Fig 11 - uploaded by Andrew Gettelman
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Schematic Figure of the Tropical Tropopause Layer as discussed in the text. Altitude of Cold Point Tropopause and Minimum Lapse Rate, and Q clear 1⁄4 0 shown. Horizontal motion indicated as vectors in and out of the page (it also occurs along the zonal dimension). Thin black arrows are radiative heating, large grey arrows are the stratospheric (Brewer-Dobson) circulation. Horizontal arrows are the detrainment from convection.
Source publication
The tropical tropopause layer (TTL) is a transition region between the troposphere and the stratosphere. In this study the vertical extent of the TTL is diagnosed from radiosonde and ozonesonde profiles in the tropics and a climatology of this layer is presented. The radiative balance in the TTL is also characterized. The TTL is locally defined as...
Contexts in source publication
Context 1
... TTL is a region of transition between tropospheric stability dominated by deep convection, and stratospheric stability caused by ozone heating. Convenient boundaries for this layer are the lapse rate minimum (LRM) at 10 – 12 km (340 – 350 K in isentropic coordi- nates), and the cold point tropopause (CPT) at 16 – 17 km (380 – 390 K). Both of these values are readily calculated from temperature soundings. The location of these levels is indicated schematically in Fig. 11. Important zonal variations in the depth of the TTL are also present, and these are noted schematically in Fig. 11. The TTL thus de fi ned is related to ozone and convection in the region. The altitude of the lapse rate minimum at 10 – 12 km corresponds to a level generally just below the minimum in ozone in the upper troposphere. The mean position of the LRM also corresponds to the main in fl uence of convection in the upper troposphere as observed from satellite cloud ob- servations and divergence of the wind fi eld. The variations of the LRM level are also signi fi cantly correlated with convective activity. In contrast, we fi nd that convective activity is not strongly correlated with the cold point, the altitude of which varies coherently with the annual cycle of the stratospheric wave driven circulation (Yulaeva et al. 1994). The cold point is actually lower over some convective regions (Seidel et al. 2001), as illustrated in Fig. 11. The spatial distribution of cold point temperatures is likely a combination of several factors related to convection, including a direct cooling effect from mixing low- y air from deep convection, an indirect dynamic response to the re- lease of latent heat in convection at lower levels (Highwood and Hoskins 1998), and the radiative cooling associated with the presence of a thick cloud deck below the cold point (Gage et al. 1991; Norton 2001). The TTL has a lower boundary dominated by tropospheric processes (convection) and an upper boundary dominated by stratospheric radiative heating and a remote response to convection. This analysis does not imply that convection is not important above the level of the LRM. As shown by Sherwood and Dessler (2001), even a small fraction of convection can have a signi fi cant effect on water vapor and ozone in the tropopause region. A small fraction of the time (less than 0.5%), convection does penetrate through the tropopause layer (Gettelman et al. 2002), as indicated in Fig. 11. In addition, we do not consider here the effect of horizontal (quasi-isentropic) motions within the TTL or between the TTL and the extratropical lower stratosphere. These motions, which have a preferential locations in space and time (Waugh and Polvani 2000), are likely to be important for determining the distribution of trace species in the TTL. The LRM at 10 – 12 km (350 K) is well below the level at which the clear sky radiative heating changes from negative to positive (16 km or 370 K). This implies that the main convective detrainment does not reach a level from which it can rise into the stratosphere. Deeper convection, which reaches within 1 km of the CPT or higher, is necessary to loft air to altitudes where it can reach the stratosphere. There are coherent variations over the year in the altitude and depth of the tropical tropopause layer. Over the Western Paci fi c, the main convective out fl ow level is higher (and at higher potential temperature) than over the Eastern Paci fi c. Variations over the seasons follow the motion of convection. However, the CPT has a single and coherent annual cycle throughout the tropics (Seidel et al. 2001). Zonal variations of the level of the LRM are around 1 km, and the LRM is lower in non-convective regions, shown schematically in Fig. 11. The level of the LRM is at lower potential temperatures in the subtropics. These regional variations in depth of layer may be signi fi cant for short lived species. In some regions short lived species will be exported into the TTL at higher altitudes than in other regions. Recognizing that isentropic motion within the TTL subsequent to convection will rapidly transport parcels horizontally, and that below 16 km parcels are subsiding, this implies that entrainment of TTL air into convection is necessary to move parcels in the TTL upward to within 1 km of the tropopause. This may have implications for the transport of, for example, the products of biomass burning into the TTL, because the burning generally occurs in dry seasons, when convection is not present (though sometimes the fi res may them- selves cause convection). Analysis of chemical budgets in the TTL (of carbon monoxide for ex- ample), might shed some light on this hypothe- sis that ef fi cient transport into the TTL re- quires active convection. Variations of the tropopause layer on interannual time scales have also been diagnosed. There is a strong signal of the El-Nin ̃ o Southern Oscillation, which has a strong affect on the organization of convection and on the height of the LRM, particularly in the Paci fi c. Over the last 40 years, the cold point tropopause has risen 200 – 400 m, consistent with previous work (Seidel et al. 2001). The altitude of the LRM is also rising at a similar rate, though the variability is higher at this level, so trends are not as signi fi cant. These trends appear consistent with increasing surface temperatures over the last 40 years, assuming that tropospheric lapse rates have remained fi xed. We may hy- pothesize that the TTL may continue to rise if the warming expected from increased green- house gases is realized. Such a change is consistent with the sensitivity experiments of Thuburn and Craig (2002) which showed little sensitivity in a radiative-convective model to expected anthropogenic changes in carbon dioxide, because their surface temperatures remained fi xed. The analysis allows us to re fi ne the sche- matics of the tropical tropopause layer recently presented by Highwood and Hoskins (1998), Folkins et al. (1999), Sherwood and Dessler (2001) and V ̈ mel et al. (2002). Figure 11 indicates some of the important fi ndings of this paper. The TTL can be de fi ned as the region between the lapse rate minimum and the cold point. It has a varying thickness changing with convective activity, and is narrower with a higher base near active convective regions. The level of zero clear sky radiative heating is quite close to the tropopause. Two of the major issues not addressed in this study are the impact of horizontal transport, and the related issue of linkages between the tropical TTL and the extratropical tropopause, which is located at 10 – 12 km, nearly coincident with the base of the TTL. This horizontal motion is also indicated by vectors in Fig. 11. While the temperature structure may not be directly affected by such isentropic motion, the distribution of trace species like water vapor and ozone (and hence the radiative heating) may be affected. A future picture of the tropical tropopause layer would need to consider the role of the subtropics and the extratropics. We would like to thank A.E. Dessler and S.C. Sherwood for advice and assistance with ac- cessing data. Thanks to J. Witte for assistance with the SHADOZ archive. We also thank J. Thuburn, H. V ̈ mel and I. Folkins for provid- ing preprints of their work. This work also bene fi tted from numerous discussions with S. Massie, L. Pan and W.J. Randel, and from the care of two anonymous ...
