Measurements of Mediterranean aerosol radiative forcing and influence of the single scattering albedo

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 114,, doi: /2008jd011037, 2009 Measurements of Mediterranean aerosol radiative forcing and influence of the single scattering albedo Claudia Di Biagio, 1 Alcide di Sarra, 1 Daniela Meloni, 1 Francesco Monteleone, 2 Salvatore Piacentino, 2 and Damiano Sferlazzo 3 Received 28 August 2008; revised 5 December 2008; accepted 28 January 2009; published 27 March [1] Ground-based measurements of aerosol optical depth and surface shortwave irradiance carried out at the Mediterranean island of Lampedusa during are used to estimate the surface aerosol direct radiative forcing for desert dust (DD), urban/ industrial-biomass burning (UI-BB), and mixed aerosols (MA). The aerosol single scattering albedo, w, at and nm is derived at 60 solar zenith angle, q, from measurements of global and diffuse radiation using radiative transfer model calculations. The shortwave forcing efficiency (FE S ) is derived, for q between 20 and 75, for the three identified classes of aerosol and for all the observed data (AD). The absolute value of FE S decreases for increasing q for all the aerosol types. FE S varies between 185 and 81.7 W m 2 for DD, 168 and 84 W m 2 for UI-BB, 251 and W m 2 for MA, and 208 and W m 2 for AD. The daily average forcing efficiency (FE d ) at the equinox is 67.2 W m 2 for DD, 59.0 W m 2 for UI-BB, and 93.2 W m 2 for MA. The forcing efficiency of DD, UI-BB, and MA at q =60 was calculated for three intervals of single scattering albedo (0.7 w < 0.8, 0.8 w < 0.9, 0.9 w 1) at and nm. The absolute value of FE S decreases with increasing w at nm for all aerosol types, while it decreases with increasing w at nm for UI-BB and MA and increases for DD. A 0.1 increment in the single scattering albedo at nm produces a reduction in FE S by W m 2, and a reduction by W m 2 in FE d. Citation: Di Biagio, C., A. di Sarra, D. Meloni, F. Monteleone, S. Piacentino, and D. Sferlazzo (2009), Measurements of Mediterranean aerosol radiative forcing and influence of the single scattering albedo, J. Geophys. Res., 114,, doi: /2008jd Introduction [2] Atmospheric aerosols are among the key parameters in the climate system [Forster et al., 2007]. They influence the Earth s energy budget both directly, because they absorb and scatter shortwave and longwave radiation, and indirectly, because they act as cloud condensation nuclei modifying the structure and the properties of clouds [Twomey, 1974; Albrecht, 1989; Pincus and Baker, 1994]. Several measurement campaigns carried out all over the world and modeling analyses have been dedicated during recent years to study the aerosol radiative effects, in particular the aerosol direct radiative forcing (RF) [Yu et al., 2006]. Because of the variability of the aerosol optical properties and spatialtemporal distribution, the determination of RF, both at the regional and the global scale, is still characterized by a great uncertainty [Forster et al., 2007]. The determination of the 1 Dipartimento Ambiente, Cambiamenti Globali e Sviluppo Sostenibile, Ente per le Nuove Tecnologie, l Energia e l Ambiente, Rome, Italy. 2 Dipartimento Ambiente, Cambiamenti Globali e Sviluppo Sostenibile, Ente per le Nuove Tecnologie, l Energia e l Ambiente, Palermo, Italy. 3 Dipartimento Ambiente, Cambiamenti Globali e Sviluppo Sostenibile, Ente per le Nuove Tecnologie, l Energia e l Ambiente, Lampedusa, Italy. Copyright 2009 by the American Geophysical Union /09/2008JD aerosol absorption is one of the key issues, since it largely influences the aerosol radiative effects [e.g., Stier et al., 2007]. [3] The aerosol forcing may play a large role on regional scales, and may contribute significantly to the determination of the regional climate. The Mediterranean region is characterized by a large variability in aerosol properties: different types of aerosol, mixtures of them, and elevated amounts of particles can be found over the basin. The particles are both locally produced and transported by air masses from Europe, North Africa, and the North Atlantic Ocean [i.e., Mihalopoulos et al., 1997; Sciare et al., 2003; Pace et al., 2006; Lyamani et al., 2006]. Marine and polluted particles are most common in winter, while cases of Saharan dust and biomass burning aerosols are often found in summer [Pace et al., 2005, 2006; Meloni et al., 2007; Lyamani et al., 2006]. The seasonal distribution of the different types of particles in the Mediterranean is modulated by local, synoptic, and large-scale circulation [di Sarra et al., 2001; Kallos et al., 2007]. [4] The evolution of the aerosol chemical and optical properties was studied at different sites in the Mediterranean basin. Measurements from small islands are particularly valuable because local sources are generally small, and medium- and long-range transport, together with aerosol 1of12

2 aging, play a central role, thus providing information on aerosol properties over a relatively large scale. Regular measurements of aerosol properties [e.g., Pace et al., 2006; Meloni et al., 2006, 2007, 2008; Di Iorio et al., 2009] and radiative fluxes [e.g., di Sarra et al., 2002, 2008] are carried out at Lampedusa (35.5 N, 12.6 E), a small island (a surface area of about 22 km 2 ) in the southern sector of the central Mediterranean sea. Studies on the aerosol properties in similar conditions were carried out at Crete [e.g., Mihalopoulos et al., 1997; Sciare et al., 2003; Gerasopoulos et al., 2006], in the eastern Mediterranean basin. [5] A limited number of studies have been dedicated to the investigation of the aerosol radiative effects in the Mediterranean region in the recent years [i.e., Formenti et al., 2002; Markowicz et al., 2002; Roger et al., 2006; di Sarra et al., 2008; Saha et al., 2008]. [6] In a recent study, di Sarra et al. [2008] derived estimates of the direct surface aerosol forcing in the central