PSCs and heterogeneous chemistry
Normally, the stratosphere is very dry and cloudless even though a thin layer of aerosol (liquid-phase binary H2SO4/H2O droplets) is present in the lower stratosphere. But during polar night, when temperatures reach below 195 K in 15–25 km the background aerosols take up HNO3 and H2O and evolve into ternary HNO3/H2SO4/H2O droplets, referred to as PSCs (Carslaw et al., 1994). Figure 1.7 shows a photograph of the PSCs in the Antarctic stratosphere. PSCs are classified into three types, Type Ia, Type Ib and Type II, according to their physical state or optical properties and chemical composition. This classification is based on air-borne lidar measurements (lidar backscatter and depolarisation ratios for PSCs) as the instrument is sensitive to the state of polarisation of the backscattered light (Browell et al., 1990; Felton et al., 2007). Type Ia PSCs are made up of crystals of ni-tric acid trihydrate [NAT – (HNO3. 3H2O)] and Type Ib consists of supercooled ternary solutions (STS) of HNO3/H2SO4/H2O. Type II PSCs are frozen water ice non-spherical crystalline particles. During austral winter heterogeneous reactions occur on the surface of PSC particles and convert the reservoir species such as ClONO2 and BrONO2 into more active species. The principal heterogeneous reactions are given below.
H Cl(s) + ClON O2 (g) → HNO3 (s) + Cl2 (1.35).
BrON O2(g) + H2O(l) → HNO3 (g) + H OBr(g) (1.36).
H Cl(s) + BrON O2 (g) → HNO3 (g) + BrCl(g) (1.37).
H Cl(l) + H OBr(g) → H2O(l) + BrCl(g) (1.38).
Other heterogeneous reactions occurring on the surfaces of PSCs releasing Cl2 are:
ClON O2(g) + H2O(s) → H N O3(s) + H OCl(g) (1.39).
H Cl(s) + H OCl(g) → H2O(s) + Cl2(g) (1.40).
Another heterogeneous reaction releasing photolytically active Cl2 by the photolysis of ClNO2 is:
N2O5(g) + H Cl(s) → ClN O2(g) + H N O3(s) (1.41).
However, for the production of Cl atoms sunlight is required. When the sun starts to shine on the polar stratosphere at the beginning of austral spring, Cl2 is photolysed to Cl, that enters into catalytic destruction of ozone. This is the reason why ozone is depleted in spring, when sunlight returns. Thus, above reactions accumulate the ClO concentration in the polar lower stratosphere, which then initiate the ozone destruction cycle (Reactions 1.30–1.34).
Mid-latitude ozone loss
By the discovery of springtime polar ozone depletion linked to the heterogeneous reactions on the cloud surfaces, ozone amount in other latitudes were also analysed carefully and significant ozone loss over mid-latitudes were found (WMO, 1992), although not so large as in the polar regions. In the mid-latitudes also, gas phase and heterogeneous chemistry as well as the dynamical processes are tied with the ozone reduction. Regarding chemical processes, the ClOX gas phase catalytic cycle (Reactions 1.21–1.23) is most eﬀective in de-pleting ozone in the upper stratosphere (35–45 km). The heterogeneous reactions occurring on the sulfate aerosols (Tolbert, 1996) are the main source of ozone depletion in the lower stratosphere (Hofmann and Solomon, 1989) following major volcanic eruptions. The most important heterogeneous reaction occurring on sulfate aerosols is the hydrolysis of N2O5 to a stable HNO3. N2O5(g) + H2O(aerosol) → 2H N O3(g) (1.42).
