Volcanic emissions of molecular chlorine

Available online at www.sciencedirect.com

Geochimica et Cosmochimica Acta 87 (2012) 210–226 www.elsevier.com/locate/gca

Volcanic emissions of molecular chlorine Michael Zelenski a,⇑, Yuri Taran b,1 a

Institute of Experimental Mineralogy, Russian Academy of Sciences, Chernogolovka 142432, Moscow Region, Russia b Institute of Geophysics, Universidad Nacional Autonoma de Mexico, Mexico D.F. 04510, Mexico Received 22 August 2011; accepted in revised form 27 March 2012; available online 5 April 2012

Abstract Up to 60 ppmv (180 mg/m3) of Cl2 together with 40–80 ppmv HCl were measured in gas emissions from the Tolbachik scoria cones, Kamchatka, which are still hot after the 1975–1976 eruption. Other gas components were atmospheric air (94–99 vol %), water vapour (1–6 vol %) and other acid species (HF, CO2 and H2SO4, total less than 0.1 vol %). Two different processes can account for the existence of Cl2 in the Tolbachik emissions. The catalytic oxidation of volcanic HCl by oxygen is probably the main source of Cl2. Fine crystals of Fe2O3, and oxides and chlorides of other transition metals on the surface of altered basalt can serve as catalysts. The oxidative decomposition of Na, K and Mg chloroferrates formed as a result of basalt acid leaching, can also create high concentrations of molecular chlorine in volcanic gases. The processes described represent a previously unknown case of abiogenic heterogeneous catalysis in nature and examples of gas–rock interactions that affect the composition of volcanic gases. Ó 2012 Elsevier Ltd. All rights reserved.

1. INTRODUCTION Volcanoes emit significant amounts of acid gases, supplying a substantial portion of the sulphur and halogen burden to the atmosphere and serving as an important source of atmospheric acidity (Devine et al., 1984; Symonds et al., 1988, 1994; Fischer, 2008). In most cases, HCl constitutes 0.1–1 mol % of gas emissions from quiescently degassing or erupting volcanoes (Symonds et al., 1994; Gerlach, 2004; Fischer, 2008). Other chlorine species in volcanic gases are usually negligible (Symonds et al., 1988; Gerlach, 1993, 2004). The exception is the possible existence of atomic chlorine in concentrations from 10 ppb to 1 ppm at 800–1200 °C that can be inferred from thermochemical modelling (Gerlach, 2004). When hot volcanic gas mixes and interacts with ambient air, HCl undergoes partial oxidation in such a way that socalled reactive halogen species appear in a gas–air mixture.

⇑ Corresponding author. Tel.: +7 9262555948.

E-mail addresses: [email protected] (M. Zelenski), taran@geofisica.unam.mx (Y. Taran). 1 Tel.: +52 556511720. 0016-7037/$ - see front matter Ó 2012 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.gca.2012.03.034

Thermochemical modelling has shown that atomic chlorine occurs in the initial plumes at concentrations of up to 0.07% of HCl content at 1000 °C (Martin et al., 2006) or even 0.4% at 1200 °C (Gerlach, 2004). Modelling of the gas–air mixture also indicated low amounts of ClO and Cl2 (Symonds et al., 1988; Gerlach, 2004; Martin et al., 2006). The recent development of spectroscopic methods has made direct measurements possible, and small amounts of reactive halogen species (BrO, ClO and OClO) have been detected in volcanic plumes (Bobrowski et al., 2003, 2007; Lee et al., 2005) although the existence of ClO and OClO in the plumes is now disputed (Kern et al., 2008; von Glasow, 2010). As of August 2011, to the authors’ knowledge, molecular chlorine (Cl2) and atomic chlorine (Cl) have not been directly measured either in fumarolic gases or in volcanic plumes. Despite their low concentrations, reactive halogen species are of special interest. They affect the oxidation potential of air (Graedel and Keene, 1995; von Glasow and Crutzen, 2003) and are involved in a multitude of reactions, including ozone depletion/production processes (Platt and Ho¨nninger, 2003; Finley and Saltzman, 2006; Simpson et al., 2007). When a volcanic eruption ceases, cooling volcanic cones, domes and lava flows continue to slowly release residual volatiles, mainly H2O, HCl and HF. Usually, the only

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

effects of such post-eruptive gas discharges are rock alteration and secondary mineralization (for example, Naboko and Glavatskikh, 1984). However, in some cases, emissions from cooling volcanic cones can contain exotic components. Here we present an account of gas emissions with a measurable content of Cl2 from the New Tolbachik scoria cones, Central Kamchatka. Two primary goals of this work were to detail the parameters of Cl2-containing gas emissions and to determine possible sources of Cl2 inside the cooling volcanic cones. The first, semi-quantitative assessment of Cl2 emissions was made in the 2008 field season using gas detector tubes (Zelenski et al., 2008). In 2009– 2010, we studied the Tolbachik emissions using quantitative in situ measurements and gas sampling with subsequent laboratory analyses. Thermochemical modelling and a set of laboratory experiments were performed to reveal the mechanisms of chlorine generation inside the cones. The proposed main source of Cl2 is the oxidation of HCl by oxygen under the catalytic influence of basaltic scoria, namely, fine crystals of Fe2O3 on the surface of the basaltic glass. This paper describes an example of previously unknown abiogenic catalytic process in nature and the influence of solid rocks on the composition of volcanic gas. Most often, volcanic gases are considered as an “isolated phase” after they are released from a magma body, with little or no influence of conduit walls on the gas. Nevertheless, there are a number of examples in which rocks affect gas compositions. The influence of rocks on gas composition has been considered in connection with hydrothermal gases (Giggenbach, 1987; Taran, 1988), the gaseous transport of trace elements (Reed, 1982; Symonds et al., 1992), oxygen fugacity buffering by cooling lava (Gerlach, 1993) and the presence of SiF4 in fumarolic gases (Francis et al., 1996). The case of the Tolbachik gases shows that gas–rock interaction can also affect volcanic gas composition through heterogeneous catalysis.

211

2. GEOLOGICAL BACKGROUND The New Tolbachik cones appeared during the Great Tolbachik fissure eruption in 1975–1976, Kamchatka, Far East Asia (Fedotov and Markhinin, 1983). The cones form two separate groups. The North Group (55°410 N, 160°140 E, 1200 m asl) consists of three cones, while the South Cone (55°360 N, 160°110 E, 540 m asl) is 10 km to the southwest. Two of the four eruptive cones, the 1st Cone and the 2nd Cone in the North Group, have relative heights of 300 m and volumes of approximately 0.1 km3 (Fig. 1). The remaining two cones are 110–150 m high. The cones are composed of basalts (SiO2 49.8–50.8%) that are enriched in MgO (up to 9.88% in the North Group) and Al2O3 (up to 17.1% in the South Cone). An extensive description of gas discharges during the Tolbachik eruption and soon after the eruption was made by Menyailov et al. (1980). Gases with a highest maximum temperature of 1135 °C were sampled from hornitos over a lava flow from the South Cone (Table 1). Samples from the South Cone are the least contaminated by air and correspond to gases separating from a partially degassed lava flow. All samples taken at the North Cones contain air. In addition to eruptive gases, Menyailov et al. sampled fumaroles on the cones during the three years after the eruption. Fumarolic gases collected in 1975–1978 consisted mainly of air, with variable amounts of water and acid species, and did not contain H2S. In general, the sulphur content in the gases rapidly decreased with time, so that late gases were relatively enriched in HCl and HF. Scoria and massive basalts at the cone summits underwent intensive post-eruptive alteration and mineralization. Three main types of rock alteration were observed. White or yellowish porous crust, consisting primarily of SiO2nH2O, covered massive basalts that were subjected to hydrochloric leaching (Fig. 2a). There were also areas covered by thick solid encrustations, consisting primarily of fluorides.

