Reduction of Tropical Cloudiness by Soot
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V. Ramanathan et al. Indian Ocean Experiment (INDOEX) A Multi-Agency Proposal for a Field Experiment in the Indian Ocean (C 4 Scripps Institution of Oceanography UCSD La Jolla CA 1996) p. 83.
Taylor J. P., Ackerman A. S., Q. J. R. Meteorol. Soc. 125, 2643 (1999).
S. K. Satheesh and V. Ramanathan [ Nature 405 60 (2000)] report direct radiometric observations at the surface and top-of-atmosphere over Kaashidhoo a Maldivian island in the Indian Ocean obtained during February and March 1998. According to their observations the diurnal-average aerosol-induced solar absorption was ∼9 W/m 2 for an average aerosol optical depth of 0.2 (at 0.5 μm). They also report a range of 0.88 to 0.90 for measurements of the aerosol single scattering albedo at 0.5 μm (which take into account the effects of relative humidity on particle size).
Satellite retrievals of monthly diurnal-average low-cloud fractional coverage during February over the years 1989 to 1993 range from ∼0.2 to 0.3 at the equator to ∼0.1 to 0.2 at 10°N [
]. METEOSAT-5 retrievals by VR yield comparable values for 1999.
A. S. Ackerman and O. B. Toon [ Nature 380 512 (1996)] artificially increased the solar heating within cloud droplets to explore the implications of enhanced solar absorption by clouds (as reported by references therein). The additional solar absorption required to match the observed enhancement at noon was 125 W/m 2 equivalent to a diurnal average of ∼50 W/m 2 .
B. Stevens et al. 13th Symposium on Boundary Layers and Turbulence 10 to 15 January 1999 Dallas TX (American Meteorological Society Boston 1999) 269–270. The simulations here are derived from a model intercomparison setup described therein.
The dynamics model [
] solves the anelastic Navier-Stokes equations in conservative form using 5 s time steps on a domain spanning 6.4 km by 6.4 km horizontally and 3 km vertically which is uniformly discretized into 32 × 32 × 75 grid cells. The boundary conditions are doubly periodic in the horizontal and rigid at the top and bottom. Surface fluxes are parameterized through similarity relations [
]. A sponge layer at the top of the model dampens trapped buoyancy waves at altitudes >500 m above the trade inversion which is defined as the horizontal-average height where the total water mixing ratio is 6.5 g of water per kg of air. First-order turbulence closure is used for subgrid-scale mixing [
] with a stability-dependent mixing length [
] modified to account for the effects of evaporation [
]. Large-scale subsidence is calculated as the product of the divergence rate (assumed constant) and altitude. Subsidence and radiative forcings are linearly attenuated to zero in the 300 m above the trade inversion to prevent drift of the overlying atmospheric properties due to any imbalanced forcings.
Radiative transfer is calculated for each column every 2.5 min using a two-stream model [
] in which water vapor continuum absorption has been modified [
]. Cloud water from the dynamics model is fit to a log-normal droplet size distribution with a prescribed number concentration and a geometric SD of 1.5. Aerosol and cloud optical properties are computed through Mie calculations [
] using the complex index of refraction for liquid water compiled by A. S. Ackerman O. B. Toon P. V. Hobbs [ J. Atmos. Sci. 52 1204 (1995)]. We specify an overlying water vapor column of 1 g/cm 2
the ocean surface albedo is assumed to be independent of wavelength and is determined from the wind speed at 10 m (averaging ∼8 m/s) and the solar zenith angle using the parameterization of J. Hansen et al. [ Mon. Weather Rev. 111 609 (1983)].
We generally follow (15) for surface properties and initial soundings which are based on the measurements and average position of the upstream ship (the R/V Planet ): sea surface temperature fixed at 298 K (16) surface pressure fixed at 1015 mbar (1 mbar = 1 hPa) (16) latitude fixed at 15°N geostrophic (and initial) winds from figure 4 of (19) and soundings of temperature and water vapor mixing ratio averaged by B. Albrecht to preserve the jumps at the base and top of the transition layer. The large-scale divergence rate is fixed at 4.2 × 10 −6 s −1 the mean value derived from the observations (16). Pseudo-random perturbations of temperature and water vapor mixing ratio are imposed below the inversion to promote turbulence initially; the amplitudes of the perturbations (which horizontally average to zero) are 0.1 K and 0.025 g/kg respectively (15).
