Atmospheric turbidity over Egypt

ARTICLE IN PRESS AE International – Africa & the Middle East Atmospheric Environment 38 (2004) 1579–1591

Atmospheric turbidity over Egypt A.S. Zakeya,*,1, M.M. Abdelwahabb, P.A. Makara a

Air Quality Modelling and Integration Division, Meteorological Service of Canada, 4905 Dufferin Street, Toronto, Canada M3H 5T4 b Astronomy and Meteorology Department, Faculty of Science, Cairo University, Cairo, Egypt Received 17 August 2003; accepted 24 December 2003

Abstract Atmospheric turbidity conditions were monitored during the period 1989–1995 in rural and urban sites in Egypt. The rural site (Aswan) is in the south region of Egypt (desert climate), while the urban site is greater Cairo (Cairo and its surroundings). The measurements were used to recover the aerosol optical depth, Linke turbidity factor, and Angstrom turbidity indices, from the total direct solar radiation measurements and knowledge of the atmospheric ozone and water vapor content. Solar direct radiation measurements were carried out with an Eppley pyrheliometer, following World Meteorological Organization recommendations, improving on previous work that made use of Voltz sun photometers. An examination of the methodologies for calculating aerosol optical depth and comparisons to observations suggests that the Davies– Hay method is superior to the method of Unsworth–Monteith. A comparison of the seasonal cycle of aerosol optical characteristics at both sites showed: aerosol of photochemical origin in the summer; a significant impact of temperature, relative humidity and dust storms in the autumn; low values of aerosol optical characteristics in the winter (due to precipitative removal as well as relative humidity-impacted deposition); and high values in the spring resulting from seasonal dust storms. An examination of the inter-annual variability of the global direct solar radiation and the anomalous aerosol optical depth at both rural and urban sites showed the signature of both the El-Chichon and Mt. Pinatubo eruptions. The Pinatubo eruption was studied in detail: anomalously high aerosol optical depth (AOD) values and low global direct radiation was observed in mid-1991 over Cairo (AOD average: 0.04) and early 1992 over Aswan (AOD average: 0.039). The annual behavior of the Angstrom turbidity coefficient and Linke turbidity factor recorded maximum values (0.27 and 7.0, respectively) over Cairo in late 1991/early 1992, with maximum values over Aswan being (0.18 and 5.57, respectively) in late 1991. r 2004 Elsevier Ltd. All rights reserved. Keywords: Turbidity; Aerosol optical depth; Pinatubo; Irradiance; Solar radiation

1. Introduction Turbidity is a dimensionless measure of the opacity of a vertical column of the atmosphere. Parameters such as the Angstrom turbidity indices ða; bÞ; aerosol optical *Corresponding author. E-mail address: [email protected] (A.S. Zakey). 1 Visiting Post-Doctoral Fellow, on leave from the Egyptian Meteorological Authority, Cairo, Egypt.

depth (AOD), and Linke turbidity factor ðTL Þ are typically used as measures of turbidity. AOD is a numerical quantification of the total column extinction of transmitted radiation by atmospheric aerosols for broadband in the solar electromagnetic spectrum (WMO, 1994a,b). The importance of aerosols in the climate system are due to their radiative impact, where radiative heating effects are produced by aerosol absorption of short and long wave radiation (cf. Penner et al., 2001). The total global radiation forcing due to the direct and indirect

1352-2310/$ - see front matter r 2004 Elsevier Ltd. All rights reserved. doi:10.1016/j.atmosenv.2003.12.017

