Diurnal cycle of ozone throughout the troposphere over Frankfurt as measured by MOZAIC-IAGOS commercial aircraft

Ozone (O₃) is a trace gas of major importance in the atmosphere due to its adverse effects on health, vegetation and climate change. In the troposphere, O₃ is produced photochemically by the oxidation of methane, carbon monoxide (CO) and non-methane hydrocarbons (NMHCs) in the presence of nitrogen oxides (NO_x=NO+NO₂) and sunlight, or transported downward from the stratosphere. It is removed by dry deposition, titration by NO and reactions with hydrogen oxide radicals (HO_x=OH+HO₂) (Monks, 2005). Ozone is lost in low-NO_x environments, such as the marine boundary layer. Halogens also play a role in the ozone chemistry, directly through a reaction with O₃ itself, or indirectly through a reaction with NMHCs (production or destruction of O₃ depending on the conditions) (Monks et al., 2015).

During the last several decades, numerous studies have investigated the causes and characteristics of the diurnal variation of surface O₃ at a wide range of locations (Singh et al., 1978; Oltmans and Levy, 1994; Pryor and Steyn, 1995; Zhang et al., 2004). At low altitudes over continental regions under the influence of anthropogenic emissions, ozone generally exhibits minimum mixing ratios during night, mainly due to dry deposition and titration by NO in the shallow nocturnal boundary layer (BL). During daytime, mixing ratios increase under the combined effect of photochemical production and mixing with O₃-rich air masses from the residual BL, reaching a maximum in the afternoon. The diurnal cycle typically has its maximum amplitude in summer when photolysis rates are the highest, while diurnal cycles observed in winter are less pronounced. Over regions with a more complex orography, the local dynamics can play a more important, if not dominant, role in the diurnal variability of O₃. At high-altitude mountain sites, observations usually show a reverse cycle, with maximum mixing ratios during night (e.g. Bonasoni et al., 2000; Cristofanelli et al., 2013; Ezcurra et al., 2013; Zhang et al., 2015). This pattern is the result of the mountain peaks being heavily influenced by the lower free troposphere at night where ozone mixing ratios are typically greater than at the surface if the mountain is located in a rural or remote region. During the day upslope flow caused by the heating of the mountain slopes brings ozone depleted air from the BL up to the summit (Oltmans and Komhyr, 1986; Oltmans et al., 2013). In the marine BL, daytime O₃ can be destroyed by photolysis, reactions with HO_x and halogens or dry deposition (Ganzeveld et al., 2009). During summertime, when low-NO_x conditions prevail over oceans, these sinks dominate the O₃ production by photolysis of NO₂ and the entrainment from aloft in the free troposphere (FT), usually leading to minimum mixing ratios in the afternoon (Ayers et al., 1992; Monks et al., 1998, 2000).

As investigating the O₃ diurnal variability requires a sufficiently high frequency of measurements (typically 1h), all the previous studies have been restricted to the surface. In the FT, neither satellite observations nor ozone soundings are frequent enough to derive a complete diurnal profile of O₃. In principle, Lidar measurements over a sufficiently long period may be used to investigate the diurnal cycle of O₃ at multiple levels in the troposphere (although with more uncertainties than considering in-situ observations) but to our knowledge, no such studies exist (probably because such Lidar observations are complex to operate routinely during long periods). The changes of diurnal variability with altitude have been investigated during some campaigns with several instruments deployed at different altitudes along a mountain slope, showing a decrease of the O₃ diurnal variability with altitude, and a slight reverse at the highest elevations with greater mixing ratios during night (Ezcurra et al., 2013; Burley et al., 2015; Brodin et al., 2010). However, the previously mentioned strong influence of the local dynamics in such environments prevents extrapolation of these results to the FT.

Airborne observations of O₃ (among other species) have been measured by the MOZAIC-IAGOS program since 1994 using commercial aircraft from several airlines. Vertical profiles from the surface to 11–12 km are obtained as the aircraft operationally descend and ascend from airports. Among the 200 airports visited worldwide, the Frankfurt, Germany airport is by far the most frequently visited, with approximately 21,600 flights between 1994 and 2012, which corresponds to 98 flights per month, on average. During this long period, airline schedules have varied enough to have yielded observations that are reasonably well-distributed throughout the day. In this paper, we take advantage of these high-frequency measurements to investigate the diurnal variations of O₃ throughout most of the troposphere (0–12 km).

The observational dataset is described in Sect. 2. Section 3 presents the results, including the diurnal profiles of O₃ in the troposphere and their change over the last two decades. Results are discussed in Sect. 4.

