Satellite systems for atmospheric ozone observations

International Journal of Remote Sensing

ISSN: 0143-1161 (Print) 1366-5901 (Online) Journal homepage: http://www.tandfonline.com/loi/tres20

Satellite systems for atmospheric ozone observations A.P. Cracknell & C.A. Varotsos To cite this article: A.P. Cracknell & C.A. Varotsos (2014) Satellite systems for atmospheric ozone observations, International Journal of Remote Sensing, 35:15, 5566-5597 To link to this article: http://dx.doi.org/10.1080/01431161.2014.945013

Published online: 28 Aug 2014.

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Date: 07 November 2016, At: 11:57

International Journal of Remote Sensing, 2014 Vol. 35, No. 15, 5566–5597, http://dx.doi.org/10.1080/01431161.2014.945013

REVIEW ARTICLE Satellite systems for atmospheric ozone observations A.P. Cracknella and C.A. Varotsosb* a

Division of Electronic Engineering and Physics, University of Dundee, Dundee DD1 4HN, Scotland, UK; bDepartment of Applied Physics, University of Athens, 15784 Athens, Greece (Received 27 August 2013; accepted 16 November 2013) A review of the satellite-borne instrumentation used for the total ozone and ozone vertical profile watch is presented. The principle of the observational method for the monitoring of the atmospheric ozone content is used to group the basic satellite systems. Along these lines the direct-absorption, scattering, and emission-measuring instruments are presented, besides providing a short discussion on the spatio-temporal variability of the ozone content mainly derived from a few outstanding field campaigns.

1. Introduction One of the important atmospheric parameters that is extensively studied by satellite remote sensing is ozone concentration, both total ozone content (TOC) and the ozone vertical profile (OVP). A TOC value at any location is often reported in Dobson units (DU) expressing the sum of all the ozone in the atmosphere directly above that location and it is generally the lowest over the equator and the highest in polar regions. TOC patterns change on various timescales due to natural air motions taking place in both stratosphere and troposphere. Another reason for TOC variability is the changes occurring in the balance of chemical production and loss processes. In general, ozone (O3) is mainly produced in the stratosphere by natural processes (involving the break-up of O2 into atomic oxygen (O), by the solar ultraviolet radiation (UV), with wavelengths less than 0.240 μm and the subsequent rapid combination of O with O2). It drifts downwards by mixing processes, producing a maximum in its concentration near 25 km. Apart from its production (in total, 3O2 + UV = 2 O3) it is also destroyed naturally by absorbing UV radiation with wavelengths of less than 0.320 μm and through chemical reactions with other atmospheric components. Among them are several catalytic cycles, such as those of NO and NO2, according to which: O3 + NO = NO2 + O2 and NO2 + O = NO + O2, or overall: O3 + O = O2 + O2 (WMO 2010). The natural balance between production and destruction leads to the maintenance of the stratospheric ozone layer, which, however, is now subject to substantial depletion due to human-produced ozone-depleting substances (ODSs), such as chlorofluorocarbons (CFCs) (e.g. Newman et al. 2009). Much smaller amounts of ozone, of the order of 10% of TOC, are residing in the troposphere (Figure 1), where several human activities (such as fossil fuel combustion) result in its gradual increase. Monitoring of the global ozone layer identified a thinning of it, which becomes most severe in springtime over Antarctica, that is commonly called the ‘ozone hole’. However, the replacement of CFCs with ‘ozone-friendly’ chemical substitutes (under the provisions *Corresponding author. Email: [email protected] © 2014 Taylor & Francis


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Figure 1. The vertical distribution of atmospheric ozone partial pressure of ozone (mPa) and height (km) (Earth Observing System (EOS) Science Plan. Chapter 7. Ozone and Stratospheric Chemistry, http://eospso.gsfc.nasa.gov/science_plan/).

of the Montreal Protocol) resulted in a decrease of the global accumulation of ODSs and subsequently there is no longer an increase of the global ozone depletion (Fahey and Hegglin 2011). Both the Vienna Convention for the Protection of the Ozone Layer and its Montreal Protocol are the most widely ratified treaties in United Nations history (with 197 parties), and indicate the necessity of the routine monitoring of the ozonesphere. Satellite platforms provide the advantage of the required spatial coverage for the effective monitoring of the global ozone layer that it is not feasible with the employment of the ground-based network only, because its geographical distribution is not uniform over the globe. The up-to-date review of the satellite systems for atmospheric ozone observations is the main objective of the present article. This gathered information along with the already identified weaknesses of the available ozone observations would reveal the need for defining the present and future strategies of ozone satellite observation as a component of an optimized ozone network, which is the scope of this review. It is now slightly over 50 years since the satellite era began after 4 October 1957, when the former Soviet Union launched Sputnik 1, the world’s first artificial satellite (a 55 cm diameter sphere that weighed 83 kg with four antennae) (Figure 2). In the intervening years since then, satellites have revolutionized the Earth sciences, including the study of atmospheric trace gases in general and ozone in particular. With time, satellite-flown instruments became more complicated, whereas larger satellites carrying more and more instruments were launched. One of the largest of these was the Environmental Satellite (Envisat), which was launched on 1 March 2002 by the European Space Agency (ESA), with a mass of 8211 kg and dimensions in orbit, i.e. with its solar panels deployed, of 26 m × 10 m × 5 m. This was a polar-orbiting Sunsynchronous satellite flown at an altitude of about 800 km, with an inclination of 98° and an orbital period of 101 minutes. It carried ten main instruments for monitoring the Earth’s atmosphere, biosphere, cryosphere, geosphere and the oceans (Figure 2) (http://envisat. esa.int/) and three of these were particularly relevant to atmospheric ozone. Envisat successfully reached and exceeded (doubling) its nominal 5 year mission lifetime, having

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A.P. Cracknell and C.A. Varotsos

Figure 2. Envisat with various instruments aboard. In its orbit configuration, its dimensions are 25 m × 10 m × 5 m, and its total mass is 8211 kg (source: ESA). Sputnik on the same scale as Envisat (a basketball in diameter with mass 83 kg).

orbited the Earth more than 52,000 times. Figure 2 shows Sputnik on the same scale as Envisat. Recently, there has been a trend of moving away from large multifunctional Earthobserving satellites, such as Envisat, to small satellites dedicated to one particular observational task. The advantages of this include simple and speedy design and manufacture, many more launch opportunities, low cost, non-competition between the requirements of different instruments, risk reduction, etc. Small satellites also provide opportunities for developing countries to become involved in technology transfer and the development of indigenous space-related capabilities. Thus, apart from the large Earth-observing satellites, we foresee a role for the small ones in the future (Cracknell and Varotsos 2007). The traditional ways to study the concentration of atmospheric ozone were by specially designed ground-based spectrophotometers and with ozonesondes. TOC was measured from the ground at a few sites in the late 1920s and the early 1930s using Dobson spectrophotometers, and records of TOC have been kept at various sites since then and they continue to be kept in parallel with satellite measurements. There are some contrasts to be noted. The ground-based spectrometers and the ozonesondes measure the concentration of ozone directly. The satellite-flown instruments generally are less direct and rely on algorithms that need to be validated to extract the ozone data. In other words, satellite instruments perform observations of the emitted or backscattered radiation in the ozone spectrum, from which, knowing the ozone spectroscopy parameters (line intensities or cross-sections) and modelling the radiative transfer in the atmosphere, we can use inversion methods to retrieve the ozone total column or vertical profiles. Generally speaking only ozonesondes make the most reliable direct (in situ) measurements. Before considering satellite remote sensing of atmospheric ozone, we should mention the use of aircraft. There have been a few studies of the atmosphere using instruments


