The two campaigns took place in June of 2012 and 2013. Both were of short duration (i.e., about 2 weeks). A description of the O3HP site is available in Kalogridis et al. (2014) and Santonja et al. (2015). In short, the site (43.93∘ N, 5.71∘ E) is located on the premises of Observatoire de Haute Provence, about 60 km north of Marseille, France, at an elevation of 680 m above mean sea level. It is implemented in a forest area that has remained untouched since at least 1945. The climate is sub-Mediterranean with dry, warm-to-hot summers.

The O3HP observatory is characterized by a highly heterogeneous karstic limestone with soil pockets developing between compact and hard limestone bedrock. The soils, which never exceed 1 m depth, range from shallow calcaric Leptosol to deeper calcaric Cambisols (IUSS Working Group WRB, 2014). The litter overlying the A horizons (O horizons) is 1–7 cm strong. The A horizons of 2–10 cm depth are clayey, calcareous, and show high organic carbon content (Table 1). These horizons have a strong, crumbly-to-fine subangular blocky structure likely due to high earthworm burrowing activity and numerous fine roots. The humus is an “active oligomull or dysmull type” (Brêthes et al., 1995). The A/C horizon consists of thin layers of a clayey and fine blocky soil material between limestone rocks of a decametric size. Roots are observed inside the thin soil layers.

Downy oak (Quercus pubescens) and Montpellier maple (Acer monspessulanum L.) represent 75 and 25 %, respectively, of the foliar biomass of the dominant tree species (Kalogridis et al., 2014). The coppice, typically constituted by multiple stems sprouting from the same rooting system, is about 70 years old. Mean tree height is 5 m, and mean diameter at breast height is 10 cm, ranging from 0.9 to 18.6 cm. European smoke bush (Cotinus coggygria Scop.) and many thermophilic and xerophilic herbaceous and grass species compose the understorey vegetation (Kalogridis et al., 2014). A network of soil sensors beneath and above the canopy continuously record environmental parameters, including global radiation, air and soil temperature profiles, air and soil moisture, wind speed, and rainfall, which are made accessible through the COOPERATE database (http://cooperate.obs-hp.fr/db).

Our understanding of the atmospheric dynamics over the O3HP sampling site does not rely solely on meteorological parameters recorded at ground level by basic weather stations. The transport and dispersion of air pollutants in the southeastern part of France was extensively investigated during the “Expérience sur Site pour Contraindre les Modèles de Pollution atmosphérique et de Transport d'Emissions” (ESCOMPTE) experiment, which took place in June–July 2001 (Cros et al., 2004; Kalthoff et al., 2005). As shown by these authors for June 2001 and in Fig. S1 in the Supplement for June of 2012, 2013, and 2015, the sea breeze is a general characteristic of the atmospheric dynamics at the site in June. It flows from the W-SW in the afternoon and carries with it the photosmog of the city of Marseille. During the night and early morning hours the wind is oriented from other directions with a strong N-NE component (Fig. S1). However, one fundamental aspect of air circulation over the area is the existence of a nocturnal jet flowing at 800–1000 m of altitude, also with a strong N-NE component, observed in the sodar (vertical wind profiler) measurements performed by Kalthoff et al. (2005). This is of crucial importance for the interpretation of our results.

The cumulated precipitations before the campaigns were about 400 and 500 mm since the beginning of the year (Fig. 1a). As few precipitation events of small intensity took place during the campaigns, the volumetric soil water content (measured at 5 cm depth) was in a decreasing phase from about 0.3 m3 m-3 during the wet season to about 0.1 m3 m-3 during the dry season (Fig. 1b). Soil temperatures went the opposite way (Fig. 1b) and were in the range of 14–19 and 14–17 ∘C during the 2012 and 2013 campaigns, respectively (Fig. 1c, d).

