JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. D3, 4114, doi: /2001jd001077, 2003

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1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. D3, 4114, doi: /2001jd001077, 2003 Meridional distributions of aerosol particle number concentrations in the upper troposphere and lower stratosphere obtained by Civil Aircraft for Regular Investigation of the Atmosphere Based on an Instrument Container (CARIBIC) flights M. Hermann, J. Heintzenberg, and A. Wiedensohler Institute for Tropospheric Research, Leipzig, Germany A. Zahn and G. Heinrich Forschungszentrum Karlsruhe/Universität Karlsruhe, Karlsruhe, Germany C. A. M. Brenninkmeijer Max Planck Institute for Chemistry, Mainz, Germany Received 9 July 2001; revised 20 July 2002; accepted 19 November 2002; published 8 February [1] Number concentrations of submicrometer aerosol particles at altitudes of km were measured along a single flight route between Germany (50 N, 10 E) and the Indic (5 N, 80 E) over 3 years using a commercial aircraft platform (project Civil Aircraft for Regular Investigation of the Atmosphere Based on an Instrument Container (CARIBIC)). During 41 intercontinental flights (380 flight hours), more than one million individual particle concentration measurements were made, yielding a comprehensive and unique aerosol data set. Using these data, the first meridional and seasonal probability distributions for ultrafine particles and for Aitken mode plus accumulation mode particles in the upper troposphere (UT) and lower stratosphere (LS) were derived. High particle number concentrations were observed in summer at tropical latitudes over the Arabian Sea and at midlatitudes over Europe, in contrast to lower values in the subtropics over the Middle East. This distribution primarily reflects the vertical transport pattern of the atmosphere. In winter, number concentrations were generally lower. The Intertropical Convergence Zone (ITCZ) was identified as a region of pronounced particle formation, but particle nucleation also occurred in the upper troposphere at midlatitudes. At tropical latitudes, convective transport and photochemistry appear to be the main driving forces for particle formation in the free troposphere. Consistent with these observations, integral length scales calculated from the aerosol data ranged from 6 to 12 km. Compared to particle concentrations in the midlatitudinal upper troposphere, those in the lowermost stratosphere were lower in winter, but equal or higher in summer. This seasonality is likely caused by stronger in-mixing of upper tropospheric air in summer. INDEX TERMS: 0305 Atmospheric Composition and Structure: Aerosols and particles (0345, 4801); 0365 Atmospheric Composition and Structure: Troposphere composition and chemistry; 0368 Atmospheric Composition and Structure: Troposphere constituent transport and chemistry; 3309 Meteorology and Atmospheric Dynamics: Climatology (1620); KEYWORDS: upper troposphere, lower stratosphere, submicrometer aerosol particles, particle formation, in situ measurements, aerosol spatial distributions Citation: Hermann, M., J. Heintzenberg, A. Wiedensohler, A. Zahn, G. Heinrich, and C. A. M. Brenninkmeijer, Meridional distributions of aerosol particle number concentrations in the upper troposphere and lower stratosphere obtained by Civil Aircraft for Regular Investigation of the Atmosphere Based on an Instrument Container (CARIBIC) flights, J. Geophys. Res., 108(D3), 4114, doi: /2001jd001077, Introduction [2] Aerosol particles in the atmosphere affect global climate, directly by absorbing and scattering solar radiation [Schwartz, 1996] and indirectly by influencing radiative Copyright 2003 by the American Geophysical Union /03/2001JD properties and the lifetime of clouds [Twomey et al., 1984]. Furthermore, atmospheric particles provide surface area and volume for heterogeneous and multiphase chemical processes and thus play a significant role in atmospheric chemistry [Ravishankara, 1997]. A reliable assessment of these influences requires detailed information on particle properties such as particle number concentration, size distribution, composition, and lifetime throughout the atmos- AAC 10-1

2 AAC 10-2 HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS phere. Because of the low temperatures and the usually low preexisting particle surface area (i.e., a small condensation sink for precursor gases), the upper troposphere (UT) is assumed to be a region of frequent new particle formation [Brock et al., 1995]. Such fresh nuclei can grow to form cloud condensation nuclei (CCN) for cirrus clouds [Laaksonen et al., 2000] and thereby can influence the climate. In particular, in the Intertropical Convergence Zone (ITCZ), deep convective clouds transport particles as well as particle precursor gases (e.g., DMS and SO 2 ) of both natural and anthropogenic origin from the surface to higher altitudes [Thornton et al., 1997]. During transport, these precursor gases are oxidized, thus promoting homogeneous particle nucleation of the oxidation products at higher altitudes (mainly sulfuric acid/water particles) [Brock et al., 1995; Clarke et al., 1998; Weber et al., 1999; de Reus et al., 2001]. Because in-cloud particle removal processes during upward transport reduce the available particle surface area, and because particle formation is favored when particle surface area is low [Clarke et al., 1998; de Reus et al., 1998; Andreae et al., 2001], particle nucleation in the