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flown on aircraft to study ozone and related chemicals directly, most notably in the early study of the Antarctic ozone hole during the Airborne Antarctic Ozone Experiment (AAOE) in 1987 and later on during the Airborne Polar Experiment – Geophysica Aircraft In Antarctica, in 1999 (APE-GAIA) (Carli, Cortesi, and DeRossi 1999; Giovanelli et al. 2005). Considerable insight was gained into the chemical processes involved in the destruction of stratospheric ozone (see, for instance, chapter 5 of Cracknell and Varotsos (2012)). However, in many cases of the satellite-flown instruments that we shall discuss in this article similar instruments have been flown on aircraft as well. There has usually been one of two reasons for this. The first is for calibrating and evaluating prototypes of the instruments before flying them in space. The second is for simultaneous flying of similar (nominally identical) instruments in aircraft and in space for the validation of data from the space-borne instrument. However, there is another rather new aspect of airborne remote sensing that should be mentioned. Actually this is only a new type of aircraft observation, but the objective is the same, i.e. validation of satellite instruments or verification of proto-flight instruments before flying them on a satellite. Similar to how remote-sensing satellites for environmental studies can be regarded as a spin-off development from military spy satellites, there is a similar spin-off from unmanned military aircraft in a new mission called Global Hawk Pacific (GloPac), and one or two other unmanned aircraft National Aeronautics and Space Administration (NASA) projects. NASA has staged environmental (manned) research flights from aircraft previously, but none has had the reach and duration of Global Hawk, a high-altitude, unmanned aircraft, see Figure 3(a). The aircraft, which is distinguished by its bulbous nose and 35.4 m wingspan, can travel about 18,500 km in up to 31 hours, carrying almost 1 t of instrument payload (http://gsfctechnology.gsfc.nasa.gov/GlobalHawk.htm). The Global Hawk aircraft used in the GloPac mission carried a suite of 12 instruments. The first operational flights of the Global Hawk for the GloPac were conducted in support of the Aura Validation Experiment (AVE). Aura is one of the NASA A-train of satellites (Figure 3(b)). This is a convoy of Earth-observing spacecraft following one another in a polar orbit. 2. Satellite remote sounding of the TOC and OVP There are three different principles involved in satellite remote sensing of ozone, two of which involve looking down at the Earth below the orbit of the spacecraft and one of which involves limb sounding, i.e. observing solar radiation by looking towards the horizon from the satellite. The choice of using limb or nadir geometry depends on whether we focus on tropospheric or stratospheric ozone monitoring, i.e. it depends on the mission goal, being either ozone hole or climate change and air-quality studies. 2.1. TOVS The first method, which has been used for many years by the Television Infrared Observation Satellite (TIROS) Operational Vertical Sounder (TOVS) and its successors, involves looking directly down at the Earth’s surface and using Earth’s surface leaving radiation. TOVS is a set of three instruments – the High Resolution Infrared Sounder (HIRS), the Stratospheric Sounding Unit (SSU) and the Microwave Sounding Unit (MSU) – providing 27 spectral channels, which are all, except one, in the infrared (IR) and

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Figure 3.


(a) A NASA Global Hawk lands at Dryden and (b) A-Train satellites (source: NASA).

microwave region of the spectrum with one panchromatic channel (for more details see, for instance, section 1.4 of Cracknell (1997)). There is one particular spectral channel, channel 9 of HIRS, at a wavelength of 9.7 µm, which is particularly well suited to monitoring stratospheric ozone concentration; this is a (general) window channel, except for absorption by ozone, i.e. ozone is the only principal atmospheric constituent that strongly absorbs radiation at this wavelength. The radiation received at this wavelength by the HIRS instrument was emitted from the Earth’s surface, but it is attenuated by the ozone in the atmosphere. The lesser the ozone, the greater the amount of radiation reaching the satellite. A 3 DU drop in lower stratospheric ozone produces a measurable (circa 0.2 °C) increase in the brightness temperature in this channel. It should be clarified, however, that the surface emission in the IR dominates at nadir geometry in this spectral window, but the radiative process includes not only atmospheric absorption, but also emission by ozone along the atmospheric column, although this may be masked by the higher surface emission.


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It appears that, strictly speaking, what is measured is the lower stratospheric ozone and a correction has to be applied to obtain TOC. Images are now regularly produced from TOVS data giving hemispherical daily values of TOC. This means that there is now a long time series of such data available. TOVS data have been used to determine atmospheric ozone concentration from 1978 to the present (Neuendorffer 1996; Kondratyev 1998) and the standard deviation is approximately 7%. Nowadays, TOVS data are also used when the more reliable Total Ozone Mapping Spectrometer (TOMS) data are not available (http://www.theozonehole.com/2010oct.htm). Li et al. (2001) discussed the potential for using Geostationary Operational Environmental Satellite (GOES) sounder radiance measurements to monitor TOC with a statistical regression using GOES sounder spectral bands 1–15 radiances. The advantage is the high temporal frequency of the availability of the data. Hourly GOES ozone products have been generated since May 1998. GOES ozone estimates were compared with TOMS TOC data and ozone measurements from ground-based Dobson spectrometer ozone observations. The results showed that the percentage root-mean-square (RMS) difference between instantaneous TOMS and GOES ozone estimates ranged from 4% to 7%. 2.2. TOMS and (S)BUV The second method also involves looking down from the spacecraft and, similar to the ground-based spectrophotometers, uses the wavelength dependence of the absorption of solar UV radiation in the atmosphere. There are two slightly different systems, namely TOMS and the (Solar) Backscatter UltraViolet ((S)BUV) type. TOMS has six UV wavelength bands from 312.5 to 380 nm (312.3, 317.4, 331.1, 339.7, 360, and 380 nm) and TOC is determined using the ozone absorption spectrum of the solar ultraviolet radiation reflected from the Earth’s surface. Solar Backscatter UV (SBUV) instruments are nadir-viewing instruments that are able to determine TOC and OVP by measuring sunlight scattered from the atmosphere in the UV spectrum. Similar to TOVS, TOMS is a scanning instrument. The first four wavelength regions are used in pairs, to provide estimates of ozone concentration by the differential absorption method, whereas the other two (without ozone absorption) are employed to estimate the effective background albedo. A detailed description of the adaptations to TOMS for Meteor-3 is provided by Herman et al. (1997) and the pre-launch and post-launch calibrations are described by Jaross et al. (1995). Of the five TOMS instruments that were built, four entered orbit successfully. It should be stressed that TOMS allows day/ night coverage, but is less accurate in case the Earth’s surface is too cold or too hot or in the presence of clouds; TOVS and SBUV measurements are restricted to daylight conditions. After a one-and-a-half year period when the programme had no in-orbit capability, the Japanese Advanced Earth Observing Satellite (ADEOS) TOMS was launched on 17 August 1996 and collected measurements until the satellite that carried it lost power on 29 June 1997. For ADEOS/TOMS TOC, the absolute error was ±3%, the random error is ±2%, and the drift over the 9 month data record was less than ±0.5%. The ADEOS/TOMS observations of TOC are approximately 1.5% higher than a 45-station network of groundbased measurements (McPeters and Labow 1996; Seftor et al. 1997; Herman et al. 1997; Torres and Bhartia 1999). The successor of ADEOS, Earth Probe TOMS (launched on 2 July 1996), failed in December 2006 and was subsequently replaced by the Ozone Monitoring Instrument

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A.P. Cracknell and C.A. Varotsos Table 1. Channel 1 2 3 4 5 6 7 8 9 10 11 12

SBUV/2 wavelengths. Wavelength (nm) 252.0 273.6 283.1 287.6 292.3 297.5 301.9 305.8 312.5 317.5 331.2 339.8

(OMI). The historical and current daily and monthly ozone data are widely available at the NASA website http://ozoneaq.gsfc.nasa.gov/. Spatially, a good global coverage in TOMS (OMI) data is combined with a high resolution of l° latitude and 1° longitude. It should be noted that TOC observations made by TOMS and OMI exhibit some minor differences (e.g. on the so-called O3 ghost column), thereby making it difficult to evaluate trends by merging these datasets. The O3 ghost column is a result of the fact that the zenith sky twilight measurements are strongly weighted by the stratosphere with limited sensitivity to tropospheric ozone (e.g. McPeters, Labow, and Logan 2007; Hendrick et al. 2011). The BUV and SBUV instruments have more UV channels than the TOMS and many of these are at shorter wavelengths than in the TOMS, see Table 1. The SBUV instrument (on board the National Oceanic and Atmospheric Administration-16 (NOAA-16)) provided observations of the scattered UV radiation (252–340 nm). The data at the shortest wavelengths are employed for the estimation of the ozone concentration as a function of height, whereas those at the longer wavelengths are used to provide TOC observations. The other spectral bands are at shorter wavelengths than on TOMS and this radiation is reflected back to the satellite by the upper layers of the atmosphere before it can ever reach the surface of the Earth. The ozone observations derived from satellite data have been performed by the BUV instrument since April 1970 (i.e. launch of Nimbus-4). The advanced SBUV instrument (launched in November 1978) has been replaced by the SBUV/2 (NOAA and TIROS) (Table 2). It should be noted that the stratospheric ozone began to be monitored operationally in 1985 (NOAA-9, NOAA-11, NOAA-14, NOAA-16, NOAA-17, NOAA-18, and NOAA-19). The Shuttle SBUV (SSBUV) was designed and developed at the NASA Goddard Space Flight Center (GSFC) to calibrate the Nimbus and NOAA solar backscatter UV instruments. In late 1989, the Space Shuttle Atlantis carried the instrument for the first time, in an appropriate orbital flight path to assess performance by directly comparing data from identical instruments (SBUV) on board the NOAA and Nimbus-7 spacecraft. SSBUV was flown on eight Shuttle missions between October 1989 and January 1996 and provided regular checks on the individual satellite instruments’ calibrations (Heath et al. 1993). Observations of around 30 trace gases over northern tropics and subtropics, as well as in the vicinity of the polar vortex over Antarctica, were collected by the Atmospheric Trace Molecule Spectroscopy Experiment (ATMOS) Fourier transform


