Difference between revisions of "Atmospheric chemistry changes over the 21st century"

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This page is part of the topic Atmospheric change over the next 100 years

Antarctic stratospheric ozone over the next 100 years

The success of the Montreal Protocol in constraining production of ozone depleting substances (CFCs, Halons, and organic chlorides and bromides) has meant that their amounts in the stratosphere are now decreasing at about 1%/year. Models predict the future evolution of ozone based on specific scenarios for future emissions of ozone-depleting substances. Both two and three-dimensional models have been used to predict future ozone loss, but because of the more complex stratospheric dynamics near the poles, polar ozone is best simulated by 3-D models. For comparison with past measurements, Chemistry Transport Models that use prescribed wind fields from past measurements have been very successful, but they clearly cannot be used for predictions of the future. Instead, coupled Chemistry-Climate Models are used. Those whose results are described below vary in their skill in representing the atmosphere, but there is sufficient general agreement between them and observations that we can have some confidence in their predictions.

Table 5.1 Models whose results appear in Figures 5.11, 5.12 and 5.13; * bromine chemistry not included

In the figures below, four diagnostic quantities are discussed:

1. The minimum ozone in the Southern Hemisphere during September to October, which is a common diagnostic of the maximum depth of the ozone hole, and so of the maximum ozone loss.
2. The ozone mass deficit (OMD), which is the mass of ozone that would be required to elevate ozone columns above 220 Dobson Units (DU), as observed by satellite instruments (Huck et al., 2007^[13]). OMD can be calculated daily or averaged over some period, and it is the most accurate diagnostic of total Antarctic ozone loss.
3. The ozone hole area, which is the area with ozone less than 220 DU as measured by satellite instruments. Because they can only observe in sunlight, in early September there can be an unobserved area of more than 220 DU close to the pole, whose area is not subtracted from the area within the outer contour. Nevertheless this is a common diagnostic of ozone hole size.
4. The total inorganic chlorine (Cly), which equals the sum of chlorine in organic chlorine compounds entering the stratosphere, after degradation by UV light and reaction with oxides of hydrogen and nitrogen. Components of Cly are the comparatively stable compounds HCl and ClNO₃, as well as the reactive Cl, ClO, OClO, Cl₂O₂ and HOCl.

Stratospheric ozone is affected by a number of natural and anthropogenic factors in addition to reactive halogens: temperature, transport, volcanoes, solar activity, hydrogen oxides, and nitrogen oxides. In any discussion of future ozone, it is important to separate the effects of these factors, particularly if considering the future success or otherwise of the Montreal Protocol. For example, when Cly has decreased to pre-ozone hole values, continued cooling of the stratosphere due to increased greenhouse gases warming the troposphere may lead to amounts of stratospheric ozone quite different to those of pre-ozone hole days. The net result is not that simple, however. Increases in greenhouse gases affect polar ozone via processes acting in opposing directions, making model predictions less certain near the poles than elsewhere:

1. Increased greenhouse gases act to cool the stratosphere, which will slow gas-phase ozone loss reactions and so tend to increase stratospheric ozone.


5.12 (a) September to October average daily ozone mass deficit, and (b) the maximum ozone hole area, for each year from each model. Curves, shading and source of observations are as in Figure 5.11. Adapted from WMO (2007^[14]).
2. The same cooling acts to increase the amounts of PSCs, on which the reactions leading to ozone loss occur, so we can expect reduced ozone given the same Cly - opposite to the effect of gas-phase chemistry in 1 (above). This is particularly likely in the edge region of the vortex (Lee A et al., 2001), because PSCs are not ubiquitous there. Ozone loss in this edge region defines the ozone hole area diagnosed in Figure 5.12.
3. Increased GHGs act to increase the strength of the Brewer-Dobson circulation by changing the wave driving that causes it. In the short term, this increases the supply to the stratosphere of (a) CH₄ and so hydrogen oxides, (b) N₂O and so nitrogen oxides, and (c) CFCs and halons, and so reactive chlorine and bromine. Each of these helps to remove ozone, so this process also acts to reduce ozone in the short term. In the long term, the increased removal of CFCs and halons via the actions of the Protocol would more rapidly reduce ozone-depleting substances in the whole atmosphere, so that the effects from reduced CFCs and halons would oppose those arriving via CH₄ and N₂O.

5.11 Minimum total column ozone in September to October predicted by various models, plus observations from the National Institute of Water and Atmosphere Research (NIWA) total ozone database in New Zealand (Bodeker et al., 2005^[15]). Solid and dashed curves show smoothed values. Light gray shading shows when CFCs and halons are expected to return to 1980 values. Adapted from WMO (2007^[14]).

