**12. Influence of dynamics in the latitude variation of particulates in the UTLS region**

Distribution of Particulates in the Tropical UTLS over

the Asian Summer Monsoon Region and Its Association with Atmospheric Dynamics 137

ITCZ along with the development of deep convection over the Bay of Bengal (particularly over its northern parts) during this period aid the formation of abundant dense STCs in the UT region. The annual variation of high altitude cloud amount over the Indian region also shows the presence of large deep convective cloud systems reaching very high altitudes and even penetrating the tropopause has been reported [81]. As the strength of the tropospheric convection decreases significantly during winter and the ITCZ shifts to the southern hemisphere sector, the probability of occurrence of dense STC over the southern hemisphere increases significantly. The deep convective clouds over this region get confined to a small geographical region between equator and 10°. Most of the STCs occurring over the Indian region during this period will be of *in situ* origin [78]. These clouds will be either ultra-thin

The latitude variation of the annual pattern of τP in the LS1 and LS2 regions are presented in Figure 18b and Figure 18c respectively. In general, the mean τP in the LS1 (18-21 km) region is in the range of 0.0008-0.003, which is one order in magnitude less than that in UT. In this altitude region, τP clearly shows relatively high values in the off-equatorial (north of 15°) regions and low values in the equatorial regions. Studies [82] on stratospheric aerosol optical depth during the decay phase of the volcanic aerosol at mid-latitudes using SAGE-II data have shown the presence of sinusoidal variation in the aerosol optical depth superimposed on an exponential decay with maximum and minimum occurring during the local winter and summer respectively. Over tropics the amplitude of these oscillations are significantly small and hence gets submerged in the disturbances caused by subsequent minor volcanic eruptions. As the period selected for the present analysis is volcanically quiescent, these oscillations are well discernable over the tropics. In the region 10-15°N, the annual variation of τP shows a relatively high value during winter and low value during summer. Note that, this variation is very similar to the annual variation of particulate backscatter observed in the lidar data from Gadanki [7]. Above 15°N, high values of τP are observed in the LS1 region during the May-August period centered around 20-25°N. The particulate extinction just above the tropopause also shows a significant enhancement in the latitude region between 20-25°N (Figure 18b). This could be due to the penetration of the top of the high altitude semitransparent cirrus clouds above the cold point (tropopause). The two peaks observed in CAPE (and TSD) near equator (Figure 17) during April and October becomes more prominent and gets closer with increase in latitude (towards north) and merges to become a strong broad peak centered around July. This shows that the convective activity in the 20-30°N is very strong and the outflow occurs very close to the tropopause. The frequency of occurrence of cloud top altitude (observation from CALIPSO) shows a maximum value of 17±0.5 km in this latitude region during the June to September period which is at least 1 km larger than the maximum value observed at other latitudes (between 30°S to 30°N) at any period [81]. Thus convection over the Indian land mass during the summer monsoon period in the 20-30°N latitude band is the strongest one in the entire latitude region 30°S to 30°N at any time during the year. This is a characteristic feature for the Indian longitude region. Strong convection plays a major role in transporting particulates from the upper tropospheric cirrus cluster to lower stratosphere causing a pronounced increase in τP in the LS1 region. Note that, such a feature in convection and

or sub-visual type cirrus with very low values of optical depth (<0.03).

To delineate the latitude variation of αp in the UTLS region over the Indian longitude sector (70-90°E), the annual variation of monthly mean τP for a latitude band of 5°width in the latitude region 30°S to 30°N, for the three altitude regions 12-16 km (UT), 18-21 km (LS1) and 21-30 km (LS2) are examined separately. Figure 18 shows the contour plots of the mean τP in the UT, LS1 and LS2 regions with month along x-axis and latitude along y-axis. In the UT region, τP shows a general decrease with increase in latitude from equator, with its gradient showing a pronounced variation from month-to-month. The summer-winter contrast (with relatively low values during winter and high values during summer) is well discernable in the UT region (Figure 18a). Relatively high values of τP observed in the UT region between 15°S-15°N during May to February period are mostly due to presence of dense STCs resulting from the outflow of convective anvils. These are the periods when convective activity in the troposphere is very strong (associated with the southwest and northeast monsoons) in this region. Above 15°N, relatively high values of τP are more-or-less confined to the June-August period when the monsoon trough usually reaches its extreme north over the continent. However, beyond 15°S, the values of τP decreases significantly with increase in latitude, compared to that observed in the northern hemisphere.