Fig. 1. New Tolbachik scoria cones (300 m high, North group), which appeared during the Great Tolbachik fissure eruption in 1975–1976, Central Kamchatka, Far East Asia. Arrows show sampling sites. The 1st Cone is on the right and the 2nd Cone is on the left.

212

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

Table 1 Chemical compositions of early Tolbachik fumaroles (Menyailov et al., 1980). O2 H2O CH4 H2 HF HCl SO2 H2S CO2 NH3 T N2 (°C) (vol %) (vol %) (vol %) (vol %) (ppmv) (ppmv) (ppmv) (ppmv) (ppmv) (ppmv) (ppmv)

Location

Date

South Cone, hornito South Cone, hornito 2nd North Cone, hornito 1st North Cone, fumarole 2nd North Cone, fumarole “Cupreous”

11.06.1976 1135 0.16 03.08.1976 1020 0.045 16.08.1975 930 13.8 19.08.1975 440 35.8 15.06.1976 700 57.4 05.08.1976 650 79.5 20.08.1976 710 69.0 05.04.1977 450 77.3

0.02 0.005 0.04 5.0 10.9 16.9 17.2 19.0

97.2 97.8 81.8 46.2 31.5 2.4 13.3 3.3

The third type of alteration exhibits “red scoria” covered by a thin layer of fine iron oxides (hematite + maghemite, Fig. 2a and b). The list of Tolbachik post-eruptive minerals contains more than 90 species (Vergasova et al., 2000; Pekov, 2007), including oxides, halides, sulphates and other oxysalts. No traces of sulphides or native sulphur exist now on the cones, nor were these minerals found earlier since the eruption (Naboko and Glavatskikh, 1984). Fumarolic activity at the cones has persisted until now. In 2010, 35 years after the eruption, there were still many gas vents at the cone summits with temperatures of up to 480 °C. Gas was discharged from holes and fissures in rocks and through loose altered scoria (Figs. 2b and 3a). Most of present-day gas discharges, rock alteration and secondary mineralization are located at the 2nd North Cone. 3. METHODS This study includes field works with in situ measurements of gas compositions, gas and rock sampling, subsequent laboratory analyses of samples, thermochemical modelling and laboratory experiments. Three seasons of fieldwork (2008–2010) were undertaken at the New Tolbachik cones, Kamchatka, Far East Asia. In 2008, chlorine emissions were roughly estimated using gas detector tubes. In 2009 and 2010, 54 in situ measurements and five timeseries measurements of Cl2 were made at the 2nd North Cone; and eight in situ measurements were made at the South Cone. Twenty-two samples of gas and condensate were collected and analysed in the laboratory. Rock samples were also collected, including unaltered scoria that was used in experiments. 3.1. Sampling and analyses 3.1.1. In situ Cl2 and HCl measurements For semi-quantitative assessment of chlorine we used Russian colorimetric gas detector tubes with a measurement range of 0.5–20 mg/m3 (0.15–8.2 ppmv) and 25% accuracy. The required amount of gas was pumped through the tube, and the chlorine concentration was calculated from the tube readings and the number of pump strokes. The principle of this method is the formation of pink eosin from yellow fluorescein and KBr in the presence of Cl2 (Frumina et al., 1983):

0.14 0.08 3.8 10.5 0.17 0.37 0.24 0.27

– 10,400 1500 0.004 7100 1900 0.7 1800 400 21 1500 400 0.001 7 43 0.001 3000 370 0.049 450 310 – – 25

5100 8600 1100 700 280 4800 1300 200

4900 1400 9 100 1100 180 34 4

3000 1200 100 – – – – –

0.1 – 500 400 – – – –

Cl2 þ 2KBr ¼ Br2 þ 2KCl; ð1aÞ 2Br2 þ C20 H12 O5 ðfluoresceinÞ ¼ C20 H6 Br4 O5 ðeosinÞ þ 2Hþ ð1bÞ High HCl and HF content in the gas suppresses the reaction (1b), thus decreasing the accuracy and detection limit of the method. A “Comet-M” gas detector (Russia) equipped with Membrapor SO2/S-20, Membrapor HCl/M-20, and Sixth Sense SureCell Cl2 electrochemical sensors was used for quantitative measurements of Cl2, HCl and SO2. Gas was sampled at flow rates of 0.5–1 l/min through a 25–30 cm Teflon suction tube inserted directly into a fissure (Fig. 3). The gas cooled naturally when passing through the suction tube, and the temperature decreased from 200 °C at the sampling point to 25–30 °C at the gas detector inlet. It is important to note that the composition of the gas from the Tolbachik fumaroles is unique. It consists mainly of air, making it quite different from common fumarolic gases consisting of water vapour and acid species. The components measured (HCl, Cl2) existed in the Tolbachik fumaroles at ppm levels. Low temperatures, the domination of air in the gas composition and the low concentrations of acid species made electrochemical sensors applicable for direct measurements. According to the manufacturer’s specification sheet, the SureCell Cl2 sensor has a measurement range of 0– 20 ppm wt (0–8.2 ppmv), with a possible overload to 50 ppm wt and a relative error of 25%. The sensor also has a cross-sensitivity to HCl that produces a response equal to approximately 13–15% of the HCl concentration. This cross-sensitivity was taken into account as follows: Cl2 (corr.) = Cl2 (measured)  0.15HCl (measured). The Membrapor HCl sensor has a high cross-sensitivity (+150%) to H2S and the Sixth Sense SureCell Cl2 has a high negative cross sensitivity to H2S (65%). Although Tolbachik gases have extremely low H2S contents (see below), we considered the possibility of incorrect readings. Calibration of the SO2, HCl and Cl2 sensors was performed by the manufacturer of the gas detector (in Moscow) using a permeation tube system and nitrogen as a dilution gas. The calibration was performed 2 weeks prior to the 2009 field works. No further calibration of the gas detector was done because many in situ measurements of Cl2 were carried out simultaneously by the gas detector and by the DPD method (Fig. 3), which was regarded as