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Table 2. Satellite instrumentation for BUV ozone observations (Hilsenrath et al. 1997). Instrumentation BUV SBUV SBUV/2 SBUV/2 SBUV/2 SBUV/2 SBUV/2 SBUV/2 SBUV/2 SSBUV TOMS TOMS TOMS TOMS GOME Aura

Satellite

Observation period

Nimbus-4 Nimbus-7 NOAA-9 NOAA-11 NOAA-14 NOAA-16 NOAA-17 NOAA-18 NOAA-19 Shuttle Nimbus-7 Meteor-3M Earth Probe ADEOS ERS-2 OMI TO3

1970–1975 1979–1990 1985–1998 1989–2003 1995–2007 2000–present 2002–present 2005–present 2009–present 8 flights, 1989–1996 1978–1993 1991–1994 1996–2005 1996–1997 1995–2011 2004–present

spectrometer (flown on the Space Shuttle as part of the Atmospheric Laboratory for Applications and Science (ATLAS)-3 instrument payload) for around 200 sunrise and sunset occultation events during November 1994 (Gunson et al. 1996).

2.3. Limb sounding The third general system used for the observation of ozone – and of other trace gases in the atmosphere – is limb sounding. This involves looking towards the horizon from the spacecraft, rather than looking vertically downwards, and observing the light from the Sun, or occasionally from a star. The reason for this is to obtain a longer path through the atmosphere and therefore a greater number of trace gas molecules in the path of the radiation. It should also be noted that the detection and retrieval of minor trace gases are not the only reason for using limb measurements, i.e. they allow one to study with greater vertical resolution the chemical processes within the stratosphere, especially in regions of high convection, e.g. the upper troposphere – lowermost stratosphere. On the contrary limb measurements do not allow one to monitor the lowermost troposphere (i.e. from the surface to 3 km height), which will be the region of interest in the coming future. During the second half of the 1960s, several attempts at satellite remote sounding of ozone and other minor gas components from Soviet manned spacecraft were first made, using a hand-held spectrograph and a complex set of solar spectrometers functioning in a regime of occultation geometry (Kondratyev 1998). An early set of limb sounding instruments was flown on Upper Atmosphere Research Satellite (UARS), which was launched on 15 September 1991. These instruments included the Cryogenic Limb Array Etalon Spectrometer (CLAES), the Improved Stratospheric and Mesospheric Sounder (ISAMS), the Microwave Limb Sounder (MLS), and the Halogen Occultation Experiment (HALOE). The decommissioned satellite re-entered the atmosphere on 24 September 2011, with considerable media attention, caused by NASA predictions that substantial parts of it could reach the land. In the event, it landed in a remote area of the Pacific.

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The complementary instrumentation of ADEOS included the Improved Limb Atmospheric Spectrometer (ILAS), the Interferometric Monitor for Greenhouse Gases (IMG), and the Retro-reflector in Space (RIS) instrument. IMG is a nadir-observing Michelson-type Fourier Transform Spectrometer (FTS) designed to measure the vertical profiles of CO2 and H2O, TOC, and the concentrations of CH4, N2O, and CO in the troposphere. 2.4. More recent instruments The TOVS, TOMS, and (S)BUV series of systems and the early limb sounding systems have, between them, provided quite a lengthy archive of atmospheric ozone data, but there are now a number of newer instruments that are expected to continue supplying this data for quite some time in the future, some of which combine both looking vertically downwards and limb sounding. A few of the more important ones include: THE Global Ozone Monitoring Experiment (GOME) on the ESA satellite ERS-2 (1995–2011); the OMI on the NASA Aura satellite (2004–today); Global Ozone Monitoring by Occultation of Stars (GOMOS) (Meijer et al. 2004), the Scanning Imaging Absorption Spectrometer for Atmospheric Chartography (SCIAMACHY), and Michelson Interferometer for Passive Atmospheric Sounding (MIPAS) on the European satellite Envisat (2002–2012); the Infrared Atmospheric Sounding Interferometer (IASI), and GOME-2 on the first European polar orbiting satellite Meteorological Operational (MetOp) (2006–today); and the Total Ozone Unit (TOU), which is one of the main payloads on FengYun-3 (FY-3), the Chinese second-generation polar-orbiting meteorological satellite series (2008–today) (Yang, Lu, et al. 2012; Yang, Shang, et al. 2012; Zhang et al. 2009). We shall describe some of these briefly in turn. GOME was launched on 20 April 1995 on board the ESA Earth Resources Satellite (ERS-2) in a polar Sun-synchronous orbit. It is a nadir-viewing multichannel spectrometer measuring solar irradiance and earthshine radiance (not solar irradiance) in the wavelength range 240–790 nm at moderate spectral resolution (0.2–0.4 nm). A second GOME instrument, GOME-2, was flown on MetOp-1, the first European polar-orbiting meteorological satellite, which was launched on 16 October 2006. Some early work on comparison of its TOC data with reliable ground-based measurement recorded by five Brewer spectrophotometers in the Iberian Peninsula was carried out over a period of a year from May 2007 to April 2008. This showed that GOME-2/MetOp ozone data has a very good quality; for details see Antón et al. (2009). The OMI is a nadir-viewing near-UV/imaging grating spectrometer aboard the Earth Observing System (EOS) Aura satellite of NASA. Aura flies about 15 minutes behind Aqua in the A-Train (see Figure 3(b)). Among the Aura instruments are the High Resolution Dynamics Limb Sounder (HIRDLS), the Microwave Limb Sounder (MLS), and the Tropospheric Emission Spectrometer (TES). The inter-comparison among OMI, TOMS, GOME, and SCIAMACHY shows that OMI exhibits improved spatial resolution for the routine monitoring of various trace gases from space (Hassinen et al. 2008; Ialongo, Casale, and Siani 2008). It should be noted that OMI has better spatial resolution, due to the better (evolved) charged-coupled device (CCD) used with respect to GOME or SCIAMACHY. A descendant of OMI (as well as of SCIAMACHY) is the Tropospheric Monitoring Instrument (TROPOMI), observing at the ultraviolet, the visible, the near-infrared, and the shortwave infrared, allowing monitoring of O3, NO2, SO2, CO, CH4, CH2O, and aerosol. The TROPOMI will be flown on board the ESA Sentinel-5 Precursor (S-5 P), a low Earth