**Figure 18.** Contour plots showing the latitudinal dependence of the annual variation of τp obtained from SAGE-II data in the altitude regions 12–16 km (UT), 18–21 km (LS1) and 21–30 km (LS2) for the period 1998-2005.

The latitudinal variation of the annual pattern of τP in the UT region shows fairly good correspondence with CAPE and mean thunderstorm activity (Figure 17) over the northern hemisphere. The annual variation of TSD in the latitude region 8-15°N shows two peaks; one during the April-May period and then another during October. Further north, these two peaks get closer and merge to form a single peak in the latitude belt 20-30°N during the June-September period. This corresponds well with the latitudinal variation of the annual pattern of τP in the UT region (Figure 18a). Increase in convective activity (as well as convective outflow) over the Indian region associated with the northward migration of ITCZ along with the development of deep convection over the Bay of Bengal (particularly over its northern parts) during this period aid the formation of abundant dense STCs in the UT region. The annual variation of high altitude cloud amount over the Indian region also shows the presence of large deep convective cloud systems reaching very high altitudes and even penetrating the tropopause has been reported [81]. As the strength of the tropospheric convection decreases significantly during winter and the ITCZ shifts to the southern hemisphere sector, the probability of occurrence of dense STC over the southern hemisphere increases significantly. The deep convective clouds over this region get confined to a small geographical region between equator and 10°. Most of the STCs occurring over the Indian region during this period will be of *in situ* origin [78]. These clouds will be either ultra-thin or sub-visual type cirrus with very low values of optical depth (<0.03).

136 Atmospheric Aerosols – Regional Characteristics – Chemistry and Physics

in latitude, compared to that observed in the northern hemisphere.

**Figure 18.** Contour plots showing the latitudinal dependence of the annual variation of τp obtained from SAGE-II data in the altitude regions 12–16 km (UT), 18–21 km (LS1) and 21–30 km (LS2) for the

The latitudinal variation of the annual pattern of τP in the UT region shows fairly good correspondence with CAPE and mean thunderstorm activity (Figure 17) over the northern hemisphere. The annual variation of TSD in the latitude region 8-15°N shows two peaks; one during the April-May period and then another during October. Further north, these two peaks get closer and merge to form a single peak in the latitude belt 20-30°N during the June-September period. This corresponds well with the latitudinal variation of the annual pattern of τP in the UT region (Figure 18a). Increase in convective activity (as well as convective outflow) over the Indian region associated with the northward migration of

**UTLS region** 

period 1998-2005.

**12. Influence of dynamics in the latitude variation of particulates in the** 

To delineate the latitude variation of αp in the UTLS region over the Indian longitude sector (70-90°E), the annual variation of monthly mean τP for a latitude band of 5°width in the latitude region 30°S to 30°N, for the three altitude regions 12-16 km (UT), 18-21 km (LS1) and 21-30 km (LS2) are examined separately. Figure 18 shows the contour plots of the mean τP in the UT, LS1 and LS2 regions with month along x-axis and latitude along y-axis. In the UT region, τP shows a general decrease with increase in latitude from equator, with its gradient showing a pronounced variation from month-to-month. The summer-winter contrast (with relatively low values during winter and high values during summer) is well discernable in the UT region (Figure 18a). Relatively high values of τP observed in the UT region between 15°S-15°N during May to February period are mostly due to presence of dense STCs resulting from the outflow of convective anvils. These are the periods when convective activity in the troposphere is very strong (associated with the southwest and northeast monsoons) in this region. Above 15°N, relatively high values of τP are more-or-less confined to the June-August period when the monsoon trough usually reaches its extreme north over the continent. However, beyond 15°S, the values of τP decreases significantly with increase

