Stratospheric impact on the Northern Hemisphere winter and spring ozone interannual variability in the troposphere

In this study we use ozone and stratospheric ozone tracer simulations from the high-resolution (0.5°x0.5°) Goddard Earth Observing System, Version 5 (GEOS-5) in a replay mode to study the impact of stratospheric ozone on tropospheric ozone interannual variability (IAV). We use these simulations in conjunction with ozone sonde measurements from 1990 to 2016 during the winter and spring seasons. The simulations include a stratospheric ozone tracer (StratO3) to aid 15 in the evaluation of the impact of stratospheric ozone IAV on the IAV of tropospheric ozone at different altitudes and location. The model is in good agreement with the observed interannual variation of tropospheric ozone, except for the postPinatubo period (1992-1994) over the region of North America. Ozonesonde data show a negative ozone anomaly in 19921994 following the Pinatubo eruption, with recovery thereafter. The simulated anomaly is only half the magnitude of that observed. Our analysis suggests that the simulated Stratosphere-troposphere exchange (STE) flux deduced from the analysis 20 might be too strong over the North American (50°N-70°N) region after the Mt. Pinatubo eruption in the early 1990s, masking the impact of lower stratospheric ozone concentration on tropospheric ozone. European ozonesonde measurements show a similar but weaker ozone depletion after the Mt. Pinatubo eruption, which is fully reproduced by the model. Analysis based on the stratospheric ozone tracer (StratO3) identifies differences in strength and vertical extent of stratospheric ozone impact on the tropospheric ozone interannual variation (IAV) between North America and Europe. Over North America 25 stations, the StratO3 IAV has a significant impact on tropospheric ozone from the upper to lower troposphere and explains about 60% and 66% of the simulated ozone IAV at 400 hPa, and ~11% and 34% at 700 hPa in winter and spring respectively. Over European stations, the influence is limited to the middle to upper troposphere, and becomes much smaller at 700 hPa. The MERRA2 assimilated fields exhibit strong longitudinal variations over northern hemisphere (NH) mid-high latitudes, with lower tropopause height and lower geopotential height over North America than Europe. These variations 30 associated with the relevant variations in the location of tropospheric jet flows are responsible for the longitudinal differences in the stratospheric ozone impact, with stronger effects over North America than over Europe.


Introduction
Tropospheric ozone plays an important role in the oxidative capacity of the troposphere. In the upper troposphere, it is the third most important greenhouse gas after carbon dioxide and methane, and affects the radiative balance of the atmosphere 35 (Forster et al., 2007). Unlike the well-mixed greenhouse gases, tropospheric ozone and its radiative forcing are spatially and temporally inhomogeneous (Lacis et al., 1990;Forster and Shine, 1997;Joiner et al., 2009;Worden et al., 2008;Bowman et al., 2013). Stratosphere-troposphere exchange (STE) has been shown to impact the tropospheric ozone distribution (e.g., Holton et al., 1995;Terao et al., 2008;Hess et al., 2015;Williams et al., 2019). Liu et al. (2017) showed that stratospheric ozone input plays a dominant role in driving the interannual variation (IAV) of upper tropospheric ozone over the southern 40 hemisphere ocean, where its radiative impact is largest. Considering the observed and expected net global decrease in emissions of ozone precursors and the predicted increase in ozone STE (e.g., Collins et al., 2003;Sudo et al., 2003;Hardiman et al., 2014;Banerjee et al., 2016), it is important to quantify the role of stratospheric ozone in comparison to that of precursor emissions in determining the tropospheric ozone distribution.
