Investigation of atmospheric boundary layer characteristics for different aerosol absorptions: Case studies using CAPS model

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Atmospheric Environment 42 (2008) 4755–4768 www.elsevier.com/locate/atmosenv

Investigation of atmospheric boundary layer characteristics for different aerosol absorptions: Case studies using CAPS model G. Pandithurai, C. Seethala, B.S. Murthy, P.C.S. Devara Indian Institute of Tropical Meteorology, Physical Meteorology and Aerology Division, NCL Post, Pune 411 008, India Received 20 June 2007; received in revised form 10 January 2008; accepted 23 January 2008

Abstract This study investigates the direct radiative effects of aerosols on the evolution of atmospheric boundary layer (ABL) over a tropical site, Anand, for 4 days, representative of each season, using land surface processes experiment (LASPEX) data sets used in a one-dimensional ABL and radiative-transfer models. Simulations with the ABL model incorporating fixed aerosol loading with three different levels of absorption and its perturbation with aerosol-free conditions were analyzed. The reduction in net available flux (NAF) increases with increase in aerosol absorption, resulting in maximum reduction for strongly absorbing type. In dry seasons, soil being dry, NAF is partitioned almost equally by latent (LE) and sensible (H) heat fluxes. In wet season, since soil moisture is abundant, LE dominates about 75% in compensating the reduction in NAF. The larger the absorption, the lesser the gradient between the surface and 2-m air temperature and hence more stable the surface layer. The reduction in vertical temperature gradient ranges from 1.74 (dry season) to 0.6 K (wet season). This stabilization of the surface layer reduces sensible heat flux and surface evaporation. Aerosol absorption decreases the turbulent heating but simultaneously increases the solar heating, and in turn increases the air temperature. This affects the inversion layer and hence the ABL height. Reduction in NAF at the surface decreases the ABL height, and the average decrease during the daytime is 167, 204 and 247 m for SSA ¼ 1.0, 0.9 and 0.8, respectively, during dry seasons. It is also found that absorbing aerosols delay the growth and promote early collapse of the ABL in all seasons. r 2008 Elsevier Ltd. All rights reserved. Keywords: Aerosols; Radiation budget; Climate change; Land surface feedback

1. Introduction Aerosols have significant impact on local weather as well as on synoptic scale phenomena like Indian summer monsoon (Patra et al., 2005; Ramanathan et al., 2005; Lau et al., 2006; Niyogi et al., 2006). Aerosols affect the radiation and hence the atmospheric boundary layer (ABL) by three ways, Corresponding author. Tel.: +91 20 2589 3600x263; fax: +91 20 2589 3825. E-mail address: [email protected] (G. Pandithurai).

1352-2310/$ - see front matter r 2008 Elsevier Ltd. All rights reserved. doi:10.1016/j.atmosenv.2008.01.038

directly by scattering and absorption of solar radiation; indirectly by acting as cloud condensation nuclei, leading to an increase in cloud reflectivity and cloud lifetime; and semi-direct/ indirectly by altering the cloud cover through shortwave heating/cooling rate (Twomey, 1974; Hansen and Lacis, 1990; Charlson et al., 1992; Haywood and Boucher, 2000). The indirect and semi-direct effects of aerosols are not discussed in this paper. In general, the scattering and absorption of aerosols reduces the solar radiation reaching the surface, inhibiting sensible and latent heat fluxes