The MOZAIC (Measurements of OZone and water vapour on Airbus In-service airCraft) program, initiated in 1994 and incorporated into the IAGOS (In-service Aircraft for a Global Observing System; www.iagos.org) program since 2011 takes advantage of commercial aircraft to provide worldwide in-situ measurements of several trace gases (e.g., O₃, CO) and meteorological parameters (e.g. water vapour) throughout the troposphere and the lower stratosphere (Marenco et al., 1998; Petzold et al., 2015; Nédélec et al., 2015). Ozone measurements are performed using a dual-beam UV-absorption monitor (time resolution of 4 seconds) with an accuracy of ±(2 ppbv+2%) (Thouret et al., 1998). More details on the new IAGOS instrumentation can be found in Nédélec et al. (2015). The continuity of the dataset between the MOZAIC and IAGOS programs has been demonstrated based on the 2-year overlap (2011–2012) (Nédélec et al., 2015).

In this study, we focus on the vertical profiles above the Frankfurt airport where observations are the most frequent. The coverage and density of the datasets are described in Sect. 2.3. Note that the availability of observations at different times of day is due to several MOZAIC-IAGOS aircraft simultaneously serving this destination. Some other airports have also been frequently visited by MOZAIZ-IAGOS aircraft (e.g. New York, Paris, Windhoek, Vienna, Tokyo), but usually by only one aircraft during specific time frames, resulting in too few profiles to properly investigate the diurnal variations of O₃.

During landing and take-off phases, observations are obtained through most or all of the troposphere, and often part of the lower stratosphere. The time resolution of 4 seconds roughly corresponds to a vertical distance of 30 m. As landing and take-off do not exactly correspond to vertical profiles, only the observations within the 400 km around the Frankfurt airport are considered, which limits the possible spread of O₃ mixing ratios due to horizontal heterogeneity. This study focuses on tropospheric O₃, and observations at the tropopause level and in the stratosphere are thus ignored. However, stratospheric intrusions within the troposphere are retained (i.e. vertical profiles are not purely tropospheric). Here, the tropopause is defined as the 30 hPa-width layer centred on the 2 pvu (potential vorticity units) iso-surface, following the method of Thouret et al. (2006). The PV values are extracted along each vertical profile from the European Centre for Medium-Range Weather Forecasts (ECMWF) operational analysis (00:00, 06:00, 12:00, 18:00 UTC) and forecasts (03:00, 09:00, 15:00, 21:00 UTC). More details on the procedure are given in Petetin et al. (2015).

Data are aggregated on 50-hPa pressure surfaces from 1000 to 300 hPa. For the construction of the diurnal profile, we consider time intervals of 3 hours: 0–3, 3–6, 6–9, 9–12, 12–15, 15–18, 18–21, 21–24 UTC. For convenience, hours in this paper are expressed in UTC; note that in Germany, the local solar time (LST) corresponds to UTC+1 from November to March (included), and UTC+2 otherwise (the 0–3 UTC time interval thus roughly corresponds to the closing times of the airport). Diurnal profiles are calculated at both the seasonal (DJF: December-February, MAM: March-May, JJA: June-August, SON: September-November) and annual scale (ANN: annual). Note that observations are far too sparse during the 0–3 UTC time interval and are therefore ignored in this paper. The procedure is applied to several metrics, including the mean and 5^th, 25^th, 50^th, 75^th and 95^th percentiles.

The number of O₃ observations available over the 1994–2012 period for the calculation of the diurnal cycles is reported in Table S1 in the Supplement. On average, each point of the diurnal cycle (corresponding to a given 3-hour time interval and a given 50-hPa pressure level) relies on about 52,000 observations. However, this density strongly varies depending on the time of day. The densest sampling is available between 3 and 12 UTC (above 80,000 observations), followed by 12–15 UTC (around 54,000) and 15–18 UTC (around 29,000). A substantially lower density occurs between 18 and 24 UTC (below 16,000 observations). The sampling is reasonably well balanced between the different seasons, with 29, 22, 29 and 20% of the data obtained in winter, spring, summer and autumn, respectively.

In terms of spatial coverage, observations available before 18 UTC are mostly obtained from flight routes north of Frankfurt, while observations after 18 UTC are typically south of Frankfurt. Considering that (i) the tropopause height decreases as one moves northward and that (ii) there are some uncertainties associated with the estimation of PV, this sampling bias may result in more frequent occurrences of stratospheric contamination before 18 UTC than after. However, due to the fact that only the troposphere is considered in this study and that we take into account only the observations within 400 km of the Frankfurt airport, this source of uncertainty is expected to be low.

