Impact of the stratosphere on solar radiative forcing of climate

5.5 Impact of the stratosphere on solar radiative forcing of climate

In Section 3.2 it was noted that the most accurate estimates of climate radiative forcing are those which take into account the effect of changes in the stratosphere on the radiative flux at the tropopause. Thus a better estimate of radiative forcing to solar irradiance changes should incorporate the effects of the influence of variations in UV on stratospheric temperature and composition (as first noted by Haigh, 1994 ). Figure 31

(above) shows solar irradiance at winter mid-latitudes as a function of wavelength and altitude calculated using a 2D atmospheric model with fairly complete representations of photochemical and radiative processes. At the top of the atmosphere most energy is at visible wavelengths and this is transmitted almost unaffected through to the surface. At wavelengths shorter than about 300 nm, however, most radiation is bsorbed by the time it reaches an altitude of around 40 km. There is also some absorption at longer visible wavelengths.

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Figure 31:

(Above) Solar spectral irradiance as a function of wavelength and altitude for 57 N at the winter solstice. Values range from $<$ 10^–4 (violet) to $>$ 6 (red) MWm^–2 cm^–1. (Below) as above but showing the difference between values at maximum and minimum periods of the 11-year solar cycle. Contour interval is 500 Wm^–2 cm^–1 ranging from $<$ –1500 (violet) to $>$ 2500 (red). After Haigh (1994 ).

Table 3:

A summary of published estimates of solar radiative forcing. 1st column: reference; 2nd: nominal solar variability; 3rd and 4th: solar UV radiative forcing at the top of atmosphere and at the tropopause; 5th solar-induced ozone change; 6th, 7th and 8th: impact of ozone change on shortwave and longwave components of radiative forcing and the net effect; 9th: percentage amplification of solar forcing due to change in ozone.


Author	Solar change	$Δ$ S RF (toa)	$Δ$ S RF (tpse)	$Δ$ O₃	O₃ SW effect	O₃ LW effect	Net O₃ effect	Amplification (%)

Haigh (1994 )	11-year amp	0.13	0.11	+ve peak near 40 km	–0.03	+0.02	–0.01	–9

Hansen et al. (1997 )	11-year amp	0.13	0.11	+ve 10 – 150 hPa			+0.05	+45

Myhre et al. (1998 )	11-year amp	0.13	0.11	+ve	–0.08	+0.06	–0.02	–18

Wuebbles et al. (1998 )	c1680 – c1990	0.49 to 0.70	0.42 to 0.60	+ve peak near 40 km			–0.13	–30 to –21

Larkin et al. (2000 )	11-year amp	0.13	0.11	+ve (as Haigh, 1994 )	–0.06	+0.11	+0.05	+45
		0.13	0.11	+ve (SBUV/ TOMS)	–0.03	+0.08	+0.05	+45

Shindell et al. (2001 )	1680 – 1780	0.30 to 0.39	0.26 to 0.33	–ve (upper strat)			+0.02	+6 to +8

The lower panel of Figure 31 shows the difference in spectral irradiance between maximum and minimum periods of the 11-year solar cycle. At the top of the atmosphere there is more energy at all wavelengths but this is not perpetuated throughout the depth of the atmosphere. At wavelengths $<$ 330 nm and $>$ 500 nm there is actually less radiation reaching the troposphere at solar maximum than solar minimum because the enhanced concentrations of stratospheric ozone are resulting in greater absorption at these wavelengths. This is a strongly non-linear effect which varies with latitude and season and thus its impact on the value of solar radiative forcing is not easy to predict. Estimates of the net effect of solar-induced ozone increases on solar radiative forcing vary widely, as can be seen in Table 3: even the sign of the ozone effect is not ascertained.

To calculate radiative forcing a knowledge of changes to the temperature of the stratosphere are also required, adding a further complication. The next section considers how well the effects of solar ultraviolet variability on the thermal structure of the middle atmosphere are understood.