The latitude variation of the annual pattern of τP in the LS1 and LS2 regions are presented in Figure 18b and Figure 18c respectively. In general, the mean τP in the LS1 (18-21 km) region is in the range of 0.0008-0.003, which is one order in magnitude less than that in UT. In this altitude region, τP clearly shows relatively high values in the off-equatorial (north of 15°) regions and low values in the equatorial regions. Studies [82] on stratospheric aerosol optical depth during the decay phase of the volcanic aerosol at mid-latitudes using SAGE-II data have shown the presence of sinusoidal variation in the aerosol optical depth superimposed on an exponential decay with maximum and minimum occurring during the local winter and summer respectively. Over tropics the amplitude of these oscillations are significantly small and hence gets submerged in the disturbances caused by subsequent minor volcanic eruptions. As the period selected for the present analysis is volcanically quiescent, these oscillations are well discernable over the tropics. In the region 10-15°N, the annual variation of τP shows a relatively high value during winter and low value during summer. Note that, this variation is very similar to the annual variation of particulate backscatter observed in the lidar data from Gadanki [7]. Above 15°N, high values of τP are observed in the LS1 region during the May-August period centered around 20-25°N. The particulate extinction just above the tropopause also shows a significant enhancement in the latitude region between 20-25°N (Figure 18b). This could be due to the penetration of the top of the high altitude semitransparent cirrus clouds above the cold point (tropopause). The two peaks observed in CAPE (and TSD) near equator (Figure 17) during April and October becomes more prominent and gets closer with increase in latitude (towards north) and merges to become a strong broad peak centered around July. This shows that the convective activity in the 20-30°N is very strong and the outflow occurs very close to the tropopause. The frequency of occurrence of cloud top altitude (observation from CALIPSO) shows a maximum value of 17±0.5 km in this latitude region during the June to September period which is at least 1 km larger than the maximum value observed at other latitudes (between 30°S to 30°N) at any period [81]. Thus convection over the Indian land mass during the summer monsoon period in the 20-30°N latitude band is the strongest one in the entire latitude region 30°S to 30°N at any time during the year. This is a characteristic feature for the Indian longitude region. Strong convection plays a major role in transporting particulates from the upper tropospheric cirrus cluster to lower stratosphere causing a pronounced increase in τP in the LS1 region. Note that, such a feature in convection and

hence in τP (in the LS1 region) is not observed in the southern hemisphere. In LS2, τP shows a distinct latitude variation with relatively high values near the equator up to 15° in both the hemispheres and low values over the off-equatorial regions (>15°). Relatively high values of τP are observed in the LS2 region during the January-June period in the latitude region 15°S-10°N and low values during the rest of the period. This pattern slowly reverses with increase in latitude. Beyond 10°N, τP shows a pronounced winter peak with low values during the March-May period. The annual variation of τP in the LS2 region over the southern hemisphere is quite similar to that in the northern hemisphere. High values of τP observed between 15°S and 15°N in the LS2 region confirms the presence of a Tropical Stratospheric aerosol Reservoir (TSR) during the study period. Earlier studies carried out by several investigators [29,83] revealed the presence of this band structure (with high aerosol loading) in the 21- 30 km altitude region over the equatorial region during volcanically perturbed period. Examining the aerosol climatology in the LS region using the SAGE-II data both during the volcanically perturbed period as well as during the near background conditions, *Bauman et al.* [84] reported maximum aerosol optical depth near the tropics and minimum between 15-45°�latitudes. *Trepte and Hitchman* [28] were the first to propose the existence of a low-latitude maximum in lower stratospheric aerosol optical depth which they referred to as the 'tropical aerosol reservoir'. They also examined the post-volcanic aerosol distribution in tropics and observed that the 18-21 km (LS1) region experiences a rapid pole ward transport, while in the upper regime (LS2) aerosol lofting and subsequent accumulation occurs within 20°S-20°N. This can fairly well explain the observed low values of τp in the lower regime (LS1 region) and high values in the upper regime (LS2 region) in the equatorial region between 15°S to 15°N in the present study. The variation of τP in the UT region in both these hemisphere are more-or-less complementary to each other indicating similar seasonal dependence in both the hemispheres, except for the fact that decrease in τP with increase in latitude (from equator) towards south is much faster than that in northern hemisphere. This is quite expected because the southern hemispheric sector is mostly occupied by ocean. The seasonal influence is rather insignificant in the LS1 and LS2 regions.