In this study we use a long-term, full chemistry simulation of ozone and a 'stratospheric ozone tracer' (StratO3) by the 45 Goddard Earth Observing System V5 (GEOS-5) -chemistry climate model (CCM), as well as the analyzed meteorological fields, to interpret the tropospheric ozone IAV derived from the ozonesonde measurement in the northern hemisphere midhigh latitudes. In doing so, we examine the vertical and longitudinal distribution of the stratospheric ozone impact on the IAV of tropospheric ozone and their linkage to transport. STE has been the subject on many studies for several decades (Danielsen, 1968;Holton et al., 1995;Olsen et al., 2002Olsen et al., , 50 2003Sprenger and Wernli, 2003;Stohl et al., 2003a;2003b;Thompson et al., 2007;Lefohn et al., 2011;Skerlak et al., 2014;Williams et al., 2019). It contributes significantly to ozone in the upper troposphere, where ozone has a strong radiative effect. Observations, assimilations and simulations from high resolution models show that deep STE events occasionally reach ground level, adversely affecting the air quality near the surface (e.g., Haagenson et al., 1981;Davies and Schuepbach, 1994;Lefohn et al., 2001;Langford et al., 2012;Lin et al., 2012;Ott et al., 2016;Knowland et al., 2017;Akritidis et 55 al., 2018). In addition, various chemistry climate models project increased STE leading to a higher contribution of stratospheric ozone to tropospheric ozone (Collins et al., 2003;Sudo et al., 2003;SPARC-CCMVal, 2010;Zeng et al., 2010).
Limitations in the representation of small-scale stratospheric intrusions lead to uncertainties in the calculated stratospheric contribution to concentrations and variations of tropospheric ozone at the spatial scales of a global model. These limitations also lead to uncertainty in their relative magnitudes compared to the effects of increased or decreased emissions of ozone 60 precursors. The uncertainties in stratospheric contribution to tropospheric ozone variations lead to similar uncertainties in resulting ozone radiative forcing, a key area of focus in climate change studies.
Various studies have used chemistry transport models (CTMs) to examine the response of tropospheric ozone to changes in stratospheric input and in surface emissions; these models have used a simple treatment of stratospheric-tropospheric flux, adopting either the SYNOZ (synthetic ozone) approximation developed by McLinden et al. (2000) to specify the 65 stratosphere-to-troposphere flux (e.g., the GEOS-Chem model in Fusco and Logan, 2003;Hess and Zbinden, 2013), or specifying ozone in the lower stratosphere (LS) (the GISS model in Fusco and Logan, 2003;Karlsdottir et al., 2000). Hess et al. (2015) analyzed the effects of stratospheric input to tropospheric ozone variations over the northern hemisphere midlatitudes with four ensemble simulations of the free running Whole Atmosphere Community Climate Model (WACCM) for 1953 to 2005. Their model used a standard stratospheric chemical mechanism and simple CH4-NOX chemistry in the 70 troposphere with constant surface emissions of ozone precursors. The study reproduced well the observed tropospheric ozone IAV, suggesting that natural variability in transport and stratospheric ozone plays a significant role in the tropospheric ozone IAV over the northern hemisphere. Williams et al (2019) used a nudged CCM simulation with the ERA-Interim reanalysis and a stratospheric tagged ozone tracer to assess the role of stratospheric ozone in influencing both regional and seasonal variations of tropospheric ozone. Their study showed that the stratosphere has a much larger influence than 75 previously estimated, though some differences from other studies may be due to different definitions of the stratospheric tracer.
In this study, we use a long-term full chemistry GEOS-CCM replay simulation with a stratospheric ozone tracer at a horizontal resolution of 0.5°. This is the suggested minimum model resolution needed to resolve the structure of deep STE events (Ott et al., 2016). We focus on 1990 -2016, a period of considerable IAV in STE (James et al., 2003), varied trends in 80 emissions of ozone precursors, and greater availability of reliable ozone observations than in prior periods. We examine the vertical extents of STE impact on tropospheric ozone using model simulations and ozonesonde measurements sampled over North America and Europe. We rely on the StratO3 tracer simulation to quantify the contribution of stratospheric ozone to tropospheric ozone at different levels, as well as its contribution to the IAV.