This hydrolysis of N2O5 reduces the amount of NOX in the lower stratosphere and hence the impact of NOX catalytic cycle on ozone destruction (Fahey et al., 1993). In contrast, formation of ClONO2 is also decreased by the reduction of NOX , which indirectly enhances the ClOX and HOX catalytic ozone destruction cycles (Brasseur et al., 1999). Another het-erogeneous reaction predominant in the colder conditions of the mid-latitude stratosphere is Reaction 1.39. This reaction provides a path to produce Cl2 from an active HOCl (Reac-tion 1.40). Another most eﬃcient ozone loss cycle in the lower stratosphere is the BrO–ClO cycle (Reactions 1.24–1.29) (Daniel et al., 1999).
The dynamical processes also strongly influence mid-latitude ozone. They include the advection of polar air activated within the vortex mixed with the mid-latitude air and the spreading of ozone depleted air from the polar vortex to the mid-latitudes in spring, referred to as the dilution (Andersen and Knudsen, 2006). That is, when vortex breakup the filaments of ozone poor air move to the low latitudes and dilute the ozone concentration in the mid-latitudes, as demonstrated in Fig. 1.4. Hadjinicolaou and Pyle (2004) also pointed out that maximum ozone depletion in the mid-latitudes coincides with the date of polar vortex breakdown and the abundance of ozone depleted polar air in the mid-latitudes. The ozone loss in the mid-latitudes thus depends on the strength and persistence of the polar vortex and severity of ozone loss.
Equivalent Eﬀective Stratospheric Chlorine
As seen in Sect. 1.3, Cly and Bry originated from the CFCs and halons are the main cause for the stratospheric ozone depletion. The quantification of the combined impact of Cly and Bry to destroy ozone is defined as the Equivalent Eﬀective Stratospheric Chlorine (EESC). EESC = Cly + αBry (1.43).
where α is the weighting factor that accounts for the greater eﬀectiveness of Br in destructing ozone compared to that of Cl on a per-atom basis. It varies with latitude, altitude and time.
As described in Sect. 1.3, even though the amount of Bry compounds is less than that of Cly , the Br catalytic cycle is more eﬃcient in destroying ozone eventhough PSC surface is not required for the activation of Br catalytic cycle. In order to account for this, Bry contribution is scaled by a factor α (WMO, 2007). The value of column α varies in a range from 50 to 130 from the equator to the poles depending on season. The global annual mean value of α is about 66. In the mid-latitudes, α is estimated to be about 60 in the lower stratosphere and 5 in the upper stratosphere (Sinnhuber et al., 2009).
Stratospheric EESC at diﬀerent regions is calculated from the surface measurements of tropospheric ODS abundances and taking into account for the transit times (or ages) and conversion of Cly and Bry . The extent of degradation of ODSs in the stratosphere is described by considering the mean stratospheric age of an air parcel (Newman et al., 2007). The mean age of air in the lower stratosphere is ∼3 years in the mid-latitudes and ∼5.5 years in the polar latitudes (Waugh and Hall, 2002; Newman et al., 2006). Therefore, EESC values are directly linked with the ODS emissions in the atmosphere. The discovery of the cause of polar ozone depletion led to the implementation of the Montreal Protocol in 1987 for controlling the vast emission of human-produced CFCs. Studies by Rinsland et al. (2003) and Lary et al. (2007) reported that levels of Cl and Br radicals are decreasing from the past decade onwards, which is clear from the EESC values too. same, its growth and declining amount varies at diﬀerent regions, i.e., in the mid-latitudes, the EESC values are smaller than those in the polar latitudes. Because, in the mid-latitudes the amount of inorganic halogen is lower while in the polar stratosphere, as air ages ODSs photochemically decompose so that a higher fraction of inorganic halogen is available. From the Figure, it is obvious that the EESC values increased from the 1980s and showed its peak value in 1996 and 2000 in the mid and polar latitudes, respectively and started to slowly decrease afterwards. This diﬀerence in the EESC peak year at diﬀerent latitude regimes is due to the diﬀerence in the transit time or mean stratospheric age of air.
