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We have already mentioned briefly that there are two components to stratospheric ozone depletion arising from human activities. One is the slow but steady depletion by a few percent per decade, least serious in the tropics but increasing with latitude. In the northern hemisphere the decrease is larger in winter and spring (11% since 1979) than in summer or autumn (4% since 1979). A negative trend of the annual variation amplitude (~0.1 ppm yr-1) was observed in the upper stratosphere of southern subpolar latitudes. Interannual changes in amplitudes of both annual and semi-annual variations are small as a rule (except in the tropical mid-stratosphere, where the influence of the El Chichón eruption was substantial, and the subpolar upper stratosphere in the southern hemisphere). The vertical distribution of the ozone trend shows distinct negative trends at about 18 km in the lower stratosphere with largest declines over the poles, and above 35 km in the upper stratosphere. A narrow band of large negative trends extends into the tropical lower stratosphere (Brunner et al. 2006).
Bojkov and Fioletov (1995), based on re-evaluated TOC observations from over 100 Dobson and filter radiometer stations from pole to pole, presented the following TOC trend results for the steady depletion of ozone:
Up to 1995 TOC continued its decline (which started in the 1970s), with statistically significant year-round and seasonal trends except over the equatorial belt.
The cumulative year-round TOC reduction over the 35°-60° belts of both hemispheres from the early 1970s until the mid 1990s was up to 8%. While in the southern mid-latitudes, it is difficult to distinguish seasonal dependence of TOC trends, the cumulative decline in the northern mid-latitudes in winter and spring is about 9% and 4-6% for summer and autumn.
At that time observations from 12 Dobson polar stations had demonstrated that the northern polar region shows the same ozone decline as northern mid-latitudes or even a slightly stronger one (the cumulative decline is about 7% year-round and 9% for winter and spring).
Nowadays, according to the recent scientific ozone assessment (WMO 2010) the average TOC values in 2006-2009 remain at roughly 3.5% and 2.5% below the 1964-1980 averages, respectively for 90°S-90°N and 60°S-60°N. Midlatitude (35°-60°) annual mean TOC 1996-2005, at ~6% (~3.5%) below the 1964-1980 average.
The second component consists of the two seasonal ozone holes in the polar regions, the Antarctic ozone hole having appeared earlier and being larger than the Arctic ozone hole. In winter, in each of the polar regions air becomes trapped in a circumpolar vortex and its temperature drops and when it becomes very cold (below about 195 K or -78 oC) polar stratospheric clouds (PSCs) form. These are sometimes called nacreous clouds or mother of pearl clouds because of the colours they exhibit. They consist, not of water, but of frozen particles of nitric acid and water, especially nitric acid trihydrate, HNO3.3H2O. When the sun begins to return at the end of the winter free chlorine atoms are released as a result of photodissociaion, by the solar UV light, of CFCs or chemicals derived from the CFCs. Each chlorine atom acts as a catalyst to the destruction of tens of thousands of ozone molecules. The Cl-catalysed ozone destruction occurs in the gas phase, but it is accelerated by heterogeneous catalysis on the surfaces of the particles in the polar stratospheric clouds. Given the longevity of CFC molecules, the recovery times for the ozone layer are of the order of several decades. For instance, a CFC molecule needs five to seven years to be transported from the ground level up to the upper atmosphere, and resides there for several decades, destroying huge quantities of ozone molecules. There are two reasons why the Antarctic ozone hole has been known for longer, and is larger than the Arctic ozone hole; first the Antarctic is a bit colder than the Arctic and secondly the (normal) ozone concentration is slightly higher in the Arctic.
In general, the Arctic experiences high extreme cold as well as sudden stratospheric warmings (SSWs) at times. As a result the degree of ozone loss is mostly controlled by the strength of the vortex and magnitude of the air temperature within it. For instance, according to Kuttippurath et al. (2010) the winters of 1995, 1996, 2000, and 2005 were very cold and the cumulative total ozone loss was as high as ~25- 35% (WMO 2007). On the other-hand, the winters of 1997, 1998, 1999, 2001, 2002, 2006, and 2009 were relatively warm and the loss was minimal about 10-15%, while the winters of 2003, 2007, and 2008 were moderately cold and hence, the loss was in an average scale of about 15-20% (WMO 2007; Goutail et al. 2010).
As far as the future projection is concerned, the global annually averaged TOC is expected to reach to 1980 levels before 2050. The simulated changes in tropical TOC from 1960 to 2100 are generally small.
Chandra et al. (1996) using measurements of TOC by Nimbus-7 TOMS version 7, suggested that the trends over the latitudes centred at 40°N of the Northern Hemisphere vary from -3 to -9% per decade during winter and within -1 to -3% per decade during summer. They also found, using 30-hPa temperature as an index of dynamical variability, that the large negative trends in column ozone are reduced in magnitude by 1 to 3% per decade. In order to establish the dynamical component of interannual variability, they constructed similar time series of stratospheric temperature from the National Meteorological Center (NMC) analyses and channel 4 (50-150 hPa) of MSU on the TIROS series of NOAA operational satellites.
