

[

] The collision rate of an oxidant with aerosol along the

flight track was estimated using 1-minute averages of the

measured aerosol surface area and oxidant mixing ratios:

collisions

=

8RT

=

ðÞ

=

ðÞ

where the factor of

=

converts aerosol surface area to cross

sectional area, (8RT/

=

is the thermal speed of the

oxidant (R is the universal gas constant, T is temperature, M

is the oxidant molar mass), S the aerosol surface area, [O

]

the concentration of oxidant (converted to a 24-hour

average in the case of OH), and f(D) the correction applied

due to gas-phase diffusion limitations to large particles (for

OH only) [

Fuchs and Sutugin

, 1971].

[

] For each point along the DC-8 flight track, we

estimate a 24-hour average [OH] by scaling the observed

[OH] by the ratio of the diurnally-averaged to instantaneous

[OH] calculated from the highly-constrained NASA LaRC

photochemical box model [

Crawford et al.

, 1999]. The

main source of uncertainty in determining the diurnal

average is that cloud effects, which during INTEX-NA

normally altered actinic flux

20% relative to clear sky

conditions, are assumed to be constant throughout the day,

though they are generally transient. We expect that our large

data set sufficiently captures the variability of cloud effects

such that they will not significantly bias our estimate.

[

] Following the studies of

Bertram et al.

[2001] and

Molina et al.

[2004] we assume that each OH collision is

reactive (

= 1) and volatilizes 6 organic carbons. Although the

alkane used in the

Molina et al.

[2004] study is not represen-

tative of all OA surfaces, long chain fatty acids may comprise a

significant fraction [

Tervahattu et al.

, 2002a, 2002b, 2005].

These parameters, which also assume full aerosol surface

coverage by organic substrates, represent by far the largest

uncertainties in our analysis. Because our assumptions imply a

unity accommodation coefficient (

= 1), we account for

diffusion limitations in calculating the OH collision rate.

[

] Estimating the OVOC flux from aerosol collisions with

and H

is significantly more challenging and is not

attempted here. While OH is highly reactive with many classes

of organic compounds, O

reactivity and product yield are

very substrate dependent [

Rudich

, 2003;

Thornberry and

Abbatt

, 2004]; such selectivity also means that

would be

time dependent as reactive sites are depleted [

Poschl et al.

2001;

Ammann et al.

, 2003]. Also, unlike for OH [

Molina et

al.

, 2004], O

reactivity with solid organic surfaces may

depend on relative humidity [

Poschl et al.

, 2001]. For H

we found no experimental studies allowing us to constrain

the product yield. For these oxidants, though, estimating the

collision rate with aerosols is a first step for assessing their

potential contributions to OVOC flux. We expect that the

accommodation coefficients for these two oxidants are signif-

icantly less than unity –

Berkowitz et al.

[2001], for example,

estimate

– and thus do not consider diffusion

limitations for their collision rates.

3. Results and Discussion

[

] Calculated 24-hour average collision rates are plotted

as a function of elevation in Figure 1. For OH (Figure 1a)

and O

(Figure 1b), most of the variability is driven by the

(dry) aerosol surface area, which varies from

in the upper troposphere to

150

in the lowest

elevation bin. For OH, we estimate the upper tropospheric

collision rate to be

collisions cm

. Our

assumption of carbon volatilization from

Molina et al.

[2004] thus yields an estimated OVOC source of

70 pptv

C/day in the upper troposphere. A significantly larger source

Figure 1.

Mean profiles of aerosol collision rates with

(a) OH, (b) O

, and (c) H

. Horizontal lines show the

interquartile range and x’s are the elevation bin medians.

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KWAN ET AL.: OVOC FLUX FROM AEROSOL OXIDATION

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would exist in the lower troposphere (

500 pptv C/day in

the lowest 2 km), but the contribution of this mechanism to

OVOC concentrations will be less significant because of the

much larger OVOC sources from gas phase oxidation of

larger hydrocarbons in this region.

[

] Assuming a product yield for O

similar to that of OH,

the ratio of OH to O

collisions gives a value of

necessary for the OVOC flux from O

to equal that from OH.

Achieving this value would require only

0.1% surface

coverage by typical alkene liquids (

[

de Gouw

and Lovejoy

, 1998;

Rudich

, 2003]). Thus, it is plausible that

for certain aerosols reaction with O

can be a significant

oxidation mechanism. This is particularly true in the lower

troposphere, where most primary (i.e., more unsaturated) OA

is expected to reside. For H

(Figure 1c), the equivalent

H2O2

ranges from

in the lower troposphere to

in the upper troposphere, so H

’s contribution to OVOC

flux is likely confined to the lower troposphere or in cloud.