Mediterranean using the direct method [Satheesh and Ramanathan, 2000]. The estimates are based on measurements of aerosol optical properties and surface shortwave fluxes carried at Lampedusa. They classified the aerosol in two different classes, namely, desert dust and biomass burning/industrial aerosols, on the basis of their optical properties using the method developed by Pace et al. [2006], and derived estimates of the forcing separately for the two classes and for the whole data set. The derived forcing efficiencies, i.e., the forcing produced by a unit optical depth at 500 nm, at the equinox were 69 W m 2 for desert dust, and 54 W m 2 for urban/industrial particles. Given the large variability in Mediterranean aerosol properties and the limited number of studies on this topic in the Mediterranean, there is the need to assess the robustness of these results. [7] Kim et al. [2005] and Ramana and Ramanathan [2006] have shown that the aerosol absorption significantly influences the surface direct forcing and that the dependence on the aerosol single scattering albedo (the ratio between the aerosol scattering and extinction) may be identified using the direct method. Several studies have shown that Mediterranean aerosols display a large variability in the aerosol single scattering albedo [e.g., Bryant et al., 2006; Meloni et al., 2006; Saha et al., 2008]. Thus, the investigation of the role played by the aerosol absorption on the forcing is crucial for a better understanding of the Mediterranean regional climate. [8] In this study the analysis made by di Sarra et al. [2008] is improved and expanded to provide more robust information on the forcing produced by different types of aerosols in the Mediterranean, and to investigate the effects produced by changes in the aerosol single scattering albedo. The analysis by di Sarra et al. is here applied to a different and longer time interval ( ), and improvements in the methodology are implemented. The improvements include a better characterization of the shortwave irradiance measurements, by adding corrections for the thermal offset and the cosine response of the pyranometers, and an extended classification of the aerosol types. The dependence of the forcing on the aerosol single scattering albedo is studied using observations and estimates of the single scattering albedo derived from measurements. To our knowledge, this is the first analysis of this kind in the Mediterranean. 2. Instrumentation [9] This study is based on measurements obtained at Lampedusa during the period June 2004 to August 2007 with a Multifilter Rotating Shadowband Radiometer (MFRSR), three Eppley Precision Spectral Pyranometers (PSP) and two Eppley Precision Infrared Radiometers (PIR). The instrumentation is installed at the Ente per le Nuove Tecnologie l Energia e l Ambiente, Agrigento, Italy (ENEA), Station for Climate Observations on the northeastern coast of Lampedusa. The instruments are placed on the roof of the station and have the horizon free of significant obstacles. [10] The MFRSR [Harrison et al., 1994] measures global and diffuse irradiances at six 10 nm wide channels centered at 415.6, 495.7, 614.6, 672.6, 868.7, and nm and at one broadband channel ( nm). The data are acquired every 15 s and averaged over 1 min. For the different channels, the direct irradiances are calculated as the difference between global and diffuse irradiances and are used to derive, by applying the Beer-Lambert law, the total atmospheric optical depth. The MFRSR calibration and aerosol optical depth derivation is described by Pace et al. [2006]. For the different channels, also the diffuse to direct irradiance ratio (DDR) is calculated. [11] The PSP measures global downward irradiance (I SW# ) with an uniform spectral responsivity throughout the region mm; the data are acquired every 30 s. During the period June 2004 to August 2007, three different PSPs (s/n 33504, 33600, and 34891) have been used. The three PSPs are all ventilated and are run, in certain periods, side by side. PSPs and were calibrated at Eppley in August PSP was recalibrated at the World Radiation Centre (WRC) at Davos (Switzerland) in September The PSP was calibrated at Eppley in February The sensitivity of PSP changed by less than 6% between 2002 and The comparison with PSP shows that the PSP sensitivity changed by less than 1% between 2005 and The calibration made in 2005 is then applied throughout the investigation period, in addition to instrument-to-instrument corrections dependent on solar zenith and azimuth angles (section 3.1). [12] The PIR measures global downward irradiance (I IR# ) with an uniform spectral responsivity throughout the region mm. The data are acquired every 30 s. During the period June 2004 to August 2007 two different PIRs (s/n and 33500) were installed at Lampedusa. The two PIRs were both calibrated at Eppley in PIR was calibrated at WRC again in August [13] The ENEA station is also equipped with an automatic weather station that performs continuous measurements of standard meteorological parameters. 3. Data Analysis 3.1. PSP Measurements [14] As a first step, the measurements of the three different PSPs are corrected for the thermal offset, i.e., the negative output signal due to the radiative cooling of the different instrumental components [Dutton et al., 2001]. To 2of12

