Reanalysis of total ozone measurements at Dombå s and Oslo, Norway, from 1940 to 1949

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. D24, 4750, doi: /2003jd003963, 2003 Correction published 22 May 2004 Reanalysis of total ozone measurements at Dombå s and Oslo, Norway, from 1940 to 1949 Tove M. Svendby Department of Physics, University of Oslo, Oslo, Norway Received 9 July 2003; revised 3 September 2003; accepted 3 October 2003; published 17 December [1] Total ozone measurements from Dobson spectrometer number 8 at Dombås, Norway (62.1 N, 9.1 E), and Oslo, Norway (59.9 N, 10.7 E), from 1940 to 1949 have been examined and reanalyzed. New sets of Bass-Paur absorption and Rayleigh scattering coefficients have been created, and total ozone values have been recalculated using the new coefficients. Approximately half of the ozone registrations at Dombås were based on direct Sun observations calculated from the CC 0 wavelength pairs. The long C 0 pair has provided valuable information about the effect of atmospheric aerosols on the ozone measurements. A method for determining aerosol corrections is presented which demonstrates that the monthly mean aerosol error can reach 4% total ozone, normally with a higher correction in the summer than in the winter. Also, the influence of SO 2 on the Oslo ozone measurements is estimated. The D8 ozone series from 1940 to 1949 has been compared to ozone records from other European stations and to the D56 ozone series from Oslo Studies of old and new Dobson data demonstrate that the annual variation in the ozone layer has changed during the last years. The comparison indicates that the ozone decrease is relatively small for the summer months (2.9 ± 1.8%), whereas the average winter and spring values have deceased by 6.1 ± 3.9% from the 1940s to the 1990s. If the postwar increase in tropospheric ozone is taken into account, the depletion of the ozone layer is considerably higher, probably 8 9% for winter/spring. All the reanalyzed total ozone data from D8 for the period 1940 to 1949 are available at the World Ozone and Ultraviolet Radiation Data Center. INDEX TERMS: 0305 Atmospheric Composition and Structure: Aerosols and particles (0345, 4801); 1610 Global Change: Atmosphere (0315, 0325); 1704 History of Geophysics: Atmospheric sciences; 1794 History of Geophysics: Instruments and techniques; KEYWORDS: stratospheric ozone, trends, Dombås, historical data, Dobson Citation: Svendby, T. M., Reanalysis of total ozone measurements at Dombås and Oslo, Norway, from 1940 to 1949, J. Geophys. Res., 108(D24), 4750, doi: /2003jd003963, Introduction [2] Total ozone measurements performed during the last two decades have revealed a negative ozone trend in most places on earth. Since the mid 1970s the destruction of the ozone layer by anthropogenic gas release has been discussed repeatedly [e.g., Molina and Rowland, 1974; Stolarski and Cicerone, 1974; Farman et al., 1985; Solomon, 1999]. The most dramatic effect of this destruction can be observed over Antarctica from September until November: the so called ozone hole [Farman et al., 1985]. The atmosphere in the Northern Hemisphere is different from the Antarctic and the conditions needed for catalytic polar ozone loss (namely, cold temperature and the polar vortex) are generally less effective at northern latitudes than over the South Pole. In order to obtain information about possible ozone destruction in the northern stratosphere caused by release of CFC gases, we have Copyright 2003 by the American Geophysical Union /03/2003JD reanalyzed 60-year-old total ozone measurement from southern Norway and compared them with modern total ozone data from Oslo. [3] The ozone measurements were performed with Dobson spectrophotometer number 8 (D8) which was in continuous operation at Dombås, Norway (62.1 N, 9.1 E) from 20 March 1940 until 18 June The instrument was lent to Norway from the International Union of Geodesy and Geophysics as a part of a planned international investigation on atmospheric ozone. Einar Tønsberg, Director of The Auroral Observatory in Tromsø, Norway, supervised the ozone work at Dombås. The ozone observations were carried out by Sigurd Einbu and his son Per Einbu [Langlo, 1952]. In June 1946 the instrument was moved from Dombås to Oslo and set up at the Norwegian Meteorological Institute (59.9 N, 10.7 E) under the supervision of Kaare Langlo. Ozone observations are available from Oslo for the period 19 July 1946 to 30 April [4] In 1978 Søren H. H. Larsen started a new series of daily ozone registrations at the University of Oslo with Dobson spectrophotometer number 56. The ACL 3-1

2 ACL 3-2 SVENDBY: TOTAL OZONE IN SOUTHERN NORWAY, Table 1. Number of Direct Sun (DS) and Zenith Sky (ZS) Total Ozone Measurements Performed at Dombås/Oslo From 1940 to 1949 a Year Number of DS Observations Number of ZS Observations Missing Months Jan., Feb., Dec Jan., Dec Jan., Feb., Oct., Nov none Jan., Feb none none 1947 none 1948 none 1949 May Dec. a The Dombås series terminated in June There is no information about daily total ozone values from July 1946 to May ozone series from Oslo has recently been reevaluated [Svendby and Dahlback, 2002] and in this paper we compare the latest Oslo data with the ozone readings obtained 60 years earlier. [5] In order to make an accurate reanalysis of Dobson total ozone measurements, detailed instrumental and calibration history of the instrument should be known. This has been a major challenge in the D8 reanalysis work. Information regarding the instrument calibration was lost during a major clean up several years ago and our knowledge about D8 is more or less entirely based on two papers from Langlo [1952] and Tønsberg and Langlo (Olsen) [1944]. In addition the original D8 registration forms are available and in excellent condition, and they have provided valuable information about the D8 instrument. [6] According to the ozone registration forms the slits of the old D8 spectrophotometer were slightly different from the slits in modernized instruments. Consequently, the 1992 Bass-Paur absorption and scattering coefficients [Bass and Paur, 1985; Komhyr et al., 1993] could not be applied directly when recalculating D8 raw data to modern scale total ozone values. Instead we have created a new set of absorption and scattering coefficients which assumingly improve the precision of the reanalyzed Dobson data. This will be described in section 3. [7] Although the detailed calibration history of D8 is unknown we