vicinity of cloud tops is likely [Perry and Hobbs, 1994; Clarke et al., 1998, 1999; Ström et al., 1999; de Reus et al., 2001; Weber et al., 2001]. [3] Despite their importance for atmospheric processes, experimental data on submicrometer particles in the free troposphere are rare. In 1989, Heintzenberg concluded that there are no data available to characterize fully the submicrometer aerosol above the boundary layer [Heintzenberg, 1989]. In the 1990s, several measurement campaigns were carried out in the UT, typically over relatively short periods of a few weeks [e.g., Hagen et al., 1995; Brock et al., 1995; Schröder and Ström, 1997; de Reus et al., 1998; de Reus et al., 2001]. Only a handful of studies provided data with larger temporal [e.g., Hofmann, 1993; Hofmann et al., 1998] or spatial coverage [e.g., Clarke, 1993; Clarke et al., 1999]. Therefore, even today the knowledge of submicrometer particles in the UT is limited [cf. Raes et al., 2000]. [4] One technique for making aerosol measurements in the free troposphere is to use remote sensing systems on satellites [cf., e.g., Kent et al., 1995; King et al., 1999]. Although such systems have the advantage of permitting global measurements, they have disadvantages, such as clouds limit their field of vision, the spatial resolution of the data is relatively coarse, and obtained data need validation by in situ measurements. Furthermore, satellite-based instruments use optical measurement techniques that can only detect optically efficient scattering (i.e., particles with diameters larger than 100 nm) and therefore they are not useful for studying particle-formation processes. Aircraftbased in situ measurements are more appropriate for measuring aerosol particles, because instruments for measuring sub-100-nm sized particles can be readily accommodated aboard aircraft. However, aircraft-based aerosol measurements are rare due to (1) high costs of operating specialized research aircraft, (2) difficulties associated with high-speed aerosol inlet systems, (3) the fact that instruments must operate under variable and low pressure conditions, and (4) mechanical and electrical aviation security requirements that must be met by the instruments [Daum and Springston, 1993]. [5] A promising approach for collecting long-term data sets with near global coverage is the use of commercial aircraft [Fleming, 1996]. Compared to research aircraft, commercial aircraft cover larger areas of the globe and can be regularly used for long times (on the order of years), thus providing extensive spatial and temporal coverage. Furthermore, operational costs are much lower than those for specialized aircraft. The development of fully automated measurement devices for commercial aircraft, however, is also complicated and expensive. Since 1997, a commercial aircraft of LTU International Airways is used as atmospheric measurement platform in the project Civil Aircraft for Regular Investigation of the Atmosphere Based on an Instrument Container (CARIBIC) [Brenninkmeijer et al., 1999] (see also In this study, we present a comprehensive data set of number concentration measurements of submicrometer aerosol particles in the upper troposphere and lower stratosphere (UT/ LS) measured along a single flight route within CARIBIC. Using this data set comprising three years of fairly regular measurements, we derived the first meridional and seasonal probability distributions of ultrafine particles (diameter <10 nm) and for Aitken mode ( nm) plus accumulation mode ( nm) particles in the UT and LS, respectively. Furthermore, we report typical length scales for the spatial distribution of these particles. 2. Experimental [6] Particle number concentrations were measured in two particle size ranges (4 nm 1.3 mm and 12 nm 1.3 mm) by two condensation particle counters (CPCs, TSI Model 7610, St. Paul, MN). The measurement time resolution was set to 2 s, corresponding to a spatial resolution of 0.5 km. The CPCs were flown aboard a passenger aircraft (B ER, LTU International Airways) as part of the CARIBIC project. In this project, a modified standard air-freight container (Figure 1) is used to house instruments for measuring trace gases (O 3, CO, and air samples) and aerosol particles (particle concentration and particle samples). For each measurement flight, the container is placed in the forward cargo compartment of the B 767. Separated, a dedicated aerosol and trace gas inlet is permanently installed on the aircraft fuselage directly below the container. The aerosol inlet tube starts with a 5 diffuser section followed by a tube section, where a backwards facing stainless steel tube with 4 mm inner diameter is used to aspirate an aerosol flow of 11 l/min [Hermann et al., 2001]. To prevent flow separation at nonzero angles of attack, blunt inlet lips according to a NASA code (NACA 1-series [Soderman et al., 1991]) are used. The inlet is connected to the particle counters inside the container by a 2 m stainless-steel sampling line. To obtain accurate number concentrations, the raw data were corrected for particle losses in the inlet system. Therefore, the inlet sampling efficiency and the transport efficiency through the sampling line were determined by calculations and by making wind-tunnel measurements [Hermann et al., 2001]. The obtained overall sampling efficiencies are nearly independent of cruise altitudes and are about 73% (±24%) for ultrafine particles and 90% (±4%) for Aitken mode plus accumulation mode particles, respectively.