3 Figure 1. Time series of (a) daily mean Ångström exponent and (b) daily mean aerosol optical depth at nm measured at Lampedusa during the period June 2004 to August Monthly averages (thick gray line) and standard deviations of the monthly averages are also shown. estimate the thermal offset we use the method developed by Dutton et al. [2001] which is based on the linear relationship between the PSP nighttime offset and the net thermopile output of a collocated PIR. The PSP offset and the net thermopile output of the PIR of each night are associated and fitted linearly; the resulting linear fit is rotated around the centroid of the data (T. L. Alberta and T. P. Charlock, A comprehensive resource for the investigation of shortwave fluxes in clear conditions: CAGEX version 3, paper presented at Tenth Conference on Atmospheric Radiation, American Meteorological Society, Boston, Massachusetts, 1999) to force it through the origin. The slope of the fit is used to estimate the thermal offset. The slopes of the retrieved linear fits are comprised, for the three different PSPs, between 0.05 and 0.1; these values are slightly larger than those reported by Dutton et al. [2001]. The obtained thermal offset varies between 0 and 25 W m 2 for PSP 33600, between 0 and 20 W m 2 for PSP 33504, and between 2 and 15 W m 2 for PSP The absolute value of the thermal offset is larger in daytime and cloud free conditions, and smaller in nighttime for heavily cloudy sky. The retrieved thermal offset is subtracted from the irradiance measurements to obtain a corrected PSP signal. [15] In several periods the PSPs were run side by side (two or three together). These periods of simultaneous measurements are used to compare and verify the performance of the three instruments, and to transfer the calibration from freshly calibrated to the other pyranometers. From the available days of simultaneous measurements, only those with a mean daily aerosol optical depth at 500 nm less than 0.1 and cloud-free conditions are selected; this is done to minimize the dependence of the measured irradiance on the presence of atmospheric aerosols. Four days satisfy these criteria for the pair PSP PSP 34891, and 18 days satisfy these criteria for the pair PSP PSP [16] The comparison between the irradiances measured with PSP and shows a different behavior, depending on the solar zenith angle, indicative of different cosine responses of the pyranometers: for a solar zenith angle q 55 the two data sets agree within about ±1.5%, while for 55 < q <80 the difference reaches about ±6.5%. The morning-to-afternoon differences in the measured irradiances are smallest during low-aerosol days for PSP 34891, and this instrument is assumed to produce the reference cosine response. From the comparison of the two data sets an updated calibration of PSP is derived. The calibration includes a correction with a linear function as a function of the solar zenith angle for q 55, and a polynomial function for q > 55. The corrected irradiances were again compared with the PSP irradiances and the two data sets agree within about ±1% for q 60, and within about ±2% for q >60. The PSP corrected irradiances were then compared with the PSP irradiances in the selected 18 days of simultaneous measurements. The two instruments show different cosine and azimuthal responses: the two data sets agree within about ±12% for q < 80, showing a different behavior between morning and afternoon. The data from 3of12

4 Figure 2. Behavior of the daily mean Ångström exponent as a function of the daily mean aerosol optical depth at nm measured at Lampedusa during the period June 2004 to August The selected three aerosol classes, desert dust, DD, urban/industrial-biomass burning aerosols, UI-BB, and mixed aerosols, MA, are shown. the PSP have been corrected for absolute responsivity, and for solar zenith/azimuth angle dependence of the responsivity with two simple correction linear functions of q. The first one is applied to morning data, the second one to afternoon data. After the calibration and correction the two data sets agree within about ±1% for q 60, and within about ±2% for q >60. Thus, we obtain a continuous record of uniform downward shortwave irradiances throughout the period , under the assumption that PSP has the correct cosine response. [17] The uncertainty associated with the determination of I SW# is the combination of the PSP measurement uncertainty (about ±2.2% for q 70, and ±3.6% for q >70 ), and the uncertainty associated with the transfer of the calibration (only for PSP and PSP 33504). The total uncertainty, different for each of the three PSPs, grows with q, reaching a maximum of ±4.6% for the PSP for q > 70. Following di Sarra et al. [2008], an additional ±2% random error is added to take into account the dependence of the downward shortwave irradiance on the water vapor atmospheric content. [18] An additional important parameter that affects the accuracy of the irradiance measurement is the pyranometer dome cleanness. All PSPs are ventilated and cleaned regularly; nevertheless, occasionally, a strong deposition of dust, sea salt, rain, etc., may occur. To minimize the effect of these events, measurements corresponding to days when the PSPs domes are found dirty by the local operator, as noted on the experiment log book, are eliminated from the data set Aerosol Optical Properties [19] The aerosol optical depth, t, at 415.6, 495.7, 614.6, and nm is calculated subtracting the contributions of Rayleigh scattering and ozone absorption (only for the channels at 495.7, and nm) to the total atmospheric optical depth. The uncertainty associated with the determination of the aerosol optical depth is 0.02 [Pace et al., 2006]. The Ångström exponent, a, is defined as the negative slope of t versus wavelength, l, in logarithmic scale and is calculated from the value of t at and nm. The uncertainty on a is calculated with the error propagation formula. Figure 1 shows the behavior of the daily mean Ångström exponent and the daily mean aerosol optical depth at nm measured at Lampedusa during the period June 2004 to August Monthly averages of t at nm and a and standard deviations of the monthly averages are also shown. The aerosol optical depth is generally lower than 0.6, with some isolated cases at 0.8 and 0.9, and displays an evident annual cycle, with a maximum in late spring and summer. The Ångström exponent is comprised between about 0.1 and 1.8, indicating the presence of aerosols both in accumulation and coarse mode. Largest values are observed in summer. Large day-today