Recent changes in the air&hyphen;sea gas exchange of methyl chloroform

Recent changes in the air-sea gas exchange of methyl chloroform

Paul O. Wennberg,

1,2

Synte Peacock,

James T. Randerson,

1,2,4

and Rainer Bleck

Received 7 May 2004; revised 14 June 2004; accepted 16 July 2004; published 24 August 2004.

[

] Atmospheric measurements of methyl chloroform

provide important constraints on the rate of oxidation of

hydrocarbons in Earth’s atmosphere. Estimates of the loss of

methyl chloroform to the oceans play a small but important

role in these calculations. Here, we examine the ocean-

atmosphere interaction of methyl chloroform in a global

ocean model. Contrary to previous assumptions, these

simulations suggest that the high-latitude oceans are

currently a source of this chemical to the atmosphere. If

confirmed, this finding alters estimates of the change in the

atmospheric oxidation rate of hydrocarbons. We highlight

the potential usefulness of methyl chloroform as a tracer of

ocean circulation.

NDEX

ERMS

0322 Atmospheric

Composition and Structure: Constituent sources and sinks; 0312

Atmospheric Composition and Structure: Air/sea constituent fluxes

(3339, 4504); 0365 Atmospheric Composition and Structure:

Troposphere—composition and chemistry; 0368 Atmospheric

Composition and Structure: Troposphere—constituent transport

and chemistry.

Citation:

Wennberg, P. O., S. Peacock, J. T.

Randerson, and R. Bleck (2004), Recent changes in the air-sea gas

exchange of methyl chloroform,

Geophys. Res. Lett.

, L16112,

doi:10.1029/2004GL020476.

1. Introduction

[

] Manufacture of the solvent 1, 1, 1-trichloroethane

(CCl

, methyl chloroform, here after MC) was first

controlled and later essentially banned by amendments to

the Montreal Protocol on Substances that Deplete the Ozone

Layer. Because its atmospheric lifetime is about five years,

the concentration of MC in the atmosphere has fallen from

its peak value of



130 pptv in 1991/1992 to less than

35 pptv today [e.g.,

Prinn et al.

, 2001]. The decline in the

concentration of MC is responsible for most of the decrease

in stratospheric chlorine levels observed to date [

Montzka et

al.

, 1999].

[

] Atmospheric measurements of the global distribution

of MC have provided important constraints on the rate of

oxidation of hydrocarbons in Earth’s troposphere. Because

the location, quantity, and timing of MC release to the

atmosphere are well known [

McCulloch and Midgley

2001], measurements of the atmospheric concentration

allow its lifetime to be reasonably estimated [e.g.,

Prinn

et al.

, 2001]. As with many hydrocarbons, this lifetime is

determined primarily by gas phase oxidation initiated by

hydroxyl radicals (OH). With knowledge of the oxidative

lifetime of MC, the lifetimes of numerous other compounds

with similar chemistry, such as methane [

Hein et al.

, 1997]

and hydrochlorofluorocarbons (HCFCs), can be estimated.

[

] Recently, inversion of the atmospheric MC record by

Prinn et al.

[2001] suggested that the atmospheric lifetime of

MC had increased markedly in the 1990s. Prinn et al.

concluded that the increase in the lifetime resulted from a



10–15% decrease in the concentration of OH during this

time period. Such a finding is alarming because it implies a

significant increase in the atmospheric lifetime of numerous

greenhouse gases, including the CFC replacements (HCFCs).

[

] We report here simulations of the uptake and redis-

tribution of MC by the ocean. As hypothesized by

Krol and

Lelieveld

[2003], our simulations suggest MC began to

outgas from the high-latitude oceans following its rapid

decrease in the atmosphere in the mid and late 1990s. The

change in ocean exchange described here reduces the

inferred trend in the atmospheric lifetime of MC.

Prinn et

al.

[2001] suggested that a ‘missing source’ of about 117 Gg

(



moles) during the five year period from 1995 to

2000 would be required to yield no trend in the atmospheric

lifetime. The change in the ocean-atmosphere interaction

described here accounts for



30% of this deficit.

2. Interaction of Methyl Chloroform

With Ocean Water

[

]

Wine and Chameides

[1989] first described the inter-

action of MC and ocean water. Ship-based measurements of

its concentration in surface water and in the overlying air

confirmed that the oceans were an important sink for MC

[

Butler et al.

