**8. Effect of tropospheric convection on UTLS particulates**

During the ASM period, the upper troposphere is significantly influenced by the tropospheric convection. This in turn could influence the microphysical properties of UTLS particulates. The outgoing long wave radiation (OLR), which is directly influenced by cloud cover and in turn by convection, could be effectively used as a proxy to the strength of tropospheric convection [59] and the seasonal variation of deep convection. The value of mean OLR is generally high for clear sky condition (~460 W m-2) and decreases with increase in cloud cover and cloud vertical extent. A threshold value [60] of 200 Wm-2 for OLR which when declines below could be treated as in index for deep convection. The percentage of occurrence of daily mean OLR with its value < 200 Wm-2, within a grid size of 2.5° × 2.5° over a geographical region 12.5° –15° N in latitude and 77.5° –80° E in longitude obtained from NCEP/NCAR reanalysis provided by the Climate Diagnostic Center through their Web site at http://www.cdc.noaa.gov/ averaged in different months for the period 1998–2003, which is almost negligible during the December to April period, starts increasing from May and shows a broad peak during the period June to October (Figure 11a).

The strength of tropospheric convection also can be assessed from the convective available potential energy (CAPE), defined as the altitude integrated buoyancy of lifted air parcel from the level of free convection to the level of neutral buoyancy [61], is estimated from the altitude profile of potential temperature as CAPE = ∫ [g(qv − qT)/qT]dz, where qv is the virtual potential temperature of the air parcel, qT is the environmental potential temperature and g is acceleration due to gravity. The integral is taken from the level of free convection to the level of neutral buoyancy. The mean values of CAPE, which is a measure of stability in the atmosphere as far as vertical displacements are concerned [62], in different months for Chennai (13°N, 80.2°E), a station located very close to Gadanki, obtained through India Meteorology Department (Web site http://www.weather.uwyo.edu/upperair/ sounding.html.) are presented in Figure 11b. The CAPE also shows a prominent peak during the period April to May along with a small secondary peak in September. Intense convection is closely related to thunderstorm activity. Climatologically averaged thunderstorm activity in different latitude belts over the Indian subcontinent was investigated in detail in earlier study [63]. The annual variation of the mean number of thunderstorm days for the latitude belt 10–15°N (Figure 11c) shows two prominent peaks in May and September–October period. Deep convection and high thunderstorm activity leads to the formation of thick convective clouds. The main reason for low lidar observation statistics during the monsoon period is the presence of these convective clouds which impede the observations. Though these clouds will have a large spatial extent and persist for several days, some gap region which is devoid of thick clouds (favouring lidar observation) starts developing after a few days The outflow from adjacent convective anvils spreading over to this gap region leads to the formation of STCs.

126 Atmospheric Aerosols – Regional Characteristics – Chemistry and Physics

**8. Effect of tropospheric convection on UTLS particulates** 

and shows a broad peak during the period June to October (Figure 11a).

During the ASM period, the upper troposphere is significantly influenced by the tropospheric convection. This in turn could influence the microphysical properties of UTLS particulates. The outgoing long wave radiation (OLR), which is directly influenced by cloud cover and in turn by convection, could be effectively used as a proxy to the strength of tropospheric convection [59] and the seasonal variation of deep convection. The value of mean OLR is generally high for clear sky condition (~460 W m-2) and decreases with increase in cloud cover and cloud vertical extent. A threshold value [60] of 200 Wm-2 for OLR which when declines below could be treated as in index for deep convection. The percentage of occurrence of daily mean OLR with its value < 200 Wm-2, within a grid size of 2.5° × 2.5° over a geographical region 12.5° –15° N in latitude and 77.5° –80° E in longitude obtained from NCEP/NCAR reanalysis provided by the Climate Diagnostic Center through their Web site at http://www.cdc.noaa.gov/ averaged in different months for the period 1998–2003, which is almost negligible during the December to April period, starts increasing from May

The strength of tropospheric convection also can be assessed from the convective available potential energy (CAPE), defined as the altitude integrated buoyancy of lifted air parcel from the level of free convection to the level of neutral buoyancy [61], is estimated from the altitude profile of potential temperature as CAPE = ∫ [g(qv − qT)/qT]dz, where qv is the virtual potential temperature of the air parcel, qT is the environmental potential temperature and g is acceleration due to gravity. The integral is taken from the level of free convection to the level of neutral buoyancy. The mean values of CAPE, which is a measure of stability in the atmosphere as far as vertical displacements are concerned [62], in different months for Chennai (13°N, 80.2°E), a station located very close to Gadanki, obtained through India

altitude.