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and inducing feedbacks through various parameters such as air temperature, ABL stratification, etc. The consequent perturbations of surface energy balance and atmospheric radiative heating/cooling can alter the evolution of the ABL (Yu et al., 2002). The available energy at the ground, NAF [(FnetG), where Fnet is the net radiation and G the soil heat flux], mostly controls the evolution of ABL through sensible (H) and latent heat (LE) fluxes during fair weather conditions. Several investigators studied ABL characteristics using different models and observations (e.g., Yamada, 1976; Chaughey and Wyngaard, 1979; Noilhan et al., 1991; Potty et al., 1996; Rao, 2001), but little attention was paid to the effect of aerosols. Satyanarayana et al. (2000) studied the ABL over Anand, India during winter by employing a multilevel ABL model with a soil heat and moisturetransport scheme. Yu et al. (2002) reported a sensitivity study on the effect of aerosols on the evolution of ABL over an extra-tropical site (401N) using the ABL model and a Fu–Liou (Fu and Liou, 1993; Fu et al., 1997) radiative transfer (RT) model. Recently, Wang and Christopher (2006) studied the direct radiative impacts of Central American biomass burning smoke aerosols on the surface energy budget, air temperature and ABL processes using a meso-scale model. Such studies are scarce over tropical regions influenced by monsoonal systems. The present study investigates the effect of aerosol absorption on surface energy fluxes and associated boundary-layer evolution for a tropical semi-arid site with observed land surface and atmospheric data sets representative of winter, summer, monsoon and post-monsoon seasons. 2. Description of models 2.1. Coupled atmospheric plant soil (CAPS) model Numerical modeling of land surface–atmosphere interactions is an effective way to understand the physical mechanisms of exchange processes in the ABL. Here, we used the Oregon State University 1D Coupled Atmospheric Plant Soil (OSU 1-D CAPS) model (Chang et al., 1999). This was developed to simulate the interactions of ABL, vegetation and soil. In this, a planetary boundary layer (BL) model is coupled with an active two-layer soil model and a primitive plant canopy model. This is a high-resolution model with 70 layers, and its first 46 layers are within the lowest 2 km with fine

resolution. This model has also been used as a stand-alone model for a number of numerical experiments under various geophysical conditions (Ek and Mahrt, 1991a, 1994; Holtslag and Ek, 1996; Chang et al., 1999; Yu et al., 2002). The equations used in this model are comprehensive enough to approximate the physical processes thought to be most important, also under a variety of diverse atmospheric conditions for many different locations around the globe. The model treats turbulent mixing with a non-local K approach (Troen and Mahrt, 1986; Holtslag and Moeng, 1991) and its surface layer with Monin–Obukhov similarity. The depth of ABL is diagnosed in terms of the modified bulk Richardson number with an inclusion of temperature excess of thermals (Troen and Mahrt, 1986; Holtslag and Boville, 1993). The total evapotranspiration is estimated as the sum of (i) direct evaporation from the soil, (ii) plant transpiration and (iii) evaporation of water from the plant canopy. A surface energy balance module calculates the surface temperature and potential evaporation (Ek and Mahrt, 1991b). Only vertical diffusion terms due to BL turbulent mixing and advection terms due to a prescribed vertical motion field are considered in the BL equations for the turbulent mixing of the potential temperature, specific humidity and the horizontal component of wind. The moisture and temperature within the soil layer are calculated by the diffusion equations for water and heat transport, respectively (Mahrt and Pan, 1984). The land surface scheme consists of multiple soil layers and a simple plant canopy modified to include the effect of vegetation using a ‘big leaf’ Jarvis–Stewart approach for canopy conductance (Pan and Mahrt, 1987; Noilhan and Planton, 1989). A detailed description of the model can be found elsewhere (Ek and Mahrt, 1991b; Chang et al., 1999; Ek, 2005). 2.2. Santa Barbara discrete-ordinate atmospheric radiative transfer (SBDART) The radiation scheme in the CAPS model is a simple one, in which aerosol radiative effects cannot be included. Hence, a detailed radiative-transfer (RT) code SBDART (Ricchiazzi et al., 1998) has been used. SBDART is a 1-D RT model that computes plane-parallel RT in clear and cloudy conditions within the earth’s atmosphere and at the surface. This code is designed for the analysis of a wide variety of RT problems encountered in satellite