variations in t and a are evident. [20] Figure 2 shows the daily mean Ångström exponent versus the daily mean aerosol optical depth at nm measured in the period June 2004 to August Pace et al. [2006] have shown that different regions in the Ångström exponent optical depth plane correspond to aerosols of different origin and properties, and developed a method to discriminate among different aerosol types on the basis of their optical properties. Applying this method, the aerosol observations at Lampedusa were grouped into three classes: desert dust, DD (corresponding to cases with t at nm 0.15 and a 0.5), urban/industrial biomass burning aerosols, UI-BB, (cases with t 0.1 and a 1.5), and mixed aerosols, MA (for t > 0 and 0.5 < a < 1.5, t < 0.15 and a 0.5, and t < 0.1 and a 1.5). DD and UI-BB represent aerosols transported from the Sahara desert and the central eastern Europe to Lampedusa, respectively. MA includes cases of pure marine, marine polluted, and mixtures of various aerosol types. As discussed by Pace et al. [2006], cases of MA are observed both in correspondence with a rapid variation of the origin of air masses arriving at Lampedusa, and in presence of air masses of different origin at different heights; both situations favor the contemporary presence of different types of particles in the air column. The data points corresponding to the three aerosol classes are shown in Figure 2. Pure marine conditions are rarely observed because related to few cases with very limited influence from the European and African continents. For the pure marine aerosol class, Pace et al. [2006] report optical depths between 0.07 and 0.11 and an Ångström exponent of 0.8. The low optical depth, the limited number of cases, and the small interval of variability in optical depth make the application of the direct method impractical for pure marine aerosols, which will not be considered in this analysis. Table 1 reports averages of the aerosol optical properties for the three aerosol classes and for the whole data set over the period June 2004 to August [21] Large differences in size distribution and amount exist between the different classes: large particles are dominant in DD, while small- and medium-sized particles are dominant in UI-BB and MA, respectively. DD, UI-BB, and MA represent 26%, 6% and 68% of all the observed 4of12

5 Table 1. Averages of the Aerosol Optical Properties Over the Period June 2004 to August 2007 a Aerosol Type t a at nm at nm at nm at nm DD 0.33 ± ± ± ± UI-BB 0.20 ± ± ± ± MA 0.14 ± ± ± ± AD 0.18 ± ± a Averages of the aerosol optical depth at nm, Ångström exponent, single scattering albedo at nm and nm, and asymmetry factor at nm and nm for different aerosol types (desert dust, DD, urban/industrial biomass burning aerosols, UI-BB, mixed aerosols, MA, and all the data set, AD). Standard deviations of the averages are also shown. w g data (AD). DD and UI-BB are predominantly observed in spring and summer, while MA is more frequent in winter. [22] The single scattering albedo, w, is defined, at a given wavelength, as the ratio between the aerosol scattering and extinction coefficients. Through the combined use of radiative transfer model calculations and the measurements of DDR and t at and nm, Meloni et al. [2006] retrieve the single scattering albedo at these two wavelengths at the solar zenith angle of 60. As shown by Meloni et al. [2006], DDR is strongly dependent on the aerosol optical depth and, for a fixed value of t, depends on the single scattering albedo, which influences the amount of diffuse radiation reaching the ground. At a solar zenith angle of 45 the value of DDR at nm is 0.53 with an optical depth of 0.11, and becomes 1.64 with an optical depth of At 60 solar zenith angle we measure a DDR at nm of 0.4 for a desert dust optical depth of 0.17, and a DDR of 3.48 for a desert dust optical depth of The method by Meloni et al. [2006] is used here to retrieve w at and nm throughout the period June 2004 to August The estimated uncertainty on the retrieved single scattering albedo is <0.05 at nm and <0.07 at nm [Meloni et al., 2006]. Figure 3 shows the estimated values of w at and nm for the three aerosol classes during the investigation period. Averages of the single scattering albedo for the different aerosol classes are reported in Table 1. [23] The single scattering albedo of desert dust generally increases with wavelength. The retrieved values of w at nm for desert dust display a large variability, with values comprised between 0.7 and 0.8 during 2005, early 2006 and 2007, and values between 0.8 and 0.9 during 2004 and in the central part of The dust single scattering albedo at nm presents values between 0.7 and 0.8 in few cases in 2005, 2006, and Five-day back trajectories arriving at Lampedusa on the days when w is determined are analyzed to identify the origin of air masses arriving at Lampedusa, with the scope to explain the large variability in single scattering albedo. The Hybrid Single- Particle Lagrangian Integrated Trajectory (HYSPLIT) model [Draxler and Rolph, 2003, available at which uses meteorological model vertical velocity fields, is used to calculate the 5 days air mass back trajectories ending at Lampedusa at 750, 2000, and 4000 m. About 57% of the DD cases with w at nm lower than 0.75 correspond with trajectories at 750 m originating from central Europe, and trajectories ending at 2000 and 4000 m coming from Africa. About 70% of the trajectories ending at 750 m for DD cases with w at nm lower than 0.85 come from central Europe. Thus, polluted absorbing particles may be present in the lower atmospheric level (1 km), in addition to desert dust, in these cases. [24] For the DD cases with w at nm between 0.8 and 0.9, 75% of the trajectories ending both at 750 m and Figure 3. Evolution of the single scattering albedo at (solid circles) and nm (open circles) for cases of desert dust, DD, urban/industrial-biomass burning aerosols, UI-BB, and mixed aerosols, MA, observed during the period June 2004 to August of12