Despite these complications, the general characteristics of future Antarctic ozone as shown in Figure 5.11 are similar in all models, and similar to projections in earlier models in WMO (2003^[16]): minimum ozone occurs around 2000, followed by a slow increase. The increase is slow because the near-total destruction of ozone in the core of the ozone hole means that there is low sensitivity to Cly there, so only small changes in ozone hole depth are expected as Cly starts to decline. Larger sensitivity to changes in Cly is expected at the upper altitudes of the ozone hole (20-22 km) where ozone depletion is not complete, and this is a possible region in which to detect the onset of ozone hole recovery (Hofmann et al., 1997^[17]).

The minimum amount of ozone in Figure 5.11 differs widely between models, ranging from 60 DU to over 120 DU compared to the observed 80 DU, highlighting the difficulty of predictions of polar ozone by fully-coupled models. The fact that Chemistry Transport Models agree much better with observations and with each other suggests that it is transport in these fully coupled models that accounts for the differences and difficulties.

Similarly, the predicted values of maximum ozone mass deficit in Figure 5.12 vary widely between models (from 7 to over 33 million tons, compared to the observed 31 million tons), as do predictions of maximum ozone hole area. Note that because both ozone mass deficit and ozone hole area are below a 220 DU threshold, a bias in global ozone in any one model will create a bias of opposite direction in both diagnostics (e.g. the low bias in area and mass deficit in MAECHAM4CHEM is probably caused by a general high bias in global ozone).

5.13 Zonal mean values of total inorganic chlorine (ppbv) in October at 50 hPa (20-25 km) and 80°S predicted by models: (a) total, (b) difference from 1980. Open black diamonds in (a) show estimates from measurements by UARS satellite in 1992 and by Aura satellite in 2005.

Some insight into the model differences can be obtained from comparisons of Cly in the models. As shown in Figure 5.13, there is a large spread in the simulated Cly, including in the maximum value and in the date at which it decreases to 1980 values. In several models, the maximum Cly is unrealistically low with the result that its return to 1980 values is too early, which is likely to ensure that model’s return to 1980 values of ozone is too early. More weight should therefore be put on results from models with more realistic maximum Cly. AMTRAC matches the observations of Cly best, and predicts the latest return to 1980 values.

AMTRAC also predicts the latest recovery of ozone in Figure 5.11 and Figure 5.12. This is almost consistent with the study of Newman et al. (2006^[18]), who used a parametric model of spring ozone amounts that includes Cly amounts and stratospheric temperatures. Figure 5.12 shows that they predicted that a return to 1980 ozone amounts would not occur until about 2070.

Despite the differences in models, extrapolating their results suggests that by 2100 Antarctic ozone will no longer be under the influence of CFCs and halons. However, it may not have reverted to 1980 values because of changes in stratospheric temperatures and dynamics caused by increased greenhouse gases.

Antarctic Tropospheric Chemistry

In this section we discuss how the chemistry of the Antarctic troposphere is influenced by the cryosphere, and how that chemistry may change in a warmer world with a reduced cryosphere.

Trace gases in the atmosphere are split apart (photolysed) by solar radiation to generate highly reactive radicals. Over regions of snow and ice, incoming solar radiation is reflected back by the surface; the extent of this reflection is the surface albedo, which varies between 0.81 and 0.83 over snow in coastal Antarctica (Gardiner and Shanklin, 1989^[19]). The presence of snow on both land and sea ice surfaces thus significantly increases the pathway of solar radiation by reflection, so amplifying opportunities for photolysis of trace gases in the troposphere. Without this high albedo, the lifetime of trace gases would be increased and Antarctic tropospheric chemistry would be less reactive.

The cryosphere also limits tropospheric chemistry by acting as a cap to emissions from the underlying land or ocean. If warming were to reduce the extent of the snowpack, exposing land surfaces, it is likely that microbes would become active in the continental soil. Microbes are recognised sources of nitrous oxide (N₂O), which is a long-lived and powerful greenhouse gas. Emissions of N₂O from such freshly exposed soil would be likely to contribute to some extent to global warming.