Context 2
... TTL is a region of transition between tropospheric stability dominated by deep convection, and stratospheric stability caused by ozone heating. Convenient boundaries for this layer are the lapse rate minimum (LRM) at 10 – 12 km (340 – 350 K in isentropic coordi- nates), and the cold point tropopause (CPT) at 16 – 17 km (380 – 390 K). Both of these values are readily calculated from temperature soundings. The location of these levels is indicated schematically in Fig. 11. Important zonal variations in the depth of the TTL are also present, and these are noted schematically in Fig. 11. The TTL thus de fi ned is related to ozone and convection in the region. The altitude of the lapse rate minimum at 10 – 12 km corresponds to a level generally just below the minimum in ozone in the upper troposphere. The mean position of the LRM also corresponds to the main in fl uence of convection in the upper troposphere as observed from satellite cloud ob- servations and divergence of the wind fi eld. The variations of the LRM level are also signi fi cantly correlated with convective activity. In contrast, we fi nd that convective activity is not strongly correlated with the cold point, the altitude of which varies coherently with the annual cycle of the stratospheric wave driven circulation (Yulaeva et al. 1994). The cold point is actually lower over some convective regions (Seidel et al. 2001), as illustrated in Fig. 11. The spatial distribution of cold point temperatures is likely a combination of several factors related to convection, including a direct cooling effect from mixing low- y air from deep convection, an indirect dynamic response to the re- lease of latent heat in convection at lower levels (Highwood and Hoskins 1998), and the radiative cooling associated with the presence of a thick cloud deck below the cold point (Gage et al. 1991; Norton 2001). The TTL has a lower boundary dominated by tropospheric processes (convection) and an upper boundary dominated by stratospheric radiative heating and a remote response to convection. This analysis does not imply that convection is not important above the level of the LRM. As shown by Sherwood and Dessler (2001), even a small fraction of convection can have a signi fi cant effect on water vapor and ozone in the tropopause region. A small fraction of the time (less than 0.5%), convection does penetrate through the tropopause layer (Gettelman et al. 2002), as indicated in Fig. 11. In addition, we do not consider here the effect of horizontal (quasi-isentropic) motions within the TTL or between the TTL and the extratropical lower stratosphere. These motions, which have a preferential locations in space and time (Waugh and Polvani 2000), are likely to be important for determining the distribution of trace species in the TTL. The LRM at 10 – 12 km (350 K) is well below the level at which the clear sky radiative heating changes from negative to positive (16 km or 370 K). This implies that the main convective detrainment does not reach a level from which it can rise into the stratosphere. Deeper convection, which reaches within 1 km of the CPT or higher, is necessary to loft air to altitudes where it can reach the stratosphere. There are coherent variations over the year in the altitude and depth of the tropical tropopause layer. Over the Western Paci fi c, the main convective out fl ow level is higher (and at higher potential temperature) than over the Eastern Paci fi c. Variations over the seasons follow the motion of convection. However, the CPT has a single and coherent annual cycle throughout the tropics (Seidel et al. 2001). Zonal variations of the level of the LRM are around 1 km, and the LRM is lower in non-convective regions, shown schematically in Fig. 11. The level of the LRM is at lower potential temperatures in the subtropics. These regional variations in depth of layer may be signi fi cant for short lived species. In some regions short lived species will be exported into the TTL at higher altitudes than in other regions. Recognizing that isentropic motion within the TTL subsequent to convection will rapidly transport parcels horizontally, and that below 16 km parcels are subsiding, this implies that entrainment of TTL air into convection is necessary to move parcels in the TTL upward to within 1 km of the tropopause. This may have implications for the transport of, for example, the products of biomass burning into the TTL, because the burning generally occurs in dry seasons, when convection is not present (though sometimes the fi res may them- selves cause convection). Analysis of chemical budgets in the TTL (of carbon monoxide for ex- ample), might shed some light on this hypothe- sis that ef fi cient transport into the TTL re- quires active convection. Variations of the tropopause layer on interannual time scales have also been diagnosed. There is a strong signal of the El-Nin ̃ o Southern Oscillation, which has a strong affect on the organization of convection and on the height of the LRM, particularly in the Paci fi c. Over the last 40 years, the cold point tropopause has risen 200 – 400 m, consistent with previous work (Seidel et al. 2001). The altitude of the LRM is also rising at a similar rate, though the variability is higher at this level, so trends are not as signi fi cant. These trends appear consistent with increasing surface temperatures over the last 40 years, assuming that tropospheric lapse rates have remained fi xed. We may hy- pothesize that the TTL may continue to rise if the warming expected from increased green- house gases is realized. Such a change is consistent with the sensitivity experiments of Thuburn and Craig (2002) which showed little sensitivity in a radiative-convective model to expected anthropogenic changes in carbon dioxide, because their surface temperatures remained fi xed. The analysis allows us to re fi ne the sche- matics of the tropical tropopause layer recently presented by Highwood and Hoskins (1998), Folkins et al. (1999), Sherwood and Dessler (2001) and V ̈ mel et al. (2002). Figure 11 indicates some of the important fi ndings of this paper. The TTL can be de fi ned as the region between the lapse rate minimum and the cold point. It has a varying thickness changing with convective activity, and is narrower with a higher base near active convective regions. The level of zero clear sky radiative heating is quite close to the tropopause. Two of the major issues not addressed in this study are the impact of horizontal transport, and the related issue of linkages between the tropical TTL and the extratropical tropopause, which is located at 10 – 12 km, nearly coincident with the base of the TTL. This horizontal motion is also indicated by vectors in Fig. 11. While the temperature structure may not be directly affected by such isentropic motion, the distribution of trace species like water vapor and ozone (and hence the radiative heating) may be affected. A future picture of the tropical tropopause layer would need to consider the role of the subtropics and the extratropics. We would like to thank A.E. Dessler and S.C. Sherwood for advice and assistance with ac- cessing data. Thanks to J. Witte for assistance with the SHADOZ archive. We also thank J. Thuburn, H. V ̈ mel and I. Folkins for provid- ing preprints of their work. This work also bene fi tted from numerous discussions with S. Massie, L. Pan and W.J. Randel, and from the care of two anonymous ...
Context 3
... TTL is a region of transition between tropospheric stability dominated by deep convection, and stratospheric stability caused by ozone heating. Convenient boundaries for this layer are the lapse rate minimum (LRM) at 10 – 12 km (340 – 350 K in isentropic coordi- nates), and the cold point tropopause (CPT) at 16 – 17 km (380 – 390 K). Both of these values are readily calculated from temperature soundings. The location of these levels is indicated schematically in Fig. 11. Important zonal variations in the depth of the TTL are also present, and these are noted schematically in Fig. 11. The TTL thus de fi ned is related to ozone and convection in the region. The altitude of the lapse rate minimum at 10 – 12 km corresponds to a level generally just below the minimum in ozone in the upper troposphere. The mean position of the LRM also corresponds to the main in fl uence of convection in the upper troposphere as observed from satellite cloud ob- servations and divergence of the wind fi eld. The variations of the LRM level are also signi fi cantly correlated with convective activity. In contrast, we fi nd that convective activity is not strongly correlated with the cold point, the altitude of which varies coherently with the annual cycle of the stratospheric wave driven circulation (Yulaeva et al. 1994). The cold point is actually lower over some convective regions (Seidel et al. 2001), as illustrated in Fig. 11. The spatial distribution of cold point temperatures is likely a combination of several factors related to convection, including a direct cooling effect from mixing low- y air from deep convection, an indirect dynamic response to the re- lease of latent heat in convection at lower levels (Highwood and Hoskins 1998), and the radiative cooling associated with the presence of a thick cloud deck below the cold point (Gage et al. 1991; Norton 2001). The TTL has a lower boundary dominated by tropospheric processes (convection) and an upper boundary dominated by stratospheric radiative heating and a remote response to convection. This analysis does not imply that convection is not important above the level of the LRM. As shown by Sherwood and Dessler (2001), even a small fraction of convection can have a signi fi cant effect on water vapor and ozone in the tropopause region. A small fraction of the time (less than 0.5%), convection does penetrate through the tropopause layer (Gettelman et al. 2002), as indicated in Fig. 11. In addition, we do not consider here the effect of horizontal (quasi-isentropic) motions within the TTL or between the TTL and the extratropical lower stratosphere. These motions, which have a preferential locations in space and time (Waugh and Polvani 2000), are likely to be important for determining the distribution of trace species in the TTL. The LRM at 10 – 12 km (350 K) is well below the level at which the clear sky radiative heating changes from negative to positive (16 km or 370 K). This implies that the main convective detrainment does not reach a level from which it can rise into the stratosphere. Deeper convection, which reaches within 1 km of the CPT or higher, is necessary to loft air to altitudes where it can reach the stratosphere. There are coherent variations over the year in the altitude and depth of the tropical tropopause layer. Over the Western Paci fi c, the main convective out fl ow level is higher (and at higher potential temperature) than over the Eastern Paci fi c. Variations over the seasons follow the motion of convection. However, the CPT has a single and coherent annual cycle throughout the tropics (Seidel et al. 2001). Zonal variations of the level of the LRM are around 1 km, and the LRM is lower in non-convective