The main constituents of volcanic plumes are H2O, CO2 and sulfur gases (dominated by SO2, H2S). Apart from these species, volcanoes also emit a certain amount of halogen species which are mainly released in the rather unreactive form of hydrogen halides such as HCl, HF, HBr, HI (e.g. ; ; ) and which are largely dominated by the chlorine emissions (HCl). reviewed past measurements (∼ 1980–2008) of arc-related volcanic halogen emissions around the globe and found that HCl emissions contribute most with an estimated flux of 4.3 (±1) Tga-1. HBr and HI emissions are orders of magnitude smaller with fluxes of 5–15 and 0.5–2 Gga-1 respectively.

In the case of Mt. Etna, SO2 / HCl ratios between 2 and 7 were found in past measurements (e.g. ; ; ). SO2 appears to dominate the total sulfur emissions of Etna with SO2 / H2S ratios of the order of 101–103 (; ).

A certain amount of the emitted hydrogen halides is converted into RHS, whereas the conversion from HBr into BrO appears to be much more efficient than the analogous reactions for volcanic chlorine. A key question related to these conversion mechanisms is the production of the halogen radicals (i.e. Br, Cl, I) in the plume. Once these are provided, oxidised halogens such as BrO and ClO are formed in reaction with ozone (O3).

A certain amount of RHS (e.g. Cl, Br) can be produced in the hot initial plume via high-temperature oxidative dissociation processes as suggested by model studies (e.g. ; ). Furthermore, Br can be formed via the reaction of HBr with OH in the very young plume (). However, the corresponding amounts are by far too small to explain the BrO amounts observed. In fact, the largest part of BrO is formed in atmospheric reactions, including the photolysis of Br2 and BrCl (e.g. ; ; ). This is further supported by direct observations showing a strong increase of the BrO levels in the young plume (e.g. ; ) and the virtual absence during night-time (). Nowadays, the underlying chemical reaction processes of the BrO formation in the young plume are mostly understood and likely driven by a heterogeneous and partly auto-catalytic reaction mechanism often referred to as “bromine explosion” (e.g. ; ), which includes the following reactions (note that the subscript “aq” denotes species in the aqueous phase on particles). BrO+HO2→HOBrgas+O2 HOBrgas→HOBraq HBrgas→HBraq→Braq-+Haq+ HOBraq+Braq-+Haq+→Br2, gas+H2O Br2, gas+hν→2Br 2Br+2O3→2BrO+2O2

The “bromine explosion” encompasses the uptake of hypobromous acid (HOBr) from the gas into the aqueous phase. After the reaction of HOBr with bromide, Br2 is released into the gas phase where it is rapidly photolysed, forming BrO in reaction with O3. Once formed, the self-reaction of BrO induces a catalytic destruction of O3. Noteworthy in this context are the similarities to observations of bromine emissions in polar regions (e.g. ; ). Measurements performed at Mt. Etna and Stromboli volcano (Aeolian islands, Italy) indicate that up to 11 % of the total emitted bromine is converted into BrO already within the first 5 min downwind ().

Potential formation processes of reactive chlorine species from the emitted HCl are still little studied. Apparently, the activation of chlorine is much weaker compared to bromine. This is indicated by the comparatively low Cl and ClOy abundances we found (relative to the BrO ratios), indicating that less than 1 % of the emitted HCl is converted into reactive chlorine (ClOy) in the Etna plume. In other words, ClOy / HCl is much smaller than BrO / HBr. In our opinion this phenomenon is mainly due to the fact that Br oxidation (conversion of bromide to Br, BrO) is a self-amplifying process (the bromine explosion) while Cl oxidation has no such properties. The reason why Br “explodes” but Cl does not is due to the relatively fast reaction of Cl atoms with CH4 (; ). Moreover, the dissolved chloride ions are less reactive compared to bromide ions (see Reaction ) (). Thus, Cl release is rather likely to be a by-product of the bromine explosion via formation of BrCl in the reaction of HOBr with chloride. However, the efficiency of this chlorine release channel strongly depends on the Cl- / Br- ion ratio in the condensed phase (). A significant release of BrCl is only likely for Cl- / Br- ratios exceeding 104; for instance a Cl-/Br- ratio of 2× 104 would yield a release of 50 % BrCl and 50 % Br2. Direct sampling measurements at Mt. Etna revealed Cl-/Br- ratios of the order of 102 () up to 103 (). For these values the HOBr uptake yields a release of more than 90 % Br2. Note that this favoured Br2 release is probably even enhanced in volcanic plumes due to the acid environment (low pH in the aerosol, for details see , i.e. the pH dependency of the discussed mechanisms). To our knowledge, there are no measurements indicating Cl-/Br- ratios of the order of 104 or larger at Mt. Etna. Hence, a significant BrCl release due to the HOBr uptake is relatively unlikely in the case of Mt. Etna.