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

Fig. 2. (a) “Chlorine Field” on the 2nd Cone, North group. Massive blocks of bleached basalt alternate with red scoria, covered by iron oxides. The 1st Cone is in the background. (b) Incrustations around the chlorine-emitting hole on the South Cone, where the highest concentrations of Cl2 (60 ppmv) were measured.

repeated calibration. The SureCell Cl2 sensor had a sensitivity at 0.25 ppm wt Cl2 that corresponded to 0.1 ppmv Cl2. The sensitivity of the DPD method was also estimated during calibration at 0.1 ppmv Cl2, with a reproducibility of approximately 5%. Both methods essentially showed similar readings (within a 25% error interval). However, in many cases the electrochemical sensors were overloaded. The second quantitative method of Cl2 measurement at the cones was the bubbling of the gas through a 1% solution of DPD (N, N-diethyl-p-phenylenediamine) in a phosphate buffer at pH 6.86 (Palin, 1957). The colourless DPD solution turns magenta in the presence of Cl2. The gas was collected directly from fissures in rocks through a Teflon tube using a diaphragm pump from the gas detector and bubbled through a trap filled with the DPD solution (Fig. 3). In cases of high chlorine concentrations, a syringe was used instead of the pump. After sampling, the optical densities of the solutions were immediately measured at 525 nm using an “Expert-003” portable photocolorimeter (Russia). We usually pumped 0.1–1 l of the gas to obtain a proper optical density (D = 0.1–1) in the photometer cell. The phosphate buffer eliminates possible influences from hydrochloric acid that can weaken or eliminate colouration. Chlorine was almost entirely (95–99%) absorbed by the DPD solution,

213

Fig. 3. (a) Simultaneous measuring of Cl2 concentrations by the gas detector and by the DPD solution in the gas emitting from a fissure. The 2nd Cone, North group. (b) Scheme of the installation shown in label (a).

even in a single trap. Calibration of the photocolorimeter was performed with a Hach free chlorine standard solution. As a rule, a quick assessment of Cl2 concentrations was made at each sampling site with the gas detector. Openings with high Cl2 were additionally studied by a series of 3–4 DPD measurements, with the highest value selected as a representative. Both electrochemical and DPD methods of Cl2 measurements have 100% cross-sensitivity to NO2 and other strong oxidisers. To evaluate the possible existence of oxidising gases other than Cl2, thermochemical modelling was performed. 3.1.2. Measuring of Cl2 and HCl in experiments To measure Cl2 concentrations in experiments with a flow reactor, gas was passed through a bubbler filled with 5–10 ml of a solution of 30 g/l CdI2 + 5 g/l of soluble starch (Frumina et al., 1983; and references therein). Chlorine was almost completely absorbed by the solution in a single trap. In the case of high Cl2 concentrations, gas was sampled using a syringe filled with the same solution. The optical density of the iodide-starch solution at 590 nm is almost linearly dependent on the quantity of Cl2 absorbed over three orders of magnitude. Coupled with a variable amount of sampled gas (0.5–600 ml), this method provided a measurement range of six orders of magnitude (0.05 ppmv–5% mol Cl2). Instead of the commonly used KI-starch solution, we

214

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

used a CdI2-starch solution that is stable for a period of several weeks (Frumina et al., 1983) and does not change colour in the presence of HCl or atmospheric oxygen (as the KI-starch solution does). The technique described differs from that applied in the field. We used the iodide–starch method in our experimental work because of its wide measuring range of six orders of magnitude and its insensibility to high HCl concentrations. High concentrations of HCl in experimental gases (0.5 vol % = 5000 ppmv) instantly suppressed the colouration of the DPD solution, even when buffered, and were deleterious to sensitive electrochemical sensors. Electrochemical sensors were used in the field because of their convenience and their quick responses in cases of low chlorine concentrations. Both the DPD method and sensors were used, as these two methods are both well established and recognised. To measure Cl2 and HCl simultaneously, a solution of 30 g/l CdI2 + 5 g/l of soluble starch + 6 g/l KIO3 was used. In an acidic environment, iodide reacts with iodate to form free iodine: 12HCl + 5CdI2 + 2KIO3 = 6I2 + 2KCl + 5CdCl2 + 6H2O (Frumina et al., 1983). I2 then forms a blue starch-iodine complex. The iodide–starch solution only reacts with Cl2, whereas the iodide–iodate–starch solution turns blue in the presence of both HCl and Cl2 with the same optical density per equivalent concentration of chlorine. The concentration of HCl can be calculated by subtracting the chlorine concentration from the total measurement. Unlike Cl2, hydrochloric acid was only partially absorbed by the trap solution. We used a 400-mm long titanium capillary with a 1-mm I.D. and a 1-mm wall that served as a condenser, quantitatively collecting both gaseous HCl and acidic mist. Two to twenty milliliters of sampling gas were pumped through the capillary. Then the capillary was flushed with the iodide–iodate–starch solution. This technique exhibited good reproducibility (5%) and sensitivity (10 ppmv HCl). 3.1.3. Gas sampling and analysis Gas components other than Cl2 were analysed in samples collected into empty evacuated glass ampoules and in condensates. Gas samples were collected via a Teflon tube connected by a T-joint to the ampoule; the third opening of the T-joint was used for purging the system of air before sampling. Concentrations of N2, O2, Ar, CO2, CH4 and H2 were determined in samples using gas chromatography. Separation of Ar from O2 was performed using an Altech CTR-III packed column with He as a carrier gas. This column and the thermal conductivity detector were also used for the analysis of N2, O2, CO2, and CH4. Ar was used as a carrier gas with a packed column with Molecular Sieves 5A for the determination of H2. The condensate was collected into a U-shaped silica glass tube immersed in liquid nitrogen. Air was pumped through the tube at a rate of 3 l/min. With the water content in the gas at 10–50 mg/m3, 1–5 h were needed to collect an appropriate amount of the condensate. In the laboratory, the condensates were analysed by ion chromatography to detect Cl, F and SO2 4 . The water content of the gas was calculated from the weight of the condensate and

