Investigation of the summer 2018 European ozone air pollution episodes using novel satellite data and modelling

Abstract

In the summer of 2018, Europe experienced an intense heatwave which coincided with several persistent large-scale ozone (${\mathrm{O}}_{\mathrm{3}}$) pollution episodes. Novel satellite data of lower-tropospheric column ${\mathrm{O}}_{\mathrm{3}}$ from the Global Ozone Monitoring Experiment-2 (GOME-2) and Infrared Atmospheric Sounding Interferometer (IASI) on the MetOp satellite showed substantial enhancements in 2018 relative to other years since 2012. Surface observations also showed ozone enhancements across large regions of continental Europe in summer 2018 compared to 2017. Enhancements to surface temperature and the ${\mathrm{O}}_{\mathrm{3}}$ precursor gases carbon monoxide and methanol in 2018 were co-retrieved from MetOp observations by the same scheme. This analysis was supported by the TOMCAT chemistry transport model (CTM) to investigate processes driving the observed ${\mathrm{O}}_{\mathrm{3}}$ enhancements. Through several targeted sensitivity experiments we show that meteorological processes, and emissions to a secondary order, were important for controlling the elevated ${\mathrm{O}}_{\mathrm{3}}$ concentrations at the surface. However, mid-tropospheric ($\sim$ 500 $\mathrm{hPa}$) ${\mathrm{O}}_{\mathrm{3}}$ enhancements were dominated by meteorological processes. We find that contributions from stratospheric ${\mathrm{O}}_{\mathrm{3}}$ intrusions ranged between 15 %–40 %. Analysis of back trajectories indicates that the import of ${\mathrm{O}}_{\mathrm{3}}$-enriched air masses into Europe originated over the North Atlantic, substantially increasing ${\mathrm{O}}_{\mathrm{3}}$ in the 500 $\mathrm{hPa}$ layer during summer 2018.

1 Introduction

Over the past few decades there have been several intense summertime heatwaves over Europe (e.g. 2003 over continental Europe, Scott et al., 2004; 2006 over north-western Europe, Rebetez et al., 2008; and 2010 across eastern Europe and Russia, Matsueda, 2011). With current and future climate change, increasing average global surface temperature is expected to trigger more frequent and intense heatwaves (Lhotka et al., 2017; Guerreiro et al., 2018). The summertime 2018 heatwave across predominantly north-western and central Europe and Scandinavia generated temperature anomalies of approximately 2.0–4.0 $\mathrm{K}$ (Li et al., 2020; Drouard et al., 2020). Dynamically, it was caused by a combination of intense anticyclonic blocking systems, Rossby wave dynamics and the positive phase of the summertime North Atlantic Oscillation (NAO$+$) (Li et al., 2020; Liu et al., 2020; Drouard et al., 2020). Environmentally, the summer 2018 heatwave caused severe drought conditions with decreased precipitation and soil moisture content (Bastos et al., 2020; Dirmeyer et al., 2020) while negatively impacting natural vegetation (e.g. decreased gross primary productivity (Smith et al., 2020; Bastos et al., 2020)). From a human health perspective, the 2018 heatwave caused 863 temperature-related excess deaths in the UK (PHE, 2019).

As well as meteorological and vegetation responses, enhancements in atmospheric pollutants from heatwaves can lead to a degradation in air quality (AQ) across Europe. Blocking systems (anticyclonic conditions) have been shown to increase the level of air pollution such as carbon monoxide (CO; Thomas and Devasthale, 2014), nitrogen dioxide (${\mathrm{NO}}_{\mathrm{2}}$; Pope et al., 2014) and particulate matter (i.e. ${\mathrm{PM}}_{\mathrm{2.5}}$; Graham et al., 2020) to hazardous levels. Pope et al. (2016) focused on the 2006 UK heatwave and detected enhancements in surface ${\mathrm{O}}_{\mathrm{3}}$ through the accumulation of pollutants (i.e. atmospheric blocking) but also the higher temperatures yielding more active atmospheric chemistry (i.e. ozone formation). Papanastasiou et al. (2015) found that Greek heatwave conditions (2001–2010) typically yielded an increase in ${\mathrm{NO}}_{\mathrm{2}}$, ${\mathrm{PM}}_{\mathrm{2.5}}$ and ${\mathrm{O}}_{\mathrm{3}}$ by 14 %–29 %, 25 %–38 % and 12 %, respectively. Rasilla et al. (2019) found that heatwaves in Madrid only moderately increased ${\mathrm{NO}}_{\mathrm{2}}$ and ${\mathrm{O}}_{\mathrm{3}}$ but significantly increased ${\mathrm{PM}}_{\mathrm{10}}$ concentrations. However, they associated this with enhanced long-range transport of African dust and then accumulation under heatwave conditions. García-Herrera et al. (2020) provided a review of the 2003 European heatwave, finding that the Alpine region had substantially elevated surface ozone levels (peaking at 417 $\mathrm{\mu }\mathrm{g}{\mathrm{m}}^{-\mathrm{3}}$ with 68 % of sites from 23 countries reaching concentrations above 180 $\mathrm{\mu }\mathrm{g}{\mathrm{m}}^{-\mathrm{3}}$) when compared with the previous 12 summers. Biogenic volatile organic compound (BVOC) emissions from vegetation are known to increase under drought conditions from temperature stress (e.g. in the 2003 European heatwave; Rennenberg et al., 2006). Churkina et al. (2017) found that heatwave conditions (2006) in Berlin yielded an increase in BVOC emissions which contributed up to 12 % of the surface ozone formation. Heatwaves can also trigger wildfires, which emit primary air pollution and can form secondary gases such as surface ozone on a regional and hemispheric scale (Honrath et al., 2004). Overall, elevated surface ${\mathrm{O}}_{\mathrm{3}}$ is associated with adverse health impacts (Doherty et al., 2017; Heal et al., 2013; Jerrett et al., 2009) with ailments such as asthma, reduced lung function and disease (WHO, 2021). It also has adverse impacts on the natural biosphere (Sitch et al., 2007) and agriculture (Hollaway et al., 2012; van Dingenen et al., 2009), in turn reducing deposition of surface ozone on vegetation. In this study, we use surface and satellite observations of ${\mathrm{O}}_{\mathrm{3}}$, in combination with the well-evaluated TOMCAT global chemical transport model (CTM), to investigate the impact of the summer 2018 heatwave on European AQ and determine the key processes driving observed surface/tropospheric ${\mathrm{O}}_{\mathrm{3}}$ enhancements. We describe the observations and model we have used in Sect. . Sections  and  contain our results and discussion/conclusions, respectively.

2 Observations and model

2.1 Satellite and surface observations

We use satellite observations of lower-tropospheric ${\mathrm{O}}_{\mathrm{3}}$ (i.e. subcolumn ${\mathrm{O}}_{\mathrm{3}}$ (${\mathrm{SCO}}_{\mathrm{3}}$) between the surface and 450 $\mathrm{hPa}$) from the Global Ozone Monitoring Experiment (GOME-2) and the Infrared Atmospheric Sounding Interferometer (IASI) instruments on board ESA's MetOp-A satellite, which was launched in 2006 into a sun-synchronous polar orbit with Equator crossing times of 09:30 LT (day) and 21:30 LT (night). GOME-2 is a nadir-viewing spectrometer with spectral coverage in the ultraviolet–visible (UV–Vis) of 240–790 $\mathrm{nm}$ (Riese et al., 2012) and a ground footprint of 40 $\mathrm{km}$ $\times$ 80 $\mathrm{km}$ in the first part of the mission and 40 $\mathrm{km}$ $\times$ 40 $\mathrm{km}$ from 2013 (once MetOp-B was commissioned). IASI is a Michelson interferometer which observes the infrared spectral range 645 to 2760 ${\mathrm{cm}}^{-\mathrm{1}}$ with a spectral sampling of 0.25 ${\mathrm{cm}}^{-\mathrm{1}}$ (Illingworth et al., 2011). It measures simultaneously in four fields of view (circular at nadir with a diameter of 12 $\mathrm{km}$) which are scanned across track to sample a 2200 $\mathrm{km}$ wide swath (Clerbaux et al., 2009).