Measurements of reactive chlorine species in volcanic plumes are still rare. Thus, the underlying chemistry is only partly understood and still bears large uncertainties, especially regarding the question of possible abundances of Cl atoms in the gaseous phase and the corresponding release mechanisms. However, once Cl atoms are provided, ClO is formed in the reaction with O3 and OClO is then formed in the reaction of ClO with BrO:

BrO+ClO→OClO+Br.

The corresponding reaction rate coefficient is k7=6× 10-12cm3s-1 (at 298 K; ). Further possible reaction channels for the OClO formation are orders of magnitude slower (e.g. ClO+O3, ClO+ClO; ) and were not considered within this study. The main daytime sink of OClO is its photolysis:

OClO+hν⟶JClO+OJ≈6× 10-2s-1.

Both and detected OClO in the plume of Mt. Etna. The corresponding OClO/SO2 ratios were between 3 and 6× 10-5 (for spectra related to the plume centre). Simultaneous BrO measurements indicate an OClO/BrO ratio of approximately 0.25 for Mt. Etna in both studies. Further detections of volcanic OClO are reduced to satellite measurements (Puyehue-Cordón Caulle volcano, Chile) after an eruption in 2011 () and most recently the detection of OClO in the plume of Soufriére Hills volcano (Montserrat) during a hiatus in 2011 (). In the latter study, comparatively large OClO/SO2 ratios (4–6 × 10-4) are reported as well as large OClO/BrO ratios showing values up to 5 (i.e. about 20 times larger compared to Mt. Etna).

A key parameter for the OClO formation in volcanic plumes is the prevailing availability of ClO and BrO molecules. Previous studies reported relatively large amounts of volcanic ClO measured with passive DOAS instruments (; ). The corresponding ClO/SO2 ratios were of the order of 5 % hence, almost 3 orders of magnitude more ClO than OClO. However, these measurements have to be treated cautiously due to difficulties and uncertainties in the DOAS evaluation of ClO.

Furthermore, to our knowledge it has not yet been possible to reproduce these measurements in model studies (e.g. ; ). investigated ClO and OClO abundances at the vent of Masaya Volcano (Nicaragua) using an active long-path DOAS instrument. They did not detect any of both species most likely due to the proximity of the measurement to the crater (i.e. early stage of the RHS formation). In addition, the halogen content of Masaya volcano is probably smaller compared to Mt. Etna (). While the OClO / ClO ratios should typically be of the same order of magnitude in the case of Mt. Etna and Masaya (; this study), this seems not to be the case for the Puyehue–Cordón Caulle eruption in 2011 which actually indicates a large excess of ClO compared to OClO and even BrO ().

The focus of this article is the temporal and spatial evolution of RHS in volcanic plumes (especially BrO, ClO, OClO) and potential impacts on the atmosphere in the vicinity of volcanic plumes. In particular, we use MAX-DOAS data to study the formation of BrO and OClO in the young plume in great detail and to infer typical formation times of these species (for the conditions at Mt. Etna in September 2012). We furthermore estimate mean plume abundances of BrO, ClO and OClO. These results are used to derive Cl-atom concentrations in the plume in order to address the question of a potential – chlorine-induced – depletion of atmospheric methane (CH4) in the plume environment.

