Background Stratospheric Aerosol Variations Deduced from Satellite Observations

Transcription

1 APRIL 2012 L I U E T A L. 799 Background Stratospheric Aerosol Variations Deduced from Satellite Observations YU LIU Chinese Academy of Meteorological Sciences, Beijing, China XUEPENG ZHAO National Climatic Data Center, NOAA/NESDIS, Asheville, North Carolina WEILIANG LI AND XIUJI ZHOU Chinese Academy of Meteorological Sciences, Beijing, China (Manuscript received 7 December 2010, in final form 13 November 2011) ABSTRACT The Stratospheric Aerosol and Gas Experiment II (SAGE II) aerosol products from 1998 to 2004 have been analyzed for the tendency of changes in background stratospheric aerosol properties. The aerosol extinction coefficient E has apparently increased in the midlatitude lower stratosphere (LS) in both hemispheres, at an annual rate that is as great as 2% 5%. Positive changes in the aerosol surface area density S in the midlatitude LS are most distinct, with a rate of increase that is as high as 5% 6% annually. At the same time, there has been a secular decrease in aerosol effective radius R, especially in the tropical LS, at a rate of up to 22.5% yr 21. Corresponding to these trends, the aerosol number concentration is inferred to have increased by roughly 5% 10% yr 21 in the tropical LS and by 4% 8% yr 21 in the midlatitude LS. Changes in aerosol mass are also deduced, with rates of increase in the midlatitude LS that are in the range of 1% 5% yr 21. The large uncertainty in operational S product is the major factor influencing the trend in S, aerosol number concentrations, and mass. The authors global assessment supports the speculation of Hofmann et al. on the basis of local observations that the cause of an increase in lidar backscatter over a similar period was a consequence of aerosol particle growth due to enhanced anthropogenic sulfur dioxide emissions. Moreover, it is found that an increase in the injection rate of condensation nuclei from the troposphere to the stratosphere at tropical latitudes is required to sustain the increase in stratospheric aerosol concentrations identified in this analysis. 1. Introduction The stratospheric aerosol layer consists of submicrometersized particles composed primarily of liquid solutions of sulfuric acid and water, with traces of other materials, such as ammonium sulfates. The layer was first identified by Junge et al. (1961); hence, it is often referred to as the Junge layer. Many studies have been performed to explain the formation, growth, and removal of stratospheric particles (e.g., Castleman et al. 1974; Hofmann et al. 1976; Turco et al. 1979; Sedlacek et al. 1983; Hofmann 1990) owing to the recognition of their radiative Corresponding author address: Dr. Yu Liu, Chinese Academy of Meteorological Sciences, 46 South Zhongguangchun Ave., Beijing, , China. liuyu@cams.cma.gov.cn and climatic effects (Pollack et al. 1976; Turco et al. 1980) and their influence on stratospheric ozone through heterogeneous chemistry (Hofmann and Solomon 1989; Turco et al. 1989; Turco and Hamill 1992; Zhao et al. 1997). Stratospheric sulfate aerosols are believed to originate from the transport of long-lived sulfur gases from the troposphere to the stratosphere, where they are oxidized to form sulfuric acid solutions (H 2 SO 4 -H 2 O) on preexisting condensation nuclei (CN) (Rosen 1971; Turco et al. 1982). Carbonyl sulfide (COS) is an important source of background stratospheric sulfate (Crutzen 1976; Turco et al. 1980; Chin and Davis 1995). COS is the most abundant gaseous sulfur species in the ambient atmosphere. Being relatively stable in the troposphere, it can be transported to the stratosphere in sufficient amounts to dominate the sulfate layer in certain regions. Sulfur dioxide (SO 2 ) is also occasionally injected directly into the DOI: /JAMC-D Ó 2012 American Meteorological Society

2 800 J O U R N A L O F A P P L I E D M E T E O R O L O G Y A N D C L I M A T O L O G Y VOLUME 51 stratosphere during strong convective events (Dickerson et al. 1987; Crawford et al. 2003). Note that a dense layer of stratospheric particles typically forms through homogeneous nucleation of H 2 SO 4 -H 2 O following major volcanic injections of SO 2 (Hamill et al. 1982; McKeen et al. 1984). During volcanically quiescent periods, the degree of supersaturation of H 2 SO 4 vapor in the stratosphere is usually insufficient for particle formation by homogeneous nucleation (Hamill et al. 1982; Golombek and Prinn 1993). In the upper tropical troposphere, however, binary homogeneous nucleation of H 2 SO 4 -H 2 O can provide nuclei upon which oxidized sulfur gases can later condense in the stratosphere (Brock et al. 1995). There have apparently been no major volcanic eruptions of the magnitude of El Chichón or Mount Pinatubo since Accordingly, the stratosphere has attained a relatively persistent volcanically quiescent period, in which variations and trends in the background stratospheric aerosol can be effectively investigated (Thomason et al. 2008). Deshler et al. (2006) studied trends in the nonvolcanic component of the stratospheric aerosol for three volcanically quiescent periods ( , , and ) by using in situ measurements, lidar observations, and satellite remote