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orbit polar satellite that will monitor key parameters of air quality, climate, and the ozone layer during 2015–2022 (Veefkind et al. 2012). TROPOMI will have even better spatial resolution, since CCD technology is always improving. It is also crucial to highlight that TROPOMI will allow a daily coverage with a spatial resolution of up to 7 km × 7 km, allowing monitoring from space the air quality of an urban area at a resolution never obtained from a satellite and on a global scale. The NOAA series of polar-orbiting meteorological satellites that have operated continuously since 1978 carrying the TOVS and also the Advanced Very High Resolution Radiometer (AVHRR), and in some instances the SBUV/2, were planned to be replaced by the National Polar-orbiting Operational Environmental Satellite System (NPOESS). NPOESS was to be a replacement for both the US Department of Defense’s Defense Meteorological Satellite Program (DMSP) and the NOAA Polar Operational Environmental Satellites (POES) series. The NPOESS Preparatory Project (NPP) was planned as a pathfinder mission for NPOESS. It was launched five years behind schedule, on 28 October 2011. The NOAA/NASA portion is called the Joint Polar Satellite System (JPSS) and the European Organisation for the Exploitation of Meteorological Satellites (EUMETSAT) operates the MetOp polar orbiting weather satellite programme. The five-instrument suite for JPSS includes the Visible/Infrared Imager Radiometer Suite (VIIRS), the Cross-track Infrared Sounder (CrIS), the Clouds and the Earth Radiant Energy System (CERES), the Advanced Technology Microwave Sounder (ATMS), and the Ozone Mapping and Profiler Suite (OMPS). The OMPS consists of two sensors: a nadir sensor and a limb sensor. Observations performed by the nadir sensor are employed to develop TOC data, whereas observations derived from the limb sensor produce OVP of the along-track limb scattered sunlight. OMPS was launched aboard NASA’s Suomi National Polar-orbiting Partnership (NPP) satellite on 28 October 2011. The JPSS algorithm is an extension of the TOMS Version 7 TOC algorithm (McPeters and Labow 1996). It should be stressed at this point that for the investigation of the ozone depletion and recovery simultaneous long-term observations of O3 and NO2 are necessary. Moreover, simultaneous observations of OH and HO2 are very important for the upper stratospheric ozone budget, because both participate in catalytic cycles for ozone destruction. In this context, MLS on Aura is the first instrument able to measure OH and HO2 simultaneously. The two satellite instruments the Optical Spectrograph and Infra-Red Imager System (OSIRIS) and the Atmospheric Chemistry Experiment (ACE) have been taking O3 and NO2 measurements since 2001 and 2003, respectively (Adams et al. 2012). OSIRIS, which was launched aboard the Odin spacecraft (in February 2001), is concerned with observing limb-radiance vertical profiles that range between 10 km and 100 km and has latitudinal coverage 82.2°N–82.2°S. ACE consists of two instruments: ACE-FTS (infrared) and ACE-MAESTRO (UV–visible-near-IR), aboard the solar occultation satellite SCISAT-1 (Canadian Space Agency), launched in August 2003 (Kerzenmacher et al. 2005). Of course, NO2 can be measured also with SCIAMACHY, GOME, OMI, and several of the other satellites. The main advantage of ACE is the high vertical resolution (due to limb geometry), the extremely good signal-to-noise ratio (due to Sun-occultation), and the stability of the radiometric calibration (self-calibrating); ACE allows the retrieval of a high number of minor trace gases (not only NO2) and also the study of chemical processes in the stratosphere. Conversely, ACE has very limited temporal coverage (sunset and sunrise) and it cannot measure in the lower troposphere or below cloud level. Recently, Kursinski et al. (2012) presented the initial observational results obtained by the Active Temperature, Ozone and Moisture Microwave Spectrometer (ATOMMS),

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which is a new remote-sensing system for global observations of various atmospheric constituents. We have described thus far the main instruments that have provided archived atmospheric ozone data in the past, those that are currently providing this data, and the instruments that are expected to be providing such data in the future. There are quite a number of other instruments that have been only of ephemeral interest or that have come on the scene only recently, and there are too many of them for us to describe them all. A comprehensive list is given in Table 3. Details of many of them will be found in several sections of chapter 2 of the book by Cracknell and Varotsos (2012), where a categorization based on the measurement techniques (backscatter Sun, emission, and occultation) has been adopted. For each type of category the relevant advantages and drawbacks have been clearly outlined with respect to ozone monitoring and in particular issues that are affecting the quality of ozone data for each technique, i.e. vertical sensitivity and dependence on the solar zenith angle have been discussed. The general principles, however, have been covered by the instruments we have described here.

3. Inter-calibration and validation of atmospheric ozone data sets We now move on to consider the results obtained from all the various instruments that we have described and the ground-based instruments that are used to validate and augment the satellite-derived data. Great importance is attached to the question of temporal changes in atmospheric ozone concentration, especially the question of stratospheric ozone depletion, both globally and particularly in the polar regions. Therefore, in looking for longterm trends in atmospheric ozone concentration, one needs to have accurate long-term datasets. As mentioned before, the satellite-flown instruments do not directly give the values of the required ozone concentrations. It is necessary to use an appropriate retrieval algorithm for each type of instrument and these various algorithms have to be validated, and if necessary refined, by comparison with data from ground-based instruments and ozonesondes. Among the ground-based instruments are lidars, which are very extensively used for validation of ozone stratospheric profiles, usually to extend ozonesonde data that are restricted to below 35 km (e.g. Meijer et al. 2004). Also, if there are successive versions of instruments that are nominally the same, e.g. the several TOMS instruments, it is necessary to inter-compare these datasets so that one consistent dataset can be constructed with no artefacts that arise from the transition from one instrument in the series to another. The problem of instrumental bias has been discussed throughout the whole history of ozone measurements, but it acquired a special importance after satellite observations started. Even at the early stages of satellite TOC observations, very serious attention was paid to this problem. The most reliable information about the spatial and temporal variation of the concentration of atmospheric ozone will be obtained by making use of all sources of data from ozonesondes, from ground-based instruments and from satellite-flown instruments. Not only is it necessary to inter-calibrate different instruments that are operating at any given time, but in order to construct a long-term archive it is necessary to inter-calibrate different versions of the same instrument, such as Dobson spectrometers, which have been operated over different periods. The whole question of inter-comparisons between different atmospheric ozone datasets is discussed at great length in chapter 3 of the book by Cracknell and Varotsos (2012). We shall consider next the question of inter-calibration of different data sources and the validation of ozone extraction algorithms.

Echo-1

SAMOS-9

Ariel-2

Kosmos-45 (experimental weather satellite)

OV1–1 Kosmos-65 (experimental weather satellite) Kosmos-92 (experimental weather satellite) Kosmos-122 (pre-Meteor weather satellite) OV1–10 OGO-4 OAO-2 Nimbus-3 Nimbus-4 OAO-3 Nimbus-6 Explorer-55 (AE-5 or AE-E) DMSP-5D1-F1 DMSP-5D1-F2 DMSP-5D1-F3 Nimbus-7 (longest-lasting mission; data 1978–1993) Explorer-60 (Applications Explorer Mission (AEM-B)) DMSP-5D1-F4 NOAA-6 DMSP-5D1-F5 (launch failed)

18 July 1962

27 March 1964

13 September 1964

21 January 1965 17 April 1965 16 October 1965 25 June 1966 11 December 1966 28 July 1967 7 December 1968 13 April 1969 8 April 1970 21 August 1972 12 June 1975 20 November 1975 11 September 1976 4 June 1977 1 May 1978 24 October 1978 18 February 1979 6 June 1979 27 June 1979 14 July 1980

Carrier satellite/space shuttle mission

12 August 1960

Launch date

(Continued )

Ground-based analysis of sunlight reflected by the balloon (passive method, inaccurate results; solar occultation; sunlight from limb was reflected to ground from the balloon) UV radiometer (first experiment with active sensor; solar occultation limb scanning) Broadband filter photometer, simple prism spectrometer (solar occultation limb scanning) BUV radiometer (first known BUV (backscatter UV) instrument; downwards looking) BUV radiometer (downwards looking, 30° left of nadir) BUV radiometer BUV radiometer BUV radiometer BUV spectrophotometer Ebert–Fastie scanning spectrometer UV telescope/spectrophotometer (starlight occultation limb scanning) IRIS BUV spectrometer, IRIS UV telescope/spectrometer (starlight occultation limb scanning) Limb Radiance Inversion Radiometer (LRIR) (limb scanning) BUV (solar occultation limb scanning rather than downwards looking) SSH (multichannel filter radiometer) SSH (multichannel filter radiometer) SSH (multichannel filter radiometer) LIMS (limb scanner (an update of the LRIR)), TOMS-1, SBUV Stratospheric Aerosol and Gas Experiment (SAGE)-1 SSH (multichannel filter radiometer) HIRS/2 (atmospheric sounder with 9.7 µm ozone channel) SSH (multichannel filter radiometer)

Ozone monitoring instrument description

Table 3. Chronological summary of ozone-monitoring satellites (or Space Shuttle missions) (source: Colorado State University, http://rammb.cira.colostate.edu/ dev/hillger/satellites.htm).