3 HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS AAC 10-3 Figure 1. Location of the CARIBIC container in the forward cargo compartment of the B 767 (LTU International Airways), directly above the aerosol inlet, and cross-sectional view of the container (direction out of the page is the flight direction). [7] Before the initial use of the CPCs, they were checked for vacuum tightness and were modified to comply with aviation requirements (e.g., 28 V DC power supply, replacement of combustible material). Extensive laboratory tests at operating pressures of hpa were made to ensure the proper functioning of the instruments at flying altitudes, and to obtain pressure-dependent counting efficiency curves [Hermann and Wiedensohler, 2001]. To measure different particle size ranges with the CPCs, the lower threshold diameter in one CPC was shifted to a smaller particle size by increasing the temperature difference between the saturator and the condenser block inside the instrument from the original setting of 17 C to 27 C. Therefore, an additional Peltier element was installed in the CPC. The temperature difference was adjusted by a potentiometer and two additional temperature sensors were used to monitor the temperature difference between the saturator and the condenser block. Despite the higher temperature difference, no homogeneous nucleation was observed in the CPC during extensive laboratory tests. As a result of the modification, the two CPCs have respective lower threshold diameters (50% counting efficiency) of 4 and 12 nm at flight level pressures. The upper threshold diameter of the measurements is determined by the inlet system, which only collects particles smaller than 1.3 mm diameter (50% efficiency diameter) [Hermann et al., 2001]. By subtracting the readings of the two CPCs, the number concentration of ultrafine particles (here defined as particles with diameters between 4 and 12 nm, N 4 12 ) can be determined. Concentrations measured by the second CPC, which measures all particles larger than 12 nm, can be considered as the sum of Aitken mode and accumulation mode particles (N 12 ). Since spring 1999, a third CPC with a lower threshold diameter of 18 nm is used in the CARIBIC payload. However, because the number of available flights with this CPC was too low for a reliable statistical analysis, only measurements from the first two CPCs are presented in this study. [8] During data processing, measured number concentrations were corrected for particle losses in the inlet and in the sampling line, for pressure-dependent CPC flow rates, for CPC counting efficiencies, and for coincidence in the CPC optics [cf. Hermann, 2000]. Together, these corrections account for a particle-size- and number-concentrationdependent increase in raw particle concentrations by a factor of 1.2 to 2. To report pressure-independent number concentrations, all number concentrations given in this study are converted to standard conditions (STP: hpa and K). [9] Uncertainties in the derived number concentrations occur due to uncertainties in the CPC flow rate, the particle coincidence correction, the CPC counting efficiency, and the inlet sampling efficiency. Combining all errors results in an overall average uncertainty of approximately ±10% for N 12 [Hermann, 2000]. Because N 4 12 is calculated by subtracting two CPC readings, each containing counting errors, the N 4 12 uncertainty is higher, but less than 35% when concentrations are above 1000 particles/cm 3. However, due to statistical fluctuations it can exceed 100% when N 4 12 is below 50 particles/cm 3. [10] For this study, data of 45 intercontinental flights from either Munich or Dusseldorf (Germany) to either Male (Maldives) or Colombo (Sri Lanka) were available, covering central and south-east Europe, the Middle East, and the Arabian Sea between 50 N, 10 E and 5 N, 80 E (Figure 2). These flights were carried out between June 1997 and November A detailed description of the meteorological conditions encountered during the flights can be found in the work of Zahn et al. [2002]. Figure 3 shows a typical time-altitude profile for the flight on 5 June 1998 as an example. On this flight, the aircraft started at 0600 UTC at Colombo, flew in the middle tropical troposphere at 31,000 feet for several hours, and changed altitude once to 35,000 feet. At higher latitudes, the aircraft flew just below the tropopause (indicated in Figure 3 by potential vorticity (PV) isolines between 1.5 and 3.5 PVU ( potential vorticity units)), entered the LS for 80 min, and finally returned to the midlatitude UT. At 1530 UTC the aircraft landed in Munich. 3. Data Analysis 3.1. Data Set [11] An example of the aerosol data measured by CARI- BIC is shown in Figure 4. It depicts particle number concentrations and ozone mixing ratios for the flight already discussed in Figure 3. N 12 (black area) and N 4 12 (gray

4 AAC 10-4 HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS Figure 2. Flight routes of 45 CARIBIC flights carried out between summer 1997 and autumn area) were highly variable along the flight and reached highest values between 0600 and 0900 UTC at tropical latitudes, i.e., over the Arabian Sea, and between 1300 and 1500 UTC at midlatitudes, i.e., over Europe. In contrast, particle number concentrations were lower in the subtropics over the Middle East. A deeper discussion of this feature and the underlying processes based on statistical evaluation is given in section 4. Further information on this individual flight can be found in the work of Zahn et al. [2000]. [12] For statistical analysis of the data set, only measurements between 8.5 and 11.3 km altitude (28,000 to 37,000 feet, or flight levels (FL) 280 to 370) were used. Approximately 85% of these measurements were made at FL310, FL330, and FL350 (i.e., in a relative small altitude band from 9.4 to 10.7 km), independent of season. However, there is a tendency to higher flight levels in summer compared to winter. Within a season, the three flight levels were almost uniformly distributed over latitude. Data measured during aircraft ascents and descents were not included for two reasons. First, due to aircraft operational restrictions the CPCs are not powered until a preset, minimum pressure level is reached, which was in the range of 500 hpa to 850 hpa. Because the instruments need several minutes to reach the set-point temperature difference, the lower threshold diameter of the CPCs during aircraft ascent is not constant. Second, because the aircraft pitch angle during ascent and descent can exceed 10, the performance of the inlet system under these extreme angles may vary, making data analysis difficult. Flight level changes, however, were included in the analysis, because pitch angles during altitude changes were small, less than 5, and aerosol data in these periods did not indicate any sampling artifacts. [13] For distinguishing between tropospheric and stratospheric air at midlatitudes, in other studies frequently relatively low-resolved PV fields were used. Although calculated PV fields are based on weather-balloon and satellite measurements, these fields show the spatial and temporal resolution of the underlying transport model. Air masses with PV values above PVU are attributed to the stratosphere and below this threshold to the troposphere [World Meteorological Organization (WMO), 1986; Hoerling et al., 1991]. However, PV fields calculated by the currently available models (e.g., ECMWF, UKMO) are not suited to account for fine dynamical structures at the tropopause (e.g., stratospheric filaments), because at the tropopause they show spatial resolutions of only km in Figure 3. Representative time-altitude profile for a CARIBIC flight (flight Colombo-Munich on 5 June 1998). The thick solid line denotes the flight track. Potential temperature (in K, dashed dotted lines) and potential vorticity isolines (in PVU, thin solid lines) were calculated by H. Cuijpers and P. van Velthoven, KNMI, de Bilt, Netherlands. Black areas near the x axis represent overflown mountain ranges.