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The land surface processes experiment LASPEX was conducted during 1997–1998 in the Sabarmati river basin of Gujarat, located in the semi-arid part of the western Indian region, by the Indian Institute of Tropical Meteorology (IITM), Pune, India and Gujarat Agricultural University (GAU), Anand, India. The main objective of this experiment is to conduct a micrometeorological experiment and generate surface and subsurface atmospheric and hydrological data sets to parameterize land surface processes in numerical models over a grid scale of 100 km  100 km. During this experiment, micrometeorological observations were collected at five sites, with Anand (22.351N, 72.551E) as the central site. The LASPEX experimental setup consists of measurements of meteorological parameters, viz., temperature, humidity, wind speed and wind direction, at heights of 1, 2, 4 and 8 m above the surface using a micrometeorological tower. Slow rising radiosonde ascents were taken using low ascent rate balloons synoptic 3-hourly from 0000 GMT to 1200 GMT. Soil temperature was measured at the surface, and at depths of 0.05, 0.10, 0.20, 0.40 and 1.0 m. Radiation components, viz., incoming shortwave (solar) radiation, reflected shortwave radiation, incoming longwave radiation and outgoing longwave radiation, are measured at a height of approximately 2 m on a separate platform near the tower. Turbulence measurements were made by using sonic anemometers at 4 and 9 m levels. Soil moisture at different depths was measured by gravimetric method, neutron probe and capacitance

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soil moisture probe. Experimental setup and details of the observations have been reported elsewhere (Vernekar et al., 2003). The data have been utilized by various scientific groups in India (Pillai et al., 1998; Satyanarayana et al., 2000; Sadani and Kulkarni, 2001; Murthy and Parasnis, 2002). Simulations were done for Anand, central site, as this site only has radiosonde ascents for temperature and humidity profiles to initialize the model. The soil type over the experimental site is loamy sand. The observed soil moisture on 14 May is quite similar to those available from the NCEP/NCAR reanalysis (Kalnay et al., 1996). On other days, due to paucity of data, we make use of the NCEP/ NCAR reanalysis data for initial soil moisture used in the model. On 16 July, NCEP soil moisture (0.36 m3 m3 in 0–10 cm soil layer) represents a supersaturated state. Soil moisture and temperature at two levels on all four study days are shown in Fig. 1. The observed broadband surface albedo is 0.23 in May and about 0.20 in February, July and October. The primary seasonal difference in

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remote sensing and atmospheric radiation budget studies. All radiatively active species such as water vapor, O3, N2, O2, CO2, CH4, N2O, etc. are taken into account. The RT equations are integrated with the discrete-ordinate method. The program is based on a collection of well-tested and reliable physical models. In the RT model, aerosol optical properties can be specified through vertical aerosol optical depth (AOD, aerosol loading), single scattering albedo (o or SSA, fraction of absorption), asymmetry parameter (g, magnitude of forward scattering) and Angstrom exponent (size distribution) to include aerosol radiative effects. Radiative forcing and atmospheric radiative heating rates derived from the RT model are coupled in the CAPS model to study the aerosol radiative effects on ABL parameters.

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vegetation is the fractional coverage, which is approximately 0.8, 0.6, 0.7 and 0.9 for February, May, July and October, respectively. 4. Methodology We selected 4 days from the LASPEX Intensive Observational Period data sets representative of each season covering different surface conditions to study the effect of aerosols on ABL parameters. The days selected were 14 February (winter), 14 May (pre-monsoon), 16 July (monsoon) and 13 October (post-monsoon) in the year 1997. A cloud-free environment was preferred in order to avoid complex land–atmosphere interactions with the presence of clouds, and hence clear sky conditions are considered in the RT calculations. The momentum roughness of 0.1 m was obtained from the modified classification table as it was found to be in agreement with that derived from gustiness analysis of wind (Wieringa, 1977, 1980). It takes into account both nearby and far-off obstacles, vegetation, etc. in the upward direction within a distance of 3 km over a sector of 20–301 width. The roughness length for heat is taken as one order of magnitude smaller than that of momentum. The CAPS model has been validated for the experimental site with observations by tuning the soil and vegetation parameters. The land surface parameters in the model are changed according to the soil and vegetation at the site, Anand. Soil type is ‘loamy sand’ and vegetation is ‘crops’. Soil porosity ¼ 0.41 m3 m3; field capacity ¼ 0.25 m3 m3; wilting point ¼ 0.07 m3 m3; minimum canopy resistance ¼ 13.3 sm1 (Murthy et al., 2004). Soil moisture and temperature observations taken at

