Airborne measurements of aerosol, ozone, and ... - Wiley Online Library

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 110, D09202, doi:10.1029/2004JD005397, 2005

Airborne measurements of aerosol, ozone, and solar ultraviolet irradiance in the troposphere Costas Varotsos1 Department of Applied Physics, University of Athens, Athens, Greece Received 29 August 2004; revised 18 January 2005; accepted 15 February 2005; published 7 May 2005.

[1] Measurements of the distributions of the aerosol characteristics and solar ultraviolet

irradiance were conducted by using instrumentation flown on a Falcon aircraft over the entire Greek area from the sea up to the middle troposphere. To study the impact of the aerosol abundance to the solar ultraviolet irradiance at various altitudes, complementary observations of ozone concentration, relative humidity, and temperature have been also performed using instrumentation flown on free balloons. In addition to a detailed description of the observed aerosol characteristics we attempt an improvement of the earlier proposed theoretical algorithms regarding the absolute values of the solar ultraviolet irradiance at various tropospheric altitudes. This is achieved by taking into account the measured accumulation aerosol particle and cloud droplet size distributions. The measurements showed that a double-layer structure in the vertical distribution of the number density of the cloud droplets and particle size aerosol was prevalent. In particular, over the area west of Crete the aerosol particles possessed a bimodal size distribution with mean diameters of 0.11 and 0.225 mm. Moreover, upon using the obtained aerosol distribution data in the radiative transfer calculations, we find that the calculated values of the solar ultraviolet irradiance correlate well with the observed ones, showing an 4.3 ± 0.1% km1 increase for altitudes ranging from the ground to 6.2 km. Citation: Varotsos, C. (2005), Airborne measurements of aerosol, ozone, and solar ultraviolet irradiance in the troposphere, J. Geophys. Res., 110, D09202, doi:10.1029/2004JD005397.

1. Introduction [2] It is a truism that Earth’s atmosphere as a component of the climatic system behaves as a colloidal medium with the clouds and aerosols as main suspended particles. Among the main sources of aerosols are the wind-driven erosion, human activities, biomass combustion, gas to particle conversion, and volcanoes [Cracknell, 2001; Dey et al., 2004]. As an example, it is well established that dust storms affecting the Greek territory primarily originate from the central Sahara in spring, the eastern Sahara in summer, and the Middle East/Arabian Peninsula in autumn. Furthermore, it has been observed that urban industrial aerosols from the north over the Balkan region, Ukraine, and Anatolia are transported to the Greek territory at low altitudes whereas summer and autumn dust intrusions usually occur at higher altitudes (above 700 hPa) [Kubilay et al., 2003]. [3] In the framework of the experimental campaign of the radiation field in the troposphere (RAFT), observations of the vertical distribution of the aerosol number density and size, erythemal ultraviolet irradiance (EUVI), ozone mixing ratio, relative humidity, and temperature have been carried 1 Visiting professor at Department of Meteorology, University of Maryland, College Park, Maryland, USA.

Copyright 2005 by the American Geophysical Union. 0148-0227/05/2004JD005397

out in Greece during 7 – 14 June 1997. RAFT was a subproject of the scientific training and access to aircraft for atmospheric research throughout Europe (STAAARTE) project. Within the STAAARTE-RAFT campaign the specially equipped Falcon 20-E5 D-CMET research aircraft of the German Aerospace Research Establishment (Deutsches Zentrum fur Luft und Raumfahrt (DLR)) carried the experimental instrumentation approximately over most of the part of the Greek area from the sea level up to 6.2 km. The measurements have been conducted during the aforementioned 7 – 14 June period because weak synoptic flow over the greater areas of Athens and Thessaloniki and the area west of Crete often prevails during June, and therefore high pollution levels are usually observed. [4] The EUVI observations obtained during STAAARTERAFT have been analyzed by Varotsos et al. [2001]. They reported an average increase of 7.2 ± 1.2% km1 in EUVI with altitude throughout the troposphere. This figure was found to agree fairly well with the theoretical estimations, although the absolute values of the experimentally and theoretically derived EUVI (at the same altitude) exhibit large deviations. The latter could be attributed to the following fact: In the theoretical estimation of EUVI, typical values for the aerosol characteristics have been considered since real aerosol measurements were not available at that time. [5] The purpose of the present paper is to present detailed measurements over almost the entire Greek area of the

D09202

1 of 10

D09202

VAROTSOS: AEROSOL MEASUREMENTS OVER GREECE

vertical distribution of the aerosol number density and size as a function of height in conjunction with the corresponding distributions of EUVI, ozone mixing ratio, relative humidity, and temperature.