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effects of atmospheric aerosols is thought to be about 1:3 W m2 with an uncertainty of at least a factor of 2 (Shine and Forster, 1999). The upper bound of the uncertainty range for aerosol radiative forcing implies the possibility of a net global forcing of zero (Houghton et al., 2001). The range of the uncertainty illustrates the need for accurate turbidity measurements. Different aerosols have very different optical radiative forcing (characterized through the use of turbidity parameters). Accurate measurements of these parameters in a sufficiently large number of locations are necessary to characterize the global aerosol impact. Monitoring of the turbidity aerosol parameters ða; b; AOD; TL Þ of the background aerosols of natural origin and urban aerosols of industrial origin (Hansen and Lacis, 1990; D’Almeida et al., 1991; Lacis and Mischenko, 1994; Utrillas et al., 2000; Cachorro et al., 2000) are therefore of great interest for environmental and climatological studies (Preining, 1991; Gueymard, 1998; Holben et al., 1998). The Angstrom turbidity coefficient ðbÞ has long been used as a turbidity index and can be determined by different methods from spectral and broadband radiation measurements (Angstrom, 1929, 1961; Masmoudi et al., 2003). The amount of aerosol present in the atmosphere in the vertical direction can be represented by an index called the Angstrom turbidity coefficient, b; which varies typically from 0.0 to 0.5 according to Angstrom (1929, 1961, 1964), Kondratyev (1968), Gueymard (1998), and Grenier et al. (1995). In order to calculate b; we will need to define another parameter, a; the so-called wavelength exponent which is obtained from the spectral measurements of the solar radiation. The wavelength exponent is related to the size distribution of aerosols, with size-related fluctuations in a appearing in past work (e.g. Fouquart et al., 1984; Ichoku et al., 1999; Zakey, 1996, 2001). A closer examination of the results reported in the literature (Deepak and Gerber, 1983; Fouquart et al., 1984; Longtin et al., 1988; Shettle and Fenn, 1979; Shettle, 1989) shows that the size distribution of desert aerosols is highly dependent on wind speed. Under calm conditions, the size distribution and the corresponding a for desert aerosols are relatively close to those for continental aerosols ðaB1:3Þ; but under windy conditions their a values become very small or may be negative. AOD is the sum of aerosol absorption and scattering. Monitoring of the AOD of the atmosphere, in principle, requires simple measurements of the relative solar spectral irradiance and the application of the Beer– Lambert–Bouguer law. AOD is an appropriate variable for studies related to anthropogenic aerosols, global impact on radiative transfer, AOD being an integrated measure of this impact. AOD can also be used to compare the attenuation of direct solar radiation over all

time scales for which data are available. AOD measurements over long time periods and at different locations of different climate systems are therefore of great importance in environmental studies and applications. The calculation of AOD by using broadband irradiance measurements has been recorded by Utrillas et al. (2000), Andreae et al. (2002), Gueymard (1998), Kambezidis et al. (2000) and Qiu (1998). Two mathematical approaches used here to calculate AOD are based on the work of Davies and Hay (1980) and Unsworth and Monteith (1972). The Davies and Hay (1980) approach has subsequently been adopted by Freund (1983), Hay and Darby (1984) and Galindo et al. (1996). The origin of the broadband turbidity method can be traced to the introduction of Linke’s broadband turbidity factor, TL (Linke, 1922), which represents the number of hours of clean dry atmosphere that would be radiatively equivalent to the actual atmosphere under investigation. The TL factor has been computed here to allow comparisons with the numerous investigations based on this coefficient (e.g. Jacovides et al., 1994; Louche et al., 1986; Cucumo et al., 1999; Grenier et al., 1994, 1995; Polavarapu, 1978). Marani et al. (2002) used Voltz sun-photometer measurements to illustrate the effects of the 1992 Pinatubo volcanic eruption on the physical and radiative properties of stratospheric aerosols. El-Hussainy and Omran (1998) used different experimental methods to detect the behavior of El Chichon’s eruption on Cairo’s atmosphere in 1982– 1983; they found that the overburden caused by El Chichon’s eruption was about 0.06 in b and 0.9 in TL : This increase in TL related to volcanic eruptions is comparable to the large increase observed during the rapid urbanization of Mexico City (Galindo, 1984). The major physical difference between TL and AOD is that TL includes the attenuation of the solar rays due to water vapor, while AOD refers to an atmosphere with specified water vapor content (Kambezidies et al., 1993). Therefore, TL is a slight function of zenith angle and is far more strongly a function of water content than AOD (Gueymard, 1998). In other words, TL cannot be considered as a strict measure of atmospheric aerosol turbidity (Polavarapu, 1978; Yamamoto et al., 1968). Some investigators have used TL data to derive the Angstrom turbidity coefficient ðbÞ from empirical relationships (e.g. Cucumo et al., 1999; Dogniaux, 1975; Grenier et al., 1994; Jacovides et al., 1994). However, such empirical relationships may not provide a good estimate of b in all cases, with RMS errors of up to 50% (Cucumo et al., 1999; El-Hussainy and Omran, 1998). In this study, following WMO (1994a, b) recommendations, the calculation of AOD has been performed using broadband irradiance measurements with an Eppley pyrheliometer, instead of the traditional Voltz