For GOME-2, the Rutherford Appleton Laboratory (RAL) scheme uses an optimal estimation algorithm (Rodgers, 2000) to retrieve height-resolved ozone distributions spanning the stratosphere and troposphere (Miles et al., 2015). The scheme applied to GOME-2 has been developed from that used first for GOME-1 on board ERS-2 (Munro et al., 1998; Forster et al., 2007). This is a multi-step scheme in which profile information is first retrieved in the stratosphere by exploiting wavelength-dependent absorption in the ${\mathrm{O}}_{\mathrm{3}}$ Hartley band (270–307 $\mathrm{nm}$) and is then extended into the troposphere by exploiting temperature-dependent spectral structure in the ${\mathrm{O}}_{\mathrm{3}}$ Huggins bands (325–335 $\mathrm{nm}$). For IASI, ${\mathrm{O}}_{\mathrm{3}}$ profiles are retrieved using an extended version of RAL's Infrared Microwave Sounding (IMS) scheme, which is described in Pope et al. (2021), Palmer et al. (2022) and Pimlott et al. (2022). The IMS core scheme was originally developed to retrieve temperature, water vapour and stratospheric ${\mathrm{O}}_{\mathrm{3}}$ profiles along with surface spectral emissivity and cloud jointly from co-located measurements by IASI, the Microwave Humidity Sounder (MHS) and the Advanced Microwave Sounding Unit (AMSU-A) on MetOp (RAL Space, 2015). GOME-2 and IMS ${\mathrm{O}}_{\mathrm{3}}$ data were filtered for a geometric cloud fraction less than 0.2, a solar zenith angle less than 80${}^{\circ }$, a cost function less than 200.0 and a convergence flag equal to 1.0. Examples of the vertical sensitivity to retrieving ozone (i.e. averaging kernels) from GOME-2 and IMS are shown in SM 1 in the Supplement.

We also use surface ${\mathrm{O}}_{\mathrm{3}}$ observations from the European Monitoring and Evaluation Programme (EMEP) network for May–August 2017 and 2018. The EMEP network contains $>$ 100 surface measurement sites measuring information on a range of air pollution types (e.g. ozone, ${\mathrm{NO}}_{\mathrm{2}}$ and ${\mathrm{PM}}_{\mathrm{2.5}}$). EMEP surface data can be used for multiple scientific applications such as trend analysis (Yan et al., 2018) and atmospheric chemistry model evaluation (Schultz et al., 2017; Archibald et al., 2020) and are hosted by the EBAS database infrastructure, developed by the Norwegian Institute for Air Research. In total, we used 125 spatial collocated EMEP sites in both years across Europe. Here, data at individual sites were selected where the corresponding data flag was set to 0.0.

2.2 Modelling and sensitivity experiments

In this study, the TOMCAT CTM (Chipperfield, 2006) is forced by European Centre for Medium-Range Weather Forecasts (ECMWF) ERA-Interim reanalysis meteorology (Dee et al., 2011) and is run at a horizontal resolution of 2.8${}^{\circ }$ $\times$ 2.8${}^{\circ }$. The model has with 31 vertical levels from the surface to 10 $\mathrm{hPa}$ with 5–7 (approximately 10) levels in the boundary layer (mid-troposphere), depending on latitude. The model includes detailed tropospheric chemistry, including 229 gas-phase reactions and 82 advected tracers (Monks et al., 2017), and heterogeneous chemistry driven by size-resolved aerosol from the GLOMAP (Global Model of Aerosol Processes) module (Mann et al., 2010). Anthropogenic emissions used in this study come from MACCity (Granier et al., 2011). The original data set in Granier et al. (2011) derived emissions up to 2010. Therefore, the Representative Concentration Pathway 8.5 (RCP 8.5) was used by Granier et al. (2011) to generate emissions for later years (e.g. 2017 and 2018 as used in this study). Fire emissions are from the Global Fire Assimilation System (GFAS, Kaiser et al., 2012) for 2017 and 2018. Year-specific offline biogenic volatile organic compound (VOC) emissions for acetone, methanol, isoprene and monoterpenes were simulated by the Joint UK Land Environment Simulator (JULES – Pacifico et al., 2011; Best et al., 2011; Clark et al., 2011). All other biogenic VOC emissions are climatological values and provided by the Chemistry-Climate Model Initiative (CCMI) (Morgenstern et al., 2017). The global budgets of the JULES VOC emissions are low in comparison to the climatological CCMI emissions, so they were scaled up on a regional basis while retaining the 2017–2018 step change related to the 2018 summer heatwave. The full details of JULES VOC emissions scaling can be found in SM 4 in the Supplement. Lightning emissions of ${\mathrm{NO}}_x$ are coupled to convection in the model, which is derived from the meteorological reanalyses. Therefore, they vary in space and time according to the seasonality and spatial pattern of convective activity (Stockwell et al.,1999). The model was run for 2017 and 2018 with output at 6-hourly intervals (i.e. 00:00, 06:00, 12:00 and 18:00 LT). Here, each year was run with its respective meteorology and emissions and given the labels Met17_Emis17 (representing 2017) and Met18_Emis18 (representing 2018).

To explore the importance of emission and meteorological processes behind the elevated European summer 2018 tropospheric ${\mathrm{O}}_{\mathrm{3}}$ levels, a 1-year model sensitivity experiment was performed using 2018 meteorology but 2017 emissions (i.e. Met18_Emis17). Therefore, the difference between Met18_Emis17 and Met17_Emis17 highlights the impact of fixed emissions (i.e. 2017 emissions for both years including BVOC emissions), while the Met18_Emis18 minus Met18_Emis17 highlights the impact of fixed meteorology (i.e. 2018 meteorology for both years). These are compared with the control differences for 2018–2017 (Met18_Emis18–Met17_Emis17). From here on in, we refer to the control differences, fixed emission differences and the fixed meteorology differences as CTL_DIFF, FIXED_EMIS_DIFF and FIXED_MET_DIFF, respectively. TOMCAT also includes a stratospheric ${\mathrm{O}}_{\mathrm{3}}$ tracer, a common approach to tag stratospheric ${\mathrm{O}}_{\mathrm{3}}$ (e.g. Roelofs et al., 2003; Akritidis et al., 2019), which can be used to investigate the impact of stratospheric ${\mathrm{O}}_{\mathrm{3}}$ intrusion into the troposphere. The tracer is set equal to the model-calculated ${\mathrm{O}}_{\mathrm{3}}$ in the stratosphere. The only tropospheric source of ${\mathrm{O}}_{\mathrm{3}\mathrm{S}}$ is transport from the stratosphere, while its sinks are via photolysis, reactions with ${\mathrm{HO}}_{\mathrm{2}}$, $\mathrm{OH}$ and ${\mathrm{H}}_{\mathrm{2}}\mathrm{O}$ through $\mathrm{O}{(}^{\mathrm{1}}\mathrm{D})$ produced from ${\mathrm{O}}_{\mathrm{3}\mathrm{S}}$ and surface deposition (Monks et al., 2017). The tracer does not have a fixed lifetime, but the loss rate in the troposphere depends on the modelled local $\mathrm{OH}$, ${\mathrm{HO}}_{\mathrm{2}}$, ${\mathrm{H}}_{\mathrm{2}}\mathrm{O}$ and photolysis. Any ${\mathrm{O}}_{\mathrm{3}}$ that gets into the stratosphere will be labelled as stratospheric before it returns. This was used to investigate the impact of stratospheric ${\mathrm{O}}_{\mathrm{3}}$ intrusion into the troposphere.