6 2000 m and 50% of the trajectories ending at 4000 m travel from the Sahara region through southern Italy, and spend a long time over the sea before arriving at Lampedusa. Thus, a significant contribution by marine particles, leading to high values of w, may be expected in these cases. [25] The single scattering albedo decreases with wavelength for UI-BB. Some cases with w at nm close to 1.0 are observed in mid The mean value of the single scattering albedo of MA is 0.81 at nm and 0.82 at nm, showing an almost neutral behavior of w with l. The values of w at and nm are spread between 0.7 and 0.9. Some cases with w at and nm greater than 0.9, probably corresponding to cases dominated by marine particles, are observed. [26] Our estimates of the single scattering albedo at nm for desert dust are in agreement, within the associated uncertainty, with those reported in previous studies, while result appear smaller at nm [i.e., D Almeida et al., 1991; Dubovik et al., 2002; Cattrall et al., 2003; Collaud Coen et al., 2004; Kim et al., 2005; Zhou et al., 2005]. For anthropogenic/mixed aerosols our results are in good agreement with those reported in previous studies [i.e., Dubovik et al., 2002; Horvath et al., 2002; Markowicz et al., 2002; Eck et al., 2003; Saha et al., 2008]. [27] The asymmetry factor, g, at nm and nm has been calculated for the three identified particle types using the Mie theory and assuming the aerosol properties given by D Almeida et al. [1991]. For DD, UI-BB, and MA, the desert background, the urban-industrial and the maritime polluted type are considered, respectively. The adopted values of g at and nm are reported in Table Determination of the Surface Shortwave Aerosol Radiative Forcing 4.1. Methodology [28] The surface shortwave aerosol radiative forcing, RF S, is defined as the difference between the observed, F net,s, and the aerosol-free, (F net,s ) a f, surface shortwave net flux: RF S ¼ F net;s F net;s a f where the surface shortwave net flux is the difference between the downward and the upward irradiances at the Earth s surface. The net fluxes in the presence and absence of atmospheric aerosols can be written as ð1þ F net;s ¼ ð1 A S ÞI SW# ð2þ F net;s a f ¼ 1 A ð SÞ I SW# where A S is the surface shortwave albedo and I SW# and (I SW# ) a f are the downward shortwave irradiances observed in presence and in absence of aerosols, respectively. [29] Different methods can be used to estimate RF S.In this study we use the direct method [e.g., Satheesh and Ramanathan, 2000] to derive the surface shortwave forcing efficiency, FE S, which is the RF S produced by aerosols with a f ð3þ an optical depth equal to 1. At fixed solar zenith angle FE S can be determined by a linear fit of F net,s (q) versus the aerosol optical depth: FE S ðþ¼df q net;s ðþ=dt q ðþ q RF S (q), at a given solar zenith angle, is calculated multiplying FE S (q) byt. [30] The value of FE S is a critical function of the optical properties of atmospheric aerosols. Beside size, the aerosol absorption is expected to play an important role. In the direct method the aerosol radiative forcing is directly determined from the observations, without further assumptions on the radiative fluxes in aerosol-free conditions. The aerosol vertical distribution may also influence the aerosol forcing. Meloni et al. [2005] have shown that differences in the aerosol vertical distribution have a small impact on the radiative fluxes at the surface in the visible spectral region, while produce a significant effect on the fluxes at the top of the atmosphere. From now on t indicates the aerosol optical depth at nm Parameterization of the Surface Shortwave Albedo [31] Direct measurements of the surface albedo are not available at Lampedusa, and a parameterization of this quantity is needed. The ENEA Station for Climate Observation is located along the northeastern coast of Lampedusa, about 15 m from the coastline (which is about 50 m above the sea level). Thus, both land and ocean contribute to the surface reflectance. [32] T. P. Charlock et al. (Validation of the archived CERES surface and atmosphere radiation budget at SGP, paper presented at the Thirteenth Atmospheric Radiation Measurements Science Team Meeting, Broomsfield, Colorado, 31 March to 4 April, 2003) show that the relevant scale for the determination of the albedo in condition of clear sky is 10 km. The surface shortwave albedo is calculated as the weighted average of land and ocean albedo over a 5 km radius circle centered at the measurement site at Lampedusa. Land and ocean contribute by 28% and 72% to the surface albedo, respectively. [33] For the land albedo the monthly mean Moderate Resolution Imaging Spectroradiometer (MODIS) measurements performed at Lampedusa between 2003 and 2004 are considered. The shortwave integrated land albedo has a minimum of 0.14 in February March and a maximum of 0.22 between August and October. [34] The ocean albedo is calculated, for the spectral interval mm, following Jin et al. [2004]. They provide values of the ocean albedo dependent on solar zenith angle, aerosol optical depth at 500 nm, and wind speed, w. The dependence of the ocean albedo on the wind speed is small for q <60, and it is taken into account only for q 60. The ocean albedo is calculated using measured values of t and w for the different values of q. It increases with q and decreases with w; forq <60 the ocean albedo increases with t, while decreases with t for q 60. The shortwave integrated ocean albedo varies from a minimum of 0.03 (q =20, t = 0) to a maximum of 0.20 (q =75, t = 0, w = 0 3 m s 1 ). ð4þ 6of12