It is worth noting that some uncertainties may arise from the temporal discontinuity of this database. Indeed, measurements are not performed continuously throughout the day (as occurs at a surface station), but sporadically between 1994 and 2012 with frequencies that vary from one 3-hour time interval to another. The data coverage at 800 hPa is shown in Fig. 1 for the different time intervals. Between 3 and 15 UTC, observations are available during almost all the 1994–2012 period. At 15–18 UTC, observations are sparse between 2005 and 2010 (inclusive). The least dense coverage concerns the 18–21 UTC time interval, and to a lesser extent the 21–24 UTC. Considering the strong inter-annual variability (IAV) of O₃ (Thouret et al., 2006; Zbinden et al., 2006), this may introduce some variability unrelated to the diurnal cycle. For instance, at 800 hPa, contrary to the other time intervals, no observations are available at 18–21 UTC during the summer 2003 when a severe heat wave induced a strong positive O₃ anomaly over Europe, mostly in the BL (Solberg et al., 2008; Tressol et al., 2008; Ordóñez et al., 2005). The absence of MOZAIC-IAGOS observations during that season thus produced a low O₃ bias. Depending on the extent of this bias, it may lead to an artificial diurnal variation.

In order to assess the potential importance of this bias, sensitivity tests were performed based on continuous O₃ measurements at the surface in the region of Frankfurt. We considered 3 stations of the German Federal Environment Agency (Umweltbundesamtes, UBA) – the Rannheim (50.01°N, 8.45°E, 90 m; station code DEHE018; located 5 km westward from the Frankfurt airport) and Darmstadt (49.97°N, 8.66°E, 158 m; DEHE001; 12 km southward) urban background stations, and the Spessart regional background station (50.16°0N, 9.40°E, 502 m; DEHE026; 50 km eastward) – as well as the Waldhof (52.80°N, 10.77°E, 74 m; 300 km north-eastward) regional background station from the Global Atmosphere Watch (GAW) network. For each of these surface stations, hourly observations of O₃ are available continuously from October 1994 to December 2012. The seasonal and annual diurnal cycles are then calculated over 3-hour time intervals (as explained in Sect. 2.2) using this surface O₃ dataset, but with three different types of data selection. First, as a reference case, we use all the available data between 1994 and 2012 (hereafter referred as REF). Second, we consider only the months with at least 1 observation available at 1000 hPa at the corresponding 3-hour time interval in the Frankfurt MOZAIC-IAGOS database (hereafter referred as SUB1). Third, similarly to SUB1, we consider only the months with at least 50 observations available at 1000 hPa n the Frankfurt MOZAIC-IAGOS database (hereafter referred as SUB50). In both the SUB1 and SUB50 cases, we thus artificially create the same data gaps present in the MOZAIC-IAGOS dataset in our surface O₃ dataset. Comparing the REF diurnal cycles with the SUB1 and SUB50 allows us to quantify the bias introduced by the absence of observations during some periods. Results are shown in Fig. S1 in the Supplement. Whatever the season and the station, the differences of 3-hour mean mixing ratios between the REF and SUB1 (SUB50) diurnal cycles remain below ±3% (±13%). They are usually greatest at 18–21 UTC, when the MOZAIC-IAGOS data coverage is sparsest. These low differences demonstrate that the data gaps in the MOZAIC-IAGOS dataset likely have a limited impact on the diurnal cycles calculated at Frankfurt, at least close to the surface where these stations are expected to be reasonably well representative (in particular the rural background ones). Similar results are obtained when considering the data coverage of MOZAIC-IAGOS at the other pressure levels. However, it is worth noting that at altitude, these surface observations cannot be considered to be representative.

The diurnal profiles of mean tropospheric O₃ over the 1994–2012 period are shown in Fig. 2 at both the seasonal and annual scale. Note that compared to the uncertainty associated with each individual measurement (estimated at ±(2 ppb+2%), see Sect. 2.1), the uncertainties associated with the mean O₃ mixing ratios are largely reduced due to the large number of points (see Table S1 in the Supplement), all values remaining below 0.1 ppb.