In June of 2012, CO2 presented a clear and reproducible diurnal cycle with a maximum during the night (Fig. 2c). This maximum, an increase of 10–20 ppm, correlates to the decrease of global radiation (Fig. 2a). This increase occurred between the period of maximum atmospheric turbulence (u∗ > 0.4 m s-1, Fig. 2b), a few hours after the maximum solar radiation (Fig. 2a), and the nocturnal period when atmospheric turbulence is reduced (u∗ < 0.2 m s-1, Fig. 2b) and strong temperature gradients above ground level form (∼ -0.5 ∘C m-1, Fig. 2a). The temperature gradient is a proxy of low atmospheric mixing and boundary layer stability. During this period, the variability in OCS was relatively low as compared to CO2 (10 ppt at the most). The strongest temperature gradients above ground level (∼ -1 ∘C m-1, Fig. 2a) were observed after sunrise (04:00 UTC), for about 2 h. The diel cycle in the atmospheric boundary layer exhibited a much steeper decline in OCS after sunrise than at night (Fig. 2c); the same holds for ozone (Fig. 2d). The amplitude of the early morning drop of OCS was in the 60–100 ppt range. That of O3 was in the range of 15–30 ppb. It is worth noting that the large nocturnal maximum of CO2 was followed by a secondary one in the early morning but of shorter duration and smaller amplitude (10 ppm at the most, Fig. 2c). Hence, important variations in CO2 were observed during the period of lowest OCS concentrations. In general, OCS and O3 diel variations were in phase except in the late afternoon when we never observed a peak of OCS associated with the peaks of O3 and CO (Figs. 3a and 2d).

Figure 4 compares the mean diel patterns in ambient OCS mixing ratios at 2 m height in June 2012 and June 2013, constructed from data presented in Figs. 2c and 3b, respectively. Data show that the OCS concentrations were more stable at night than during the day since a drop of ∼ 50 ppt was observed in the early morning hours, down to ∼ 450 ppt, followed by a rise up to ∼ 520 ppt in June 2012 and ∼ 650 ppt in June 2013. These huge diurnal variations, with amplitudes in the range of 150–250 ppt (Fig. 3b), were confirmed by independent measurements carried out with the LGR CO2–OCS–CO–H2O analyzer, which was running in parallel (Fig. 3b). The concomitant decrease of OCS and O3 in the early morning hours was confirmed in the 2013 records (Fig. 3b). Furthermore, the air masses richest in O3, which were transported over O3HP by strong winds in the late afternoon, were not the richest in OCS throughout the campaign (Fig. 3b). Our ground-based meteorological and ozone observations from June 2012 and June 2013 (650 MSL), presented in Figs. 2 and 3, are highly consistent with data reported by Kalthoff et al. (2005).

Diel variations in near-surface OCS and CO2 vertical gradients were documented twice in June 2012 from data collected alternatingly at 2 and 10 m (Fig. 5). Both time series show no apparent OCS gradient at night, whereas CO2 data showed strong vertical gradients with CO2 at 2 m being higher by approximately 5 ppm than at 10 m. During the day, the CO2 gradient reversed, with CO2 mixing ratios that are lower at 2 m than at 10 m and a back-reversal of the CO2 gradient occurring in the late afternoon at 17:00–18:00 UTC. During the day, OCS mixing ratios were systematically lower at 2 m than at 10 m by a few ppt in the morning and up to 10–20 ppt in the afternoon. Hence, CO2 and OCS were consistently lower at 2 m than at 10 m during the day. At night however, CO2 had a gradient in line with the respiratory production of CO2, whereas OCS showed no measurable gradient.

It should be noted here that the CO2 and water fluxes are not strictly linked at the ecosystem level because the non-foliar contribution is different for CO2 (non-green plant biomass, and soil respiration) and H2O (evaporation from soil and tree surfaces). Furthermore, the gas exchange between the substomatal cavity and the atmosphere has drivers that impact biological and physical processes differently (e.g., the temperature effect on photosynthesis and respiration for CO2 and transpiration for water). However, it is known that soil water content will impact litter decomposition processes and other microbial and rooting activity that determine soil respiration. The presence of a non-stomatal water flux is an indication of the wetness of upper soil layers; hence, it is a proxy of increased respiration rate. Negative water fluxes at dew point temperature indicate dew formation that may cause non-stomatal fluxes due to the dissolution of gases. The latent heat and CO2 fluxes (GPP and NEE) followed a clear diurnal cycle well correlated with global radiation, indicating that there was no significant water stress that would tend to lower the flux in the afternoon (Fig. 6a, b). However, the latent heat flux was significantly higher on 13 June than for later days (Fig. 6a). Higher water fluxes were also measured on 11 and 12 June, which were likely due to the evaporation of precipitation of low intensity (2 mm at the most) that occurred on 10, 11, and 12 June as well as the water that was deposited as dew the nights of 11 and 12 June, which was clearly shown by the air temperature reaching the dew point temperature and the sensible heat flux being highly negative at night (data not shown). The stomatal conductance for water vapor also followed a clear diurnal cycle (Fig. S2). Significant positive isoprene fluxes were only observed during the daytime, following diel cycles with midday maxima ranging from 10 to 35 nmol m-2 s-1 (Fig. 6c redrawn from Kalogridis et al., 2014).

