Knowledge of the vertical distribution of aerosols in the atmosphere is important in estimating its radiative forcing. While aircraft based measurements over two locations in India have provided valuable information, the temporal coverage of measurements was limited. In this paper, we examine the vertical distribution of aerosols over a continental, urban location, Bangalore in southern India, using a micro pulse lidar (MPL) operated for about two years (2004 and 2005), and infer the effects of the boundary layer dynamics. Early morning hours are characterized by a shallow aerosol layer, a few hundred meters thick. As day advances, the strong convective eddies are seen to transport the aerosols vertically up to more than 1500 m. Seasonal changes in the aerosol vertical structure, contribution of the boundary layer aerosols to the column optical depth as well as the frequency of occurrence of clouds within aerosol layer are examined.
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 The aerosol radiative forcing still remains one of the largest sources of uncertainty in predicting future climate [e.g., Hansen et al., 1998; Kaufman et al., 2002]. Aerosols consisting of sulfates, nitrates, sea-salt, and organic carbon scatter sunlight back to space and cause a regional cooling effect, while soot, and to a certain extent, mineral dust absorb sunlight and contribute to global warming. The question of whether aerosols warm or cool the planet thus depends on its chemical composition and relative contribution of various chemical species, which constitute the aerosol [Intergovernmental Panel on Climate Change, 2001]. The radiative forcing due to aerosols change from region to region, depending on the relative strengths of various sources and sinks. However, when the amount of absorbing aerosols, such as soot, is significant, aerosol optical depth and chemical composition are not the only determinants of aerosol radiative effects, but the altitude of the aerosol layer, as well as the altitude and type of clouds also become important [Heintzenberg et al., 1997; Satheesh, 2002]. Thus, the vertical distribution of aerosols assumes significance. It is the usual practice to collect aerosol samples on filters and chemically analyze to obtain the mass concentration of different aerosol species. These are then converted to number distribution and subsequently to optical depths using Mie scattering theory [Satheesh et al., 1999; Satheesh and Ramanathan, 2000]. Here, surface-measured properties are to be converted to column properties by making assumptions about vertical profiles. In many cases, the surface aerosol properties are entirely different from column aerosol properties because of the presence of distinct aerosol layers aloft [Ramanathan et al., 2001; Müller et al., 2001]. Thus, the assumption of the same vertical profile of aerosols for different days can result in large errors (as much as by a factor of two) [Satheesh, 2002]. Such vertical layers of aerosols would occur as a consequence of strong thermal convection, which lifts aerosols near ground to greater heights, where they might get entrained when the boundary layer collapses either during the night or due to reduction in convective activities [Stull, 1999]. It could also occur due to long-range transport of aerosols from other places.
 The information on the vertical distribution of aerosols is sparse over Indian region. There exist a few recent studies on the vertical distribution of aerosols using aircraft based measurements [Moorthy et al., 2004; Tripathi et al., 2005]. While Moorthy et al.  carried out measurements over south central India, Tripathi et al.  measured vertical distribution over Indo-Gangetic region. While, these experiments have provided valuable information about the vertical structure, the temporal coverage of measurements was limited. In this paper, we examine the vertical distribution of aerosols over a continental, urban location, Bangalore (13°N, 77°E; 960 m above mean sea-level), using a micro pulse lidar (MPL), and infer the effects of the boundary layer dynamics on the vertical distribution of aerosols.
2. Micro Pulse Lidar (MPL) and Data
 The micro pulse lidar (model MPL1000 of Science and Engineering Services Inc., USA) uses an AlGaAs diode pumped Nd-YLF laser, converting the primary radiation at 1047 nm to its second harmonic at 523.5 nm, at a pulse energy of 10 μJ and a pulse repetition frequency (PRF) of 2.5 kHz. A Schmidt-Cassegrain telescope (20 cm diameter) is used for transmission and reception, in a coaxial, collocated configuration making the MPL a very compact and sturdy system for field operations. The receiver uses silicon avalanche photodiode (Si-APD) in the photon counting mode and this provides higher quantum efficiency (40–50%) than the photo multiplier tube (PMT) based systems. The laser beam is conditioned and directed to the telescope. Since the laser beam is polarized, a polarizing beam splitter is used in combination with a depolarizer to maximize the return signal, which then falls on the detector, the electrical pulses from which are counted after time gating.