2 Data and Model 85

Ozonesondes
We select 17 ozonesonde sites, eight in North America and nine in Europe, all of which have a record of at least 3 profiles every month between 1990 and 2016 (Figure 1 and Tables 1 & 2). The data are obtained from the World Ozone and Ultraviolet Data Center (WOUDC, http://www.woudc.org). Observations over most stations are obtained using electrochemical concentration cells (ECC), which rely on the oxidation reaction of ozone with potassium iodide in solution 90 (Komhyr et al., 1995). At Hohenpeissenberg, Germany, measurements are obtained using the Brewer/Mast instrument. The sondes ozone measurements have a vertical resolution of ~150 m, with an accuracy of ±5% in the troposphere (WMO, 2014).

MERRA2-GMI
We use a replay simulation (http://acd-ext.gsfc.nasa.gov/Projects/GEOSCCM/MERRA2GMI) of the GEOSCCM with the 95 Global Modeling Initiative (GMI) chemical mechanism (Duncan et al., 2007;Strahan et al., 2007) for gas chemistry, which includes a complete treatment of stratospheric and tropospheric chemistry, and the Goddard Chemistry Aerosol Radiation and Transport (GOCART) module (Chin et al., 2002;Colarco et al., 2010) for aerosols. The replay simulation follows the replay methodology as described in Orbe et al. (2017) and uses the RAs3 setting, which best represents overall transport.
The model inputs the three-hourly time-averaged output of MERRA-2 meteorology (U, V, T, pressure) and recomputes the 100 analysis increments, which are used as a forcing to the meteorology at every time step over the 3 h replay interval. The replay simulation is run at a MERRA-2 native resolution of ~50 km in the horizontal dimension and 72 vertical levels. This replay simulation is referred to as the 'MERRA2-GMI" simulation.  (Giglio et al., 2013(Giglio et al., ) after 1997 Prior to 1997, biomass burning emissions are based on a GFED4s climatology with year-to-year variability imposed using regional scale factors derived from the Total Ozone Mapping Spectrometer (TOMS) aerosol index (Duncan et al., 2003).
The simulation used the Model of Emissions of Gases and Aerosols from Nature (MEGAN) (Guenther et al., 2006) to simulate biogenic emissions, including isoprene, within the model. The lightning parameterization in the model (Allen et al., 2010) is constrained by the MERRA-2 detrended cumulative mass flux, with seasonal constraints from the Lightning 115 Imaging Sensor (LIS) / Optical Transient Detector (OTD) v2.3 climatology (Cecil et al., 2014). Methane is specified using latitude and time dependent surface observations from the NOAA Earth System Research Laboratory (ESRL) Global Monitoring Division (GMD) network (Dlugokencky et al., 2011). Our initial model evaluations suggest that the MERRA2-GMI ozone simulation are in good agreement with the means and variability of the total and tropospheric column ozone from satellite observations ( Figure S1 and S2). 120 A StratO3 tracer is included in the model to diagnose the stratospheric ozone influence on the troposphere. StratO3 is set equal to simulated ozone in the stratosphere and is removed in the troposphere based on interannually-varying monthly mean loss rates and surface deposition fluxes archived from a standard full chemistry simulation. StratO3 tracer is defined relative to a dynamically varying tropopause, which is derived from an artificial tracer, e90, introduced by Prather et al. (2011). The e90 tracer is set to a uniform mixing ratio (100 ppb) at the surface with a 90-day e-folding lifetime everywhere in the 125 atmosphere. This lifetime is long enough for the tracer to be well mixed throughout the troposphere but short compared to the transport time scales in the stratosphere, resulting in sharp e90 tracer gradients across the tropopause. The e90 tropopause is optimal in effectively separating stratospheric from tropospheric air from a chemical composition perspective compared to other traditional definitions of the tropopause (Prather et al. 2011), and offers the advantage of being able to be calculated online in the model. In our simulations, the e90 tropopause value is set to 90 ppb. The e90 tracer has been used in many 130 studies of STE as an accurate tropopause definition and an ideal transport tracer in the UTLS (e.g., Hsu and Prather, 2014;Randel et al., 2016).