The MAX-DOAS instrument used in this study analyses the solar spectrum in the ultraviolet (UV) and the visible (VIS) ranges using two spectrographs (UV: Avantes AVA AvaBench-75-ULS2048x64; VIS: Avantes AVA AvaBench-75-ULS2048L) covering a spectral range of 292–578 nm (UV: 292.1–456.1 nm; VIS: 434.7–577.8 nm). Scattered sunlight was collected using a small telescope consisting of a quartz lens (f=100mm) which focuses incoming light onto an optical fibre bundle. The latter consists of seven individual fibres each with a diameter of d=100µm. Six of these were coupled into the UV spectrograph, while the seventh fibre was connected to the VIS spectrograph. The measured spectral resolution of both spectrographs was ΔλUV=0.51 nm and ΔλVIS=0.39 nm. A SCHOTT BG-3 filter was placed behind the entrance slit of the UV spectrograph to reduce stray light. The telescope was focused such that both spectrographs have approximately the same field of view (UV: 0.15∘, VIS: 0.16∘, full aperture angle). The optical benches of the spectrographs were thermally insulated and temperature stabilised using a Peltier element controlled by a Supercool PR-59 temperature controller. During the whole measurement campaign, both spectrographs were stabilised to a temperature of Tmeas=10∘C. The air-tight instrument box was mounted onto a tripod. Two motors (azimuth and elevation) allowed to control the viewing direction of the telescope, geo-locations were recorded using a GPS receiver. All hardware elements were remotely controlled using an embedded PC. The software MS-DOAS was used for data acquisition. MS-DOAS was developed by U. Frieß at the Institute of Environmental Physics in Heidelberg and is designed to control standard hardware components used in DOAS instruments (e.g. spectrographs, motors, temperature controller, GPS receiver). Furthermore, it includes a scripting feature making it easily possible to automatise measurement and scanning routines.

Mt. Etna is the largest and most active volcano in Europe and is situated in the eastern part of Sicily, an island south of the Italian mainland. The activity of Mt. Etna shows a distinct variability, including quiescent degassing periods as well as eruptive periods. During the measurement campaign in September 2012, Etna showed a stable quiescently degassing behaviour from the four active craters – north-east (NE), Bocca Nuova (BN), Voragine (VOR) and south-east (SE) crater – which are located in the summit region at an altitude of approximately 3300 ma.s.l. The first 3 days of the campaign (11–13 September 2012) took place at the Etna observatory (Pizzi Deneri) which is located approximately 2.5 km north-east of the active summit at an altitude of 2800 m. Figure  shows a photo of the volcano and the emission plumes from the different craters. The photo was taken from the observatory and shows the NE crater (right) and the SE crater (faint in the background) on 13 September at 07:24 UTC. The plume was slightly condensed (see Fig. ) during most of the measurements performed in September 2012 and showed no visible indications of any ash emissions. In Fig. , all measurement locations of the campaign (11–26 September 2012) are indicated. One of the main objectives of this study was to investigate the temporal evolution of oxidised halogens in the volcanic plume. Therefore, the measurements were performed at different locations in order to cover a large variety of different plume ages in the spectra.

Three different plume scanning routines (“scans”) were performed in order to study the chemical variability of the measured species in the volcanic plume (see sketch in Fig. ). One “scan” typically consists of a set of plume spectra plus a subsequently recorded solar reference spectrum with the telescope pointing into a volcanic gas free atmosphere (for details see Sect. ).