, 1991]. Although only sparingly soluble in

water [

Gossett

, 1987], a significant amount of MC has

entered the ocean during the last thirty years. Significant

transfer from the atmosphere to the ocean occurs in the

warm tropical oceans, driven mainly by chemical loss

(hydrolysis) in the surface waters. Uptake has also occurred

at mid- and high- latitudes, due to the higher solubility of

MC in colder waters and two distinct high-latitude water-

mass transformation processes: high-latitude deep convec-

tion, and wind-forced lateral subduction into the main

thermocline of the ocean.

[

] It is well established that MC is chemically removed

in warm ocean water. Observations by

Butler et al.

[1991] in

the tropics showed significantly lower dissolved MC than

expected from equilibrium with the atmosphere. They

demonstrated that the flux of MC from the atmosphere to

the ocean was broadly consistent with physical uptake

[

Gossett

, 1987] and subsequent hydrolysis [

Jeffers et al.

1989;

Gerkens and Franklin

, 1989]. The hydrolysis lifetime

in ocean water, however, is strongly dependent on temper-

GEOPHYSICAL RESEARCH LETTERS, VOL. 31, L16112, doi:10.1029/2004GL020476, 2004

Division of Geological and Planetary Science, California Institute of

Technology, Pasadena, California, USA.

Also at Division of Engineering and Applied Sciences, California

Institute of Technology, Pasadena, California, USA.

Department of Geophysical Sciences, University of Chicago, Chicago,

Illinois, USA.

Now at the Department of Earth System Science, University of

California, Irvine, California, USA.

Los Alamos National Laboratory, Los Alamos, New Mexico, USA.

L16112

1of4

ature. At 25

°

C, the lifetime against hydrolysis is about one

year. This increases to



50 years at 3

°

C, though this value

is quite uncertain because it is determined by extrapolation

of the laboratory data from warmer temperatures [

Jeffers et

al.

, 1989;

Gerkens and Franklin

, 1989].

[

] In previous studies of the atmospheric lifetime of MC,

loss from the atmosphere to the ocean has been treated as a

simple first order process. Consistent with the recommen-

dation of

Butler et al.

[1991], a lifetime (1/k

a-o

)of



60 years

with respect to the atmospheric reservoir is generally

prescribed. Such a simple parameterization is adequate for

the component of MC lost by hydrolysis in warm ocean

water, where the lifetime in ocean water (



1year)is

shorter than in the atmosphere (



5 years). It is, however,

inadequate for loss via physical uptake in the cold (



°

C),

high latitude ocean. In this process, the flux is sensitive to

both the time derivative of the atmospheric concentration and

the rate at which mixing of the surface waters into the deep

ocean occurs – as both determine the saturation properties of

the surface waters. Thus, when the atmospheric concentration

of MC was growing exponentially, the flux into the ocean

was proportional to [CH

CCl

]. After 1990, when the atmo-

spheric concentration abruptly began to decrease, loss to the

high latitude ocean was a complicated function of both

atmospheric MC concentration and ocean dynamics.

3. Global Ocean Model

[

] In this study, we have used a 60-year simulation of

ocean transport from the Miami Isopycnal Coordinate

Ocean Model (MICOM) to investigate the uptake of MC.

MICOM is a primitive equation numerical general circula-

tion model that describes the evolution of momentum, mass,

heat and salt in the ocean [

Bleck et al.

, 1992, 1989]. Further

details of the model are provided in auxiliary material

[

] We have forced MICOM with observed (and into the

future with modeled) concentrations of atmospheric MC.

We used the emissions inventory of

McCulloch and Midgley

[2001] along with the tropospheric and stratospheric loss

rates suggested by

Prinn et al.

[2001] to calculate the

concentration in a three box model (northern extra tropics,

the tropics and southern extratropics). The exchange

rate between the hemispheres was picked such that the

latitudinal gradient observed between 1970 and 1990 was

captured. In this model, atmospheric measurements of MC

are reproduced to better than 10%.

[

] Tracers (MC and CFC-11) were advected offline

using 30-day averages of isopycnal and diapycnal mass

fluxes. These were obtained by transforming instantaneous

horizontal fluxes in the two non-isopycnic model layers

representing the surface mixed layer and a ‘‘fossil’’ mixed

layer to density coordinates, integrating over time, and

inferring diapycnal fluxes from mass conservation. Air-sea

gas exchange was parameterized using the protocol

described by

Dutay et al.

[2002].