believe that the old D8 Dobson series from Dombås is a high-quality record. The observers were excellent astronomers and meteorologists and Prof. Langlo and Tønsberg had regular contact with G. M. B Dobson [Tønsberg and Langlo (Olsen), 1944]. The high frequency of ozone registrations (see Table 1) and handwritten notes on the registration forms indicate that the instrument was well maintained. [8] For the period July 1946 to April 1949 all the registration forms from Oslo are missing and only monthly mean total ozone values are available [Langlo, 1952]. These data have larger uncertainties and are treated with extra care in our reanalysis work. Furthermore, we believe that SO 2 pollution in Oslo was considerable in the late 1940s, which possibly influenced the measurements and gave fictitiously high total ozone values [de Muer and de Backer, 1992; Svendby and Dahlback, 2002]. This will be discussed more carefully in section Observation Method 2.1. Ozone Measurements From a Single Wavelength Pair [9] Detailed information concerning derivation of the mathematical equations used for total ozone calculations are given by Dobson [1957]. However, a short summary of the theory is presented in this chapter to offer a better overview of the reevaluation process. [10] Monochromatic radiation of intensity I 0, passing through the atmosphere with extinction coefficient k and slant thickness s, will be reduced according to Lambert- Beers law, log I=log I 0 ks. The extinction coefficient k describes the Rayleigh scattering for the whole atmosphere, b [(atm) 1 ], extinction due to aerosol scattering, d [(atm) 1 ], and ozone absorption, a [(atm cm) 1 ]. The relative path length of light passing through the ozone layer (m) and the atmosphere (m) will change in accordance with the solar zenith angle Z. Thus monochromatic light passing through the atmosphere can be expressed as log I ¼ log I 0 amx þ bm p þ d sec Z ; p 0 where p is observed station pressure and p 0 is mean sea level pressure. The thickness of the ozone layer is derived from the measured intensity ratio of at least two wavelengths, denoted l and l 0, which are scattered and absorbed differently in the atmosphere. For ozone observations made on single pair wavelengths the general data reduction equation is x ¼ log I 0=I0 0 log I=I 0 ðb b ð Þ 0 Þm p p 0 ðd d0þsec Z ða a 0 Þm ða a 0 Þm ða a 0 Þm ; ð1þ where x is total ozone in cm at STP (Standard Temperature and Pressure). The values log(i/i 0 ) and log(i 0 /I 0 0) are often denoted L and L 0, respectively, and are related to the dial reading R through calibration charts (N tables) or polynomials characteristic for the specific instrument. By defining N = 100(L 0 L), the thickness x of the ozone layer can be expressed in Dobson units (DU): x ¼ 10N ða a 0 A : ð2þ Þm The last two terms in equation (2) are denoted atmospheric corrections, where A = 1000(b b 0 ) m( p/p )/(a a 0 )m and the last term = 1000(d d 0 ) secz/(a a 0 )m represents the aerosol correction Direct Sun Measurements in the 1940s [11] The old Dobson spectrophotometer number 8 at Dombås was a photoelectric double monochromator. Compared to modern Dobson instruments equipped with photomultipliers, the old detector device had relatively low response to UV signals shorter than approximately 310 nm.

3 SVENDBY: TOTAL OZONE IN SOUTHERN NORWAY, ACL 3-3 discussed by Karandikar [1948] and Ramanathan and Karandikar [1949]. They found that on hazy days during the cold winter period equation (3) would in general underestimate the ozone values, whereas the ozone values normally were overestimated on hazy days in the hot season. Figure 1. Slit functions for the modernized D83 (solid line) and the old D8 (dotted line) Dobson spectrophotometers: (a) S2 slit functions for the short wavelengths and (b) S3 slit functions for the longer wavelengths. According to the D8 registration forms three different wavelength selections were possible, centered at l = nm, l 0 = nm and l 00 = nm. Thus two different wavelength pairs could be selected from the spectrophotometer: the C pair (l = nm and l 0 = nm), and the C 0 pair (l 0 = nm and l 00 = nm). [12] Calculations of total ozone at Dombås in the 1940s were based on Dobson s original formula where it was believed that only small errors in total ozone were introduced by assuming that the entire atmospheric scattering could be described as the inverse fourth power of the wavelength. The ozone values were calculated from the expression x ¼ L 0 L K L 0 L 0 0 =mp; ð3þ where K =(l 4 l 0 4 )/(l 0 4 l 00 4 ) = and P =(a a 0 ) Ka 0 = The values L 0 and L 0 0 represent log(i 00 /I 0 ) and log(i 00 0/I0), 0 respectively. The error in total ozone caused by the simplified scattering terms in equation (3) is 3. Reanalysis Method 3.1. Calculations of Absorption Coefficients [13] In Dobson spectrophotometer number 8 was sent to Oxford for a modernization [Langlo, 1952]. The instrument was equipped with a photomultiplier and it was rebuilt to allow for ozone registrations with other than the C and C 0 wavelength pairs. Such modifications might influence the absorption and Rayleigh scattering coefficients for the C(Direct Sun) ozone calculations. Modernized Dobson instruments are normally comprised of the slits S2 with slit widths of 0.40 mm ± 0.01 mm and S3 with width of 1.20 mm ± 0.02 mm [Komhyr et al., 1993]. The two slits are used for selecting the C wavelengths pair, the short wavelength l = nm and the long wavelength l 0 = nm. In contrast, the original D8 instrument had corresponding slit widths of 0.62 mm and 1.20 mm [Dobson, 1931], providing wavelength selections l = nm and l 0 = nm, respectively. The wavelength shift represents a significant change in the absorption and scattering coefficients for Dobson spectrophotometer number 8. Thus when reanalyzing the old D8 ozone data we have created a new set of coefficients that are more in line with the original instrument. As a quality check we have attempted to reproduce the 1992 Bass-Paur coefficients to test the reliability of our calculation method. [14] The wavelengths l and l 0 for the C wavelength pair represent center wavelengths for the radiation passing through the instrument slits. When calculating Dobson absorption and scattering coefficients, however, the entire slit functions must be taken into account. In Figures 1a and 1b the slit functions for S2 and S3, respectively, are presented. The slit functions for the modernized Dobson instruments, S D83 (l) and S 0 D83(l), are determined experimentally for World Primary Standard Dobson spectrophotometer number 83 [Komhyr et al., 1993]. Unfortunately, detailed descriptions of the slits in the old instrument, S D8 (l) and S 0 D8(l), are not available. To overcome this problem we have assumed that the slit functions for the original D8 instrument can be derived from D83, provided that the center wavelengths and widths are in accordance with the information described above (summarized in Table 2). Table 2. Slit Widths and Center Wavelengths for the Old Dobson Spectrophotometer Number 8 and the World Primary Standard D83 Instrument Width, mm Slit S2 l (Center), nm Width, mm Slit S3 l 0 (Center), nm Modern Dobson (D83) 0.40 ± ± Old Dobson (D8) 0.62 ± ±

