Depletion of the Ozone Layer and Climate Change - Research Paper Example

Ozone Introduction Most of the oxygen in the Earth’s atmosphere is in the form of molecules containing two oxygen atoms, known by the familiar chemical symbol O2. In certain circumstances, three atoms of oxygen can bond together to form ozone, a gas with the chemical symbol O3. Ozone occurs naturally in the Earth’s atmosphere where its concentration varies with altitude. Concentration peaks in the stratosphere at around 25–30 kilometres from the Earth’s surface and this region of concentration of the gas is known as the ozone layer (Barnett, Houghton and Pyle). The ozone layer is important because it absorbs certain wavelengths of ultraviolet (UV) radiation from the Sun, reducing their intensity at the Earth’s surface. High doses of UV radiation at these wavelengths can damage eyes and cause skin cancer, reduce the efficiency of the body’s immune system, reduce plant growth rates, upset the balance of terrestrial and marine ecosystems, and accelerate degradation of some plastics and other materials. A number of man-made chemicals are known to be harmful to the ozone layer. They all have two common properties: they are stable in the lower atmosphere and they contain chlorine or bromine. Their stability allows them to diffuse gradually up to the stratosphere where they can be broken down by solar radiation. This releases chlorine and bromine radicals that can set off destructive chain reactions breaking down other gases, including ozone, and thus reducing the atmospheric concentration of ozone. This is what is meant by ozone depletion (Bird). Another important environmental impact of a gas is its contribution to global warming. Global Warming Potential (GWP) is an estimate of the warming of the atmosphere resulting from release of a unit mass of gas in relation to the warming that would be caused by release of the same amount of carbon dioxide. Some ODS and some of the chemicals being developed to replace them are known to have significant GWPs. For example, CFCs have high GWPs and the non-ozone-depleting hydrofluorocarbons (HFCs) developed to replace CFCs also contribute to global warming. GWP is an increasingly important parameter when considering substances as candidates to replace ODS. During past decades, sufficient quantities of ODS have been released into the atmosphere to damage the ozone layer significantly. The largest losses of stratospheric ozone occur regularly over the Antarctic every spring, resulting in substantial increases in UV levels over Antarctica. A similar though weaker effect has been observed over the Arctic. At present, scientists predict that, provided the Montreal Protocol is implemented in full, ozone depletion will reach its peak during the next few years and will then gradually decline until the ozone layer returns to normal around 2050 (Allaart, Valks and van der). Data Climate change and ozone depletion have both been the subject of recent assessments describing the advances in understanding made in recent years. It is now well established that observed polar and mid-latitude ozone depletion is a consequence of the increase in stratospheric chlorine and bromine concentrations. The chemical processes involving these compounds and leading to polar ozone depletion are well understood. Similarly, the important role of ozone and the ODSs in the climate system has been documented. Stratospheric ozone depletion during recent decades has represented a negative radiative forcing of the climate system; in contrast, the increase in ODSs has been a positive radiative forcing (Dvortsov). The distribution of ozone in the atmosphere is maintained by a balance between photochemical production and loss, and by transport between regions of net production and net loss. A number of different chemical regimes can be identified for ozone. In the upper stratosphere, the ozone distribution arises from a balance between production following photolysis of molecular oxygen and destruction via a number of catalytic cycles involving hydrogen, nitrogen and halogen radical species (Chipperfield). Figure 1. Vertical profiles of ozone-related quantities. (a) Typical mid-latitude ozone mixing ratio profile, based on an update of Fortuin and Langematz (1994); (b) atmospheric temperature profile, based on Fleming et al. (1990), showing the stratosphere bounded by the tropopause below and the stratopause above; (c) schematic showing the ultraviolet (UV) radiative flux through the atmosphere (single-headed arrows) and infrared (IR) emission at around 9.6 μm (the ozone absorption band, double-headed arrows), and the heating in the ultraviolet (solid curve) and infrared (dashed curve) associated with these fluxes; (d) schematic of the change in surface temperature due to a 10% change in ozone concentration at different altitudes (based on Figure 6.1 of IPCC, 2001). Discussion Global and hemispheric-scale variations in stratospheric ozone can be quantified from extensive observational records covering the past 20 to 30 years. There are numerous ways to measure ozone in the atmosphere, but they fall broadly into two categories: measurements of column ozone (the vertically integrated amount of ozone above the surface), and measurements