the sampling flow rate. Mass spectrometric measurements of d18O and dD were carried out using a DELTA Plus mass spectrometer (Finnigan) coupled an H/D device. Accuracies were approximately 0.2 & for d18O and approximately 1& for dD. 3.1.4. Analyses of solid samples and solutions The Na, Mg, Al, P, S, K, Ca, Fe and Cu contents of basaltic scoria and hydrochloric leachates from the scoria were determined using ICP-AES (ICAP-61, Thermo Jarrell Ash, USA). The measurements were made with the following spectrometer parameters: a generator output power of 1200 W; an angular nebuliser; a plasma-forming Ar flow rate of 18 l/min; an auxiliary Ar flow rate of 0.9 l/min; an Ar flow rate into the atomiser of 0.6 l/min; an analysed sample flow rate of 1.5 ml/min; and a plasma observation zone height of 14 mm. The integration time of the spectrum during measurements was 5 s. Determination of the element contents in aqueous solutions was made by the quantitative method using calibration solutions (High Purity Standards, USA) of 0.5 and 10 mg/l of each element. The relative standard deviation did not exceed 5%. The content of SiO2 in scoria was calculated by difference. Phase analyses were performed using XRD (Bruker D8 Advance) at 60 °C in a vacuum to avoid deliquescence of hygroscopic samples. XRD patterns were interpreted with the 2009 ICDD database and Bruker Eva software. Electronic microphotography and local microanalyses were made using a Tescan Vega TS 5130MM scanning electron microscope equipped with an INCA energy-dispersive Xray spectrometer. 3.2. Experimental studies 3.2.1. Reactor design The oxidation of HCl without a catalyst and the catalytic properties of the basaltic scoria, reference Fe2O3 and CuCl2 catalysts were studied in a flow reactor (Fig. 4) made of high purity silica glass with a SiO2 content that exceeded 99.988%. Glass contained less than 75 ppm Al, 15 ppm Na, 15 ppm Ca and less than 3 ppm of transition metals. The reaction was conducted in the central zone of the reactor (50% of its length). Each end of the reactor tube was filled with inert silica glass granules. The temperature was controlled by two type-K thermocouples. One thermocouple was placed directly on a coil, and the other was placed in an isolated channel inside the reactor. Variations of the temperature field within the reaction zone were less than 5° in the axial direction and less than 2° in the radial direction during the heating/cooling of the reactor at a rate of 0.4°/min. In experiments at constant temperature, temperature variations were less than 2 °C in both the axial and radial directions. Air entering the reaction zone was fed through a 3-mm I.D. silica tube inside the reactor. This tube also served as a gas heater. A silica fibre thread inside this tube ensured the smooth evaporation of the hydrochloric solution that was supplied by a peristaltic pump. Typically, we used air at a flow rate of 10 cc/s together with 0.5 mm3/s of 15% HCl to produce gas containing 20 mol % O2, 5 mol %

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

215

Fig. 4. Scheme of the flow reactor used in experiments. Not to scale, without heat insulation.

H2O and 0.5 mol % HCl. The same reactor design without the silica fibre thread, was used to study the temperature-induced decomposition of Na, K and Mg chloroferrates (products of basalt acid leaching that we consider as another possible source of chlorine). Reaction products were let out from the reaction zone through a 3-mm I.D. silica tube and quenched to 150 °C within 0.1 s. The estimated residence time for gas within the reactor, taking into account reactor filling, was 5 s at 300 °C and 3 s at 600 °C. 3.2.2. Experimental materials To prepare “basaltic catalyst”, a 3–5 mm scoria fraction was put in the reactor. Air with an additional 5 mol% H2O and 0.5 mol% HCl was passed through the reactor at a constant flow rate of 10 cc/s at 600 or 400 °C for 1-2 days. The reference Fe2O3 catalyst was prepared by drying a solution of Fe (III) nitrate together with supporting granules at 110 °C. Dried granules were then calcinated at 450 °C for 24 h and purged with air to eliminate traces of NO2. Alternatively, Fe2O3 catalyst was deposited onto supporting granules from a suspension of Fe2O3 powder. The reference CuCl2 catalyst was prepared similarly by drying a solution of Cu (II) chloride with supporting granules; the dried granules were then calcinated at 200 °C to evaporate the associated water. Chloroferrates were made by the evaporation of hydrochloric leachate from basaltic scoria. Chloroferrates were also made from laboratory chemicals. To prepare hydrochloric leachate from the scoria, 3- to 5-mm scoria fractions were put into sealed polypropylene beakers with 0.5-3 N solution of HCl at a scoria to solution weight ratio of nearly 1:1. The beakers were heated to 85–90 °C and vortexed. The duration of the reaction was 24 h. A longer reaction time produced no significant difference in the results. The leachate and the 3- to 5-mm silica glass granules was then dried at 110 °C. These granules served as support for the investigated substance that was exposed to the heating and air fluxes inside the reactor. Reference chloroferrates were also prepared from laboratory chemicals. NaCl, KCl, MgCl26H2O and CaCl2 and equivalent amounts of FeCl36H2O were dissolved in distilled water. This solution was dried at 110 °C with the

support of silica glass granules. Dry residue was examined under an electronic microscope with an EDS spectrometer and by XRD. All materials were prepared immediately before experiments because chloroferrates (except for potassium ones) are highly hygroscopic. 3.2.3. Implementation of experiments Three types of experiments were performed: (1) oxidation of HCl by air without a catalyst, (2) oxidation of HCl by air in the presence of different catalysts, and (3) oxidative decomposition of chloroferrates, including dried hydrochloric leachate from basaltic scoria. The oxidation of HCl by air without a catalyst (homogeneous process) was studied for gas containing 20 mol % O2, 5 mol % H2O and 0.5 mol % HCl at 500 to 1100 °C. Measurements were performed using an empty reactor, for three different ratios of the reactor volume to the reactor surface, to reveal possible catalytic activity of the reactor walls. Chlorine concentrations were found to be almost independent of the volume/surface ratio of the reactor, which indicates the homogeneous oxidation of HCl. The catalytic oxidation of HCl by air in the presence of different catalysts was carried out with the same initial gas compositions from 180 to 700 °C. For these experiments, the central 50% inner zone of the reactor was filled with “basaltic catalyst” or silica glass granules covered with reference catalysts (see Section 3.2.2). In each experiment, the temperature range was chosen to ensure measureable chlorine concentrations at the reactor outlet. To study the oxidative decomposition of dried hydrochloric leachate and chloroferrates, pure air, with a constant flow rate of 10 cc/s was passed through the reactor filled with the substance to be examined supported on granules of silica glass. The temperature was increased at a rate of 0.4°/min (24°/h). Reaction products were sampled and measured at the reactor outlet as described in Section 3.1.2. 3.3. Thermochemical modelling We used the Outokumpu HSC 6.1 Chemistry computer code for modelling (Roine, 2007; see also www.outotec.com). Equilibrium calculations were made using the

216

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

Table 2 Chemical compositions of Tolbachik gas discharges. Sample

Date

T (°C)

N2 (vol %)

O2 (vol %)

Ar (vol %)

H2O (vol %)

CO2 (ppmv)

CH (ppmv)

H2 (ppmv)

HF (ppmv)

HCl (ppmv)

S total (ppmv)

Cl2 (ppmv)*

L1 L2 L3 H1 H2 H3 H4 L4 L5 Air**

07.07.2009 08.07.2009 08.07.2009 07.07.2009 07.07.2009 25.08.2010 25.08.2010 22.09.2010 22.09.2010

185 175 212 427 330 480 338 210 186 20

77.1 77.2 77.2 73.4 72.8 75.2 75.9 77.5 77.4 77.30

20.6 20.5 20.6 19.2 19.6 20.1 20.2 20.7 20.7 20.74

0.92 0.91 0.92 0.88 0.87 0.90 0.90 0.92 0.93 0.922

1.30 1.26 1.28 5.9 6.7 3.7 2.9 0.82 0.87 1.0

780 1080 630 990 650 540 580 520 540 370

2.4 3.4 3.4 2.0 2.9 4.5 3.5 7.9 12 1.7

10 19 24 77 52 84 69 5 7 0.55

0.05 0.14 0.09 19 14 22 12 0.10 0.11

80 55 65 35 6.5 31 6.7 41 48 <0.0012

0.015 0.016 0.013 0.011 0.014 0.011 0.048 n. d. n. d.

