N2o O1d No No

Model studies imply a total stratospheric sink of around 12.6 TgN y-1 (Prather et al. 2001), which, combined with the above rate of atmospheric accumulation, suggests an approximately balanced N2O budget. As for CH4, however, the uncertainties in individual source and sink estimates are large, and hence the task of defining (for example) the precise reasons for the observed steady rise in N2O remains to be accomplished.

Halocarbons and SF6

Current concerns about the atmospheric trends in chlorofluorocarbons (CFCs), hydrochlorofluorocarbons (HCFCs), hydrofluorocarbons (HFCs), hydrochlorocar-bons (HCCs), chlorocarbons, perfluorocarbons (PFCs), and sulfur hexafluoride (SF6) are based on either their deleterious effects on the ozone layer (where many of them are sources of destructive Cl and ClO), and /or their radiative forcing of climate. Their GWPs are generally extremely large (Table 9.1), so that even very small emissions of these gases relative to CO2 can have a nonnegligible climate impact.

The budgets of this wide-ranging set of gases, whose sources with few exceptions are almost exclusively industrial, have been recently reviewed (Kurylo and Rodriguez 1999; Prinn and Zander 1999; Prather et al. 2001; Montzka and Fraser 2002), and there is not space here to discuss all of the details. The total amount of chlorine contained in all the major chlorine-containing halocarbons peaked at 3.7 parts per billion (ppb) at the surface around 1992—1994 owing to the decreases in emissions of these gases mandated by the Montreal Protocol for the Protection of the Ozone Layer (Montzka et al. 1999; Prinn et al. 2000). Presuming these decreases continue in the future, attention to climate impacts and regulation of these gases in the Kyoto Protocol of the FCCC will increasingly be on halogen-containing gases not covered by the Montreal Protocol (specifically HFCs, PFCs, and SF6; see Tables 9.1 and 9.2).

The major sources for all of these halogen-containing gases are anthropogenic. They are used primarily as refrigerant fluids, foam-blowing agents, or solvents. The primary sinks for CFCs, PFCs, and SF6 are in the upper atmosphere (photodissociation, ionospheric reactions), while HFCs and HCFCs are destroyed by reaction with OH in the troposphere as well as being photodissociated in the stratosphere.

The most abundant HFCs and PFCs are, respectively, CHF3 (a byproduct of production of the refrigerant CHF2Cl) and CF4 (a byproduct of aluminum production). Also rapidly rising are the HFC species CF3CH2F (a common refrigerant), the PFC compound C2F6 (aluminum industry), and SF6 (leaked from transformers, etc.). The latter gas, with one of the highest known GWPs (22,200) and a 6 percent y-1 rate of increase, is indicative of the importance of this set of non-CO2 gases.

Nonmethane Hydrocarbons, CO, NOx, and O3

The ability of the troposphere to chemically transform and remove trace gases depends on complex chemistry driven by the relatively small flux of energetic solar ultraviolet radiation that penetrates through the stratospheric ozone layer (see, e.g., Ehhalt 1999). This chemistry is also driven by emissions of NO, CO, and hydrocarbons (RH) and leads to the production of ozone, which is a potent greenhouse gas. Globally, about 3.4— 4.6 Pg y-1 of ozone are produced in the troposphere with a lifetime of 28—37 days (Ehhalt 1999; Lelieveld and Thompson 1999; Prinn 2003). The most important cleansing chemical in the troposphere, however, is the hydroxyl free radical (OH), and a key measure of the capacity of the atmosphere to oxidize trace gases injected into it is the local concentration of hydroxyl radicals (Prinn 2003).

As noted earlier, the OH radical serves as the principal sink for the greenhouse gases CH4, HFCs, and HCFCs. It removes about 3.65 Pg y-1 of trace gases from the atmosphere (Ehhalt 1999; Prinn 2003). The principal catalyzed reactions creating OH and O3 in the troposphere are:

O3 + ultraviolet S O2 + O(1D) O(1D) + H2O s 2OH Net effect: O3 + H2O S O2 + 2OH and

OH + RH s R + H2O R + O2 s RO2 RO2 + NO s RO + NO2 NO2 + ultraviolet s NO + O O + O s O,

In theoretical models, the global production of tropospheric O3 by the pathway beginning with CO is typically about twice that beginning with RH. For most environments, these catalytic processes pump the majority of the OH and O3 into the system. If the concentration of NO2 gets too high, however, its reaction with OH to form HNO3 ultimately limits the OH concentration.

How stable is the cleansing capability of the troposphere? If emissions of gases that react with OH, such as CH4, CO, and SO2, are increasing then, keeping everything else constant, OH levels should decrease. Conversely, increasing NOx emissions from combustion should increase tropospheric O3 (and thus the primary source of OH), as well as increase the recycling rate of HO2 to OH (the secondary source of OH). Also, if the oceans are increasing in temperature, one would expect increased water vapor in the lower troposphere. Because water vapor is part of the primary source of OH, climate warming also increases OH. This increase could be lowered or raised if changes in cloud cover accompanying the warming lead to more or less reflection of ultraviolet back toward space. Rising temperature also increases the rate of reaction of CH4 with OH, thus lowering the lifetimes of both chemicals. Opposite conclusions apply if trace gas emissions increase or the climate cools. Finally, decreasing stratospheric ozone can also increase tropospheric OH, as discussed later in this chapter. Determining whether OH concentrations, and hence the cleansing capacity of the atmosphere, are changing is a major focus of recent research (Thompson 1992; Wang and Jacob 1998; Prinn 2003).

The ozone in the stratosphere is maintained by a distinctively different set of chemical reactions than in the troposphere. Stratospheric ozone is produced and destroyed at a massive global rate of about 120 Pg y-1 (Warneck 1988). The fundamental stratospheric ozone source is photodissociation of O2. The ozone sink is driven by chemicals like water vapor, chlorofluorocarbons, and nitrous oxide that are transported from the troposphere to the stratosphere, where they produce chlorine-, nitrogen-, and hydrogen-carrying free radicals, which catalytically destroy ozone.

There is an important link between stratospheric chemistry and climate. The precursor gases for the destructive free radicals in the stratosphere are greenhouse gases, as is stratospheric ozone. Therefore, while increases in the concentrations of the source gases will increase radiative forcing of warming, these increases will, at the same time, lead to decreases in stratospheric ozone, lowering the radiative forcing. As in the troposphere, there are important feedbacks between chemical processes and climate, through the various greenhouse gases. It is not just manmade chemicals that can change stratospheric ozone. Changes in natural emissions of N2O and CH4, and changes in climate that alter H2O flows into the stratosphere, undoubtedly led to changes in the thickness of the ozone layer in the past.

There is also a strong link between the thickness of the stratospheric ozone layer and concentrations of OH in the troposphere. Because O(1D) can only be produced from O3 at wavelengths less than 310 (nanometers) nm, and the current ozone layer effectively absorbs all incoming ultraviolet at less than 290 nm, a very narrow window of

290—310 nm radiation is driving the major oxidation processes in the troposphere. Decreasing stratospheric ozone, as in recent decades, can therefore increase tropos-pheric OH by increasing the flux of radiation less than 310 nm reaching the troposphere to produce O(1D) (Madronich and Granier 1992).

Global Warming Potentials

As mentioned earlier, GWPs provide a tool for ranking the potential for a given mass of an emitted greenhouse gas to influence climate, relative to the same emitted mass of CO2. Formally the GWP is defined as follows:

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