sensing records. They did not find any evident long-term changes in the background sulfate layer either over the entire period or during any of the individual volcanically quiescent periods. Liu et al. (2007) analyzed the Stratospheric Aerosol and Gas Experiment (SAGE) II aerosol products over China since 1997 and noted a 4% 5% per year (% yr 21 ) increase in stratospheric aerosol surface area density over the Tibetan Plateau and in the same latitudinal regions of eastern China. The maximum increase occurred at about 23-km altitude. Hofmann et al. (2009) more recently analyzed the backscatter ratio of stratospheric aerosols using lidar observations at Mauna Loa, Hawaii, and Boulder, Colorado and found a 4% 7% yr 21 increase in the aerosol backscatter between 20- and 30-km altitudes after 2000 in both cases. They suggested that these local trends are associated with an increase in anthropogenic SO 2 emissions worldwide, which may have other implications, including secular changes in stratospheric chemistry and radiative forcing. These somewhat inconsistent results among the trend studies cited above suggest that further analysis is needed to clarify the differences. In this paper, we analyze possible trends in the stratospheric aerosol layer on a global scale using SAGE II satellite observations collected during the volcanically quiescent period from 1998 to 2005, and we explore the cause of such changes. As indicated in Deshler et al. (2006) and Thomason et al. (2008), this period is perhaps the first epoch of aerosol measurements (since regular observations began in the early 1970s) that is truly nonvolcanically influenced. There are actually still some smaller volcanic eruptions in the tropics during this so-called volcanically quiescent period that may have somewhat influenced the tropical stratosphere (Vernier et al. 2009), including the Reventador volcanic eruption at 08 latitude in September of 2002 and the Manam volcanic eruption at 48S in January of Our expectation that the effect of the small Reventador eruption on the following trend analysis is limited since its stratospheric impact is concentrated in the first 6 months and below the minimum velocity level of approximately km in the tropical stratosphere will be examined later by checking the time series of SAGE II zonal mean extinction ratio (see Fig. 8 of Vernier et al. 2009). The effect of the Manam eruption may be bigger on the trend analysis, however, since it is at the end of the SAGE II data time series and therefore the well-known endpoint problem associated with trend analysis (see Weatherhead et al. 1998) may enhance the impact. Thus, we will confine our trend study to (without doing any further data filtering) to avoid the potentially amplified effect of the Manam eruption at the endpoint of the time series. Global SAGE II data also provide additional useful aerosol information, such as the effective particle size, allowing further interpretation of observed trends on a global scale. 2. SAGE II aerosol products Version 6.2 of the SAGE II satellite dataset (Mauldin et al. 1985; McCormick 1987) is used in this study. The data (obtained online at include profiles of aerosol extinction coefficient E at 1.02, 0.525, 0.435, and mm, number density profiles for ozone (O 3 ) and nitrogen dioxide (NO 2 ), and molecular density and water vapor mixing ratio profiles. All profiles have 0.5-km vertical resolution, cover the time period from October 1984 to August 2005, and are nearly global in coverage, spanning 808S 808N. The dataset also includes profiles of aerosol surface area density S and particle effective radius R derived from the solar transmission measurements (Thomason et al. 1997, 2008; Thomason and Peter 2006). These latter parameters are central to the current study, as described below. SAGE II aerosol observations have been widely used for providing/studying the climatological behavior of stratospheric aerosols (e.g., Hitchman et al. 1994; Thomason et al. 1997). The aerosol surface area density S and volume density V are defined, respectively, by S 5 ð 0 4pr 2dn(r) dr dr and V 5 ð pr3dn(r) dr, (1) dr

3 APRIL 2012 L I U E T A L. 801 FIG. 1. Height latitude distribution of the zonal-mean difference between annual and seasonal means of stratospheric aerosol extinction coefficient (km 21 ) at 1.02 mm, averaged over the 7-yr period from 1998 to where dn(r)/dr is the number of particles per unit volume per unit radius r in the interval between r and r 1 dr.the nth size distribution moment M n is defined as M n 5 ð 0 dn(r) dr rn dr. (2) A parameter that is frequently used to describe particle size is the effective, or area-weighted, radius R, defined as the ratio of the third moment to the second moment of the size distribution: R 5 M 3 /M 2 5 3V/S. (3) The sensitivity of operational SAGE II aerosol measurements at the lowest levels of stratospheric aerosols during a volcanically quiescent period may be a concern in our analysis. This sensitivity issue has been carefully examined by Thomason et al. (2008), who found that SAGE II aerosol extinction measurements, especially at 1.02 mm, remain robust and reliable at the observed lower aerosol levels for nonvolcanic conditions. The operational aerosol surface area density product may have an uncertainty on the order of at least a factor of 2, however, because of the lack of sensitivity to particles with radii of less than 0.1 mm.theimpactofthisuncertainty on aerosol surface area density S is especially evident in the regions where strong particle nucleation occurs, such as the lower tropical stratosphere and the winter polar stratosphere. Thus, we are cautious with regard to the results of the S trend analysis (especially for the tropical lower stratosphere), as is addressed further in the discussion section. 