4 ACL 3-4 SVENDBY: TOTAL OZONE IN SOUTHERN NORWAY, Figure 2. Absorption cross sections (solid line) and Rayleigh scattering cross sections (dashed line) in the (a) short wavelength region and in the (b) longer wavelength region. [15] Calculations of ozone absorption coefficients for the Dobson spectrophotometers are based on the following equation: Z AðlÞSðlÞIðlÞdl a ¼ Z ; ð4þ SðlÞIðlÞdl where A(l) is the spectral ozone absorption coefficients (see Figures 2a and 2b) determined by Bass and Paur at a temperature of 45 C[Bass and Paur, 1985; Paur and Bass, 1985], adjusted to temperature 46.2 C according to the temperature dependence described by Mateer [Komhyr et al., 1993]; S(l) represents the Dobson instrument slit function; and I(l) is the solar spectrum at the ground, calculated from a radiative transfer model [Stamnes et al., 1988; Dahlback and Stamnes, 1990] at air mass m = 2 and total ozone of 325 DU. Equation (4) is used for calculating both a and a 0. When referring to a 0 the corresponding slit function is denoted S 0 (l). The indices D8 and D83 will be used for describing the coefficients for the old D8 and modernized D83 Dobson instruments, respectively. [16] The results calculated from Dobson ozone absorption coefficients in equation (4) are presented in Table 3. The ozone absorption coefficients for the C wavelength pair adopted for use with all modern Dobson instruments from 1 January 1992 [Komhyr et al., 1993] are included in Table 3 as a comparison. [17] The recalculated value (a a 0 ) D83 = presented in Table 3 is in good agreement with the 1992 absorption coefficients adopted for use by the WMO, implying that our method of calculation is reliable. By assuming that the slit functions S D8 (l) and S 0 D8(l) illustrated in Figures 1a and 1b are realistic, the absorption coefficients (a a 0 ) D8 = should be adopted for Dobson spectrophotometer D8. [18] As seen from Table 3 the uncertainty of (a a 0 )is relatively large. In reality the uncertainty reflects the sensitivity to changes in center wavelengths and bandwidths (shifting the center wavelength by 0.1 nm and adjusting the slit width by 0.01 mm). When the slit function S D8 (l) is shifted 0.1 nm towards longer wavelengths, i.e., the center wavelength is moved to nm, the Dobson absorption coefficients is reduced from to This represents up to 2% increase in calculations of total ozone. The effect of adjusting the slit widths by 0.01 mm is insignificant compared to the uncertainty associated with the center wavelength. The sensitivities in a and a 0 are in accordance with tests performed by Basher [1980] Calculations of Scattering Coefficients [19] The slit-weighted Rayleigh scattering coefficients are derived from the Rayleigh scattering cross sections of Bates [Bates, 1984; Komhyr et al., 1993]. They are calculated in a similar way as the absorption coefficients described in the previous section: Z BðlÞSðlÞIðlÞdl b ¼ Z : ð5þ SðlÞðlÞdl Here B(l) are the spectral scattering coefficients (see Figures 2a and 2b). Results from calculating Dobson Rayleigh scattering coefficients using equation (5) and slit functions S D83 (l), S D8 (l), S 0 D83(l) and S 0 D8(l) from Figures 1a and 1b, are presented in Table 4. The Rayleigh scattering coefficients for the C wavelength pair adopted for use with all modern Dobson instruments from 1 January 1992 [Komhyr et al., 1993] are also included in the Table 4. [20] The calculated value (b b 0 ) D83 is exactly the same as the coefficient adopted for use by the WMO [Komhyr et al., 1993]. The implication here again is that the scattering Table 3. Recalculated Dobson Absorption Coefficients and Standard Coefficients Recommended by the World Meteorological Organization (WMO) From 1992 Instrument a, a 0, (a a 0 ), (atm cm) 1 (atm cm) 1 (atm cm) 1 Recalculated D8 values ± ± ± Recalculated D83 values ± ± ± WMO

5 SVENDBY: TOTAL OZONE IN SOUTHERN NORWAY, ACL 3-5 Table 4. Recalculated Rayleigh Scattering Coefficients and Standard Coefficients Recommended by WMO From 1992 Instrument b, atm 1 b 0,atm 1 (b b 0 ), atm 1 Recalculated D8 values ± ± ± Recalculated D83 values ± ± ± WMO coefficient calculated for the old instrument, namely, (b b 0 ) D8 = 0.100, can be relied on. The uncertainty or sensitivity in (b b 0 ) is calculated in the same way as the uncertainty of the absorption coefficients. [21] As seen from Tables 3 and 4 the absorption coefficients for D8 (developed for the original D8 slits) is 2.6% higher than the 1992 Bass-Paur coefficients, whereas the Rayleigh scattering coefficient is 9.0% lower. For air mass 1.5 and ozone 300 DU the D8 absorption coefficients yield total ozone values that are approximately 2.8% lower than values calculated from the WMO 1992 coefficients. On the other hand the 9.0% reduction of the Rayleigh scattering coefficients for D8 increases the total ozone values by 2.9%. Consequently, the overall effect of introducing the new set of scattering and absorption coefficients for D8 is very small. On average, when recalculating the entire data set from Dombås , the D8 coefficients increase the annual mean total ozone values by 0.4% compared to ozone calculations from the WMO 1992 coefficients. The difference varies from 0.5% to 2.3% depending on air mass and total ozone Aerosol Scattering [22] The reevaluated D8 ozone values are calculated from equation (1) using the C wavelength pair. One important source of error is associated with aerosols in the atmosphere. The aerosol amount will in general vary, and consequently the aerosol term in equation (2) will also vary. A method to determine based on L 0 measurement is described below. A very similar method was suggested by Ramanathan and Karandikar [1949]. [23] The measured intensity ratio (I 0 /I 00 )forl 0 = nm and l 00 = nm wavelength channels