International Journal of Remote Sensing 5577

(Continued ).

NOAA-7 Explorer-64/Solar Mesosphere Explorer (SME)

NOAA-8 Earth Radiation Budget Satellite (ERBS) (launched from STS-41G) NOAA-9 NOAA-10 NOAA-11 STS-34 STS-41 NOAA-12 STS-43 Meteor-3–5 (Meteor-TOMS) UARS (launched from STS-48)

STS-45 STS-56 NOAA-13 Système Pour l’Observation de la Terre (SPOT)-3

STS-62 STS-66 NOAA-14 ERS-2 TechSat-1a microsatellite (failed) FASat-Alfa microsatellite (failed to separate from Sich-1) STS-72 MSX TOMS-Earth Probe (EP) ADEOS-1 STEP-4

28 March 1983 5 October 1984

12 December 1984 17 September 1986 24 September 1988 18 October 1989 6 October 1990 14 May 1991 2 August 1991 15 August 1991 12 September 1991

24 March 1992 8 April 1993 9 August 1993 26 September 1993

3 April 1994 3 November 1994 30 December 1994 20 April 1995 28 March 1995 31 August 1995 11 January 1996 24 April 1996 2 July 1996 17 August 1996 23 October 1997


23 June 1981 6 October 1981

Launch date

Table 3.

(Continued )

SBUV/2 (based on the SBUV/TOMS-1 flown on Nimbus-7), HIRS/2 HIRS/2 SBUV/2, HIRS/2 SSBUV-1 (GAS canister) SSBUV-2 (GAS canister) HIRS/2 SSBUV-3 (GAS canister) TOMS-2 CLAES, ISAMS, HALOE, SUSIM, MLS, Solar-Stellar Irradiance Comparison Experiment (SOLSTICE) SSBUV-4 (GAS canister, ATLAS-1 payload) SSBUV-5 (GAS canister, ATLAS-2 payload) SBUV/2, HIRS/2 Polar Ozone and Aerosol Monitor (POAM)-2 (solar occultation limb scanning) SSBUV-6 (GAS canister) SSBUV-7 (GAS canister, ATLAS-3 payload) SBUV/2, HIRS/2 GOME-1 (BUV technique) OM-2 UV radiometer (SBUV technique) OLME SSBUV-8 (GAS canister) UVISI TOMS(−3?) TOMS(−4?), ILAS-1 OOAM

HIRS/2 UV ozone experiment, Four-Channel Infrared Radiometer, Airglow Instrument HIRS/2 SAGE-2


5578 A.P. Cracknell and C.A. Varotsos

(Continued ).

STS-87

SPOT-4 NOAA-15 FASat-Bravo microsatellite Meteor-3M-2 (this Meteor-3M-1 follow-up was cancelled in 1999) TechSat-1b/OSCAR-32 microsatellite NOAA-16 Odin QuikTOMS (launch failed) Meteor-3M-1 Envisat NOAA-17 ADEOS-2 STS-107 (lost on re-entry) SORCE SciSat EOS-Aura (formerly EOS-Chem)

NOAA-18 MetOp-2 (MetOp-A) FY-3A GOSAT (formerly GCOM-A1)

24 March 1998 13 May 1998 10 July 1998 1999

10 July 1998 21 September 2000 20 February 2001 21 September 2001 10 December 2001 1 March 2002 24 June 2002 14 December 2002 16 January 2003 25 January 2003 13 August 2003 15 July 2004

20 May 2005 19 October 2006 27 May 2008 6 February 2009


13 November 1997

Launch date

Table 3.

(Continued )

OM-2 UV radiometer (SBUV technique) SBUV/2, HIRS/3 OSIRIS, SMR TOMS-5 SAGE-3 GOMOS (starlight occultation limb scanning) SBUV/2, HIRS/3 ILAS-2 SOLSE-2 (HH payload), LORE-2 (Limb-viewing spectrometers) SOLSTICE ACE-FTS OMI, HIRDLS, TES, MLS (OMI-SBUV technique; data similar to earlier TOMS data, but higher horizontal resolution; HIRDLS and MLS-limb scanners; TES-limb and nadir scanner) SBUV/2, HIRS/4 GOME-2, HIRS/4 TOM or TOU, and SBUS TANSO-FTS (similar to SciSat’s ACE-FTS) (ODUS-Ozone Dynamics UV Spectrometer (an SBUV instrument) was planned but not used)

Shuttle Ozone Limb Sounding Experiment (SOLSE)-1 (HH payload), Limb Ozone Retrieval Experiment (LORE)-1 POAM-3 (solar occultation limb scanning) HIRS/3 OLME SAGE-3?


International Journal of Remote Sensing 5579

2012? 2016? 2016–2030 2018–2026

6 February 2009 17 September 2009 15 June 2010 4 November 2010 28 October 2011 Future satellites with 2012–? 201? 201? 201?


NOAA-19 Meteor-M1 (or Meteor-M) (replacement for Meteor-3M) Picard FY-3B NPP/JPSS series (formerly NPOESS series) ozone equipment FY-3C Meteor-M2 CX-1 microsatellite (University of Colorado) DSCOVR (formerly Triana, or ‘Gore Sat’) (to rest at Lagrange point) MetOp-B MetOp-C MTG-I MTG-S

(Continued ).

Launch date

Table 3.

GOME-2, HIRS/4 GOME-2, HIRS/4 IRS IRS, UVN

TOM or TOU, and SBUS MTVZA imager/sounder; IRFS-2 advanced IR sounder (BUV?) spectrophotometer BUV?

SBUV/2, HIRS/4 MTVZA imager/sounder; IRFS-1 PREMOS TOM or TOU, and SBUS OMPS (nadir-viewing UV sensor and limb-viewing UV/visible sensor)


5580 A.P. Cracknell and C.A. Varotsos


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As mentioned in the Introduction ozonesondes provide direct measurements of ozone concentrations as a function of height (up to 35 km) in the atmosphere (OVP) and therefore may also provide estimates of TOC, assuming an upper atmospheric column using climatological ozone data. But these measurements are made only at a few geographical locations and often also somewhat infrequently. Dobson spectrophotometers, and other ground-based instruments (e.g. Brewer and Bruker Fourier transform infrared– Fourier transform infrared (FTIR) spectrometers), have been used for several years and provide quite long-term datasets, but the question of the inter-comparison of data from the various early laboratories is very important and it has been discussed by Staehelin, Vogler, and Brönnimann (2009) and Strong et al. (2013). However, Dobson spectrophotometers and other ground-based instruments only provide data for rather few geographical locations. On the other hand, as we have already noted, satellite-flown instruments provide frequent data over enormous areas, approaching global coverage. All these instruments need some algorithm to calculate the OVP or TOC from the actual measurements made by the instruments. We may loosely describe this as validation of the measurements obtained by a remote-sensing instrument. Then, once an algorithm has been established for a given instrument the results obtained using the algorithm need to be validated. Ozonesonde data and data from ground-based spectrophotometers therefore play a key role in algorithm development and validation. Ground-based measurements constitute a key component of the Global Ozone Network, both on their own account and by providing the ground truth for satellite-based instruments. The benefit of ground-based instruments is that it is easy to maintain them in good condition, whereas the benefit of satellite-based instruments is that they provide much better spatial coverage and resolution. One crucial parameter for the reliability of the ozone measurements is the selection of the ozone absorption coefficients for the in-field ozone observations. Usually, knowledge of the accurate absorption cross sections of ozone in the UV and IR band is of crucial importance for the reliability of the results obtained from the remote-sensing applications (Orphal 2003). For this reason, in the last decade several inter-comparisons of the ozone intensity were carried out and the results obtained were not consistent, thus raising a debate among different groups of atmospheric remote-sensing scientists (e.g. Gratien et al. 2010). Additionally, the inconsistency between UV cross-section and IR line intensities reach up to 5% and can lead to inconsistencies in the retrieval of ozone, especially in the lowermost troposphere (Cuesta et al. 2013). The validation of some of the satellite-flown instruments that we discussed in section 2 is discussed in some detail in chapter 3 of Cracknell and Varotsos (2012). There we consider the inter-comparison of Dobson, ozonesonde, and satellite data and – as an example – we consider in some detail the case of the Dobson spectrophotometer of the University of Athens. Scatter diagrams for TOMS data and Athens Dobson spectrophotometer data and for SBUV data and Athens Dobson spectrophotometer data from March to December 1991 are shown in Figures 4 and 5, where R2 and the square root of the mean-square error (RMSE) are 0.89 (3.8) and 0.76 (4.3), respectively. In both figures only total ozone data collected during clear sky days were considered to avoid degradation of the data obtained by uncertainties in particle-scattering coefficients (Komhyr and Grass 1972). For such a comparison, the ozone data obtained under cloud-free conditions should be used, because clouds cause an underestimation of total ozone stemming from the redirection of photons scattered within the cloud and emerging downwards from its base (Slusser et al. 1999). It should be stated that the current version of the TOMS total ozone dataset is version 8. Then, as an example of the validation of ozone data obtained from a satellite-flown instrument, we consider in some detail the validation of ozone data from MIPAS, which

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Figure 4. Scatter diagram for TOMS (version 7) and Dobson observations at Athens, Greece (1 March 1991–31 December 1991). Only cloud-free days were included (Varotsos, Alexandris, and Chronopoulos 1999).

