NO ppt

FIGURE 6.7 Calculated 24-h average 03 production and loss rates for the free troposphere above Hawaii during the MLOPEX as a function of the NO* mixing ratio (Liu et al. 1992).

Figure 6.7 shows calculated, 24-hour-average production and loss rates for the free troposphere above Hawaii during the MLOPEX (Mauna Loa Photochemistry Experiment) as a function of NO* mixing ratio. The 03 loss rate is seen to be almost independent of NO*, at about 5 x 105 molecules cm-3 s_1. For an 03 mixing ratio of 40 ppb, this loss rate gives an 03 lifetime of 17 days. Data for upper tropospheric concentrations over Hawaii indicate that [NO*] is typically ~ 30 ppt, with midday [NO] at ~ 10 ppt (Ridley et al. 1992). From Figure 6.7 we see that at these levels 03 production and loss are just about in balance, with loss predicted to be slightly greater.

Ozone lifetimes in the troposphere vary significantly depending on altitude, latitude, and season. Lifetimes are shorter in the summer than in the winter as a result of the higher

solar fluxes in the summer. Lifetimes are also shorter at the surface because of the higher water vapor concentration near the surface. At higher latitudes, lifetimes increase because of the reduced solar intensity. At 20°N, for example, it is estimated that Oa lifetimes at the surface are about 5 days in summer and 17 days in winter, whereas at 403N these increase to 8 days in summer and 100 days in winter. At 20°N, at 10 km altitude, estimated summer and winter lifetimes are 30 days and 90 days, respectively, increasing by a factor of from those at the surface.

The Troposphere-Stratosphere Transition from a Chemical Perspective The transition from troposphere to stratosphere is traditionally defined on the hasis of the reversal of the atmospheric temperature profile. That transition is also dramatically reflected in how the concentrations of trace species vary with altitude both below - and above the tropopause. Of these, H02 and OH exhibit perhaps the most profound differences across the tropopause (Wennberg et al. 1995). The major difference between the stratosphere and the troposphere is what controls the interc on version between OH and H02 within the HO.r family. In the lower stratosphere the cycle within the HOA family is

In the lower stratosphere the H02/0H ratio is described by the extension of (5.28):

This ratio varies from about 4 to 7 and decreases as [NOJ increases. [OH] Itself is essentially independent of [NO] and depends almost entirely on solar zenith angle. This independence of OH on NO is a result of the fact that the increase of OH that results from the H02 + NO reaction is offset almost exactly by a decrease of the rates of reactions that generate OH. This occurs because the HO; that participates ¡n the H02 4- NO reaction is not otherwise available for other reactions.

Whereas OH to HOz conversion in the lower stratosphere occurs mainly by OH + 03, that in the upper troposphere occurs mainly by OH + CO (Lanzendorf et al. 2001),

The HO2/OH ratio in the upper troposphere can be approximated as

As one proceeds up in the troposphere, the N02/N0r ratio decreases (recall Section 6.6.1), achieving its lowest value at the tropopause, and then increases on moving into the stratosphere. The increase of N02 relative to NO in the lower stratosphere is the result of the HO2 + NO reaction. {The NOi(/NOv ratio is more or less constant in the upper troposphere, falling off as one goes into the stratosphere. This failoff reflects the influence of the stratospheric aeroso! layer in promoting the heterogeneous formation of HN0>) Hydroxy! radical levels in the upper troposphere vary from ~0,01 to 0.1 ppt. In the lower stratosphere OH depends on solar zenith angle and ranges up to ~ 1 ppt. As noted above, OH is essentially independent of NO in the lower stratosphere, whereas in the upper troposphere OH dccrcases as NO decreases. This fundamentally different behavior of OH with respect to changes in NO also characterizes the troposphere-stratosphere transition.

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