The spectra were analysed using the software package DOASIS (v. 3.2.4422, ). Details on the scanning routines can be found in Sect. . In order to improve the detection sensitivity, several hundred up to 1500 individual spectra were co-added for the DOAS analysis. A standard DOAS fit (see ) was performed for the UV and VIS spectra in order to retrieve slant column densities (SCDs) of the chemical species in the plume (in this study mainly: OClO, BrO, SO2, IO, OBrO, OIO). A Fraunhofer reference spectrum (FRS, I0(λ)) was included in the fitting routines to account for solar absorption lines in the spectra (Fraunhofer lines) and atmospheric background absorption. The FRS was recorded with the telescope pointing in the direction of a volcanic gas-free atmosphere and close in time to the corresponding plume spectra (usually subsequently to each scan). The latter is important to keep potential additional stratospheric signals at a minimum (for details see Sect. ). Each potential FRS was pre-evaluated regarding its SO2 content using a literature solar background spectrum as FRS () which was convolved with the instrumental line spread function (LSF). Only FRS candidates showing SO2 SCDs (SSO2) smaller than SSO2,FRS<5× 1016moleculescm-2 were used as FRS.

In the following, the implemented steps to retrieve the SCDs from the raw spectra are described. Further details to individual topics regarding the data evaluation can be found in Appendix .

Prior to the DOAS evaluation, all FRS and plume spectra were corrected for electronic offset and dark current. Two ring spectra (R,R4) were included into the fitting routine to account for inelastic scattering effects (Raman scattering) in the atmosphere (see e.g. ). The first ring spectrum (R) was calculated in the usual way from the respective FRS using the function of the evaluation software DOASIS (). The second ring spectrum (R4) was determined following the suggestions from (for details see Appendix ). Improvements due to the R4 correction are discussed in Sect. , and a fit example with a strong R4 signal is shown in Fig. . Literature cross sections of the individual absorbers (σi, see Table ) were convolved with the LSF of the respective spectrograph. During the convolution, the σi were corrected for the solar I0 effect and for spectral saturation () using the corresponding functions in DOASIS. The latter was performed assuming typical SCDs for the respective species (e.g. SSO2=2×1018moleculescm-2). In order to correct for any misalignment of the spectrograph, a slight shift (±0.1 nm) and squeeze (±5 %) was allowed for all fitted species (i.e. FRS, R, R4, σi). Shift and squeeze of all σi were linked to the strongest absorber and the two ring spectra were linked to the corresponding FRS in order to minimise the degrees of freedom during the fit-process. A third-order polynomial was included in the fitting routine to remove broad band extinction. An additional zero-order polynomial residing in intensity space was included (also referred to as offset polynomial) to account for intensity offsets in the spectra (e.g. due to stray light, for details see ; ). The measurement uncertainty (δmeas) was estimated conservatively by multiplying the retrieved fit errors (δfit) with a factor of U=4 to account for potential abundances of fit residuals structures (, see e.g. Fig. ). In the case of good fit results (which were assessed by the peak-to-peak values of the fit residuals Δ‾res) the correction factor was reduced down to U=3 (i.e. for Δ‾res≤ 1.2×10-3, see e.g. Fig. ). Details regarding the error treatment are discussed in Sect. . The detection limits of the SCDs were defined to be twice the measurement uncertainty (2×δmeas), thus representing a detection certainty of 95 %.

The focus of this study was regarding the data collected with the UV spectrograph (i.e. the DOAS evaluation of OClO, BrO, SO2, IO). In order to find the optimum evaluation range for each species, detailed sensitivity studies were performed including DOAS fit contour plots (“retrieval wavelength mapping”, for details see ). For the VIS data (i.e. the OBrO and OIO evaluation) these sensitivity studies were not performed since these data were of secondary interest. Therefore, we used a fixed correction factor of U=5 for the estimation of the corresponding measurement uncertainties of OBrO and OIO (for details see Sect. ). All evaluation routines used in this study are summarised in Table , including the corresponding wavelength ranges, additional absorbers and the used correction factors U for the DOAS fit errors. An overview of all corresponding literature cross sections (σi) used is given in Table . Note that in the case of O4, two different literature cross sections were used because we found that different cross sections for the UV and VIS spectral ranges give the best results.