Because of clouds’ high aerosol surface area, increased

actinic flux, and potential for aqueous phase chemistry,

aerosol oxidation in clouds may warrant particular attention.

[

] Photolysis can, in principle, also lead to the decom-

position of OA [

Kieber et al.

, 1990]. To be competitive with

the OH chemistry, however, photolysis rates would have to

be relatively fast (J

), as photolysis is likely less

efficient at degrading organic molecules in the condensed

phase than in the gas phase: caging effects can stabilize the

intermediates and aid their recombination or polymeriza-

tion, preventing volatilization.

4. Atmospheric Implications

[

] A dispersed and previously unconsidered source of

OVOC from OA oxidation may have important implications

for tropospheric photochemistry. For example, measure-

ments of acetaldehyde (CH

CHO) have routinely exceeded

model predictions [

Singh et al.

, 2001, 2004;

Wisthaler et al.

2002]; this was also true during INTEX-NA (Figure 2a).

For each acetaldehyde measurement, we divide the differ-

ence between measured and box model-predicted mixing

ratios by the photochemical lifetime to estimate the flux

necessary to reconcile the difference. We find that a source

90 pptv/day (

180 pptv C/day) is required to sustain

the observed acetaldehyde concentrations in the upper

troposphere. INTEX-NA also marked the first extensive

airborne measurements of peroxyacetic acid

(CH

C(O)OOH, PAA), which significantly exceeded mod-

el-predicted values in the upper troposphere as well

(Figure 2b); a similar analysis for PAA yields an estimated

missing source of

20–200 pptv C/day (J. Crounse et al.,

manuscript in preparation, 2006). Even considering only

acetaldehyde, oxidation by OH alone is likely too small to

explain the upper tropospheric discrepancies between

OVOC measurements and models; other oxidation mecha-

nisms or alternative explanations are needed.

[

] A significant flux of OVOC from aerosol would have

consequent impacts on HO

and peroxyacyl nitrates. In fact, if

the observations and our understanding of the subsequent

chemistry of acetaldehyde are correct, peroxyacetyl nitrate

(CH

C(O)OONO2, PAN) formation is very fast in the upper

troposphere. Observations of PAN are not, however, consistent

with those of acetaldehyde, based on current understanding of

the chemistry that links these compounds [

Staudt et al.

, 2003].

[

] A large OVOC flux from aerosol is also incompat-

ible with current estimates of OA budgets. The

Intergov-

ernmental Panel on Climate Change

[2001] estimate of OA

production (and loss) is

150 Tg ‘‘organic matter’’/yr, or

100 Tg C/yr. If our estimated OA oxidation rate from OH

were representative of the entire atmosphere, the global

flux of organic carbon from aerosol would be as large as

150 Tg C/yr (integrating with bin median collision rates,

100 Tg C/yr; interquartile range,

50–200 Tg C/yr).

This is clearly an upper limit due to the assumptions of

unity

and that the continental summertime conditions of

INTEX-NA are representative. Nonetheless, even a fraction

of this large flux would imply that oxidation may need to be

included in OA models, which currently consider only

depositional loss. From our carbon flux (and assuming

internally mixed aerosols with density 1 g cm

and organic

carbon fraction 0.5), we estimate the median lifetime of

aerosol organic carbon to be

10 days, similar to estimates

of carbonaceous aerosol lifetime against deposition of

5–

10 days [

Kanakidou et al.

, 2005]; thus, oxidation may be an

important sink, particularly in the upper troposphere and

regions with minimal precipitation.

[

] Consideration of an additional significant sink of OA

from oxidation would dramatically increase the required

global OA production inferred from top-down analyses

Figure 2.

Modeled (circle) vs. measured (x) profiles for

(a) acetaldehyde and (b) peroxyacetic acid. Marks are bin

medians, lines are interquartile range. Model predictions,

offset by 0.25 km for clarity, are based on the photo-

chemical box model of

Crawford et al.

[1999], and assume

a PAA lifetime of

20 days (upper limit).

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(generally calculated from measured OA burdens and esti-

mates of the depositional loss rate). Bottom-up estimates,

deduced from emission inventories and secondary organic

aerosol (SOA) yields for precursor gases, may also be too

low:

Holzinger et al.