2. Instrumentation [ 6 ] The DLR research aircraft, as mentioned, was specially equipped with the necessary standard atmospheric instrumentation [Fimpel, 1991], which will be shortly presented below. The EUVI detector of Athens University was mounted on the top of the aircraft with an inclination of 6 relative to the horizon in order to compensate for the aircraft angle of attack. In order to make the data comparable, independently of the time of the day in which the observations were carried out, the data obtained have been corrected in accordance to the solar zenith angle, considering perpendicular incidence of the solar beam on the instrument. [7] Ozone mixing ratio was measured on board by using a well-calibrated chemiluminescence analyzer with an ozone detection limit of 2 ppb and accuracy of 5% (95%). For the sake of comparison, simultaneous ozone, relative humidity, and temperature profile observations were also carried out by means of free balloon ascents. Details concerning the operation principle and the specific characteristics of the EUVI, ozone, relative humidity, and temperature sensors are given by Varotsos et al. [2001]. [8] The aerosol measurements with a recording frequency set to 1 s integrating interval have been performed by using the spectrometers Forward Scattering Spectrometer Probe (FSSP)-300 and passive cavity aerosol spectrometer probe (PCASP)-100X mounted under the wings of the DLR research aircraft. [9] The FSSP-300, built by Particle Measurements Systems (Boulder, Colorado), samples particles in the free stream so that they are not altered in size or shape by an inlet system. Individual particles are detected and their size is determined on the basis of the intensity of light that they scatter as they pass through the focused beam of a helium-neon laser [Hallar et al., 2004]. The size of the particle is determined using Mie theory, assuming that the particles are spherical with a known refractive index. The measurements of the particle size distributions conducted by FSSP-300 are expressed as the number density of cloud droplet size (CS) aerosols with mean bin diameters 3.5, 6.5, 9.5, 12.5, 15.5, 18.5, 21.5, 24.5, 27.5, 30.5, 33.5, 36.5, 39.5, 42.5, and 45.5 mm [Strapp et al., 1992; Cutten et al., 1996]. [10] The PCASP-100X provides measurements of the accumulation mode particles, expressed as the number density of particle size (PS) aerosols with mean bin diameters 0.11, 0.13, 0.155, 0.185, 0.225, 0.275, 0.35, 0.45, 0.6, 0.8, 1.05, 1.35, 1.75, 2.25, and 2.75 mm. Aerosol sample flow rates in PCASP were recorded and used in the data processing to calculate ambient concentrations. Furthermore, it is worth noting that the PCASP measurements are more reliable at relative humidities low enough to ignore size differences resulting from water associated with the sampled aerosol [Guibert et al., 2003]. Also, Haywood et al. [2003] showed that the PCASP-100X sizing of supermicron particles becomes more uncertain because of differ-

D09202

ent scattering responses to ambient aerosol and calibration sphere refractive indices.