5 HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS AAC 10-5 Figure 4. Particle number concentrations from the CARIBIC flight Colombo-Munich on 5 June N 12 (black area) indicates the particle number concentration of Aitken plus accumulation mode particles. N 4 12 (gray area) indicates the particle number concentration of the ultrafine particles. The thick dashed line denotes the flight pressure and the thin solid line the ozone mixing ratio (50 s moving average) measured with an UV absorption photometer [cf. Brenninkmeijer et al., 1999]. The arrow shows the position of the ozone tropopause (TP) for the flight date according to Zahn et al. [2002]. the horizontal direction and km in the vertical direction. Therefore, for this study an ozone tropopause and in situ measured ozone mixing ratios were used to distinguish between tropospheric and stratospheric air at midlatitudes. According to recent findings by Zahn et al. [2002], at midlatitudes the ozone tropopause threshold value (O 3 TP ) along the CARIBIC flight route (Figure 2) undergoes a uniform seasonal cycle that can be well approximated by a sinusoidal function, which minimizes near 1 November at 71 ppb and maximizes near 1 May at 123 ppb. Accordingly, measurements with ozone mixing ratios either above or below O 3 TP are attributed to either the LS or the UT, respectively. The advantage of this tropopause definition can be clearly seen by comparing the resolutions of PV and ozone in Figures 3 and 4. [14] A further subdivision of the data into seasonal data sets for spring, summer, autumn, and winter was considered, but was not realized, because the measurement flights are not uniformly distributed over the year. In particular, there are fewer flights in spring and autumn, leading to poor statistics during those seasons. Instead, two expanded seasons were defined: summer, spanning from May through September, and winter, covering November through March. Consequently, four intermediate flights during April and October were omitted, and only 41 flights (23 for summer and 18 for winter) were included in the statistical analysis. The resulting data set still comprises 380 flight hours and more than one million individual particle concentration measurements in the UT/LS Probability Distributions and Integral Length Scales [15] To investigate the meridional distribution of UT/LS particles, probability distributions for the particle number concentration were calculated for each of the two atmospheric regions, two size ranges, and two seasons. Each subset was grouped into 1 latitudinal bands, and then within each of these bands, grouped into 30 concentration bins of equal logarithmic width (0.065 and for N 12 and N 4 12, respectively). The resulting distributions give the probability of encountering a particular particle number concentration at a given latitude in the altitude range km. We do not report information on the altitude dependency of UT particle number concentrations within this altitude range, because even at coarser horizontal resolution, the number of individual flights within each latitudinal band and at each flight level is too small to obtain statistically meaningful results. Moreover, as shown in section and Figure 11, for LS data the change in particle concentration with altitude was mainly determined by the original relative altitude of the air mass above the tropopause, and not by the flight altitude. To obtain statistically significant results, for the meridional probability distributions, latitude bands were included only if a mini-

6 AAC 10-6 HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS mum of 1100 data points from at least eight flights were available for tropospheric data and from at least five flights for stratospheric. Similar distributions were generated to investigate the correlation between the percentage of ultrafine particles and local daytime. These distributions were calculated by using time bins of 1.5 hours and percentage bins of 10% width. Compared to the meridional distributions, however, statistics were worse, because for most time bins, data from only four flights were available, and for one bin, data from only two flights were available. [16] Because the particle number concentrations obtained by CARIBIC can be regarded as time series, autocorrelation analysis can be applied to each flight. Autocorrelation analysis investigates the relationship between data points of a variable x with other data points of the same variable x at a former or later times [cf., e.g., Kaimal and Finnigan, 1994]. As result, the autocorrelation function is obtained, which is a measure of the relationship of values measured at different times. The autocorrelation function can be used to calculate an integral timescale, which can be converted to an integral length scale by accounting for the aircraft velocity. This length scale gives information on the spatial dimensions of underlying physical processes and can be derived from the integral over the autocorrelation function, calculated between a lag of zero and the first zero-line crossing [Kaimal and Finnigan, 1994]. For most latitudinal bins, the probability distribution for particle number concentrations is biased toward higher concentration values and has approximately a lognormal shape (as shown in Figure 5 where the spread of the probability is nearly equal to both sides of the maximum values on a logarithmic concentration scale). Because normally distributed data were desired for autocorrelation analysis, the logarithm of the concentrations was used. Furthermore, the individual flights were divided into subsets for tropical, subtropical, and midlatitudinal as well as for UT and LS data. To fulfill the stationarity requirement for autocorrelation analysis, the data were detrended using a high-pass Gauss filter with a cut-off wavelength of 250 km. With respect to time series, stationarity demands that statistical properties, such as the mean or the variance, reach stable values as the investigated period is increased [cf. Kaimal and Finnigan, 1994]. The detrending implies, however, that spatial features in the particle number concentration larger than 250 km, such as those caused by different flight altitudes, are not represented in the obtained results. For analysis, eventually only those sections of flight were used, where at least 30 min (respectively 430 km) of measurement within one of the above stated domains were available. This criterion was fulfilled by about 50% of all data points. 