5.9 9.8 12.6

* **

Concentrations of Cl2 are shown for the same vents where full analyses were made. At 20 °C and 60% relative humidity.

Table 3 Summary of chlorine measurements at Tolbachik cones. Location

Method

Number*

Concentrations (ppmv) Mean**

2nd North Cone

Sensor DPD

South Cone

DPD

Min

Max ***

54 14

5.8 9.7

0.24 3.6

8.2 12.6

8

30.3

1.2

60

St. dev. 2.3 3.6 25.7

*

Only measurements with detectable amounts of chlorine are included. Time-series measurements are not included. Only openings with the highest Cl2 levels were studied by the DPD method. *** The upper detection limit of the gas detector was 8.2 ppmv Cl2. **

Gibbs energy minimization method in the H-N-O-Cl and H-N-O-Cl-S systems with models consisting of 67 or 105 species respectively, from 100 to 1100 °C in 20 °C intervals. Measured gas compositions L5 (the highest Cl2 concentration) and H3 (the highest temperature) were accepted as initial compositions for the modelling. We also made calculations for early gas compositions from the “Cupreous” fumarole (Menyailov et al., 1980). We did not include CO2 and fluorine species in the model because of their low reactivities in the gas phase under the conditions studied. Modelling was necessary to estimate equilibrium Cl2 concentrations and to evaluate the possible presence of oxidizing gases other than Cl2. 4. RESULTS AND DISCUSSION 4.1. Tolbachik gases 4.1.1. General gas composition Tolbachik gases consisted of N2 (73–77 vol %), O2 (19– 21 vol %), Ar (0.88–0.93 vol %), water vapour (1–6 vol %) and acid species (HCl, HF, CO2, H2SO4, Cl2, total less than 0.1 vol %). The cones emitted two types of gas with different temperatures, water vapour contents and acid species contents (Table 2). The most obvious difference was in temperature (175–210 vs. 330–480 °C), so we denoted the low-temperature gas as L-type and the high-temperature gas as H-type. The water content in the L-type gas was 4–6 times lower and the HF content was one order of mag-

nitude lower than in the H-type gas, while all other measured components, except for molecular chlorine (Table 3), were almost equal. The isotopic composition of the water vapour (dD = 148& to 106&, Table 4) indicated that it had primarily meteoric origin. Water in the L-type gas exhibited a very large positive d18O-shift of approximately 22&, which could be explained, at least partially, by complete isotopic exchange with Tolbachik basalts (d18O = +5.7, Bindeman et al., 2004) and strong fractionation during the evaporation of meteoric water. At the same time, water vapour from the H-type gas possessed the same isotopic composition as local meteoric water. The composition of the Tolbachik gases corresponds to hot air (94–99 vol %) mixed with variable amounts of water vapour and acid species. The origin of such gases can be explained as follows. Because the scoria cone has a loose, porous and fractured structure, ambient air penetrates the cone at the bottom. The air is heated and mixes with meteoric water vapour and acid species (HF, HCl, HBr, CO2 and sulphur species) that are slowly released from the hot scoria. This “volcanic air” is discharged at apical parts of the cone until the latter is completely cold. The two different gas types at the 2nd North Cone most likely arose due to the complex internal structure of the cone. The L-type gas discharges from more peripheral parts of the cone that are composed mainly of loose scoria while the H-type fumaroles are confined to the central zone, which is built of large blocks of massive lava. The large

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

217

Table 4 Isotopic compositions of water in condensates from Tolbachik gases, and meteoric waters. Sample*

Date

dD

d18O

L1 L3 H1 M1 M2 M3 M4

07.07.2009 07.07.2009 07.07.2009 04.07.2009 07.07.2009 09.07.2009 09.07.2009

–148 –106 –110 –120 –100 –144 –124

+11.2 +9.8 11.8 14.5 11.9 19.0 16.6

* L1, L2 and H1 – gas, the 2nd North Cone; M1–M2 – rain, M3 – snow, M4 – cold spring.

d18O-shift indicates high extent of isotopic exchange and probably derives from a high residence time of the L-type gas within a hot zone. This is true at least for water vapour from this gas. It is worth noting that the gas emission from a “Cupreous” fumarole in 1977, two years after the eruption (Table 1, sampling date 05.04.1977; Menyailov et al., 1980), was similar in its composition to the emission from the same vent in 2009–2010, 32–33 years later (Table 2, samples H2 and H4). Although concentrations of acid species decreased by a factor ranging from 2 (HF) to 30 (HCl), temperature and N2, O2 and H2O contents were almost identical. At the same time, gases sampled by Menyailov from hornitos over lava flows were similar to gases from arc volcanoes, partly depleted in CO2 (Gerlach, 2004; Fischer, 2008). 4.1.2. Molecular chlorine in Tolbachik gases The most prominent feature of the Tolbachik emissions was the very high concentration of Cl2 that reached 60 ppmv, or 0.015 wt.% (Tables 2 and 3). Concentrations of Cl2 in the low temperature (L-type) gas were typically 6–12 ppmv but ranged from 0.24 to 60 ppmv. The high temperature (H-type) gas had Cl2 concentrations below the detection limit. The highest concentration of Cl2 (60 ppmv, or 0.015 wt%) was measured in a small fumarole vent (Fig. 2b) located on the South Cone of the Great Tolbachik fissure eruption. This value is more than five orders of magnitude higher than the highest concentrations of Cl2 ever measured in a natural environment in a marine boundary layer (Spicer et al., 1998; Pszenny et al., 1999; Keene et al., 2007). 4.1.3. Temporal fluctuations of Cl2 Five time-series in situ measurements of Cl2 and HCl with durations from five min to one day were made in the 2009 field season. The “Comet-M” gas detector allows for continuous measurements with a time resolution of 15 s, although the response times of the electrochemical sensors, according to manufacturer’s specification sheets, are within 1–2 min. We found that our Cl2 sensor had a fast response time of less than one minute, while the HCl sensor had much worse performance with a response time of approximately 5 min. Continuous measurements have shown that Cl2 emissions could be highly unstable, with fluctuation periods from one to several minutes (Fig. 5). The most

Fig. 5. Temporal variations of Cl2 and HCl in the gas emission. Before measurements, sensors were equilibrated with the gas for 15 min. The dashed line shows the moment after which the gas detector was exposed to fresh air. The response time of the Cl2 sensor was less than one minute, while response time of the HCl sensor was approximately 5 min.