The ozone lidar system at OHP
The ozone lidar system at OHP (43.93 N, 5.71 E) uses DIAL technique for measuring stratospheric ozone since 1986. Since then, various improvements have been made in the experimental set-up among which the most important one is the implementation of a new optical and electronic detector in 1993. The schematic picture of OHP ozone lidar system is shown in Fig. 2.1 and a detailed description of the new experimental set-up is discussed in the following sections.
The lidar system includes a Lambda Physik EMG 200 excimer laser for the ozone absorbed laser radiation at 308 nm and the third harmonic of a continuum Nd:YAG laser for the reference wavelength at 355 nm. Excimer laser operates at 100 Hz and the output energy is 200 mJ/pulse. The Nd:YAG laser operates at 50 Hz and its output energy is adapted to provide a return signal equivalent to the on-line signal at 40 km altitude, which results in an emitted pulse energy of ∼60 mJ. Additionally, two beam expanders are used to reduce the divergence of both lasers to 0.2 and 0.1 mrad at 308 and 355 nm respectively.
The back scattered radiations of the emitted laser pulses is collected by an optical receiver consisting of four Newtonian telescopes having F/3 mirrors of 0.53 m diameter. These mir-rors collect a fraction of the backscattered light that is transmitted to the optical analysing device through the optical fibers mounted on the focal plane of the mirrors, which can be moved vertically for focusing and horizontally for the alignment. As the transmission of the emitted laser beams is in the center of the mirrors, the system can be considered as quasi-coaxial. The optical device includes the imaging optics, a mechanical chopper and a spectrometer. The chopper consists of a 40 W cooled motor, rotating at 24000 rpm in primary vacuum.
The spectrometer comprises a collimated mirror for the incoming light, a holographic grating having 3600 grooves/mm providing a dispersion of 0.3 nm mm 1 with an eﬃciency of 52%. The spectrometer separates the backscattered radiations into the emitted laser beams at 308 and 355 nm (Rayleigh signals) and at 331.8 and 386.7 nm (Raman signals). To account for the high dynamic range of the lidar signals the elastically scattered Rayleigh signals are separated into a high and low energy channels so that the lidar set up consists of 6 optical channels.
Detection and acquisition
The 6 optical channels are detected by bialkali Hamamatsu photomultiplier tubes (PMTs), characterised by a quantum eﬃciency of ∼20% in the 300–400 nm spectral range. The PMT provides current pulses when photons strike the photocathode. They are amplified by a 250 MHz bandwidth amplifier resulting in its broadening up to 5 ns and the current signals are converted to voltage signals. In addition to the chopper, electronic gating of the PMT is used for the high energy Rayleigh signal detection that reduces the signal-induced noise in these channels.
These signals are then directed to the acquisition system that uses photon counting method to process the electronic lidar signals. For that, high speed counters (250 MHz) operated with 1024 time gates of 1µs corresponding to a sampling vertical resolution of 150 m are used. Each channel is equipped with two counters in parallel to avoid the dead time between two memory bins. The overlap between the pulses of finite duration restricts the linearity of the counter, which mostly aﬀects the Rayleigh signals. It is managed by the use of two optical channels for Rayleigh signals. All these processes are controlled by a computer program and the master clock is set at 800 Hz and is provided by a mechanical chopper or a quartz crystal. The trigger of the counter is set by the laser light pulse, which is detected by a fast photodiode and converted to a transistor-transistor-logic signal. The six counting channels are transferred simultaneously to the computer with the acquisition time of 1024 µs. The data are averaged over 1000 shots corresponding to a temporal resolution of 200 s.