TOMCAT has been used in a number of previous studies to investigate air quality and tropospheric composition (e.g. Richards et al., 2013; Emmons et al., 2015; Pope et al., 2018; Pope et al., 2020) whose results give confidence in the model's ability to simulate European tropospheric ${\mathrm{O}}_{\mathrm{3}}$ in this study. Overall, when compared with observations, TOMCAT has good spatial agreement with both GOME-2 and IASI and can reasonably reproduce the 2018 ${\mathrm{SCO}}_{\mathrm{3}}$ enhancement in 2018 versus 2017 (SM 5). The model also has good agreement, both in magnitude and seasonality, with the EMEP-observed surface concentrations (SM 5). TOMCAT surface ozone was also compared with higher-resolution modelling (reanalysis) data from the Copernicus Atmosphere Monitoring Service (CAMS), which showed good spatial agreement between the modelling data sets and in the simulated surface ozone absolute values during the European summer 2018 pollution episode (SM 5).

Figure 1

Surface temperature ($\mathrm{K}$) over Europe for May to August in 2017 (left column), 2018 (centre column) and 2018–2017 difference (right column) retrieved from MetOp-A IASI, MHS and AMSU by the IMS scheme.

[Figure omitted. See PDF]

2.3 ROTRAJ back-trajectories

We use the Reading Offline Trajectory Model (ROTRAJ) to generate air mass back-trajectories (Methven et al., 2003) to assess the import of tropospheric ${\mathrm{O}}_{\mathrm{3}}$ into Europe. ROTRAJ is a Lagrangian atmospheric transport model driven by meteorology from the same ECMWF ERA-Interim reanalyses (horizontal resolution of 1.0125${}^{\circ }$) as used by TOMCAT. Velocity fields at the Lagrangian particle positions are determined by cubic Lagrange interpolation in the vertical, bilinear interpolation in the horizontal and linear interpolation in time. This method accounts for large-scale advection since the winds are resolved, but it does not resolve small-scale subgrid turbulent transport. Kinematic back-trajectories were released at 6-hourly intervals (i.e. at 00:00, 06:00, 12:00 and 18:00 LT) from Paris and Berlin, both central locations over Europe in the region of summertime 2018 ${\mathrm{O}}_{\mathrm{3}}$ enhancements, between 1 May and 31 August for both 2017 and 2018. The trajectories were released at the surface and at approximately 500 $\mathrm{hPa}$ and integrated for 10 $\mathrm{d}$ with 6-hourly output (i.e. 41 trajectory points including the starting location) to investigate the origin of air masses arriving in these altitude regions of enhanced summertime ${\mathrm{O}}_{\mathrm{3}}$ in 2018. In total, ROTRAJ was therefore run eight times (2 years $\times$ 2 altitudes $\times$ 2 locations).

To quantify the import of tropospheric ${\mathrm{O}}_{\mathrm{3}}$ into Europe, for each trajectory, all the trajectory points were co-located with corresponding TOMCAT ${\mathrm{O}}_{\mathrm{3}}$ mixing ratio values (i.e. the horizontal and vertical grid box the trajectory point sits within and corresponding time stamp) and then the average ${\mathrm{O}}_{\mathrm{3}}$-weighted back-trajectory (${\mathrm{O}}_{\mathrm{3}}$-WBT) determined (i.e. back-trajectories with larger ${\mathrm{O}}_{\mathrm{3}}$-WBT values come from air masses with larger ${\mathrm{O}}_{\mathrm{3}}$ content). This follows a similar approach to Graham et al. (2020) and Stirling et al. (2020), though using a model chemical tracer and not emission inventories.

3 Results

3.1 Surface temperature

Several studies (e.g. Li et al., 2020; Liu et al., 2020; Drouard et al., 2020) have documented the intense heatwave across Europe in the summer of 2018. This is further shown in Fig. 1, which compares surface temperature, co-retrieved with ozone and other variables from MetOp-A by the IMS scheme, between 2017 and 2018. In May, higher temperatures occur across Scandinavia (5.0–10.0 $\mathrm{K}$), eastern Europe (3.0–7.0 $\mathrm{K}$) and the UK (1.0–3.0 $\mathrm{K}$), but temperatures are lower ($-$3.0 to $-$1.0 $\mathrm{K}$) across the Iberian Peninsula. In June, a similar spatial distribution occurs, but the magnitude of the differences is smaller. In July, the largest temperature increases range from 6.0–8.0 $\mathrm{K}$ in Scandinavia to 2.0–6.0 $\mathrm{K}$ in the UK and France. The Iberian Peninsula continued to experience temperatures lower by $-$2.0 to 0.0 $\mathrm{K}$. In August, there are near-zero differences over the UK, the Iberian Peninsula and most of Scandinavia but with increases of 1.0–3.0 $\mathrm{K}$ over eastern Europe and Finland.

Figure 2

Subcolumn ozone (${\mathrm{SCO}}_{\mathrm{3}}$, surface – 450 $\mathrm{hPa}$), in Dobson units ($\mathrm{DU}$), retrieved from GOME-2 on MetOp-A averaged across May to August between 2012 and 2018 (top left panel) and the corresponding difference from the 2012–2018 mean for each year, respectively. The green-polygon-outlined regions show where the year-specific seasonal average is significantly different (95 % confidence level based on the Wilcoxon rank test, WRT) from the long-term (2012–2018) seasonal average. The “Sig Pixel %” label indicates the number of pixels in the domain with significant differences.

[Figure omitted. See PDF]

Figure 3

${\mathrm{SCO}}_{\mathrm{3}}$ ($\mathrm{DU}$) from GOME-2 over Europe for May to August in 2017 (left column), 2018 (centre column) and 2018–2017 difference (right column).

[Figure omitted. See PDF]

3.2 Satellite ozone

We investigate the longer-term variability in tropospheric ${\mathrm{O}}_{\mathrm{3}}$ (i.e. ${\mathrm{SCO}}_{\mathrm{3}}$) to determine whether 2017 is a suitable comparator for the 2018 summer ${\mathrm{O}}_{\mathrm{3}}$ enhancements as it is for temperature. Figure 2 shows the 2012–2018 ${\mathrm{SCO}}_{\mathrm{3}}$ average between May and August for a domain over the Atlantic and Europe and the difference for the same season between specific years and the 2012–2018 average. In 2012 and 2013, there are significant positive differences from the average between 1.0 and 5.0 $\mathrm{DU}$ over much of the domain. Over continental Europe, the differences are smaller. Here, the significance of differences between the year-specific and long-term averages are determined using the Wilcoxon rank test (Pirovano et al., 2012) at the 95 % confidence level. In 2014 and 2015, there are negative differences across Europe ($-$4.0 to $-$1.0 $\mathrm{DU}$). In 2016, similar negative differences are primarily across the north and south-east of the domain. In 2017, there are near-zero differences across the Atlantic, UK and western Europe. Over eastern Europe and Mediterranean, there are significant negative differences of between $-$2.0 and $-$1.0 $\mathrm{DU}$. In 2018, across continental Europe there are significant positive differences between 2.0 and 4.0 $\mathrm{DU}$. As the 2017 differences are relatively small in magnitude with a low proportion of significant pixels (i.e. Sig Pixels % $=$ 32.7 is the lowest across the 7 years), it is representative of average conditions for comparison with 2018. For 2018, the summer ${\mathrm{SCO}}_{\mathrm{3}}$ enhancements across continental Europe are the largest for the years shown with a coherent cluster of significant differences. This illustrates that the summer 2018 ${\mathrm{SCO}}_{\mathrm{3}}$ enhancements are a substantial deviation from the average conditions (which we represent as 2017 herein) and that this is an intense ${\mathrm{O}}_{\mathrm{3}}$ event.

