Impact of Dust Radiative Forcing upon Climate - Mineral Dust: A Key Player in the Earth System - page 339

at the surface and TOA is altered. At the surface, the radiative and turbulent fluxes
adjust until F S is fully compensated. Only if heat transport by the ocean circulation
adjusts to the aerosol forcing can the net energy flux across the surface remain
nonzero.
At TOA, however, an imbalance between the forcing and anomalous OLR
may persist even after the climate returns to equilibrium. Consider the example
of negative forcing at TOA due to a layer of reflective aerosols. To balance this
reduction of insolation, OLR and temperature decrease, and because of dynamical
adjustment, this decrease generally extends beyond the forcing region. The net
radiative anomaly at TOA is positive beyond the aerosol layer due to the decrease of
OLR. Conversely, within the region of aerosol forcing, the net radiative anomaly at
TOA is negative because this anomaly must be zero when summed over the extent
of the response, as expressed by ( 13.1 ). Thus, above the aerosol layer, the reduction
of OLR cannot fully compensate the forcing. The mismatch between the regional
extent of the forcing and the equilibrium response creates a horizontal contrast in
the TOA energy flux that must be balanced by the atmospheric transport of energy.
In our example, the continual net radiative loss at TOA above the aerosol layer is
balanced by anomalous energy import.
The coupling of TOA forcing and anomalous atmospheric energy transport can
lead to regional anomalies of precipitation. Here, we discuss precipitation anomalies
over the Sahel and their relation to TOA forcing in this region (Figs. 13.1 and 13.7 ).
Following an argument by Chou et al. ( 2005 ), we consider the anomalous moisture
budget for the convecting region:

ıP c D ıE c C ıM.q S;d q T;c / C Mı.q S;d q T;c /; (13.7)

where ıP c and ıE c are the anomalous precipitation and evaporation within the
convecting region, respectively, and ıM is the circulation anomaly linking the
convecting region to the neighboring regions of subsidence. The quantity q S;d

q T;c
is the difference between the unperturbed specific humidity entering the convecting
region from the surface of the subsiding region (q S;d ) and the upper tropospheric
humidity q T;c where the rising air detrains (Fig. 13.8 a). Equation ( 13.7 )saysthat
anomalous precipitation ıP c exceeds the local evaporative anomaly ıE c as a result
of two contributions to the anomalous moisture flux. The first is the product of the
anomalous overturning ıM and the net inflow of unperturbed moisture .q S;d q T;c /

into the column. The second is the unperturbed circulation M multiplied by the
anomalous net inflow ı.q S;d q T;c /.

Negative TOA forcing within a region of climatological convection must be
balanced by a decrease in energy export. Chou et al. ( 2005 ) assume that this
decrease occurs due to weakening of the direct circulation (ıM < 0), neglecting
a possible contribution by a changing contrast in moist static energy between the
convecting and subsiding regions. They also assume that the primary contribution
to anomalous precipitation is by the anomalous moisture flux resulting from the
weakened overturning (the penultimate term in Eq. 13.7 ). Thus, negative TOA