We justify our assumption of a fixed sea surface temperature on the basis that its diurnal range is negligible: for our simulations (with maximum insolation of 950 W/m 2 an average precipitation rate of ∼0.1 mm/day and 10-m wind speed of ∼8 m/s) the parameterization of P. J. Webster C. A. Clayson J. A. Curry [ J. Clim. 9 1712 (1996)] yields an amplitude of only 0.2 K for the diurnal variation in sea surface temperature.
Advective fluxes for the ATEX triangle were calculated by (19) from which advective forcings can be computed (as small differences between large terms): in the lower 60 mbar of the atmosphere a drying tendency of 1.4 g/kg/day and a cooling tendency of 0.9 K/day are indicated. Following (15) we parameterize these forcings to fade linearly from a maximum at the surface to zero at the trade inversion. The surface drying tendency we use (1.5 g/kg/day) is taken directly from (15). We double the surface cooling tendency of (15) to 2 K/day which compensates for a diurnal-average clear-sky cooling rate (1.7 K/day infrared cooling 0.7 K/day solar heating) that is half the value imposed in (15).
In the real atmosphere changes in aerosol-induced heating rates (tending to decrease cloud coverage) are linked to changes in cloud droplet concentrations (tending to increase cloud coverage) through microphysical details of the aerosol population (specifically their chemical composition and size distribution). Any net effect of these opposed tendencies depends on such microphysical details as well as the meteorology. Rather than attempt to comprehensively evaluate any net effects we instead decouple the forcings by varying the haze properties and cloud droplet concentrations separately.
The model domain is initially cloudless; the cited droplet concentrations apply only to grid cells in which clouds appear. In all simulations the number concentration of haze particles increases linearly from zero at the ocean surface up to 600 m maintains a uniform value up to the trade inversion and vanishes linearly in the overlying 300 m. The haze particle size distribution is log-normal with a geometric mean radius of 0.1 μm and a geometric SD of 1.8.
For clear-sky conditions at 5°N on March 1 a 1.8-km-deep layer of our idealized INDOEX 1998 haze absorbs 7.4 W/m 2 of solar radiation (diurnally averaged) which is comparable to the absorption measured during INDOEX 1998 (10). Optical properties of black carbon (soot) are taken from [
To address the sensitivity of the simulations to small variations in initial conditions for the baseline we ran an ensemble of four simulations that differed only in the pseudo-random distribution of initial perturbations of temperature and water vapor. Output from one member of the ensemble is shown in Figs. 2 and 3.
Vertically resolved cloud fraction defined as the fraction of cells in each layer with cloud water >0.05 g/kg is distinct from the fractional cloud coverage we define subsequently which is evaluated from vertically integrated columns.
During the first 5 days of ATEX (the observation period upon which our meteorology is based) surface reports of the cloud fractional coverage which may include clouds above the boundary layer ranged from an early morning maximum ∼0.9 on two separate days to a minimum of ∼0.2 one afternoon; the average value over the period was 0.5 [
]. For comparison the simulated cloud coverage depends on a number of factors including the criterion used to count cloudy grid columns the model resolution subgrid-scale mixing assumptions and the water vapor above the inversion. With our model setup the diurnal average (0.23) and range (0.1 to 0.4) from the baseline simulations are more comparable to those found over the Indian Ocean during the northeast monsoon (11).