Figure 4

${\mathrm{SCO}}_{\mathrm{3}}$ ($\mathrm{DU}$) for May to August in 2017 (left column), 2018 (centre column) and 2018–2017 difference (right column) over Europe retrieved from MetOp-A IASI, MHS and AMSU by the IMS scheme.

[Figure omitted. See PDF]

Investigation of ${\mathrm{SCO}}_{\mathrm{3}}$ retrieved from both GOME-2 (Fig. 3) and the IMS scheme (Fig. 4) shows consistent enhancements in summer 2018. In 2017, between May and August, GOME-2 typically observed ${\mathrm{SCO}}_{\mathrm{3}}$ values between 20.0–30.0 $\mathrm{DU}$ across continental Europe. Peak ${\mathrm{SCO}}_{\mathrm{3}}$ values occurred over the Mediterranean (30.0–38.0 $\mathrm{DU}$); relatively high ozone is a typical feature of the Mediterranean in summer (Richards et al., 2013). In 2018, the seasonality is consistent with 2017, but the continental European ${\mathrm{SCO}}_{\mathrm{3}}$ values ranged between 25.0 and 35.0 $\mathrm{DU}$. For the 2018–2017 difference, ${\mathrm{SCO}}_{\mathrm{3}}$ enhancements occur across continental Europe in all 4 months but peaked in May and July between 3.0 and 8.0 $\mathrm{DU}$ while typically 1.0–5.0 $\mathrm{DU}$ in June and August. The spatial distribution of IMS-retrieved ${\mathrm{SCO}}_{\mathrm{3}}$ is similar to that of GOME-2 in 2017 and 2018, although the absolute values tend to be systematically lower by 3.0–4.0 $\mathrm{DU}$. However, despite this systematic offset, the 2018–2017 differences are reasonably consistent with GOME-2. Across continental Europe, IMS ${\mathrm{SCO}}_{\mathrm{3}}$ shows 2018 enhancements in all months investigated but peaks in May and July, like GOME-2, between 3.0 and 6.0 $\mathrm{DU}$. The differences range from 1.0 to 3.0 $\mathrm{DU}$ in June and are approximately 1.0 $\mathrm{DU}$ in August (though a peak enhancement of 3.0–5.0 $\mathrm{DU}$ occurs over the Mediterranean). Spatial correlations between the GOME-2 and IASI difference (i.e. 2018–2017) maps for the months investigated ranged between 0.21 and 0.47 (see SM 5 in the Supplement).

The GOME-2 and IASI instruments observe UV–Vis and IR wavelengths, with peak vertical sensitivities to tropospheric ${\mathrm{O}}_{\mathrm{3}}$ in the lower and mid/upper troposphere, respectively. Consistency in the 2018 enhancements in ${\mathrm{SCO}}_{\mathrm{3}}$ indicates that these extend over the bulk of the troposphere and increases confidence in the detected enhancements for both sensors.

Investigation of several satellite-retrieved ${\mathrm{O}}_{\mathrm{3}}$ precursor gases (see SM 2) showed enhancements in total column methanol (${\mathrm{TCCH}}_{\mathrm{3}}\mathrm{OH}$, Fig. S2), especially linked to May and July temperature enhancements (Fig. 1), minor increases in tropospheric column ${\mathrm{NO}}_{\mathrm{2}}$ (${\mathrm{TCNO}}_{\mathrm{2}}$, Fig. S3) in May and July over central Europe, and widespread enhancements (weakest in July and strongest in August) in total column carbon monoxide ($\mathrm{TCCO}$, Fig. S4). Investigation of the GOME-2 and IASI total column ${\mathrm{O}}_{\mathrm{3}}$ (${\mathrm{TCO}}_{\mathrm{3}}$) differences between 2017 and 2018 (Figs. S5 and S6) showed these to be in close agreement. Some spatial structure is similar to that of the ${\mathrm{SCO}}_{\mathrm{3}}$ difference patterns (Figs. 3 and 4), with correlations of approximately 0.5 between ${\mathrm{TCO}}_{\mathrm{3}}$ and ${\mathrm{SCO}}_{\mathrm{3}}$ for each instrument (see SM 3). Given the complex relationship between tropospheric ${\mathrm{O}}_{\mathrm{3}}$, precursor gases, atmospheric chemistry (e.g. ${\mathrm{NO}}_x$ or VOC-limited regimes), surface deposition and meteorological conditions (e.g. atmospheric temperatures and transport), a detailed chemistry transport model is required to assess the key processes leading to the observed ${\mathrm{SCO}}_{\mathrm{3}}$ enhancements over Europe.

Figure 5

European surface ozone ($\mathrm{ppbv}$) for (a) May–June–July–August (MJJA) 2017, (b) MJJA 2018, (c) regional mean time series (dotted lines show mean $\pm$ standard deviation) for MJJA 2017 (blue), MJJA 2018 (red) and the 2018–2017 difference (black) and (d) box-and-whisker plots for MJJA 2017 and 2018. In panel (d) the median, 25th and 75th percentiles and 10th and 90th percentiles are shown by the red, green and blue lines, respectively.

[Figure omitted. See PDF]

3.3 Surface ozone

Increased temperatures during heatwaves have been shown to enhance surface ${\mathrm{O}}_{\mathrm{3}}$ concentrations (e.g. Jacob and Winner, 2009; Vieno et al., 2010; Pyrgou et al., 2018). In the summer (May–June–July–August, MJJA) of 2018, EMEP recorded larger ${\mathrm{O}}_{\mathrm{3}}$ mixing ratios across most of Europe in comparison to 2017 (Fig. 5a and b). Over central Europe, surface ${\mathrm{O}}_{\mathrm{3}}$ mixing ratios ranged from approximately 45.0 $\mathrm{ppbv}$ to over 60.0 $\mathrm{ppbv}$, while in 2017 it was 35.0 to 50.0 $\mathrm{ppbv}$. Over the UK and north-western Europe, surface ${\mathrm{O}}_{\mathrm{3}}$ mixing ratios ranged from 20.0 to 30.0 $\mathrm{ppbv}$ and then 25.0 to 35.0 $\mathrm{ppbv}$ in MJJA 2017 and 2018, respectively. In Scandinavia and eastern Europe, surface ${\mathrm{O}}_{\mathrm{3}}$ mixing ratios ranged from 20.0 to 35.0 $\mathrm{ppbv}$ in MJJA 2017 while increasing to 25.0 $\mathrm{ppbv}$ to approximately 40.0 $\mathrm{ppbv}$ in MJJA 2018. Figure 5c highlights these widespread enhancements where domain-average surface ${\mathrm{O}}_{\mathrm{3}}$ mixing ratios are larger by typically 5.0–10.0 $\mathrm{ppbv}$ in May and from mid-June to mid-August in 2018. Figure 5d shows that the domain median surface ${\mathrm{O}}_{\mathrm{3}}$ concentration across MJJA was larger by 2.0–3.0 $\mathrm{ppbv}$ in 2018, but the 2018 extremes were greater with 75th and 95th percentiles of 45.0 and 55.0 $\mathrm{ppbv}$ in 2017 and 48.0 and 59.0 $\mathrm{ppbv}$ in 2018. Therefore, surface observations of ${\mathrm{O}}_{\mathrm{3}}$ recorded widespread enhancements in MJJA 2018 compared to 2017 with peak site differences $>$ 10.0 $\mathrm{ppbv}$. This is generally consistent with the 2018 layer-averaged enhancements in the satellite-retrieved ${\mathrm{SCO}}_{\mathrm{3}}$ for regions where both data sets have spatial coverage.

3.4 Model simulations

We use the TOMCAT model to investigate different factors potentially driving the observed enhancements in tropospheric ${\mathrm{O}}_{\mathrm{3}}$. In comparison with the observations (see SM 5 in the Supplement), the model reproduces the sign and spatial distribution of observed 2018–2017 differences reasonably well. Although it has a tendency to underestimate the absolute magnitude, we are confident in the model's ability to simulate the tropospheric ${\mathrm{O}}_{\mathrm{3}}$ enhancements relative to 2017.