Ozone retrieval algorithm
The retrieval algorithm is generally based on the diﬀerential lidar equation formalism. First, to increase the signal-to-noise ratio, the lidar signals are time averaged during the measure-ment period (3–4 h in general), taken as the temporal resolution of the measurement. Then, a certain number of corrections such as background correction and dead time correction are applied to the averaged signals. The former term is related to the estimation of background light using linear or polynomial regression in the 80–150 km altitude range, where the lidar signal is negligible. The latter term is linked with the saturation of photon counting used for the signal acquisition in the lower ranges. In addition to the diﬀerentiation, a low pass filter is used in the DIAL technique to account for the rapid decrease of signal-to-noise ratio in the higher altitudes (above 40 km). Generally, the ozone number density is computed from the diﬀerence in the derivative of the logarithm of each lidar signal fitted to a straight line, or to a second order polynomial or to higher order polynomials. At OHP, the second order polynomial fit is used to derive the ozone number density (Godin et al., 1999).
The Rayleigh high and low energy and Raman signals optimise the accuracy of the ozone profile in the upper, middle-low and lower stratosphere respectively. In background aerosol loading conditions, the low energy Rayleigh signals provide more vertically resolved profiles than the Raman signals in the lower stratosphere, but the use of these signals in the lowermost stratosphere is prevented by the saturation of the photon counters. Hence, a correction called pulse pile-up correction is applied to correct for this saturation. The equation used to compute the true photon count rate from the observed count rate is Pc = 1 + [(1 − x)Pr − 1] exp(−xPr ) (2.11).
where Pc is the observed photon count rate, Pr is the true count rate and x = 1/Pmax with Pmax is the maximum observed count rate.
For high energy Rayleigh signals, x is adjusted for each wavelength in order to obtain the best agreement between the slopes of both low and high energy Rayleigh signals. For the low energy Rayleigh signals, Raman signal is used by computing reference Rayleigh slopes from the Raman slopes, the derived Raman ozone profile and the Rayleigh extinction correction. So the final ozone profile is retrieved by combining the slopes of low and high energy Rayleigh signals at first and then by combining the Raman and the composite Rayleigh ozone profiles. The altitude range, where both profiles are combined depends on the aerosol content. It is around 14–15 km in the case of background aerosol conditions while ment (Reproduced from Godin-Beekmann et al., 2003). high energy Rayleigh channels are used from 18 to 22 km. At the end, both Raman and composite Rayleigh profiles are corrected from the Rayleigh extinction using the pressure-temperature profiles obtained from nearby radio soundings performed in Nîmes in the lower stratosphere and the COSPAR International Reference Atmosphere 1985 (CIRA-85) model in the upper stratosphere.
Features of OHP lidar measurements
The optical receiver installed in 1993 enabled the ozone lidar system to measure in the lowermost stratosphere even in the presence of volcanic aerosols. The simultaneous acquisi-tion of all lidar signals improved the observational capacity in terms of temporal resolution and accuracy. Thus, the average number of measurements per year increased from ∼40 in 1986–1993 to ∼110 from 1994 onwards, with a maximum of 190 in 1997. The typical duration of an ozone measurement in the whole stratosphere with the present DIAL system at OHP is 4 hours. The precision and vertical resolution of ozone measurement is shown in Fig. 2.2. The precision ranges from about 5% below 20 km to more than 20% above 45 km. The vertical resolution ranges from 0.5 km at 20 km to about 2 km at 30 km, and it increases to ∼4.5 km at 45 km. The average accuracy ranges from ∼5% below 20 km to more than 10% above 45 km and the best accuracy of 3% is found in the 20–45 km altitude range (Godin-Beekmann et al., 2003). The altitude range of each profile varies depending on the presence of clouds in the lower stratosphere and varied signal-to-noise ratio in the upper stratosphere. The profiles are cut when 80% statistical error is reached.