Distribution of Particulates in the Tropical UTLS over

the Asian Summer Monsoon Region and Its Association with Atmospheric Dynamics 139

stratospheric background level in the recent past [26]. An overview of the current understanding on stratospheric aerosol science can be found in *Thomason and Peter* [24].

Though, in general, most of the particles generated through the gas-to-particle conversion process in the stratosphere will be small and nearly spherical in nature, there could be a few larger size particles in the lower stratosphere associated with moderate and intense volcanic eruptions, leading to a pronounced enhancement in δ and αp in this region. These volcanic perturbations are clearly distinguishable from STCs in the UT region based on the amount of enhancement, its temporal structure as well as the duration of enhancements. Moreover, while δ (and αp) of STCs vary significantly at shorter time scales [57,87], the stratospheric cloud formed through volcanic emissions will be stable for a longer period. In addition, the values of δ associated with STC will be significantly larger (as they are mostly composed of non-spherical ice crystals) than those of volcanic clouds. The vertical structure of δ for the volcanic cloud will remain fairly stable at short time scales typically over a night. Thus a long lasting enhancement of δ in the LS region observed by the lidar is an indicator for assessing the volcanic impact on stratospheric aerosols. In order to illustrate this in detail and depict the difference in the nature of perturbations due to STC and volcanic cloud, a contour map of δ for a typical night (25 November 2002) during the eruption period of Reventador (started in November 2002 and lasted up to January 2003) is presented in Figure 19a. In order to accommodate the large variations, the contouring of δ in this figure is performed in two bands; one from 0.04 to 0.2 in steps of 0.02 and the other above 0.2 at 0.2 interval. Two enhanced layers (of δ) one between 15 and 16 km and the other around 19km, are distinctly seen in this figure. The value of δ and its temporal variations are very large for the lower layer while these are very small for the upper layer. The lower layer disappears after mid-night while the upper layer continues to persist up to the end of lidar observations. This shows that the lower layer is an STC while the upper one is the volcanic cloud. This volcanic cloud will persist on subsequent night also while STC may or may not

**Figure 19.** Altitude- cross-section of volume depolarization ratio (δ) during the night of 25 November 2002 (a). Altitude profiles of δ on different nights (averaged for one hour around 22:00 IST) for the period November 2002 to March 2003, during the active and post -active phases of Mt. Reventador (b).

## **13. Influence of moderate volcanic eruptions on** α**p in the LS region**

The background stratospheric aerosol layer usually referred to as the Junge layer [85] consists of liquid droplets composed of a mixture of sulfuric acid and water. This layer will be quite prominent subsequent to major volcanic eruptions, such as El Chichón (Mexico, 1982) and Mount Pinatubo (Philippines, 1991), which are powerful enough to inject large amount of SO2 into the stratosphere [86]. After oxidation, sulfate aerosols are formed at these altitudes. These particles are subsequently distributed globally depending on the latitude of the eruption. Removal of these aerosols is rather difficult. It takes several months to years (depending on size) to scavenge these volcanic aerosols. The last major eruption (Mount Pinatubo) took place in 1991 and the stratospheric aerosol layer returned to its "background" level around 1997. There is no major increase in stratospheric aerosol loading after 1997 [13,25]. However, increase in the anthropogenic SO2 emission has been proposed as a plausible mechanism responsible for the observed small increasing trend in stratospheric background level in the recent past [26]. An overview of the current understanding on stratospheric aerosol science can be found in *Thomason and Peter* [24].