Figure 6

TOMCAT ozone ($\mathrm{ppbv}$) 2018–2017 differences for May to August for the surface (left column), 500 $\mathrm{hPa}$ (centre column) and the stratospheric contribution (%) to the 500 $\mathrm{hPa}$ layer (right column).

[Figure omitted. See PDF]

At the surface (Fig. 6), TOMCAT CTL_DIFF (i.e. Met18_Emis18–Met17_Emis17) suggests that ${\mathrm{O}}_{\mathrm{3}}$ is enhanced in May over Scandinavia (2.0 to $>$ 5.0 $\mathrm{ppbv}$), north-western Europe (0.0–2.0 $\mathrm{ppbv}$), the Arctic Ocean ($>$ 5.0 $\mathrm{ppbv}$) and off the coast of the Iberian Peninsula (3.0–5.0 $\mathrm{ppbv}$). However, negative values exist over eastern Europe ($-$3.0 $\mathrm{ppbv}$ to $-$1.0 $\mathrm{ppb}$) and the Atlantic west of Ireland ($-$3.0 $\mathrm{ppbv}$ to $-$1.0 $\mathrm{ppb}$). In June, the negative differences persist in eastern Europe ($-$3.0 $\mathrm{ppbv}$ to $-$1.0 $\mathrm{ppb}$), but positive differences are located over northern Scandinavia (1.0–2.0 $\mathrm{ppbv}$) and the North Atlantic (2.0–4.0 $\mathrm{ppbv}$). For July, CTL_DIFF shows the largest enhancements over continental Europe (i.e. Po Valley, France, Benelux region and Iberian Peninsula) and the UK ($>$ 5.0 $\mathrm{ppbv}$). Negative differences of between $-$3.0 and $-$1.0 $\mathrm{ppbv}$ remain over eastern Europe. In August, the only clear differences are over the Iberian Peninsula and the western Mediterranean, ranging between 3.0 and $>$ 5.0 $\mathrm{ppbv}$. Overall, TOMCAT simulates subregional surface ${\mathrm{O}}_{\mathrm{3}}$ enhancements over Europe, which are generally consistent with EMEP observations apart from several sites over eastern Europe.

At 500 $\mathrm{hPa}$, TOMCAT CTL_DIFF shows larger-scale ${\mathrm{O}}_{\mathrm{3}}$ enhancements in 2018 compared to 2017 ($>$ 5.0 $\mathrm{ppbv}$) throughout May to August. In May and August, there are, however, a few negative differences ($-$5.0 to $-$3.0 $\mathrm{ppbv}$) over far eastern Europe. In June and July, the full domain is more or less dominated by ${\mathrm{O}}_{\mathrm{3}}$ enhancements in 2018. In Figs. 3 and 4 (and SM 5 in the Supplement), GOME-2 and IASI (and TOMCAT with the instrument averaging kernels (AKs) applied to account for the vertical sensitivity of the retrievals; see SM 5 in the Supplement for more information) show ${\mathrm{SCO}}_{\mathrm{3}}$ enhancements during these months of 2018. Given the vertical extents and peak heights of their retrieval sensitivities and consistency in spatial patterns (Figs. S9 and S11), it is evident that the ${\mathrm{O}}_{\mathrm{3}}$ enhancements detected by GOME-2 and IASI extend over the free troposphere. The model shows large-scale ${\mathrm{O}}_{\mathrm{3}}$ enhancements in the free troposphere and similar patterns to GOME-2 and IASI when averaging kernels applied. So, the model corroborates this finding from the satellite retrievals. Signals from EMEP and TOMCAT at the surface, on the other hand, are more mixed across the domain.

The right-hand column of Fig. 6 shows the relative difference in the stratospheric ${\mathrm{O}}_{\mathrm{3}}$ contribution to the 500 $\mathrm{hPa}$ ${\mathrm{O}}_{\mathrm{3}}$ layer (i.e. Strat % at 500 $\mathrm{hPa}$), from CTL_DIFF, between 2017 and 2018. Here, the percentage of stratospheric ${\mathrm{O}}_{\mathrm{3}}$ contributing to the ${\mathrm{O}}_{\mathrm{3}}$ concentration at the 500 $\mathrm{hPa}$ is calculated for 2017 and 2018 and then the 2018–2017 difference determined. The largest enhancement to the 500 $\mathrm{hPa}$ layer was in July where the stratospheric ${\mathrm{O}}_{\mathrm{3}}$ contribution increased by 3.0 % to $>$ 5.0 % across Europe. In June and August, the spatial patterns are similar with stratospheric ${\mathrm{O}}_{\mathrm{3}}$ contribution enhancements of 3.0 %–5.0 % across southern Europe, Scandinavia and the North Atlantic (above the UK). In the North Atlantic, UK and northern Europe, there are near-zero changes in June and August. In May, there are enhancements $>$ 5.0 % across the northern region of the domain and northern Africa, while there are smaller enhancements (1.0 %–3.0 %) over the UK and near-zero changes over eastern Europe. This is partially supported by analysis of ${\mathrm{TCO}}_{\mathrm{3}}$ (see SM 3 in the Supplement) where there are reasonable spatial correlations ($\sim$ 0.5 to 0.6) between the ${\mathrm{SCO}}_{\mathrm{3}}$ 2017–2018 summer differences and the equivalent for ${\mathrm{TCO}}_{\mathrm{3}}$. Therefore, these results indicate a potentially enhanced contribution of stratospheric ${\mathrm{O}}_{\mathrm{3}}$ into the mid-troposphere during the summer of 2018 across Europe.

Figure 7

TOMCAT ozone ($\mathrm{ppbv}$) 2018–2017 differences for May to August for the fixed emissions simulation (Fixed_EMIS) for the surface (left column), 500 $\mathrm{hPa}$ (centre column) and the stratospheric contribution (%) to the 500 $\mathrm{hPa}$ layer (right column).

[Figure omitted. See PDF]

Figure 8

TOMCAT ozone ($\mathrm{ppbv}$) 2018–2017 differences for May to August for the fixed meteorology simulation (Fixed_MET) for the surface (left column), 500 $\mathrm{hPa}$ (centre column) and the stratospheric contribution (%) to the 500 $\mathrm{hPa}$ layer (right column).

[Figure omitted. See PDF]

To quantify the separate importance of precursor emissions and meteorology in governing the summer 2018 ${\mathrm{O}}_{\mathrm{3}}$ enhancements, we compare the sensitivity experiments with the control runs. Figure 7 (left column) shows the results for the fixed emissions differences (i.e. FIXED_EMIS_DIFF) between years (i.e. Met18_Emis17 – Met17_Emis17). At the surface, the FIXED_EMIS_DIFF shows similar spatial patterns to that of CTL_DIFF (Fig. 6 – left column). The domain spatial difference correlations between these simulations are greater than 0.96 for all months considered. However, FIXED_EMIS_DIFF (Fig. 7 – left column) tends to be lower than CTL_DIFF (Fig. 6 – left column) by approximately 0.0–2.9 $\mathrm{ppbv}$ (i.e. positive red regions are weaker and negative blue regions stronger in intensity). Therefore, the Met18_Emis17 run struggles to reproduce the absolute surface ${\mathrm{O}}_{\mathrm{3}}$ enhancements in the Met18_Emis18 run. When the fixed meteorology differences (FIXED_MET_DIFF, i.e. Met18_Emis18 – Met18_Emis17; Fig. 8 – left column) are compared with CTL_DIFF, the surface 2018–2017 differences are substantially different.