4. Results

[

] Figure 1 shows the calculated, annually averaged

flux of MC into and out of the ocean in 1990, 2000, and

2010. In 1990, the flux is everywhere into the ocean. The

largest fluxes are at low-latitudes. Outside the tropics, the

uptake pattern reflects the thermal asymmetry of the anti-

cyclonic subtropical gyres. In 2000, while significant losses

to the ocean are calculated in the tropics, these are nearly

balanced by a net source to the atmosphere from the high

latitude oceans - particularly in the southern hemisphere. By

2010 the net air-sea flux over most of the low- and mid-

latitude ocean is close to zero, reflecting the expectation that

the atmospheric burden has dropped to nearly zero. The

high-latitude northern and southern oceans remain net

sources to the atmosphere, although smaller than in 2000.

[

] The calculated time series of the net MC flux into the

ocean, and an inert tracer (no hydrolysis) with the same

Figure 1.

Yearly time-averaged net air-sea flux of methyl

chloroform for 1990, 2000 and 2010 in moles km

Positive fluxes are into the ocean. The zero contours are

highlighted in black.

Auxiliary material is available at ftp://ftp.agu.org/apend/gl/

2004GL020476.

L16112

WENNBERG ET AL.: METHYL CHLOROFORM AND THE OCEANS

L16112

2of4

solubility, is shown in Figure 2. Also shown is the ocean

loss parameterization used in previous inversions [e.g.,

Prinn et al.

, 2001]. Prior to 1980, when atmospheric

MC concentrations were increasing rapidly, the first-order

loss calculated with an assumed 1/k

a-o

of 60 years agrees

well with the total loss calculated here. Between 1980 and

1992, the first-order loss parameterization is slightly larger

than the best estimate calculated here. Between 1992 and

2000, the difference between the MICOM-calculated loss

and the first-order loss estimate gradually increases.

Between 1980 and 2000, the integrated difference is 5

moles with much of the difference (3

moles)

occurring between 1992 and 2000. We have attempted to

estimate the uncertainty in this calculation due to poor

knowledge of the hydrolysis rate at cold temperatures. As

described in the auxiliary material, the error-bars in the upper

panel of Figure 2 correspond to our estimate of the lower and

upper bounds of the hydrolysis rate. In the lower panel of

Figure 2, the fluxes are shown for various latitude bands.

Most of the chemical loss occurs equatorward of 45

°

while

physical uptake, and later release, occurs at high latitudes.

5. Model Evaluation

[

] There are only a few published measurements of MC

in the ocean to evaluate the model representation of the

coupled chemistry and physics. For evaluation of the

calculated chemical loss at warm temperature, we rely on

the tropical and subtropical MC surface saturation measure-

ments described by

Butler et al.

[1991] and

Lobert et al.

[1995]. In the tropics, the sub-saturation calculated in this

work is similar to that reported by

Lobert et al.

[1995] and

somewhat smaller than those of

Butler et al.

[1991]. In cold

water, depth profile measurements of MC illustrate that MC

is quite conservative (very slow biological or abiological

hydrolysis loss). Profiles from the north Pacific are de-

scribed by

Lobert et al.

[1995]; additional profiles were

obtained during the 1993, 1994, 1995, 1996, and 1998

World Ocean Circulation Experiment (WOCE) repeat cruise

tracks in the Labrador Sea (AR07W). In both cases, the

profiles are similar to the inert CFC-11 and CFC-12,

showing (unlike CCl

) no evidence of rapid loss. See

auxiliary material for further discussion and comparisons

of our calculations with these and other data.

[

] To evaluate the calculation of the physical uptake of

MC and its subduction through the thermocline we have

compared calculations of the air-sea exchange of CFC-11

with measurements of CFC-11 made during the WOCE.

Because the atmospheric temporal trajectory of MC (until

1992) and CFC-11 are very similar this is a reasonable

comparison. The calculated inventory of CFC-11 in 1995

was 5.6

moles, remarkably close to the 5.5 ± 1

mole inventory estimated from WOCE [

Willey et al.

2004]. This suggests that, in the absence of chemical loss,

the model MC inventory would also be quite accurate.

[

] In summary, the losses of MC to the tropical ocean

appear to be reasonably well represented by this calculation.

The fluxes out of the high latitude Southern Ocean in the

period 1990–2000 are more uncertain. The Southern Ocean

is, in general, poorly simulated by most ocean models

[

Dutay et al.