138 Atmospheric Aerosols – Regional Characteristics – Chemistry and Physics

hence in τP (in the LS1 region) is not observed in the southern hemisphere. In LS2, τP shows a distinct latitude variation with relatively high values near the equator up to 15° in both the hemispheres and low values over the off-equatorial regions (>15°). Relatively high values of τP are observed in the LS2 region during the January-June period in the latitude region 15°S-10°N and low values during the rest of the period. This pattern slowly reverses with increase in latitude. Beyond 10°N, τP shows a pronounced winter peak with low values during the March-May period. The annual variation of τP in the LS2 region over the southern hemisphere is quite similar to that in the northern hemisphere. High values of τP observed between 15°S and 15°N in the LS2 region confirms the presence of a Tropical Stratospheric aerosol Reservoir (TSR) during the study period. Earlier studies carried out by several investigators [29,83] revealed the presence of this band structure (with high aerosol loading) in the 21- 30 km altitude region over the equatorial region during volcanically perturbed period. Examining the aerosol climatology in the LS region using the SAGE-II data both during the volcanically perturbed period as well as during the near background conditions, *Bauman et al.* [84] reported maximum aerosol optical depth near the tropics and minimum between 15-45°�latitudes. *Trepte and Hitchman* [28] were the first to propose the existence of a low-latitude maximum in lower stratospheric aerosol optical depth which they referred to as the 'tropical aerosol reservoir'. They also examined the post-volcanic aerosol distribution in tropics and observed that the 18-21 km (LS1) region experiences a rapid pole ward transport, while in the upper regime (LS2) aerosol lofting and subsequent accumulation occurs within 20°S-20°N. This can fairly well explain the observed low values of τp in the lower regime (LS1 region) and high values in the upper regime (LS2 region) in the equatorial region between 15°S to 15°N in the present study. The variation of τP in the UT region in both these hemisphere are more-or-less complementary to each other indicating similar seasonal dependence in both the hemispheres, except for the fact that decrease in τP with increase in latitude (from equator) towards south is much faster than that in northern hemisphere. This is quite expected because the southern hemispheric sector is mostly occupied by ocean. The seasonal influence is rather insignificant in the LS1 and LS2 regions.

**13. Influence of moderate volcanic eruptions on** α**p in the LS region** 

The background stratospheric aerosol layer usually referred to as the Junge layer [85] consists of liquid droplets composed of a mixture of sulfuric acid and water. This layer will be quite prominent subsequent to major volcanic eruptions, such as El Chichón (Mexico, 1982) and Mount Pinatubo (Philippines, 1991), which are powerful enough to inject large amount of SO2 into the stratosphere [86]. After oxidation, sulfate aerosols are formed at these altitudes. These particles are subsequently distributed globally depending on the latitude of the eruption. Removal of these aerosols is rather difficult. It takes several months to years (depending on size) to scavenge these volcanic aerosols. The last major eruption (Mount Pinatubo) took place in 1991 and the stratospheric aerosol layer returned to its "background" level around 1997. There is no major increase in stratospheric aerosol loading after 1997 [13,25]. However, increase in the anthropogenic SO2 emission has been proposed as a plausible mechanism responsible for the observed small increasing trend in Though, in general, most of the particles generated through the gas-to-particle conversion process in the stratosphere will be small and nearly spherical in nature, there could be a few larger size particles in the lower stratosphere associated with moderate and intense volcanic eruptions, leading to a pronounced enhancement in δ and αp in this region. These volcanic perturbations are clearly distinguishable from STCs in the UT region based on the amount of enhancement, its temporal structure as well as the duration of enhancements. Moreover, while δ (and αp) of STCs vary significantly at shorter time scales [57,87], the stratospheric cloud formed through volcanic emissions will be stable for a longer period. In addition, the values of δ associated with STC will be significantly larger (as they are mostly composed of non-spherical ice crystals) than those of volcanic clouds. The vertical structure of δ for the volcanic cloud will remain fairly stable at short time scales typically over a night. Thus a long lasting enhancement of δ in the LS region observed by the lidar is an indicator for assessing the volcanic impact on stratospheric aerosols. In order to illustrate this in detail and depict the difference in the nature of perturbations due to STC and volcanic cloud, a contour map of δ for a typical night (25 November 2002) during the eruption period of Reventador (started in November 2002 and lasted up to January 2003) is presented in Figure 19a. In order to accommodate the large variations, the contouring of δ in this figure is performed in two bands; one from 0.04 to 0.2 in steps of 0.02 and the other above 0.2 at 0.2 interval. Two enhanced layers (of δ) one between 15 and 16 km and the other around 19km, are distinctly seen in this figure. The value of δ and its temporal variations are very large for the lower layer while these are very small for the upper layer. The lower layer disappears after mid-night while the upper layer continues to persist up to the end of lidar observations. This shows that the lower layer is an STC while the upper one is the volcanic cloud. This volcanic cloud will persist on subsequent night also while STC may or may not