Unfortunately, the fast-O3 sensor that was used to assess the O3 deposition velocity had some sporadic down times that occurred frequently during the 12 to 18 June sampling period. During that period, the analyzer only performed well during one night. Good-quality data, however, were recorded continuously from 29 May to 3 June and from 7 to 9 June (Fig. 7). Stomatal conductance for O3 (gsO3) assessed with the method of Lamaud et al. (2009) followed diel cycles with midday maxima throughout the whole month of June 2012 in the range 6 to 8 mm s-1 (data not shown but Fig. 7 provides an illustration of the typical diel pattern of gsO3) for late May and the first week of June 2012. The shape of these diel cycles provides another indication that the canopy was never under water stress and the gsO3 is mostly light driven. The ozone deposition velocity (VdO3) exhibited diurnal variations with generally larger deposition before midday (Fig. 7a). Since the stomatal conductance showed a much more symmetrical feature during daytime (Fig. 7b), it indicates that non-stomatal ozone deposition occurred preferentially during the morning. However, estimates of gnsO3 were less numerous in the afternoon than in the morning because of inconsistencies between gcO3 and gsO3 values noticed during the afternoons of 29–31 May and 9 June, where gsO3 was higher than gcO3 (Fig. 7b). Nevertheless, in five cases out of six, a peak in gnsO3 was observed during the period between 29 May and 3 June. Data show a shift in the relative importance of both pathways since from 7 June the ozone deposition in the morning in all cases was predominantly through the stomatal pathway. Unfortunately, we have no indication about ozone deposition pathways during the periods where OCS was monitored in the atmosphere. However, the shift towards higher O3 deposition through the stomatal pathway during the second week of June (Fig. 7b) and the strong similarities between OCS and O3 diurnal patterns in June 2012 (Fig. 3a) suggest that the non-stomatal pathway lost importance throughout the month of June.

OCS diel variations presented here (Fig. 3) resemble those reported by Berkelhammer et al. (2014) at two sites in central North America where steep rises in OCS also occurred after sunrise (see their Figs. 7b and S11). The authors suggested that this morning rise was related to boundary layer dynamics when air from above, richer in OCS than the air from the nocturnal boundary layer, was entrained downwards. This is also the case at O3HP as shown in the vertical profiles of water vapor (Fig. S3). Entrainment of dry air from the nocturnal boundary layer is evidenced by the decrease in water vapor concentrations about 2 h after sunrise. This decrease is generally more important at 10 m than at 2 m. However, diurnal variations with amplitudes over 200 ppt as observed at the O3HP in June 2013 were never reported before. This raises the question of the origin of air masses rich in OCS advected over O3HP in mid-June 2013. It is highly unlikely that long-range transport of biomass burning gases and aerosols between North America and the Mediterranean region was responsible for OCS contamination because the transport of biomass burning material occurred in late June 2013 after the end of our OCS surveys (see Fig. 4 in Ancellet et al., 2016). Since the O3-rich air masses that reach the O3HP in the late afternoon lag behind those rich in OCS by ∼ 4 h (Fig. 3b), it is clear that the OCS and O3 peaks have distinct origins. Backward trajectories at 300 m above ground level ending at 12:00 UTC (Stein et al., 2015), when OCS levels at the O3HP in June 2013 were over 600 ppt (Fig. 3b), show that the circulation of the air masses during the 2012 and 2013 periods was at low altitude (below about 500 m a.g.l., i.e., below 1100 m a.s.l.); thus, they were generally in the boundary layer. The back trajectories show that the air masses were in closer contact with the continent in June 2013 than in June 2012 and that the transport in June 2013 was from the N–NW along the Rhône Valley (Fig. S4). South of the city of Lyon, the Rhône Valley is highly industrialized, and it is therefore likely that the O3HP site is impacted by anthropogenic direct or indirect emissions of OCS (i.e., from the oxidation of CS2 since the largest production of CS2 in western Europe is located in the Rhône Valley; Campbell et al., 2015). Polluted air very likely propagates southwards in the upper layers within the nocturnal jet, which was observed in the sodar measurements performed nearby at Cadarache (Kalthoff et al., 2005), and is entrained downwards in the morning when turbulence recovers. Moreover, we can also demonstrate that the source of OCS pollution is persistently from the same direction using data gathered in Fig. S5, which show the full June 2013 OCS record, starting from 8 June, and the corresponding back trajectories. It is clear that there is no sign of pollution in OCS when air masses, advected from the Mediterranean Sea, reach the OHP site at noon, 300 m a.g.l. Finally, Fig. S6 demonstrates that advection of pollutants from the combustion of fossil fuels (and from biomass burning, see above) is unlikely in the OHP area except for on the night of 15 June when CO levels went up to 250 ppb. A CO pollution event was also recorded the next morning but data show no impact on OCS levels. In the afternoon, polluted air from the metropolitan area of Marseille is transported by the sea breeze thus leading to an increase of ozone at elevated layers above the convective boundary layer as demonstrated in the air circulation study of Kalthoff et al. (2005). The highest ozone concentrations above 100 ppb can be found about 50 km further downwind north and northeast of Marseille both in the mountainous areas of Luberon and above (Kalthoff et al., 2005; see Fig. 6 of that paper). We can therefore conclude that the photosmog of the city of Marseille is not a source of OCS.