7 Figure 4. Surface shortwave net flux versus aerosol optical depth at nm at solar zenith angles of (top) 30 and (bottom) 60 for (a, d) desert dust, DD, (b, e) urban/industrial-biomass burning aerosols, UI-BB, and (c, f) mixed aerosols, MA. Least squares linear fits are also shown. [35] The overall (considering the contributions by land and ocean) shortwave surface albedo varies from a minimum of 0.06 to a maximum of Determination of Surface Aerosol Shortwave Forcing Efficiency [36] Downward shortwave irradiance measurements, I SW#, corresponding to the period June 2004 to August 2007, together with the simultaneous determinations of t and a, are used in this analysis. For the determination of FE S using observations taken in different seasons, all values of I SW# measured by PSP are scaled to the mean Sun-Earth distance. [37] Only cloud-free periods are used in this analysis. The cloud-free intervals are identified using the cloud-screening algorithm described by Meloni et al. [2007]. [38] Averages of t, I SW#, and a are calculated over ±1.5 intervals at q = 20, 30, 40, 50, 60, 65, 70 and 75. The ratio R between the average downward irradiance and its standard deviation is calculated over the ±1.5 interval; to eliminate residual contamination by thin clouds, cases with R < 100 (variability greater than 1% of the average) are discarded. To remove possible further contamination by thin clouds and cases of rapid changes in aerosol optical depth, cases displaying a standard deviation of t over the ±1.5 interval larger than 0.03 were discarded is an empirical threshold for the variability of t identified by Pace et al. [2006] on the basis of the estimated uncertainties and on the typical variability of the optical depth. [39] FE S (q) is calculated, for the three classes of aerosols and for the whole data set, by calculating the least squares linear fit between F net,s and t (both averaged over ±1.5 of solar zenith angle) at q = 20, 30, 40, 50, 60, 65, 70 and 75. The uncertainty on the retrieved FE S (q) is estimated as the least squares linear fit error, taking into account the measurement errors on F net,s and t. The uncertainties on q and A S are small and have been neglected. The uncertainty on I SW# is calculated taking into account the PSP measurement uncertainty, the uncertainty due to the water vapor atmospheric variability (2%), and the standard deviation of the average over the solar zenith angle interval (<1%). The uncertainty on t is taken as the largest between the measurement error and the standard deviation over the ±1.5 solar zenith angle interval. [40] The retrieved values of FE S at different solar zenith angles are fitted with a third degree polynomial as a function of q (assuming FE S (90 ) = 0), and then integrated over 24 h to obtain the daily forcing efficiency, FE d.fe d is calculated, taking into account the evolution of q with time, for the summer solstice, the winter solstice, and the equinox. The daily aerosol RF S is obtained by multiplying the retrieved FE d by the daily average t. [41] The forcing efficiency at q = 60 is also calculated for different values of w at and nm. For each aerosol type, the data are divided in three intervals of single scattering albedo (0.7 w < 0.8, 0.8 w < 0.9, 0.9 w 1), both for w at and nm, and the surface 7of12

8 Figure 5. Surface aerosol shortwave forcing efficiency (calculated with respect to the aerosol optical depth at nm) versus solar zenith angle for desert dust, DD, urban/industrial-biomass burning aerosols, UI-BB, mixed aerosols, MA, and for all the observed data (AD). aerosol shortwave forcing efficiency is derived for each class separately. 5. Results and Discussion [42] Figure 4 shows the behavior of the shortwave surface net flux versus the aerosol optical depth at nm at q = 30 and 60 for DD, UI-BB, and MA. Least squares linear fits are also shown. For t < 0.2 the data display a larger spread, owing to the large influence of the measurement uncertainty on the low optical depth values and to a possible residual contamination by thin clouds, which also may become significant at low optical depths. The MA class displays the largest spread because of the lowest values of t in this class and the possible presence of aerosol with different properties. [43] Figure 5 shows the behavior of the surface aerosol shortwave forcing efficiency versus solar zenith angle for the different aerosol classes and for all the data set. FE S decreases (in absolute value) for increasing q for all the aerosol types. The uncertainties on the retrieved FE S decrease with q for all the aerosol types. This decrease is attributed to the progressive increase with q in the number of data points (a minimum of 282 at q = 20, and a maximum of 625 at q =65 ) due to the seasonal evolution of the solar zenith angle. A larger number of data points generally produce a better linear correlation between the data and a smaller uncertainty on the slope. The surface forcing efficiency is largest for MA at all values of q. FE S is smallest for UI-BB. di Sarra et al. [2008] show that the differences in size distribution and wavelength dependence of the single scattering albedo between UI-BB and DD may produce the observed partial overlap of the forcing efficiencies for these two types of particles at large values of q. For the whole data set the forcing efficiency is comprised, for all the solar zenith angles, between the values of DD and MA. This effect is due to the large influence of MA and DD within the data set: MA constitutes the large majority of the data while DD includes the largest optical depth; both these effects concur to determine the slope of the fitting line for AD. Within the estimated uncertainties, the results in Figure 5 agree with those reported by di Sarra et al. [2008] for the period It must be emphasized that data from a different time interval and from different pyranometers are used in this study. This result points toward the robustness of the obtained information. [44] Table 2 shows the retrieved values of the daily forcing efficiency FE d calculated as outlined in section 4.3, for the different aerosol classes at the equinox, summer solstice and winter solstice. As for FE S (q), the daily forcing efficiency