of the vertical profile of ozone. Approximately 90% of the vertically integrated ozone column resides in the stratosphere. There are more independent measurements, longer time-series, and better global coverage for column ozone. Regular measurements of column ozone are available from a network of surface stations, mostly in the mid-latitude NH, with reasonable coverage extending back to the 1960s. Near-global, continuous column ozone data are available from satellite measurements beginning in 1979. The different observational data sets can be used to estimate past ozone changes, and the differences between data sets provide a lower bound of overall uncertainty. The differences indicate good overall agreement between different data sources for changes in column ozone, and thus we have reasonable confidence in describing the spatial and temporal characteristics of past changes (EC). Human activities have led to changes in the atmospheric concentrations of several greenhouse gases, including tropospheric and stratospheric ozone and ODSs and their substitutes. Changes to the concentrations of these gases alter the radiative balance of the Earth’s atmosphere by changing the balance between incoming solar radiation and outgoing infrared radiation. Such an alteration in the Earth’s radiative balance is called a radiative forcing. Positive radiative forcings are expected to warm the Earth’s surface and negative radiative forcings are expected to cool it. Changes in carbon dioxide (CO2) provide the largest radiative forcing term and are expected to be the largest overall contributor to climate change. In contrast with the positive radiative forcings due to increases in other greenhouse gases, the radiative forcing due to stratospheric ozone depletion is negative. Halocarbons are particularly effective greenhouse gases in part because they absorb the Earth’s outgoing infrared radiation in a spectral range where energy is not removed by carbon dioxide or water vapour (sometimes referred to as the atmospheric window). Halocarbon molecules can be many thousands of times more efficient at absorbing the radiant energy emitted from the Earth than a molecule of carbon dioxide, which explains why relatively small amounts of these gases can contribute significantly to radiative forcing of the climate system. Because halocarbons have low concentrations and absorb in the atmospheric window, the magnitude of the direct radiative forcing from a halocarbon is given by the product of its tropospheric mixing ratio and its radiative efficiency. In contrast, for the more abundant greenhouse gases (carbon dioxide, methane and nitrous oxide) there is a nonlinear relationship between the mixing ratio and the radiative forcing. Observed changes in stratospheric ozone Global and hemispheric-scale variations in stratospheric ozone can be quantified from extensive observational records covering the past 20 to 30 years. There are numerous ways to measure ozone in the atmosphere, but they fall broadly into two categories: measurements of column ozone (the vertically integrated amount of ozone above the surface), and measurements of the vertical profile of ozone. Approximately 90% of the vertically integrated ozone column resides in the stratosphere. There are more independent measurements, longer time-series, and better global coverage for column ozone. Regular measurements of column ozone are available from a network of surface stations, mostly in the mid-latitude NH, with reasonable coverage extending back to the 1960s. Near-global, continuous column ozone data are available from satellite measurements beginning in 1979. The different observational data sets can be used to estimate past ozone changes, and the differences between data sets provide a lower bound of overall uncertainty (Dvortsov). Figure 2 (a) Time-series of de-seasonalized global mean column ozone anomalies estimated from five different data sets, including ground-based (black line) and satellite measurements (colour lines). Anomalies are expressed as percentage differences from the 1964–1980 average, and the seasonal component of the linear trend has been removed. (b) Time-series of deseasonalized global mean lower-stratospheric temperature anomalies estimated from radiosondes (colour lines) and satellite data (black line). Anomalies (in °C) are calculated with respect to the 1960–1980 average. Observed changes in ODSs As a result of reduced emissions because of the Montreal Protocol and its Amendments and Adjustments, mixing ratios for most ODSs have stopped increasing near the Earth’s surface. The response to the Protocol, however, is reflected in quite different observed behaviour for different substances. By 2003, the mixing ratios for CFC-12 were close to their peak, CFC-11 had clearly decreased, while methyl chloroform (CH3CCl3) had dropped by 80% from its maximum. Halons and HCFCs are among the few ODSs whose mixing ratios were still increasing in 2000. Halons contain bromine, which is on average 40 to 50 times more efficient on a per-atom basis at destroying stratospheric ozone than chlorine. However, growth rates for most halons have steadily decreased during recent years. Furthermore, increases in tropospheric bromine from halons have been offset