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sun-photometers. Measurements using Voltz sun-photometers may be inaccurate, due to the lack of calibration information, the need for temperature and zenith angle corrections, and the sensitivity of the measurement results to manual adjustments made to the device by different observers; i.e. observer bias. Our results are therefore likely to be more accurate than earlier studies making use of Voltz sun-photometer (Zakey, 2001; Marani et al., 2002). Here, Eppley pyrheliometer measurements were acquired in a rural area representative of the desert climate of south Egypt (Aswan), and in urban area representative the climate of greater Cairo. These data are used to determine the general features of the atmospheric turbidity by means of the AOD as the turbidity index, and TL factor and Angstrom turbidity parameters as simple preliminary aerosol characterizations. The appropriate methods to determine AOD, TL and b are discussed together with the analysis of the input parameters in Section 2. In Section 3, we review the experimental measurements and discuss the resulting turbidity parameters, prior to the concluding section of this work.

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M.Matruh

El-Arish

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El-Kharga

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Aswan

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25.00 26.00 27.00 28.00 29.00 30.00 31.00 32.00 33.00 34.00 35.00

Fig. 1. The locations of the Egyptian radiation network monitoring sites.

between 0 and 2 oktas (1 okta ¼ 18 of sky covered by cloud) and occasionally 3 or 4 oktas. 2.1. Calculation of the TL

2. Experimental measurements Hourly measurements of direct solar radiation over Aswan and Cairo were taken in the period 1989–1995 to estimate AOD, TL factor and b: Greater Cairo (16,000,000 inhabitants) is facing, like most big cities in the world, serious air pollution problems. The rapid increase in population, large industrial areas, many traffic jams, and commercial activities have resulted in high air pollution levels during the last several years. In comparison, Aswan is a smaller (population 974,000) and less polluted city surrounded by desert areas. The radiation network station in Egypt consists of 10 stations, as shown in Fig. 1. The network started measuring global radiation in 1969 at four stations using Eppley pyranometers. In the beginning of 1980, the system was connected to digital integrators (CJ10). The measured values were recently corrected from the International pyrheliometric scale 1956 (IPS 56) to the World Radiometric Reference (WRR). Hourly direct solar radiation measurements (280– 400 nm) were recorded at Cairo and Aswan by normal incidence pyrheliometer (NIP) manufactured by Eppley. These instruments are calibrated each year against a reference instrument traceable to the WRR maintained at Davos, Switzerland. The time series of the calibration factors of the NIPs used shows no appreciable drift over long periods of time, confirming the remarkable stability of this instrument. Therefore, an accuracy of about 2% can be achieved in routine operation. Cloudiness at the times of the present pyrheliometeric observations varied

The basic relation of the extinction of solar radiation is of the form: Z N I ¼ E0 ISC I0 ðlÞ expðkðlÞÞ dl; ð1Þ 0

where l denotes the radiation wavelength, I is the solar radiation intensity measured by the instrument; E0 is the eccentricity factor; ISC is the solar constant; I0 ðlÞ is the solar radiation intensity outside the atmosphere, at the mean sun–earth distance; exp ðkðlÞÞ is the transmission of the whole atmosphere as a function of the wavelength, and the exponent kðlÞ is composed of two components: kðlÞ ¼ mtt ;

ð2Þ

where ðtt Þ is the optical depth of the total spectrum direct beam observed at the earth’s surface, and m denotes the corrected relative optical air mass. The corrected relative optical air mass depends on the solar zenith angle ðZÞ; on the actual pressure ðpÞ; on the standard pressure ðp0 ¼ 1013:25 hPaÞ; and consequently m at the altitude of the site is calculated via (Kasten and Young, 1989)   1 P 0:50572 cos Z þ : ð3Þ m¼ P0 ð96:07995  ZÞ1:6364 The optical depth of the total spectrum direct beam is calculated as the sum of five terms tt ¼ tR þ tO3 þ tg þ tw þ ta ;

ð4Þ

where tR is absorption transmittance of Rayleigh scattering; tO3 is the absorption transmittance of ozone;

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tg is the absorption transmittance of CO2 and SO2 ; tw is the absorption transmittance of water vapor; and ta is the transmittance of aerosol. The Linke turbidity factor ðTL Þ is defined as the ratio of the total optical depth to that of a clean, dry (Rayleigh) atmosphere: TL ¼ tt =tR :

ð5Þ

As a function of the extraterrestrial solar radiation at normal incidence ðH0 Þ; the total optical depth can be calculated as   1 I tt ¼  ln ; ð6Þ m H0