4. Results and Discussion 4.1. Tropospheric Particle Distributions Aitken Plus Accumulation Mode Particles (N 12 ) [17] The probability distributions from 7 N to 49 N derived for tropospheric N 12 in summer and winter are shown in Figure 5. Colors indicate the percentage of all data points in a particular latitudinal band that fall into a certain concentration bin. The summer distribution of N 12 (Figure 5a) shows the same significant meridional variation already visible in Figure 4. This feature can be described by three distinct geographic regions with different ranges for the particle number concentration as follows. First, over the Arabian Sea at tropical latitudes south of 22 N, N 12 was relatively high and variable, mostly between 2000 and 8000 particles/cm 3. Second, in contrast, over the Middle East at subtropical latitudes (25 38 N) N 12 was lower and less variable, mostly in the range of particles/cm 3. Third, toward higher latitudes, the particle number concentration became more variable again and peaked around particles/cm 3. [18] Compared to summer, in winter the differences in N 12 between the tropical, subtropical, and midlatitudinal regions was not as distinct (Figure 5b). In the tropics, N 12 was less variable than in summer, on average only 60% of the summer value, which is still high. The subtropical region, which had relatively low, stable N 12 in summer, was less clearly marked in winter, but is still discernable, located somewhat more south. Toward higher latitudes, N 12 again became relatively high and more variable than at the more southern latitudes. In contrast to summer, however, N 12 over Europe decreased north of 42 N and was on average only half of the summer values. [19] In general, sources of UT N 12 particles can be attributed to two kinds of processes: (1) transport, primarily import from lower altitudes due to vertical uplift, and to a lesser degree injection from the LS; and (2) in situ formation, e.g., caused by mixing processes, atmospheric waves, and aircraft emissions. The majority of these processes are triggered by upward transport of air masses, particularly due to convection and cyclonic lofting. Therefore, the meridional distribution of UT particles shown in Figure 5 should primarily reflect the frequency of such vertical transport processes. [20] In summer, near the Indian subcontinent, the ITCZ is located at latitudes up to 25 N [cf.zahn et al., 2002, Plate 1]. In the vicinity of the ITCZ, deep convective clouds carry particles as well as their precursor gases into the middle and upper troposphere, where new particles are probably formed [Brock et al., 1995; Raes et al., 2000; de Reus et al., 2001]. Therefore, as shown in Figure 5, N 12 is high in the tropics. The contrasting low number concentrations observed at latitudes south of 14 N can most likely be attributed to cleaner air masses originating south of the ITCZ. This is shown in Figure 6, which depicts data of a flight from Male to Dusseldorf in summer The trajectories in the upper graph indicate that N 12 values below 1000 particles/cm 3 (lower graph) were measured in air masses that mostly originated at lower altitudes to the south. These air masses were transported by the prevailing south-west trade winds of the Indian summer monsoon northward, and were finally uplifted within the ITZC to flight altitudes. An even more convincing indicator for clean air is the CO mixing ratio during this period (lower graph), near 65 ppb, a value typical for the southern-hemispheric free troposphere [Seinfeld and Pandis, 1998]. [21] Over the Middle East at subtropical latitudes, i.e., in the downward branch of the Hadley circulation, large-scale subsidence prevails, and uplifting of polluted boundary layer air to higher altitudes and particle formation is less likely. This explains the observed smaller particle number concentrations in Figure 5. The increased concentrations and large variability at midlatitudes over Europe can be

7 HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS AAC 10-7 Figure 5. Probability distribution of Aitken plus accumulation mode particles (N 12 ) in the troposphere in (a) summer and (b) winter. Colors indicate the percentage of all particle measurements in a latitudinal band that fall into a particular concentration bin. Number concentrations are given at standard conditions (STP: hpa, K). attributed to several sources, such as aircraft emissions [Schlager et al., 1997; Hofmann et al., 1998; Anderson et al., 1999; Schröder et al., 2000], upward transport of polluted surface air [Talbot et al., 1998; Hermann, 2000; Wang et al., 2000], in particular by warm conveyor belts [Stohl, 2001], or in situ particle formation processes (cf. section 4.1.2). [22] The lower N 12 found at tropical latitudes in winter can probably be attributed to the ITCZ being located in the southern hemisphere [cf. Zahn et al., 2002, Plate 1] and hence to deep convective clouds being less frequent over the Arabian Sea, as indicated by Meteosat-5 satellite images. Nevertheless, particle formation related to such clouds can still be found in this region and during this season, as shown by de Reus et al. [2001]. Similarly, the decrease in N 12 observed over Europe at latitudes north of 42 N might be explained by the reduced convective activity over continents at midlatitudes in winter (cf. Hov and Flatøy [1997] and the discussion by Stohl [2001]). The decrease of aircraft emissions over Europe of about 20% during winter compared to summer [Friedl, 1997] might also contribute to the observed difference.