3. Results and Discussion: Aircraft and Free Balloon Observations [11] In this section we present the combined measurements of the vertical profiles of aerosols (size and number density), EUVI, and the other atmospheric parameters collected during the intensive campaign of the STAAARTE-RAFT. We first start from the days dedicated to the experiments around the greater Athens area (GAA); next, we turn to the greater Thessaloniki area (GTA); and finally, we turn to the area west of Crete (AWC). The most characteristic feature of the synoptic situation that prevailed over the three experimental sites was typically cloudless sky, thus avoiding interference with EUVI by clouds. 3.1. Description of the Measurements Obtained Over the Greater Athens Area [12] The vertical distributions of ozone partial pressure and air temperature deduced from the free balloon ascents over the GAA on 13 June 1997 are depicted in Figure 1. They resemble typical profiles that are usually observed in this region under the aforementioned specific weather conditions. Specifically, the ozone partial pressure was 70 mPa near the surface while it reaches 30 hPa at the pressure level of 450 hPa (7 km height) following the conventional gradual decrease versus height. Concerning the relative humidity observations, the noticeable point is a maximum extended in the altitude range 2 – 3.7 km, which is further discussed later in this section. [13] On 13 June 1997 the DLR aircraft was also flown over the GAA from the surface up to 6.2 km, performing measurements (from 1107 to 1226 UTC) of the vertical profiles in EUVI, aerosols (size and number density of CS and PS), ozone mixing ratio, and other meteorological variables. [14] Figure 2a illustrates the vertical distribution of the CS aerosol number density for four selected diameter intervals (with mean bin diameters 3.5, 6.5, 9.5, and 12.5 mm) as deduced from FSSP-300 in situ measurements. Figure 2a shows that in the range 3.5 –12.5 mm the CS aerosol number density exhibits two main maxima (each 1 km deep) at altitudes 2.2 and 3.2 km, which indicates a double-layer structure. The latter also holds in the range 3.5– 21.5 mm (not shown in Figure 2). On the contrary, as can be clearly seen in Figure 2b, the CS aerosol number density with mean bin diameters 30.5, 36.5, 39.5, and 45.5 shows one (main) maximum near the surface and a secondary one at the altitude of 2.2 km. [15] Similar to Figure 2b, main and secondary maxima are also observed (Figure 3a) in the vertical distribution of the PS aerosol number density for mean bin diameters 0.11, 0.185, and 0.225 mm obtained from the observations of the PCASP instrument. [16] In view of the fact there are well-defined common features in Figures 2a, 2b, and 3a, it is of interest to plot (Figure 3b) the vertical distribution of the total aerosol number density measured by each FSSP and PCASP instrument along with the measured vertical EUVI distribution. The conclusion drawn from Figure 3b is that an

2 of 10

D09202


D09202

Figure 1. Observations of the ozone vertical distribution obtained from free balloon ascents at the Athens ozone station on 13 June 1997 (solid line indicates ozone partial pressure, dashed line indicates air temperature, and dash-dotted line indicates relative humidity).

intermediate layer of aerosol particles exists, which separates the above-mentioned two distinct layers of aerosol cloud droplets in a sandwich scheme (‘‘aerosol sandwich’’). It seems highly probable that this aerosol sandwich structure produces the scale-shaped structure in the EUVI profile shown in Figure 3b. Furthermore, if we recall the vertical distribution of the relative humidity shown in Figure 1, we notice that this aerosol sandwich coexists with the maximum of relative humidity (75%) located in the altitude

range 2 –4 km. A plausible explanation for the latter pattern could be the growth of hygroscopic particles at that atmospheric region. [17] Figure 4 illustrates the combined in situ measurements of the aerosol mean diameter, the aerosol total number density, the ozone mixing ratio, and the EUVI as a function of height. An inspection of Figure 4 shows that both the aerosol mean diameter and the ozone mixing ratio decrease with altitude, thus contributing to the EUVI

3 of 10

D09202


D09202

Figure 2. Distribution of cloud droplet size (CS) aerosol number density for the following mean bin diameters: (a) 3.5 mm (dashed line), 6.5 mm (solid line with crosses), 9.5 mm (solid line with circles), and 12.5 mm (solid line) and (b) 30.5 mm (dashed line), 36.5 mm (solid line with crosses), 39.5 mm (solid line with circles), and 45.5 mm (solid line). increase with altitude. Furthermore, the observed vertical EUVI decrease, from 875 hPa toward the surface, seems to become much more intense below the starting point of the increase in the aerosol number density (at 875 hPa). An additional interesting remark about Figure 4 is the following: The observed ozone depletion is located at 675 hPa and hence at the upper part of the aerosol sandwich shown in Figure 3. This probably indicates the ozone uptake by the aerosol particles and cloud droplets. In this respect, Ghosh and Varotsos [1999] have found that in the lower troposphere the ozone uptake rates are high because of a preponderance of large cloud droplets, which is reflected in the ozone concentration decrease in the region over Athens, Greece (see, e.g., the ozone depletion at 700 hPa in Figure 4). They have also suggested that the diffusional contribution to the ozone uptake rate increases upon increasing the droplet size. [18] On the previous day, i.e., 12 June 1997, the aircraft was flown over the GAA. The measurements have been taken in the time interval from 1012 to 1027 UTC during the aircraft’s descent from the altitude of 6.3– 0.2 km. [19] The aerosol data collected during this flight are shown in Figure 5. In particular, Figure 5a depicts the vertical distribution of the PS aerosol number density with