The absorption transmittance of Rayleigh scattering is given by tR ¼ expfð0:0903m0:84 Þð1 þ m  m1:01 Þg;

ð11Þ

where m is corrected relative optical air mass according to Eq. (3). The absorption transmittance of ozone is considered as tO3 ¼ 1  f0:161x0 ð1 þ 139:48x0 Þ0:3035  0:002715x0 þ 0:0003x20 g2 ; x0 ¼ O3 m:

ð12Þ

where H0 ¼ E0 ISc cosðZÞ: Following Louche et al. (1986) and Grenier et al. (1994, 1995), the total optical depth of a clean, dry (Rayleigh) atmosphere becomes

‘‘O3 ’’ represents the actual total amount of ozone over Cairo and Aswan taken from 1989 to 1995, measured in Dobson units. The absorption transmittance of gases and water vapor are given by

tR ¼ ð6:6296 þ 1:7513m  0:1202m2

tg ¼ expð0:0127m0:26 Þ;

3

4 1

þ 0:0065m  0:00013m Þ : From Eqs. (5)–(7), the TL factor is calculated as   1 H0 TL ¼ : ln mtR I

tw ¼ 1  2:4959Um mfð1 þ 79:034Um mÞ0:6828 þ 6:385Um mg1 : ð8Þ

2.2. Calculation of broadband AOD 2.2.1. Unsworth and Monteith (1972) approach ‘‘tM ’’ AOD was introduced by Unsworth and Monteith (1972) (represented by ‘‘tM ’’ here) and expresses the absorption of the solar rays by a dust-laden atmosphere relative to a dust-free one with specified water vapor content. Typically, tM varies from 0 to 1. An expression for tM is given by ðln½Id =Id  þ ln SÞ ; ð9Þ tM ¼  m where Id represents normal incidence total solar radiation at the earth’s surface ðWm2 Þ; Id represents the normal incidence total solar radiation at the bottom of a dust-free atmosphere ðWm2 Þ and S is the correction term for the earth–sun distance. The major physical difference between what is measured by TL and tM is that TL is focused on the attenuation of the solar rays produced by a water-vaporfree atmosphere, while tM is referenced to an atmosphere with specified water vapor content. Following Bird and Hulstrom (1981), the direct solar beam reaching the surface of the earth and passing through such an atmosphere can be estimated through the relationship I  ¼ 0:9751ðI t t t t Þ; ð10Þ d

ð13Þ

ð7Þ

0 R O3 g w

where I0 is extraterrestrial solar radiation beam as before.

ð14Þ ðH=22 000Þ

According to Gates (1962), Um ¼ 2:3eM 10 ; where H is height a.m.s.l (m) and eM ¼ es RH=100; where RH is the ambient relative humidity and eS is the water vapor pressure (mbar). 2.2.2. Davies and Hay (1980) approach ‘‘tD ’’ AOD according to the Davies and Hay (1980) approach (represented by ‘‘tD ’’ here) is calculated from: tD ¼

1 ; fm lnðId =H0 ðtO3 tR  aw ÞÞg

ð15Þ

where the units of Id and H0 in Eq. (15) are MJ m2 h1 ; and aw is the water vapor absorption. Here, ozone attenuation is calculated using empirical formulae of Lacis and Hansen (1974). The relevant equations are tO3 ¼ 1:0  avis  auv ; 0:002118X avis ¼ ; ð1:0 þ 0:0042X þ 0:00000323X 2 Þ 0:1082X 0:00658X auv ¼ ; þ 0:805 ð1:0 þ ½10:36X 3 Þ ð1 þ 13:86X Þ X ¼ O3 m:

ð16Þ

Again, the ozone amount is taken to be the actual measured value in Dobson units. The water vapor absorption is determined according to Lacis and Hansen (1974) using   2:9w0 aw ¼ ; ½ð1 þ 141:5w0 Þ0:635 þ 5:925w0   p 0:75 ; ð17Þ w0 ¼ mw 1013:25

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where p is the atmospheric pressure (mbar) and w is the total water column in mm, is estimated using the equation suggested by Won (1977): w ¼ 0:1 expð2:257 þ 0:05454Td Þ;

ð18Þ

where Td is the dew point temperature in C: The Rayleigh transmittance term tR is obtained using a polynomial fitted to the data of Davies and Hay (1980): tR ¼ 0:98552  0:10345m þ 0:01733m2 þ 0:00198m3 þ 0:00011m4  0:000002m5 :