Figure 5. Scatter diagram for SBUV (version 7) and Dobson observations at Athens, Greece (1 March 1991–31 December 1991). Only cloud-free days were included (Varotsos, Alexandris, and Chronopoulos 1999).

was flown on Envisat (Cortesi et al. 2007). This involves data from a wide variety of other instruments (e.g. for temperature see Ridolfi et al. 2007). In this context, according to Fioletov et al. (2008) the seasonal median difference between all Brewer direct Sun measurements for a decade over the northern mid-latitudes and GOME and OMI overpasses were around ±0.5% and the systematic differences between the analysed satellite instruments are within ±2%–±3%. Finally, in Cracknell and Varotsos (2012) we present other inter-comparisons between data sets obtained recently from various ozone monitoring systems. We also consider a few examples of what is involved in the inter-calibration of different ozone data sources and the validation of satellite-derived ozone datasets. In particular, we discuss the inter-calibration of Dobson spectrophotometer data with the validation of TOMS, SBUV, OMI, and SCIAMACHY data. Detailed information about


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the contents of our book (Cracknell and Varotsos 2012) along with a few chapters of it may be found in its electronic version: books.google.co.uk/books?isbn=3642103340.

4. Observations of TOC variability TOC is characterized by significant temporal and spatial variability. At any given location TOC varies rapidly from day to day, see Figure 6. As far as temporal variability over the longer term is concerned, it consists of large periodic and aperiodic components (Antón et al. 2011; Cracknell and Varotsos 1994, 1995; London 1982; Varotsos 1989, 2005; Varotsos, Ondov, and Efstathiou 2005). A very important feature of global TOC distribution is the strong latitudinal gradient with lower values over the equator and tropics and higher values over mid- and high latitudes. This gradient is characterized by a well-pronounced annual cycle, reaching a maximum in spring and a minimum in autumn. The amplitude of this annual cycle is a function of latitude, with a maximum at about 60° north and south latitude. In the tropics, seasonal variations are small, with ozone maxima in summer. Such a latitudinal distribution results from the relatively long lifetime (months to years) of ozone in the lower stratosphere and the Brewer–Dobson circulation that transports stratospheric ozone from the tropics towards the poles and downwards at high latitudes. Superimposed on the annual variations that have, presumably, been present since preindustrial times is the recent ozone depletion due to human activities. There are two aspects to ozone depletion arising from the escape of manufactured ozone destroying substances into the atmosphere. The first is a general depletion all over the globe; the

Figure 6. TOC observations at the NOAA/CMDL, Pt. Barrow Observatory during March 1997. The blue values are daily averages for days when weather permitted observations. The mean March TOC for the past 10 years (black) and the previous March daily low values (red) are indicated in the figure (NOAA).

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Figure 7. The TOC at 30 November 1999 assimilated using observations (GOME on-board ESA’s ERS-2). The ‘mini-hole’ is the patch of very low ozone values (just under 200 DU, coloured dark blue) above the North Sea, around 10° E and 55° N; 0° is pointing downwards. The grey patch over the North Pole is an area marked as ‘no data’ (source: GOME Fast Delivery Service © KNMI/ESA, 2000).

extent of this varies according to latitude, being quite small in the tropics rising to several per cent per decade in high latitudes, see Figures 7 and 8. Second, there is the dramatic ‘ozone hole’, which appears in the Antarctic spring and persists for a few months and more recently it has begun to appear in the Arctic spring too (e.g. Varotsos 2002, 2003; Varotsos, Cracknell, and Tzanis 2012; Manney et al. 2011). This is a massive ozone loss, falling from around 400 DU to below 230 DU and taking several months to recover.

5. Ozone-vertical profile (OVP) variability In the troposphere, the ozone concentration and solar UV radiation fall (on the average) with increasing altitude until the tropopause is reached (Katsambas et al. 1997; Varotsos, Kalabokas, and Chronopoulos 1994; Varotsos et al. 1995), see Figure 1. In the stratosphere, ozone concentration increases rapidly with altitude to a maximum near 50 hPa, with a secondary maximum often appearing in the lower stratosphere near 100 hPa. Reliable information on OVP variability in the stratosphere is very important for solving a number of problems, such as ozone impact on climate change and ozone


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Figure 8. Latitudinal distribution of TOC depletion as derived from observations between the periods 1964–1980 and 1984–1993 (after Bojkov and Fioletov (1995)) in comparison with the IPCC modelled values.

depletion due to emissions of man-made chemicals (see Section 6). One feature is a laminar structure. Dobson (1973) examined the occurrence of a laminar ozone structure in the stratosphere over a wide latitudinal and longitudinal range. In Dobson’s analysis the criterion for the detection of laminae within a certain height interval was the change in ozone partial pressure to a value greater than 3 mPa. Dobson (1973) also observed a characteristic ozone minimum in the OVP at 14–17 km, but no explanation was given for both the incidence of minima at the preferred height region and its constancy with latitude. It was stated though that there is a strong correlation between the existence of the characteristic ozone minimum at 14–17 km and the occurrence of the double tropopause, especially at latitudes around 40°N. The classification scheme for laminations used by Dobson (1973) has been extended by several workers (Križan and Laštovička 2006). Subsequently, special features in the OVP structure, especially in the lower stratosphere, have received significant scientific attention in the past few decades. The analysis of TOMS data from October 1978 to September 1987 made it possible to identify basic features of ozone annual and semi-annual variations in the lower, mid-, and upper stratosphere for the latitude interval 65°S–65°N. Further studies of lamination were made using ozonesonde data from Athens; details of the ozonesondes used are given by Varotsos, Kalabokas, and Chronopoulos (1994); Varotsos et al. (1995); Varotsos, Alexandris, and Chronopoulos (1999), and Orsolini, Simon, and Cariolle (1995). Substantial contributions to OVP studies were presented at the International Ozone Symposium in 1996. Grainger and Atkinson (1998) carried out the first comprehensive global three-dimensional 6 hourly analyses of ozone mixing-ratio data from satellite observations (MLS HALOE, SAGE II, SBUV/2, TOVS), which were interpolated on to a 2.5° grid with 19 levels between 1.000 and 0.1 hPa (data for October 1994 were considered). There are still, however, some problems with the harmonization of observational data from various platforms. Chan, Liu, and Lam (1998) discussed OVP observations (ozonesonde data) in Hong Kong, which demonstrate a complicated OVP structure (basically bimodal with a laminated structure when observed in the lower stratosphere in winter and spring).