regions, shown schematically in Fig. 11. The level of the LRM is at lower potential temperatures in the subtropics. These regional variations in depth of layer may be signi fi cant for short lived species. In some regions short lived species will be exported into the TTL at higher altitudes than in other regions. Recognizing that isentropic motion within the TTL subsequent to convection will rapidly transport parcels horizontally, and that below 16 km parcels are subsiding, this implies that entrainment of TTL air into convection is necessary to move parcels in the TTL upward to within 1 km of the tropopause. This may have implications for the transport of, for example, the products of biomass burning into the TTL, because the burning generally occurs in dry seasons, when convection is not present (though sometimes the fi res may them- selves cause convection). Analysis of chemical budgets in the TTL (of carbon monoxide for ex- ample), might shed some light on this hypothe- sis that ef fi cient transport into the TTL re- quires active convection. Variations of the tropopause layer on interannual time scales have also been diagnosed. There is a strong signal of the El-Nin ̃ o Southern Oscillation, which has a strong affect on the organization of convection and on the height of the LRM, particularly in the Paci fi c. Over the last 40 years, the cold point tropopause has risen 200 – 400 m, consistent with previous work (Seidel et al. 2001). The altitude of the LRM is also rising at a similar rate, though the variability is higher at this level, so trends are not as signi fi cant. These trends appear consistent with increasing surface temperatures over the last 40 years, assuming that tropospheric lapse rates have remained fi xed. We may hy- pothesize that the TTL may continue to rise if the warming expected from increased green- house gases is realized. Such a change is consistent with the sensitivity experiments of Thuburn and Craig (2002) which showed little sensitivity in a radiative-convective model to expected anthropogenic changes in carbon dioxide, because their surface temperatures remained fi xed. The analysis allows us to re fi ne the sche- matics of the tropical tropopause layer recently presented by Highwood and Hoskins (1998), Folkins et al. (1999), Sherwood and Dessler (2001) and V ̈ mel et al. (2002). Figure 11 indicates some of the important fi ndings of this paper. The TTL can be de fi ned as the region between the lapse rate minimum and the cold point. It has a varying thickness changing with convective activity, and is narrower with a higher base near active convective regions. The level of zero clear sky radiative heating is quite close to the tropopause. Two of the major issues not addressed in this study are the impact of horizontal transport, and the related issue of linkages between the tropical TTL and the extratropical tropopause, which is located at 10 – 12 km, nearly coincident with the base of the TTL. This horizontal motion is also indicated by vectors in Fig. 11. While the temperature structure may not be directly affected by such isentropic motion, the distribution of trace species like water vapor and ozone (and hence the radiative heating) may be affected. A future picture of the tropical tropopause layer would need to consider the role of the subtropics and the extratropics. We would like to thank A.E. Dessler and S.C. Sherwood for advice and assistance with ac- cessing data. Thanks to J. Witte for assistance with the SHADOZ archive. We also thank J. Thuburn, H. V ̈ mel and I. Folkins for provid- ing preprints of their work. This work also bene fi tted from numerous discussions with S. Massie, L. Pan and W.J. Randel, and from the care of two anonymous ...
Context 4
... TTL is a region of transition between tropospheric stability dominated by deep convection, and stratospheric stability caused by ozone heating. Convenient boundaries for this layer are the lapse rate minimum (LRM) at 10 – 12 km (340 – 350 K in isentropic coordi- nates), and the cold point tropopause (CPT) at 16 – 17 km (380 – 390 K). Both of these values are readily calculated from temperature soundings. The location of these levels is indicated schematically in Fig. 11. Important zonal variations in the depth of the TTL are also present, and these are noted schematically in Fig. 11. The TTL thus de fi ned is related to ozone and convection in the region. The altitude of the lapse rate minimum at 10 – 12 km corresponds to a level generally just below the minimum in ozone in the upper troposphere. The mean position of the LRM also corresponds to the main in fl uence of convection in the upper troposphere as observed from satellite cloud ob- servations and divergence of the wind fi eld. The variations of the LRM level are also signi fi cantly correlated with convective activity. In contrast, we fi nd that convective activity is not strongly correlated with the cold point, the altitude of which varies coherently with the annual cycle of the stratospheric wave driven circulation (Yulaeva et al. 1994). The cold point is actually lower over some convective regions (Seidel et al. 2001), as illustrated in Fig. 11. The spatial distribution of cold point temperatures is likely a combination of several factors related to convection, including a direct cooling effect from mixing low- y air from deep convection, an indirect dynamic response to the re- lease of latent heat in convection at lower levels (Highwood and Hoskins 1998), and the radiative cooling associated with the presence of a thick cloud deck below the cold point (Gage et al. 1991; Norton 2001). The TTL has a lower boundary dominated by tropospheric processes (convection) and an upper boundary dominated by stratospheric radiative heating and a remote response to convection. This analysis does not imply that convection is not important above the level of the LRM. As shown by Sherwood and Dessler (2001), even a small fraction of convection can have a signi fi cant effect on water vapor and ozone in the tropopause region. A small fraction of the time (less than 0.5%), convection does penetrate through the tropopause layer (Gettelman et al. 2002), as indicated in Fig. 11. In addition, we do not consider here the effect of horizontal (quasi-isentropic) motions within the TTL or between the TTL and the extratropical lower stratosphere. These motions, which have a preferential locations in space and time (Waugh and Polvani 2000), are likely to be important for determining the distribution of trace species in the TTL. The LRM at 10 – 12 km (350 K) is well below the level at which the clear sky radiative heating changes from negative to positive (16 km or 370 K). This implies that the main convective detrainment does not reach a level from which it can rise into the stratosphere. Deeper convection, which reaches within 1 km of the CPT or higher, is necessary to loft air to altitudes where it can reach the stratosphere. There are coherent variations over the year in the altitude and depth of the tropical tropopause layer. Over the Western Paci fi c, the main convective out fl ow level is higher (and at higher potential temperature) than over the Eastern Paci fi c. Variations over the seasons follow the motion of convection. However, the CPT has a single and coherent annual cycle throughout the tropics (Seidel et al. 2001). Zonal variations of the level of the LRM are around 1 km, and the LRM is lower in non-convective regions, shown schematically in Fig. 11. The level of the LRM is at lower potential temperatures in the subtropics. These regional variations in depth of layer may be signi fi cant for short lived species. In some regions short lived species will be exported into the TTL at higher altitudes than in other regions. Recognizing that isentropic motion within the TTL subsequent to convection will rapidly transport parcels horizontally, and that below 16 km parcels are subsiding, this implies that entrainment of TTL air into convection is necessary to move parcels in the TTL upward to within 1 km of the tropopause. This may have implications for the transport of, for example, the products of biomass burning into the TTL, because the burning generally occurs in dry seasons, when convection is not present (though sometimes the fi res may them- selves cause convection). Analysis of chemical budgets in the TTL (of carbon monoxide for ex- ample), might shed some light on this hypothe- sis that ef fi cient transport into the TTL re- quires active convection. Variations of the tropopause layer on interannual time scales have also been diagnosed. There is a strong signal of the El-Nin ̃ o Southern Oscillation, which has a strong affect on the organization of convection and on the height of the LRM, particularly in the Paci fi c. Over the last 40 years, the cold point tropopause has risen 200 – 400 m, consistent with previous work (Seidel et al. 2001). The altitude of the LRM is also rising at a similar rate, though the variability is higher at this level, so trends are not as signi fi cant. These trends appear consistent with increasing surface temperatures over the last 40 years, assuming that tropospheric lapse rates have remained fi xed. We may hy- pothesize that the TTL may continue to rise if the warming expected from increased green- house gases is realized. Such a change is consistent with the sensitivity experiments of Thuburn and Craig (2002) which showed little sensitivity in a radiative-convective model to expected anthropogenic changes in carbon dioxide, because their surface temperatures remained fi xed. The analysis allows us to re fi ne the sche- matics of the tropical tropopause layer recently presented by Highwood and Hoskins (1998), Folkins et al. (1999), Sherwood and Dessler (2001) and V ̈ mel et al. (2002). Figure 11 indicates some of the important fi ndings of this paper. The TTL can be de fi ned as the region between the lapse rate minimum and the cold point. It has a varying thickness changing with convective activity, and is narrower with a higher base near active convective regions. The level of zero clear sky radiative heating is quite close to the tropopause. Two of the major issues not addressed in this study are the impact of horizontal transport, and the related issue of linkages between the tropical TTL and the extratropical tropopause, which is located at 10 – 12 km, nearly coincident with the base of the TTL. This horizontal motion is also indicated by vectors in Fig. 11. While the temperature structure may not be directly affected by such isentropic motion, the distribution of trace species like water vapor and ozone (and hence the radiative heating) may be affected. A future picture of the tropical tropopause layer would need to consider the role of the subtropics and the extratropics. We would like to thank A.E. Dessler and S.C. Sherwood for advice and assistance with ac- cessing data. Thanks to J. Witte for assistance with the SHADOZ archive. We also thank J. Thuburn, H. V ̈ mel and I. Folkins for provid- ing preprints of their work. This work also bene fi tted from numerous discussions with S. Massie, L. Pan and W.J. Randel, and from the care of two anonymous ...