mean bin diameters 0.11, 0.185, and 0.225 mm obtained from the observations of PCASP instrument on 12 June 1997. [20] Figure 5a clearly shows that the vertical distribution of the PS aerosol number density is unimodal. In addition, the gradual decrease of the PS aerosol number density with altitude is the common feature in the three diameter intervals selected. The regions with enhanced aerosol number density are more evident in Figure 5b, where the vertical distribution of the total aerosol number density measured by each FSSP and PCASP instrument are shown along with the observed vertical distribution of EUVI. Two pronounced maxima of the total CS aerosol number density at 0.7 and 3.2 km are evident at the vertical distribution of FSSP observations: The first one at 0.7 km coincides with that of the total PS aerosol number density, which both contribute to the observed anomaly in EUVI profile (Figure 5b). The second maximum in the profile of the CS total aerosol number density (at 3.2 km) is accompanied by the very weak maximum of the PS total aerosol number density at 2.5 km. [21] It is of great importance to mention that when measuring the aerosol size distribution using the PCASP and employing Mie theory combined with suitable refrac-

4 of 10

D09202


D09202

Figure 3. (a) Distribution of the particle size (PS) aerosol number density for the following mean bin diameters: (a) 0.11 mm (dashed line), 0.185 mm (solid line with crosses), and 0.225 mm (solid line with circles) as derived from passive cavity aerosol spectrometer probe (PCASP) on 13 June 1997. (b) Distribution of erythemal ultraviolet irradiance (EUVI) (solid line with circles), the CS aerosol total number density from Forward Scattering Spectrometer Probe (FSSP) (solid line), and the PS aerosol total number density from PCASP (solid line with crosses) on 13 June 1997. tive indices (to determine the aerosol extinction) and then integrating the measurements over height during the profile descent and profile ascent, the aerosol optical depth can be derived [Haywood et al., 2003]. However, one should consider that the PCASP number concentration is significantly affected by small changes (e.g., 2.5) in the pitch of the aircraft. Finally, although the PCASP may not accurately count the total number of aerosol particles, the size distribution is well represented [Haywood et al., 2003]. 3.2. Description of the Measurements Obtained Over North Greece [22] The observations during the experimental day of 9 June 1997 were obtained when the aircraft was flown at the altitude range between 0.12 and 4.2 km over the GTA, which is the second largest Greek city (with 1 million inhabitants), being nowadays one of the largest urban agglomerations in the Balkans [Gusten et al., 1997]. [23] The vertical distribution of the PS aerosol number density for the three selected diameter intervals (with mean diameters 0.11, 0.185, and 0.225 mm) observed on 12 June 1997 are shown in Figure 6a, where a prominent maximum in the boundary layer is depicted. The total PS aerosol number density in Figure 6b is very high compared to the total PS aerosol number density measured at the GAA

(Figure 3b). This may be attributed to the fact that the GTA is mostly characterized by intense industrial activity. In this respect, Gerasopoulos et al. [2003] have experimentally found that the local pollution contributes as much as 35 ± 10% to the average aerosol optical depth at the Thessaloniki site during summer. Furthermore, Formenti et al. [2002] have measured near Thessaloniki city aerosol sulfate levels up to 500 ppt. Of specific interest is also the observed profile of EUVI into and above the PC and CS layers, especially if it is compared with the corresponding profiles in the other experimental days. [24] It should be clarified that there are two kinds of uncertainties in the FSSP-300 measurements, i.e., concentration uncertainties and sizing errors, which both affect the calculation of total condensed material. In particular, the concentration uncertainties are 25%, while the various contributions to sizing errors amount to 60% uncertainty in mass concentration. Furthermore, the asphericity of ice particles adds a further complication, as does the potential for large ice crystals to shatter on the FSSP inlet, although recent studies suggest that this causes an overestimation of number concentration by about a factor of 2 or less [Hallar et al., 2004]. In this respect, Min et al. [2003] proposed that correction algorithms of probe-dependent and distributiondependent optical coincidence effects, Mie curve adjust-