The MERRA2-GMI simulation has hourly output for ozone and three-hourly output for StratO3 at each model level. When comparing to the ozonesonde measurements, the model outputs are sampled at the nearest grid point and launch time for each sonde. 135

Winter and spring ozone IAV in the lower stratosphere and troposphere over North American and European sites
Previous studies have shown that the relative contribution of stratospheric ozone to tropospheric ozone is greatest in the free 140 troposphere during winter (e.g., Holton et al., 1995;Stohl et al., 2000;Skerlak et al., 2014; and at the surface during spring (e.g., Lin et al., 2012;. In summer, the relative contribution of stratospheric ozone is low due to the increased chemical ozone production in the troposphere. Here, we focus on the winter (DJF) and spring (MAM) seasons to examine the interannual variations of the strength and vertical extent of stratospheric ozone impact on the tropospheric ozone. and spring seasons from 1990 to 2016 averaged over sites from North America and Europe. Anomaly at each site is calculated by removing the respective seasonal mean climatology from 1995 to 2016, and then averaged over all sites for each region (see Supplement). The shaded area represents the 95% CI of the calculated mean from daily observations over all the selected stations. To quantify the magnitude of IAVs, we adopt the standard deviation (SD) of these ozone anomalies.
We perform the standard statistical F-test to assess the regional and seasonal differences in the ozone IAVs. The calculated 150 standard deviations and F-test statistics are shown in Table S1 and S2.
At 200 hPa, the model reproduces well the observed IAV in both seasons over both regions (r ³ 0.91, Figure 2 a-d). There are no significant differences in the magnitude of ozone IAV between North America and Europe (Table S1). But over both regions, the ozone IAV show significant seasonal differences with greater magnitude in spring than in winter (Table S2).
Negative ozone anomalies occur in the early 1990s and at the end of the record from 2014 to 2016, while positive anomalies 155 are obtained for most years between 1998 and 2013. The negative ozone anomalies during the period of 1992-1996 are consistent with the chemical and dynamical perturbations following the June 15 th 1991 eruption of Mt. Pinatubo (Hadjinicolaou et al., 1997;Stenchikov et al., 2002;Rozanov et al., 2002). The negative ozone anomaly in 2015-2016 is associated with stratospheric circulation changes caused by the unusually warm ENSO event coinciding with a disrupted Quasi-Biennial Oscillation (QBO) during that period (Tweedy et al., 2017;Diallo et al., 2018). The maximum positive anomaly, observed in 1990-1991, is not reproduced by the model. The model-observation correlation coefficients increase significantly if we omit these two years (from 0.18 to 0.58 in the winter and from 0.43 to 0.58 in the spring). At 700 hPa, the model reproduces the magnitude of observed ozone depletion after the Mt. Pinatubo eruption.
Unlike at 200 hPa, the differences in the magnitude of ozone anomalies between North America and Europe are significant at 175 400 hPa and 700 hPa, with smaller ozone anomalies over European sites than over North American sites in both seasons (Table S1). The magnitude of ozone anomalies doesn't show significant seasonal difference between DJF and MAM, except over North America at 400 hPa (Table S2).
We use explained variance (square of correlation coefficient: r 2 ) to determine the fraction of the ozone variance that in the troposphere that can be attributed to the variance in stratospheric ozone. Table 3  hot spots along western U.S. in the spring season (Lin et al., 2012;Skerlak et al., 2014;Langford et al., 2015).