3. Seasonal variations and time series Latitude height distributions of the zonal-mean difference between annual and seasonal means of extinction coefficient E at 1.02 mm, the aerosol surface area density S, and the effective radius R all averaged over the 7-yr period from 1998 to 2004 are plotted for the seasonal intervals December February (DJF), March May (MAM), June August (JJA), and September November (SON) in Figs. 1, 2, and 3, respectively. Distinct seasonal variations are observed in these distributions. Values in the lower stratosphere (LS) are typically higher

4 802 J O U R N A L O F A P P L I E D M E T E O R O L O G Y A N D C L I M A T O L O G Y VOLUME 51 FIG. 2. As in Fig. 1, but for surface area density (mm 2 cm 23 ). in winter/spring than in summer/autumn. This seasonal variation is consistent with the stratospheric aerosol lidar backscatter measurements collected by Hofmann et al. (2009), which show a distinct peak for each parameter in winter and reach a minimum in summer. This result suggests that the SAGE II aerosol products still effectively capture seasonal variations in the background stratospheric aerosols even though there may be some residual invisible cirrus cloud effects contained in the extinction coefficient E in the tropical LS as shown in Fig. 1. Figure 4 shows the time series of more than 20 yr of extinction measurements E (1.02 mm), at 20-km altitude for the latitude belt of S. The impact of the 1991 Mount Pinatubo eruption is significant but diminishes almost completely within several years. The impact from relatively small volcanic eruptions, such as Reventador in September of 2002 and Manam in January of 2005, is not discernible in Fig. 4 because of their much smaller magnitude relative to the Mount Pinatubo eruption and the fact that their impact is mainly confined in the tropical LS. To minimize the effects of the Pinatubo perturbation (due to its big magnitude) and Manam (due to its endpoint effect) in masking secular changes in the background stratospheric aerosol, we restrict our trend analysis to the period from January of 1998 to December of 2004 for the following studies. Figure 5 then illustrates the time series of E (at 0.386, 0.452, and 1.02 mm), S, and R at 20-km altitude in the latitude belt of S for this truncated time interval. It can be seen that in each case there is a distinct (although noisy) trend in each parameter. Thus, in terms of a linear regression, extinction at all of the wavelengths shown, as well as the total particle surface area, increases after 1998 while the effective radius decreases. Because measurements of E at the longer wavelengths in the ensemble (0.525 and 1.02 mm) are more accurate (Thomason et al. 1997, 2008; Thomason and Peter 2006), our analysis focuses on these wavelengths. The time series of E (at and 1.02 mm), S, and R at 20-km altitude are also displayed for 58S 58N and N in Figs. 6 and 7, respectively. The trends in these parameters in the northern midlatitude LS (Fig. 7) are similar to the trends in the southern midlatitude LS (Fig. 5), but magnitudes are somewhat reduced. The trends in the tropical LS, shown in Fig. 6, are significantly different,

5 APRIL 2012 LIU ET AL. 803 FIG. 3. As in Fig. 1, but for effective radius (mm). however. For example, the effect of the minor volcanic eruption of Reventador is evident. Note that E at 1.02 mm tends to decrease rather than increase and that there is no evident trend in S or E at mm. The effect of the minor volcanic eruption on R is not evident, probably because R depends only on the ratio of V and S as given in Eq. (3). The irregular changes of these parameters in the tropical LS are mainly due to the effect of relatively small volcanic eruptions in the tropical LS, such as Reventador in September of 2002, and the relatively larger uncertainty in aerosol surface area density at the lowest levels of stratospheric aerosols during volcanically quiescent periods. Thus, we are cautious about the results of the trend analysis for the tropical LS. In general, the trends in E and S are more evident in the midlatitude LS than in the tropical LS but are reversed regarding the trend in R. Further analysis and interpretation are provided below. and II ozone products (e.g., McCormick et al. 1992; McPeters et al. 1994; H. J. Wang et al. 1996). Here, we will follow the ozone trend analysis and use SAGE II aerosol products to investigate zonal mean long-term 4. Zonal mean trend In the past, zonal-mean long-term changes in stratospheric ozone have been investigated using SAGE I FIG. 4. Time series from October 1984 to August 2005 for zonalmean extinction coefficient (km21) at 1.02 mm between 358 and 458S at 20-km altitude.