At the O3HP, OCS concentration gradients showing lower concentrations at 2 m than at 10 m were observed during the daytime (Fig. 5), especially during the afternoon so when turbulent mixing was strongest (Fig. 1b). Gradients were nonexistent at night. This implies that the forest ecosystem was essentially a net sink of OCS. Measured CO2 vertical gradients indicate that the forest ecosystem was a net sink of CO2 during the daytime and a net source at night, features that were confirmed by the eddy covariance data showing NEE to range between -15 and -20 µ mol m-2 s-1 around midday and 0–5 µ mol m-2 s-1 at night (Fig. 6). However, the sharp rise in OCS concentrations between 06:00 and 12:00 UTC (Fig. 2) and the reversal of the CO2 gradients at 17:00–18:00 UTC (Fig. 5) reduce the time window to a few hours in the afternoon where the ecosystem relative uptake of OCS (ERU), which is the ratio of the relative vertical gradients of OCS and CO2, can be assessed. ERU is an important parameter since it is proportional to GPP and NEE scaled by the ratio of relative leaf exchange rates (LRU) following Eq. (2). Therefore, we anticipate that this approach to partition measured NEE will hardly be applicable at O3HP, not only because the amplitude of the diurnal variations in LRU is unknown at O3HP but also because vertical gradients of OCS cannot be calculated from measurements carried out throughout the whole period of illumination. In 2012, only data collected in the afternoon were exploitable and the mean OCS / CO2 ratio at 2 m height was 1.33 ± 0.02 ppt ppm-1, n=27. In June 2013, polluted air masses produced dramatic variation in atmospheric OCS / CO2 ratios in the morning and the afternoon, leaving no time window for ERU calculation. These air masses were not related to urban photosmog episodes since there was a gap of ∼ 4 h between the peaks of OCS (up to 700 ppt) and O3 (up to 85 ppb). With these caveats in mind, the ratio of the mean relative vertical gradients of OCS and CO2 (calculated from linear OCS profiles) was equal to 4.7 and 4.3 for the afternoons of 6 and 17 June, respectively. However, it had a large relative error (≥ 50 %) and was consistent with ERUs reported by Blonquist et al. (2011) at the Harvard Forest AmeriFlux site in summer–autumn 2006 (5.7 ± 1.2 (1 SD) for short-term ERU values calculated from linear OCS profiles as we did at the O3HP).

Only when the plant uptake is the dominant flux, is the ERU proportional to the ratio of GPP / NEE, with a proportionality constant that is the LRU (Campbell et al., 2008). As discussed above, this is only the case at the O3HP site for a few hours in the afternoon (because at other moments the ecosystem is not the main driver but rather part of the boundary layer dynamics) and that ERU could only be calculated using the OCS and CO2 gradients for these few hours. When ERUs and the mean NEE / GPP ratio calculated for the period 12:00–17:00 UTC (0.78 ± 0.05, n=20) are used in Eq. (2), LRUs at the O3HP are equal to 3.7 and 3.4. These values fall in the upper range of LRUs obtained from leaf chamber studies over a large range of light conditions and tree species (1–4, Stimler et al., 2010; 1.3–2.3, Berkelhammer et al., 2014).