The correlations between the stratosphere and troposphere IAV in both observations and model simulations suggest a potential impact of stratospheric ozone on tropospheric ozone variations. Previous studies have found high correlations between ozone in the lower stratosphere with that in the middle and lower troposphere with the largest effects in late winter and spring. Correlation does not necessarily mean causality, and to date model investigations of this correlation (Terao et al., 195 2008;Hess and Zbinden, 2013) did not use a model with both stratospheric and tropospheric chemistry, and up-to-date stratospheric circulation. The MERRA2-GMI simulation has both of these attributes, detailed dynamic diagnostics, and StratO3 as described in Section 2.2. In the next section, we use the StratO3 tracer to examine the contribution of stratospheric ozone to the IAV of tropospheric ozone, as a function of altitude, season, and location. We will also use diagnostics from the model to explore the influence of transport on the stratospheric ozone contribution to the tropospheric ozone and its IAV. 200 Figure 3 shows the same comparison between the observed (black lines) and simulated ozone (red lines) anomalies as in Figure 2, but adding the anomalies of simulated StratO3 (green lines). As expected, the StratO3 anomalies at 200 hPa are almost identical to the simulated ozone anomalies, since most measurements are in the stratosphere at this level.

Impact of stratospheric ozone on tropospheric ozone IAV
The variability in the amount of tropospheric ozone that was transported from the stratosphere as inferred by StratO3 is due 205 to both the variability in the lower stratospheric ozone reservoir and the variability in the net downward mass flux (Albers et al., 2018). These two variabilities may either cancel or reinforce each other, depending on their relative phases. At 400 hPa, over the North American stations, the minimum and maximum of StratO3 tracer is highly correlated with the minimum and maximum of simulated ozone. The IAV of StratO3 explains more than 60% of simulated ozone variations (Figure 3 e, f, r = 0.77 in DJF and 0.81 in MAM), suggesting that the changes of stratospheric ozone input strongly impact the simulated ozone 210 IAV in the upper troposphere. The correlation between StratO3 and observed ozone is slightly lower (0.44) than that with simulated ozone in DJF over North America. The decreased correlation is mainly due to the model-observation discrepancy The bottom panel of Figure 3 compares the simulated StratO3 anomalies to the observed and simulated ozone anomalies at 700 hPa over North American and European ozonesonde sites during winter and spring. Over North America, the observed 225 ozone anomalies stay low in the early 1990s, and increase thereafter in both seasons, which is underestimated in the model.
In the winter, StratO3 anomalies decrease slightly in contrast to increases in both observed and simulated ozone anomalies.
The winter StratO3-O3 correlation is ~ 0. In spring, both observed and simulated ozone exhibit similar IAV, which agree with the phase of the StratO3 anomalies after the Pinatubo period (1991)(1992)(1993)(1994)(1995). The StratO3-O3 correlation increases from 0.07 to 0.33 in winter and from 0.36 to 0.58 in spring from 1996 to 2016. 230 Over North America, our model results are in good agreement with the observed IAV at all levels except right after the Mt.
Pinatubo eruption. The model only reproduces about half of the observed tropospheric depletion over North America. As discussed above, this could be due to an excessive mass flux from the stratosphere in the MERRA-2 analysis during this period. Model results are in better agreement with the magnitude of observed ozone depletion after the Mt. Pinatubo eruption in the middle and lower troposphere over Europe. There is no significant relationship between StratO3 and simulated ozone 235 at 700 hPa. This is expected since the impact of stratospheric ozone decreases, and the impact of ozone production from its precursors becomes more important at lower altitudes. In summary, our model analysis identifies differences in the strength and vertical extent of stratospheric ozone impact on the tropospheric ozone IAV between North America and Europe. Over North America, the StratO3 IAV has a significant impact on the tropospheric ozone IAV from the upper to lower troposphere and explains 60% and 66% of the simulated ozone IAV at 400 hPa, and 11% and 34% at 700 hPa in winter and spring 240 respectively, after 1995. Over Europe, the influence is limited to the middle to upper troposphere, and becomes much less at 700 hPa.