can, by Lambert- Beer s law, be written as logði 0 =I 00 Þ ¼ log I =I0 a 0 mx ðb 0 b 00 Þm p ðd 0 d 00 Þsec Z; p 0 yielding ðd 0 d 00 Þ ¼ L 0 L 00 0 a0 mx ðb 0 b 00 Þm p cos Z; p 0 where L 0 = log(i 0 /I 00 ) and L 0 0 = log(i 0 0/I 00 0). The spectral dependence of aerosol optical depth, t, can be described by the Ångström formula t = bl n, where l is the wavelength and b and n are two characterizing parameters. From the Ångström formula the aerosol scattering can be expressed as ðd 0 d 00 Þ ¼ Kðl 0 n l 00 n Þ; where K is a constant proportional to the aerosol concentration. By combining the two expressions (d 0 d 00 ) the constant K can be expressed as (L 0 L 0 0 a 0 mx (b 0 b 00 )m( p/p 0 )) cos Z/(l 0 n l 00 n ). Thus the aerosol scattering term for the C pair is ðd d 0 Þ ¼ L 0 L 0 0 a0 mx ðb 0 b 00 Þm p cos Z p 0 l 0 n l 00 n ðl n l 0 n Þ: The Rayleigh scattering coefficient b 00 is calculated from the weighting procedure described in section 3.2 for a slit width of 0.50 mm [Dobson, 1931] and center wavelength nm. The following value is obtained: b 00 ¼ 0:101: By using the coefficients a 0 and b 0 in Tables 3 and 4, respectively, the final aerosol scattering term can be expressed: ðd d 0 Þ ¼ C L 0 L 0 0 0:056mX 0:251m p cos Z; ð6þ p 0 where C ¼ ðl n l 0 n Þ= ðl 0 n l 00 n Þ: [24] As demonstrated by Kylling et al. [1998] the coefficient n might vary from zero to around 1.4 within a couple of days. Large values of n may be interpreted as the presence of small aerosol particles, while small values of n are due to large aerosol particles. According to Ramanathan and Karandikar [1949] negative n values can also be observed. However, for calculating daily values of ozone, it was assumed that the particle scattering was nearly neutral and thus varied as l n with n = 0, yielding C = This C value is used when recalculating the Dombås data. [25] From equation (6) it is obvious that the ozone value x needs to be known in order to find the exact aerosol correction. However, relatively small errors in (d d 0 ) are introduced if x is within ±3% of the true ozone value. Thus total ozone is first calculated from the C(DS) method described in equation (2), ignoring the aerosol term. Secondly, the preliminary ozone value is used to find the term (d d 0 ) from equation (6), and finally the aerosol corrected ozone value is recalculated from equation (2) including the aerosol term. The aerosol scattering term,, will normally reduce the preliminary ozone value by 0 4%, depending on the season. Variations in aerosol scattering will be described more closely in section 3.6. It should also be noted that a relatively large uncertainty is associated with the extraterrestrial constant L 0 0 in equation (6). This will be discussed in the next section Extraterrestrial Constant and L Tables [26] Prior to calculations of total ozone, Dobson spectrophotometers have to be calibrated. The optical wedges are examined, normally using a two-lamp method [Dobson, 1957], enabling the dial reading R to be converted to the quantity G = log(i/i 0 )+C, where C is an unknown constant and from which L can be calculated using the relation L = log(i/i 0 )= log(i 0 /I). Finally, observations performed during clear days with constant ozone values and varying solar elevations offer information about the extraterrestrial constant L 0. The L values are plotted against air mass m, and the best fit line through these points intercept the ordinate

6 ACL 3-6 SVENDBY: TOTAL OZONE IN SOUTHERN NORWAY, axis (m = 0) at the point L 0 = log(i 0 /I 0 0), where I 0 and I 0 0 represent solar intensities outside the atmosphere. This Langley plot method was used to determine the extraterrestrial constants L 0 and L 0 0 in the 1940s. [27] An examination of the L tables used for the old D8 instrument shows that the conversion from R to L values can be described by the polynomial where L ¼ k 1 R 3 þ k 2 R 2 þ k 3 R þ k 4 ; k 1 ¼ 18: ; k 2 ¼ 7: ; k 3 ¼ 9: ; k 4 ¼ 3:1603: In other words, a dial reading R yields an intensity ratio L given by equation (7). [28] The extraterrestrial constants L 0 = and L 0 0 = were originally used at the Dombås observatory. We do not know exactly when and where these constants were derived, but the fact that the coefficients were adopted immediately after the arrival of D8 at Dombås implies that they were constructed prior to 1940, most likely by G. M.B. Dobson at Oxford. Thus we have used the original data set to derive new extraterrestrial constants that hopefully are more representative for ozone retrievals at Dombås. Studies of eight different Langley plots in 1940 to 1942 result in extraterrestrial constants with mean and standard deviation: L 0 ¼ 2:946 0:004 L 0 0 ¼ 1:541 0:020: [29] In addition to the standard deviation of L 0 and L 0 0,an error is introduced to the Langley plots by neglecting the term (b b)m and assuming that m equals m during the course of the day. By replacing m by m the extraterrestrial constant calculated from the Langley plot method typically decreases by As a rough estimate we have assumed that the error associated with the (b b)m term amounts to half this value (i.e., 0.006) which gives overall uncertainties of and in L 0 and L 0 0, respectively. This uncertainty represents up to a 3.5% error in calculations of total ozone Calibration Check [30] Modern Dobson spectrophotometers will normally undergo regular calibration checks and comparisons with other Dobson spectrophotometers to ensure the stability of the instruments. After such intercomparisons the individual G tables are transformed to N tables to obtain identical total ozone values from all Dobson spectrophotometers. However, intercomparisons and calibrations checks were not as frequent in the 1940s as they are today. [31] As demonstrated in section 3.4 the extraterrestrial constants for D8 are characterized by relatively large uncertainty. Also, the Langley plot method presupposes that the ozone and haze conditions remain constant during the measuring period (i.e., half a day). In practice this is usually not the case. In addition, instrumental drifts might affect the ð7þ instrumental readings and give inaccurate N tables. To check the calibration reliability of Dombås data we have used the statistical method described by Dütch [1984]: On average, we would expect the ozone value x to be independent of the air mass m. If we assume that the N tables, N = 100(L 0 L), are correctly determined the average difference between ozone values obtained at noon and in the