**Figure 19.** Altitude- cross-section of volume depolarization ratio (δ) during the night of 25 November 2002 (a). Altitude profiles of δ on different nights (averaged for one hour around 22:00 IST) for the period November 2002 to March 2003, during the active and post -active phases of Mt. Reventador (b).

be present. To illustrate this the sequences of lidar profiles observed on a few nights during the eruption period of Reventador are presented in Figure 19b.The features of the layer located around 19 km (in the LS region) is rather steady in all these profiles while those of the layer below ~17 km (due to STC), is highly variable. The stratospheric layer is strongest in December and started decaying in January 2003 and became almost insignificant by March 2003, while the STC layers appear at random.

Distribution of Particulates in the Tropical UTLS over

the Asian Summer Monsoon Region and Its Association with Atmospheric Dynamics 141

stratosphere for four different latitude belts (over tropics) are used mainly to have sufficient number of profiles in each belt as well as owing to the fact that in the lower stratosphere the spatial variability along longitude could be minimal (because of efficient mixing in the zonal direction and strong horizontal transport prevailing in the region). As expected the aerosol extinction decreases with increase in altitude with significant loading confined to the altitude region 18–27 km. Short-lived enhancement in αp are distinctly seen in the lower stratosphere (Figure 20a) in different years. These signatures could be attributed to the influence of various minor volcanic eruptions during the study period. Those eruptions which could make discernable impact in the stratosphere are marked at the start of the respective eruption along the X-axis of each panel and further details of these eruptions are summarized in Table 1. Only those eruptions which occurred between 30°S and 30°N are included in Table 1. The Volcanic Explosivity Index (VEI) for different eruptions [89] is obtained from the Web site of Smithsonian Institution–Global Volcanism Program (GVP). The eruptions of Ulawun (eruption d), Ruang (eruption e), Reventador (eruption f), and Manam (eruption j) are relatively strong with VEI ~4 (Table 1), and the signature of these eruptions are well discernable in the lower stratospheric (18–20 km) aerosol extinction (Figure 20a). These perturbations can be distinctly seen both over the equatorial as well as

**Table 1.** Details of volcanic eruptions in the tropics during 1998–2005 having significant stratospheric

Figure 20b shows the temporal variation of the altitude structure of αp obtained from lidar during the period 1998– 2005 in the form of a contour plot. This figure also depicts the signatures of various minor volcanic eruptions similar to that depicted in the zonal mean values obtained from SAGE-II data, in the latitude sector 0–15°N (Figure 20a). Small-scale features are more pronounced in lidar data (Figure 20b) mainly because of the fact that it corresponds to a point observation while that in Figure20a is the zonal average. In volcanically quiescent periods the sulfur bearing gases, SO2 and OCS emitted from the earth's surface are transported across the tropopause [90], photolyzed, and oxidized to sulfuric acid before condensation to form sulfuric acid and water droplets in the stratosphere [24,91], which are, by their liquid nature, spherical. The depolarization caused by these particles will be very small and hence the resultant δ in the stratosphere will be

off-equatorial regions.