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A significant contribution to studies of OVP has been made by various ground-based and aircraft remote-sensing observations with a special role played by lidar soundings. For instance, the airborne UV Differential Absorption Lidar (DIAL) system participated in the Tropical Ozone Transport Experiment/Vortex Ozone Transport Experiment (TOTE/ VOTE) in late 1995/early 1996 (Grant et al. 1998). Several options of DIAL systems have been developed, e.g. the Airborne Excimer Ozone DIAL (ABDIAL) and Tunable Optical Profiler for Aerosol and oZone lidar (TOPAZ) aboard aircrafts (Alvarez II et al. 1996) and the Ozone Profiling Atmospheric Lidar (OPAL) aboard ships (Zhao, Hardesty, and Post 1992). Before leaving the discussion of OVPs in this section there are two other points that we should mention briefly; the first is stratospheric/tropospheric exchange of ozone and the second is low ozone pockets. Numerous studies have shown that there is an increasing trend of background values of ozone in the troposphere. This is of serious concern because of the damage caused by high ozone concentrations to human beings, animals, and plants, and also because of possible climatic effects. Stratospheric ozone leaks into the troposphere forming a natural background. In addition man-made and natural emissions of nitrogen oxides and hydrocarbons lead to the production of ozone in the sunlit troposphere. Until recently, knowledge of the behaviour of ozone in the troposphere and transfer across the upper troposphere/lower stratosphere (UTLS) boundary was mostly obtained from ground stations operating ozonesonde balloons and from some data obtained from aircraft-flown instruments. More recently, information has become available from satellite systems. Changes in the stratosphere–troposphere exchange (STE) of ozone over the past few decades have altered the tropospheric ozone abundance and are likely to continue doing so in the coming century as climate changes. An important issue in determining tropospheric ozone concentrations from space-borne data is to discriminate accurately between the troposphere and the stratosphere. It is worth noting that the tropospheric ozone can only be retrieved from satellite measurements if we are able to disentangle the contribution of stratospheric column. Lately the merged usage of measurements from Thermal Infrared (TIR) and Ultra-violet Visible Near infrared Shortwave infrared (UVNS) has been shown to be essential in improving the retrieval of ozone in the lowermost troposphere (Cuesta et al. 2013). Recently, direct observation of the tropospheric ozone profile has become possible using data from systems such as GOME, HIRDLS, MLS, and OMI. Strategically designed ozonesonde networks have transformed sampling in the upper troposphere and lower stratosphere with consistent temporal and vertical (100 m) resolution. At the same time, they are essential components of satellite validation and monitoring of ozone profiles (Thompson et al. 2011). In particular two strategic networks – SHADOZ (Southern Hemisphere Additional Ozonesondes) and IONS (INTEX Ozonesonde Network Study) – have been described (see section 6.6.3 in Cracknell and Varotsos 2012). The stratosphere– troposphere exchange ozone flux in 2001–2005 has been estimated as 290 Tg year–1 in the northern hemisphere and 225 Tg year–1 in the southern hemisphere (Witte et al. 2008; Hsu and Prather 2009). We now turn to low ozone pockets. It should already be apparent that ozone spatiotemporal variability is characterized by a high level of complexity. One of the features of this complexity is the low-ozone pockets (Morris et al. 1998). They emphasized that, although the lowest ozone concentrations are typically found in mid-stratosphere within the winter circumpolar vortex, they were also frequently observed in satellite remote sounding using Limb Infrared Monitor of the Stratosphere (LIMS) and Microwave Limb Sounder (MLS) instrumentation outside the polar vortex.


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According to analysis of the LIMS data, ozone pockets are formed during mid- to late winter in conjunction with stratospheric wave-breaking events. Trajectory calculations revealed that much of the air within the pockets originated in the tropics or subtropics at higher altitudes several weeks earlier. For a detailed discussion of these low ozone pockets see Morris et al. (1998). The ozone observations in the lowermost troposphere may be conducted by the thermal infrared (TIR) instrument, on board the Geostationary Earth Orbit (GEO) platform (or GEO-TIR). The importance of probing the lowermost troposphere using GEO-TIR instrument with respect to instrument devoted to temperature sounding is discussed in Claeyman et al. (2011). Finally, it should be noted that new satellite instruments (e.g. MLS, GOMOS, and MIPAS) are able to measure ozone profiles up to the mesosphere. There, among the processes causing O3 variability is that of the Solar Proton Events (SPE) (Damiani et al. 2012). In this regard, highly energetic solar protons that are emitted in extremely large solar proton events (e.g. occurring in 1972, 1989, 2000, 2001, and 2003) caused the production of odd hydrogen (HOx) and odd nitrogen (NOy), which then led to ozone variations (Jackman et al. 2009). 6. Ozone depletion We have already mentioned briefly that there are two components to stratospheric ozone depletion arising from human activities. One is the slow but steady depletion by a few per cent per decade, least serious in the tropics but increasing with latitude. In the northern hemisphere, the decrease is larger in winter and spring (11% since 1979) than in summer or autumn (4% since 1979). A negative trend of the annual variation amplitude (around 0.1 ppm year–1) was observed in the upper stratosphere of southern subpolar latitudes. Inter-annual changes in amplitudes of both annual and semi-annual variations are small as a rule (except in the tropical mid-stratosphere, where the influence of the El Chichón eruption was substantial, and the subpolar upper stratosphere in the southern hemisphere). The OVP trend shows distinct negative trends at about 18 km in the lower stratosphere with the largest declines over the poles, and above 35 km in the upper stratosphere. A narrow band of large negative trends extends into the tropical lower stratosphere (Brunner et al. 2006). Bojkov and Fioletov (1995), based on re-evaluated TOC observations from over 100 Dobson and filter radiometer stations from pole to pole, presented the following TOC trend results for the steady depletion of ozone. – Up to 1995 TOC continued its decline (which started in the 1970s), with statistically significant year-round and seasonal trends except over the equatorial belt. – The cumulative year-round TOC reduction over the 35°–60° belts of both hemispheres from the early 1970s until the mid-1990s was up to 8%. While in the southern mid-latitudes, it is difficult to distinguish the seasonal dependence of TOC trends, the cumulative decline in the northern mid-latitudes in winter and spring is about 9% and 4–6% for summer and autumn. – At that time observations from 12 Dobson polar stations had demonstrated that the northern polar region shows the same ozone decline as northern mid-latitudes or even a slightly stronger one (the cumulative decline is about 7% year-round and 9% for winter and spring).

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Nowadays, according to the recent scientific ozone assessment (WMO 2010) the average TOC values in 2006–2009 remain at roughly 3.5% and 2.5% below the 1964–1980 averages, respectively, for 90°S–90°N and 60°S–60°N. Mid-latitude (35°–60°) annual mean TOC 1996–2005 was at around 6% (around 3.5%) below the 1964–1980 average. The second component consists of the two seasonal ozone holes in the polar regions, the Antarctic ozone hole having appeared earlier and being larger than the Arctic ozone hole (Farman, Gardiner, and Shanklin 1985; Solomon 1999). In winter, in each of the polar regions air becomes trapped in a circumpolar vortex and its temperature drops and when it becomes very cold (below about 195 K or –78°C) polar stratospheric clouds (PSCs) form. These are sometimes called nacreous clouds or mother of pearl clouds because of the colours they exhibit. They consist not of water, but of frozen particles of nitric acid and water, especially nitric acid trihydrate (HNO3.3H2O). When the Sun begins to return at the end of the winter, free chlorine atoms are released as a result of photodissociation, by the solar UV light, of CFCs or chemicals derived from the CFCs. Each chlorine atom acts as a catalyst to the destruction of tens of thousands of ozone molecules. The Cl-catalysed ozone destruction occurs in the gas phase, but it is accelerated by heterogeneous catalysis on the surfaces of the particles in the polar stratospheric clouds. The rates of the heterogeneous chemical conversions depend on both surface and bulk concentrations, diffusion measurements in or on ice (Alexopoulos and Varotsos 1981; Varotsos 2007; Varotsos and Zellner 2010). Given the longevity of CFC molecules, the recovery times for the ozone layer are of the order of several decades (e.g. Monks 2007). In general, the Arctic experiences high extreme cold as well as sudden stratospheric warmings (SSWs) at times. As a result, the degree of ozone loss is mostly controlled by the strength of the vortex and magnitude of the air temperature within it. For instance, according to Kuttippurath et al. (2010) the winters of 1995, 1996, 2000, and 2005 were very cold and the cumulative TOC loss was as high as around 25–35% (WMO 2010). On the other hand, the winters of 1997, 1998, 1999, 2001, 2002, 2006, and 2009 were relatively warm and the loss was minimal – about 10–15%, whereas the winters of 2003, 2007, and 2008 were moderately cold and hence the loss was in an average scale of about 15–20% (WMO 2010; Goutail et al. 2010). As far as the future projection is concerned, the global annually averaged TOC is expected to reach to 1980 levels before 2050. The simulated changes in tropical TOC from 1960 to 2100 are generally small. Chandra, Varotsos, and Flynn (1996), using measurements of TOC by Nimbus-7 TOMS version 7, suggested that the trends over the latitudes centred at 40° N of the Northern Hemisphere vary from –3% to –9% per decade during winter and within –1% to –3% per decade during summer. Long-term measurements of the stratospheric ozone with balloon-borne instruments allow winter ozone altitude profiles to be compared between the Antarctic and Arctic regions (WMO 2010). Inspection of Figure 9 shows that in the Antarctic at the South Pole (left panel), a normal ozone layer was observed to be present between 1962 and 1971. What is particularly interesting about the left-hand half of Figure 9 is that, as shown here for 9 October 2006 in spring over Antarctica, the ozone is almost completely destroyed between 14 and 21 km. In the past two decades (1990–2009) the average ozone concentrations in October are found to be 90% lower than the pre-1980 values at the peak altitude of the ozone layer (16 km). In contrast, the ozone layer over the Arctic region does not exhibit any depletion, as shown by the average ozone profile in March for 1991–2009 obtained over the Ny-Ålesund site


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Figure 9. Vertical distribution of Arctic and Antarctic ozone; TOC values given in brackets (Polar Ozone Depletion) (source: WMO 2010).