morning (or late afternoon) should be zero: X ð xi1 x i2 Þ ¼ 0: ð8þ The suffixes 1 and 2 refer to observations with high Sun (small m) and low Sun (large m) on day i. The sum in equation (8) is taken over a suitable number of days. If the summation differs significantly from zero it indicates inaccuracies in the N table. Assuming that the true N value is given by N 0 = N + N*, the correction N* may be found by substituting equation (2) into equation (8) using N i1 + N* and N i2 + N* as the true N values. Rearrangement of equation (8) gives X Ni1 =m i1 N i2 =m i2 þ ~a ð Ai2 A i1 Þ N* ¼ X ; ð9þ 1=mi1 1=m i2 where ã = 0.1(a a 0 ) and A is the atmospheric correction given by the last two terms of equation (2). [32] For the Dombås data set the N* corrections are calculated for all relevant days from 1940 to 1946 (i.e., all days where total ozone data exist for both noon and morning/ evening). The values N i1 and N i2 in equation (9) are based on the extraterrestrial constants L 0 = 2.946, L 0 0 = and the L polynomial of equation (7). Hence a possible N* adjustment will compensate for small discrepancies in L 0. As explained by Dütch [1984] the N* correction will be a good substitute for the standard lamp tests used today. [33] The calibration check performed at Dombås could not detect any systematic errors in the N values (positive or negative) for the period 1940 to At 95% confidence level the N* corrections of the N table are within ±1.0, which represent a calibration uncertainty of approximately 3% total ozone Recalculated Total Ozone Values From 1940 to Direct Sun Observations, [34] The coefficients needed for calculating C(DS) total ozone from equations (2) and (6) are summarized below: ða a 0 Þ ¼ 0:854; ðb b 0 Þ ¼ 0:100; L 0 ¼ 2:946; L 0 0 ¼ 1:541; L ¼ 18: R 3 7: R 2 9: R þ 3:1603; p 0 =p ¼ 0:925: The Dombås observatory is located 659 m above sea level where the pressure normally is considerable lower than the

7 SVENDBY: TOTAL OZONE IN SOUTHERN NORWAY, ACL 3-7 Figure 3. Monthly average aerosol corrections at Dombås. The corrections represent the difference between Direct Sun measurements calculated with and without the aerosol correction term,. The solid line represents corrections for neutral particle scattering (n = 0), whereas the dashed line is calculated for n = 1.4 (see equation (6) in the text). normal sea level pressure. Thus the ratio p 0 /p = is used for all the direct Sun ozone calculations at Dombås. [35] As discussed in section 3.1 to 3.5 there will always be some uncertainty associated with the individual coefficients listed above. The largest uncertainties are normally related to the extraterrestrial constants and the N tables used for D8. However, the calibration check discussed in section 3.5 reveals no serious calibration errors in the old D8 Dobson spectrophotometer. Also, the aerosol corrections in equation (6) have demonstrated that L 0 0 hardly deviates very much from Large errors in L 0 0 would normally yield unrealistically high or low aerosol corrections, which is not the case for the D8 calculations from Dombås. Altogether we believe that the uncertainties in total ozone values are within ±3%. [36] Figure 3 shows aerosol corrections calculated from equation (6). The corrections are calculated for two different values of C (i.e., n values). The solid line represent the standard method where we assume that the aerosols scatter neutral (n = 0 and C = 0.165), whereas the dashed line represent aerosol corrections with n = 1.4 and C = By using n = 1.4 instead of n = 0.0 the total ozone value are reduced by up to 1.6%. However, on average the difference amounts to 0.6% total ozone. [37] Figure 3 also shows that the aerosol corrections normally are largest in the summertime, when the monthly mean total ozone values are up to 3% lower than C(DS) measurements calculated without aerosol corrections (4% if n = 1.4 is used). Studies of day to day variations in the aerosol corrections at Dombås reveal that might exceed 20 DU on hazy days, whereas slightly negative aerosol corrections occasionally are observed on clear winter days. When reducing data from equation (6) the negative are taken to be equivalent to zero. The aerosol corrections calculated at Dombås show large similarities to the aerosol corrections obtained for the modernized D56 instrument that operated in Oslo from 1978 to 1998 [Svendby and Dahlback, 2002]. [38] The original direct Sun total ozone values retrieved at Dombås were much lower than ozone values computed from equation (2) with adoption of new absorption and scattering coefficients. Figure 4 shows the difference between reanalyzed total ozone data and the original Dombås values from the 1940s. The large data spread is mainly caused by the aerosol correction which reduces the individual reanalyzed C(DS) measurements by 0 8%. From Figure 4 no seasonal pattern can be seen for the conversion of original ozone data to the reevaluated Bass-Paur scale. The average conversion factor (and standard deviation) is C ¼ 1:402 0:023: ð10þ The corrections of for converting D8 measurements from Ny-Choong [Ny and Choong, 1933] to the Bass-Paur scale is slightly lower than the corrections reported by, for example, Brönnimann et al. [2003a]. However, the Figure 4. Ratio of reanalyzed and original total ozone values at Dombås, The average ratio (solid line) is

8 ACL 3-8 SVENDBY: TOTAL OZONE IN SOUTHERN NORWAY, Figure 5. Comparison of reanalyzed monthly mean total ozone values from CC 0 Direct Sun and CC 0 Zenith Sky measurements performed at Dombås, difference between recalculated and original ozone values increases to ± if the aerosol corrections are ignored Zenith Sky Measurements [39] About 52% of the total ozone measurements performed at Dombås are based on the CC 0 zenith sky observations. During the winter months the instrumental signals are very weak due to the low Sun and a limited amount of direct Sun observations can be carried out. In order to obtain ozone information for the winter period and cloudy days one has to rely on zenith observations. However, the calculation of CC 0 zenith ozone values will to some extent depend on the height, thickness and type of clouds [Komhyr, 1980]. The observations were originally adjusted according to these parameters, and subjective judgment was a part of the final calculation of the most reliable zenith ozone values. When reanalyzing the old ozone measurements it is very difficult to attain detailed information on the atmospheric condition, and