Surface FIXED_MET_DIFF ranges between 0.0 and 2.0 $\mathrm{ppbv}$ across the domain in May and June and is more confined to continental Europe in July and August. This shows that TOMCAT simulates lower 2018 summertime ${\mathrm{O}}_{\mathrm{3}}$ when 2017 emissions are used and indicates that emissions do have some role in controlling ${\mathrm{O}}_{\mathrm{3}}$ levels at the surface. However, as the spatial difference pattern for FIXED_MET_DIFF (Fig. 8 – left column) is different to that of CTL_DIFF (Fig. 6 – left column), spatial correlations between them range from $-$0.53 to 0.54 over the 4 months, which suggests that meteorology is important in governing the spatial distribution of CTL_DIFF. This is supported by the fact that FIXED_MET_DIFF–CTL_DIFF (Fig. 8 left column – Fig. 6 left column) yields absolute domain variations between 0.0 and 12.2 $\mathrm{ppbv}$. Therefore, the two sensitivity experiments suggest meteorology and emissions both play important roles in controlling surface ${\mathrm{O}}_{\mathrm{3}}$ during the summer of 2018, but meteorology predominantly governs the spatial pattern and absolute magnitude of the ${\mathrm{O}}_{\mathrm{3}}$ enhancements.

Figure 9

TOMCAT ozone, zonally averaged between 20${}^{\circ }$ W and 40${}^{\circ }$ E, 2018–2017 percentage differences (absolute difference ($\mathrm{ppbv}$) shown as solid lines) from the control simulation. Panels (a–d) represent the monthly averages for May, June, July and August.

[Figure omitted. See PDF]

At 500 $\mathrm{hPa}$, comparison of FIXED_EMIS_DIFF and CTL_DIFF shows very consistent spatial patterns across the 4 months with correlations all above 0.98. In terms of the absolute differences between FIXED_EMIS_DIFF and CTL_DIFF (i.e. Fig. 7 centre column – Fig. 6 centre column), it peaks at approximately 2.8 $\mathrm{ppbv}$. For FIXED_MET_DIFF, the spatial correlation with CTL_DIFF, as for the surface, is variable with values between $-$0.38 and 0.43. The absolute differences between FIXED_MET_DIFF and CTL_DIFF (i.e. Fig. 8 centre column – Fig. 6 centre column) range from 0.0 to 14.8 $\mathrm{ppbv}$. Therefore, emissions have a secondary role in controlling the ${\mathrm{O}}_{\mathrm{3}}$, while meteorology is by far the dominant factor. For Strat % at 500 $\mathrm{hPa}$, the spatial correlations between CTL_DIFF and FIXED_EMIS_DIFF are above 0.95 for all months, and the absolute differences between them (i.e. Fig. 7 right column – Fig. 6 right column) are near-zero. Comparison of FIXED_MET_DIFF and TC_CTL shows spatial difference correlations ranging between $-$0.33 and 0.71 and absolute differences (i.e. Fig. 8 right column – Fig. 6 right column) peaking at 12.9 %. Therefore, as expected, meteorological processes are dominating the influence of the stratospheric ${\mathrm{O}}_{\mathrm{3}}$ contribution (i.e. through stratosphere–troposphere exchanges) to the 500 $\mathrm{hPa}$ layer during the summer 2018 ${\mathrm{O}}_{\mathrm{3}}$ enhancements over Europe.

Figure 10

TOMCAT stratospheric ozone tracer, zonally averaged between 20${}^{\circ }$ W and 40${}^{\circ }$ E, 2018–2017 percentage differences (absolute difference ($\mathrm{ppbv}$) shown as solid lines) from the control simulation. Panels (a–d) represent the monthly averages for May, June, July and August.

[Figure omitted. See PDF]

To investigate the importance of stratospheric–troposphere exchange to the middle troposphere enhancement (i.e. as shown in the TOMCAT 500 $\mathrm{hPa}$ layer and the satellite ${\mathrm{SCO}}_{\mathrm{3}}$ data), Figs. 9 and 10 show TOMCAT control run zonal 2018–2017 difference cross-sections (for the domain longitudes) of ${\mathrm{O}}_{\mathrm{3}}$ profiles and the stratospheric ${\mathrm{O}}_{\mathrm{3}}$ contribution to each pressure layer. In May and June, in the lower troposphere (approximately surface to 800 $\mathrm{hPa}$), there are negative ($-$3.0 % to 0.0 %) and positive (0.0 % to 3.0 %) differences between 30–50${}^{\circ }$ N and 50–70${}^{\circ }$ N, respectively. During June, there are positive differences (0.0 % to 5.0 %) across most latitudes, and in August the opposite occurs to that of May/June. In the mid-troposphere (800–300 $\mathrm{hPa}$), positive differences occur in most months (0.0 %–5.0 % in May, 0.0 %–7.0 % in June, $>$ 10 % in July and 5.0 %–10.0 % in August), though in May and August negative differences ($-$5.0 % to 0.0 %) exist around 40 and 55${}^{\circ }$ N. This is consistent with the 500 $\mathrm{hPa}$ ${\mathrm{O}}_{\mathrm{3}}$ differences in Fig. 6 (centre panels). In the upper troposphere–lower stratosphere (UTLS, approximately 300–100 $\mathrm{hPa}$), there are limbs of positive ${\mathrm{O}}_{\mathrm{3}}$ differences (i.e. $>$ 10 %, 5.0–10.0 $\mathrm{ppbv}$) propagating into the mid-troposphere (30–40${}^{\circ }$ N in May, 30–50${}^{\circ }$ N in June, 40–50${}^{\circ }$ N in July and 30–40${}^{\circ }$ N and 60–70${}^{\circ }$ N in August), suggestive of stratospheric intrusion into the mid-troposphere. Using the stratospheric ${\mathrm{O}}_{\mathrm{3}}$ tracer in TOMCAT, Fig. 10 shows the enhanced proportion of ${\mathrm{O}}_{\mathrm{3}}$ originating from the stratosphere in the summer of 2018. Interestingly, for all months (apart from May between 30–45${}^{\circ }$ N), there are enhanced contributions of stratospheric ${\mathrm{O}}_{\mathrm{3}}$ (15.0 % to $>$ 50.0 %) in the lower–mid-troposphere (i.e. below 500 $\mathrm{hPa}$). In absolute terms, this is only a minor contribution typically $<$ 1.0 $\mathrm{ppbv}$ below 800 $\mathrm{hPa}$. Between 800–400 $\mathrm{hPa}$, this increases to 1.0–5.0 $\mathrm{ppbv}$ (remains relatively consistent in percentage terms) in most months and latitude bands. In the UTLS, it increases to 5.0 %–10.0 % enhancements in stratospheric ${\mathrm{O}}_{\mathrm{3}}$ contributions, which is consistent with its proximity to the stratosphere. Where there are enhancements in the stratospheric ${\mathrm{O}}_{\mathrm{3}}$ contribution but negative differences in ${\mathrm{O}}_{\mathrm{3}}$ (e.g. in July in the lower troposphere between 50 and 55${}^{\circ }$ N – Figs. 9 and 10), this is indicative of competing processes influencing the ${\mathrm{O}}_{\mathrm{3}}$ concentrations (e.g. descent of relatively small stratospheric ${\mathrm{O}}_{\mathrm{3}}$ contributions but advection of tropospheric ${\mathrm{O}}_{\mathrm{3}}$ away from the region). Overall though, in the mid-troposphere, where there are larger enhancements in ${\mathrm{O}}_{\mathrm{3}}$, there are similar responses in the stratospheric ${\mathrm{O}}_{\mathrm{3}}$ contribution. For June, the mid-troposphere ${\mathrm{O}}_{\mathrm{3}}$ enhancement is approximately 5.0–7.0 $\mathrm{ppbv}$ with a signal of 1.0–2.0 $\mathrm{ppbv}$ in the stratospheric tracer. Therefore, in the more extreme cases, the stratospheric ${\mathrm{O}}_{\mathrm{3}}$ contribution is approximately 15.0 %–40.0 % to the mid-tropospheric ${\mathrm{O}}_{\mathrm{3}}$ enhancements in summer 2018 over Europe. However, a separate study would be required to undertake a detailed assessment of the meteorological processes controlling the enhanced stratospheric intrusion of ozone in the summer of 2018 and how it compares to other years (how does it compare with years other than 2017?).