[2005], for example, demonstrate that

many biogenic SOA precursors are too short-lived to have

been previously measured, and thus have been omitted from

emission inventories. Furthermore, because of the diversity

of SOA precursor gases, the photochemistry and SOA yield

of only a few model compounds have been extensively

studied; even for relatively well studied compounds, such as

isoprene, estimates of SOA yields are undergoing significant

revision upward [

Limbeck et al.

, 2003;

Claeys et al.

, 2004a,

2004b;

Kanakidou et al.

,2005;

Kroll et al.

,2005].In

addition, our understanding of other aspects of OA chem-

istry is poor. For example,

Heald et al.

[2005], considering

only wet depositional loss, underpredict OA mass in the free

troposphere 1–2 orders of magnitude, which they cannot

attribute merely to OA flux underestimates.

5. Conclusions and Recommendations

[

] Our estimate of the OVOC source and its atmospher-

ic impact are highly speculative and uncertain due to the

complexity of the processes involved and the paucity of

laboratory and field data for quantifying key parameters.

Continued identification of OA constituents and study of

oxidant interactions with a wider range of substrates is

necessary to better constrain OVOC flux from atmospheric

aerosols. Of particular concern is that much of the OA in the

atmosphere may actually consist of highly oxidized humic-

like substances (HULIS) [

Limbeck et al.

, 2003;

Claeys et

al.

, 2004a]; whether HULIS can be volatilized in a similar

fashion as the aliphatic compounds usually studied in the

laboratory must be investigated.

De Gouw and Lovejoy

[1998] find that O

reacts with liquid aldehydes and ketones

(

), presumably with the carbonyl moeities, but do

not determine if any gas phase products form. An additional

difficulty in applying laboratory studies is that they have

utilized reaction parameters, such as low pressures and [O

and high [O

] and [OH], which are not representative of the

real atmosphere.

Moise and Rudich

[2000] find, for exam-

ple, that

drops when O

instead of He is used as a

carrier gas in their experiments; also,

Molina et al.

[2004]

propose that the carbon volatilization in their experiments

would be reduced at atmospheric [O

]. Other oxidants, such

as NO

and – in certain areas such as coastal and polar

regions – halogen atoms, may also play an important role in

aerosol oxidation. Finally, our analysis assumes that all

aerosols are completely covered with organic films; accu-

rate parameterizations of surface coverage require more

field studies of aerosol coatings. A full understanding of

OA oxidation will only result from continued, multifaceted

research endeavors.

[

]

Acknowledgments.

We thank Jonathan Abbatt (University of

Toronto) for reviewing a draft of this manuscript, Steven Howell (University

of Hawai’i at Manoa) for helpful discussions, Clyde Brown (NASA Langley

Research Center) for preparing merged data files, and the NASA Tropo-

spheric Chemistry Program office for support of the INTEX-NA campaign

(grant NNG04FA59G). This material is based upon work supported under a

National Science Foundation Graduate Research Fellowship (AJK). This

article was also developed under a STAR Research Assistance Agreement

FP-91633401-2 awarded by the U.S. Environmental Protection Agency

(JDC). It has not been formally reviewed by the EPA. The views expressed

in this document are solely those of the authors and the EPA does not endorse

any products or commercial services mentioned in this publication.

References

Ammann, M., U. Poschl, and Y. Rudich (2003), Effects of reversible

adsorption and Langmuir-Hinshelwood surface reactions on gas uptake

by atmospheric particles,

Phys. Chem. Chem. Phys.

, 351–356,

doi:10.1039/b208708a.

Berkowitz, C. M., et al. (2001), Aircraft observations of aerosols, O

, and

NOy in a nighttime urban plume,

Atmos. Environ.

, 2395–2404.

Bertram, A. K., A. V. Ivanov, M. Hunter, L. T. Molina, and M. J. Molina

(2001), The reaction probability of OH on organic surfaces of tropo-

spheric interest,

J. Phys. Chem. A

105

, 9415–9421.

Claeys, M., et al. (2004a), Formation of secondary organic aerosols through

photooxidation of isoprene,

Science

303

, 1173–1176.

Claeys, M., et al. (2004b), Formation of secondary organic aerosols from

isoprene and its gas-phase oxidation products through reaction with

hydrogen peroxide,

Atmos. Environ.

, 4093–4098.

Clarke, A. D., et al. (2004), Size-distributions and mixtures of black carbon

and dust aerosol in Asian outflow: Physio-chemistry, optical properties,

J. Geophys. Res.

109

, D15S09, doi:10.1029/2003JD004378.