consequently all the zenith observations have been recalculated using the correction function in equation (10). [40] Figure 5 shows monthly mean difference between the recalculated direct Sun and zenith sky ozone observations at Dombås. The mean and standard deviation of the difference is 0.60 ± 1.51%. By ignoring the first two years of the observational period (with extraordinary cold atmospheric conditions and relatively high aerosol corrections) the agreement improves significantly and the average difference between C 0 (DS) and CC 0 (ZS) measurements becomes less than 0.3%. Owing to the limited number of C direct Sun measurements from November until January no comparison exists for these winter months Total Ozone Values From July 1946 to April 1949 [41] As mentioned previously D8 was moved to Oslo (59.9 N, 10.7 W, 96 m) in the end of June 1946 where Kaare Langlo continued the systematic D8 ozone registrations. Unfortunately the raw data for the entire Oslo period are lost and only the original monthly mean total ozone values are available, summarized in the paper by Langlo [1952]. In order to obtain ozone values in line with the Bass-Paur 1992 coefficients, the correction factor in equation (10) is applied to the entire D8 data set from July 1946 to April [42] It should be mentioned that the total ozone values from Oslo probably were influenced by high levels of SO 2 pollution. When systematic SO 2 measurements were established in Oslo in 1959 the average SO 2 winter values were about 300 mg/m 3 in downtown Oslo [Hunnes, 1998]. Burning of fossil fuels and coal firing in private homes and industry were the main SO 2 sources. Thus it is reasonable that high SO 2 concentrations also were present ten years earlier when D8 operated in Oslo. As demonstrated by de Muer and de Backer [1992] SO 2 might influence the Dobson AD(DS) measurements and lead to overestimation of the total ozone values. The reanalyzed D56 Oslo data from 1978 to 1998 [Svendby and Dahlback, 2002] show that the average total ozone winter values in the late 1970s were overestimated by 6 ± 2 DU due to SO 2 concentrations of approximately 70 mg/m 3. [43] According to the SO 2 absorption cross sections from, for example, Vandaele et al. [1994] it is obvious that C(DS) measurements are influenced by atmospheric SO 2 pollution to a similar extent as AD(DS) total ozone measurements. The same conclusion can be drawn from a paper by Komhyr and Evans [1980]. The SO 2 absorption spectrum is essentially zero for wavelengths larger than 325 nm. Consequently, only solar intensity at the shortest wavelength will be notably influenced by SO 2 in the atmosphere. De Muer and de Backer [1992] showed that the measured SO 2 ground concentrations will overestimate Dobson total ozone values, x, by the following amount: x ¼ CðSO 2 Þ1:4B=d; ð11þ Here C(SO 2 ) is the measured SO 2 ground level concentration in mg/m 3 and B/d is a conversion factor depending on the season. For the Oslo station the B/d values are estimated to ± and ± milli-atm cm mg 1 m 3 for the winter and summer, respectively [Svendby and Dahlback, 2002]. [44] From equation (11) it can be concluded that the SO 2 concentrations monitored in Oslo in the late 1950s (approx-

9 SVENDBY: TOTAL OZONE IN SOUTHERN NORWAY, ACL 3-9 imately 300 ± 40 mg/m 3 in the winter months October March and 150 ± 20 mg/m 3 in the summer period April September) would overestimate the ozone measurements by 20 ± 9 DU and 14 ± 3 DU for the winter and summer seasons, respectively. However, the Norwegian Meteorological Institute (University of Oslo) where D8 measurements were performed in the late 1940s is located on the outskirt of Oslo city center and will normally not be as heavily polluted as the downtown area where SO 2 was monitored. Recent studies of the reanalyzed D56 measurements from Oslo [Svendby and Dahlback, 2002] suggest that the SO 2 level at the University of Oslo can be estimated as the average SO 2 concentrations measured within Oslo, i.e., approximately 2/3 of the downtown SO 2 values. By using a similar approach for the period and anticipating that the SO 2 winter and summer concentrations at the University were 200 ± 26 mg/m 3 and 100 ± 13 mg/m 3, respectively, the overestimate of Dobson total ozone measurements will be in the order of 13 ± 7 DU and 9 ± 2 DU for the winter and summer seasons, respectively. Of course large uncertainties are associated with the actual SO 2 levels. [45] Owing to the issues discussed above, i.e., the absence of daily ozone data and a possible influence of SO 2 pollution, the total ozone data from Oslo for the period July 1946 to April 1949 are believed to suffer from a relatively large uncertainty (probably up to 5%) and the data set should be treated with extra care when compared to other series of total ozone. The error is expected to lead to a positive bias rather than a negative one, i.e., the total ozone values obtained could be somewhat too high. 4. Results 4.1. Ozone Measurements in the 1940s [46] In the 1940s a few other European stations performed total ozone measurements, e.g., Arosa, Switzerland (46.8 N, 9.7 E), Tromsø, Norway (69.7 N, 18.9 E), Århus, Denmark (56.3 N, 10.6 E), Oxford, UK (51.8 N, 1.2 W), and Potsdam, Germany (52.4 N, 13.1 E). Total ozone values from the Norwegian Tromsø station for the period 1935 to 1972 are under revision and will be published shortly (T. Svenøe, personal communication, 2003). In Arosa, continuous ozone measurements have been performed since When one instrument has been replaced by another, overlapping ozone measurements have ensured a consistent and homogeneous data set [Staehelin, 1998a]. However, in the period after 1949 a break needed to be fixed in the series of Arosa. Also, total ozone records from Oxford, Potsdam, and Århus have recently been examined [Brönnimann et al., 2003b] and might be suitable for comparisons with the Dombås ozone series. [47] In Oxford total ozone observations were resumed in the mid-1930s and continued until However, with several long gaps and periodically infrequent observations, a very limited number of observations are available for comparison with the Dombås data. The same is the case for the Potsdam ozone series. From 1940 until 1946 only 7 months with overlapping Potsdam/Dombås data exist (3 months in 1942 and 4 months in 1943). [48] Århus represents the pre-igy ozone station closest to Dombås and Oslo. In Århus frequent total ozone observations were performed from 1940 to May 1944 and altogether 30 months