8 AAC 10-8 HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS Figure 6. Flight section of the CARIBIC flight Male-Dusseldorf on 7 July Upper graph indicates trajectories (gray lines) at p < 500 hpa (solid lines) and p > 500 hpa (dashed lines). Lower graph shows the Aitken plus accumulation mode concentration (N 12, solid line) and the CO mixing ratio (solid squares). [23] A comparison of the obtained Aitken plus accumulation mode particle number concentrations with results of other experimental studies is difficult, because UT particle data with such a large spatial and temporal coverage are rare. For similar altitudes, in the tropical middle troposphere particle number concentrations of particles/cm 3 were reported by Clarke [1993], Brock et al. [1995], and de Reus et al. [2001]. Clarke [1993] furthermore reported number concentrations for the subtropics in the range of particles/cm 3. At midlatitudes, concentrations of particles/cm 3 were most frequently measured [Hofmann, 1993; Brock et al., 1995; Schröder and Ström, 1997]. CARIBIC particle number concentrations shown in Figure 5 are at the upper end of these concentration ranges. There are three likely reasons for this difference, as follows. First, most published data were not corrected for all the particle-loss processes in the inlet system that were considered in the present study, mainly, because the respective knowledge was lacking. For CARIBIC these corrections account for an increase in N 12 and N 4 12 of 11% and 37%, respectively [Hermann et al., 2001]. Second, in contrast to most research aircraft, the CARIBIC aircraft always flies in flight corridors that particularly over Europe are strongly frequented. In such corridors particle number concentrations are estimated to be increased on average by approximately 10 20% over clean background values [Schlager et al., 1997; Hofmann et al., 1998; Anderson et al., 1999; Schröder et al., 2000]. For the North Atlantic Flight Corridor, in one case, even a much stronger increase was reported [Ferry et al., 1999]. Unfortunately, the CARIBIC instrumentation did not allow to separate out the contribution from aircraft emissions. We can only estimate that for the current flight route over Europe the influence due to aircraft is in the same order of magnitude of about 20%, for the less frequented tropical flight section presumably lower (cf. section 4.3). Finally, because we can not distinguish between measurements inside and outside clouds, artifacts caused by the break-up of cloud droplets at the inlet lips [Weber et al., 1998] cannot be totally excluded, and might be responsible for some of the high concentration values Ultrafine Particles (N 4 12 ) [24] The probability distributions of ultrafine particles in the UT for summer and winter are shown in Figure 7. Regions of particle formation in the middle and upper troposphere can be identified from these distributions. The N 4 12 summer distribution (Figure 7a) shows the same, even more pronounced features as the N 12 summer distribution. High and variable N 4 12 in the tropics in the range of particles/cm 3 were probably caused by deep convective transport in the vicinity of the ITCZ, which is associated with in situ particle formation. As the ozone mixing ratio generally increases with altitude [Seinfeld and Pandis, 1998], ozone mixing ratios in air masses pumped to higher altitudes by deep convective clouds should be lower

9 HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS AAC 10-9 Figure 7. winter. Probability distribution of tropospheric ultrafine particles (N 4 12 ) in (a) summer and (b) than ozone background values at the same altitude. If particles are formed by these clouds, particle concentrations should show a negative correlation with ozone. Figure 8 presents an example of a strong particle formation event over the Arabian Sea from a CARIBIC flight in summer Particle number concentrations (N 12 black area, N 4 12 gray area) show a strong negative correlation with the ozone mixing ratio (solid line), which is emphasized by the dashed lines for some peaks. The correlation coefficient for that period is This suggests that in summer deep convective clouds are responsible for the observed high particle concentrations over the Arabian Sea shown in Figure 7. For the present flight, these clouds were located over the Bay of Bengal, as indicated by satellite pictures, and air masses were transported to the measuring site by easterly winds, as indicated by trajectories. [25] In the subtropics, more than two-thirds of all N 4 12 measurements were below 200 particles/cm 3, indicating less particle formation activity than in the tropics. Generally, N 4 12 values of 50 particles/cm 3 or less might arise from statistical fluctuations in the particle concentrations measured by the two CPCs. Furthermore, it should be noted that N 4 12 values below about 100 particles/cm 3 do not necessarily indicate in situ particle formation. The rapid heating of the aerosol from ambient temperatures of approximately 55 C to +25 C in the container occurring during transport

10 AAC HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS Figure 8. Relation between particles and ozone in cloud pumped air. Shown are particle number concentrations (N 12 = black area, N 4 12 = gray area) and the ozone mixing ratio (solid line) from the CARIBIC flight Dusseldorf-Male on 14 June Ozone (1 s moving average) was measured using a fast chemiluminescence instrument [cf. Brenninkmeijer et al., 1999]. Vertical dashed lines are shown to emphasize the strong negative correlation between particle concentrations and the ozone mixing ratio. through the sampling line leads to a significant evaporation of water from the particles. This shifts the ambient particle size distribution toward smaller particle diameters. For particles larger than about 4 nm, however, the amount of sulfuric acid is almost preserved [Hermann et al., 2001]. Therefore, the CARIBIC results must be regarded as measurements of a dry sulfuric acid aerosol. Water evaporation causes, however, that measured ultrafine particles (for CARIBIC particles with 4 12 nm diameter) were not necessarily freshly nucleated. Likewise, they might have been larger