probable cause of such instability is shallow mixing of chlorine-containing gas with pure air, induced by wind gusts. These were observed during continuous measurements. Fluctuations in Cl2 concentrations show that porous basaltic scoria is highly permeable for ambient air. Additionally, chlorine emissions are sensitive to the inner structure of the cone as well as to weather conditions. 4.2. Oxidation of HCl The existence of molecular chlorine in the Tolbachik gases can be explained by the reaction between HCl and oxygen. Oxidation can occur in a gaseous phase (homogeneous reaction) or on the surface of a solid catalyst (heterogeneous reaction). The stoichiometric chemical equation of the reaction is very simple: 2HCl þ 0:5O2 ¡Cl2 þ H2 O:

ð2Þ

The mechanisms of both the homogeneous and heterogeneous oxidation processes are much more complicated than the stoichiometric equation and are beyond the scope of this article. 4.2.1. Thermodynamics and kinetics of the non-catalytic oxidation of HCl in the gaseous phase From a thermodynamic perspective, the formation of Cl2 in an H–N–O–Cl–S system (“volcanic air”) is favoured at relatively low temperatures (Fig. 6a). In an excess of oxygen, temperature and water content are two parameters that control the equilibrium composition. Both high temperature and high H2O content shift the equilibrium in the reaction (2) to the left while high concentrations of HCl result in chlorine formation. The modelling shows that equilibrium concentrations of Cl2 increase when temperature decreases. Concentrations of all other chlorine species decrease with decreasing temperature. This is a key feature of the reaction (2) that enables the formation of chlorine at low temperatures.

218

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

absence of hydrogen sulphide (Menyailov et al., 1980; Table 1). A strong evidence of low H2S content in the gases is the total absence of sulphides and native sulphur in fumarolic incrustations at the Tolbachik cones. At the same time, sulphates (chlorosulphates) are ubiquitous, thus indicating some previously higher content of sulphur in gases. Measured concentrations of Cl2 in the low-temperature L-type gas ranged from 5–15% of the calculated equilibrium values at the 2nd North Cone to almost equilibrium values at the South Cone (50–60 ppmv Cl2 together with 40–50 ppmv of HCl at 200 °C). For the high-temperature H-type gas, the calculated Cl2 equilibrium concentrations were in the range of 0.004–0.015 ppmv. This range falls well below the 0.1 ppmv Cl2 detection limits of the electrochemical gas detector and DPD solution methods. Early fumaroles at the 2nd North Cone, 1–2 years after the eruption, had temperature range of 350–710 °C and contained large amounts of air and hydrochloric acid simultaneously (Table 1). These conditions could be favourable for Cl2 formation. Calculated Cl2 equilibrium concentrations in early fumarolic gases were from 0.03 ppm for water-rich gas to 3 ppm for water-poor gas. Modelling for the sample dated 05.08.1976 indicated as much as 160 ppm equilibrium chlorine. However, this seems arguable because of the very high HCl/H2O ratio in the reported gas composition corresponding to 29% hydrochloric acid in the gas condensate. In addition to equilibrium modelling, kinetics considerations should be taken into account. The homogeneous gas-phase reaction (2) is a first order reaction with respect to HCl and to O2 (Gavrilov et al., 1975). Overall, it is a second order reaction. However, in an excess of oxygen in diluted air-HCl mixtures, the gas-phase reaction (2) can be considered as a pseudo-first-order reaction. In this case, the Arrhenius equation for the reaction rate constant k can be written as follows (Gavrilov et al., 1975): Fig. 6. Equilibrium concentrations of gaseous species in the H–N– O–Cl–S system vs. temperature. (a) Plot for all species except sulphur species; (b) plot for sulphur species. The formation of Cl2 is favoured at low temperatures. Below 300 °C, the equilibrium concentrations of NOX and ClOX are 2.5–5 orders of magnitude lower than those of Cl2. Sulphur species are shown on a separate plot for clarity.

Thermochemical modelling indicated that the possible influence of oxidising gases other than Cl2 on our measurements was negligible. Calculated equilibrium concentrations of NO and NO2 at 180–200 °C were 2.5–5 orders of magnitude lower than those of Cl2, and concentrations of ClX (X = O0.5, O, O2, OH, NO, NO2) were 3 to 5 orders of magnitude lower. Therefore, Cl2 should be the only oxidising species in the gas to affect the electrochemical gas detector or the DPD solution used for measurements. According to the modelling, sulphur exists in the Tolbachik gases below 300 °C primarily as H2SO4 and SO3, with minor quantities of SO2 and other S (IV) species (Fig. 6b). Hydrogen sulphide can persist in the air-gas mixture for kinetic reasons; however, we assumed that its role in late Tolbachik gases during the time of our sampling was insignificant. Although H2S was measured in the eruptive gases, early fumaroles on the cones have exhibited the

k ¼ 10^ ð10:1  0:3Þ  exp½ð213:8  2Þ=RTL mol1 s1 : ð3Þ The high activation energy (213.8 kJ/mol) indicates that the reaction rate rapidly decreases with decreasing temperature (Figs. 7 and 8). The conversion time of HCl to Cl2 at a temperature T can be estimated according to the following equation:   ½HCl  ln ½HCl0  t¼ ; ð4Þ k  ½O2  where [HCl0] denotes the initial HCl concentration, [HCl] denotes the concentration of HCl at a time t, and k is the reaction rate constant at temperature T, with all concentrations expressed in mol/l. Calculations according to Eq. (4) provide unrealistically long times for reaction (2) to reach the concentrations of Cl2 observed in the Tolbachik gas at low and moderate temperatures (Fig. 8). Thus, to achieve oxidation of 20% HCl (typical values for the Tolbachik cones), it takes 1.8  108 s (5.9 years) at 400 °C or 6  1011 s (19,000 years) at 280 °C. For temperatures above approximately 280 °C, equilibrium ratios of Cl2/HCl are

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

219

Fig. 8. Calculated time required to achieve a 20% oxidation of HCl in the Tolbachik gases vs. inverse temperature. Curve 1 corresponds to the Eqs. (2)–(4) for the homogeneous oxidation of HCl by O2 without a catalyst. Curves 3–6 are calculated from our experimental data for the oxidation of HCl with different catalysts: 3 – Fe2O3 powder, 4 – calcinated Fe (III) nitrate, 5 – CuCl2, 6 – “basaltic catalyst”.