As expected, the strongest diurnal variations occur in the BL due to the more active O₃ production and loss processes close to the surface. The BL diurnal cycle reaches its highest amplitude in summer (25 ppb at 1000 hPa) when the photochemical production is at a maximum, followed by spring (17 ppb) and autumn (12 ppb), while very low variations are observed in winter (4 ppb). The diurnal variation propagates higher in altitude in summer due to the development of a deeper convective BL. In particular, the O₃ maximum at 15–18 UTC extends to all pressure levels up to 800 hPa. At 1000 hPa and to a lesser extent at 950 hPa, mixing ratios quickly decrease after 18 UTC, down to 20 ppb at 3–6 UTC, close to the surface. This is likely due to the combined effect of dry deposition and titration by NO (both sinks being enhanced in the shallow nocturnal BL), as suggested by the much lower decrease in the next level. In the early morning (6–9 UTC), the O₃ vertical gradient between 1000 and 950 hPa is slightly reduced compared to the night. This is likely due to the turbulence that starts to mix low-O₃ air masses from the nocturnal BL with high-O₃ air masses from the residual BL (Zhang and Rao, 1999). Later, this O₃ vertical gradient keeps decreasing as the turbulent mixing increases, which progressively homogenizes the O₃ mixing ratios in the BL.

Higher in altitude, there is an absence of strong O₃ diurnal variations, regardless of season. Some variations are still observed at some pressure levels after 18 UTC, and in particular an increase at 18–21 UTC above 350 hPa. This may be an artefact related to the sparser observations at this time of day. Interestingly, based on the observations at the Waldhof station, we previously highlighted (in Sect. 2.3) that this lower frequency of measurement may introduce a moderate positive bias in the diurnal cycle at 18–21 UTC. However, it is not possible to draw definitive conclusions on this point since this bias was mainly observed during the winter and is likely limited to the lower troposphere (i.e. the representativeness of the Waldhof surface station). In any event, coefficients of variation (CV) – defined as the standard deviation of the diurnal cycle normalized by its mean – quickly decrease from 10–30% at 1000 hPa to less than 4% above 950 hPa in winter, 850 hPa in spring, 800 hPa in summer and 900 hPa in autumn. Above 800 hPa, CV range between 0.6 and 3.3%. The highest values are found during the winter at the uppermost pressure levels while most of the CV values elsewhere remain around 1–2%.

Although small in amplitude, these diurnal variations in the free troposphere may still be statistically significant. For each season and each pressure level, between each pair of points of the diurnal cycle (21 possible combinations among the 7 time intervals), a Smith/Welch/Satterthwaite (SWS) t-test (Smith, 1936; Welch, 1938; Satterthwaite, 1946) (also referred as the unequal variance t-test in the literature) is performed to compare the mean O₃ mixing ratios. Ruxton (2006) explained why this SWS t-test should always be used in preference to the Student’s t-test or Mann-Whitney U test (in particular, it does not assume equal variances between the two populations). Results indicate that most of the differences are statistically significant at a 95% confidence level. On average over all the pressure levels and 3-hour time intervals, about 85, 87, 79, 81 and 74% of the SWS t-tests performed gives significant differences in the DJF, MAM, JJA, SON and ANN diurnal cycles, respectively. This proportion of statistically different mixing ratios varies with altitude, from about 90–95% below 950 hPa to 72–78% around 650–350 hPa. Therefore, throughout the whole troposphere, most of the diurnal variations of O₃ mixing ratios are statistically significant, but both the relative amplitude of these variations and the proportion of significant differences are reduced higher in altitude.

The diurnal profiles of the O₃ 5^th, 25^th, 50^th, 75^th and 95^th percentiles at the annual scale are shown in Fig. 3. As for the mean O₃ mixing ratios, strong diurnal variations are observed at the lowest levels but quickly decrease with altitude. Qualitatively, a similar picture emerges at the seasonal scale, although variations are usually higher (see Fig. S2–S5 in the Supplement). This is particularly true in summer for the O₃ 95^th percentile (Fig. 4), for which the mixing ratios at 21–24 UTC are 5–15 ppb higher than during the rest of the day at several pressure levels in the FT (600–300 hPa). However, the corresponding CVs remain below 6%.

Whatever the season or the metric, most of the O₃ vertical gradients – defined as positive when O₃ increases with altitude – in the troposphere are positive or only slightly negative. However, as illustrated in Fig. 4, a noticeable exception is the O₃ 95^th percentile in summer. During the afternoon, the vertical gradients are -0.22 ppb hPa^-1 between 850 and 800 hPa (i.e. likely within the BL). At this time of year and day, the solar radiation reaches its maximum and is expected to enhance both the vertical transport and the photolysis rates. These negative gradients may thus suggest that the photochemical production of O₃ is faster than the vertical mixing.