is largest for MA, lowest for UI-BB, intermediate for DD, and comprised between MA and DD for AD. Using the average optical depth for each class the daily average radiative forcing at equinox is about 22 W m 2 for DD, 12 W m 2 for UI-BB, 13 W m 2 for MA and 14 W m 2 for AD. The radiative forcing is strongest for DD, almost twice as large as for UI-BB, because desert dust shows both a large forcing efficiency and the largest optical depth. As discussed by di Sarra et al. [2008] the aerosol size distribution plays a central role in producing a high forcing for DD (characterized by a large aerosol optical depth throughout the shortwave spectral range) and a low value for UI-BB (the optical depth rapidly decreases with wavelength). The single scattering albedo, as will be shown below, also plays a relevant role. [45] Figure 6 shows the behavior of the shortwave surface net flux versus t at nm at q =60 for cases of DD, UI-BB, and MA classified on the basis of their single scattering albedo at nm and nm. Least squares linear fits are reported for the different classes. Because of the small number of cases, it is not possible to derive FE S for some classes of aerosol and w. Table 3 shows the instantaneous shortwave surface forcing efficiency at q = 60 retrieved for the different ranges of single scattering albedo for DD, UI-BB, and MA. Uncertainties on the retrieved FE S are also shown. The surface forcing efficiency decreases (in absolute value) with single scattering albedo for all aerosol types, except for DD with w at nm. As expected, a stronger the absorption leads to a larger surface forcing. [46] The dust optical depth is almost independent on wavelength, and the spectral region producing the largest Table 2. Daily Forcing Efficiency FE d at Equinox, Summer Solstice, and Winter Solstice for DD, UI-BB, MA, and AD Aerosol Type at the Equinox FE d (W m 2 ) at the Summer Solstice at the Winter Solstice DD 67.2 ± ± ± 1.3 UI-BB 59.0 ± ± ± 3.5 MA 93.2 ± ± ± 1.5 AD 79.8 ± ± ± 0.8 8of12

9 Figure 6. Surface shortwave net flux versus aerosol optical depth at nm at q =60 for (a, b) desert dust, DD, (c, d) urban/industrial-biomass burning aerosols, UI-BB, and (e, f) mixed aerosols, MA, grouped in three classes of single scattering albedo at nm (Figures 6a, 6c, and 6e), and nm (Figures 6b, 6d, and 6f) (note changes in the x axis in the different graphs). Least squares linear fits are also shown. 9of12

10 Table 3. Instantaneous Shortwave Surface Forcing Efficiency at q =60 for DD, UI-BB, and MA a FE S ± s (W m 2 ) w DD UI-BB MA At nm ± ± ± ± ± ± ± 33 At nm ± ± ± ± ± ± ± 31 a Classified on the basis of their single scattering albedo at nm and nm. Uncertainties on the retrieved forcing efficiencies are also shown. impact is the one close to the median of the solar spectrum, i.e., around 700 nm. Consequently, the single scattering albedo around 700 nm is more representative of the dust radiative effect than w at nm. The dust single scattering albedo increases rapidly with wavelength from the UV to about 500 nm, and is relatively constant at longer wavelengths [i.e., Patterson et al., 1977; Dubovik et al., 2002; Cattrall et al., 2003; Derimian et al., 2006; Bergstrom et al., 2007]. Thus, the single scattering albedo at nm is much more indicative of the dust absorption in the most relevant portion of the spectrum than w at nm. The increase of FE S for dust with is w at nm seems to suggest that particles highly absorbing at nm display relatively low absorption at nm; this effect may be related to the influence of dust of different composition on the wavelength dependence of w [e.g., Alfaro et al., 2004], or possible mixing with other types of aerosols. For instance, mixing with carbonaceous particles, which display values of w larger than DD at nm and lower than DD at mm, might produce the observed behavior. [47] The optical depth of UI-BB and MA decreases with wavelength, and the radiatively effective portion of the spectrum is located at shorter wavelength than for DD. In addition, w displays a weaker wavelength dependence than DD [e.g., Bergstrom et al., 2007], and also w at nm gives information on the aerosol radiative effect throughout the SW range. [48] A 0.1 decrease in w at nm produces an increase in the absolute value of FE S by about W m 2. The ratio between FE d (Table 2) and FE S at 60 solar zenith angle (Figure 5) was calculated for DD, UI-BB, and MA. The ratio is comprised between 0.43 and 0.5 for the equinox, and between 0.55 and 0.63 for the summer solstice, depending on the aerosol type. Estimates of the daily average forcing efficiencies were derived for different values of the single scattering albedo using the FE d / FE S (60 ) ratios for the different aerosol types. The results for the equinox and summer solstice are reported in Table 4. The changes in FE d associated with a variation by 0.1 in w are of the same order of the changes in FE d associated with a change in aerosol type for the same value of w. Thus, size distribution and single scattering albedo play a competitive role in the determination of the forcing efficiency. [49] In Table 5, the results of our analysis are compared with recent estimates of the aerosol direct radiative forcing obtained in the Mediterranean. In general, values span a relatively large range of values. The largest daily forcing efficiencies are observed by Roger et al. [2006] and Saha et al. [2008]. These values by Roger et al. [2006] are relative to anthropogenic particles characterized by high black carbon concentration, with a low single scattering albedo (w = 0.83 at 550 nm) and a low asymmetry factor (g = 0.59 at 550 nm). Saha et al. [2008] discuss cases of continental polluted aerosol and continental dust (w at 525 nm between 076 and 0.9). The largest (negative) daily forcing efficiency ( 97.6 W m 2 ) is found for highly absorbing fine particles (w at 525 nm of 076, Ångström exponent of 1.81). The absorption by these particles is higher than