 Although MPL used a coaxial geometry and the transmitting laser beam overlaps with receiver field of view almost perfectly, there is certain range within which the scattered radiation is not completely covered by the detector field-of-view. For near surface (when operating from ground) features, this imperfect overlap needs to be corrected. Even though the manufacturers provide a geometrical formula for the overlap correction, in this study we have streamed the laser horizontally during relatively clean atmospheric conditions. Assuming horizontal homogeneity, the range corrected back-scatter (signal multiplied by R2 where R is the range) should be uniform if overlap is perfect. The changes in the backscatter signal from this are attributed to incomplete overlap. While deriving overlap function experimentally, the horizontal homogeneity was ensured by making measurements of column optical depth using Microtops sun photometer (Solar Light Inc., USA) in a mobile platform along the path length. The overlap function derived experimentally as above along with that estimated using the geometrical formula is shown in Figure 1. The theoretical overlap function is estimated using a geometrical formula, which contain information regarding optics used in the lidar. However, overlap function changes from one system to another. The theoretical overlap function need not agree with experimentally derived function as the theoretical overlap function is not realistic as far as each lidar system is concerned. Figure 1 clearly shows that the experimentally determined function approaches the limiting value “1” almost asymptotically at an altitude of ∼4 km (unlike the geometrical function which gives a discontinuity). In this study we have used experimentally derived overlap function up to an altitude of 4 km. Overlap corrected signal is obtained by dividing the signal with overlap function.
 For the analysis of MPL data, we have followed particulate-free zone approach described by Kovalev and Eichinger . This is a modified approach originally proposed by Fernald . This approach is based on the following principal elements.
 (a) The molecular extinction profile is known (for e.g., from standard atmosphere models).
 (b) A priori information is used to specify the boundary value of the particulate extinction coefficient at a specific range within the measured region. In this study we have taken the aerosol extinction to be insignificant (particulate-free zone) at an altitude of 10 km (where normalized back-scatter value is four orders of magnitude smaller compared to that at the surface) and is used as the boundary value.
 The height integrated extinction coefficient (from lidar) was compared with column optical depth measured simultaneously using a Microtops sun photometer (Solar Light Inc., USA) calibrated regularly at Mauna Loa observatory, Hawaii. The comparison between column optical depth estimated from lidar extinction profiles and sun photometer measured optical depths yielded agreement within instrumental uncertainties (∼5%).
 The MPL was operated at Bangalore, for two to three days a week, from May 2004 to December 2005. During winter (November to February), summer (March to May) and post monsoon (September–October) seasons, the schedule was followed without interruptions. During monsoon season (June, July and August) the observations were not possible due to rain. In addition to this, there was a break in observations for five months from October 2004 to February 2005 due to a technical snag. Thus, the data set used in this study comprises of 129 days. The season-wise break up the database is given in Table 1. It may be noted from Table 1 that during monsoon database is weak whereas during summer, winter and post monsoon database is reasonably good.
Table 1. Summary of MPL Data
Number of Days
Post Monsoon 2004
Post monsoon 2005
 A typical summer profile, obtained on 18th March 2005, is shown in Figure 2a, and is representative of the late night/early morning scenario. Time shown in x-axis is Indian Standard Time (IST). Aerosols are mostly confined to a shallow layer of few hundred meters thickness (∼300 m in the night). For daytime conditions, the profile in Figure 2b (for the same day) shows the presence of strong turbulent mixing, lifting the aerosols upwards from the near-surface levels. Consequently, the aerosol layer deepens considerably to reach as high as about 1500 m (from ∼300 m in the night). Plume-like structures of enhanced aerosol concentration are clearly visible above this, going as high as ∼2 to 3 km and having horizontal scales of ∼6 min (which translates to a horizontal eddy size of ∼1.5 to 1.8 km for the prevailing wind speeds of ∼5 m s−1). Winter season is mostly cloud-free for Bangalore, and the peak daytime temperature goes to ∼32°C, so that the solar heating is significant (though not as much as in summer, when the surface temperature can go up to 38°C). Figure 3a shows daytime aerosol profile for 08th November 2005 and Figure 3b that during late night of same day. During the winter, aerosol layer was found to extend beyond 1 km during daytime (Figure 3a). During nighttime the rapid cooling of the earth's surface after sunset, leads to the formation of the nocturnal stable layer, close to the surface (at ∼100–500 m for typical inland conditions) and consequently there occurs elevated layers of enhanced concentration in the residual layer. Such an enhanced layer can be seen in Figure 3b. The occurrence of such aerosol layers aloft was observed almost regularly in winter during late evening (2100 local time) following clear and cloud free daytime. A typical early morning profile is shown in Figure 4 (10th March 2005 at 5 am). Afternoon profiles were generally steep at lower levels and mean profiles for July 2004 and March 2005 are shown in Figure 5. The general features can be summarized as follows.
 (a) During early morning hours, aerosol layer was shallow and well mixed and a few hundred meters thick (for example, Figure 4).
 (b) During daytime, convective mixing of aerosols upwards from lower levels was observed.
 (c) During afternoon hours, a sharp decrease in aerosol extinction up to 2 km followed by less steep decrease to low values at around 6 to 8 km (for example, Figure 5).