ð19Þ

At a constant value of the Angstrom wavelength exponent ða ¼ 1:3Þ; the Angstrom turbidity coefficient, b; can be computed after modifying the formula of Louche et al. (1987) and Machler (1983) to be   1 tD  F b¼ ln ; ð20Þ mD C where the constants C; D; and F are defined as functions of a C ¼ 1:003  0:125a; D ¼ 1:089a þ 0:5123; F ¼ 0:1244a  0:0162:

ð21Þ

The a value used here is typical of small-sized particles as components of a background and relatively aged continental boundary layer that has been polluted by sources of non-local origin, and its value here is representative of the broadband average. Some limiting conditions have been introduced for the Linke and other atmospheric aerosol parameters calculations, such as; Linke turbidity varies from 1 to 10, with 10 being an upper limit. Also, Unsworth–Monteith AOD varies from 0 to 1, with 1 as an upper limit.

3. Results and discussion 3.1. Comparison of Aswan and Cairo seasonal cycle of aerosol optical characteristics The daily variation of post-astronomical-reduction clear-sky aerosol optical characteristics (b; tD ; tM ; and TL ) is metric-dependent. This in turn impacts on the calculated monthly and annual variation for each metric. These variations also show the sensitivities of each metric to the level of atmospheric pollution, in turn dependent on the regional climate experienced by the area of the study. Here, monthly values were obtained by averaging the hourly calculated values of b; tD ; tM ; and TL in each month. The spring season (March to May) over Cairo and Aswan is characterized by strong winds, moderately high temperature, low relative humidity, and decreasing

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surface pressure as shown in Figs. 2(a)–(f). These spring weather patterns usually carry fine dust and sand particles from the desert area to Cairo and Aswan. The number of dust storm occurrences over Cairo and Aswan maximizes in the spring (Fig. 2(f)). This leads to increases in the values of tM and tD as shown in Figs. 3(a) and (b) on monthly time scales. The monthly values of b over Aswan are higher than those in Cairo in spring season as shown in Fig. 4(a), while the monthly value of TL is higher over Cairo than Aswan in spring as shown in Fig. 4(b). All of the aerosol optical characteristic metrics (b; tD ; tM ; and TL ) thus recorded significant high values in spring season, due to the effects of sand and dust particles emitted from the surface of the surrounding desert and local sources of wind-blown dust. Comparisons between the different metrics across different sites show that spatial variations may be metric-dependent; b and TL variations do not correlate spatially within a given season. The summer season (June to August) is characterized by high temperature, low relative humidity, and low wind speeds over Cairo and Aswan. This climate situation leads to increased photochemical processing in the atmosphere and hence greater anthropogenic aerosol production. Mobile (automotive) emissions are one of the major sources for poor air quality in the Cairo area, and the aerosols formed through primary emissions and secondary reactions are thus potential causes for the observed turbidity values during this season. In addition, strong vertical temperature gradients in the Cairo area in late summer lead to enhanced vertical convection, carrying aerosols aloft and enhancing turbidity (see b values, Fig. 4(a)). The summer background aerosols in Aswan are highly variable from month to month within the summer, with rapid changes in turbidity (especially as characterized by b). This may indicate long-range transport of aerosols to Aswan. Decreases in the relative humidity in June over Aswan may relate to the observed decreases in aerosol optical characteristics (b; tD ; tM ; and TL as shown in Figs. 2(c), 3(a), (b) and 4(a), (b). In conclusion, the turbidity characteristics for the summer season implies aerosol of photochemical aerosol origin, with slight increases in the values of (b; tD ; tM ; and TL ) despite lower wind speeds from the previous season. The autumn season (September to November) differs from spring with higher values in temperature and relative humidity, and wind events leading to dust storms. As a result of this climatic situation, the values of aerosol optical characteristics still have higher values than spring until the end of this season. The lowest values of aerosol optical characteristics over Aswan are recorded in November, as opposed to the winter season as shown in Cairo: this may be a result of Aswan’s relative humidity values in this month.

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Fig. 2. Monthly variation of some meteorological parameters over Cairo and Aswan: (a) maximum temperature ð CÞ; (b) minimum temperature ð CÞ; (c) ambient relative humidity (%), (d) atmospheric pressure (mbar), (e) total amount of rainfall (mm), (f) total number of dust storms.

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AOD_ASWAN

Fig. 3. Monthly mean variation of AOD over Cairo and Aswan during the period 1989–1995: (a) AOD as tD from Davies–Hay method; (b) AOD as tM from Unsworth–Monteith method.