6 804 J O U R N A L O F A P P L I E D M E T E O R O L O G Y A N D C L I M A T O L O G Y VOLUME 51 FIG. 5. Time series from January 1998 to December 2004 for zonal-mean extinction coefficient (km 21 ) at 0.386, 0.452, 0.525, and 1.02 mm, stratospheric aerosol surface area density (mm 2 cm 23 ), and effective radius (mm) between 358 and 458S at 20-km altitude. changes in background stratospheric aerosols. Monthly SAGE II aerosol products are zonally averaged in 108 latitude bins from 508S to508n (to avoid data gaps at higher latitudes during the polar winter). The following linear-regression model commonly used in O 3 zonal trend studies (P. H. Wang et al. 1996; Cunnold et al. 2000; Nazaryan et al. 2005) is adopted for this analysis: x(t) 5 Seasonal 1 Trend 1 C 3 QBO 1 D 3 Solar 1 d(t), (4) where x is the monthly mean of any atmospheric variable of interest and t indicates the month. The Seasonal term defines that variance as demonstrated at the beginning of the section 3, and the Trend term is assumed

7 APRIL 2012 L I U E T A L. 805 FIG. 6. Time series from January 1998 to December 2004 for zonal-mean stratospheric aerosol surface area density (mm 2 cm 23 ), effective radius (mm), and extinction coefficient (km 21 ) at 1.02 and mm between 58S and 58N at 20-km altitude. to be linear. QBO is the quasi-biennial oscillation contribution, which will be represented by the zonal wind at the 30-hPa pressure level over Singapore along with a scaling parameter C. The zonal wind at 30 hpa is calculated using National Centers for Environmental Prediction reanalysis data (obtained online at cpc.noaa.gov/products/wesley/reanalysis.html#data). The QBO effect on tropical stratospheric aerosols has been confirmed by Trepte and Hitchman (1992) using SAGE I and II data and by Vernier et al. (2009) using lidar observations from the Cloud Aerosol Lidar and Infrared Pathfinder Satellite Observations (CALIPSO) platform. The Solar influence is scaled to solar variability on the basis of the radio flux at 10.7 cm with a scaling parameter D. The solar radio flux is taken from the U.S. National Geophysical Data Center database (obtained online at SOLAR/). According to Deshler et al. (2006), the effect of the 11-yr solar cycle on the aerosol in the LS can actually be neglected, which is corroborated by our analysis, given the nearly zero values of scaling parameter D that were obtained. In Eq. (4), d is the residual term. An F test is also applied to the linear trend derived from Eq. (4) to estimate the significance of the trend. Shaded areas in the following height latitude distribution plots for trends (Figs. 8 12) indicate areas with a confidence level of 95% on the basis of an F test. The trends in the zonal-mean values of the SAGE aerosol extinction coefficient E at 1.02 and mm from January of 1998 to December of 2004 are displayed in Figs. 8 and 9, respectively. For E at 1.02 mm, there is a decreasing tendency in the tropical LS at a rate from approximately 0% to 22% yr 21, which is generally below the 95% confidence level. There is an evident positive trend throughout the midlatitude LS from 19- to 26-km altitude, especially over the Southern Hemisphere (SH). The greatest rates of increase (with confidence level above 95%) occur in the latitude bands between 308 and 508S at approximately km and between 208 and 508N at altitudes of approximately km, with values up to 2% 5% yr 21. The trend in the SH midlatitude LS is somewhat more evident than its Northern Hemisphere counterpart. For E at mm, the tendency in the tropical LS becomes more ambiguous

8 806 J O U R N A L O F A P P L I E D M E T E O R O L O G Y A N D C L I M A T O L O G Y VOLUME 51 FIG. 7. As in Fig. 6, but between 358 and 458N. than that of E at 1.02 mm, at rates from 0% to 13% yr 21, which lie below the 95% confidence level. This is probably due to the fact that SAGE observations at mm are more susceptible to measurement uncertainties and smaller volcanic effects than at 1.02 mm (Thomason et al. 1997, 2008). The increasing trend in the midlatitude LS is more evident at mm than at 1.02 mm. The greatest rates of increase (with confidence levels above 95%) occur in the latitude bands 208 and 508S andbetween208 and 508N at altitudes of approximately km, with rates up to 3% 5% yr 21. These global results are consistent with the behavior observed locally by Hofmann et al. (2009) using lidar backscatter and observed regionally by Liu et al. (2007) on the basis of SAGE II data over a limited range of northern midlatitudes. The derived linear trend in the global zonal-mean aerosol surface area density S