Sea ice has a direct influence on boundary layer chemistry. Newly forming sea ice with its associated brine pools and frost flowers, as well as sea salt on snow, are all potential sources of inorganic bromine compounds. These trace gases are emitted into the boundary layer where they are potent destroyers of ozone. Under specific meteorological conditions, ozone depletion events (ODEs) are observed, where ozone concentrations can drop from a normal background amount to below instrument detection limits within a matter of minutes. These extreme events are observed within the boundary layer using ground-based instruments. Vertical profile measurements of ozone have also been made using balloon-borne sensors to assess the influence of such halogen-driven ozone loss at higher altitudes. Studies at two different coastal sites in Antarctica have reported significant reductions in ozone up to around 3 km above the snow surface (Wessel et al., 1998^[20]; Kreher et al, 1997^[21]). Ozone is a radiatively important gas whose influence varies with altitude, being more important in the free troposphere where it is colder (Lacis et al., 1990^[22]). Roscoe et al. (2001^[23]) suggested that Antarctic boundary layer ozone loss could be sustained and mixed to higher altitudes where the reduction in ozone would exert a radiative cooling that would be significant on a regional scale. They further argued that in a warmer world with reduced sea ice extent, the natural process of BrO production and therefore ozone depletion would be reduced. Ozone in the free troposphere would consequently be sustained at higher concentrations, thereby exerting an additional warming influence. They estimated the additional warming to be of the order 0.05ºK, i.e. a small but positive feedback.

Plainly, sea ice also stops emissions to the atmosphere of trace gases with an oceanic origin. A loss of sea ice would therefore enhance emissions of such gases to the atmosphere. Tropospheric trace gases released from the oceans around Antarctica include: dimethyl sulphide (DMS); alkenes such as ethene (C₂H₄) and propene (C₃H₆); and bromocarbons such as bromoform (CHBr₃) and dibromomethane (CH₂Br₂), all generated by phytoplankton. DMS plays a critical role as a source of cloud condensation nuclei (CCN) via its oxidation to sulphate (SO₂) (see e.g. Finlayson-Pitts and Pitts, 1999^[24]). Changing the number of CCN alters cloud properties and albedo, so influences the Earth’s radiation budget, surface temperature and climate. Feedback loops may exist with increased DMS emissions resulting in enhanced CCN, with consequent changes in climate that would then impact on DMS production and emission (Charlson et al., 1987^[25]). It has been suggested that chemical analyses of deep ice cores confirm Charlson's hypothesis, but this is not confirmed by recent (and apparently still unique) chemistry/climate modeling (Castebrunet et al, 2006^[26]).

Oxidation of DMS is predominantly driven by reaction with hydroxyl ions (OH) and proceeds via two channels:

OH + CH₃SCH₃ → CH₃S(OH)CH₃ (addition)

OH + CH₃SCH₃ → CH₃SCH₂ + H₂O (abstraction)

The proportion of DMS being oxidised by the respective channels depends partly upon ambient temperature; at temperatures below 12ºC the addition channel leading to dimethylsulphoxide (DMSO) is believed to dominate (Arsene et al., 1999^[27]). The abstraction channel, which dominates at warmer temperatures, ultimately results in production of SO₂, and hence CCN. In a warmer world, the proportion of DMS oxidising with OH via the abstraction channel could therefore increase, forming more CCN. However, that possibility may be countered by the influence of BrO. BrO oxidises DMS to DMSO rather than to SO₂, and Von Glasow et al. (2002^[28]) showed that by including BrO reactions in an atmospheric chemistry transport model, global concentrations of DMSO increased by 63%. Around coastal Antarctica, even present day concentrations of BrO appear to significantly influence DMS (Read et al., 2008^[29]). Future concentrations of BrO might be affected in two ways. In a future warmer world, with the potential for additional oceanic emissions of bromocarbons, background BrO concentrations in coastal Antarctica could quite likely be higher than they are today. But the sea ice source leading to BrO might become less important, thus reducing BrO concentrations. Thus, even with enhanced levels of DMS in the future it is not clear that greater numbers of CCN would result around Antarctica. Whether or not the climate was influenced by these reactions would depend on the balance between increases in the oceanic sources, changes in BrO, and ambient temperature.

Iodine compounds are also emitted into the boundary layer. Significant concentrations of iodine monoxide (IO) have been observed in the boundary layer at Halley station on the Weddell Sea coast (Saiz-Lopez et al., 2007^[30]). The seasonal maximum occurred in spring, but even during the summer, concentrations were high enough to influence tropospheric chemistry processes. The IO is thought to originate from diatoms under the sea ice. It is a major source of new particles (marine aerosols and CCN) from which clouds originate. Aerosols and clouds scatter incoming solar radiation and so cool the atmosphere. The link between IO and sea ice suggests that in a warmer world with less sea ice, less IO will be produced, hence there will be fewer aerosol and CCN particles, and less radiative scattering, encouraging additional warming at the Earth’s surface. But, if IO emanates directly from the ocean, and is not dependent upon the presence of sea ice, then reduced sea ice could increase IO emissions, hence CCN and aerosol production, encouraging cooling. To clarify whether the feedback would be positive or negative it is necessary to determine the source of boundary layer IO.