5 of 10

D09202


D09202

Figure 4. Distribution of EUVI (thick solid line), the aerosol total number density (dashed line), the ozone partial pressure (thin solid line), and the aerosol mean diameter (solid line with crosses) on 13 June 1997. ment, and time response and laser beam inhomogeneity effects have to be applied to the raw FSSP data in order to obtain results with an accuracy of 15% in the cloud drop radius and 34% in the liquid water content. In this regard, Guibert et al. [2003] documented a substantial disparity in the comparison of predicted and observed wet aerosol spectra. This was attributed either to the bias in the wet aerosol measurements made with an FSSP or to the error in the sizing of dry aerosol particles. [25] Concerning the PCASP reliability, the aerosol size distributions measured by the PCASP and derived from the Aerosol Robotic Network site were found to be in excellent agreement over the 0.05 – 1.0 mm radius range, which contains the majority of the optically active particles [Haywood et al., 2003]. 3.3. Description of the Measurements Obtained Over West of Crete [26] On the first experimental day, 7 June 1997, the measurements were carried out during the aircraft flight over the Aegean Sea (west of Crete) in the latitude of 35.4N and in the time interval between 1002 and 1203 UTC. Crete represents a rather clean area, a considerable distance downwind of continental pollution sources [Formenti et al., 2002]. [27] The analysis of the obtained observations refers to the altitude range between 0.12 and 6.2 km, and the results are illustrated in Figures 7a and 7b. Focusing first on Figure 7a, we notice that the observed vertical distribution of the PS aerosol number density in the lower troposphere for the three selected diameter intervals (with mean diameters 0.11, 0.185, and 0.225 mm) is characterized by two prominent maxima. Interestingly, the second maximum in

the distribution of PS with mean diameter 0.11 mm is shifted to higher altitudes compared to the corresponding maxima of 0.185 and 0.225 mm (bimodality effect). The second mode was more pronounced in areas of higher relative humidity, thus indicating the presence of deliquescent aerosols, but also in areas where high ozone concentrations were measured. In Figure 7b it is evident that the total PS aerosol number density agrees fairly well with the total CS aerosol number density throughout the lower and middle troposphere (Figure 6b). The results depicted in Figures 7a and 7b are consistent with the fact that in the eastern Mediterranean, in summertime, sulfate aerosols dominate the aerosol population because of low precipitation and efficient photochemistry [Mihalopoulos et al., 1997; Formenti et al., 2001]. Furthermore, it is of pertinent interest to annotate the changes in the EUVI gradient, especially into the first kilometer of the observations. Since aerosol absorption is important, we expect significant departures from estimates of clear-sky radiative forcing when considering sulfate aerosols only.

4. Association Between the Observed and Modeled Solar Ultraviolet Irradiance [28] Varotsos [1994] proposed a theoretical algorithm for the calculation of the broadband and spectral EUVI reaching Earth’s surface. This model has been already validated by using many experimental observations of EUVI [Ziemke et al., 2000; Varotsos et al., 2001]. [29] It is therefore of pertinent interest to employ this model to estimate the EUVI at various altitudes under the conditions described in section 3 [Piazena, 1996]. We first briefly describe the model.

6 of 10

D09202


D09202

Figure 5. (a) Distribution of the PS aerosol number density for the following mean bin diameters: 0.11 mm (dashed line), 0.185 mm (solid line with crosses), and 0.225 mm (solid line with circles) as derived from PCASP on 12 June 1997. (b) Distribution of EUVI (solid line with circles), the CS aerosol total number density (solid line), and the PS aerosol total number density (solid line with crosses) on 12 June 1997. [30] The calculation of the direct EUVI reaching a horizontal surface is performed by using the relationship [Bird and Riordan, 1986] Id ðlÞ ¼ Ho ðlÞDtr ðlÞta ðlÞtw ðlÞto ðlÞtu ðlÞ;