Context 5
... TTL is a region of transition between tropospheric stability dominated by deep convection, and stratospheric stability caused by ozone heating. Convenient boundaries for this layer are the lapse rate minimum (LRM) at 10 – 12 km (340 – 350 K in isentropic coordi- nates), and the cold point tropopause (CPT) at 16 – 17 km (380 – 390 K). Both of these values are readily calculated from temperature soundings. The location of these levels is indicated schematically in Fig. 11. Important zonal variations in the depth of the TTL are also present, and these are noted schematically in Fig. 11. The TTL thus de fi ned is related to ozone and convection in the region. The altitude of the lapse rate minimum at 10 – 12 km corresponds to a level generally just below the minimum in ozone in the upper troposphere. The mean position of the LRM also corresponds to the main in fl uence of convection in the upper troposphere as observed from satellite cloud ob- servations and divergence of the wind fi eld. The variations of the LRM level are also signi fi cantly correlated with convective activity. In contrast, we fi nd that convective activity is not strongly correlated with the cold point, the altitude of which varies coherently with the annual cycle of the stratospheric wave driven circulation (Yulaeva et al. 1994). The cold point is actually lower over some convective regions (Seidel et al. 2001), as illustrated in Fig. 11. The spatial distribution of cold point temperatures is likely a combination of several factors related to convection, including a direct cooling effect from mixing low- y air from deep convection, an indirect dynamic response to the re- lease of latent heat in convection at lower levels (Highwood and Hoskins 1998), and the radiative cooling associated with the presence of a thick cloud deck below the cold point (Gage et al. 1991; Norton 2001). The TTL has a lower boundary dominated by tropospheric processes (convection) and an upper boundary dominated by stratospheric radiative heating and a remote response to convection. This analysis does not imply that convection is not important above the level of the LRM. As shown by Sherwood and Dessler (2001), even a small fraction of convection can have a signi fi cant effect on water vapor and ozone in the tropopause region. A small fraction of the time (less than 0.5%), convection does penetrate through the tropopause layer (Gettelman et al. 2002), as indicated in Fig. 11. In addition, we do not consider here the effect of horizontal (quasi-isentropic) motions within the TTL or between the TTL and the extratropical lower stratosphere. These motions, which have a preferential locations in space and time (Waugh and Polvani 2000), are likely to be important for determining the distribution of trace species in the TTL. The LRM at 10 – 12 km (350 K) is well below the level at which the clear sky radiative heating changes from negative to positive (16 km or 370 K). This implies that the main convective detrainment does not reach a level from which it can rise into the stratosphere. Deeper convection, which reaches within 1 km of the CPT or higher, is necessary to loft air to altitudes where it can reach the stratosphere. There are coherent variations over the year in the altitude and depth of the tropical tropopause layer. Over the Western Paci fi c, the main convective out fl ow level is higher (and at higher potential temperature) than over the Eastern Paci fi c. Variations over the seasons follow the motion of convection. However, the CPT has a single and coherent annual cycle throughout the tropics (Seidel et al. 2001). Zonal variations of the level of the LRM are around 1 km, and the LRM is lower in non-convective regions, shown schematically in Fig. 11. The level of the LRM is at lower potential temperatures in the subtropics. These regional variations in depth of layer may be signi fi cant for short lived species. In some regions short lived species will be exported into the TTL at higher altitudes than in other regions. Recognizing that isentropic motion within the TTL subsequent to convection will rapidly transport parcels horizontally, and that below 16 km parcels are subsiding, this implies that entrainment of TTL air into convection is necessary to move parcels in the TTL upward to within 1 km of the tropopause. This may have implications for the transport of, for example, the products of biomass burning into the TTL, because the burning generally occurs in dry seasons, when convection is not present (though sometimes the fi res may them- selves cause convection). Analysis of chemical budgets in the TTL (of carbon monoxide for ex- ample), might shed some light on this hypothe- sis that ef fi cient transport into the TTL re- quires active convection. Variations of the tropopause layer on interannual time scales have also been diagnosed. There is a strong signal of the El-Nin ̃ o Southern Oscillation, which has a strong affect on the organization of convection and on the height of the LRM, particularly in the Paci fi c. Over the last 40 years, the cold point tropopause has risen 200 – 400 m, consistent with previous work (Seidel et al. 2001). The altitude of the LRM is also rising at a similar rate, though the variability is higher at this level, so trends are not as signi fi cant. These trends appear consistent with increasing surface temperatures over the last 40 years, assuming that tropospheric lapse rates have remained fi xed. We may hy- pothesize that the TTL may continue to rise if the warming expected from increased green- house gases is realized. Such a change is consistent with the sensitivity experiments of Thuburn and Craig (2002) which showed little sensitivity in a radiative-convective model to expected anthropogenic changes in carbon dioxide, because their surface temperatures remained fi xed. The analysis allows us to re fi ne the sche- matics of the tropical tropopause layer recently presented by Highwood and Hoskins (1998), Folkins et al. (1999), Sherwood and Dessler (2001) and V ̈ mel et al. (2002). Figure 11 indicates some of the important fi ndings of this paper. The TTL can be de fi ned as the region between the lapse rate minimum and the cold point. It has a varying thickness changing with convective activity, and is narrower with a higher base near active convective regions. The level of zero clear sky radiative heating is quite close to the tropopause. Two of the major issues not addressed in this study are the impact of horizontal transport, and the related issue of linkages between the tropical TTL and the extratropical tropopause, which is located at 10 – 12 km, nearly coincident with the base of the TTL. This horizontal motion is also indicated by vectors in Fig. 11. While the temperature structure may not be directly affected by such isentropic motion, the distribution of trace species like water vapor and ozone (and hence the radiative heating) may be affected. A future picture of the tropical tropopause layer would need to consider the role of the subtropics and the extratropics. We would like to thank A.E. Dessler and S.C. Sherwood for advice and assistance with ac- cessing data. Thanks to J. Witte for assistance with the SHADOZ archive. We also thank J. Thuburn, H. V ̈ mel and I. Folkins for provid- ing preprints of their work. This work also bene fi tted from numerous discussions with S. Massie, L. Pan and W.J. Randel, and from the care of two anonymous ...