Figure 11

TOMCAT temperature, zonally averaged between 20${}^{\circ }$ W and 40${}^{\circ }$ E, 2018–2017 percentage differences (absolute difference ($\mathrm{K}$) shown by black solid, dotted and dashed lines) from the control simulation. Overplotted are contours of the temporal correlation (i.e. within each grid box) between the temperature and ozone 2018–2017 differences. Panels (a–d) represent the monthly averages for May, June, July and August.

[Figure omitted. See PDF]

The two remaining factors, linked to meteorological processes (as suggested above), which may affect the ${\mathrm{O}}_{\mathrm{3}}$ enhancements in 2018, are increased summer temperatures (e.g. through enhanced kinetic rates) and the import of tropospheric ${\mathrm{O}}_{\mathrm{3}}$ from upwind (e.g. North America from the prevailing winds). Figure 11 shows the 2017–2018 zonal temperature differences (i.e. same as Fig. 9 but for temperature) with the correlation between the 2017 and 2018 temperature and ${\mathrm{O}}_{\mathrm{3}}$ differences overplotted. Qualitatively, the zonal differences in ${\mathrm{O}}_{\mathrm{3}}$ and temperature have some similarities. There are positive differences (temperature differences of 0.0 %–1.0 %) between 50–60${}^{\circ }$ N at the surface and 400 $\mathrm{hPa}$ in May and June. Then in July, collocated positive differences (peaking at 2.0 % or 3.0 $\mathrm{K}$) exist between 50–70${}^{\circ }$ N from the surface to 300 $\mathrm{hPa}$. In August, there is no clear relationship between temperature and ${\mathrm{O}}_{\mathrm{3}}$ enhancements. In all months (to a lesser extent in August), in the UTLS, there are spatial agreements with positive differences between approximately 30–45${}^{\circ }$ N and negative differences between 50/55–70${}^{\circ }$ N. In terms of correlations (i.e. temporal correlation in each grid box using the TOMCAT 6-hourly time series), the spatial agreement is relatively weak. In all months, most of domain has relatively small values ranging between $-$0.5 and 0.5. There are only a few locations with strong correlations (i.e. $>$ 0.5), which are in the UTLS or in the lower–mid-troposphere between 50–70${}^{\circ }$ N (June and August) and 45–55${}^{\circ }$ N in July near the surface. Overall, the relationship between increased temperatures and enhanced kinetic rates yielding more ozone formation is non-linear, so it is unsurprising that the direct comparisons of temperature and ozone 2018–2017 differences above show no clear pattern. Therefore, future work could include a further sensitivity experiment running TOMCAT for 2018, but with 2017 temperatures used in the chemistry routines to quantify the role of temperature in the summer 2018 ${\mathrm{O}}_{\mathrm{3}}$ enhancements.

Figure 12

The difference between May–August 2018 and May–August 2017 (i.e. 2018–2017) ROTRAJ back-trajectories (10 $\mathrm{d}$), weighted by the average TOMCAT ${\mathrm{O}}_{\mathrm{3}}$ ($\mathrm{ppbv}$) concentration along each trajectory path, gridded onto the TOMCAT horizontal resolution for (a) Paris at the surface, (b) Berlin at the surface, (c) Paris at approximately 500 $\mathrm{hPa}$ and (d) Berlin at approximately 500 $\mathrm{hPa}$. The black circles represent the location of Paris or Berlin, where the trajectories were released from.

[Figure omitted. See PDF]

To investigate the potential advection of tropospheric ${\mathrm{O}}_{\mathrm{3}}$-rich air masses into Europe, we have used ROTRAJ back-trajectories to determine the ${\mathrm{O}}_{\mathrm{3}}$-WBTs (i.e. an indicator of air mass ${\mathrm{O}}_{\mathrm{3}}$ content). As shown in SM 6, there is large variability in the ${\mathrm{O}}_{\mathrm{3}}$-WBT values and spatial distribution (i.e. Figs. S13 and S14), so they have to be gridded onto the TOMCAT horizontal resolution (see Figs. S15 and S16). While this approach does not directly account for the frequency of trajectory points in each grid box, Figs. S13 and S14 show there is widespread coverage across the North Atlantic. This results in $>$ 500 trajectory points near the receptor sites (i.e. Paris and Berlin), $\sim$ 100 trajectory points around the edge of Europe and 25–50 trajectory points in the North Atlantic (not shown here). Overall, this spatial distribution is relatively consistent and does not change substantially between years (typically 10 %); thus this approach is suitable in this study. Figure 12 shows the differences (2018–2017) between the gridded ${\mathrm{O}}_{\mathrm{3}}$-WBTs where the back-trajectories have been released at the surface from Paris (Fig. 12a), at the surface from Berlin (Fig. 12b), at approximately 500 $\mathrm{hPa}$ from Paris (Fig. 12c) and at approximately 500 $\mathrm{hPa}$ from Berlin (Fig. 12d). We selected Paris and Berlin as they are situated in central Europe, where the summer 2018 ${\mathrm{O}}_{\mathrm{3}}$ enhancements were observed, while the surface and 500 $\mathrm{hPa}$ are the altitudes of primary focus in the modelling work.

At the surface, Paris and Berlin show consistent patterns. Over the North Atlantic (i.e. origin of the prevailing winds into Europe), there are typically negative ${\mathrm{O}}_{\mathrm{3}}$-WBT values between $-$5.0 and $-$1.0 $\mathrm{ppbv}$ suggesting that advection of ${\mathrm{O}}_{\mathrm{3}}$ into Europe during the summer (i.e. May–August) was predominantly larger in 2017 and did not strongly contribute to the 2018 observed surface ${\mathrm{O}}_{\mathrm{3}}$ enhancements. Advection of ${\mathrm{O}}_{\mathrm{3}}$-rich air in 2018 did originate from Scandinavia into continental Europe, though the number of trajectories is relatively low (see Fig. S13). As both locations show similar relationships, it provides confidence in this methodology. At 500 $\mathrm{hPa}$, the 50–60${}^{\circ }$ N spatial pattern is less defined with values typically between $-$5.0 and 5.0 $\mathrm{ppbv}$ for both locations. However, in the southern North Atlantic (30–50${}^{\circ }$ N) there are positive differences of approximately 3.0–10.0 $\mathrm{ppbv}$ for both release locations. Note that as free-tropospheric winds tend to have larger horizontal velocities, the back-trajectories generally start from further away, closer to North America. Again, given the broad similarity in differences between both release locations, it provides confidence in this approach. Overall, our results indicate a larger transport of ${\mathrm{O}}_{\mathrm{3}}$ to the surface of continental Europe in 2017, while at approximately 500 $\mathrm{hPa}$ the import of ${\mathrm{O}}_{\mathrm{3}}$ into Europe is larger in 2018. Here, the positive differences originate from the southern North Atlantic (i.e. a larger range of locations, absolute values and homogeneous signal than the mixed differences between 50–60${}^{\circ }$ N).