impact

The perturbation in the lower stratosphere over the equatorial and off-equatorial regions during the period 1998-2005, when the volcanic activity was relatively small, are examined in detail [88] using the altitude profiles of αp over the tropics from SAGE-II data archive (Figure 20a) and the altitude profiles of αp and δ obtained from lidar data (Figure 20b and 20c) at Gadanki. Zonal-averaged monthly mean αp at different altitudes in the lower

**Figure 20.** (a) Zonally averaged monthly mean extinction coefficient (αP) for the equatorial (0°S–15°S and 0°N–15°N) and off-equatorial (15°N–30°N and 15°S–30°S) regions estimated from SAGE-II data, (b) altitude-time cross section of aerosol extinction coefficient (αP) from lidar and (c) altitude-time cross section of the Volume Depolarization Ratio from lidar ( at Gadanki) during the period 1998–2005. The letters along the abscissa represent the eruption of each volcano, the name and other details of which are listed in Table 1. The major ticks correspond to December of each year.

stratosphere for four different latitude belts (over tropics) are used mainly to have sufficient number of profiles in each belt as well as owing to the fact that in the lower stratosphere the spatial variability along longitude could be minimal (because of efficient mixing in the zonal direction and strong horizontal transport prevailing in the region). As expected the aerosol extinction decreases with increase in altitude with significant loading confined to the altitude region 18–27 km. Short-lived enhancement in αp are distinctly seen in the lower stratosphere (Figure 20a) in different years. These signatures could be attributed to the influence of various minor volcanic eruptions during the study period. Those eruptions which could make discernable impact in the stratosphere are marked at the start of the respective eruption along the X-axis of each panel and further details of these eruptions are summarized in Table 1. Only those eruptions which occurred between 30°S and 30°N are included in Table 1. The Volcanic Explosivity Index (VEI) for different eruptions [89] is obtained from the Web site of Smithsonian Institution–Global Volcanism Program (GVP). The eruptions of Ulawun (eruption d), Ruang (eruption e), Reventador (eruption f), and Manam (eruption j) are relatively strong with VEI ~4 (Table 1), and the signature of these eruptions are well discernable in the lower stratospheric (18–20 km) aerosol extinction (Figure 20a). These perturbations can be distinctly seen both over the equatorial as well as off-equatorial regions.

140 Atmospheric Aerosols – Regional Characteristics – Chemistry and Physics

March 2003, while the STC layers appear at random.

be present. To illustrate this the sequences of lidar profiles observed on a few nights during the eruption period of Reventador are presented in Figure 19b.The features of the layer located around 19 km (in the LS region) is rather steady in all these profiles while those of the layer below ~17 km (due to STC), is highly variable. The stratospheric layer is strongest in December and started decaying in January 2003 and became almost insignificant by

The perturbation in the lower stratosphere over the equatorial and off-equatorial regions during the period 1998-2005, when the volcanic activity was relatively small, are examined in detail [88] using the altitude profiles of αp over the tropics from SAGE-II data archive (Figure 20a) and the altitude profiles of αp and δ obtained from lidar data (Figure 20b and 20c) at Gadanki. Zonal-averaged monthly mean αp at different altitudes in the lower

**Figure 20.** (a) Zonally averaged monthly mean extinction coefficient (αP) for the equatorial (0°S–15°S and 0°N–15°N) and off-equatorial (15°N–30°N and 15°S–30°S) regions estimated from SAGE-II data, (b) altitude-time cross section of aerosol extinction coefficient (αP) from lidar and (c) altitude-time cross section of the Volume Depolarization Ratio from lidar ( at Gadanki) during the period 1998–2005. The letters along the abscissa represent the eruption of each volcano, the name and other details of which

are listed in Table 1. The major ticks correspond to December of each year.