Anand provide initial soil conditions for model simulations, while Radiosonde launched from the same location at 0530 h LT (00 UTC) provides initial atmospheric conditions. Vertical velocity data were obtained from the National Centre for Medium Range Weather Forecasting (NCMRWF, India) model analysis, and is specified at each model level. Geostrophic winds are estimated from the actual winds at approximately 1.5 km from the 0530-h LT soundings on the study day and 24 h later, and are linearly interpolated in time and height-independent. In the RT model, aerosol optical properties were specified through AOD, SSA and the asymmetry parameter at mid-visible wavelength (0.5 mm) for a mixture of scattering and absorbing components. Optical depth is scaled into vertical profile with 80% extinction in the BL with a scale height of 1.4 km. Optical depth at other wavelengths is extrapolated using Angstrom exponent. In this study, a constant AOD of 0.5 and asymmetry parameter of 0.7 at mid-visible wavelength 0.5 mm is used. An Angstrom exponent of 1.4 is used to extrapolate AOD at other wavelengths. In this study, the SSA of aerosols at mid-visible wavelength (0.5 mm) is assumed to be 1.0, 0.9 and 0.8 to represent cases of pure scattering, moderate absorption and strong absorption, respectively. The urban aerosol model of OPAC (Hess et al., 1998) is used to scale the spectral variation of SSA and asymmetry parameter. The aforementioned three values of SSA were employed to simulate the short- and long-wave radiative fluxes at the surface and radiative heating rates due to aerosol absorption at the ABL model levels. Model simulations were initiated at 0530 h LT and are integrated for 24 h with a time step of

Table 1 Input and output parameters in CAPS and SBDART models Models

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Land surface: Surface albedo, roughness length for heat and momentum, soil type, wilting point, shading factor, air-dry value, plant coefficient, soil moisture and soil temperature

Temperature, potential temperature, relative humidity, sensible, latent, soil heat fluxes, 2 m wind speed, 2 m air temperature, surface temperature, ABL height

Atmosphere: Vertical profiles of u, v, w components of wind, air temperature and mixing ratio SBDART model

Aerosol parameters: Aerosol optical depth, Angstrom exponent, single-scattering albedo and asymmetry parameter Atmosphere: Vertical profiles of temperature, relative humidity, total column ozone, precipitable water content

Total downward and upward flux at the surface and at the top-of-the atmosphere (TOA); direct downward flux at the surface and at TOA; radiative heating/cooling rates computed from fluxes at different height levels

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180s. Input and output parameters of both the models are given in Table 1. 5. Results and discussion 5.1. Radiation budget Aerosol reduces the net solar radiation absorbed by the surface depending on the SSA. The reduction increases with decreasing SSA, because aerosol absorption removes part of solar radiation that would otherwise reach the surface. Thus, the reduction is more for strongly absorbing aerosols. The daytime average (0800–1800 h) net flux on all the four simulation days decreases to nearly 33, 57 and 78 W m2 for SSA 1.0, 0.9 and 0.8, respectively. The diffuse to direct radiation ratio (DDR) increases for increase in aerosol loading. However, for a given aerosol loading, DDR is high for purely scattering aerosols (SSA ¼ 1.0) and decreases for increase in absorbing aerosols. From the RT calculations, it is found that daytime mean DDR decreases by 15% and 28% for SSA of 0.9 and 0.8, respectively, as compared to SSA ¼ 1.0, for a given aerosol loading of AOD ¼ 0.5. For semi-arid landscapes in the tropics, the diffuse radiation effect of aerosols may not be as large as those for tropical rainforests because of a reduced leaf area index (LAI), but it can still be significant (Niyogi et al., 2006). Diurnal average short-wave aerosol radiative forcing efficiency at the surface (radiative forcing per unit optical depth) over the Indian subcontinent is reported to vary from 70 to 88 W m2 (Babu et al., 2002; Ganguly et al., 2005; Jayaraman et al., 2006; Pandithurai et al., 2004; Singh et al., 2005; Tripathi et al., 2005). The short-wave aerosol forcing efficiencies at the surface computed from SBDART are 36, 63, 87 W m2 for SSA ¼ 1.0, 0.9 and 0.8, respectively. This suggests that aerosol SSA over the Indian subcontinent can lie in between moderately and strongly absorbing type. Simulated net radiation for clean and aerosols of different SSAs along with the observed values on 14 May is shown in panel (a) of Fig. 2. The reduction of solar flux at the surface changes due to factors such as aerosol type, aerosol amount, absorption fraction, backscatter fraction, Rayleigh scattering and surface albedo. The reduction in NAF during daytime (0800–1800 h) is about 27, 53 and 76 W m2 in dry season (February and May) and 30, 61 and 89 W m2 in wet season (July

