Chapter 2 | SCIENTIFIC BACKGROUND |
Clear- and cloudy-sky fluxes. Long-wave flux (also referred to as irradiance) is
represented by the letter F. The flux for clear-sky conditions is denoted by Fa, where the
subscript a represents the atmosphere; the flux for overcast cloud conditions is Fc. The
average flux for a mixture of clear and cloud conditions is F. Thus, if a region has a
fraction of area, C, covered by clouds, then F = (1-C) x Fa + C x Fc, where the fraction (1-
C) is the clear-sky fraction.
Up and down long-wave fluxes. The long-wave flux emitted in the upward direction is
denoted by the symbol F+, and the flux in the downward direction is denoted by the symbol
F-. The net flux, then, is F = F+ - F.
Before describing the role of the five factors upon which the thermostat hypothesis is based,
several terms used by Ramanathan and Collins will be reviewed.
Atmospheric greenhouse effect. The definition adopted by the global warming community
(IPCC, 1990), gives the atmospheric (cloud-free) greenhouse effect (Ga) as Ga = F+(Z=0)
- Fa+, where F+(Z=0) is the blackbody emission from the sea surface, and Fa+ is the
outgoing long-wave radiation under clear skies, as obtained from the ERBE satellite data.
Without the atmosphere the OLR would be ÓT4, where T is the surface temperature. The
atmosphere, primarily water vapor and CO2, absorbs most (70 to 95%) of the surface long-
wave radiation and re-emits to space at the much colder temperature of the atmosphere.
The effect is to reduce OLR. Ga then represents the energy trapped by the entire
atmospheric column between the surface and the top-of-the-atmosphere. The atmosphere
then heats the surface by emitting absorbed energy back to the surface. This downward
emission becomes another measure of the greenhouse effect (e.g., Vonder Haar, 1986). In
a system in radiative-convective equilibrium, both methods are self-consistent. However,
when used in a regional context, care should be taken to keep this distinction and to
measure the greenhouse effect using both methods.
Ga denotes the greenhouse effect of the entire troposphere. However, with the aid of
aircraft observations, it is possible to measure the vertical structure of Ga. This is denoted
by Ga (Z), which is estimated from Ga(Z) = F+(Z=0) - Fa+(Z). Ga(Z) is the trapping by
the atmosphere between the surface and the altitude, Z. At any altitude, Z, in the tropics,
the temporal or spatial variations in Ga(Z) are mostly due to corresponding variations in
water-vapor concentration and temperature profile. This is because temporal and spatial
variations in other radiatively active gases are negligible compared with those of water
vapor.
Cloud radiative forcing. Cloud radiative forcing denotes the effect of clouds on the
surface-atmosphere column radiation budget or on the surface energy budget. Clouds
reduce the long-wave energy emitted to space (i.e., OLR). Cloud long-wave forcing (Cl)
for the surface-troposphere column is obtained from Cl = Fa+ - F+, where the upfluxes
should be estimated at tropopause altitudes or at the top-of-the-atmosphere. Cl is the
enhancement of long-wave trapping due to the presence of clouds. The total greenhouse
effect combines the effect of the atmosphere with the effect of the clouds, so that G = Ga +
Cl = F+(Z=0) - F+.
Clouds also enhance the downward long-wave energy at the surface. This enhancement,
or cloud long-wave forcing at the surface, is given by Cl(Z=0) = Fa - Fa-. Thus, Cl(Z=0)
is the effect of clouds on the surface long-wave radiation heating; Cl is the effect of clouds
on the surface-troposphere long-wave radiation heating; and the difference, Cl - Cl(Z=0), is
the effect of clouds on the long-wave heating of the troposphere.
In addition, clouds enhance the planetary albedo, reducing the solar radiation absorbed by the
surface-atmosphere system. This effect, cloud short-wave forcing, is Cs = S - Sa, where S is the
solar radiation absorbed by the surface-atmosphere column for cloudy skies (i.e., the average of
clear plus overcast skies) and subscript a refers to clear skies. Cloud radiative forcing is the net of
the long-wave and short-wave effects.
Let us now consider the five factors involved in the thermostat hypothesis.
