Figure 7.5 Contours of the zonal average of the zonal wind velocity (in s 0 on August 21, 1995. Positive values represent westerly winds. Winds are from the UKMO assimilation system.
maximum velocity of nearly 100 m s 1 in the middle stratosphere between 30 and 35 km.
The term "polar vortex" refers to the region poleward of the polar night jet. Scientific analyses of the vortex, however, generally use a definition of the vortex edge based on the maximum gradient in potential vorticity (PV) . This definition is superior to a definition based on maximum wind speed because PV is conserved by atmospheric transport (on time-scales of a month or less) in a manner similar to trace gases. Thus, defining the vortex using PV better captures the effects of vortex isolation on the trace gas distribution. In 1992, the Antarctic polar vortex covered ~35 million km2, encompassing virtually all of the southern hemisphere poleward of about 60"S. While the vortex is subject to normal meteorological variability, there is no evidence that the overall properties of the Antarctic polar vortex have changed over the last 18 years.
It is important to note that the polar vortex and the Antarctic ozone hole are different entities. The ozone hole is the region covered by low values of column O,, while the polar vortex is defined by the meteorological conditions. Between 1980 and 1996 the area covered by the polar vortex remained roughly constant. During this same time the size of the Antarctic ozone hole increased dramatically (see Figure 7.1). By the mid-1990s the ozone hole covered about two-thirds of the vortex. As we will discuss, the increase in area! extent of the ozone hole is due almost entirely to the increase in stratospheric chlorine over this time period, and not to changes in meteorology .
The Antarctic polar vortex has two important properties without which the formation of the Antarctic ozone hole would not occur. First, the temperatures in the vortex are cold (see Figure 7.4). litis allows the formation of polar stratospheric clouds (PSCs), which play a pivotal role in the chemistry of the ozone hole. We will discuss PSCs in detail in the next section. Second, as mentioned in Chapter 5, the edge of the polar vortex is a mixing barrier that isolates the Antarctic polar stratosphere from the mid-latitudes over the altitude range between -425 and -1000 K (-18 and 35 km) [186,188,221,222], To show the isolation of the vortex, in Figure 7.6 we plot O, measured on the 465 K potential temperature surface (~20 km, in the lower stratosphere) as a function of equivalent latitude. There is a strong gradient in ozone collocated with the edge of the polar vortex at 60°S. A strong gradient in a constituent with a lifetime of weeks or longer can be maintained only if mixing is slow; otherwise, mixing would wipe out the gradient.
As we will show, the chemical composition of the polar vortex is very different from the chemical composition of the mid-latitudes. As a result, the isolation of the vortex from mid-latitude air is crucial for the development and maintenance of the Antarctic ozone hole. Without it, mixing of polar vortex and mid-latitude air would short circuit ()3 loss, and prevent the formation of the whole.
Figure 7.6 Ozone abundance versus equivalent latitude on the 465 K potential temperature surface for September 15, 1992. Dots are individual measurements, and the solid line is an average of the data. Equivalent latitude is derived from UKMO PV. The vertical line denotes the edge of the vortex, as determined by the Nash et al. |219] algorithm. Ozone data are from the 205 GHz channel of the UARS MLS (version 4).
While horizontal transport into the Antarctic polar vortex is negligible between -425 and -1000 K (18 and 35 km), vertical descent within the vortex does occur |186,221,223,224|. This descent is attributable to the mean overturning circulation of the stratosphere, which features air descending at mid- and high latitudes, as well as the fact that the polar lower stratosphere is cooling, and air occupies a smaller volume as it cools. Figure 7.7 shows a calculation of the descent of air parcels within the Antarctic polar vortex. The calculation shows that air in the Antarctic polar region descends rapidly during the fall, with the rate of descent slowing as the region enters winter. By the middle of winter (late July) the descent in the lower stratosphere has virtually stopped, and the polar vortex is close to radiative equilibrium; in the upper stratosphere, descent is still occurring, but at a greatly reduced rate.