The winter season recorded the minimum values of aerosol optical characteristics, due to the increased amount of rainfall over both sites as shown in Fig. 2(e). Although the total amount of rainfall over Cairo is

greater than Aswan, we found that the values of aerosol optical characteristics over Cairo are still greater than Aswan in winter season, indicating that rainout and washout of aerosols does not completely counteract the

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Angstrom Turbidity coefficient over Cairo

Angstrom Turbidity coefficient over Aswan

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(b)

Linke turbidity coefficient over Cairo Linke turbidity coefficient over Aswan

Fig. 4. Monthly variation of Angstrom turbidity coefficient and Linke turbidity factor over Cairo and Aswan 1989–1995: (a) Angstrom turbidity coefficient, b; (b) Linke turbidity factor, TL :

impact of local emissions and photochemical aerosol formation. Also, the winter decrease in atmospheric turbidity may result from increases in relative humidity and decreases in temperature, leading to increased water uptake in the aerosols, hence increased particle size and mass, and increased fall and deposition velocities. The two main causes of the low aerosol optical characteristics in the winter season are washout by rainfall and high relative humidities leading to aerosol size increases and subsequent deposition. 3.2. Inter-annual variability in AOD (Pinatubo contribution)

Fig. 5. Annual variation of direct solar radiation over Cairo during the period 1980–1995 ðWm2 Þ:

Volcanic eruptions eject gas and dust into the earth’s upper troposphere, where dispersion processes lead to worldwide transport of the ejected material. Sulfur dioxide gas emitted by the volcanoes is oxidized over a period of weeks to form sulfuric acid gas. This gas nucleates and/or condenses to form aerosols that impact the direct radiation. From the annual variation of direct solar radiation over Cairo during the interval 1980–1995, the eruption signatures of El-Chichon in 1982 and Pinatubo’s 1992 may be distinguished (minima). Here we will concentrate on the effect of Pinatubo’s eruption, the larger of the two signatures shown in Fig. 5. The Pinatubo eruption signal may be seen as higher values of anomalous variation of tD (7 year average subtracted from the yearly value), recorded in 1991 over Cairo and Aswan, and in 1992 over Cairo (Figs. 6(a) and (b)), and in the annual values of b in 1991 at both sites, and for Cairo in 1992 (Fig. 7). The high annual values in Cairo may also be affected by anthropogenic pollution—comparing Figs. 6(a) and (b), it can be seen that only the 1991 anomaly is consistently positive for both AOD calculation methodologies. However, comparing satellite-based measurements of aerosol mass loading (NASA, 2003) for the closest SAGE-II datapoint to Cairo in 1991 and 1992 (Figs. 6(c) and (d),

respectively), it can be seen that a strong upper atmosphere loading is present in the 1992 yearly average (Fig. 6(d)), but is considerably smaller in the 1991 yearly average (Fig. 6(c)). It should be noted that the 1991 SAGE-II values include relatively clean conditions prior to the June 1991 Pinatubo eruption and subsequent global transport. The 1989 high anomalous AOD at Aswan is likely due to intense sandstorms observed during that time. Pinatubo’s eruption was of larger magnitude than ElChichon’s (1982) and led to an estimated 50% increase in stratospheric extinction compared to El-Chichon’s (Michalskym et al., 1994). These aerosols first produced a strong depleting effect on direct radiation, as documented elsewhere (Michalskym et al., 1994), but subsequently decayed exponentially over time. Aerosol optical characteristics during the eruption for Aswan were b ¼ 0:0161; tD ¼ 0:00246; tM ¼ 0:025; and TL ¼ 0:244; compared to Cairo’s values of b ¼ 0:0497; tD ¼ 0:01389; tM ¼ 0:0707 and TL ¼ 0:746; during the same period. The lower values of aerosol parameters recorded over Aswan relative to Cairo (e.g. Fig. 7) may result from latitudinal differences in the spread of the Pinatubo aerosol cloud, as well as the impact of local pollution for the Cairo values. The former explanation

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Fig. 6. Anomalous AOD of Cairo and Aswan during the period 1989–1995: (a) based on Unsworth–Montieth method; (b) based on Davies–Hay method; (c) SAGE II aerosol mass loading ðmgm3 Þ; Cairo, 1991; (d) SAGE II aerosol mass loading, Cairo, 1992.