Features of other NDACC lidar measurements
The NDACC lidar stations considered in the study include the lidars from Meteorologi-cal Observatory Hohenpeissenberg (MOHp: 47.8 N, 11 E), Tsukuba (36 N, 140 E), Ta-ble Mountain Facility (TMF: 34.5 N, 117.7 W), Mauna Loa Observatory (MLO: 9.5 N, 155.6 W) and Lauder (45 S, 169.7 E). All lidar stations use DIAL method for measuring stratospheric ozone with 308 nm as the ozone absorption wavelength. The main diﬀer-ence among the lidars is in the selection of reference wavelength. Most lidar stations use 355 nm as the reference wavelength except MOHp and Lauder lidar, which use 353 nm over the whole period and, TMF and MLO lidars used this configuration until 2000 and then changed to 355 nm (Leblanc and McDermid, 2000). Other diﬀerences among the lidars are in the receiving system and the number of channels used to detect the lidar signal in order to increase their dynamical range. At Tsukuba (Tatarov et al., 2009) and Lauder (Brinksma et al., 2000), 6 channels (4 Rayleigh and 2 Raman) are used to measure ozone. However, only 2 receiving channels (Rayleigh) are used at MOHp (Steinbrecht et al., 2009a) and, 8 channels at TMF (4 at 308, 332 nm; 4 at 355, 387 nm) and MLO (3 at 308, 332 nm; 5 at 355, 387 nm). The precision of lidar ozone measurements decreases generally from 1% up to 30 km, 2–5% at 40 km and to 5–25% at 50 km.
Ozone absorption cross-section
Absorption cross-section is a measure of the probability of a molecule to absorb photon at a particular wavelength. It is proportional to the intensity of absorption or emission between the two energy levels and is expressed in cm2/molecule. The absorption cross-section of a substance can be determined from the Beer-Lambert law as I(λ) = I0(λ) exp(−σ(λ)N l) (2.12).
where I0 and I are intensities, in W/m2, of the incident and transmitted light, respectively, σ(λ) is the absorption cross-section in cm2, N is the number density of absorbing particles in molecules/cm3 and l is the cell optical path in cm or the distance the light travels through the material.
There are diﬀerent groups performing experiments for the determination of ozone cross-sections. Most commonly used cross-sections are the Bass and Paur (BP) and Brion-Daumont-Malicet (BDM) cross-sections. The BP cross-sections are measured over 230– 350 nm wavelength range for temperatures 203, 223, 246, 273, 276 and 280 K based on the assumption that the ozone cross-section at 253.65 nm mercury line is temperature indepen-dent (Bass and Paur, 1984). BDM ozone cross-sections are provided at 218, 223, 243, 273 and 295 K in the spectral range 195–345 nm except for the measurements at 273 K which are limited to 300–345 nm (Daumont et al., 1992). BDM cross-sections at 295 K are now available in the 345–830 nm wavelength range too (Brion et al., 1998).
Temperature and wavelength dependence of cross-section
Figure 2.3 shows the spectral variations of BDM (top) and BP (bottom) ozone cross sec-tions with wavelengths, in the Hartley (200–310 nm) and Huggins (310–345 nm) bands, at diﬀerent temperatures. The wavelength ranges used for the DIAL ozone measurements in the troposphere and stratosphere are also shown on the Figure. It is clear that the cross-section decreases as wavelength increases. That is, ozone cross-section at 355 nm is less than that at 308 nm. The cross-sections at diﬀerent temperatures show a strong continuity and are identical until 310 nm, indicating a weak temperature dependence in the Hartley band. Above 310 nm or in the region of Huggins bands, the cross-section varies with temperature. In the 305–315 nm range, temperature dependence is almost linear and ∼15% diﬀerence is observed between the cross-sections at 218 and 295 K. However, above 315 nm the temper-ature eﬀect becomes prominent and increases progressively as wavelength increases, where the diﬀerence in cross-section at the extreme temperatures reaches more than 50%. The features are the same for BP too.