 (d) Aerosol layers aloft (between 1.0 and 1.5 km) were observed during winter season during late evening. This feature was observed on almost all days during winter nights preceding sunny days. During summer and post monsoon seasons, no such layer was found in late evening.
 The question of whether aerosol cools (negative forcing) or warms (positive forcing) the planet depends on the relative dominance of absorbing aerosols (such as soot). Investigations over the tropical Indian Ocean have shown that even though soot contribute ∼11% to optical depth, it has an important role in the overall radiative forcing and contributes as much as 35% [Satheesh and Ramanathan, 2000, Babu et al., 2004]. Recently, Satheesh  and S. S. Babu et al. (Temporal heterogeneity in aerosol characteristics and the resulting radiative impacts at a tropical coastal station: 2. Direct short wave radiative forcing, submitted to Journal of Geophysical Research, 2006, hereinafter referred to as Babu et al., submitted manuscript, 2006) have demonstrated that aerosol forcing changes sign from negative (cooling) to positive (warming) when reflection from below (either due to land or clouds) is high. In this context, we examine the contribution of ‘near surface’ aerosols (we define ‘near surface’ as altitude below 1 km from ground) to column optical depth and percentage occurrence of clouds within the aerosol layer. The reason for selecting the 1 km level arises also from the fact that thin stratus or stratocumulus clouds often form above 1 km (at this location) and hence contribution of extinction above and below 1 km has implications while estimating radiative forcing. The results are given in Table 2. The numbers inside bracket in Table 2 indicate the contribution (in terms of extinction) of lowest 2 km to column optical depth. It can be seen that during winter ‘near surface’ aerosols contribute to as much as 60 to 80% to the columnar optical depth (depending on the time of the day), its share decreases to ∼30 to 40% during summer season, when the convective activity is stronger and the minimum surface temperatures rise (from ∼10°C in winter to ∼22°C in summer). This means that aerosols in the higher altitudes contribute more to the columnar optical depth during summer season. Both during morning and afternoon hours (during summer) contribution of aerosols from higher levels (>1 km) dominate that of ‘near surface’ aerosols. During post monsoon months, the contribution to column optical depth from ‘near surface’ aerosols and that above are nearly equal.
Table 2. Near Surface Aerosols and Clouds Within Aerosol Layer: Seasonal Variation
Contribution of Near Surface Aerosols to Column Optical Depth, %
Percentage of Occurrence of Clouds Within Aerosol Layer, %
Morning (7–10 hrs, IST)
Afternoon (14–17 hrs, IST)
Morning (7–10 hrs, IST)
Afternoon (14–17 hrs, IST)
 While estimating aerosol radiative forcing for realistic sky, it is important to know the altitude of the cloud. This is because radiative forcing estimates can drastically change depending on whether cloud is within or above the aerosol layer [Satheesh, 2002]. Recently, Babu et al. (submitted manuscript, 2006) have demonstrated that when cloud is embedded in the aerosol layer, aerosol forcing changes sign from negative (cooling) to positive (warming). Statistics of percentage of occurrence of clouds within aerosol layer as inferred from MPL observations is shown in Table 2. Here, cloud was detected by applying a threshold extinction of 1.0 km−1, which is much larger than aerosol extinction. This criterion obviously excludes invisible clouds. It was observed that thin stratus or stratocumulus clouds often form between 1.0 to 2.0 km. During afternoon hours of summer, occurrence of clouds within the aerosol layer was quite high. The large frequency of occurrence of clouds within the aerosol layer during summer is due to intense convective mixing of aerosols upwards to higher levels (2 to 3 km). Note that we have considered thin and semi-transparent clouds (with half to one km thickness) only. Thicker clouds (such as cumulonimbus) were not transparent to the laser. In summer, when surface is dry and hence bright, it appears that clouds are embedded within aerosol layer. Thus, part of aerosol is below cloud layer and rest is above cloud layer leading to a complicated scenario as far as radiative transfer is concerned. These issues need to be addressed further for accurate determination of regional aerosol radiative forcing.
 (a) During early morning hours, aerosol layer was shallow and well mixed with a few hundred meters thick (∼300 m).
 (b) During daytime, convective mixing of aerosols upwards from lower levels was observed. During afternoon hours, a sharp decrease in aerosol extinction up to 2 km followed by less steep decrease to low values at around 6 to 8 km.
 (c) Aerosol layers aloft (between 1.0 and 1.5 km altitude) were observed during winter season during late evenings.
 (d) A major fraction of column optical depth is contributed by aerosols below 1 km during winter months especially morning hours. During summer convective mixing of aerosols upwards from lower levels makes the contribution from higher levels dominant.
 (e) During afternoon hours of summer, when surface is dry and hence bright, it was observed that clouds are embedded within aerosol layer. This observation has implications while estimating aerosol radiative forcing.
 The authors thank Department of Science and Technology, New Delhi for supporting this work.