The differences in the stratospheric ozone impact between North America and Europe are likely due to variations in the net downward flux associated with planetary-scale waves. Previous studies have suggested that the IAV of the STE mass flux is likely correlated to changes in the tropopause height (e.g., Gettelman et al., 2011). The top panel of Figure 4 shows the 245 comparison of the observed ozone mixing ratio anomalies at 400 hPa and the tropopause pressures derived from the observed ozone profiles following the criteria in vertical gradient and ozone mixing ratio given by Browell et al. (1996). The tropopause pressure was estimated to be at the pressure where a linear regression line passing through the lower stratospheric ozone profile (150 ppb -400 ppb, lower than 100 hPa) intersects with the 100-ppb ozone level. The bottom panel of Figure 4 compares the simulated ozone and StratO3 anomalies at 400 hPa with the tropopause pressures derived from simulated ozone 250 profiles following the same criteria as for the observations. As expected, the IAV of ozone and StratO3 positively correlates with that of the derived tropopause pressure (anticorrelates with the tropopause height) in both model and observation. In general, during years with a lower tropopause, stratospheric ozone influence at 400 hPa increases and results in a positive ozone anomaly. During years with a higher tropopause, decreased stratospheric ozone influence leads to a negative ozone anomaly at 400 hPa. 255 The above high correlations between the IAV of tropopause pressure and StratO3 raise the question of what dynamical conditions control the higher/lower tropopause pressures, STE mass fluxes, and the subsequent impact of stratospheric ozone on tropospheric ozone. These questions are particularly important if these dynamical conditions change in the future as a result of climate change. In the next section, we rely on the model's 3-d dynamical diagnostics, including air mass flux and horizontal wind patterns, to examine both the vertical and horizontal transport influence of the stratospheric ozone 260 contribution on the tropospheric ozone and its IAV. We also examine the longitudinal difference in the model's dynamics to explain the identified differences in the stratospheric ozone impact between North America and Europe. associated with these wave disturbances are responsible for a large fraction of ozone temporal and spatial variability in winter and spring (e.g., Kinnersley and Tung, 1998;McCormack et al., 1998;Lozitsky et al., 2011;Zhang et al., 2015).
In this section, we examine the linkage of the vertical and horizontal transport to the stratospheric ozone contribution in the troposphere using the model's 3-d air mass flux and horizontal winds. Our analysis focuses on the year 1993, when there is a major discrepancy with the observations at 400 hPa as shown in Figure 3. 275 In the winter of 1993, the jet at 400 hPa exhibited a typical wave pattern, with a trough over eastern North America, and ridges over western North America and western Europe (Figure 5a). Strong northwesterly winds prevailed north of 50°N over western North America. They converged with the Westerlies around 45°N in eastern North America. The winds then 290 changed direction to southeasterly and flowed into the North Atlantic and Europe, bringing warmer tropospheric low ozone air into western Europe (Figure 5a). The maps of the air mass flux and its anomalies ( Figure 5 e and g) suggest that North America between 50°N and 70°N was dominated by more vigorous downward mass fluxes of stratospheric air. Meanwhile, the northwesterly winds brought ozone rich air from high latitudes, resulting in a positive anomaly of stratospheric ozone influence at 400 hPa (Figure 5c). Although the lower stratospheric ozone level decreased significantly in the winter of 1993 295 due to the Pinatubo eruption (Figure 2), the enhanced downward mass fluxes across the tropopause in the model counteracted the depletion and led to a positive ozone anomaly at 400 hPa over North America between 50°N and 70°N ( Figure S3). Over the high latitudes (> 70°N), where there is less dynamic perturbation (including both vertical and horizontal transport), the stratospheric ozone contribution at 400 hPa was largely driven by the depletion of the ozone concentration in the lower stratosphere and showed a strong negative anomaly in 1993 (Figure 5 c). Most of the European 300 region was covered by the increased downward air mass flux near the tropopause in the winter of 1993. However, a negative anomaly of the StratO3 contribution at 400 hPa was seen over western Europe. It is likely that the combined negative effects of the ozone depletion in the lower stratosphere and the downwind of the warmer tropospheric low ozone air from the subtropical North Atlantic Ocean exceeded the positive effect of the increased downward air mass flux over this region.