by the decline observed for methyl bromide (CH3Br) since 1998 (EC). Figure 3. Estimated global atmospheric mixing ratios (in ppt) of CFC-11, HCFC-22 and HFC-134a, shown separately for the NH (red) and SH (blue). Circles show measurements from the AGAGE (Advanced Global Atmospheric Gases Experiment) and CMDL (Climate Monitoring and Diagnostics Laboratory) networks, while colour lines show simulated CFC-11 concentrations based on estimates of emissions and atmospheric lifetimes (Pyle and Shepherd). Observed changes in stratospheric aerosols, water vapour, methane and nitrous oxide In addition to ODSs, stratospheric ozone is influenced by the abundance of stratospheric aerosols, water vapour, methane (CH4) and nitrous oxide (N2O). Observed variations in these constituents are summarized in this section. During the past three decades, aerosol loading in the stratosphere has primarily reflected the effects of a few volcanic eruptions that inject aerosol and its gaseous precursors (primarily sulphur dioxide, SO2) into the stratosphere. The most noteworthy of these eruptions are El Chichón (1982) and Pinatubo (1991). The 1991 Pinatubo eruption likely had the largest impact of any event in the 20th century, producing about 30 Tg of aerosol (compared with approximately 12 Tg from El Chichón) that persisted into at least the late 1990s. Current aerosol loading, which is at the lowest observed levels, is less than 0.5 Tg, so the Pinatubo event represents nearly a factor of 100 enhancement relative to non-volcanic levels. The source of the non-volcanic stratospheric aerosol is primarily carbonyl sulfide (OCS), and there is general agreement between the aerosols estimated by modelling the transformation of observed OCS to sulphate aerosols and observed aerosols. However, there is a significant dearth of SO2 measurements, and the role of tropospheric SO2 in the stratospheric aerosol budget – while significant – remains a matter of some uncertainty. Because of the high variability of stratospheric aerosol loading it is difficult to detect trends in the non-volcanic aerosol component. Trends derived from the late 1970s to the current period are likely to encompass a value of zero (Barnett, Houghton and Pyle). Conclusions The halogen loading of the stratosphere increased rapidly in the 1970s and 1980s. As a result of the Montreal Protocol and its Amendments and Adjustments, the stratospheric loading of chlorine and bromine is expected to decrease slowly in the coming decades, reaching pre-1980 levels some time around 2050. If chlorine and bromine were the only factors affecting stratosphere ozone, we would then expect stratospheric ozone to ‘recover’ at about the same time. Over this long (about 50-year) time scale, the state of the stratosphere may well change because of other anthropogenic effects, in ways that affect ozone abundance. For example, increasing concentrations of CO2 are expected to further cool the stratosphere, and therefore to influence the rates of ozone destruction. Any changes in stratospheric water vapour, CH4 and N2O, all of which are difficult to predict quantitatively, will also affect stratospheric chemistry and radiation (Pyle and Shepherd). In addition, natural climate variability including, volcanic eruptions, can affect ozone on decadal time scales. For these reasons, ‘recovery’ of stratospheric ozone is a complicated issue. The distribution of ozone depends on a balance between chemical processes, which can be affected by changes in the concentration of the ODSs, and transport processes, which can be affected by climate change. Climate change, in its broadest sense, can also affect ozone chemistry directly by modifying the rates of temperature-dependent reactions. Changes in stratospheric ozone can also affect climate. Changes in the structure of the stratosphere, caused by ozone changes, could alter the interaction between the troposphere and stratosphere and lead to further changes in stratospheric ozone (EC). Works Cited Allaart, Marc, et al. "Ozone mini-hole observed over Europe, influence of low stratospheric temperature on observations." Geophysical Research Letters (2000): 27, 4089–4092. Print. Barnett, J. J., J. T. Houghton and J. A. Pyle. "The temperature dependence of the ozone concentration near the stratopause." Quarterly Journal of the Royal Meteorological Society (1975): 101, 245–257. Print. Bird, Geoffrey. "PROTECTING THE OZONE LAYER." REFRIGERANTS (2002): 1, 3-14. Print. Chipperfield, Martyn. "A three-dimensional model study oflong-term mid-high latitude lower stratosphere ozone changes." Atmospheric Chemistry and Physics (2003): 3, 1253–1265. Print. Dvortsov, Victor. "Response of the stratospheric temperatures and ozone to past and future increases in stratospheric humidity." Journal of Geophysical Research – Atmospheres (2001): 106(D7), 7505–7514. Print. EC, European Commission. Ozone-Climate Interactions. Air Pollution Research Report No. 81, EUR 20623, European Commission. Luxembourg: Office for Official Publications of the European Communities, 2003. Print. Pyle, John and Theodore Shepherd. "Ozone and Climate: A Review of Interconnections." IPCC/TEAP Special Report: Safeguarding the Ozone Layer and the Global Climate System (2002): 1, 85-124. Print. Read More