Context 6
... TTL is a region of transition between tropospheric stability dominated by deep convection, and stratospheric stability caused by ozone heating. Convenient boundaries for this layer are the lapse rate minimum (LRM) at 10 – 12 km (340 – 350 K in isentropic coordi- nates), and the cold point tropopause (CPT) at 16 – 17 km (380 – 390 K). Both of these values are readily calculated from temperature soundings. The location of these levels is indicated schematically in Fig. 11. Important zonal variations in the depth of the TTL are also present, and these are noted schematically in Fig. 11. The TTL thus de fi ned is related to ozone and convection in the region. The altitude of the lapse rate minimum at 10 – 12 km corresponds to a level generally just below the minimum in ozone in the upper troposphere. The mean position of the LRM also corresponds to the main in fl uence of convection in the upper troposphere as observed from satellite cloud ob- servations and divergence of the wind fi eld. The variations of the LRM level are also signi fi cantly correlated with convective activity. In contrast, we fi nd that convective activity is not strongly correlated with the cold point, the altitude of which varies coherently with the annual cycle of the stratospheric wave driven circulation (Yulaeva et al. 1994). The cold point is actually lower over some convective regions (Seidel et al. 2001), as illustrated in Fig. 11. The spatial distribution of cold point temperatures is likely a combination of several factors related to convection, including a direct cooling effect from mixing low- y air from deep convection, an indirect dynamic response to the re- lease of latent heat in convection at lower levels (Highwood and Hoskins 1998), and the radiative cooling associated with the presence of a thick cloud deck below the cold point (Gage et al. 1991; Norton 2001). The TTL has a lower boundary dominated by tropospheric processes (convection) and an upper boundary dominated by stratospheric radiative heating and a remote response to convection. This analysis does not imply that convection is not important above the level of the LRM. As shown by Sherwood and Dessler (2001), even a small fraction of convection can have a signi fi cant effect on water vapor and ozone in the tropopause region. A small fraction of the time (less than 0.5%), convection does penetrate through the tropopause layer (Gettelman et al. 2002), as indicated in Fig. 11. In addition, we do not consider here the effect of horizontal (quasi-isentropic) motions within the TTL or between the TTL and the extratropical lower stratosphere. These motions, which have a preferential locations in space and time (Waugh and Polvani 2000), are likely to be important for determining the distribution of trace species in the TTL. The LRM at 10 – 12 km (350 K) is well below the level at which the clear sky radiative heating changes from negative to positive (16 km or 370 K). This implies that the main convective detrainment does not reach a level from which it can rise into the stratosphere. Deeper convection, which reaches within 1 km of the CPT or higher, is necessary to loft air to altitudes where it can reach the stratosphere. There are coherent variations over the year in the altitude and depth of the tropical tropopause layer. Over the Western Paci fi c, the main convective out fl ow level is higher (and at higher potential temperature) than over the Eastern Paci fi c. Variations over the seasons follow the motion of convection. However, the CPT has a single and coherent annual cycle throughout the tropics (Seidel et al. 2001). Zonal variations of the level of the LRM are around 1 km, and the LRM is lower in non-convective regions, shown schematically in Fig. 11. The level of the LRM is at lower potential temperatures in the subtropics. These regional variations in depth of layer may be signi fi cant for short lived species. In some regions short lived species will be exported into the TTL at higher altitudes than in other regions. Recognizing that isentropic motion within the TTL subsequent to convection will rapidly transport parcels horizontally, and that below 16 km parcels are subsiding, this implies that entrainment of TTL air into convection is necessary to move parcels in the TTL upward to within 1 km of the tropopause. This may have implications for the transport of, for example, the products of biomass burning into the TTL, because the burning generally occurs in dry seasons, when convection is not present (though sometimes the fi res may them- selves cause convection). Analysis of chemical budgets in the TTL (of carbon monoxide for ex- ample), might shed some light on this hypothe- sis that ef fi cient transport into the TTL re- quires active convection. Variations of the tropopause layer on interannual time scales have also been diagnosed. There is a strong signal of the El-Nin ̃ o Southern Oscillation, which has a strong affect on the organization of convection and on the height of the LRM, particularly in the Paci fi c. Over the last 40 years, the cold point tropopause has risen 200 – 400 m, consistent with previous work (Seidel et al. 2001). The altitude of the LRM is also rising at a similar rate, though the variability is higher at this level, so trends are not as signi fi cant. These trends appear consistent with increasing surface temperatures over the last 40 years, assuming that tropospheric lapse rates have remained fi xed. We may hy- pothesize that the TTL may continue to rise if the warming expected from increased green- house gases is realized. Such a change is consistent with the sensitivity experiments of Thuburn and Craig (2002) which showed little sensitivity in a radiative-convective model to expected anthropogenic changes in carbon dioxide, because their surface temperatures remained fi xed. The analysis allows us to re fi ne the sche- matics of the tropical tropopause layer recently presented by Highwood and Hoskins (1998), Folkins et al. (1999), Sherwood and Dessler (2001) and V ̈ mel et al. (2002). Figure 11 indicates some of the important fi ndings of this paper. The TTL can be de fi ned as the region between the lapse rate minimum and the cold point. It has a varying thickness changing with convective activity, and is narrower with a higher base near active convective regions. The level of zero clear sky radiative heating is quite close to the tropopause. Two of the major issues not addressed in this study are the impact of horizontal transport, and the related issue of linkages between the tropical TTL and the extratropical tropopause, which is located at 10 – 12 km, nearly coincident with the base of the TTL. This horizontal motion is also indicated by vectors in Fig. 11. While the temperature structure may not be directly affected by such isentropic motion, the distribution of trace species like water vapor and ozone (and hence the radiative heating) may be affected. A future picture of the tropical tropopause layer would need to consider the role of the subtropics and the extratropics. We would like to thank A.E. Dessler and S.C. Sherwood for advice and assistance with ac- cessing data. Thanks to J. Witte for assistance with the SHADOZ archive. We also thank J. Thuburn, H. V ̈ mel and I. Folkins for provid- ing preprints of their work. This work also bene fi tted from numerous discussions with S. Massie, L. Pan and W.J. Randel, and from the care of two anonymous ...