In the spring of 1993, regions with maximum StratO3/O3 at 400 hPa located further north compared to those in the winter 305 and centered around the northern part of the Labrador Sea. Southwesterly wind prevailed south of 65°N over western North America, bringing in low StratO3 oceanic air from the subtropics. The winds changed direction to northwesterly around  Pinatubo eruption and reached a regional maximum (Figure 2). In the winter of 1998, a poleward shift of the jet occurred over most of North America. The jet location as well as the regions with the maximum StratO3/O3 ratio moved to the north 320 by about 7° compared to the winter of 1993 (Figure 6 a). With the poleward shift of the jet, the StratO3/O3 anomaly exhibited a negative (positive) maximum along the equatorward (poleward) of the jet. Southwesterly winds prevailed over regions equatorward of the jet at 400 hPa, bringing in tropical oceanic low ozone air to North America. Around the tropopause, there were increased upward air mass fluxes along the west coast of North America and decreased downward air mass fluxes over western and central North America. Therefore, although there was an increase in the stratospheric ozone concentration in 325 1998, the wind patterns in the troposphere and the vertical mass flux around the tropopause associated with the northward shift of the jet system caused a less favorable transport of stratospheric ozone into troposphere. This resulted in a negative anomaly of StratO3/O3 at 400 hPa over most of North America. In the spring of 1998, a similar poleward shift of the jet occurred over North America as that in the winter, with a negative StratO3/O3 anomaly over most of North America north of 45°N (Figure 6 b, d). . 330 Our analysis suggests that significant interannual variations exist in both the regional wind patterns associated with the Westerlies waves and the strength of downward air mass fluxes across the tropopause. The IAV of stratospheric ozone influence in the troposphere reflects a combined effect of the IAV in the 3-d dynamics as well as that in the lower stratospheric ozone concentration, which may either oppose or reinforce each other.

Longitudinal difference of stratospheric ozone influence 335
Our analysis based on data and model sampled at sonde sites identified differences in the strength and vertical extent of stratospheric ozone impact on tropospheric ozone IAV between North America and Europe. Over North American sites, a significant impact of the StratO3 IAV on tropospheric ozone extends to the lower troposphere. Over European sites, the influence is limited to the middle to upper troposphere. In this section, we examine whether the longitudinal differences seen over the ozonesonde sites is a large-scale phenomenon, by extending our analysis to a broader spatial domain. Figure 7 and 340 8 show the latitudinal average (30°N to 80°N) of tropopause pressure, geopotential height at 400 hPa, and the StratO3/O3 ratio at 400 hPa at each longitude from 180°W to 180°E from 1990 to 2016 in winter and spring. The geopotential heights and tropopause pressure are good diagnostics of large-scale circulation patterns. All of them show strong longitudinal difference between North America (120°W-60°W) and Europe (10°W-26°E), with lower geopotential height, higher tropopause pressure (lower tropopause height), and greater stratospheric ozone contribution over North America than 345 Europe. The longitudinal gradients between North America and Europe are slightly weaker in spring than in winter. The spatial map of StratO3/O3 climatology at 400 hPa averaged from 1990 to 2016 suggests that the longitudinal difference is persistent over most of the mid-high latitudes ( Figure S4), and is closely related to the wavelike pattern in jets. The climatology of jet meanders to the south over central and eastern North America and brings in cold polar air with more stratospheric subsidence. The jet moves to the north over Europe and brings in warm air with less stratospheric ozone 350 influence. Skerlak et al. (2014) identified the deep STE hot spots along the western North America using the ERA-Interim reanalysis data set from the European Centre for Medium-Range Weather Forecasts (ECMWF) from 1979 to 2011. Therefore, over North America, the stratospheric subsidence inside the polar vortex as well as deep stratospheric intrusion events results in a deeper and greater stratospheric ozone influence on the tropospheric ozone than over Europe.