Crawford, J., et al. (1999), Assessment of upper tropospheric HOx sources over

the tropical Pacific based on NASA GTE/PEM data: Net effect on HOx and

other photochemical parameters,

J. Geophys. Res.

104

(D13), 16,255–

16,273.

de Gouw, J. A., and E. R. Lovejoy (1998), Reactive uptake of ozone by

liquid organic compounds,

Geophys. Res. Lett.

(6), 931–934.

de Gouw, J. A., et al. (2005), Budget of organic carbon in a polluted

atmosphere: Results from the New England Air Quality Study in 2002,

J. Geophys. Res.

110

, D16305, doi:10.1029/2004JD005623.

Ellison, G. B., A. F. Tuck, and V. Vaida (1999), Atmospheric processing of

organic aerosols,

J. Geophys. Res.

104

(D9), 11,633–11,641.

Faloona, I. C., et al. (2004), A laser-induced fluorescence instrument for

detecting tropospheric OH and HO

: Characteristics and calibration,

J. Atmos. Chem.

, 139–167.

Fuchs, N. A., and A. G. Sutugin (1971), High-dispersed aerosols, in

Topics

in Current Aerosol Research

, part 2, edited by G. M. Hidy and J. R.

Block, pp. 1–60, Elsevier, New York.

Gogou, A., et al. (1996), Differences in lipid and organic salt constituents of

aerosols from eastern and western Mediterranean coastal cities,

Atmos.

Environ.

(7), 1301–1310.

Grannas, A. M., P. B. Shepson, and T. R. Filley (2004), Photochemistry and

nature of organic matter in Arctic and Antarctic snow,

Global Biogeo-

chem. Cycles

, GB1006, doi:10.1029/2003GB002133.

Heald, C. L., D. J. Jacob, R. J. Park, L. M. Russell, B. J. Huebert, J. H.

Seinfeld, H. Liao, and R. J. Weber (2005), A large organic aerosol source

in the free troposphere missing from current models,

Geophys. Res. Lett.

, L18809, doi:10.1029/2005GL023831.

Holzinger, R., A. Lee, K. T. Paw U, and A. H. Goldstein (2005), Observa-

tions of oxidation products above a forest imply biogenic emissions of

very reactive compounds,

Atmos. Chem. Phys.

, 67–75.

Intergovernmental Panel on Climate Change (2001),

Climate Change 2001:

The Scientific Basis—Contribution of Working Group I to the Third

Assessment Report of the Intergovernmental Panel on Climate Change

Cambridge Univ. Press, New York. (Available at http://www.grida.no/

climate/ipcc_tar/wg1/.)

Kanakidou, M., et al. (2005), Organic aerosol and global climate modeling:

A review,

Atmos. Chem. Phys.

, 1053–1123.

Kieber, R. J., X. Zhou, and K. Mopper (1990), Formation of carbonyl com-

pounds from UV-induced photodegradation of humic substances in natural

waters: Fate of riverine carbon in the sea,

Limnol. Oceanogr.

, 1503–

1515.

Kroll, J. H., N. L. Ng, S. M. Murphy, R. C. Flagan, and J. H. Seinfeld

(2005), Secondary organic aerosol formation from isoprene photooxida-

tion under high-NOx conditions,

Geophys. Res. Lett.

, L18808,

doi:10.1029/2005GL023637.

Limbeck, A., M. Kulmala, and H. Puxbaum (2003), Secondary organic

aerosol formation in the atmosphere via heterogeneous reaction of

gaseous isoprene on acidic particles,

Geophys. Res. Lett.

(19), 1996,

doi:10.1029/2003GL017738.

McFiggans, G., et al. (2005), Simplification of the representation of the

organic component of atmospheric particulates,

Faraday Disc.

130

341–362, doi:10.1039/b419435.

Moise, T., and Y. Rudich (2000), Reactive uptake of ozone by proxies for

organic aerosols: Surface versus bulk processes,

J. Geophys. Res.

105

(D11), 6469–6476.

Molina, M. J., A. V. Ivanov, S. Trakhtenberg, and L. T. Molina (2004),

Atmospheric evolution of organic aerosol,

Geophys. Res. Lett.

L22104, doi:10.1029/2004GL020910.

L15815

KWAN ET AL.: OVOC FLUX FROM AEROSOL OXIDATION

L15815

4of5

Muller, J.-F., and G. Brasseur (1999), Sources of upper tropospheric HOx:

A three-dimensional study,

J. Geophys. Res.

104

(D1), 1705–1715.