The snowpack itself also influences the atmospheric boundary layer by acting as a source of highly reactive trace gases. Over parts of Antarctica, such as the polar plateau, these significantly increase the oxidising capacity of the boundary layer above what is expected. In coastal regions, the effect is less pronounced. The differences reflect both the fetch of snow and the stability of the boundary layer. If all Antarctic snow disappeared, this source of trace gases would be removed, and the atmosphere would most likely move to a more sluggish state with longer-lived and less reactive chemical species. But, given that the East Antarctic Ice Sheet is unlikely to disappear for thousands of years or longer, emissions from snow will be an important driver of local tropospheric chemistry for years to come.

Clearly the cryosphere over and around Antarctica has a significant influence on tropospheric composition and chemistry. Although there are significant uncertainties, it is clear that a substantial change in the Antarctic cryosphere would alter Antarctic tropospheric chemistry from its present day state. Such an alteration might itself further influence the climate system. These factors need to be encapsulated in climate models.

References

1. ↑ Austin, J., Wilson, J., Li, F. and Vömel, H. 2006. Evolution of water vapor and age of air in coupled chemistry climate model simulations of the stratosphere, J. Atmos. Sci., in press.
2. ↑ Akiyoshi, H., Sugita, T., Kanzawa, H. and Kawamoto, N. 2004. Ozone perturbations in the Arctic summer lower stratosphere as a reflection of NO_x chemistry and planetary scale wave activity. J. Geophys. Res., 109, D03304.
3. ↑ Beagley, S.R., De Grandpré, J., Koshyk, J.N., McFarlane, N.A. and Shepherd, T.G. 1997. Radiative-dynamical climatology of the first-generation Canadian Middle Atmosphere Model, Atmos. Ocean, 35, 293-331.
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6. ↑ Manzini, E., Steil, B., Brühl, C., Giorgetta, M.A. and Krüger, K. 2003. A new interactive chemistry-climate model: 2. Sensitivity of the middle atmosphere to ozone depletion and increase in greenhouse gases and implications for recent stratospheric cooling, J. Geophys. Res., 108, 4429.
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18. ↑ Newman, P.A., Nash, E.R., Kawa, S.R., Montzka, S.A. and Schauffler, S.M. 2006. When will the Antarctic ozone hole recover?, Geophys. Res. Lett., 33, L12814.
19. ↑ Gardiner, B.G. and Shanklin, J.D. 1989. Measurements of solar and terrestrial radiation at Faraday and Halley, British Antarctic Survey Report, ISBN 0-85665-139-7.
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21. ↑ Kreher, K., Johnston, P.V., Wood, S.W., Nardi, B. and Platt, U. 1997. Ground-based measurements of tropospheric and stratospheric BrO at Arrival Heights, Antarctica. Geophys. Res. Lett., 24(23), 3021-3024.
22. ↑ Lacis, A.A., Weubbles, D.J. and Logan, J.A. 1990. Radiative forcing of climate change by changes in the vertical distribution of ozone, J. Geophys. Res., 95, 9971-9981.
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24. ↑ Finlayson-Pitts, B.J. and Pitts, J.N. 1999. Chemistry of the Upper and Lower Atmosphere, Academic Press.
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27. ↑ Arsene, C., Barnes I. and Becker, K.H. 1999. FT-IR product study of the photo-oxidation of dimethyl sulfide: Temperature and O₂ partial pressure dependence, Phys. Chem. Chem. Phys, 1, 5463-5470.
28. ↑ Von Glasow, R., Sander, R., Bott, A. and Crutzen, P.J. 2002. Modelling halogen chemistry in the marine boundary layer. 2. Interactions with sulfur and cloud-covered MBL, J. Geophys. Res., 107, 4323.
29. ↑ Read, K.A. Lewis, A.C., Bauguitte, S., Rankin, A.M., Salmon, R.A., Wolff, E.W., Saiz-Lopez, A., Bloss, W.J., Heard, D.E., Lee, J.D. and Plane, J.M.C. 2008. DMS and MSA measurements in the Antarctic Boundary Layer: impact of BrO on MSA production, Atmos. Chem. Phys., 8, 2985-2997.
30. ↑ Saiz-Lopez, A., Mahajan, A.S., Salmon, R.A., Bauguitte, S.J.-B., Jones, A.E., Roscoe, H.K. and Plane, J.M.C. 2007. Boundary layer halogens in coastal Antarctica, Science, 317, 348-351, DOI: 10.1126/science.1141408.