**Table 1.** Details of volcanic eruptions in the tropics during 1998–2005 having significant stratospheric impact

Figure 20b shows the temporal variation of the altitude structure of αp obtained from lidar during the period 1998– 2005 in the form of a contour plot. This figure also depicts the signatures of various minor volcanic eruptions similar to that depicted in the zonal mean values obtained from SAGE-II data, in the latitude sector 0–15°N (Figure 20a). Small-scale features are more pronounced in lidar data (Figure 20b) mainly because of the fact that it corresponds to a point observation while that in Figure20a is the zonal average. In volcanically quiescent periods the sulfur bearing gases, SO2 and OCS emitted from the earth's surface are transported across the tropopause [90], photolyzed, and oxidized to sulfuric acid before condensation to form sulfuric acid and water droplets in the stratosphere [24,91], which are, by their liquid nature, spherical. The depolarization caused by these particles will be very small and hence the resultant δ in the stratosphere will be

very close to that of the molecules. In general, it is in the range 0.03 to 0.04. But, during major volcanic eruption abundant amount of precursor gases will be injected into the lower stratosphere along with a few fine particulates. Because of this influx of particles and gases there will be an increase in the number density of particles as well as an increase in the size of these particles. The size spectrum of the stratospheric particles also shifts toward the larger size regime following the volcanic eruption [92]. As some of those particles that are directly injected into the lower stratosphere during volcanic eruption could be non-spherical in nature, an increase in δ in the stratospheric aerosol would be expected.

Distribution of Particulates in the Tropical UTLS over

the Asian Summer Monsoon Region and Its Association with Atmospheric Dynamics 143

**Figure 21.** Time series of zonal mean monthly average τP in the altitude region 18–28 km obtained from SAGE-II for (a) the equatorial region (0°–15°N, 0°–15°S) and (b) the off-equatorial region (15°–30°N and

of 4.8% and 6.3% per year (with respect to its value in 2002) was reported [26] at Hawaii (19.5°N, 155.6°W) (Mauna Loa Laboratory) as well as at Boulder (40°N) (Colorado). The increase in τp at Hawaii and Boulder was attributed to the increase in global coal consumption since 2002, mainly from China, and subsequent increase in emission of SO2. Anthropogenic aerosols produced through gas-to-particle conversion of precursor gases like sulfates and ammonia transported to the upper troposphere [5] through intense convection [93] in the tropics and subsequently across the tropopause also could possibly be a contributing factor for this increase. An increase in tropical upwelling (Brewer- Dobson circulation) because of global warming also was suggested to be a plausible mechanism for

Even though on an average the stratospheric particulate loading is in its background level during the period 1998–2005, it was influenced particularly by a few moderate volcanic eruptions mainly after September 2002. While the period before September 2002 was absolutely quiet (with low particulate loading), the later period was mildly disturbed. The variation in stratospheric particulate loading need not solely be represented by a corresponding variation in tephra emissions Their could be some other causative mechanisms, such as increase in anthropogenic emissions as well as the increase in tropical

15°–30°S), along with (c) the mean τP obtained from lidar at Gadanki.

the observed increasing trend [93,94] in stratospheric βp after 2002.

upwelling, which could influence the stratospheric particulate loading.

Figure 20c shows a contour plot of δ at different altitude in the lower stratosphere over Gadanki during the period 1998–2005. This plot generated adopting the same procedure as that used for generating Figure 20b from αp profiles, clearly shows a few short-lived δ enhancements in the lower stratosphere. The sporadic increase in δ is associated with the eruption of a few moderately intense volcanic eruptions. The duration of the increase in α<sup>p</sup> also matches well that of δ. The disturbances caused by eruption of volcanoes Ruang (eruption e) and Reventador (eruption f) are relatively stronger (δ ranging from 0.05 to 0.2 in the altitude region 18–21 km during November 2002 to February 2003). In a few cases the enhancements in αp and δ does not match exactly. This could be due to the fact that the volcanic locations are at different distance from Gadanki as well as the prevailing transport process could be different at various occasions.