Long-term measurements of the stratospheric ozone with balloon-borne instruments allow winter ozone altitude profiles to be compared between the Antarctic and Arctic regions (WMO 2010). Inspection of Figure 9 shows that in the Antarctic at the South Pole (left panel), a normal ozone layer was observed to be present between 1962 and 1971. What is particularly interesting about the left hand half of Figure 9 is that, as shown here for 9 October 2006 in spring over Antarctica , the ozone is almost completely destroyed between 14 and 21 km. In the last decades (1990-2009) the average ozone concentrations in October are 90% lower than pre-1980 values at the peak altitude of the ozone layer (16 km). In contrast, the ozone layer over the Arctic region does not exhibit any depletion as shown by the average ozone profile in March for 1991-2009 obtained over the Ny-Ålesund site (right panel). No Ny-Ålesund data are available for the 1962-1971 period before significant ODS destruction began. Some March profiles do reveal significant depletion (e.g. the case of 29 March 1996). In these cases, minimum temperatures during wintertime are generally lower than normal, allowing PSC formation for longer periods. Arctic profiles with depletion similar to that shown for 9 October 2006 at the South Pole have never been observed. The number in parentheses for each profile is the total ozone value in Dobson units (DU). Ozone abundances are shown here as the pressure of ozone at each altitude using the unit "milli-Pascals" (mPa).
7. Dynamics of atmospheric ozone
From what we have said already in this article, it can be seen that there are many sources of data on atmospheric ozone concentrations using instruments on the ground, airborne instruments and instruments on satellites. However, the concentration of atmospheric ozone changes rapidly, both spatially and temporally. Two factors are involved. First, ozone is continuously being generated and destroyed by solar UV radiation. Secondly, the atmosphere is constantly in motion. It is not easy to separate these two aspects. Assessments of relative contributions of photochemical and dynamical processes to ozone concentration field formation confirmed earlier conclusions. Ozone distribution in the lower stratosphere (below the 30 hPa level) is controlled (especially in winter) mainly by large-scale atmospheric circulation, while in the upper stratosphere (above the 5-10 hPa levels) radiative and photochemical processes dominate (except at subpolar latitudes in winter when atmospheric transport is also important). In the tropics this zone is located at a lower level because of the shorter photochemical lifetime in the equatorial belt under conditions of high insolation. However, Thouret et al. (1998a) pointed out that there are still large uncertainties in the budget of atmospheric ozone and especially with regard to the relative importance of photochemical formation and exchanges between the stratosphere and troposphere, and in determining future ozone trends. In this context certain regular features of ozone spatial distribution were analysed. Thouret et al. (1998a) reached the conclusion that north of 35°N the distribution at 9-12 km altitude is dominated by the influence of ozone-rich air of stratospheric origin; farther south, ozone-poor air from the troposphere prevails. To identify the stratospheric and tropospheric components and to help in interpreting the data a classification based on a threshold of 100 ppbv of ozone was used.
While it is possible to use data from many different sources to determine the concentration of ozone at one particular point and at one particular time, to follow the movement of ozone is much more difficult. One serious attempt to do this has been made by the Match Network. This originated from the 1991-1993 European Arctic Stratospheric Ozone Experiment (EASOE) campaigns (Von der Gathen et al. 1995) because there was already substantial coverage of mid- and high-latitude ozone sounding stations in the northern hemisphere. Several dozen of these stations agreed to launch ozonesondes on schedule over the period when ozone loss is greatest in the Arctic vortex, which is typically from December to mid-March.
The basic idea of Match is to determine stratospheric polar ozone losses by observing individual air masses using ozonesondes at two stations over which the same air mass passes; the loss of ozone can then be determined by the difference between the two ozonesonde profiles. The name "Match" thus comes from the attempt to match two ozonesondings that correspond to the same air mass at different instants in its lifetime. The EASOE experiment included the launch of a large number (~ 1400) of ozonesondes in the Arctic and in mid-latitudes. This data set was used afterwards to determine Arctic ozone losses for the first time with the Match method. Since the launches were not coordinated, matches were only obtained by chance. However, because of the large number of soundings made during EASOE enough matches were found for a successful analysis and in this way the Match method was established.
The method was then developed into an active method, to eliminate and to be able to economise on the number of ozonesondes launched, for the Second European Stratospheric and Mid-latitude Experiment (SESAME, 1998, 1999) in the winter of 1994-1995. Since 1994-1995 the Alfred Wegener Institute for Polar and Marine Research (AWI) has coordinated the Match network campaign in most Arctic winters and in a single Antarctic winter in order to determine the chemical ozone depletion in the stratosphere. The objective of Match, which is to probe, i.e. to determine the ozone content of, a lot of air parcels twice during their way through the atmosphere, is achieved by coordinating the soundings. The element of chance that was present in the original EASOE experiment was eliminated by having the subsequent trajectories of the air masses which had passed over one ozonesonde station forecast by meteorologists at the Free University (FU) in Berlin. Thus when this trajectory passes over a second ozonesonde station the staff there can be asked to launch an ozonesonde in order to examine this air mass for a second time.
A decrease in the ozone concentration within the time period of the two ozonesonde flights can then be attributed to chemical ozone depletion so that Match is able to distinguish between dynamical and chemical variations. The vertical and time resolution is one of the best reported. Due to the great number of ozonesonde pairs, statistically significant ozone loss rates can be determined (Cortesi et al. 2007).
The Match technique has been picked up by Sasano and co-workers for determining ozone loss rates from data of the spacebome sensor Improved Limb Atmospheric Spectrometer (ILAS) (Sasano et al. 2000; Terao et al. 2002).