in the UI-BB class and is close to the average MA of Lampedusa. The daily forcing of MA is close to the values reported by Roger et al. [2006] and Saha et al. [2008]. [50] The results obtained for MA are also in good agreement with those for mixing of different types of particles reported by Hansell et al. [2003], Ramana and Ramanathan [2006], and Markowicz et al. [2003]. [51] Few studies, principally based on radiative transfer models, have investigated the dependence of the aerosols direct forcing on the absorbing properties of particles. Kim et al. [2005] use 3 years of ground-based measurements carried out in east Asia to retrieve estimates of the daily forcing efficiency for different aerosol types. The values they retrieve are for stations between 28 and 36.5 N and for the month April. They found values of desert dust FE d of 106, 104, and 91 W m 2 for single scattering albedo at 500 nm of 0.80, 0.84, and 0.86, respectively. For urban aerosols with single scattering albedo of 0.86, 0.87, and 0.92 they retrieve daily forcing efficiencies of 90, 82 and 69 W m 2, respectively. These values may be compared with the FE d at the equinox reported in Table 4. Our values for DD are somewhat smaller than those by Kim et al. [2005]; results for urban aerosols are closer to the MA than the UI-BB cases. [52] Using radiative transfer model calculations and measurements carried out at 6.8 N in the period October 2004 to January 2005, Ramana and Ramanathan [2006] derive daily forcing efficiencies of 42, 64, and 80 W m 2 for cases of mixing between marine and polluted aerosols with single scattering albedo at 500 nm of 1.0, 0.95, and 0.90, respectively. The differences between these values and those in Table 4 may be partly due to differences in latitude and in aerosol properties. [53] Saha et al. [2008] derive daily forcing efficiencies of 97.6, 87.7, 73.5, and 68.1 W m 2 for four anthropogenic polluted events observed at 43.1 N in southern France in June The single scattering albedo at 525 nm Table 4. Estimated Daily Average Forcing Efficiencies for Different Aerosol Types and Different Values of the Single Scattering Albedo at nm FE d (W m 2 ) at the Equinox at the Summer Solstice w at nm DD UI-BB MA DD UI-BB MA of 12

11 Table 5. Daily Average and Instantaneous Aerosol Forcing Efficiencies for Different Aerosol Types in the Mediterranean Basin Reference Aerosol Type FE d (W m 2 ) FE S (W m 2 ) Present study DD 67.2 (equinox) 185/ 82 (20 < q <75 ) UI-BB 59.0 (equinox) 168/ 84 (20 < q <75 ) MA 93.2 (equinox) 251/ 120 (20 < q <75 ) Formenti et al. [2002] biomass burning 64 Horvath et al. [2002] continental polluted 60 Markowicz et al. [2002] anthropogenic 70.7 anthropogenic + fire particles 87.9 Derimian et al. [2006] desert dust 86 urban-industrial 81 Roger et al. [2006] anthropogenic 107 Cachorro et al. [2008] desert dust 130 (53 < q <75 ) biomass burning 147 (53 < q <75 ) desert dust + biomass burning 100 (53 < q <75 ) di Sarra et al. [2008] desert dust 68.8 (equinox) biomass burning + urban-industrial 54.4 (equinox) Saha et al. [2008] continental polluted 68.1/ 97.6 for the different cases was respectively 0.76, 0.80, 0.85, and These values compare well with the FE d at the summer solstice for types UI-BB and MA. 6. Conclusions [54] Measurements of the aerosol optical properties and shortwave radiative fluxes obtained at the Mediterranean island of Lampedusa during the period June 2004 to August 2007 are used to derive the surface aerosol shortwave direct radiative forcing. Estimates of the aerosol single scattering albedo are used in combination with the ground-based observations to study the dependence of the radiative forcing on the absorbing properties of particles. The main results of our study may be summarized as follows: [55] 1. Three different aerosol types, desert dust (DD), urban/industrial-biomass burning aerosols (UI-BB), and mixed aerosols (MA), are identified on the basis of their optical properties. During the period June 2004 to August 2007 the average optical depth at nm is 0.33 for DD, 0.20 for UI-BB, and 0.14 for MA. The average Ångström exponent is 0.24 for DD, 1.64 for UI-BB, and 0.82 for MA. [56] 2. The aerosol FE S is retrieved as a function of the solar zenith angle for the three aerosol classes and for all the observed data (AD). FE S decreases in absolute value for increasing q for all the aerosol types. Estimates of the daily mean forcing efficiency are derived at the equinox, at the summer solstice, and at the winter solstice. The largest FE d is observed for MA, followed by AD, DD, and UI-BB. The daily average radiative forcing efficiency at the equinox is 93.2 W m 2 for MA, 79.8 W m 2 for AD, 67.2 W m 2 for DD, and 59.0 W m 2 for UI-BB. The daily mean radiative forcing is largest for DD ( 22 W m 2 ) and lowest for UI-BB and MA ( 12 W m 2 ). These results are in good agreement with similar analyses made at Lampedusa during [57] 3. Cases of DD, UI-BB, and MA corresponding to q = 60 were grouped in three classes of single scattering albedo (0.7 w < 0.8, 0.8 w < 0.9, 0.9 w 1), at and nm. FE S was calculated for the different classes. FE S decreases for increasing w, except for DD with respect to the single scattering albedo at nm. This unexpected behavior of DD is attributed to the dust strong wavelength dependence of w at short wavelengths and to the largest role played by radiation at longer wavelengths in the determination of the forcing for DD. A 0.1 increment in the single scattering albedo at nm produces a reduction in FE S by about W m 2 and a reduction in FE d by about W m 2. [58] Acknowledgments. This study was partly supported by the Aeroclouds and the NOMAC programs, funded by the Italian Ministry for University and Research. Contributions by Lorenzo De Silvestri and Giandomenico Pace are gratefully acknowledged. Helpful comments and suggestions by three anonymous reviewers are also acknowledged. 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