with overlapping Dombås and Århus measurements exist. No data sheets from the Dobson spectrophotometer at Århus were recovered prior to the reevaluation [Brönnimann et al., 2003a], and the monthly mean values were derived from figures in Moser [1949] and Langlo [1952]. As explained by Brönnimann et al. [2003a] the absolute values of the reanalyzed Århus data are very uncertain and, consequently, not ideal for evaluating the reliability of the Dombås total ozone series. Nevertheless, the original monthly ozone mean values from Dombås and Århus show a large degree of consistency. On average (for the months of overlapping measurements) the Århus data are 5 DU higher than the Dombås total ozone mean. As a comparison, study of TOMS (Total Ozone Mapping Spectrometer) version 7 overpass data for the period 1979 to 2002 shows that Århus annual ozone mean on average is 4 DU higher than registered at Dombås, meaning that the original Dobson ozone measurements internally agree to within 0.5% compared to recent TOMS overpass data. The TOMS values are based on data from two different satellites: Nimbus 7 operating from 1 November 1978 to 6 May 1993 and Earth Probe operating from 17 July 1996 and onwards (see TOMS homepage gov for more information). The month of December is not included in the Århus/Dombås comparison due to lack of TOMS overpass data from Dombås. [49] Owing to the relatively short Århus series and the uncertainty in the absolute level of the recalculated data we have chosen to concentrate on the comparisons with the old measurements from Arosa in the 1940s. Figure 6 shows annual mean total ozone values from Dombås/Oslo and Arosa from 1940 to The lower two curves represent the original average ozone values published by Langlo [1952]. Only months where simultaneous Dobson/Oslo and Arosa observations exists are included in the annual means, thus monthly average ozone values from December, January and February are lacking some of the years. The annual ozone differences agree to within ±5% for all years from 1940 to The upper two curves in Figure 6 represent the annual means for the reanalyzed data. On average the conversion from Ny-Choong to Bass-Paur scale yields 0.5% higher ozone values for Dombås than for the Arosa record. This is a satisfactory agreement keeping in mind that different reanalysis methods are used. [50] From Figure 6 it can be seen that the D8 ozone series increases after 1946 and exceeds the Arosa annual means, i.e., at the same time as D8 was moved from Dombås to Oslo. By studying TOMS satellite data from 1979 to 2002 one would normally expect the annual Oslo data to be approximately 1% higher than the Dombås data. However, this difference can hardly explain the major jump after 1946 where D8 is more than 15 DU (5%) high compared to the period. It cannot be excluded that the jump results from calibration errors, instrumental problems or new measuring methods introduced by the observer in Oslo. However, another possible explanation for the sudden increase in the D8 total ozone values is the SO 2 pollution level in the late 1940s. As demonstrated in section 3.6 the pollution might have caused a significant overestimation of total ozone from July 1946 to April It should be emphasized that the method of using TOMS

10 ACL 3-10 SVENDBY: TOTAL OZONE IN SOUTHERN NORWAY, Figure 6. Comparison between ozone values from Arosa and Dombås, (annual means). The lower two curves represent the original values, whereas the upper two curves are reanalyzed data. observations to derive latitudinal gradients for the ozone measurements during the 1940s is uncertain. The total ozone values from Arosa in the early 1940s are the largest ever observed, suggesting that extraordinary dynamical conditions existed [Brönnimann et al., 2000]. Thus meteorological anomalies in the 1940s might have influenced the latitudinal ozone gradient all over Europe. It should also be mentioned that TOMS has retrieval problems over snow/ice and tends to overestimate total ozone under these conditions. [51] The reanalyzed monthly mean ozone values from the Norwegian and Swiss stations are presented in Figure 7. As expected the seasonal pattern differs for the two stations - the most northern stations having the largest month to month variations. The high Oslo values after 1946 can clearly be seen from Figure 7. Except for 1946 the average winter and spring ozone values measured in Oslo are considerably higher than the Arosa values. [52] On the basis of the present comparison between Dombås/Oslo and Arosa total ozone measurement one can not draw any conclusion about the absolute level of the two ozone series. Calculations of annual ozone means from 1940 to 1946 show that the reanalyzed Dobson values from Dombås on average are 0.9 ± 0.7% lower than the Arosa measurements (0.3% if 1940 is omitted from the comparison). Contrary the annual mean ozone values from TOMS are 1.8 ± 0.4% higher for Dombås than Arosa, i.e., the absolute level of the old Dombås and Arosa series deviates by about 2 ± 1% compared to the ozone difference anticipated from TOMS overpass data for recent years. It should be emphasized that total ozone values from TOMS cannot be compared directly to the old Dobson measurements, partly because the data set consists of relatively fewer monthly mean values from the winter period than is the case for the TOMS overpass data. It should also be emphasized that the latitudinal ozone distribution might perfectly well have changed during the two time periods under consideration. However, another possible explanation for the observed difference between the Dombås and Arosa annual averages is errors in the absolute level of the ozone measurements, e.g., due to calibration inaccuracies, different measuring methods used, or interference from atmospheric aerosols Comparison to Present Ozone Values [53] The D8 data set contains two ozone series with different uncertainties: (1) data from the early 1940s which probably are most affected by the particular dynamical conditions [Brönnimann et al., 2000] and (2) the Oslo series being most affected by uncertainty related to SO 2 interference and the fact that only monthly mean ozone values exist. In order to study whether the ozone layer has changed from the 1940s until today, we have paid most attention to the Dombås series and compared the data to D56 ozone measurements from Oslo 1978 to Even if the Arosa comparison and the reanalysis method have demonstrated that the absolute level of the reanalyzed Dombås series is uncertain to within ±3%, there is no reason to question the relative ozone values (i.e., the seasonal ozone variation) obtained from the recalculation. Hence the general ozone situation in southern Norway in the 1940s can be well described and deemed appropriate for comparison with modern measurements. [54] Dombås is located approximately 260 km north of Oslo and consequently the normal annual cycle of total ozone might be different for the two locations. Hence total ozone data from Dombås and Oslo have been examined from TOMS version 7 satellite data for the period January 1979 to November Data from TOMS, presented in Figure 8 and Table 5, show that the annual ozone cycles in Oslo and Dombås are not identical. December is not included in the comparison due to lack of TOMS overpass