A modulating factor in the IAV is the Artic Oscillation (AO) -the primary mode of dynamical IAV in the troposphere 355 during winter. Several studies have examined the mechanism for downward transport from the stratosphere to the troposphere and attributed changes in the strength of lower-stratospheric polar vortex to AO anomalies at the surface, with a positive AO phase linked to a more isolated and stronger polar vortex (Ambaum and Hoskins, 2002;Perlwitz and Harnik, 2003) and lower tropopause heights. Lamarque and Hess (2004) found that the AO explains up to 50% of the IAV in tropospheric ozone over North America in January-March, but did not find any significant correlation in European sonde 360 data, with similar results from the Model for OZone And Related chemical Tracers (MOZART) model. They argued that the correlation may be caused by the influence of the AO on its modulation of STE as well as transport of ozone and its precursors. Kivi et al. (2007) found that changes in the AO explained most of the tropospheric ozone trends in January-April, based on analysis of Arctic ozonesonde data. Figure 9 shows the longitudinal variations of simulated ozone and AO correlation profiles averaged over 30°N and 80°N from 1000 hPa to 200 hPa. Over North America (120°W to 60°W), the 365 correlation between simulated ozone and the AO index is negative and stays low above 400 hPa. The anticorrelation increases with increased pressure and reaches its maximum near the surface around 90°W. The anticorrelation averaged over Europe (10°W-26°E) stays low above 400 hPa, increases slightly from 400 hPa to 700 hPa, then decreases sharply near the surface. This is similar to the correlations obtained from the ozonesonde profiles ( Figure S5). The similarity of correlation patterns over sonde sites and their surrounding broader regions indicates that the AO-related stratospheric 370 subsidence is a large-scale phenomenon.

Conclusions and discussion
In this study we used ozone and stratospheric ozone tracer simulations from MERRA-2 GMI and observations from ozonesondes to investigate the interannual variations and vertical extents of the stratospheric ozone impact on tropospheric ozone. Our work focuses on the winter and spring seasons over North America and Europe. 375 The model reproduces the observed interannual variations of tropospheric ozone over North America except for the Pinatubo period from 1991 to 1995. The ozonesonde data show a negative ozone anomaly in 1992-1994 following the Pinatubo eruption, with recovery thereafter. However, the simulated anomaly is about half the magnitude of the observed tropospheric ozone depletion. Over European regions, ozonesonde measurements show a similar but weaker ozone depletion, which was fully reproduced by the model. We use a stratospheric ozone tracer to gauge the impact of stratospheric ozone variations in 380 different regions of the troposphere. Our results based on the stratospheric ozone tracer suggest that the influence of the stratospheric IAV is significant in the upper to lower troposphere over North America, while over Europe, the stratospheric influence is limited to the middle to upper troposphere. Our analysis of the MERRA2 assimilated fields shows strong longitudinal variations in meteorological parameters over northern hemisphere mid-high latitudes, with lower tropopause height and lower geopotential height over North America than Europe. These variations associated with the relevant 385 variations in the location of tropospheric jet flows are responsible for the longitudinal change in the stratospheric ozone influence. The increase in frequency in stratospheric folds near the jets, and the strong winter subsidence inside the polar vortex lead to stronger stratospheric impact over North America than over Europe. Code/Data availability All data used for this article can be obtained by contacting J. Liu (email:Junhua.liu@nasa.gov).

Competing interests
The authors declare that they have no conflict of interest.

Acknowledgement 420
I gratefully acknowledge the financial support by NASA's Atmospheric Chemistry Modeling and Analysis Program (ACMAP) (grants NNX17AG58G). We thank the World Ozone Data Centre and the SHADOZ program for making the routine sonde data accessible. We gratefully acknowledge Dr. Jerry R Ziemke from NASA for providing the OMI/MLS TCO data and Dr. Stacey M. Frith from NASA for providing SBUV total ozone column data. Work was performed under contract with NASA at Goddard. We thank Dr. Clara Orbe for her helpful comments on the model's replay configuration. 425 Computer resources for the MERRA-2 GMI simulation were provided by the NASA Center for Climate Simulation. We thank the editor and the reviewers for their helpful comments and suggestions to improve this paper.