(right panel). No Ny-Ålesund data are available for the 1962–1971 period before significant ODS destruction began. Some March profiles do reveal significant depletion (e.g. the case of 29 March 1996). In these cases, minimum temperatures during wintertime are generally lower than normal, allowing PSC formation for longer periods. Arctic profiles with depletion similar to that shown for 9 October 2006 at the South Pole have never been observed. The number in parentheses for each profile is the TOC value in DU. Ozone abundances are shown here as the pressure of ozone at each altitude using the unit ‘millipascals’ (mPa).

7. Dynamics of atmospheric ozone Measurements of TOC and OVP provide information about ozone at a particular place and time and tell us nothing about how ozone moves around (e.g. Lee and Feldstein 2013). Olsen et al. (2008) described HIRDLS data and investigation of the stratospheric air intrusion from the tropics to high latitudes. On 26 January 2006 low mixing ratios of ozone and nitric acid were observed in a 2 km layer near 100 hPa extending from the subtropics to 55° N over North America. The subsequent evolution of the layer was simulated with the Global Modelling Initiative (GMI) model and confirmed by HIRDLS observations. Very important progress in studying both ozone dynamics near the tropopause and stratosphere–troposphere ozone exchange resulted from the work of the Measurement of Ozone and Water Vapour by the Airbus In-Service Aircraft (MOZAIC) programme over a 2 year period, from September 1994 to August 1996 (Thouret et al. 1998). Thompson et al. (2011) have recently reported on the results obtained by employing OVP data from the Arctic Intensive Ozonesonde Network (ARC-IONS) over Canada, Alaska, and the mid-upper USA, which involved 18 sites, most launching daily

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ozonesondes (http://croc.gsfc.nasa.gov/arcions). The IONS data are used for forecasting and flight planning during the field phase, for determining ozone budgets, satellite validation, and evaluation of chemical-transport models of various scales. Thompson et al. (2011) suggested that whereas the Canadian air-quality forecast models AURAMS and CHRONOS show considerable skill at predicting ozone in the planetary boundary layer and just above, they have large errors in the free troposphere, owing largely to the inadequate treatment of model domain boundaries. From what we have said already in this review, it can be seen that there are many sources of data on atmospheric ozone concentrations using instruments on the ground, airborne instruments, and instruments on satellites. However, the concentration of atmospheric ozone changes rapidly, both spatially and temporally. Two factors are involved. First, ozone is continuously being generated and destroyed by solar UV radiation. Second, the atmosphere is constantly in motion. It is not easy to separate these two aspects. Assessments of relative contributions of photochemical and dynamical processes to ozone concentration field formation confirmed earlier conclusions (e.g. Gabriel et al. 2011). Ozone distribution in the lower stratosphere (below the 30 hPa level) is controlled (especially in winter) mainly by large-scale atmospheric circulation, whereas in the upper stratosphere (above the 5–10 hPa levels) radiative and photochemical processes dominate (except at subpolar latitudes in winter when atmospheric transport is also important). In the tropics this zone is located at a lower level because of the shorter photochemical lifetime in the equatorial belt under conditions of high insolation. Thouret et al. (1998) reached the conclusion that north of 35°N the distribution at 9–12 km altitude is dominated by the influence of ozone-rich air of stratospheric origin; farther south, ozone-poor air from the troposphere prevails. To identify the stratospheric and tropospheric components and to help in interpreting the data a classification based on a threshold of 100 ppbv of ozone was used. While it is possible to use data from many different sources to determine the concentration of ozone at one particular point and at one particular time, to follow the movement of ozone is much more difficult. One serious attempt to do this has been made by the Match Network. This originated from the 1991–1993 European Arctic Stratospheric Ozone Experiment (EASOE) campaigns (Von Der Gathen et al. 1995) because there was already substantial coverage of mid- and high-latitude ozone sounding stations in the northern hemisphere. Several dozen of these stations agreed to launch ozonesondes on schedule over the period when ozone loss is the greatest in the Arctic vortex, which is typically from December to mid-March. The basic idea of Match is to determine stratospheric polar ozone losses by observing individual air masses using ozonesondes at two stations over which the same air mass passes; the loss of ozone can then be determined by the difference between the two ozonesonde profiles. The name ‘Match’ thus comes from the attempt to match two ozonesondings that correspond to the same air mass at different instants in its lifetime. In the frame of EASOE a large number of ozonesonde ascents were performed in the region from the mid-latitudes to the Arctic. The measurements collected were used for the estimation of the Arctic ozone losses by employing the Match method. Because of the large number of soundings made during EASOE enough matches were found for a successful analysis and in this way the Match method was established. The method was then developed into an active method, to reduce the number of ozonesondes launched, for the Second European Stratospheric and Mid-latitude Experiment (SESAME 1998, 1999) in the winter of 1994–1995. Since 1994–1995 the Alfred Wegener Institute for Polar and Marine Research (AWI) has coordinated the Match


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network campaign in most Arctic winters and in the first Antarctic winter (in June– November 2003) in order to determine the chemical ozone depletion in the stratosphere. The objective of Match, i.e. to probe the ozone loss of air parcels, is achieved by coordinating the soundings. The element of chance that was present in the original EASOE experiment was eliminated by having the subsequent trajectories of the air masses that had passed over one ozonesonde station forecast by meteorologists at the Free University (FU) in Berlin. Thus when this trajectory passes over a second ozonesonde station, the staff there can be asked to perform an ozonesonde ascent in order to measure again the ozone content in this air mass. An ozone decrease occurring during the two ozonesonde ascents can be attributed to chemical processes and therefore the Match technique can distinguish between dynamical and chemical components. The Match technique has been applied by Sasano and coworkers to the observations made by the Improved Limb Atmospheric Spectrometer (ILAS) for determining ozone loss rates (Sasano et al. 2000; Terao et al. 2002).

8. Conclusions and future requirements We have presented a review of the satellite systems for the atmospheric column ozone and OVP observations. The accuracy and precision that accompany the measurement of ozone concentration at various atmospheric layers depend on the instrumentation and the technique employed. This is very important in the cases where the monitoring sensor is far away from the observer as it happens when remotely sensed ozone datasets are used. There is no doubt that the recognition of the stratospheric ozone depletion from one side and the increase of the tropospheric ozone amount from the other side, as crucial environmental problems, has led to the development of advanced observing instrumentation flown on satellite platforms in order to have substantial temporal and spatial covering of the global ozonesphere. The main benefit from the extensive analyses performed to date exploiting the available ozone datasets is the current understanding about the physicochemistry of the global ozone layer that has been recognized in the award of the Nobel prize in Chemistry in 1995. It was awarded jointly to Paul J. Crutzen, Mario J. Molina, and F. Sherwood Rowland for their pioneering contributions to explaining how ozone is formed and decomposes through chemical processes in the atmosphere. Nevertheless, important problems still remain and urgently need to be addressed. One of them is the non-linear behaviour of the chemical and dynamical mechanisms that play a fundamental role in the atmospheric ozone variability. To gain experience, however, from this complex ozone system more observations with special accuracy and precision are needed. Satellites can contribute a lot to this end with the employment of new observational systems of ozone.

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