shows a pronounced oscillation with prominent peaks in May and September–January period. Above 20 km, the value of mean δ (~0.03) is very close to the molecular depolarization, it is relatively large (in the range 0.03 – 0.04) in the altitude region 18–20 km just above the cold point tropical tropopause. Thus, in the LS region, the particle size and non-sphericity decreases with increase in altitude. Note that, the value of δ encountered in this region is less than the threshold value (0.04) used for identifying the cirrus particles. The value of δ is largest in the altitude region 14–16 km, just below the mean level of tropical tropopause. The general similarity in the annual variation of δ in the altitude regions 14–16 km and 10-12 km indicates that the particle microphysics in these altitudes are strongly coupled. Even though the cold point tropopause altitude shows a small variation from month-to-month, the mean level lies around 16.5 km. Figure 10d shows the mean annual variation of the altitude weighted δ around the cold point tropopause for a slab thickness of 2 km. The observed general similarity of the annual variation of δ in the entire region from 10 to 20 km indicates that the annual variation of the particle habit (size and shape) in the UT and LS regions are strongly coupled. The observed general decrease in δ from UT to LS suggests that the particles tend to become small and more regular in shape with increase in

**Figure 11.** Annual variation of the (a) frequency distribution of outgoing long-wave radiation (<200 Wm-2) around Gadanki between 12.5°–15°N and 77.5°–80°E along with (b) convective available potential energy (CAPE) obtained from Radiosonde observations at Chennai during the period 1998– 2003. (c) Month-to-month variation of mean number of thunderstorm days (TSD) for the latitude belt 10–15°N during the period 1970–1980 [63].

Being originated from convective outflow [43,44], especially during the summer monsoon period, particles of STCs in the UT region during this period will be relatively large and highly non-spherical [8,45]. Presence of these large non-spherical particles leads to an increase of δ in this region, as is observed. Subsequent uplift of some of these particles along with tropospheric air across the tropopause leads to an increase in δ along with the integrated backscatter (ΙβP) in the region just above the cold point tropopause. As the particles and precursor gases are not directly injected into the stratosphere but diffuse

slowly across the tropopause they confine to a small region just above the cold point tropopause. Note that during the period July to August when CAPE as well as thunderstorm activity shows a small decrease, δ (Figure 12) and ΙβP (Figure 13) in the 18–20 km also shows a decrease. However, the STCs during this period shows large spatiotemporal variations [37] introducing large day-to-day variability in ΙβP and δ in the UT and LS region as indicated by the large error bars during the May-October period.

Distribution of Particulates in the Tropical UTLS over

the Asian Summer Monsoon Region and Its Association with Atmospheric Dynamics 129

circulation in the troposphere and Brewer- Dobson circulation (B−D) in the UTLS region. The B−D circulation is primarily driven by mechanical forcing (westward directed wave drag) arising from the dissipation of planetary-scale waves (breaking of Rossby waves) in the extra tropical stratosphere [64]. Because of filtering by the large-scale stratospheric winds, vertical propagation of planetary waves into the extra-tropical stratosphere occurs primarily during winter and this seasonality in wave forcing accounts for the winter maximum in the B−D circulation [65,66]. In association with these processes, the cold point tropopause at tropics is pushed to a higher altitude during winter and they play a major role

As can be seen from Figure 10a the tropopause is cooler and higher during the January– March period and warmer and lower during August–October period which is in accordance with what is reported for tropical locations [67-70]. These Seasonal changes in the cold point tropopause altitude also contribute to the mass influx to stratosphere [71-73]. Transport of constituents like water vapor, aerosols, trace gases etc across the tropopause which acts as a permeable membrane through which continuous mixing between tropospheric and stratospheric air takes place is well demonstrated by various investigators [74-77]. Over and above the penetration of deep convection, this annual adjustment of tropopause altitude is

in exchange of mass from troposphere to stratosphere.

to be in its background level.

an important mechanism for the stratosphere troposphere exchange.