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and October) for SSA of 1.0, 0.9 and 0.8, respectively. A fraction of the radiative fluxes, reduced in reaching the surface, gets absorbed in the atmosphere and in turn heats the atmosphere. The rate of change of temperature (dT/dt) in a layer due to radiative heating/cooling is called radiative heating/cooling rate (HR), defined as dT 1 dF ¼ dt rC p dZ where Cp is the specific heat capacity and r is the density of air. dF/dZ is the radiative flux divergence. The heating/cooling rates were computed using the above equation. Differences in heating rates with and without aerosols yield HR due to aerosols. The diurnal mean aerosol heating rates upto an altitude of 3 km estimated on all 4 days were about 0.01, 0.73 and 1.34 K day1 for SSA ¼ 1.0, 0.9 and 0.8, respectively. 5.2. Surface fluxes and its perturbations with aerosol absorption for different seasons The perturbation in the incoming solar radiation changes the net energy available at the surface, which determines the surface heat fluxes. As aerosols reduce the solar radiation reaching the surface, there is reduction in H and LE. The reduction increases with decrease of SSA and the partitioning of turbulent fluxes is dependent on the temperature and moisture present in the atmosphere and at the surface and aerosol radiative heating in the atmosphere. The daytime variations in simulated sensible and latent heat fluxes for clean (AOD ¼ 0.0) and different aerosol SSAs (AOD ¼ 0.5, SSA ¼ 1.0, 0.9 and 0.8) for 14 February (winter) are shown in Fig. 2(b) and (c). The observed Fnet and sensible heat fluxes are also shown in the above figure. Daytime mean perturbations in H and LE for all days as a function of aerosol SSA are shown in Fig. 3. For example, perturbation in H, DHQH (aerosol, AOD ¼ 0.5, SSA ¼ 1.0, 0.9 and 0.8)H (no aerosol, AOD ¼ 0.0), and this definition is applicable to all parameters discussed in this paper. In winter, the average reduction in H due to aerosols is about 11, 22 and 32 W m2 for aerosol SSA of 1.0, 0.9 and 0.8, respectively. But the reduction in LE is higher and is 21, 32 and 43 W m2 for SSA 1.0, 0.9 and 0.8, respectively. To quantify the indirect effect of solar heating of the atmosphere due to aerosol absorption, percentage reductions in H and LE for a given

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reduction in NAF (DH/DNAF and DLE/DNAF) at the surface are calculated for SSA 1.0, 0.9 and 0.8. For a given reduction in NAF, the percentage reduction in H increases to 35%, 39% and 42%, whereas the percentage reduction in LE decreases to 65%, 57% and 56% for SSA 1.0, 0.9 and 0.8, respectively. It suggests that absorbing aerosols (SSA ¼ 0.8) cause less reduction in surface evapora-

tion relative to that by scattering aerosols. In other words, absorbing aerosols have a moderate effect on the reduction of moisture content of ABL comparatively. For purely scattering aerosol, the reduction in the NAF largely goes to reducing the evaporation because of negligible changes in surface–air temperature difference. The drier air with absorbing aerosols increases the atmospheric demand for