is shown in Fig. 10. In general, the trend is increasing throughout the LS above 18-km altitude. The rate of increase in the tropical LS is in the range of 0% 5% yr 21 but is generally below the 95% confidence level. We found that the small Reventador volcanic eruption in the tropics interrupts the seasonal term of S in the trend analysis model in the tropical LS, which forced us to set the trend detection in the region to default missing values in the analysis package. In addition to the effects of relatively small volcanic eruptions, relatively larger uncertainty is associated with aerosol surface area density during volcanically quiescent periods, especially in the tropical LS (Thomason et al. 2008); thus, the trend derived here, especially for the tropical LS, is subject to larger uncertainties, which explains why the trend confidence level is generally low in the tropical LS. The greatest rate of increase (at a confidence level above 95%) occurs in the latitude bands that span S and N at altitudes of approximately km; here, the rates are as large as to 2% 6% yr 21. Because the operational SAGE II aerosol surface area density product has a measurement uncertainty of at least a factor of 2 in background aerosol loading conditions (Thomason et al. 2008), it is worthwhile to have a further examination of the corresponding uncertainty in the derived trend of S. According to Leroy et al. (2008), the time t that it takes for a trend T to be detected in the presence of natural variability and measurement uncertainty with an instrument signal-to-noise ratio s can be expressed as

9 APRIL 2012 L I U E T A L. 807 FIG. 8. Height latitude distribution of trends in the zonal-mean aerosol extinction coefficient at 1.02 mm from January 1998 to December The contours define trends in units of percent per year. Shading indicates a confidence level of 95% on the basis of an F test. t 5 [(12s 2 /T 2 )s 2 t] 1/3 (1 1 f 2 ) 1/3 T 5 ss(12t/t 3 ) 1/2 (1 1 f 2 ) 1/2 5 T 0 (1 1 f 2 ) 1/2, (5) where s 2 is the zero-lag variance associated with natural variability and t is autocorrelation time scale for the natural variability. Here, f 2 is the ratio of the total measurement uncertainty multiplied by its autocorrelation time scale to natural variability multiplied by its autocorrelation time scale, and T 0 can be considered to be the trend for the ideal case in which measurement uncertainty is approaching zero (or the trend associated solely with natural variability). Because the uncertainty of S in the background aerosol loading conditions is mainly due to the error from the deriving scheme rather than to calibration error, it is reasonable to assume that the total measurement uncertainty is close to the error from the deriving scheme, which is at least a factor of 2 in the LS, as indicated in Thomason et al. (2008). Considering that the SAGE II operational S product is derived by applying a regression technique to the extinction coefficient at 1.02 mm and the extinction ratio of over 1.02 mm, the error in derived S should come mainly from the incapability of the derived regression formula to capture the complete seasonal variations of S. Therefore, it is reasonable to assume the autocorrelation time scale of the imperfection deriving scheme of S is close to 1 yr, which is also the autocorrelation time scale of the natural variability since the variations of E and S in the LS are dominated by seasonal changes as shown in Figs. 1 and 2. The variance of S associated with natural variability is about or FIG. 9. As in Fig. 8, but for extinction at mm. 70% according to Thomason and Peter (2006), and the measurement uncertainty of S is at least a factor of 2 in the background aerosol loading conditions. Then, f 2 is estimated to be about 8.2 and the trend of S including the measurement uncertainty can be different from the trend of natural variability (or the ideal case) by at least a factor of (1 1 f 2 ) 1/2 3. This suggests that the large uncertainty (at least a factor of 2) in the operational S product may contribute at least two-thirds of the trend detected for S for the background aerosol loading conditions. The height latitude distribution of the zonal-mean trends in particle effective radius R is illustrated in Fig. 11. There is a distinct global-scale decrease in R at km altitudes, which maximizes over the tropics near 21-km altitude at a rate about 22.5% yr 21. This negative center is right at the tropical reservoir region identified by Hitchman et al. (1994), which holds temporally the FIG. 10. As in Fig. 8, but for stratospheric aerosol surface area density.