11 SVENDBY: TOTAL OZONE IN SOUTHERN NORWAY, ACL 3-11 Figure 7. Monthly mean total ozone values from Arosa and Dombås, data from Dombås. The average ozone value for the first 4 months of the year (January to April) is more or less identical for the two stations (367.8 ± 3.8 DU and ± 3.6 DU for Dombås and Oslo, respectively), contrary to the summer period May August where the Oslo ozone values are approximately 1.5% higher than the Dombås values. Normally, total ozone values from TOMS yield higher values than retrieved from Dobson measurements [Fioletov et al., 1999]. Consequently, it is preferable to base our analysis on Dobson spectrophotometers alone when studying possible changes in the ozone layer over a long time period. [55] In order to study the ozone layer from the 1940s and make comparisons to present years we have divided the D56 series into two time periods: a) , referred to as the 1980s, and b) , referred to as the 1990s. The mean summer and winter total ozone values for these time periods are summarized in Table 5. [56] The D56 summer ozone values from 1978 to 1987 have mean and standard error amounting to ± 2.9 DU. In contrast, the recalculated D8 summer mean at Dombås is ± 3.2 DU for the period 1940 to 1946 (see Table 5), i.e., the D56 summer values from 1978 to 1981 is 2.0% higher than the D8 measurements at Dombås. From TOMS overpass data we would expect 1.5% difference, suggesting that the summer ozone values were approximately 0.5% lower than the values measured from 1978 to For the winter months an insignificant increase of 0.2% is observed. It is well known that the annual ozone layer can change from one time period to another and fluctuations up to ±2% are not exceptional [Bojkov and Fioletov, 1995]. Thus it is not unreasonable that the summer ozone values have increased insignificantly from the 1940s until On the other hand it is meaningless to draw conclusions about ozone changes of % when the overall uncertainty in the Dombås data amounts to ±3%. Figure 8. Annual variations of TOMS total ozone overpass data, , for Oslo (solid line) and Dombås (dashed line).

12 ACL 3-12 SVENDBY: TOTAL OZONE IN SOUTHERN NORWAY, Table 5. Average Summer and Winter Total Ozone Values (±Standard Errors) From Dombås and Oslo Station and Period Winter/Spring Values, Jan. April, DU Summer Values, May Aug., DU TOMS Dombås ± ± 2.4 TOMS Oslo ± ± 2.4 D8 Dombås, ± ± 3.2 D8 Dombås/Oslo, ± ± 3.2 D56 Oslo, ± ± 2.9 D56 Oslo, ± ± 3.0 It should also be mentioned that the observed total ozone increase can be related to enhanced levels of tropospheric ozone. Measurements from Arosa [Staehelin et al., 1998b] demonstrate that large increases in tropospheric ozone took place between the end of the 1950s and the middle of the 1980s, probably around 10 DU. This aspect can be relevant when comparing data from the 1940s and the 1980s. [57] Even if no significant total ozone changes can be observed from the 1940s to the 1980s, the picture is far more dramatic when the D8 Dombås data are compared to ozone values from the 1990s. From Figure 9 (studied relative to Figure 8) it can be seen that the annual cycle in the ozone layer has changed during the last years, the changes being most pronounced for the winter and spring months. For the winter period an ozone decrease of 6.1 ± 3.9% is observed whereas the ozone decrease amounts to 2.9 ± 1.8% for the summer period. The summer ozone decrease is calculated relatively to the 1.5% Oslo-Dombås difference anticipated from TOMS overpass data. The calculated total ozone changes (presented in Table 6) do not take any possible influence of solar activity, volcanic aerosols or meteorological variations into consideration. The calculations do neither take the probable influence of tropospheric ozone into account. By anticipating that tropospheric ozone has increased by approximately 10 DU from the 1940s until the 1980s/ 1990s, the calculations of stratospheric ozone decrease is underestimated by 2 3%. Thus the winter ozone layer Table 6. Observed Total Ozone Changes for Different Time Periods and Seasons Station and Period Winter/Spring, Jan. April, % Summer, May Aug., % D56( ) D8( ) 0.2 ± ± 1.7 D56( ) D8( ) 6.1 ± ± 1.8 might have decreased by as much as 8 9% from the 1940s until the 1990s. This change is in good agreement with the 3.1 %/decade ozone winter decrease reported by the World Meteorological Organization [1999] for the period 1970 to [58] In addition to the possible influence of tropospheric ozone it should be keep in mind that the Dombås series might be 2% low (as suggested from the Arosa comparison). Consequently, it cannot be excluded that a significant ozone decline had started in the 1980s. Unfortunately we have no data from southern Norway for the years 1950 to 1978 which would make the picture more complete. Nevertheless, the observed winter/spring ozone decrease supports the theory that middle and high-latitude ozone depletion can be explained by heterogeneous processes involving halogen compounds: Heterogeneous reactions on ice crystals might facilitate conversion of inactive chlorine (or bromine) to active radicals, which subsequently destroy the ozone molecules [e.g., Solomon, 1999; Rodriguez, 1993]. 5. Conclusions [59] Total ozone measurements from Dobson spectrophotometer number 8 at Dombås and Oslo, Norway, have been recalculated. Even if detailed information about the instrumental history and calibration procedures are absent the data set is believed to be consistent and of high quality taking the existing standards into account. Calibration checks indicate that the N table to Figure 9. Annual variations of total ozone at three periods: (D8 Dombås), (D56 Oslo), and (D56 Oslo).