, 2002]. On an encouraging note, a sensitivity

run in which gas exchange was set to zero at all latitudes

south of 60

°

S (thereby preventing any tracer from entering

the ocean at high southern latitudes) yielded very similar

results to the control run. This was mainly due to anoma-

lously large air-sea gradients forming in the region just

north of 60

°

S as water moved across this boundary, which

in large part compensated the zero-air-sea-flux requirement.

However, the most important test of these conclusions

would be measurements of MC depth profiles and surface

saturation measurements in the Southern Ocean.

6. Conclusions

[

] Consistent with the conclusion of

Krol and Lelieveld

[2003], this study suggests that accurately assessing trends

in the atmospheric lifetime of MC during the period of

rapidly changing emissions (1990–2000) is challenging and

perhaps impossible. Any reservoir for MC with a lifetime

longer than



5 years, later released to the atmosphere, will

produce a significant error in attempts to infer the atmo-

spheric lifetime of this compound. Here we address one

such reservoir – the high latitude ocean. There are,

undoubtedly others.

Krol and Lelieveld

note, for example,

that a delay in the emission of just 65 Gg MC from the early

to the late 1990s (for example by stockpiling in advance of

the Montreal Protocol controls [

McCulloch and Midgley

2001]) would be sufficient to remove the entire OH trend

Figure 2.

Air-sea flux of methyl chloroform for the globe

(top) and by latitude band (bottom). Positive is into the

ocean. In the top panel, the black curve (with uncertainty

bounds) shows the calculated air-sea flux; the blue curve

shows the first-order loss parameterization used in previous

inversions and the dashed line shows the difference. Also

shown for reference is the calculated air-sea flux for a

hypothetical tracer which is methyl chloroform without the

hydrolysis loss (top - light blue; bottom dashed).

L16112

WENNBERG ET AL.: METHYL CHLOROFORM AND THE OCEANS

L16112

3of4

inferred by

Prinn et al.

[2001]. Such delayed emissions

have been observed in both Europe [

Krol et al.

, 2003] and in

North America [

Barnes et al.

, 2003;

Millet and Goldstein

2004]. Following the rapid decline in its emission in 1990,

MC alone is not likely to be a useful tracer of changes in the

gas-phase oxidation rates in Earth’s troposphere. Its future

utility may lie, however, in its pulse-like appearance in the

ocean.

[

] Measurements of the pulse of

C following atmo-

spheric bomb tests in the 1960s provided significant

constraints on ocean circulation [

Broecker et al.

, 1985].

This study suggests that the rapid rise and fall of atmo-

spheric MC may provide a new tracer against which to test

models of ocean circulation. Unlike bomb tritium, which

also had a pulse-like atmospheric source, MC is a gas whose

atmospheric dispersion does not rely on incorporation into

raindrops, and hence is more easily modeled. Thus,

ocean observations of MC combined with measurements

of CFC-11 and CFC-12 may provide a significant constraint

on diffusive/advective transport in the oceans on the 50–

100 year timescale critical for understanding how the ocean

uptake of CO

may change over the next century [

Hall et

al.

, 2002]. The interactions of MC with ocean water are in

some ways simpler than

(apparently no interaction

with the biosphere) and in other ways more complicated

(hydrolysis). Both MC and

C are similar, however, in that

the ocean is currently changing from an atmospheric sink to

an atmospheric source [

Caldeira et al.

, 1998].

References

Barnes, D. H., S. C. Wofsy, B. P. Fehlau, E. W. Gottlieb, J. W. Elkins, G. S.

Dutton, and S. A. Montzka (2003), Urban/industrial pollution for the

New York City–Washington, D. C., corridor, 1996–1998: 1. Providing

independent verification of CO and PCE emissions inventories,

J. Geo-

phys. Res.

108

(D6), 4185, doi:10.1029/2001JD001116.

Bleck, R., H. P. Hanson, D. M. Hu, and E. B. Kraus (1989), Mixed layer

thermocline interaction in a 3-dimensional isopycnic coordinate model,

J. Phys. Oceanogr.

, 1417–1439.

Bleck, R., C. Rooth, D. M. Hu, and L. T. Smith (1992), Salinity-driven

thermocline transients in a wind-forced and thermohaline-forced isopycnic

coordinate model of the North-Atlantic,

J. Phys. Oceanogr.

, 1486–

1505.

Broecker, W. S., T. H. Peng, G. Ostlund, and M. Stuiver (1985), The

distribution of bomb radiocarbon in the ocean,

J. Geophys. Res.

6953–6970.