Rayleigh scattering [Bird and Riordan, 1986] ð2Þ

Aerosol attenuation [Bird and Riordan, 1986] ta ðlÞ ¼ expðbla M Þ;

n o tw ðlÞ ¼ exp 0:2385aw ðlÞWM=½1 þ 20:07aw ðlÞWM 0:45 ; ð4Þ

ð1Þ

where Id(l) is the direct normal irradiance at the horizontal surface, Ho(l) is the extraterrestrial solar irradiance at the mean Earth-Sun distance at a given wavelength l [Neckel and Labs, 1981], and D is the correction factor arising from the change of the Earth-Sun distance throughout the year. The parameters tr(l), ta(l), tw(l), to(l), and tu(l) stand for the following transmission functions of the atmosphere at the wavelength l: molecular (Rayleigh) scattering, aerosol attenuation (scattering and absorption), water vapor absorption, ozone absorption, and uniformly mixed minor gas (CO2, CO, CH4, N2O, and O2) absorption, respectively. [31] In particular, the transmission functions used in formula (1) are as follows:

tr ðlÞ ¼ exp M 0 = l4 115:6406 1:335=l2 ;

Water vapor [Leckner, 1978]

ð3Þ

Ozone absorption [Leckner, 1978] to ðlÞ ¼ exp½ao ðlÞO3 Mo ;

ð5Þ

Uniformly mixed gases [Leckner, 1978] n o 0:45 ; tu ðlÞ ¼ exp 1:41au ðlÞM 0 =½1 þ 118:3au ðlÞM 0

ð6Þ

where M is the relative air mass; M0 is the pressure-corrected relative air mass; Mo is the relative ozone atmospheric mass; ˚ ngstrom’s turbidity coefficient; a is the wavelength b is the A exponent; W is the precipitable water (cm); O3 is the total ozone amount (atm cm) (1 atm = 105 1.01325 N/m2); and aw(l), ao(l), and au(l) are the absorption coefficients for water vapor, ozone, and uniformly mixed gases, respectively. ˚ ngstrom’s [32] On the basis of previous observations the A turbidity coefficient was taken equal to b = 0.71 for June [Varotsos, 1994]. Additionally, according to Bird and Riordan [1986] the wavelength exponent a in (3) is a =

7 of 10

D09202


D09202

Figure 6. (a) Distribution of the PS aerosol number density for the following mean bin diameters: 0.11 mm (dashed line), 0.185 mm (solid line with crosses), and 0.225 mm (solid line with circles) as derived from PCASP on 9 June 1997. (b) Distribution of EUVI (solid line with circles), the CS aerosol total number density (solid line), and the PS aerosol total number density (solid line with crosses) on 9 June 1997. 1.0274 if l < 0.5 mm and a = 1.206 if l 0.5 mm. Furthermore, tw(l) = 1 because of the fact that aw(l) = 0 in the spectral region 0.28– 0.75 mm. The values of ao(l) and au(l) are given by Bird and Riordan [1986] for various wavelengths. For mean ground albedo we assumed the value 0.3, while above the sea we used the value of 0.1. [33] Concerning the total diffuse irradiance on a horizontal surface at a given altitude from the Earth’s surface, it is usually described by Is ðlÞ ¼ Ir ðlÞ þ Ia ðlÞ þ Ig ðlÞ;

ð7Þ

where Is(l) is the total diffuse irradiance on a horizontal surface at a given altitude from the ground at wavelength l, Ir(l) is the Rayleigh scattering component, Ia(l) is the aerosol scattering component, and Ig(l) is the component which takes into account the multiple reflected radiation between the air and the ground. Consequently, the total solar irradiance on a horizontal surface at a given altitude from the ground depends on the apparent solar zenith angle z as follows: It ðlÞ ¼ Id ðlÞ cos z þ Is ðlÞ;

ð8Þ

where It(l) is the total solar irradiance on the horizontal surface at a wavelength l. Finally, by assuming the

erythema action spectrum given by McKinley and Diffey [1987] the calculated irradiance at various altitudes from Earth’s surface is converted into EUVI. [34] The results obtained from the application of the model described above are shown in Figure 8a. An inspection of Figure 8a shows that there is a persistent overestimation of 20% of the theoretically derived EUVI when compared to the corresponding experimental results. This might be due to the fact that typical values for the optical properties of the atmosphere have been used in the model. This is strengthened by the fact that when employing experimental values deduced from the above mentioned experimental observations, this overestimation vanishes, and the deviations between the two curves (