The data from plume cross-section scans were used to estimate mean concentrations (c¯i) of BrO and OClO in the plume. This was done assuming a circular plume cross section and straight line absorption light paths through the plume. Any potential deviations due to radiative transfer effects (RTE, e.g. multiple scattering, light dilution; for details see e.g. ; ) or deviations from the assumed circular shape were not considered in this estimation. The plume diameter (Øpl) was estimated from the angular extend of the SO2 SCD profile and the distance to the plume (see also e.g. ; ). The corresponding SO2 SCDs were used as a proxy for the lengths of the absorption light paths (leff,i) in the plume, whereas the largest SO2 SCD of the scan was assigned to Øpl. Based on that, the leff,i could be estimated for all scan spectra i: leff,i=Øpl/SSO2,max×SSO2,i. Using this, the mean concentrations (c¯j,i) of the measured species j (e.g. OClO, BrO) were estimated: c¯ji=Sj,i/leff,i. The corresponding uncertainties were determined from the DOAS-fit errors and the uncertainties in the estimation of the plume diameter using Gaussian error propagation.

Following , we estimated ClO concentrations from the retrieved BrO and OClO-SCDs assuming steady state between the formation of OClO (Reaction ) and its photolytic destruction (Reaction ): ClO=JOClOk7×[OClO][BrO]≈JOClOk7×SOClOSBrO.

Since BrO and OClO were evaluated in the same wavelength range, differences in the retrieved SCDs (Si) due to differences in the radiative transfer can be neglected. We therefore assume that the ratio of the OClO and BrO concentrations is approximately the same as the ratio of the respective SCDs (Eq. ). The OClO photolysis frequencies JOClO used for the calculation of the ClO concentrations were simulated for our data set by E. Jäkel (Leipzig Institute for Meteorology). For the simulation, the 1-D radiative transfer model libRadtran () was used. The photolysis frequencies were determined for a set of chosen spectra from the field campaign and were between 5.1×10-2s-1 (SZA≈62∘) and 7.1×10-2s-1 (SZA≈34∘), slightly slower than typical values found in the stratosphere (e.g. Jstr,OClO≊ 7.6×10-2s-1; ). Uncertainties in the determination of the ClO-concentrations were estimated using Gaussian error propagation.

The plume age (τ) was estimated using meteorological information (i.e. wind speed and direction) and the measurement geometry (i.e. geo-locations of instrument and craters, telescopes viewing direction). A typical measurement geometry at Mt. Etna is sketched in Fig. . The intersection of the telescopes viewing direction with the plume determines the distance l. Based on that, the plume age was estimated as follows: τ=lvwind.

The azimuthal alignment of the instrument was performed using a compass. Due to possible disturbances of the planetary magnetic field by the volcano, we estimated the instruments azimuth-uncertainty to ±3∘ (gray shaded area in Fig. ). Wind directions were estimated using own observations/notes and – on clear days – satellite pictures from the MODIS network (Aqua, Terra satellites). Wind velocities were partly retrieved from simultaneously performed SO2-camera measurements and from own observations. From the 16. September we additionally monitored meteorological data using a meteorological station, which was installed on the southern side of the craters. Uncertainties in the plume age estimation were determined using Gaussian error propagation, a detailed discussion of these, especially relative and absolute errors can be found in Appendix .

Typical vertical column densities (VCDs) of stratospheric BrO are of the order of several 1013moleculescm-2 (e.g. ; ). Therefore, MAX-DOAS measurements of volcanic BrO (using scattered sunlight) can be significantly disturbed by stratospheric BrO signals under certain conditions. Based on the geometrical air-mass factor (AMF: X=1/cos⁡(Θ)) and by assuming a constant stratospheric BrO-VCD of Vstr,BrO=4.0× 1013moleculescm-2 (; ) a correction was implemented to account for additional stratospheric BrO signals in our retrieved SCDs. A detailed discussion including simplifications and sensitivity studies can be found in Appendix .