Butler, J. H., J. W. Elkins, T. M. Thompson, B. D. Hall, T. H. Swanson, and

V. Koropalov (1991), Oceanic consumption of CH

CCl

-implications for

tropospheric OH,

J. Geophys. Res.

, 22,347–22,355.

Caldeira, K., G. H. Rau, and P. B. Duffy (1998), Predicted net efflux of

radiocarbon from the ocean and increase in atmospheric radiocarbon

content,

Geophys. Res. Lett.

, 3811–3814.

Dutay, J.-C., et al. (2002), Evaluation of ocean model ventilation with

CFC-11: Comparison of 13 global ocean models,

Ocean Model.

,89–

120.

Gerkens, R. R., and J. A. Franklin (1989), The rate of degradation of 1, 1,

1-trichloroethane in water by hydrolysis and dehydrochlorination,

Chemosphere

, 1929–1937.

Gossett, J. M. (1987), Measurement of Henrys law constants for C1 and C2

chlorinated hydrocarbons,

Environ. Sci. Technol.

, 202–208.

Hall, T. M., T. W. N. Haine, and D. W. Waugh (2002), Inferring the con-

centration of anthropogenic carbon in the ocean from tracers,

Global

Biogeochem. Cycles

, 1131, doi:10.1029/2001GB001835.

Hein, R., P. J. Crutzen, and M. Heimann (1997), An inverse modeling

approach to investigate the global atmospheric methane cycle,

Global

Biogeochem. Cycles

, 43–76.

Jeffers, P. M., L. M. Ward, L. M. Woytowitch, and N. L. Wolfe (1989),

Homogeneous hydrolysis rate constants for selected chlorinated

methanes, ethanes, ethenes, and propanes,

Environ. Sci. Technol.

965–969.

Krol, M., and J. Lelieveld (2003), Can the variability in tropospheric OH be

deduced from measurements of 1, 1, 1-trichloroethane (methyl chloro-

form)?,

J. Geophys. Res.

108

(D3), 4125, doi:10.1029/2002JD002423.

Krol, M. C., et al. (2003), Continuing emissions of methyl chloroform from

Europe,

Nature

421

, 131–135.

Lobert, J. M., T. J. Baring, J. H. Butler, S. A. Montzka, R. C. Myers, and

J. W. Elkins (1995), OAXTC 92: Ocean/atmosphere exchange of trace

compounds 1992,

NOAA Tech. Memo., ERL CMDL-9

, 43 pp.

McCulloch, A., and P. M. Midgley (2001), The history of methyl chloro-

form emissions: 1951–2000,

Atmos. Environ.

, 5311–5319.

Millet, D. B., and A. H. Goldstein (2004), Evidence of continuing methyl-

chloroform emissions from the United States,

Geophys. Res. Lett.

doi:10.1029/2004GL020166, in press.

Montzka, S. A., J. H. Butler, J. W. Elkins, T. M. Thompson, A. D. Clarke,

and L. T. Lock (1999), Present and future trends in the atmospheric

burden of ozone-depleting halogens,

Nature

398

, 690–694.

Prinn, R. G., et al. (2001), Evidence for substantial variations of atmo-

spheric hydroxyl radicals in the past two decades,

Science

292

, 1882–

1888.

Willey, D. A., R. A. Fine, R. E. Sonnerup, J. L. Bullister, W. M. Smethie Jr.,

and M. J. Warner (2004), Global oceanic chlorofluorocarbon inventory,

Geophys. Res. Lett.

, L01303, doi:10.1029/2003GL018816.

Wine, P. H., and W. L. Chameides (1989), AFEAS report in scientific

assessment of stratospheric ozone, vol. 2,

Rep. 20

, pp. 273–295, Global

Ozone Res. and Monit. Proj., World Meteorol. Org., Geneva.

R. Bleck, Los Alamos National Laboratory, Los Alamos, NM 87545,

USA. (bleck@lanl.gov)

S. Peacock, Department of Geophysical Sciences, University of Chicago,

Chicago, IL 60637, USA. (synte@geosci.uchicago.edu)

J. T. Randerson, Department of Earth System Science, University of

California, Irvine, Irvine, CA 92697, USA. (jranders@uci.edu)

P. O. Wennberg, Division of Geological and Planetary Science, California

Institute of Technology, Pasadena, CA 91125, USA. (wennberg@caltech.

edu)

L16112

WENNBERG ET AL.: METHYL CHLOROFORM AND THE OCEANS

L16112

4of4