Context 7
... TTL is a region of transition between tropospheric stability dominated by deep convection, and stratospheric stability caused by ozone heating. Convenient boundaries for this layer are the lapse rate minimum (LRM) at 10 – 12 km (340 – 350 K in isentropic coordi- nates), and the cold point tropopause (CPT) at 16 – 17 km (380 – 390 K). Both of these values are readily calculated from temperature soundings. The location of these levels is indicated schematically in Fig. 11. Important zonal variations in the depth of the TTL are also present, and these are noted schematically in Fig. 11. The TTL thus de fi ned is related to ozone and convection in the region. The altitude of the lapse rate minimum at 10 – 12 km corresponds to a level generally just below the minimum in ozone in the upper troposphere. The mean position of the LRM also corresponds to the main in fl uence of convection in the upper troposphere as observed from satellite cloud ob- servations and divergence of the wind fi eld. The variations of the LRM level are also signi fi cantly correlated with convective activity. In contrast, we fi nd that convective activity is not strongly correlated with the cold point, the altitude of which varies coherently with the annual cycle of the stratospheric wave driven circulation (Yulaeva et al. 1994). The cold point is actually lower over some convective regions (Seidel et al. 2001), as illustrated in Fig. 11. The spatial distribution of cold point temperatures is likely a combination of several factors related to convection, including a direct cooling effect from mixing low- y air from deep convection, an indirect dynamic response to the re- lease of latent heat in convection at lower levels (Highwood and Hoskins 1998), and the radiative cooling associated with the presence of a thick cloud deck below the cold point (Gage et al. 1991; Norton 2001). The TTL has a lower boundary dominated by tropospheric processes (convection) and an upper boundary dominated by stratospheric radiative heating and a remote response to convection. This analysis does not imply that convection is not important above the level of the LRM. As shown by Sherwood and Dessler (2001), even a small fraction of convection can have a signi fi cant effect on water vapor and ozone in the tropopause region. A small fraction of the time (less than 0.5%), convection does penetrate through the tropopause layer (Gettelman et al. 2002), as indicated in Fig. 11. In addition, we do not consider here the effect of horizontal (quasi-isentropic) motions within the TTL or between the TTL and the extratropical lower stratosphere. These motions, which have a preferential locations in space and time (Waugh and Polvani 2000), are likely to be important for determining the distribution of trace species in the TTL. The LRM at 10 – 12 km (350 K) is well below the level at which the clear sky radiative heating changes from negative to positive (16 km or 370 K). This implies that the main convective detrainment does not reach a level from which it can rise into the stratosphere. Deeper convection, which reaches within 1 km of the CPT or higher, is necessary to loft air to altitudes where it can reach the stratosphere. There are coherent variations over the year in the altitude and depth of the tropical tropopause layer. Over the Western Paci fi c, the main convective out fl ow level is higher (and at higher potential temperature) than over the Eastern Paci fi c. Variations over the seasons follow the motion of convection. However, the CPT has a single and coherent annual cycle throughout the tropics (Seidel et al. 2001). Zonal variations of the level of the LRM are around 1 km, and the LRM is lower in non-convective regions, shown schematically in Fig. 11. The level of the LRM is at lower potential temperatures in the subtropics. These regional variations in depth of layer may be signi fi cant for short lived species. In some regions short lived species will be exported into the TTL at higher altitudes than in other regions. Recognizing that isentropic motion within the TTL subsequent to convection will rapidly transport parcels horizontally, and that below 16 km parcels are subsiding, this implies that entrainment of TTL air into convection is necessary to move parcels in the TTL upward to within 1 km of the tropopause. This may have implications for the transport of, for example, the products of biomass burning into the TTL, because the burning generally occurs in dry seasons, when convection is not present (though sometimes the fi res may them- selves cause convection). Analysis of chemical budgets in the TTL (of carbon monoxide for ex- ample), might shed some light on this hypothe- sis that ef fi cient transport into the TTL re- quires active convection. Variations of the tropopause layer on interannual time scales have also been diagnosed. There is a strong signal of the El-Nin ̃ o Southern Oscillation, which has a strong affect on the organization of convection and on the height of the LRM, particularly in the Paci fi c. Over the last 40 years, the cold point tropopause has risen 200 – 400 m, consistent with previous work (Seidel et al. 2001). The altitude of the LRM is also rising at a similar rate, though the variability is higher at this level, so trends are not as signi fi cant. These trends appear consistent with increasing surface temperatures over the last 40 years, assuming that tropospheric lapse rates have remained fi xed. We may hy- pothesize that the TTL may continue to rise if the warming expected from increased green- house gases is realized. Such a change is consistent with the sensitivity experiments of Thuburn and Craig (2002) which showed little sensitivity in a radiative-convective model to expected anthropogenic changes in carbon dioxide, because their surface temperatures remained fi xed. The analysis allows us to re fi ne the sche- matics of the tropical tropopause layer recently presented by Highwood and Hoskins (1998), Folkins et al. (1999), Sherwood and Dessler (2001) and V ̈ mel et al. (2002). Figure 11 indicates some of the important fi ndings of this paper. The TTL can be de fi ned as the region between the lapse rate minimum and the cold point. It has a varying thickness changing with convective activity, and is narrower with a higher base near active convective regions. The level of zero clear sky radiative heating is quite close to the tropopause. Two of the major issues not addressed in this study are the impact of horizontal transport, and the related issue of linkages between the tropical TTL and the extratropical tropopause, which is located at 10 – 12 km, nearly coincident with the base of the TTL. This horizontal motion is also indicated by vectors in Fig. 11. While the temperature structure may not be directly affected by such isentropic motion, the distribution of trace species like water vapor and ozone (and hence the radiative heating) may be affected. A future picture of the tropical tropopause layer would need to consider the role of the subtropics and the extratropics. We would like to thank A.E. Dessler and S.C. Sherwood for advice and assistance with ac- cessing data. Thanks to J. Witte for assistance with the SHADOZ archive. We also thank J. Thuburn, H. V ̈ mel and I. Folkins for provid- ing preprints of their work. This work also bene fi tted from numerous discussions with S. Massie, L. Pan and W.J. Randel, and from the care of two anonymous ...
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Citations
... Though the main divergence occurs much lower for SC cases, secondary divergence will occur closer to the CPT , taken as the COT. Because COT-H marks the altitude where radiation begins to influence the temperature as it departs from the moist adiabat (Gettelman & Forster, 2002). Thus, even though SC lies much below the tropopause, COT forms at the altitude near the level of neutral buoyancy (LNB) above which the influence of radiation dominates. ...
Knowledge of the tracer characteristics, such as water vapor and ozone in the tropical tropopause layer (TTL), is vital in quantifying the radiation budget and critical to understanding the exchange processes between the troposphere and stratosphere. In this study, we have characterized the tropical tropopause parameters such as the cold point tropopause (CPT) height (CPT-H) and temperature (CPT-T), convective tropopause (COT) height (COT-H) and temperature (COT-T), and the tropical tropopause layer (TTL) using radiosonde observations during 2014–2019 over Chennai (13.0° N, 80.06° E) located in the northeast (NE) monsoon region. The water vapor and ozone data from the microwave limb sounder (MLS) simultaneous to the radiosonde observations are also utilized to understand their roles on the CPT variations for different convective conditions obtained from Infrared brightness temperature (IRBT) data. CPT over Chennai becomes higher (17.6 ± 0.3 km) and colder (189.7 ± 0.9 K) during the winter season and lower (16.6 ± 0.2 km) and warmer (192.1 ± 1.0 K) during the summer monsoon season, however, not in the same month. The water vapor (CPT-W) and ozone (CPT-O) mixing ratios at CPT are found to be lower (~ 70 ± 1.4 ppmv and 3.1 ± 0.4 ppmv) during the winter season and higher (153 ± 4.2 ppbv and 4.8 ± 0.6 ppmv) during summer monsoon season. COT, however, becomes lower (12.4 ± 0.3 km) and higher (13.3 ± 0.3 km) during pre-monsoon and summer monsoon seasons, respectively. The TTL thickness is lesser (3.5 ± 0.6 km) during the winter and greater (4.8 ± 0.8 km) during the summer monsoon seasons. Over Chennai, the seasonal variation of the upper troposphere and lower stratospheric temperature, water vapor, and ozone anomalies are in phase. We have categorized tropical convections as non-penetrative and penetrative using IRBT data. It is observed that the TTL temperature warms with the increasing strength of the non-penetrative convections and cools for the penetrative convection.
... The CV exhibited higher values in tropical and subtropical regions compared to other regions. Numerous research studies (e.g., Gettelman, 2002;RavindraBabu et al., 2019;Xie et al., 2012) have confirmed the influence of various factors, such as monsoons and El Niño, on the tropopause characteristics within the tropical and subtropical regions. The presence of these influencing factors and their periodic variations could potentially contribute to the occurrence of extreme CV values in these particular geographical areas. ...