Table of contents :
1.1 Vertical structure of the atmosphere
1.2 Stratospheric ozone
1.2.1 Stratospheric chemistry
1.2.2 Dynamical processes
1.3 Ozone depletion issue
1.3.1 Antarctic ozone loss
1.3.2 Arctic ozone loss
1.3.3 Mid-latitude ozone loss
1.4 Equivalent Effective Stratospheric Chlorine
1.5 Present state of the ozone layer
1.5.1 Ozone total column measurements
1.5.2 Ozone vertical profile
1.6 Ozone recovery : different stages of ozone evolution
1.7 Ozone and climate
2 Ozone lidar measurements
2.2 Ozone DIAL system
2.2.3 Error analysis
2.3 The ozone lidar system at OHP
2.3.2 Optical receiver
2.3.3 Detection and acquisition
2.3.4 Ozone retrieval algorithm
2.3.5 Features of OHP lidar measurements
2.4 Features of other NDACC lidar measurements
2.5 Sensitivity tests
2.5.1 Ozone absorption cross-section
2.5.2 Temperature and wavelength dependence of cross-section
2.5.3 Comparison between BP and BDM cross-sections
2.5.4 Comparison between BP and BDM ozone number densities
2.5.5 Temperature dependence of ozone retrieval
2.6 OHP lidar ozone retrieval using NCEP data
3 Stability of ozone measurements at OHP
3.1 Ozone Measurements
3.1.4 SAGE II
3.2.1 Data screening
3.2.2 Coincidence criteria
3.2.3 Data conversion
3.2.4 Data analysis
3.3 Vertical distribution of mean bias
3.3.1 Long-term data sets
3.3.2 Short-term data sets
3.4 Temporal evolution
3.4.1 Comparison of Umkehr with lidar
3.4.2 Comparison of ozonesondes with lidar
3.4.3 Comparison of SAGE II and HALOE with lidar
3.4.4 Comparison of SBUV(/2) with lidar
3.4.5 Comparison of MLS and GOMOS with lidar
3.5 Drift in ozone differences
3.5.1 Sensitivity of standard deviations
3.5.2 Significance of the drifts in terms of the chosen standard deviation .
4 Stability of ozone observations over NDACC lidar stations
4.1 Ozonesonde measurements
4.2 Data analysis
4.2.1 Relative difference and mean bias
4.2.2 Data conversion
4.3 Average biases: comparison with lidar measurements
4.3.1 Correction factor
4.4 Relative drifts
4.4.1 Comparison with ozone lidar as reference
4.4.2 Comparison of lidar with SBUV(/2), SAGE II and HALOE as references
4.4.3 Comparison of SBUV(/2), SAGE II and HALOE
4.4.4 Average of the drifts of long-term measurements
4.5 Combined data: SAGE II, HALOE and Aura MLS
4.5.1 Time series
4.5.2 Relative drifts of the combined time series
5 Stratospheric ozone evolution in the northern mid-latitudes
5.1 Explanatory variables
5.1.1 Quasi Biennial Oscillation
5.1.2 Solar flux
5.1.4 Eddy heat flux
5.1.5 North Atlantic Oscillation
5.1.6 PWLT and EESC : Ozone trend estimation methods
5.2 Multiple regression model and method
5.3 Ozone total column measurements
5.3.1 Evolution of ozone total column
5.3.2 Ozone anomaly
5.3.3 Comparison between Dobson and SAOZ at OHP: bias and drift .
5.4 Multiple regression analysis of ozone total column at OHP
5.4.1 Contribution of proxies to ozone variability
5.4.2 Trends in ozone total column
5.5 Multiple regression analysis of ozone total column at MOHp
5.5.1 Contribution of proxies to ozone variability
5.5.2 Trends in ozone total column
5.6 Vertically resolved ozone observations at OHP
5.6.1 Stratospheric ozone evolution
5.6.2 Stratospheric ozone anomaly
5.6.3 Application of multiple regression
5.6.4 Contribution of proxies to the variability of ozone profiles
5.6.5 Trends in stratospheric ozone vertical profiles
5.7 Connection between ozone profile and column measurements
6 Summary, conclusions and perspectives
6.1 Summary and conclusions