particles that shrank into the ultrafine particle size window during sampling due to water evaporation. These particles might initially have been as large as 20 nm, depending on chemical composition and relative humidity [Hermann et al., 2001]. After passing through the sampling line, they can become sufficiently small that they are only registered by the CPC with the lowest threshold diameter. However, if the measured N 4 12 is much higher than N 12, the maximum of the ambient particle size distribution must have already been located at particle diameters smaller than 15 nm before the evaporation process, and hence in situ particle formation is more likely. [26] As shown in Figure 7a, in summer, N 4 12 again became more variable toward midlatitudes and increased to particles/cm 3. Several processes were identified to contribute to ultrafine particle concentrations in the midlatitude UT, such as aircraft emissions [Anderson et al., 1999; Schröder et al., 2000], vertical transport in convective cloud systems [Perry and Hobbs, 1994; Ström et al., 1999], lightning events [Yu and Turco, 2001; Huntrieser et al., 2002], particle formation related to frontal systems [de Reus et al., 2000; Weber et al., 2001], mixing of tropospheric and stratospheric air masses [de Reus et al., 1998; Zahn et al., 2000], and atmospheric waves, such as lee and mountain waves [Bigg, 1997; Nilsson et al., 2000]. The importance of the individual processes, however, is not known. [27] Compared to N 12, the differences in the regional structure of N 4 12 between summer and winter are stronger. In the winter, N 4 12 in the tropics typically was below 750 particles/cm 3, and about 20% of the measurements did not show ultrafine particles at all. In the subtropics, N 4 12 mostly was below 250 particles/cm 3. Mid-latitudinal N 4 12 values typically were higher and more variable again, most frequently in the range of particles/cm 3. [28] Ultrafine particle number concentrations for the midlatitudinal UT were reported by Schröder and Ström [1997] and de Reus et al. [1998] (STREAM campaigns) and range from particles/cm 3. Ultrafine particle number concentrations presented in this study are up to a factor of two higher, which might partly reflect differences in the measuring site, i.e., over the continent (CARIBIC) and over the sea or coastal regions (STREAM). Measurements made in the tropical middle troposphere in winter by

11 HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS AAC Figure 9. Probability distribution for the percentage of ultrafine particles (N 4 12 ) versus local time of day for tropical latitudes between 5 and 20 N over the Arabian Sea in (a) summer and (b) winter. Gray areas mark periods where either no data or data from only a single flight were available. de Reus et al. [2001] and in summer by Clarke [1993], however, agree well with our measurements, or even show larger values. [29] Because vertical transport and photochemistry are processes that promote the formation of particles in the free troposphere, a correlation between the existence of ultrafine particles and the local time of day can be expected. Figure 9 shows this dependency for flight sections over the Arabian Sea (between 5 N and 20 N). The analysis is restricted to this tropical region, because particle formation is less frequent in the subtropics, and because at midlatitudes the influence of the heavier air traffic on the ultrafine particle number concentration can not be excluded (cf. section 4.3). Figure 9 shows probability distributions for ultrafine particles, given as a percentage of all particles, i.e., N 4 12 divided by the sum of N 4 12 and N 12. Gray areas indicate time bins where either no data or data from only one flight were available. In summer (Figure 9a), the percentage of ultrafine particles shows a clear diurnal variation. At night, only 20% of all measurements had more than 25% ultrafine particles. During the day, this probability increased to a maximum of 60% at noon local time. This indicates that during daytime particle formation occurred close to the sampling point in the tropical middle troposphere in the vicinity of the ITCZ. The connection between particle formation and daylight hours suggests photochemistry to be a driving force for particle formation in the free troposphere, in agreement with results obtained in case studies at lower altitudes by Clarke et al. [1998] and Weber et al. [2001]. During the day, solar radiation also supports convective transport by heating the ascending air. [30] In winter (Figure 9b), the percentage of ultrafine particles is generally lower and the diurnal variations are smaller. Measurements with more than 25% ultrafine particles represent only between 13% and 20% of all measurements, suggesting that nearly independently of daytime aged aerosols were measured. Interpreting these numbers one has to be careful since they are based on poor statistics (cf. section 3.2). This can be seen in the winter graph for the UTC bin by the nonnegligible probabilities for measurements with more than 90% ultrafine particles. Nevertheless, the difference between summer and winter is so strong that we believe it is significant. The likely explanation for the observed seasonal difference is again that in winter local convection over the Arabian Sea is less frequent, as indicated by Meteosat-5 satellite images. Instead, the measured aerosols seem to be formed in the vicinity of clouds located further south, as shown by the satellite images, and to be transported to the northern measuring site, as indicated by trajectories. Hence, deep convection and the associated vertical transport of precursor gases as well as the associated cloud processing seem to be the prime requisites for particle formation in the tropical middle troposphere. Because the reduction of solar radiation in winter is not strong at tropical latitudes, the difference in the available solar radiation between summer and winter is less likely to cause the observed difference in the occurrence of ultrafine particles Stratospheric Particle Distributions Aitken Plus Accumulation Mode Particles (N 12 ) [31] The probability distribution of stratospheric N 12 particles during winter is