The tropopause serves a critical role in shaping global and regional weather and climate dynamics. Changes in tropopause characteristics can significantly impact other atmospheric components, thereby influencing Earth’s climate systems. In the long run, variations in tropopause features can lead to shifts in the thermal, dynamic, and chemical properties of the tropospheric layer. This study aims to investigate the descriptive attributes of tropopause pressure levels (TPLs) during different months, as well as the temporal and spatial trends in TPL across the Northern Hemisphere spanning from 1979 to 2022. Utilizing ERA5 temperature data for the 700 to 50 hPa range, the tropopause was identified using the lapse rate of tropopause (LRT), and its changes were analyzed employing the linear regression model with the least squares error approach. The results indicated that the spatial pattern of TPLs changed across various latitudes varies seasonally. Generally, the changes in TPLs did not exhibit a linear relationship with latitude, and in most observed months, the highest and lowest TPLs did not correspond to the lowest and highest latitudes, respectively. Examination of the trend in TPLs revealed that in numerous significant areas across different seasons, the trends were statistically insignificant. Where significant, the trends predominantly indicated negative changes (decreases), suggesting a reduction in pressure and potentially an increase in tropopause altitude in these regions, possibly reflecting the influence of global warming.
... The lower limit of the TTL is the Level of Main Convective Outflow (LMCO) which is also the Level of Neutral Buoyancy (LNB) at around 14-15 km. Additionally, the level of minimum stability (LMS) [72] or the lapse rate minimum (LRM) [73] typically around 10-12 km, is also important and associated with cloud base height in certain months such as April and May, see Section 3.1 and Figure 6. . Note that the y axis of (b) for LNB is the potential temperature (K) and the double y axis is for CPT in different units: K and • C. The monthly mean and with 1-σ standard error of the data are shown by a red line and error bar, respectively. ...
... The meteorology profiles are from radio soundings by the National Weather Service of Palau, see Section 2.2.2. The level of the Cold Point Tropopause (CPT), which is the level of temperature minimum [73,74] or the Lapse Rate Tropopause (LRT; World Meteorological Organization, 1957), is used to identify the upper TTL, which is about 17 km. Fueglistaler et al. [8] proposed a Level of Highest Convective Overshooting (LHCO) tops that can reach about 18.5 km (70 hPa) as the upper limit of the TTL. ...
This study presented an analysis of the geometric and optical properties of cirrus clouds with data produced by Compact Cloud-Aerosol Lidar (ComCAL) over Koror, Palau (7.3°N, 134.5°E), in the Tropical Western Pacific region. The lidar measurement dataset covers April 2018 to May 2019 and includes data collected during March, July and August 2022. The results show that cirrus clouds occur approximately 47.9% of the lidar sampling time, predominantly between altitudes of 15 and 18 km. Seasonal variations in cirrus top height closely align with those of the cold point tropopause. Most cirrus clouds exhibit low cloud optical depth (COD < 0.1), with an annual mean depolarization ratio of 31 ± 19%. Convective-forming cirrus clouds during the summer monsoon season exhibit a larger size by notably lower values in terms of color ratio. Extremely thin cirrus clouds (COD < 0.005) constituting 1.6% of total cirrus occurrences are frequently observed at 1–2 km above the cold point, particularly during winter and summer, suggesting significant stratosphere–troposphere exchange. The coldest and highest tropopause over Palau is persistent during winter, and related to the pathway of tropospheric air entering the stratosphere through the cold trap. In summer, the extremely thin cirrus above the cold point is likely correlated with equatorial Kelvin waves induced by western Pacific monsoon convection.
... Most of the air reaching Palau has crossed the Pacific and has thus been cleaned of O 3 precursors from anthropogenic or other continental pollution. Palau has a hot, humid and wet climate all year, resulting in a continuously high convective activity (Gettelman and Forster, 2002;Müller et al., 2024). This persistent high convective activity creates a well-mixed profile of uniformly low O 3 VMR throughout the free troposphere, as illustrated in Fig. 2a. ...
... Recognized as a characteristic feature of tropical profiles, low UT O 3 is often used as an indicator for deep convective detrainment (e.g., Kley et al., 1996;Folkins, 2002;Solomon et al., 2005;Gettelman et al., 2009;Paulik and Birner, 2012). Satellite observations confirm year-round convective activity in Palau, with some variability due to the traverse of the ITCZ twice a year (Gettelman and Forster, 2002;Müller et al., 2024). Müller et al. (2024) associate the occurrence of the strongest winds below 15 km when the ITCZ is furthest south or north of Palau with the poleward branch of the Hadley circulation. ...
The transport history of tropospheric air masses above the tropical western Pacific (TWP) is reflected by the local ozone and relative humidity (RH) characteristics. In boreal winter, the TWP is the main global entry point for air masses into the stratosphere and therefore a key region of atmospheric chemistry and dynamics. Our study aims to identify air masses with different pathways to the TWP using ozone and radio soundings from Palau from 2016–2019. Supported by backward trajectory calculations, we found five different types of air masses. We further defined locally controlled ozone and RH background profiles based on monthly statistics and analyzed corresponding anomalies in the 5–10 km altitude range. Our results show a bimodality in RH anomalies. Humid and ozone-poor background air masses are of local or Pacific convective origin and occur year-round, but they dominate from August until October. Anomalously dry and ozone-rich air masses indicate a non-local origin in tropical Asia and are transported to the TWP via an anticyclonic route, mostly from February to April. The geographic location of origin suggests anthropogenic pollution or biomass burning as a cause for ozone production. We propose large-scale descent within the tropical troposphere and radiative cooling in connection with the Hadley circulation as being responsible for the dehydration during transport. The trajectory analysis revealed no indication of a stratospheric influence. Our study thus presents a valuable contribution to the discussion about anomalous layers of dry ozone-rich air observed in ozone-poor background profiles in the TWP.
... In this case, TEJ peak (H TEJ ) lies $2 km below the H CPT coinciding with UT temperature inversion (Fujiwara et al., 2003) which is in thermal wind balance. The frequent convection (vertical motion) with an average outflow level $12-14 km (Gettelman & de F. Forster, 2002;Mehta et al., 2008) occurs during the ISM that would cause the tropospheric layer away from the hydrostatic equilibrium, however, not the layer above the top of the maximum convective heating which will remain in the thermal wind balance. ...
This paper presents the first quantitative relationship between the cold point tropopause (CPT) and tropical easterly jet (TEJ) using radiosonde observations over Gadanki (13.45 N, 79.2 E) during the Indian summer monsoon season 2006-2014. CPT and TEJ peak altitudes (H CPT and H TEJ) show amalgams of two categories of variability on the day-today scale. In category1 H TEJ occurs close to H CPT and they show in-phase variation. While in Category2 H TEJ occurs far apart from H CPT and they do not show any relationship. For Category1 H CPT and H TEJ are strongly correlated (0.70), as well as H CPT and T CPT (CPT temperature) are moderately anticorrelated (À0.55) significant at a 95% confidence level, indicating the dominance of adiabatic processes. Whereas in Category2 H CPT and T CPT are not significantly anti-correlated. Thus, when TEJ and CPT are close to each other, it may serve as an indicator for the prevalence of the synoptic-scale effect.
... It is beneficial to understand the formation mechanisms of tropopause, formatted by the combination of tropospheric and stratospheric processes, and to further research STE processes by defining the tropopause from various perspectives (Gettelman and Forster, 2002;Bian, 2009;Chen et al., 2006). In early atmospheric models, the tropopause is characterized as a 45 discontinuous interface featuring a sharp vertical gradient. ...
... It is more reasonable to consider the tropical tropopause as a transition layer than a discontinuous interface (Highwood and Hoskins, 1998). Gettelman and Forster (2002) comprehensively considered both radiation and convection, and separated cold point tropopause (CPT) and potential temperature lapse minimum rate tropopause (LRM) as the upper and lower 55 boundaries of tropical tropopause layer, respectively. A primary characteristic of the tropopause is the drastic alteration of the atmospheric static stability when crossing this transitional layer. ...
... The atmospheric stability within the TTL is affected by convection in the troposphere and radiation in the stratosphere (Gettelman and Forster, 2002;Thuburn and Craig, 2002). Following further analysis and researches, two commonly used thermal tropopause definitions were emerged, namely CPT (in Fig. 2(a)) and LRM (in Fig. 2(c)). ...