10 808 J O U R N A L O F A P P L I E D M E T E O R O L O G Y A N D C L I M A T O L O G Y VOLUME 51 FIG. 11. As in Fig. 8, but for aerosol effective radius. constituents (such as injected CN particles in this case) transported from the troposphere to the LS by tropical upwelling before being further transported to mid- and high latitudes. The rate contours follow closely the concentration isopleths of tracers having a tropospheric origin, such as methane, suggesting aging of the aerosol with time as particles are transported meridionally. This is consistent with the ongoing oxidation of injected sulfate precursor gases followed by condensation on particles, along with coagulation and sedimentation, which tend to normalize the size distribution over time (or, equivalently, increasing latitude) in counterbalance to the initial decrease in injected particle size above the tropics. The decrease in the effective radius of the particles as derived from SAGE II data suggests a decrease in particle terminal fall velocity u (and the aerosol removal rate due to sedimentation). According to Kasten (1968), u at 20 km is roughly 0.09 mm s 21 for particles of ;0.2-mm radius and is 0.04 mm s 21 for particles of ;0.1 mm. The average upwelling velocity at 100 hpa in the tropics varies from approximately 0.2 to 0.6 mm s 21 over the annual cycle (Randel et al. 2006), which is nearly one order of magnitude larger than the estimated terminal fall speeds. Accordingly, the effect of changes in sedimentation rates in the tropics can probably be neglected. Toward higher latitudes, however, the cumulative effect of continued settling, combined with weakly descending mean circulation and particle growth, may lead to more significant effects on the aerosol distribution; such effects are not quantifiable in this analysis. Because the SAGE II extinction coefficient measurement is sufficiently reliable at and 1.02 mm for background stratospheric loading conditions (Thomason et al. 2008), we also examine the trend for the ratio of E at and 1.02 mm to further confirm that our trend analysis result of R is also sufficiently reliable since this FIG. 12. As in Fig. 8, but for the ratio of extinction at and 1.02 mm. ratio of E at and 1.02 mm can be considered to be an effective surrogate of aerosol particle size. A large (small) value of the ratio indicates a small (large) particle size. The height latitude distribution of the zonalmean trends for the ratio is displayed in Fig. 12. There is a distinct global-scale increase in the ratio at km altitudes, which maximizes over the tropics near 21-km altitude at a rate of ;8% yr 21. These distribution features are consistent with the features observed in Fig. 11 for R, which suggests that the trend results of R from our analysis should be reliable. 5. Increase of total aerosols The data analysis determines the relative changes in S and R as DS/S 5 0% 5% yr 21 and DR/R 522.5% yr 21 in the tropical LS. For the midlatitude LS the relative changes in S and R are DS/S 5 2% 6% yr 21 and DR/R 5 21% yr 21. From Eq. (3), we can determine that DV/V 5 (DS/S 1DR/R). (6) Then the corresponding relative changes in V, DV/V, are from 22.5 to 12.5% yr 21 in the tropical LS, and DV/V 5 1% 5% yr 21 in the midlatitude LS. Because the uncertainty of DR/R in Eq. (6) is much smaller than the uncertainty of DS/S according to our above analyses, the uncertainty of DV/V should be close to that of DS/S and should be associated with the large uncertainty in S for background stratospheric aerosol loading conditions. If the aerosol particle density is roughly constant (e.g., 75% of H 2 SO 4 and 25% of H 2 O), the aerosol mass trends would be the same as the volume trends. The bulk aerosol mass changes estimated here for the midlatitude LS agrees reasonably well with the estimate (5% 6 2%)

11 APRIL 2012 L I U E T A L. 809 of Hofmann et al. (2009). To sustain this change, it is required that 0.6% 0.8% of tropospheric sulfur enter the stratosphere each year as suggested by Hofmann et al. (2009), who also provided an excellent analysis on the ballpark relationship between the anthropogenic sulfur emission and the fraction eventually transported to the stratosphere. The number distribution of stratospheric aerosols is generally represented as a bimodal lognormal distribution, with one mode typically at,0.1-mm radius and a larger mode at ; mm. The larger mode tends to be more pronounced in aerosols perturbed by a volcanic eruption. The particle number and surface area are dominated by the small mode, whereas the total particle mass can be concentrated in the larger mode (Thomason and Peter 2006). In practice, spherical particles and a unimodal lognormal distribution have often been assumed to represent the size distribution of nonvolcanic stratospheric aerosols (Turco et al. 1982; Yue et al. 1986; Bauman et al. 2003). A self-consistent change in the aerosol number concentration N 0 can then be inferred approximately from the analyzed changes in S and R. Note that an increase in S and a decrease in R, as observed, necessarily imply an increase in N 0. This can be demonstrated using the relationships derived below. First note that the dependence of the total aerosol sulfate mass M s on measured parameters is readily expressed [using Eq. (3)] as M s 5 rwv } RS, (7) where r is the particle density, W is the acid (sulfate) weight fraction of the aerosol solution, and the product rw is assumed at this