Deep convection threshold temperature (Tc). The thermostat mechanism operates in
regions of deep convective clouds, with bases at between the surface and 3 km and tops at between
12 and 18 km. The horizontal scale of an individual, non-entraining cell is anywhere from a few to
several tens of kilometers. Such clouds typically have an anvil that begins as a thick (5Ð8 km)
cloud deck near the convective region, then thins to a few kilometers or less at distances that are
sometimes several hundreds of kilometers from the convective source (top panel of Figure 10).
Systems of deep cumulonimbus-cirrus anvil clouds, referred to as super cloud clusters (Nakazawa,
1988), can grow to a size as large as 106 km2 in a matter of a day or less (e.g., WMO, 1972;
Houze and Betts, 1981). The radiative effect of these clusters is dominated by the anvils and
stratiform cirrus (bottom panel of Figure 10), because of their spatial extent. The albedo of these
cloud systems can reach as high as 60Ð80% (Harrison et al., 1991), and they can reduce OLR by
as much as 200 Wm-2.
These convective systems are triggered over warm oceans only when the SST > 300 K (Tc)
(Gadgil et al, 1984; Graham and Barnett, 1987), where the subscript c represents convection.
Satellite data of OLR clearly reveal this threshold effect (Figure 11 shows highly reflective clouds
[HRC] and OLR in relation to SST). It is important to note that SST > Tc is a necessary, but not
sufficient, condition for convection. In addition, convergence of moisture in the boundary-layer is
required to initiate convection (Holton, 1992). In regions with SST > Tc, and where deep
convection persists, the horizontal gradients in SST diminish. The maximum SST remains within
a few degrees of Tc.
Super-greenhouse effect. The phrase super-greenhouse effect was first used by Vonder
Haar (1986). From observations of net (upward minus downward) long-wave flux at the surface
of the tropical Atlantic Ocean during GATE (GARP [Global Atmospheric Research Program]
Atlantic Tropical Experiment), Vonder Haar noted that the net long-wave flux, for clear as well as
for cloudy skies, decreased with an increasing SST (i.e., as SST increases, the downward
emission to the surface increases at a rate faster than the increase of upward emission from the surface). The super-greenhouse effect is defined, then, as:
Fa - (Z=0) / SST > F + (Z=0) / SST, where F+(Z=0)= Ó (SST)4.
The ERBE satellite data used by Ramanathan and Collins (1991) revealedÑfor the tropical oceans
in general and over the warm oceans (SST > Tc) in particularÑthat the rate of increase in Ga with
increasing SST indeed exceeds the rate of increase of F+(Z=0), the blackbody emission by the sea
surface, ¶Ga / ¶SST > ¶F+(Z=0) / ¶SST. Ramanathan and Collins deduced this super-greenhouse
behavior from analyses of spatial, seasonal, and interannual variations of Ga and SST. In
addition, the cloudy sky (G) also increased significantly with SST, so that OLR decreases with an
increase in SST (Figure 11). The coupling between SST and the atmospheric greenhouse effect
represents a potentially unstable feedback loop, especially when cloud long-wave effects are taken
into account (Figure 12). The total greenhouse effect, G = Ga + Cl, increases exponentially when
SST > 300 K. The upper troposphere cirrus clouds contribute the most to this increase. Because
tropical cirrus result from deep convection, and because deep convection increases with SST, a
strong negative feedback is required to limit maximum SSTs and counter the super-greenhouse
effect.
Radiative effects of cirrus and the f factor. The large magnitude of short-wave cloud
forcing over the warm pool, as revealed by the ERBE data, indicates that the surface-atmosphere
system in the warm pool absorbs significantly less solar radiation than that in the colder oceans
(Figure 13) in the tropics. Ramanathan and Collins hypothesize that the oceans become warmer,
and convection gets deeper, until the cirrus become sufficiently thick to reduce solar energy
reaching the oceans and prevent their further warming. The large changes in planetary albedos
observed in the central Pacific during the 1982 El Niño (Chertock et al., 1991) and during the 1987
El Niño (Figure 14) support this idea. That a negative feedback between surface temperature and
cloud reflection of solar radiation may regulate SSTs had been suggested earlier by Monin (1972)
and Sikka and Gadgil (1980) to explain ocean-atmospheric oscillations in temperature, cloudiness,
and rainfall; by Graham and Barnett (1987) to explain the maximum SSTs; and by Barnett et al.