The return of sunlight at the end of winter begins heating the polar vortex, leading to higher temperatures there (see Figure 7.4). This in turn reduces the latitudinal pressure gradient, which in turn reduces wind speeds around the vortex. As the speed of the jet around the vortex is reduced, so too is the stability and isolation of the vortex. The vortex wind system abruptly transitions to the summer wind regime by way of an event called the final stratospheric warming [222,225-226], which typically occurs in late November to late December, depending on altitude. During this final warming the Antarctic vortex breaks into smaller air masses which then mix with mid-latitude air , It is this break-up of the Antarctic polar vortex that marks the end of the Antarctic ozone hole.
Finally, it should be noted that the area covered by, the range of temperatures of, and the longevity of the Antarctic polar vortex exhibit year-to-year variability (see WMO [13), Figure 3.3). As will be shown later, the meteorological variability of the polar vortex has important implications for the size and depth of the. ozone hole.
Polar stratospheric clouds At the cold temperatures found within the polar vortex, water vapor and nitric acid condense to form particles known as PSCs. PSCs are crucial to the chemistry of the polar region because, much like sulfate aerosols, they provide surfaces on which heterogeneous reactions can take place. The exact temperature at which PSCs fomt is a function of pressure and trace gas abundances, but a canonical formation temperature, valid near 20 km for typical lower stratospheric conditions, is ~196 K. Because of the requirement for such low temperatures, PSCs exist only in the winter and early spring polar regions from the tropopause to altitudes as high as -26 km (see WMO [ I3J, Figure 3.10).
PSCs are subdivided into two classes, designated imaginatively as type 1 and type II. Type I PSCs form at temperatures several degrees above the frost point (where water vapor condenses into ice). Because of this, it was suggested that type I PSCs were composed of a mixture of water and nitric acid [228-230], Laboratory work
showed that crystalline HNO, 3H20 (nitric acid trihydrate or NAT) was thermo-dynamically stable for stratospheric conditions below -196 K , This work, combined with in situ observations [2321, led the field initially to identify NAT as the composition of type I PSCs.
Subsequent atmospheric observations, however, could not be reconciled with the NAT composition. By the mid-1990s it became apparent that the real situation is more complex. Other compositions, such as liquid H20/HN0.,/H2S04 aerosol , also known as supercooled ternary solution or STS, or crystalline HN0.-2H.0 (nitric acid dihydrate or NAD) , are now thought likely to play a role [234-239],
Further, measurements have identified two classes of type 1 PSC, designated as type la and type lb. Type la particles are crystalline, likely NAT or NAD. while type lb particles are liquid, likely STS [240,241]. These particles have typical characteristic radii of 0.5-1 pm.
Type II PSCs form at temperatures below the frost point, which, for typical stratospheric values, is -189 K. For this reason, they are believed to be ice crystals. Type II PSCs are predicted to have characteristic radii of 5-20 pm [242,243].
Chlorine activation As discussed in Chapter 4, normally -95% of total inorganic chlorine, Cl„, in the lower stratosphere during the day is in the form of HC1 and C10N02, with HC1 the slightly more abundant reservoir. CIO typically makes up most of the rest. However, it has been observed that there are regions in the Antarctic lower stratosphere where nearly all of the CI, is in the form of CIO and ClOOCl [244,245]. As we will show, ClOOCl is involved in catalytic destruction of O,, so we include it in our definition of CI,:
[CIJ = [CIO] + [CI] + 2[CIOOCl] = [CIO] + 2[C100C1] (7.1)
Thus, observations suggest that virtually all of the CI, has been converted to CI,. Our present understanding of polar chemistry suggests that this repartitioning occurs because the following reactions occur on the surfaces of PSCs:
In the parlance of atmospheric chemistry, the conversion of chlorine reservoirs HC1 and ClONO, into CI, is known as "chlorine activation".
Reactive uptake coefficient y for the various reactions on NAT (type la), STS (type lb), and ice PSCs (type II) are listed in Table 7.1. Reactions on NAD have not been studied as thoroughly as those on NAT: what laboratory measurements have been made suggest that the rates on NAD are comparable to those on NAT [246.247],
Extinction measurements suggest that PSCs exist only briefly (hours to days), and air masses are only occasionally exposed. The reactive uptake coefficients, however, are generally so large that even intermittent exposure to PSCs can cause significant repartitioning of the Cl„ family. In many cases, even a brief exposure to a PSC causes reactions (7.2)—(7.5) to run to completion—i.e. run until one reactant is depleted.
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