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water vapor, compensating the reduced evaporation due to the cooling effect of decreased radiative flux and resulting in a small change in the latent heat flux. This may be due to increase in stability of the surface layer (due to increase in air temperature for SSA ¼ 0.8), which results in relatively large decrease in H and less reduction in LE. This effect is reflected in Bowen ratio, averaged for the period (0800–1800 h LT), which is found to be 0.25, 0.20, 0.16 and 0.12 for clean and SSA 1.0, 0.9 and 0.8, respectively (Fig. 2(d)). In May (summer), the soil temperature is 301 and 295 K in the surface layer 0–10 and 10–100 cm, respectively. The soil moisture is 0.1 m3 m3 in the 0–10 cm layer and 0.08 m3 m3 in the 10–100 cm layer. The reduction in NAF is 22, 50 and 75 W m2 for SSA 1.0, 0.9 and 0.8, respectively. Consequently, reduction in H is about 10, 27 and 41 W m2 and the reduction in LE is about 12, 24 and 36 W m2 for SSA 1.0, 0.9 and 0.8, respectively. Changes in land surface characteristics can alter the surface Bowen ratio and have a feedback via BL processes. If soil moisture is abundant, evaporation takes place at potential rate (a function of air temperature and moisture) that results in more moisture in atmosphere and drying of soil (top layer just below the ground). As soil dries up gradually, evaporation

depends on moisture flux from deep soil layer to the surface. Humid atmosphere with less temperature increases the probability of cloud formation and consequent precipitation, which in turn increases soil moisture. The daytime averaged Bowen ratio (ratio of sensible to latent heat flux) decreases from 0.70 (no aerosol case) to 0.67, 0.64 and 0.61 for aerosol SSA of 1.0, 0.9 and 0.8, respectively. In July (monsoon), the reduction in NAF is 27, 61 and 91 W m2 for SSA 1.0, 0.9 and 0.8, respectively. In contrast to dry season, reduction in NAF is partitioned into 20% by H and 80% by LE. The reduction in H is 3, 14 and 20 W m2 for SSA of 1.0, 0.9 and 0.8, respectively. The reduction in LE is about 21, 43 and 69 W m2 for SSA of 1.0, 0.9 and 0.8, respectively. Bowen ratio during noontime is found to be 0.08 for clean case and 0.07, 0.05 and 0.03, for SSA of 1.0, 0.9 and 0.8, respectively. In October (post-monsoon), the lower soil layer is wetter and warmer than the top layer. The reduction in NAF with respect to the clean case is 33, 62 and 87 W m2 for SSA 1.0, 0.9 and 0.8, respectively. This reduction is compensated by H (10%) and LE (76%) for SSA of 1.0, whereas the compensation by H is of about 18% and that by LE of 72% for SSA of 0.9 and 0.8, respectively.

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For a given NAF, H dominates for dry soil whereas LE dominates for wet soil. Similar changes apply for a given reduction in NAF due to aerosols. Thus, during early morning, soil moisture being relatively more compared to noon, DLE is more than DH for a given DNAF. By noon, soil moisture below the ground surface (dries up) reduces, which results in DH4DLE for a given DNAF. The reduction of NAF increases with the decrease of SSA, resulting in maximum reduction for the strongly absorbing type of aerosols on all the study days. In May, soil being dry, DNAF is partitioned mostly into DH as compared to DLE. Early morning the soil is relatively wet, and hence evaporation dominates over sensible heat flux. In wet season, since soil moisture is abundant, DLE dominates relative to DH in compensating DNAF throughout the day from 0600 to 1800 LT. As the soil type is loamy sand (upper soil layer is wetter than the lower layer), most of the excess energy is released from the surface through the surface evaporation (LE) than the surface heating (H) on 16 July. 5.3. Perturbation in surface as well as 2 m air temperature Fig. 4 shows how the aerosol and its effect on surface energy balance modify the surface temperature (Tsfc) and 2 m air temperature (T2m) on all 4 days representative of each season. From the above figure, it is evident that the surface temperature decreases from the clean atmosphere to aerosolladen atmosphere. The mean decrease in Tsfc during the period 0800–1800 h LT is about 2.5 K in February and May (dry season), and it is about 0.4–0.7 K in July and October (wet season) for strongly absorbing aerosols with respect to clean conditions. The 2 m air temperature varies between the skin temperature and ABL temperature. It increases for moderately and strongly absorbing type of aerosols in the afternoon hours because of strong aerosol radiative heating. There is a slight decrease in 2 m air temperature for purely scattering aerosols due to radiative cooling. The impact of aerosols on the stratification of surface layer can be readily derived by taking the difference between Tsfc and T2m. Fig. 5(a) shows an average change in surface minus 2 m air temperature with clean conditions for the period from 0800 to 1800 LT for different SSA. The larger the absorption (lower SSA), the larger the reduction in surface minus 2 m