Our OCS measurements were carried out during the period of maximum gross primary productivity of Mediterranean oak forests (Allard et al., 2008; Maselli et al., 2014). At the O3HP, the maximum of Q. pubescens net photosynthetic assimilation also occurs in June (Genard-Zielinski et al., 2015). The O3HP site appears to be ideal for the use of OCS uptake by plants as a tracer for GPP in a Mediterranean oak forest because the soil is neither a source nor a sink of OCS when GPP fluxes culminate. The lack of net uptake of OCS at night is a specific feature of the O3HP site that is not shared by other open oak woodlands characterized by a Mediterranean climate (Kuhn et al., 1999; Sun et al., 2015). The study of Kuhn et al. (1999) was performed in June 1994 at the Hastings Natural History Reservation in Monterey County, central coastal California (490 m a.s.l.), which is located in a side valley of the Carmel Valley, approximately 40 km from the coast. These authors reported a nocturnal drop in the OCS ambient mixing ratio by about 150 ppt corresponding to a nocturnal OCS deposition rate of up to -7.6 pmol m-2 s-1, which was estimated by a nocturnal boundary layer depletion model. The range of fluxes reported by Kuhn et al. (1999) is consistent with those measured using soil chambers at Stunt Ranch in southern California in April 2013 (0.1 to -6.5 pmol m-2 s-1; Sun et al., 2015). OCS fluxes at Stunt Ranch exhibited clear diurnal variations with higher uptakes during the night than during the day (Sun et al., 2015). Unfortunately, the signature of these fluxes in the nocturnal boundary layer in terms of nocturnal drop in OCS mixing ratio was not reported in that paper. To give an illustration of what might be the atmospheric signature during stable nocturnal conditions of OCS uptake events of such intensity, we extracted data from a set of observations where the role that soil, leaf, and atmospheric dynamics have on surface OCS exchange is investigated from OCS diurnal cycles (as at O3HP) and nocturnal fluxes calculated using the radon-tracer method (Belviso et al., 2013). Figure S7 shows an 8-day time series of ambient mixing ratios of OCS, CO2, CO, and O3 carried out in mid-April 2015 (after bud break and almost complete leaf expansion) in a suburban area of the Saclay Plateau (Paris region) in relation to incoming global radiation, thermal stratification, and wind speed (as at the O3HP). Periods of low atmospheric turbulence over the Saclay Plateau were evaluated using 222Rn accumulations. In April 2015, hourly variations show nighttime and early morning decreases of OCS mixing ratios (Fig. S7c) and corresponding 222Rn increases (Fig. S7b). The amplitude of OCS diurnal variations is in the 40–80 ppt range. OCS minima coincide with calm meteorological conditions with wind velocities lower than 6 km h-1 (Fig. S7b), which are favorable to thermal stratification (Fig. S7a), with CO2 maxima sometime up to ∼ 480 ppm (Fig. S7c) and with O3 minima down to a few ppb (Fig. S7d). However, it is worth noting here that the amplitude of CO2 and O3 nocturnal variations over the Saclay Plateau in early spring are higher than those at O3HP due to anthropogenic emissions of CO2, which can be traced using CO mixing ratios (Fig. S7d), and to NOx emissions, which accelerate the chemical removal of O3 (O3 reacts with NO, data not shown). OCS fluxes calculated using the radon-tracer method during stable nocturnal conditions ranged from -4.8 pmol m-2 s-1 (night of 14 April) to -14.2 pmol m-2 s-1 (night of 11 April, Fig. S7c). They fall in the upper range of fluxes reported by Kuhn et al. (1999) and Sun et al. (2015), but the comparison should be made with caution because three different methods were used to estimate the OCS fluxes (i.e., a boundary layer model, soil chambers, and the radon-tracer method). Qualitatively, it is clear that uptake rates of several pmol m-2 s-1 lead to drops in the OCS ambient mixing ratio by several tens of ppt during periods of low atmospheric turbulence. Hence, a major difference between these woodlands and the O3HP site during springtime is that soil of the Mediterranean forest ecosystem of southern France is not a net sink of OCS. Soil OCS uptake has been shown to be dependent on soil physical properties like soil structure, water content, water-filled pore space, and temperature (Van Diest and Kesselmeier, 2008; Ogée et al., 2016) but also on soil biological properties like microbial activity (Kato et al., 2008; Ogawa et al., 2013), active root density (Maseyk et al., 2014), or the presence of a litter layer (Berkelhammer et al., 2014; Sun et al., 2015). Away from a range of optimum uptake, which varies between soils, changes in soil water content and temperature can markedly reduce OCS uptake by soils (Van Diest and Kesselmeier, 2008). However, the soil temperature and water content at the O3HP (Fig. 1c, d) are typically in the range of optimum uptake published by Van Diest and Kesselmeier (2008). A limitation of OCS uptake by soils due to a poor OCS diffusion is also unlikely considering that the soils from the O3HP are strongly structured and are far from being water saturated. Finally, the only physical property of soil differing among the three open oak woodlands is the soil texture, with a fine clayey texture at the O3HP but a coarse sandy loam texture at Hastings Reservation (Kuhn et al., 1999) and at stunt Ranch (Sun et al., 2015). OCS uptake by fine-textured soils has already been reported (Maseyk et al., 2014), this result pointed out the need for measurements of OCS uptake for a greater diversity of soils. Concerning the biological soil properties, the soil at the O3HP is covered by a relatively thick litter layer that may induce a change from OCS uptake to OCS emission (Berkelhammer et al., 2014). However, at Stunt Ranch Sun et al. (2015) measured that the litter was responsible for OCS uptake. The surface horizons at the O3HP showed organic carbon content ranging from 167 to 43 g kg-1 in the surface soil horizons (Table 1) but only 24 g kg-1 at Hastings Reservation (no data on soil organic carbon are available for Stunt Ranch). Being richer in organic carbon, soils at the O3HP show very likely higher microbial activity, a factor that should stimulate uptake of OCS by soils but apparently does not. If the capacity of soils to consume OCS is more related to specific enzymatic activities (carbonic anhydrase (CA) and OCS hydrolase) than to the general variables presented above, our observations would highlight deficiencies in these enzymatic activities in the calcium-carbonate-rich soils of O3HP. However, this hypothesis is not consistent with the suggestion that CA performs an essential role in microbial organisms surviving periods of osmotic stress such as drought at the surface of Mediterranean soils (Wingate et al., 2008). Finally, as roots and associated rhizosphere have been found to produce OCS, a greater abundance of roots in the surface soils at O3HP by comparison with the two other oak woodlands may explain why the soils at O3HP are not a sink of OCS. In other words, the lack of nocturnal net uptake of OCS would indicate that gross consumption of this gas in soil is compensated for by emission processes that remain to be characterized. However, no data on root abundance are available at Hastings Reservation or Stunt Ranch to confirm such a hypothesis.