point to remain constant. The variation in M s is then also given by Eq. (6). Now, by assuming that the aerosol particles are spherical, with a lognormal distribution having a constant standard deviation of the radius (or sigma), and with an equivalent effective radius R, it can be shown (ignoring the form factors that arise in taking lognormal moments, since these are constant because sigma is fixed) that from which it follows immediately that S } N 0 R 2, (8) DN 0 /N 0 5DS/S 2 2(DR/R). (9) From the analysis in section 4, the relative changes in S and R in the tropical LS are DS/S 5 0% 5% yr 21 and DR/R 522.5% yr 21. Accordingly, the corresponding relative change in N 0 can be estimated from Eq. (9) as DN 0 /N 0 5 5% 10% yr 21. In the midlatitude LS, the relative changes in S and R are DS/S 52% 6% yr 21 and DR/R 521% yr 21, yielding the corresponding relative change in N 0 as DN 0 /N 0 5 4% 8% yr 21. Again the uncertainty of DR/R in Eq. (9) is much smaller than the uncertainty of DS/S. Thus, the uncertainty of DN 0 /N 0 should be close to that of DS/S and should be associated with the large uncertainty in S for background stratospheric aerosol loading conditions. The inferred changes in N 0 in the tropical LS are somewhat larger than those in the midlatitude LS. 6. Some discussions Because the effective radius of the particles as derived using SAGE observations shows a decreasing trend during recent volcanically quiescent periods, the corresponding secular increases in S and E suggest that there has also been an increase in the total stratospheric aerosol number concentration. Hofmann et al. (2009) had speculated that the cause of an increase in lidar backscatteroverasimilarperiodhaditsbasisinthe growth of the aerosol particles due to enhanced anthropogenic SO 2 emissions, under the influence of the temperature and humidity variations observed by Randel et al. (2006). To back up this supposition, Hofmann et al. (2009) characterized anthropogenic SO 2 emissions from 1970 to 2005, along with fossil-fuel consumption rates from 1980 to 2007, and found an evident increase in anthropogenic SO 2 emissions since The enhancement in the stratospheric aerosol appears to have begun earlier than 2002 in the SAGE II data, however, leading us to explore other potential causes, as described below. There are two microphysical processes that may partially explain an increase in aerosol surface area. The first is related to changes in ambient temperatures and water vapor content, which affect the composition of the sulfate particles. As temperature decreases, or water vapor increases, the aerosol will tend to grow in size for the same condensed sulfur mass. This would result in an increase in both S and R, which is inconsistent with the observed increase in S and decrease in R derived from the SAGE II data. Moreover, the water vapor content in the LS has tended to decrease rather than increase after 2001, as determined on the basis of both satellite and balloonborne measurements (Randel et al. 2006). Hence, this mechanism is not likely to be the explanation. The second mechanism would involve a general increase in both the total sulfur mass and the concentration of CN through tropical upwelling. Enhanced nucleation of new particles in the tropical LS would initially increase the total particle surface area and number, although the aerosol would likely to be closer to the original CN size (smaller) prior to photochemical

12 810 J O U R N A L O F A P P L I E D M E T E O R O L O G Y A N D C L I M A T O L O G Y VOLUME 51 conversion of the sulfur precursors to sulfates during transport to higher latitudes. Because of continuing condensational growth along with coagulation and sedimentation during transport, one would expect DN 0 in the midlatitude LS to be somewhat smaller than in the tropical LS, which is consistent with the preceding analysis. Accordingly, our analysis supports the speculation of Hofmann et al. (2009) that a major cause of the recorded increase in lidar backscatter over a similar period at two midlatitude locations was aerosol growth due to enhanced anthropogenic SO 2 emissions. Moreover, we conclude that an enhanced injection rate of CN from the troposphere at tropical latitudes would be required to sustain the observed increase in stratospheric aerosol concentrations. The exact cause of the CN source enhancement is unknown. Binary homogenous nucleation of new H 2 SO 4 -H 2 O particles in the upper tropical troposphere would be injected into the tropical LS and would serve as a source of nuclei for stratospheric aerosol formation, according to Brock et al. (1995). New CCN particle formation was also observed in the tropical/subtropical thin cirrus clouds by airborne in situ measurements, but the formation mechanisms were unidentified, according to Lee et al. (2004). Higher rates of biomass burning and enhanced fine-dust generation perhaps caused by changes in surface conditions related to global warming may also contribute to the CN enhancement. For example, recent CALIPSO lidar profile observations (Vernier et al. 2009) reveal evident small particle formation and enhancement extending vertically from the tropical upper troposphere to the tropical LS and horizontally from northern India to southern Arabia and North Africa between July and August (the most convective season). Vernier et al. (2009) speculated that desert dust particle transport by the strong convective