One potentially important factor is dry deposition of ${\mathrm{O}}_{\mathrm{3}}$ to the land surface. Due to the heatwave, stress on the biosphere and the associated dieback of vegetation could potentially reduce the efficiency of ${\mathrm{O}}_{\mathrm{3}}$ deposition, decreasing the ${\mathrm{O}}_{\mathrm{3}}$ sink (i.e. ${\mathrm{O}}_{\mathrm{3}}$ is more likely to deposit onto land covered by vegetation than bare soil). Investigation of the normalized difference vegetation index (NDVI), from the IMS scheme, between the summers of 2017 and 2018 did not highlight any spatially coherent changes (not shown here). As a result, there is no obvious large-scale spatial vegetation dieback in 2018 due to the heatwave and thus the impact this would have on ozone deposition in TOMCAT. Therefore, we ran two further experiments where the bare-soil fraction for each grid box over Europe was increased and decreased by 25 % in summer 2018. This was to investigate the sensitivity of surface ozone deposition to changes in the land surface. For the increase in bare-soil fraction there was a moderate systematic increase in European summer ozone by 0.0–1.5 $\mathrm{ppbv}$ (i.e. less ozone deposition). When the bare-soil fraction was decreased by 25 %, this yielded a small decrease in surface ozone by approximately 0.5 $\mathrm{ppbv}$. Overall, a sizable level of vegetation dieback would be required for decreased ozone dry deposition to substantially contribute to the summer 2018 surface ozone enhancements.

4 Discussion and conclusions

The summer of 2018 produced an intense heatwave across most of Europe with a substantial impact on tropospheric temperatures, droughts, stress on vegetation and human mortality. Observations of surface temperature, precursor gases and total column ${\mathrm{O}}_{\mathrm{3}}$ (${\mathrm{TCO}}_{\mathrm{3}}$) experienced enhancements in 2018 relative to 2017. In this paper, we have demonstrated a strong enhancement in surface and tropospheric ${\mathrm{O}}_{\mathrm{3}}$ during the heatwave between May and August 2018. The EMEP surface data suggest an average European enhancement, relative to 2017, peaking at approximately 10.0 $\mathrm{ppbv}$ in July and August. Investigation of lower-tropospheric ${\mathrm{O}}_{\mathrm{3}}$ (i.e. surface – 450 $\mathrm{hPa}$ sub-column ${\mathrm{O}}_{\mathrm{3}}$ – ${\mathrm{SCO}}_{\mathrm{3}}$) from the GOME-2 and IASI instruments also showed enhancements, peaking at 5.0–10.0 $\mathrm{DU}$, relative to 2017. Analysis of the long-term GOME-2 ${\mathrm{SCO}}_{\mathrm{3}}$ record indicates 2017 to be a suitably neutral/average reference year and the enhancement in 2018 to be anomalously large. Our comparisons were therefore made between the summers of 2017 and 2018.

Consistency between the UV (GOME-2) and IR (IASI) sounders was important to our analysis because their vertical sensitivities peak in the lower and mid-upper troposphere, respectively. The similar patterns of ${\mathrm{SCO}}_{\mathrm{3}}$ enhancement detected by the two sounders therefore indicate that these extend over the bulk of the troposphere, supportive of surface/lower-tropospheric ozone enhancements. This consistency also provides confidence that the complementary vertical sensitivities of GOME-2 and IASI ozone retrievals could be exploited in further investigations of tropospheric ozone in the future (e.g. long-term trends from multiple platforms/retrieval schemes have shown large-scale inconsistencies in other studies; e.g. Gaudel et al., 2018).

Tropospheric ${\mathrm{O}}_{\mathrm{3}}$ behaviour is complex, and the summer 2018 enhancements over Europe could potentially have been caused by various factors: atmospheric chemistry, meteorology (e.g. temperature, advection of ${\mathrm{O}}_{\mathrm{3}}$-rich air masses), anthropogenic and natural precursor emissions, dry deposition, and stratospheric intrusion. To investigate the interactions between these processes, potentially leading to the summer 2018 ${\mathrm{O}}_{\mathrm{3}}$ enhancements, we used the well-evaluated TOMCAT 3D CTM. Evaluation of the model in this study showed that it could accurately capture the spatial pattern, temporal evolution and sign (i.e. positive 2018–2017 ${\mathrm{O}}_{\mathrm{3}}$ differences) of the ${\mathrm{O}}_{\mathrm{3}}$ enhancements and that, although it underestimated the observed enhancements, TOMCAT is an adequate tool to investigate them.

The results of several model simulations showed that the surface ozone enhancements (mainly in north-western Europe) in the summer of 2018 were predominantly driven by meteorological processes with emissions acting as a secondary factor. As the ROTRAJ back-trajectories suggest that advection of summertime ${\mathrm{O}}_{\mathrm{3}}$ was larger in 2017, the 2018 European ${\mathrm{O}}_{\mathrm{3}}$ enhancements at surface level were likely from in situ processes. The TOMCAT stratospheric ${\mathrm{O}}_{\mathrm{3}}$ tracer indicated a negligible contribution of stratospheric ${\mathrm{O}}_{\mathrm{3}}$ to these surface enhancements. At 500 $\mathrm{hPa}$, the enhancement in tropospheric ${\mathrm{O}}_{\mathrm{3}}$ is much larger spatially across Europe and dominated by meteorological processes. Intrusion of stratospheric ${\mathrm{O}}_{\mathrm{3}}$ into the mid-troposphere has a moderate influence on the observed/modelled ${\mathrm{O}}_{\mathrm{3}}$ enhancements with contributions of up to 15.0 %–40.0 %. Correlations between TOMCAT temperature and ${\mathrm{O}}_{\mathrm{3}}$ enhancements show broad agreement at some latitudes (e.g. 50–70${}^{\circ }$ N in the lower–mid-troposphere). However, this relationship is non-linear and difficult to quantify without further simulations/model tracers, which was beyond the scope of this study. ROTRAJ back-trajectories suggest that in 2018, relative to 2017, there is the advection of more ${\mathrm{O}}_{\mathrm{3}}$-rich air masses into the European mid-troposphere contributing to the summer 2018 ${\mathrm{O}}_{\mathrm{3}}$ enhancements at this altitude. Therefore, in the summer of 2018 over Europe, in situ meteorological processes appear to be predominantly driving surface ${\mathrm{O}}_{\mathrm{3}}$ enhancements over Europe, while advection of tropospheric ${\mathrm{O}}_{\mathrm{3}}$-rich air and stratospheric intrusion are driving the corresponding tropospheric ${\mathrm{O}}_{\mathrm{3}}$ enhancements

Overall, through our study focusing on the European summer 2018 air pollution episode, we have demonstrated the use of novel satellite data sets and a modelling framework (i.e. targeted sensitivity experiments and model tracers) suitable to investigate the air quality impacts from future European heatwaves such as that which occurred in summer 2022.

Data availability

The EMEP surface ${\mathrm{O}}_{\mathrm{3}}$ data were obtained from http://ebas-data.nilu.no/default.aspx (EMEP, 2023). The GOME-2 tropospheric column ${\mathrm{NO}}_{\mathrm{2}}$ data were downloaded from EUMETSAT at https://acsaf.org/nrt_access.php (ACSAF, 2023). The TOMCAT and RAL Space satellite data have been uploaded to the Zenodo open-access portal (10.5281/zenodo.10001068, Pope and RAL Space, 2023).

The supplement related to this article is available online at: https://doi.org/10.5194/acp-23-13235-2023-supplement.

Author contributions

RJP, MPC and BJK conceptualized and planned the research study. RJP performed the TOMCAT model simulations with support from MPC and WF. The JULES BVOC emissions were provided by ECP and GDH. RJP analysed the satellite data provided by RAL Space (BJK, RS, BGL and LJV) with support from BJK, RS and BGL. RJP undertook the EMEP analysis. RJP ran ROTRAJ with technical support from SRA and AMG. RJP prepared the manuscript with contributions from all co-authors.

Competing interests

The contact author has declared that none of the authors has any competing interests.

Disclaimer

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Acknowledgements

This work was funded by the UK Natural Environment Research Council (NERC) by providing funding for the National Centre for Earth Observation (NCEO, award reference NE/R016518/1).

Financial support

This research has been supported by the National Centre for Earth Observation (grant no. NE/R016518/1).

Review statement

This paper was edited by Tim Butler and reviewed by two anonymous referees.