air temperature difference and more stable the surface layer. This stabilization of the surface layer reduces the sensible heat flux and the surface evaporation (in the demand control stage). In dry season, the average reduction in TsfcT2m is about 0.42, 1.1 and 1.5 K for SSA of 1.0, 0.9 and 0.8, respectively. But in wet season, the average reduction in TsfcT2m is about 0.15, 0.46 and 0.69 K for SSA of 1.0, 0.9 and 0.8, respectively, which is lower as compared to dry season. It can be noted that the reduction in TsfcT2m is more for absorbing aerosols during dry season, which stabilizes the surface layer and reduces the ABL height. 5.4. Perturbation in ABL height Diurnal variation of the ABL height is determined by the surface buoyancy flux and the strength of capping inversion, both of which are affected by aerosols. Perturbation in daytime mean ABL height, buoyancy flux and entrainment heating for all days as a function of aerosol SSA is given in Fig. 5(b–d). Daytime diurnal variations in ABL heights simulated for clean and different aerosol SSAs for all study days are shown in Fig. 6. Buoyancy flux decreases due to aerosols and the reduction is more for strongly absorbing aerosols. Thus in February and October, as buoyancy flux decreases, the height of the BL also decreases from clean to different SSAs of 1.0, 0.9 and 0.8. The average decrease in ABL height is 167, 204 and 247 m for aerosol SSA of 1.0, 0.9 and 0.8, respectively, in dry seasons (February, May and October). In July, the average reduction in daytime ABL height is about 30, 13 and 5 m for SSA of 1.0, 0.9 and 0.8, respectively, in contrast to other days in dry season, where ABL height decreases more for moderately and strongly absorbing aerosols. However, the ABL temperature increment during the day depends on balance between the surface sensible heat flux, atmospheric absorption of solar radiation (heating rate) and the entrainment heat flux. Dry air entrainment (Cq) approximated as C q ¼ w0 q0 h =w0 q0 s , where w0 q0 is the heat flux; subscripts h and s refer to the levels just below the BL top and the surface, respectively. Cq40 means entrainment of dry air from above the BL and Cqo0, entrainment of moist air from above the BL. Entrainment drying into the ABL from above the ABL decreases from clean to different aerosol absorption in February and October. In case of purely scattering type of aerosols, reduction in both the sensible heating and the

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Fig. 4. Temporal variation in simulated surface temperature Tsfc and 2 m air temperature T2m for clean and different aerosol absorption conditions on all 4 days representative of each season. (a,b) 14 February, (c,d) 14 May, (e,f) 16 July and (g,h) 13 October.

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Fig. 5. Perturbations in (a) TsfcT2m, (b) ABL height, (c) buoyancy flux, (d) entrainment heating (Cq) as a function of aerosol SSA.

entrainment heating makes comparable contributions to lowering the ABL temperature. When there is significant absorption due to aerosols, the sensible heat flux decreases, while the absorbed solar radiation increases, causing net increase in ABL temperature. In addition, the increased entrainment heating with elevation of ABL causes the ABL temperature to rise. For purely scattering aerosols,

radiative cooling due to aerosols as well as reduced buoyancy flux causes ABL to shrink than in the clean case, while for strongly absorbing type of aerosols strong solar heating overcomes the effect of reduction in buoyancy flux and allows the ABL to grow relative to that of the clean atmosphere in May and July. This agrees with previous report over a dry subsiding region in extra-tropics (Yu et al.,

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2500

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Fig. 6. Simulated diurnal variation of the ABL height for clean and different aerosol absorption cases on all four study days. (a) 14 February (winter), (b) 14 May (pre-monsoon), (c) 16 July (monsoon) and (d) 13 October (post-monsoon).