systems might be the possible source of the new particles over southern Arabia and North Africa because of the depolarizing property of the particles detected by CALIPSO lidar. They also suggested that the new particles observed over southern Asia during the monsoon season could be related to the transport of sulfur dioxide in convective systems near the cold-point tropopause where oxidation by the hydroxyl radical (OH)couldleadtonewparticleformation. Alternatively, stronger tropical upwelling after 2001 associated with a colder and dryer LS identified by Randel et al. (2006) (a result also supported by our analysis of the SAGE II temperature and water vapor data, but not elaborated here) could also increase the transport of CN into the LS. Each of these processes may play a role in augmenting CN abundances. More conclusive results will require a deeper investigation of the tropical dynamics during the periods of interest and of the possible sources of surface particulates that can reach the LS. As mentioned in several places in the above sections, relatively larger uncertainties are associated with our trend analyses (especially for S, V, and N 0 ) in the tropical LS as compared with in the midlatitude LS because of two factors: 1) the impact of relatively small volcanic eruptions, such as Reventador in September of 2002, and 2) the relatively larger uncertainty in aerosol surface area density (and other derived aerosol microphysical properties) due to reduced sensitivity of operational SAGE II aerosol measurements for aerosol surface area density at the lowest abundances of stratospheric aerosols during volcanically quiescent periods. This is why the trends detected for these variables in the tropical LS generally have a lower confidence level than in the midlatitude LS. The impact of these two factors on the stratospheric aerosol effective radius R may be somewhat alleviated considering that R depends only on the ratio of S and V [Eq. (3)] and is confirmed by the trend analysis for the ratio of E at and 1.02 mm. The trends in Figs were further confirmed by removing the effect of small volcanic eruptions. The results are almost identical to the current Figs except for very small differences in the tropical LS. In summary, more than two-thirds of the trends detected for S, V, and N 0, especially in the tropical LS, are probably due to the large uncertainty (at least a factor of 2) in the operational S product, but the natural variability should be the major cause for the trend in the effective radius. 7. Summary and conclusions The Stratospheric Aerosol and Gas Experiment II aerosol products from 1998 to 2004 have been examined to identify trends in the background stratospheric aerosol layer during the most volcanically quiescent periods over the last two decades. Our major findings include the following: 1) Background aerosol extinction coefficients E have tended to increase in the midlatitude lower stratosphere, at a rate as large as 2% 5% yr 21 ; the total particle surface area density S has also increased, especially in the midlatitude LS, where the rate is in the range of 2% 6% yr 21. More than two-thirds of the trend detected for S is probably due to the large uncertainty (at least a factor of 2) in the operational S product according to our analysis. 2) The effective particle radius R has tended to decrease, mainly in the tropical LS, at a rate of up to 22.5% yr 21. 3) Correspondingly, stratospheric aerosol number concentrations are inferred to have increased at a rate in the range of 5% 10% yr 21 in the tropics and 4% 8% yr 21 at midlatitudes. 4) An overall increase in aerosol mass has occurred mainly in the

13 APRIL 2012 L I U E T A L. 811 midlatitude LS at a rate in the range of 1% 5% yr 21. The impact of measurement uncertainties on the trend of stratospheric aerosol number concentrations and mass is comparable to that on particle surface area density, but the corresponding impact on the trend of effective radius is evidently alleviated. Our trend analysis is at a higher confidence level in the midlatitude LS than in the tropical LS. Factors contributing to these fundamental changes in the ambient stratospheric aerosol layer include an enhancement in anthropogenic SO 2 emissions since 2002, which has been identified by Hofmann et al. (2009), and enhanced upwelling transport associated with a colder and dryer LS, as suggested by Randel et al. (2006). The work presented here also suggests an increasing rate of injection of tropospheric condensation nuclei at tropical latitudes, followed by meridional transport and aging, as a possible explanation for increases in stratospheric aerosol concentrations derived here. The exact cause of this CN source enhancement is still unresolved but may involve higher rates of biomass burning and fine-dust generation at the surface in the tropics and subtropics or may involve enhanced binary homogeneous nucleation of H 2 SO 4 and H 2 O in the upper tropical troposphere (which is a source of nuclei for the stratospheric aerosol). Efficient transport of these CN into the LS is consistent with the strengthening of tropical upwelling that has been indicated by recent observations (e.g., Randel et al. 2006). Acknowledgments. The authors acknowledge the NASA Langley Climate Data Center for providing operational SAGE II aerosol products and acknowledge communications with Dr. Larry W. 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