For our data set, we found that deviations in the volcanic BrO column due to superimposed stratospheric BrO signals are smaller than 5 % in 85 % of the analysed spectra. Only 8 % of the retrieved BrO SCDs showed deviations exceeding the corresponding fit uncertainty. All of these spectra were either recorded before 08:15 or after 16:45 LT (64.6∘<SZA<83.2∘), which shows the importance of this correction especially for measurements performed in the early morning and late afternoon.

In order to study spatial (and temporal) variations of the retrieved halogen species XmOn, molar SO2 ratios of these species were analysed (i.e. XmOn/SO2 ratios), whereas SO2 was treated as volcanic plume proxy due to its comparatively long tropospheric lifetime (e.g. ; ; ). This is a common method to avoid signal variations due to atmospheric dilution effects (e.g. ; ; ; ). Furthermore, compared to the individual SCDs the XmOn/SO2 ratios are much less affected by RTE such as light dilution or multiple scattering (). We therefore neglect any potential influences of these effects on our retrieved trace gas ratios. The discussion of our results mostly relates to the measurements performed at the Etna observatory. Especially for these data, potential influences due to RTE on the retrieved ratios should be negligible because of the proximity to the plume, the relatively high altitude (i.e. low plume dilution; see e.g. ; ) and the fairly good visibility during most of the measurements (i.e. low aerosol scattering; for details see Sect. ). The errors and detection limits of the XmOn/SO2 ratios were calculated from the SCD errors using Gaussian error propagation.

In Fig.  we plotted all retrieved OClO and BrO SCDs as a function of the corresponding SO2 SCDs (A, B) and furthermore OClO vs. BrO (C). Both BrO and OClO show a good correlation to SO2 (Fig. a, b), indicating that these species could only be detected in volcanic plume spectra. Average ratios of 1.65× 10-4 for BrO/SO2 (Fig. a) and 3.17× 10-5 for OClO/SO2 (Fig. b) were found (linear regression). These values are in good agreement with previous findings (e.g. ; ; ). For the linear regression, only significant detections at plume ages exceeding 3 min were considered (blue dots in Fig. ). Measurements at plume ages smaller than 3 min (green dots) were excluded because in this plume age range the formation of BrO and OClO is not yet fully developed and therefore the XmOn/SO2 ratios are smaller (for details see Sect. ). The corresponding average OClO/BrO ratio (at plume ages exceeding 3 min) is 0.16±0.08 and shows a very good correlation between both species in this plume age range (R2=0.9447). Young plume measurements (green dots, τ<3 min), however, rather indicate stronger fluctuations of the OClO/BrO ratio (R2=0.4717).

We investigated the presence of IO (in the UV spectral range), OIO and OBrO (in the VIS spectral range) but did not detect any of these species. Here, we give both the detection limits for plume ages smaller and larger than 3 min (Table ) since it appears reasonable to assume that these species – if abundant in the plume – show a similar plume evolution as was observed in the case of BrO and OClO (for details see Sect. ). For plume ages larger than 3 min, upper limits of 5.2×10-6 (IO / SO2), 2.8×10-5 (OIO / SO2) and 1.8×10-5 (OBrO / SO2) were found. Note that the UV spectrograph showed a better performance (S/N ratio) than the VIS spectrograph, which is indicated by the lower detection limits for IO compared to OIO and OBrO.

Once Cl atoms are produced in a volcanic plume, they will rapidly react either with CH4 or with O3:

Cl+CH4→HCl+CH3, Cl+O3→ClO+O2.

The corresponding reaction rate coefficients are k9=1.0× 10-13cm3s-1 (Reaction  at 298 K) and k10=1.2× 10-11cm3s-1 (Reaction  at 298 K; ). Note that Reaction () is 16 times faster than the OH+CH4 reaction (at 298 K) and has a strong positive temperature dependence (). All other Cl sink reactions are much slower and are therefore neglected here.