0.28 0.26 0.24 0.22 0.2 0.18 0.16 0.14 0.12 0.1 1989

1990

1991

1992

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Years Beta-extinction coefficient over Cairo

season). It should be noted that the highest difference values of the aerosol optical characteristics recorded in early spring and late autumn (i.e. the transitional seasons) also include the effects of seasonal dust and sand storms on the variability of aerosols, factors not directly related to the volcanic eruption. Overall, the AOD determined from the pyrheliometer measurements carried out at Cairo and Aswan was found to be higher following the volcanic eruption, likely the result of volcanic aerosol particles suspended at 16.5–24:5 km (Figs. 6(c) and (d)).

Beta-extinction coefficient over Aswan

Fig. 7. Detail of AOD by month (Davies–Hay) for 1991 and 1992: (a) Cairo, (b) Aswan.

3.3. Comparison of different AOD metrics and previous aerosol measurements

has some merit because a strong latitudinal gradient existed during the whole life of the Pinatubo aerosols. Using the average of 1995 as a reference of unperturbed year, an estimated overburden was calculated by comparing monthly averages of AOD in 1991 and 1992 to the same months in 1995, as shown in Fig. 8 (we assumed that the Pinatubo effect did not extend for more than 3 years). The 1991 and 1992 values of anomalous AOD from both methods indicate greater aerosol overburden relative to the same months in 1995, aside from November, December and January (rainy

The comparison between Davies–Hay (1980) and Unsworth–Monteith (1972) approaches over Cairo and Aswan show considerable variability in their predictions of AOD. The two AOD methods (Davies–Hay and Unsworth–Monteith) differ in several key ways. The Davies–Hay method makes use of measurements of dewpoint temperature and surface pressure as inputs, while the Unsworth–Monteith method requires relative humidity and saturation vapor pressure of water as inputs, for example. The first two quantities are directly measured, while the second two are indirect, requiring calculations from the measured quantities to obtain

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Fig. 8. AOD differences between 1991 ((a), (b)) and 1995, and 1992 and 1995 ((c), (d)), for Cairo ((a), (c)) and Aswan ((b), (d)), Davies–Hay methodology.

0.5

Davies-Hay AOD Unsworth-Monteith AOD

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0.55 0.5 0.45

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AOD over Cairo by two different methods

AOD over Aswan by two different methods

0.05 2

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12

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2

3

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Fig. 9. Comparison of AOD values calculated by Unsworth–Monteith and Davies–Hay methods, for Aswan (a) and Cairo (b).

values for the respective AOD formulae. More importantly, the Davies–Hay and Unsworth–Monteith methods differ in the formulae used to obtain the Rayleigh scattering (compare formulas (19) and (11), respectively), the absorption transmittance of ozone (compare formulas (16) and (12), respectively), and the absorption due to water vapor (compare formulas (17) and (14), respectively). The two methods also differ in that Unsworth–Monteith includes a term for the absorptive transmittance due to other gases in the atmosphere

(Eq. (13)), while total air mass effects in the Davies–Hay method enter into the equation via other terms. The most important differences between the two methods can be seen in Fig. 9, and this is the key in determining which method is the recommended one for future use. Several researchers (WMO, 1994b; Zakey, 1996; Cachorro et al., 2000; Jacovides et al., 1994), based on direct AOD measurements of the atmosphere, note a typical seasonal sinusoidal variation in AOD; with maxima in the summer and minima in the winter.

ARTICLE IN PRESS Angstrom Turbidity coefficient (Beta) over Cairo

0.450 0.400 0.350 0.300 0.250 0.200 0.150 0.100 0.050 0.000 Jan

Feb

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Angstrom Turbidity coefficient (Beta) over Aswan

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0.5 0.45 0.4 0.35 0.3 0.25 0.2 0.15 0.1 0.05 0

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El-Hussainy 75-79

Omar 75-79

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Months El-Syed 1980-1992

This study

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Soudy 1991-1992

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(b) Fig. 10. Monthly Values of b from other studies compared to this study: (a) Cairo, (b) Aswan.