We now investigate how the O₃ diurnal profile has changed over the past two decades. The dataset is split into two periods, 1994–2003 and 2004–2012. Fig. 5 shows the diurnal profiles of mean O₃ at the annual scale for both periods, and the relative differences between them (the difference being normalized by the mixing ratios observed during the first period). Note that in this section, we investigate only the changes of O₃, not the trends for which uncertainties would have to be estimated. Overall, the mean O₃ mixing ratios have slightly increased in most of the troposphere and during most of the day. The highest increases are found during the night close to the surface (changes between +21 and +41%). The much lower increase between 9 and 18 UTC (< +7%) leads to a strong decrease of the amplitude of the diurnal cycle, from 15 to 10 ppb. These changes quickly attenuate with altitude, the increase at 950 hPa remaining below +7%. In the second part of the day (after 15 UTC), a slight decrease of the O₃ mixing ratios is observed at some pressure levels, with values ranging between -3 and -1%.

Results at the seasonal scale (Fig. S6–S9 in the Supplement) indicate that the increase of night-time O₃ is observed during all seasons, at least at the first pressure level. It is more accentuated in winter/autumn (from +27 to +95%) than in spring/summer (from +3 to +28%). This increase is likely mainly due to the reduction of NO_x emissions in Europe over the last decades, leading to a lower titration of O₃ by NO (Derwent et al., 2003; Solberg et al., 2005; Jonson et al., 2006). In fact, the wintertime O₃ has increased in the whole troposphere and during the whole day, although at very different extents (from +1 to +20% above the 950-hPa pressure level). A similar picture emerges in spring, although with lower changes. Conversely, the increase in autumn appears confined to the BL (below the 600-hPa pressure level) and is observed throughout the day except at 15–18 UTC where mixing ratios have decreased (from -3 to -9% above the 950-hPa pressure level). Contrary to the other seasons, the O₃ mixing ratios have decreased during summer (apart from the night-time increase at 1000 hPa), in particular below the 700-hPa pressure level where changes from 1994–2003 to 2004–2012 range between -16 and -2%. Above, the decrease is observed only after 18 UTC, although not at all pressure levels. This decrease of summertime O₃ in the BL during daytime partly compensates the increase observed during the other seasons.

The changes of the 5^th and 95^th percentiles at the annual scale are given in Fig. S10–S11 in the Supplement. Interestingly, the relative increase of the 5^th percentile is substantially higher than for the mean O₃ (in particular close to the surface), while the 95^th percentile is shown to decrease in large parts of the troposphere.

The diurnal variations of O₃ mixing ratios are well characterized at the surface in many different environments (e.g. urban, coastal, mountain, marine). However, due to the absence of continuous measurements at higher altitudes, both in the boundary layer and the FT, our knowledge of the O₃ diurnal variations relies on assumptions and/or numerical simulations. One generally assumes that O₃ mixing ratios in the free troposphere remain relatively constant throughout the day (in comparison to the surface). This assumption relies on the fact that the O₃ production rates are expected to be lower at altitude. In the free troposphere, a major part of the NO_x is tied up as peroxyacyl nitrate (PAN) where it is stable at cold temperatures (Singh et al., 1990; Ridley et al., 1990). Besides narrow plumes of lightning NO_x in which mixing ratios can reach several ppbv, NO_x mixing ratios remain low, typically ranging between 10 and 1000 pptv (Emmons et al., 1997; Huntrieser et al., 1998; Brunner et al., 2001; Cooper et al., 2006). Over central-western Europe, the NO_x mixing ratios measured close to thunderstorms in summer were mostly below 150 pptv at 3–8 km and below 400 pptv at 8–11 km (Huntrieser et al., 2002). As highly reactive hydrocarbons also have relatively low mixing ratios in the free and upper troposphere (Singh et al., 2000; Balzani Löv et al., 2008; Helmig et al., 2008), a larger portion of the O₃ production (or consumption depending on the NO_x mixing ratio) relies on the slower oxidation of CO or CH₄. Using the GEOS-Chem model, Zhang et al. (2009) showed O₃ production rates of only about 1 ppbv day^-1 in pollution plumes crossing the North Pacific in the FT. Even above the south-eastern United States where a large amount of lightning NO_x is available in summer, Cooper et al. (2006) reported O₃ production rates of only 3–4 ppbv day^-1 in upper troposphere.