Cl atom concentrations in the plume were estimated using the ClO and OClO concentrations inferred from our measurements (see Sect. ) and the corresponding young plume formation times τ0 (see Sect. ). For the estimation we assumed that the total amount of ClOy (i.e. [ClOy]=[ClO]+[OClO]), observed after the levelling of OClO (i.e. at plume age τ0) was produced from Cl atoms via Reaction () (ClO) and Reaction () (OClO). Assuming a linear increase of [ClOy] the corresponding formation rate of Cl atoms was estimated as follows: ddt[Cl]obs≈ddt[ClOy]≈[ClOy](τ=τ0)τ0.

Actually the true rate of Cl atom production d/dt[Cl] is larger since a fraction of the Cl atoms reacts with CH4 (Reaction ) and never shows up as ClOy (possible reaction of Cl with other hydrocarbons is likely to be unimportant and was therefore neglected here): ddt[Cl]≈ddt[Cl]obs×1K with K=[O3]×k10[O3]×k10+[CH4]×k9.

The corresponding Cl-atom concentration is then given by [Cl]=ddt[Cl]×τCl≈ddt[ClOy]×1K×τCl, whereas the lifetime of Cl (τCl) was estimated by τCl=1[O3]×k10+[CH4]×k9.

Introducing the expression for τCl in Eq. () and using Eq. (), an estimate of the Cl-atom concentration can be obtained: [Cl]=ddt[ClOy][O3]×k10≈[ClOy]τ0×[O3]×k10.

Based on this the CH4 lifetime in the plume (due to reaction with Cl, see Reaction ) can be derived: τCH4≈1[Cl]×k9≈τ0×[O3]×k10[ClOy]×k9.

Note that τ0 denotes the formation time of OClO as introduced in Sect. , whereas τCH4 corresponds to the methane lifetime in the plume.

For the estimation of ClOy, we determined mean ClO and OClO concentrations from our retrieval considering only plume ages between 120 and 240 s (see also Sect. ). We retrieved values of ClO‾=2.0×109 cm-3 and OClO‾=3.7×109 cm-3 respectively and hence ClOy‾=5.7×109 cm-3. Based on our findings discussed in Sect.  we estimated the ClOy formation duration to be τ0=142 s. The OClO concentrations (used to estimate ClOy) might carry potential uncertainties due to RTE or non-circular plumes. However, as discussed in Sect. , these deviations should be small (i.e. ≲ 2) for the majority of data points recorded in the young plume and should therefore not significantly influence the outcome of this analysis. The typical tropospheric O3 background is 60–80 ppb for this (relatively polluted) region and altitude (). The expected CH4 lifetime in the plume is directly proportional to the prevailing O3 concentration (Eq. ). Since O3 is most likely depleted in the plume (; ) we determined [Cl] and τCH4 under variation of the O3 concentration (assuming O3 mixing ratios between 1 and 80 ppb). Our results are summarised in Table  and show relatively small Cl-atom concentrations (i.e. 106–108cm-3) and CH4 lifetimes between 14 h and up to 47 days. These lifetimes are more than 2 orders of magnitude shorter than the average atmospheric lifetime of CH4.

However, if O3 is not strongly depleted, CH4 destruction by Cl atoms will probably not lead to a detectable loss of CH4 in the plume since the Cl levels derived from our measurements (at plume age τ=142 s) should decrease rapidly as the plume disperses. Even if these Cl levels would prevail for a few hours downwind, only a small fraction (less than 1 %) of the CH4 would be destroyed (assuming that the mean O3 levels in the plume exceed 10 ppb). However, in regions of very low O3 concentrations (i.e. possibly in the plume centre; ), a significant loss of CH4 could be present but the atmospheric impact would probably still be negligible since the effective volume of this potential methane-depleting environment would be very small. Nevertheless, we want to remark that our calculations are based on the volcanic conditions at Mt. Etna in September 2012, and we therefore want to stress that it is absolutely possible that CH4 depletion may become detectable in plumes of other volcanoes or at different conditions (e.g. due to varying volcanic activity, stronger chlorine emissions, larger Cl-/Br- ratios, low NMHC (non-methane hydrocarbons) concentrations or the presence of volcanic particles favouring the chloride oxidation).