Fig. 10 shows that the Davies–Hay method exhibits this expected behavior with a small degree of variability within each month (max–min values are shown in the figure in addition to the average value, for each method). The Unsworth–Monteith method results show less of the expected seasonal behavior, have a higher degree of variability, and thus are much more sensitive to the input conditions. Based on these considerations, we recommend the use of the Davies–Hay method over the Unsworth–Monteith method for future work. The work by several other researchers in the same region used other technologies that were not recommended in the recent WMO review. El-Sayed (1996) and El-Hussainy (1986) used the measurements of Voltz sunphotometer; Omar (1988) and Omran (1989) used the measurements of Georgi-type Actinometer. Fig. 10 shows the comparative results of Angstrom turbidity coefficient ðbÞ between this study and previous Voltz photometer work over Cairo (Fig. 10(a)) and Aswan (Fig. 10(b)) (El-Sayed, 1996; El-Hussainy, 1986), as well as Georgi Actinometer measurements at these locations (Omar, 1988; Omran, 1989). It is interesting to note that the measurements reported here for Cairo have the lowest variability throughout the year, while those by earlier instrumentation showed stronger monthly variability. Some of the total variation between the measurement data sets may be due to the effects of water and other gas extinction, ignored in the Omar, Omran, and El-Hussainy measurements (which also may have had lower precision, based on the WMO recommendations). The Aswan values show less scatter, with the exception of the Georgi measurements of Omran, which are anomalously high relative to the others depicted here. The Omran data did not filter out the effects of water vapor and other trace gas extinction. The Souady (1998) measurements and all others at Aswan aside from those performed as part of the current work were made using

Georgi Actinometers. Statistical analysis of the different data sets was used to determine that the lowest values of b were recorded in the present study, and are accompanied by the lowest value of standard deviation over both sites. The range of values of b recorded here (0.218–0.252) is similar to other Eppley measurements (e.g. Grenier et al., 1995; 0.01–0.13), while the range from the four data sets discussed above is 0.08–0.43, and the range reported in other published work using Voltz instrumentation is also larger (e.g. Cachorro et al., 2000; 0.05–0.8). This is in accord with the WMO recommendation, with the new values expected to be more accurate than those recorded in earlier work with other, non-recommended technologies and methodologies. Further improvements to accuracy could be obtained with more detailed spectral measurements, and are recommended for future work.

4. Conclusions and recommendations The turbidity measurements presented here were obtained from broadband measurements using an Eppley pyrheliometer, at two sites in Egypt; Cairo and Aswan. The level of turbidity was found to depend greatly upon the emissions and climate system of each area. The measurements performed as part of this work made use of more accurate measurement technology than previous work in the same region (Eppley pyrheliometer rather than Voltz sun-photometers), where a variety of measurement methodology factors led to higher levels of scatter in the earlier measurement records. Our analysis suggests that the AOD estimation method of Davies and Hay (1980) exhibits seasonal dependencies more in line with observations than the method of Unsworth and Monteith (1972). The former

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method also has a much lower sensitivity to input conditions than the latter, and is recommended over the latter for future work on AOD. The comparison of the Angstrom turbidity coefficient ðbÞ for the previous work in the same region by Souady (1991–1992), El-Sayed (1980–1992), Omran (1975–1983), Omar (1975–1979), El-Hussainy (1975–1979) and the results of this study showed that the new work has a smaller range of values, suggesting that the high scatter in the earlier work may have been the result of instrumentation errors. The current work makes use of WMO (1994a, b) recommendations for broadband measurements of turbidity— detailed spectral measurements would be of benefit in future work, in order to study the size distribution impact of aerosols on net turbidity. It was found that in the transitional seasons (spring and autumn), the aerosol optical characteristics (b; tD ; tM ; and TL ) recorded significant high values, due to the effects of sand and dust particles emitted from the desert, and also from road-dust sand emitted from the road inside the towns itself. These effects were accentuated in autumn by higher relative humidities and temperatures. Cairo city is also surrounded by Mokatm Hill which provides Cairo with fine sand during strong spring and autumn winds. In the summer season photochemical processes became the main origin of the aerosol and this leads to slight increases in the values of aerosol optical characteristics, despite lower wind speeds (hence less wind-blown dust) relative to other seasons. In the winter season, the lowest aerosol optical characteristics were measured over Cairo and Aswan. This was due to the washout by rain and high relative humidity (the latter responsible for increases in aerosol size and deposition especially in the morning hours and late night). From the measurements of direct solar radiation over Cairo, it was noticed that the effects of Pinatubo’s eruption in 1991 were stronger than El-Chichon’s in 1982. Anomalous aerosol optical characteristics determined from 7 years of pyrheliometer measurements carried out at Cairo and Aswan were found to be higher following the volcanic eruption, suggesting that solar radiation extinction was mainly produced by volcanic aerosol particles suspended at elevations between 16.5 and 24:5 km:

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