We assume that all the enhanced solar absorption occurs only within the haze though some fraction of the soot is likely to be incorporated into cloud droplets through nucleation and coagulation/coalescence. Cloud droplets (of typical radius 10 μm) collect significantly more sunlight than do haze (of typical radius 0.1 μm) and hence a fixed amount of soot will absorb more sunlight when embedded within cloud droplets (particularly when most of the haze lies below the bulk of cloud cover); such an effect could be expected to increase the impact of the soot on boundary-layer dynamics. Yet this expectation is not borne out by simulations in which all the soot is assumed to be within cloud droplets (when present within a grid cell) because such a small volume of boundary-layer air is occupied by cloud in our simulations.
“Cloud-burning” is the response of clouds to increased atmospheric heating which includes reductions in cloud coverage and liquid and water path.
The reductions in cloud coverage due to solar absorption in the INDOEX hazes are not completely independent of droplet concentration. The strong increase in daytime cloud coverage between droplet concentrations of 50 and 100 cm –3 only for the simulations without soot implies that the effect of soot on cloud cover is maximum at a droplet concentration of 100 cm –3 . However the departure of that maximum from the average cloud-burning effect is not significant compared to the noise found for the baseline ensemble.
As seen in Fig. 5B the average liquid water path is roughly independent of droplet concentration for any particular aerosol. The increases of cloud coverage with droplet concentration in these simulations (Fig. 5A) are largely due to enhancement of total droplet cross-sectional area and therefore optical depth (which vary as the cube root of droplet concentration holding liquid water path fixed). Because cloud coverage is defined as the fraction of columns exceeding an optical depth threshold (2.5) columns do not need as much liquid water path at increased droplet concentrations to be counted as cloudy.
Simulations with moisture enhanced above the inversion layer (at 6 g/kg up from 4.5 g/kg used in our other simulations) produce moister clouds with greater fractional coverage that are more strongly influenced by both the soot cloud-burning and the conventional indirect aerosol effects.
An increase in cloud coverage results in more reflection of solar energy (a cooling effect) while at the same time allowing less infrared energy to escape to space (a warming effect). The solar forcing dominates any infrared compensation in trade cumulus which therefore exert a net cooling influence (compare clear-sky to cloudy net fluxes in Fig. 5C).
J. T. Houghton et al. Climate Change 1995: The Science of Climate Change (Cambridge Univ. Press Cambridge 1996) pp. 112–117.
S. K. Satheesh et al. [ J. Geophys. Res. 104 27421 (1999)] estimate that 60% of the aerosol optical depth in the INDOEX 1998 haze was anthropogenic (accounting for an optical depth of 0.12 at 0.5 μm) which is 70% of the optical depth in our baseline haze.
The scope of our calculations (17 simulations of 30 hours) demand a number of computational efficiencies which include parameterized cloud microphysics moderate grid resolution and domain area and no treatment of horizontal radiative transfer.
Distributed over a 2.5-km-deep boundary layer (not in a trade cumulus regime) the diurnal-average aerosol forcings reported by P. B. Russel et al. [ J. Geophys. Res. 104 2289 (1999)] and P. Hignett J. P. Taylor P. N. Francis M. D. Glew [ J. Geophys. Res. 104 2279 (1999)] yield clear-sky heating rates of 0.5 and 0.6 K/day respectively. In comparison the diurnal-average aerosol-induced heating rates in the boundary layer range from 0.5 to 1 K/day for our idealized INDOEX hazes.
R. R. Draxler and G. D. Hess NOAA Tech. Memo. ERL ARL-224 (1997).
V. Manghani et al. Boundary-Layer Meteorol. in press.
Supported by NASA grants NAG5-6504 and NAG5-8362 the U.S. Department of Energy and NSF. The Micro Pulse Lidar measurements were supported by NASA contract NA55-31363 and the NASA SIMBIOS project. We thank J. Coakley for performing radiative transfer calculations for comparison against ours C. Twohy for helpful discussions regarding droplet nuclei measurements from INDOEX the NOAA Pacific Marine Environmental Laboratory (PMEL) and the crews of the R/V Ronald H. Brown and the National Center for Atmospheric Research (NCAR) C-130 aircraft for making the INDOEX measurements possible and two anonymous reviewers for comments that improved the manuscript.