Murphy, D. M., D. S. Thomson, and M. J. Mahoney (1998), In situ mea-

surements of organics, meteoric material, mercury, and other elements in

aerosols at 5 to 19 kilometers,

Science

282

, 1664–1669.

Poschl, U., T. Letzel, C. Schauer, and R. Niessner (2001), Interaction of

ozone and water vapor with spark discharge soot aerosol particles coated

with benzo[a]pyrene: O

and H

O adsorption, benzo[a]pyrene degrada-

tion, and atmospheric implications,

J. Phys. Chem. A

105

, 4029–4041.

Quinn, P. K., et al. (2005), Impact of particulate organic matter on the

relative humidity dependence of light scattering: A simplified parameter-

ization,

Geophys. Res. Lett.

, L22809, doi:10.1029/2005GL024322.

Robinson, A. L., N. M. Donahue, and W. F. Rogge, (2006), Photochemical

oxidation and changes in molecular composition of organic aerosol in the

regional context,

J. Geophys. Res.

111

, D03302, doi:10.1029/

2005JD006265.

Rudich, Y. (2003), Laboratory perspectives on the chemical transformations

of organic matter in atmospheric particles,

Chem. Rev.

103

, 5097–5124.

Singh, H. B., M. Kanakidou, P. J. Crutzen, and D. J. Jacob (1995), High

concentrations and photochemical fate of oxygenated hydrocarbons in the

global troposphere,

Nature

378

, 50–54.

Singh, H., et al. (2001), Evidence from the Pacific troposphere for large

global sources of oxygenated organic compounds,

Nature

410

, 1078–

1081.

Singh, H. B., et al. (2004), Analysis of the atmospheric distribution,

sources, and sinks of oxygenated volatile organic chemicals based on

measurements over the Pacific during TRACE-P,

J. Geophys. Res.

109

, D15S07, doi:10.1029/2003JD003883.

Staudt, A. C., D. J. Jacob, F. Ravetta, J. A. Logan, D. Bachiochi, T. N.

Krishnamurti, S. Sandholm, B. Ridley, H. B. Singh, and B. Talbot (2003),

Sources and chemistry of nitrogen oxides over the tropical Pacific,

J. Geophys. Res.

108

(D2), 8239, doi:10.1029/2002JD002139.

Tervahattu, H., et al. (2002a), New evidence of an organic layer on marine

aerosols,

J. Geophys. Res.

107

(D7), 4053, doi:10.1029/2000JD000282.

Tervahattu, H., J. Juhanoja, and K. Kupiainen (2002b), Identification of an

organic coating on marine aerosol particles by TOF-SIMS,

J. Geophys.

Res.

107

(D16), 4319, doi:10.1029/2001JD001403.

Tervahattu, H., et al. (2005), Fatty acids on continental sulfate aerosol

particles,

J. Geophys. Res.

110

, D06207, doi:10.1029/2004JD005400.

Thornberry, T., and J. P. D. Abbatt (2004), Heterogeneous reaction of ozone

with liquid unsaturated fatty acids: Detailed kinetics and gas-phase pro-

duct studies,

Phys. Chem. Chem. Phys.

, 84–93.

Wennberg, P. O., et al. (1998), Hydrogen radicals, nitrogen radicals, and the

production of ozone in the upper troposphere,

Science

279

, 49–53.

Wisthaler, A., A. Hansel, R. R. Dickerson, and P. J. Crutzen (2002), Or-

ganic trace gas measurements by PTR-MS during INDOEX 1999,

J. Geophys. Res.

107

(D19), 8024, doi:10.1029/2001JD000576.

B. E. Anderson, M. A. Avery, and J. H. Crawford, NASA Langley

Research Center, Hampton, VA 23681, USA.

W. H. Brune, Department of Meteorology, Pennsylvania State University,

University Park, PA 16802, USA.

A. D. Clarke, C. S. McNaughton, and Y. Shinozuka, Department of

Oceanography, University of Hawai’i at Manoa, Honolulu, HI 96822, USA.

J. D. Crounse, Division of Chemistry and Chemical Engineering,

California Institute of Technology, Pasadena, CA 91125, USA.

A. J. Kwan and P. O. Wennberg, Division of Engineering and Applied

Science, California Institute of Technology, Pasadena, CA 91125, USA.

(kwan@caltech.edu)

H. B. Singh, NASA Ames Research Center, Moffett Field, CA 94035,

USA.

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