Data show strong similarities during the night and early morning hours between OCS and O3 diel variations at the O3HP suggesting a similar sink during that period (Fig. 3). At the O3HP, volatile organic compounds (VOCs) produced by the vegetation are essentially in the form of isoprene (Kalogridis et al., 2014; Genard-Zielinski et al., 2015). Isoprene is oxidized in the atmosphere by the hydroxyl radical (OH), O3, and the nitrate radical (NO3), but in-canopy chemical oxidation of isoprene at the O3HP was found to be weak and did not seem to have a significant impact on isoprene concentrations and fluxes above the canopy (Kalogridis et al., 2014). Hence, ozone deposition at the O3HP was essentially through leaf uptake via stomata and surface deposition, without a strong contribution from chemical reactions. In late May and early June 2012, the non-stomatal contribution to the ozone flux was in general markedly higher than the stomatal one in the morning hours (before 10:00 UTC), but it became much less significant in the afternoon (Fig. 7b). Conversely, during the second week of June, although there were still signs of non-stomatal loss of ozone in the morning, the major contribution to ozone deposition was through the stomatal pathway (Fig. 7b). The analogy with OCS at nighttime and in the early morning suggests that soil did not contribute much to the O3 flux and that the deposition flux of O3 in mid-June was essentially the result of leaf uptake. However, it is difficult to evaluate the soil ozone pathways without turbulence measurements inside the canopy. It would be worth looking further into how OCS could be used to partition ozone fluxes near the ground between soil and leaf deposition processes. The applicability of OCS to characterize the strength of ozone sinks would be reduced in situations where NOx would significantly impact the chemical production or destruction of ozone in the canopy or when background air is contaminated by primary or secondary anthropogenic sources of OCS (Fig. 3b).