2002). In May, buoyancy flux decreases in the afternoon hours and more so for strongly absorbing aerosols. The entrainment of dry air from above the BL and associated decrease in relative humidity enhance the development of BL height in the

afternoon hours. Temporal evolution of ABL height shows nearly 1 h delay in growth and is 1 h early in collapse. ABL height has not increased significantly in the afternoon hours for absorbing aerosols (except in 14 May) as reported by Yu et al. (2002)

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for extra-tropics. For moderately absorbing type of aerosols, reduction of buoyancy flux overcomes the effect of direct heating and reduces the ABL height as compared to the clean case. In July, the reduction of buoyancy flux and the effect of direct heating compensate each other and hence no change in the ABL height. It can be noted that the later growth and earlier collapse of ABL for absorbing type of aerosols is due to the reduction of buoyancy flux. Thus, the effect of aerosol absorption on the evolution of ABL height can differ from dry to wet season depending on the atmospheric as well as soil conditions. 6. Summary and conclusions The utilization of the high-resolution ABL model in this study simulates effectively the responses of important ABL processes to the aerosol radiative perturbation including surface heat release, surface evaporation, turbulent mixing, and entrainment over a tropical semi-arid site. In the vegetated surface (July and October), more soil moisture can promote most of the excess energy from the surface through evaporation than sensible heating, but hot summer can reverse this situation. Thus, the partitioning of energy between sensible heat flux and evaporation depends not only on the aerosol absorption of solar radiation but also on the land surface characteristics. As aerosol absorption removes part of the radiation reaching the surface, surface temperature decreases and the decrease is high for strongly absorbing aerosols (SSA ¼ 0.8), irrespective of season. But the change in air temperature depends on atmospheric conditions. On all the days, air temperature increases for absorbing aerosols during late morning to early afternoon. Thus, the difference TsfcT2m decreases with increase in absorption and the decrease is more for strongly absorbing aerosols. This reduces the sensible heat flux, which in turn reduces the heating of the overlying atmosphere. However, aerosols directly heat the atmosphere by absorption. Thus, purely scattering (absorbing) aerosols decrease (increase) the temperature of the ABL. In addition, entrainment heating also contributes during daytime, either positively for strongly absorbing type of aerosols or negatively for purely scattering type of aerosols. During May and July, the reduced ABL temperature along with the reduced buoyancy flux increases the strength of capping inversion for

purely scattering type of aerosols. On the other hand, the increased ABL temperature due to the dry air entrainment in the BL for SSA of 0.8 raises the top of the ABL in spite of the reduction in the buoyancy flux. The combination of the reduced surface evaporation and the change in entrainment drying determines the sign and magnitude of the perturbation of water vapor in the ABL. Changes in radiative redistribution of surface energy fluxes (via Bowen ratio changes), and a potential for warming/ cooling as a feedback to the aerosol loading suggest that there is a need to consider the role of aerosols and their absorption levels in land–atmosphere forcing. This serves as a sensitivity study for different aerosol absorption cases during dry and wet seasons with observed land surface processes experiment data sets. It can be revisited with observed collocated/simultaneous aerosol, atmospheric boundary layer and land surface processes data to delineate the effect of aerosols on the ABL evolution. Acknowledgments The authors would like to thank the IITM Director for encouraging this work. This work is supported under ISRO-GBP Aerosol Radiation Budget Studies projects. Thanks are due to Dr. Michael Ek, Oregon State University for providing OSU 1-D CAPS model and discussions. The SBDART code from Dr. Paul Ricchiazzi, UCSB and discussions with Dr. Hongbin Yu, NASA/ GSFC are also acknowledged. Thanks are due to the Department of Science and Technology for LASPEX data sets. Constructive comments/suggestions from the three anonymous reviewers are gratefully acknowledged.

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