**9. Contribution of STC to particulate scattering in the UTLS region** 

Fine ice crystals of STCs originating either through the outflow from deep convective anvils or through freeze drying of moist air lifted up to the tropopause by normal convection contribute significantly for scattering in the UTLS region. In addition, other aerosol particles originating from the surface (mainly through bulk to particle conversion and those of vegetative origin) as well as those formed in the upper troposphere through gas- to-particle conversion (mainly of various industrial gases) also contribute for the particulate loading in the UTLS region. It would be worth examining the relative contributions of the two components in the UTLS region to make a quantitative assessment of the contribution of STCs to the scattering properties in the UTLS region. Figure 8a shows the contour plots of month-to-month variation of mean βp derived from lidar data during the period 1998–2003 (including both STC-contaminated and STC-free). Note that, as this study period was mostly devoid of major volcanic eruptions [24] the stratospheric aerosol loading can be considered

While βp in the UT is a maximum during the May –September period and minimum during October–November, in the lower stratosphere it is minimum during summer (July and August) and maximum during winter. However, the winter high in the LS region was attributed [7] to the transport of tropospheric air (containing aerosols and precursor gases) in conjunction with the tropical upwelling and B− D circulation while the observed high during May–June is due to the upward influx of particles (including ice crystals of STCs) from the UT region. As the uncertainty associated with lidar-derived βp is small compared to αp [42], for the lidar based study the altitude profile of βp and ΙβP for a desired altitude

**Figure 12.** Mean annual variation of integrated volume depolarization ratio (VDR) normalized to the slab thickness of 2 km in the altitude region 14-16 km (UT) and 18-20 km (LS) superposed on the annual variation of CAPE and number of thunderstorm days. Vertical bars represent the respective standard error.

**Figure 13.** Mean annual variation of Integrated Particulate Back scatter (IβP) in UT (a) and LS (b) superimposed on the annual variation of CAPE for the period 1998-2003.Vertical bars represent the standard error.

Though during the December–February period CAPE is relatively small and thunderstorm activity also is less, a broad peak in ΙβP at 18–20 km is observable during this period (Figure 13).This feature though appears to be contradicting to the above hypothesis proposed for the summer monsoon period, the real mechanism responsible for the transport of tropospheric air into the stratosphere during winter is different from that during the monsoon period. During winter the intrusion of tropospheric air into the lower stratosphere occurs through vertical ascent of air driven by the strength of the ascending branch of the Hadley cell circulation in the troposphere and Brewer- Dobson circulation (B−D) in the UTLS region. The B−D circulation is primarily driven by mechanical forcing (westward directed wave drag) arising from the dissipation of planetary-scale waves (breaking of Rossby waves) in the extra tropical stratosphere [64]. Because of filtering by the large-scale stratospheric winds, vertical propagation of planetary waves into the extra-tropical stratosphere occurs primarily during winter and this seasonality in wave forcing accounts for the winter maximum in the B−D circulation [65,66]. In association with these processes, the cold point tropopause at tropics is pushed to a higher altitude during winter and they play a major role in exchange of mass from troposphere to stratosphere.

128 Atmospheric Aerosols – Regional Characteristics – Chemistry and Physics

error.

standard error.

slowly across the tropopause they confine to a small region just above the cold point tropopause. Note that during the period July to August when CAPE as well as thunderstorm activity shows a small decrease, δ (Figure 12) and ΙβP (Figure 13) in the 18–20 km also shows a decrease. However, the STCs during this period shows large spatiotemporal variations [37] introducing large day-to-day variability in ΙβP and δ in the UT and

**Figure 12.** Mean annual variation of integrated volume depolarization ratio (VDR) normalized to the slab thickness of 2 km in the altitude region 14-16 km (UT) and 18-20 km (LS) superposed on the annual variation of CAPE and number of thunderstorm days. Vertical bars represent the respective standard

**Figure 13.** Mean annual variation of Integrated Particulate Back scatter (IβP) in UT (a) and LS

(b) superimposed on the annual variation of CAPE for the period 1998-2003.Vertical bars represent the

Though during the December–February period CAPE is relatively small and thunderstorm activity also is less, a broad peak in ΙβP at 18–20 km is observable during this period (Figure 13).This feature though appears to be contradicting to the above hypothesis proposed for the summer monsoon period, the real mechanism responsible for the transport of tropospheric air into the stratosphere during winter is different from that during the monsoon period. During winter the intrusion of tropospheric air into the lower stratosphere occurs through vertical ascent of air driven by the strength of the ascending branch of the Hadley cell

LS region as indicated by the large error bars during the May-October period.

As can be seen from Figure 10a the tropopause is cooler and higher during the January– March period and warmer and lower during August–October period which is in accordance with what is reported for tropical locations [67-70]. These Seasonal changes in the cold point tropopause altitude also contribute to the mass influx to stratosphere [71-73]. Transport of constituents like water vapor, aerosols, trace gases etc across the tropopause which acts as a permeable membrane through which continuous mixing between tropospheric and stratospheric air takes place is well demonstrated by various investigators [74-77]. Over and above the penetration of deep convection, this annual adjustment of tropopause altitude is an important mechanism for the stratosphere troposphere exchange.