shown in Figure 10a. The histogram in the top graph indicates the frequency of stratospheric measurements given as percentage of all measurements within a particular latitudinal band. South of 29 N stratospheric encounters were too rare to yield significant results. In winter, stratospheric N 12 decreased slightly with increasing latitude and became more variable, similar to the tropospheric winter data. Concentration values in the LS, however, are on average lower by 25% compared to UT values. The percentage of stratospheric measurements shows in winter a local maximum near 33 N and continuously increased north of 40 N. Vertical PV cross-sections of the atmosphere along the flight track show that over Arabia between 29 N and 37 N, a depression of the tropopause is frequently found, which is likely caused by orography and which leads to the observed maximum (cf. Figure 3). [32] The probability distribution of stratospheric N 12 particles in summer is shown in Figure 10b. Because in summer the subtropical and midlatitudinal tropopause is at higher altitudes, only a few stratospheric measurements south of 39 N were available and data in these latitudinal bands had to be omitted (indicated by the gray area). In contrast to winter N 12, measurements of N 12 in summer were on the same order of magnitude as those in the

12 AAC HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS Figure 10. Probability distribution of stratospheric N 12 in (a) winter and (b) summer. Colors indicate the percentage of all particle measurements in a latitudinal band that fall into a particular concentration bin. Histograms show the percentage of stratospheric measurements within a particular latitudinal band. The gray area marks latitudinal bands where the statistical requirement of a minimum of 1100 data points from at least five flights were not fulfilled. troposphere. This seasonality can be likely explained by stronger in-mixing of tropospheric air into the lowermost stratosphere in summer, whereas in winter tropospherestratosphere-exchange is reduced due to the strong potential vorticity gradient along potential temperature surfaces [Haynes and Shuckburgh, 2000]. [33] Number concentrations in the LS were reported in the range of particles/cm 3 by Wilson et al. [1991], Hofmann [1993], Brock et al. [1995], and Schröder and Ström [1997]. For summer, CARIBIC data were higher again, whereas for winter data fell into the reported range. [34] To illustrate the variation of the LS particle number concentration with increasing altitude, Figure 11 shows scatterplots of N 12 versus in situ measured ozone for all data points with ozone mixing ratios larger than 65 ppb. Only data points north of 25 N were considered. Because ozone shows a steep vertical gradient above the tropopause, ozone mixing ratios can be used as vertical coordinate in the LS. To obtain an approximate measure for corresponding altitudes, an ozone climatology provided by the German weather service (DWD, Hohenpeissenberg, 48 N) was used (dashed lines in Figure 11). This ozone altitude function

13 HERMANN ET AL.: DISTRIBUTIONS OF AEROSOL CONCENTRATIONS AAC Figure 11. Scatterplot of N 12 versus in situ measured ozone for (a) December January and (b) June July. The tropopause and approximate altitudes above the tropopause (dashed lines) were calculated using an ozone climatology for Hohenpeissenberg, Germany, 48 N, which was kindly provided by H. Claude, DWD, Hohenpeissenberg. varies strongly with season, particularly in summer and winter, wherefore data were restricted to two-month intervals of December January and June July, respectively. [35] In winter (Figure 11a), N 12 decreased clearly with increasing altitude above the tropopause, in agreement with previous measurements [Wilson et al., 1991; Brock et al., 1995]. This decrease is again likely caused by the reduced transport of UT air into the LS [Haynes and Shuckburgh, 2000]. In contrast to winter N 12, the summer N 12 decreased only slightly within the first few kilometers above the tropopause (Figure 11b). This reflects again the stronger mixing of particles and precursor gases from the UT into the LS in summer. A similar behavior was observed by Hofmann [1993] Ultrafine Particles (N 4 12 ) [36] The distributions of stratospheric N 4 12 for winter and summer are displayed in Figure 12. In winter (Figure 12a), stratospheric N 4 12 in the first few kilometers above the tropopause were again lower compared to the tropospheric N 4 12, whereas in summer stratospheric N 4 12 were either the same or higher. This seasonality is likely caused by the increased in-mixing of tropospheric air into the LS in summer and associated particle formation processes, as already shown by Zahn et al. [2000] using CARIBIC data. Convective lifting of the tropopause in summer can also initiate particle formation in the LS, as shown by de Reus et al. [1999]. The majority of CARIBIC N 4 12 values for the LS were in the range of particles/cm 3, higher than those reported by Schröder and Ström [1997] and de Reus et al. [1998], who found only between ultrafine particles/cm 3. However, de Reus et al. [1999] reported ultrafine particle number concentrations of up to 10,000 particles/cm 3 in the lowermost stratosphere in connection with convective lifting of the tropopause Autocorrelation Analysis and Integral Length Scales [37] Integral length scales for the CARIBIC particle data were calculated by using autocorrelation analysis (section 3.2). The calculated mean lengths scales, shown in Table 1, ranged from 6 to 12 km, but within their standard deviations no significant difference can be inferred. This implies that independent of particle size (ultrafine or Aitken plus accumulation mode), season (summer or winter), atmospheric domain (UT or LS), and geographical region (tropics, subtropics, or midlatitudes), the main processes controlling the number concentration of submicrometer particles seem to act on similar scales. It should be noted, however, that integral length scales give only information on the spatial range over which variations in the data occur, but they do not indicate the magnitude of these variations. For example, although the tropospheric N 12 summer dis-