Fig. 7. The logarithms of kinetic constants vs. inverse temperature (Arrhenius plot) for the homogeneous oxidation of HCl by oxygen (curves 1 and 2) and for catalytic reactions in the presence of the basaltic scoria, Fe2O3 and CuCl2 (curves 3–6). A constant flow of air with 0.5 mol % HCl was passed through the reactor. The basaltic scoria in the reactor was exposed to 600 °C for 24 h before measurements were taken. Fe2O3 and CuCl2 were deposited onto support silica granules.

lower than the ratios measured; therefore, they cannot be achieved at all. Low rates of the direct (homogeneous) oxidation of HCl invoke the catalytic nature of the chlorineproducing process at the Tolbachik cones. 4.2.2. Basalt as a catalyst The oxidation of HCl is greatly accelerated in the presence of chlorides, oxychlorides and oxides of transition metals. The catalytic oxidation of HCl is known as the Deacon process (Hisham and Benson, 1995, and references therein). In our experiments, in the presence of CuCl2, known as a catalyst with a moderate activity (Hisham and Benson, 1995), the rate of oxidation of HCl increased by 11.5 orders of magnitude at 200 °C compared to the rate of reaction without a catalyst. The oxidation of HCl at 200 °C with a copper chloride catalyst occurs at a noticeable rate and can be easily measured (Fig. 7, curve 5). At the same time, the oxidation of HCl without a catalyst is so slow at low temperatures that the reaction rate can only be obtained through extrapolation of data from the 500 to 1100 °C range. The basaltic scoria treated by air with the addition of 5 mol % H2O and 0.5 mol % HCl at 600 or 400 °C for 1– 2 days showed strong catalytic properties comparable to those of the reference catalysts CuCl2 and Fe2O3 below 400 °C (Fig. 7, curve 6). At higher temperatures, the catalytic activity of the scoria decreased. The rate of oxidation

HCl with “basaltic” catalyst exceeded the corresponding rates of the gas–phase reaction (2) by three orders of magnitude at 600 °C and by 10 orders of magnitude at 200 °C. It should be noted that the activity of reference catalysts CuCl2 and Fe2O3 crucially depended on how they were prepared (actually, the influence of the preparation on the specific surface area of the catalyst). After treatment with air-HCl gas, particles of the altered scoria were covered with NaCl and fine Fe2O3 crystals (Fig. 9a). A smaller amount of CuO crystals was also observed (Fig. 9b). The colour of the altered scoria indicated that c-Fe2O3 crystals (maghemite, brown powder) most likely appeared below 400 °C while a-Fe2O3 (hematite, red powder) prevailed above 500 °C. For Fe2O3 powder prepared by calcination of Fe (III) nitrate, XRD analysis showed that it was a mixture of a-Fe2O3 and c-Fe2O3. However, we failed to instrumentally determine whether a-Fe2O3 or c-Fe2O3 crystals were on the scoria surface. The fine Fe2O3 crystals (either a-Fe2O3 or c-Fe2O3 or both) were possible main substances to catalyse oxidation of HCl in our experiments. First, the reference Fe2O3 catalyst prepared from Fe (III) nitrate provided approximately the same effect, at least at the low-temperature end. Second, the main visible effect on the scoria particles after the experiment was the appearance of fine Fe2O3 crystals on their surfaces. Other oxides, such as CuO (which can be easily converted to CuCl2), and, in general, oxides (chlorides, oxychlorides) of all transition metals that are contained in basaltic glass, could also have contributed to the catalytic activity of the altered basalt. The observed decline in catalytic activity of the altered scoria above 400 °C (Fig. 7) could have been caused, at least partially, by volatilisation of the crystals of Fe2O3 or their recrystallization in to larger crystals. Both processes decrease the free surface of the catalyst and, therefore, its activity. Other possible explanations are the volatilisation

220

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

of some highly active catalysts, such as vanadium-based substances (Tarabanko et al., 2009 and references therein), cannot be excluded. Seven copper vanadates were observed at the Tolbachik cones (Zelenski et al., 2011). 4.2.3. Influence of sulphur concentration The low concentration of sulphur (Table 2) promotes the formation of Cl2 inside the cones. Although Cl2 and ClOX can coexist with SO2 in diluted air mixtures, they can be reduced back to HCl by sulphur species with oxidation states lower than +6 if concentrations of the latter are high enough: SO2 þ Cl2 þ 2H2 O ¼ H2 SO4 þ 2HCl:

ð5Þ

Direct investigations of the SO2/Cl2 interaction (Ryan et al., 2006 and references therein) have shown that the homogeneous conversion of Cl2 to HCl in the presence of SO2 is a slow reaction at temperatures below 750 °C. It is more likely that the suppression of HCl oxidation by SO2 is related to the reduced concentrations of Fe and Cu oxides/chlorides because of their conversion into sulphates (Ryan et al., 2006). In any case, the S/Cl ratios for the L-type Tolbachik gases were 0.016–0.00025, resulting in an insignificant influence of sulphur on the total oxidation balance. 4.3. Acid leaching of basalt as a possible source of Cl2

Fig. 9. Scanning electron micrographs of altered scoria. (a) Crystals of Fe2O3 (fine, white) with NaCl (rounded, grey) and (b) crystals of CuO (white) with NaCl (rounded grey) on a surface of basaltic scoria after exposure at 600 °C for 24 h in air with 0.5 mol % HCl. SEM BSE images.

of copper chlorides that formed on the scoria surface at lower temperatures (200–300 °C), and a decrease in the activation energy of the reaction due to some unknown changes in the reaction mechanism. Volatilisation of Fe and Cu compounds from the scoria surface during the experiments, resulting in the subsequent deposition of FeCl3, CuCl2 and KCuCl3 at the reactor outlet, was actually observed. The following arguments support the hypothesis that the catalytic oxidation of HCl is a source of molecular chlorine in Tolbachik emissions. Above 280 °C, calculated equilibrium concentrations of Cl2 (more exactly, Cl2/HCl ratios) for the L-type gas were lower than measured values. Below 280 °C, too much time is required to achieve the measured concentrations of Cl2 if we consider only the non-catalytic oxidation of HCl. There are plenty of substances under the cone surface that can play the role of catalyst. Even a catalyst with a moderate activity (Fe2O3) greatly decreased the reaction time (Fig. 8). Fine Fe2O3 crystals (hematite/maghemite) are abundant at the Tolbachik scoria cones, both at the surface and under the surface, at least to a depth of 50 cm (Figs 2a, b). Copper compounds, such as anhydrous CuCl2 (tolbachite), CuCl (nantokite), CuO (tenorite) and Cu2OCl2 (melanothallite), together with variable amounts of NaCl, occur within the scoria subsurface layer (Vergasova et al., 2000; Pekov, 2007 and our observations). The contributions