The tropopause is important as a diagnostic of the upper troposphere and lower stratosphere structures, with unique atmospheric thermal, dynamic structures. A comprehensive understanding of the evolution of the fine tropopause structures is necessary and primary to further study the complex multi-scale atmospheric physicochemical coupling processes in the upper troposphere and lower stratosphere. Utilizing the bi-Gaussian function, a novel method is capable of identifying the characteristic parameters of tropopause vertical structures, as well as providing the information of double tropopauses (DT) structures. The new method improves the definition of cold point tropopause, and detects one (or two) most significant local coldest point(s) in mathematical statistics by fitting the temperature profiles to the bi-Gaussian function, which is (are) defined as the tropopause height(s). The bi-Gaussian function exhibits remarkable potential for explicating the variation trend of temperature profiles. The recognition results of the bi-Gaussian method and lapse rate tropopause, as defined by World Meteorological Organization, are compared in detail for different cases. Results indicate that the bi-Gaussian method possesses a lower missed detection rate and false detection rate than lapse rate tropopause, because it is not restricted by thresholds, even in the presence of multiple temperature inversion layers at higher elevations. Five-year (from 2012 to 2016) historical radiosondes in China revealed that the occurrence frequency and thickness of DT, as well as the single tropopause height, and the first and second DT height displayed significant meridional monotonic variations. The occurrence frequency (thickness) of DT increased from 2.93 % (2.61 km) to 72.45 % (6.84 km) in the latitude range [16° N, 50° N]. At mid-latitudes [30° N, 40° N], the meridional gradients of tropopause height were relatively large, essentially corresponding to the climatological location of the subtropical jet and Tibetan Plateau. The average DT thickness reported in this study is approximately 1–2 km thicker than that in previous studies, particularly in the mid-high latitudes [45° N, 50° N], which may be related to the different vertical resolution of temperature profiles provided by various data sources. DT structure occurs most frequently and has the largest meridional gradient in the mid-latitudes, formatted by a combination of poleward advection in the low-latitude upper troposphere and equatorward advection in the high-latitude lower stratosphere. In addition, although DT is thick in winter, the DT temperature difference is small, even the case of the first tropopause temperature is lower than the second tropopause temperature happens occasionally.
... The stratospheric water vapour (SWV) mainly originates from the troposphere: the moist air parcels at the bottom of the troposphere ascend, reaching the tropical tropopause layer (TTL) between 14-18.5 km, experiencing a severe dehydration process at the TTL (because the TTL has the coldest temperature in the lower atmosphere) (Gettelman and Forster 2002;Fueglistaler et al. 2009), then arriving in the stratosphere. The SWV is suggested to contribute significantly to global climate change by altering the infrared opacity of the atmosphere (e.g., Soden and Held 2006), providing a strong positive feedback at +0.3 W/(m 2 ·K) to global warming (Dessler et al. 2013). ...
... With O 3 as an indicator for its precursor NO x (Crawford et al., 1997;Gao et al., 2014;Rex et al., 2014), we would expect a decrease from east to west, while crossing the remote Pacific and following the trade winds within the ITCZ (Thompson et al., 2012), as more and more NO x is lost by conversion to HNO 3 and subsequent washout in convection is outweighing NO x production (Graedel et al., 1994). However, this lateral ozone gradient is not apparent in any of our statistical measures. ...
The tropical western Pacific is recognized as an important region for stratosphere–troposphere exchange but lies in a data-sparse location that had a measurement gap in the global ozone sounding network. The Palau Atmospheric Observatory (PAO, approx. 7.3∘ N, 134.5∘ E) was established to study the atmospheric composition above the remote tropical western Pacific with a comprehensive instrumental setup. Since 2016, two laboratory containers in Palau host a Fourier transform infrared spectrometer; a lidar (micro-lidar until 2016, cloud and aerosol lidar from 2018); a Pandora 2S photometer; and laboratory space for weather balloon soundings with ozone, water vapor, aerosol, and radiosondes. In this analysis, we focus on the continuous, fortnightly ozone sounding program with electrochemical concentration cell (ECC) ozonesondes. The aim of this study is to introduce the PAO and its research potential, present the first observation of the typical seasonal cycle of tropospheric ozone in the tropical western Pacific based on a multiannual record of in situ observations, and investigate major drivers of variability and seasonal variation from January 2016 until December 2021 related to the large-scale atmospheric circulation. We present the PAO ozone (O3) volume mixing ratios (VMR) and relative humidity (RH) time series complemented by other observations. The site is exposed to year-round high convective activity reflected in dominating low O3 VMR and high RH. In 2016, the impact of the strong El Niño is evident as a particularly dry, ozone-rich episode. The main modulator of annual tropospheric O3 variability is identified as the movement of the Intertropical Convergence Zone (ITCZ), with the lowest O3 VMR in the free troposphere during the ITCZ position north of Palau. An analysis of the relation of O3 and RH for the PAO and selected sites from the Southern Hemispheric Additional Ozonesondes (SHADOZ) network reveals three different regimes. Palau's O3 / RH distribution resembles the one in Fiji, Java and American Samoa but is unique in its seasonality and its comparably narrow Gaussian distribution around low O3 VMR and the evenly distributed RH. A previously found bimodal distribution of O3 VMR and RH could not be seen for the full Palau record but only during specific seasons and years. Due to its unique remote location, Palau is an ideal atmospheric background site to detect changes in air dynamics imprinted on the chemical composition of the tropospheric column. The efforts to establish, run and maintain the PAO have succeeded to fill an observational gap in the remote tropical western Pacific and give good prospects for ongoing operations. The ECC sonde record will be integrated into the SHADOZ database in the near future.
... Any reference to the tropopause in the text hereafter consequently refers to the average cold-point tropopause. Although this vertical resolution will not allow for a replication of radio-sounding measurements (Gettelman & Forster, 2002), we chose it as Dauhut and Hohenegger (2022) showed that it is possible to compute a moisture budget at this grid spacing. Figure 1 shows the inner tropical (potential) temperature and lapse rate profiles of the Control (CTL, blue) and Perturbed (PTB, red) simulations. ...
Plain Language Summary
The stratosphere is a dry region since moisture entering it from below has to pass the cold‐point, a temperature minimum between troposphere and stratosphere. The low temperatures lead to ice formation and sedimentation of moisture. Frozen moisture within clouds rising above the cold‐point tropopause can pass this temperature barrier and be injected into the stratosphere, where temperatures increase again, promoting the melting and sublimation of ice crystals. However, little is known about the sensitivity of the split of moisture entering the stratosphere into frozen and non‐frozen moisture, especially under external influences, like heating by volcanic aerosol or stratospheric geoengineering efforts. Convective parameterizations in conventional simulations can lead to biases. The emerging km‐scale simulations, which explicitly resolve the physical processes, offer the unique possibility to study moisture fluxes under external forcing while circumventing the downsides of parameterizations. Here, the sensitivity of the moisture flux partitioning into non‐frozen and frozen components to an additional heating source is studied for the first time in global storm‐resolving simulations. The analysis reveals an unaltered flux partitioning even at an average cold‐point warming exceeding 8 K. In the control and perturbed simulations, water vapor contributes around 80% of the moisture entering the stratosphere.
... The tropical tropopause plays a crucial role in this regard. Recent studies have shown that the tropical tropopause is not just a material surface but more of a transition layer named the 'tropical tropopause layer (TTL)' which extends from the top of the convectively well-mixed troposphere to the relatively stable stratified stratosphere (Gettelman and Forster, 2002). The TTL is the gateway for the entry of minor/trace constituents into the lower stratosphere and plays a key role in the Stratosphere-Troposphere Exchange (STE) (Holton et al., 1995;Fueglistaler et al., 2009 and reference therein). ...