However, several O₃ sources including lightning NO_x and stratospheric intrusions may introduce some diurnal variability in the free and upper troposphere. The lightning NO_x emissions represent only 10% (4%) of the total NO_x emissions at the global scale (at northern midlatitudes) (Schumann and Huntrieser, 2007), but are known to contribute substantially to the formation of O₃ in the FT (DeCaria et al., 2005; Cooper et al., 2006, 2007, 2009). At midlatitudes above continents, a large part of these emissions occur in the middle and upper troposphere (Ott et al., 2010). Based on a lightning detection network of 30 antennas in Germany, Wapler (2013) highlighted a strong diurnal variation of the lightning activity, with a frequency of occurrence peaking at 14–18 UTC (whatever the season). Concerning troposphere-stratosphere exchange, to our knowledge, no previous study has investigated the existence of diurnal variations. However, during the Deep Convective Clouds and Chemistry (DC3) experiment (central United States, summer 2012), Pan et al. (2014) recently reported downward transport of stratospheric O₃-rich air masses in the vicinity of overshooting mesoscale convective systems. The thunderstorm activity is generally higher in the afternoon. Based on observations over Europe from the Spinning Enhanced Visible and Infrared Imager (SEVIRI) instrument aboard the Meteosat satellites, Bedka (2011) showed that over land, the frequency of occurrence of these overshooting cloud tops strongly varies throughout the day, with a maximum around 14–17 UTC. As for the formation of O₃ from lightning NO_x, this source may thus increase the O₃ mixing ratios during the late afternoon.

Based on the frequent profiles of MOZAIC-IAGOS aircraft above Frankfurt since 1994, the diurnal variations of the mean O₃ mixing ratios appear statistically significant regardless of pressure level. However, we demonstrate that these diurnal variations quickly decrease with altitude in the troposphere, and remain small in the middle and upper troposphere. Besides the fact that the photochemically-driven diurnal cycle of O₃ in the BL extends higher in altitude during the summer, we show that these low diurnal variations in the free and upper troposphere are observed regardless of season. Similar results are obtained for the different percentiles (5^th, 25^th, 50^th, 75^th, 95^th) of the O₃ distribution.

Comparing the diurnal cycles between 1994–2003 and 2004–2012, we found an increase of O₃ in most of the troposphere and during most of the day. This change appears stronger close to the surface during the night, which leads to a decrease of the amplitude of the diurnal cycle in the BL. At the seasonal scale, this increase occurs mainly in winter, and to a lesser extent in spring and autumn while O₃ is shown to decrease during the summer, although only in the BL. Such changes are consistent with the shift of the O₃ seasonal cycle toward earlier maxima recently reported at several surface stations over Europe (Parrish et al., 2013) and observed mainly in the lower troposphere around Frankfurt (Petetin et al., 2015).

This result applies to central/western Europe but may be different at other locations, for instance in the inter-tropical convergence zone (ITCZ) where the deep convection is stronger and more frequent. While no other location has the high sampling frequency of Franfkurt, in the near future, more aircraft will join the IAGOS program, which may allow investigations of the diurnal ozone variations in other regions of the world.

No new measurements were made for this review article. All ozone datasets mentioned in the text were obtained from existing databases. The MOZAIC-IAGOS data are available on http://www.iagos.fr. The O₃ observations at surface stations around Frankfurt are available on request from the German Federal Environment Agency (Umweltbundesamt, UBA; contact [email protected]).

© 2016 Petetin et al. This is an open-access article distributed under the terms of the Creative Commons Attribution License, which permits unrestricted use, distribution, and reproduction in any medium, provided the original author and source are credited.

The authors acknowledge the strong support of the European Commission, Airbus, and the Airlines (Lufthansa, Air-France, Austrian, Air Namibia, Cathay Pacific, Iberia and China Airlines so far) who carry the MOZAIC or IAGOS equipment and perform the maintenance since 1994. In its last 10 years of operation, MOZAIC has been funded by INSU-CNRS (France), Météo-France, Université Paul Sabatier (Toulouse, France) and Research Center Jülich (FZJ, Jülich, Germany). IAGOS has been additionally funded by the EU projects IAGOS-DS and IAGOS-ERI. The MOZAIC-IAGOS database is supported by ETHER (CNES and INSU-CNRS). Data are also available via Ether web site www.pole-ether.fr. The authors are also grateful to the Air Monitoring Network of the Federal Environment Agency of Germany for providing the O₃ measurements at the Rannheim, Darmstadt, Spessart and Waldhof monitoring sites, and to the Global Atmosphere Watch (GAW) program for hosting and delivering the Waldhof data.