(1991) to explain El Niño.
However, there is an important detail that was not addressed by these earlier studies. This
concerns the long-wave-trapping effect of cirrus. As mentioned earlier, there is a near cancellation
between short-wave and long-wave cloud forcing. If the warm pool were in strict radiative-
convective equilibrium, the net effect of anvils on warm pool SST would be small. For the cirrus
thermostat to be effective in arresting the SST increase, the long-wave energy trapped by cirrus has
to be exported away from the warm pool. Ramanathan and Collins assume that between 80 and
100% of the trapped energy is exported away, and this assumption is based upon the so-called f
factor. One of the key assumptions in their model is that the atmospheric long-wave trapping (Ga)
warms the surface directly, whereas most of the cirrus long-wave trapping contributes to the
radiative heating of the troposphere. They describe this as the f factor, where f = Cl(Z=0) / Cl.
Recall that Cl is the cloud radiative forcing of the surface-troposphere column, and Cl(Z=0) is that
of the surface; i.e., f describes that fraction of the long-wave cloud forcing that arrives at the sea
surface.
The following example is given to clarify the terminology. Consider a 50% cloud fraction, and the
clear-sky downward flux at the surface is Fa-(Z=0) = 400 Wm-2; underneath the cloud deck, the
downward flux is Fc- = 440 Wm-2. The flux for the mixture of clear plus overcast skies is then
F+(Z=0) = (400 x .5) + (440 x .5) = 420, and the cloud forcing at the surface is Cl(Z=0) = 420 -
400 = 20 Wm-2. Thus, long-wave cloud forcing is simply the change in radiative heating due to
clouds. Now, if the corresponding values for upflux at the tropopause altitudes are Fa = 300 Wm-2
and Fc = 200 Wm-2, then F = 250 Wm-2 and Cl = 50 Wm-2. The factor for this case is f = 20 / 50
= 0.4. This implies that, out of the surface-troposphere total heating (due to clouds) of 50 Wm-2,
the surface heating is 20 Wm-2, and the balance of 30 Wm-2 is the net long-wave energy heating the
tropospheric column.
Ramanathan and Collins assumed that f < 0.2 for the warm-pool region, because they asserted that
water-vapor opacity is large in such regions and that clouds do not influence the long-wave energy
at the surface. For example, if the water vapor is optically thick in the entire long-wavelength
region, then Fc-(Z=0 ) ~ Fa-(Z=0) and Cl(Z=0) ~ 0 , resulting in f ~ 0. If the f factor for cirrus is
close to 1, the cirrus by itself will not be effective in limiting SSTs.
Rapid heat advection in the troposphere. A remarkable feature of the tropical atmosphere
is the near-horizontal homogeneity of atmospheric temperature above the boundary layer (see
Figures 1 and 2 in Wallace, 1992). Between about 750 mb and 150 mb, temperature varies by less
than five degrees across an entire latitudinal circle of about 40,000 km, as seen, for example, in
Figure 15 for the 200-mb level. This is remarkable, because heat sources and sinks in the tropics
have significant longitudinal and latitudinal asymmetry. The temperature homogeneity is due to
transport by the Walker and Hadley circulations and planetary-scale wave disturbances. This
suggests that atmospheric motions and disturbances in the tropics are significantly efficient in
exporting heat away from regions of heating and maintain a spatially uniform temperature profile
(Wallace, 1992).
A unique feature of tropical cirrus is that the long-wave cloud effect leads to a radiative heating of
the troposphere, with very little direct effect on the surface. This result (based largely upon model
studies), when considered in conjunction with the efficient transport in the troposphere, leads to
the deduction that the long-wave heating effect of cirrus is exported away from the warm-pool
region. Thus, the only local effect of cirrus is the shielding of the sea surface from solar radiation.
The magnitude of this effect is large (see Ramanathan and Collins, 1991, ERBE data) and can be in
the range of 75125 Wm-2 in the monthly mean. This represents about 3050% of the total solar
heating of the oceans under clear-sky conditions. The question remaining is, can SST respond to
this reduction locally? The answer to this question depends upon the efficiency of dynamic heat
transport in the oceans.