The catalytic oxidation of HCl is the most plausible process for the generation of Cl2 inside the Tolbachik cones. However, it may not be the only one. Some products of acid leaching of basalt, namely, so-called chloroferrates, can also produce Cl2 when heated in the presence of air. Chloroferrates are complex chlorides of Fe and alkali (alkali earth) metals with the common formula of mMeClnFeCl3, Me = Na, K, 0.5 Mg, etc. In our experiments, chloroferrates of Na, K and Mg instead of individual chlorides crystallised when the polycation hydrochloric leachate solution of rock-forming elements (Na, K, Mg, Ca, Al and Fe) evaporated (Fig. 10a). The identification of chloroferrates was challenging. The dried polycation hydrochloric leachate was studied using XRD and SEM-EDS. In addition, dried residues from the aqueous solutions of MgCl2–FeCl3 and KCl–MgCl2–FeCl3 were studied as reference samples. The interpretation of the XRD patterns of all of the samples was handicapped by the fact that only few of the numerous chloroferrates are included in the ICDD database. Only KFeCl4, K2FeCl5H2O and MgCl22H2O were recognised in the XRD patterns. At the same time, several unrecognised strong peaks were found to be identical both in the XRD pattern of the dried leachate and in the patterns of the reference samples. In all samples, semi-quantitative EDS microanalyses yielded Mg chloride and components such as K2FeCl5 – (K,Mg0.5)2FeCl5 – MgFeCl5, (K,Mg0.5)FeCl4 and (Mg,Ca)(Fe,Al)0.33Cl3, probably with some amount of crystallization water. From this data, we concluded that K–Mg chloroferrates and Mg–Ca–Fe–Al complex chlorides produced unrecognized strong peaks in all XRD patterns. The polycation hydrochloric solution forms within a surface scoria layer at the Tolbachik cones (and also in

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

221

Table 5 Elemental compositions of Tolbachik basaltic scoria and hydrochloric leachates from this scoria, ppm. Only rock-forming elements and copper are included.

Fig. 10. (a) Scanning electron micrograph of dried hydrochloric leachate from the basaltic scoria consisting of K2FeCl5H2O (k), MgCl22H2O (m), MgFeCl5 (f) and Al chlorohydrate (h). Phases were distinguished according to XRD and semi-quantitative EDS analyses. (b) Dried hydrochloric leachate from the basaltic scoria after decomposition in air flux at 300 °C for 6 h. The two main phases are hematite globules and a matrix composed mainly of MgOCl. SEM BSE images.

Element

Basalt

3 N HCl

1 N HCl

0.5 N HCl

Na Mg Al Si K Ca Fe Cu

20,300 61,000 81,000 232,000 8600 94,000 81,000 155

1520 9150 4800 165 800 3930 8580 20

590 3300 1970 140 295 1450 3150 8.1

220 1480 540 720 95 520 1140 3.9

Fig. 11. EDS profile of rock-forming oxides for the outer layer of leached basalt. A value of zero on the horizontal axis corresponds to the sample surface. All rock-forming elements except Si are depleted in the bleached crust in almost equal proportion. SiO2nH2O makes up the rest of the analyses (not shown).

Competitive decomposition reactions without evolution of Cl2 also proceed: laboratory experiments) as a result of hydrochloric acid leaching. During the experimental leaching of basalt by hydrochloric acid at very low pH values, rock forming and trace elements, with the exception of Si, migrated to the fluid in equal proportions, and the latter acquired a composition that reflected close to congruent rock dissoluP tion (Table 5). The (cations)/Cl ratio (in terms of equivalents) in the leachate was near 0.6 independently of the HCl concentration. When such leaching is complete, basalt turns to a yellowish-white porous rock that consists almost solely of SiO2nH2O (Fig. 11). The described process differs from the well-known acid leaching in volcanoes (Getahun et al., 1996; Africano and Bernard, 2000) because HCl is almost solely responsible for leaching, while the concentration of sulphuric acid, common on volcanoes, is very low. In the presence of air oxygen, chloroferrates undergo the following two-stage thermal decomposition: K2 FeCl5  H2 O ¼ K2 FeCl5 þ H2 O ð> 100  CÞ ð6aÞ K2 FeCl5 þ 0:5O2 ¼ 2KCl þ FeOCl þ Cl2 ð210  300  CÞ ð6bÞ 2FeOCl þ 0:5O2 ¼ Fe2 O3 þ Cl2 ð> 300  CÞ

ð6cÞ

K2 FeCl5  H2 O ¼ 2KCl þ FeOCl þ 2HCl ð200  300  CÞ ð7aÞ 3FeOCl ¼ Fe2 O3 þ FeCl3 ð> 360  CÞ ð7bÞ Reactions (6a–c) occurred sequentially if the heating rate exceeded 20°/h. Reactions (7a) and (7b) prevailed when the heating rate was below l0°/h. Reactions (6a), (6b), (6c) and (7a) were derived from this experimental study through analyses of starting materials and products; reaction (7b) was studied by Dai et al. (2003). Independent of the reaction mechanism, the final products of the decomposition included hematite Fe2O3 and anhydrous chlorides (oxychlorides) of K, Na, Mg, Ca, Al (Fig. 10b) and sometimes FeCl3. The oxidative temperature-induced decomposition of the dried hydrochloric leachate from basalt as well as reference Na, K and Mg chloroferrates was studied in the flow reactor (Fig. 4). The substance was heated in the reactor with a constant airflow and a linear temperature increase of 0.4°/min. During heating, the dried leachate emitted Cl2 with several noticeable peaks (Fig. 12a). The positions of these peaks correlate well with the corresponding peaks

222

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

Fig. 12. Temperature-induced decomposition of chloroferrates. (a) Production of Cl2 during decomposition of the dried hydrochloric leachate from the basaltic scoria. (b) Peaks of Cl2 for individual Na, K and Mg-chloroferrates, synthesized from pure chlorides, and for FeCl36H2O; decomposition under the same conditions. Numbers denote: 1 – Mg-K-chloroferrates; 2 – K-chloroferrates; 3 – Na3FeCl6; 4 – NaFeCl4, 5 – FeCl36H2O; 6 – FeOCl.

from individual chloroferrates and FeCl36H2O (Fig. 12b) heated in the same conditions. To explain the mechanism of possible chloroferrate formation and subsequent decomposition at the volcanic cones, we considered a model of cyclic scoria leaching and flushing (Fig. 13). The model contains two layers of scoria. The upper (surface) layer is exposed to precipitation; it is warm (20–80 °C) and wet. The lower layer is 20–40 cm under the surface; it is hot (200–240 °C) and dry. A flux of air with little admixture of HCl constantly percolates from underneath through the loose scoria. In wet weather, particles of the scoria from the upper layer (Fig. 13a) are covered by a thin water layer that easily absorbs HCl, even trace concentrations of the latter, from the air, until the water layer is in equilibrium with the air. The equilibrium pressure of HCl under hydrochloric acid can be calculated as follows from a table of partial vapour pressures of aqueous solutions of HCl (Washburn, 1926–1930; 2003): log P ðmm HgÞ ¼ 0:1723  ðHCl; wt:%Þ  4:1531ðfor 20  CÞ; ð8aÞ 

log P ðmm HgÞ ¼ 0:1456  ðHCl; wt:%Þ  2:276ðfor 60 CÞ ð8bÞ

For example, air containing only 60 ppmv of HCl (the value measured in the cones) is in equilibrium with 6% (1.7 N) hydrochloric acid at 60 °C or 16% (4.7 N) at 20 °C. Thus, water in a thin layer on scoria particles turns in to hydrochloric acid that easily leaches cations from porous basalt. The leaching continues until intensive precipitation (30 mm rain during several hours) flushes leachate down to the underlying hot layer (Fig. 13b). Simultaneously, the temperature of the underlying hot layer decreases. After the precipitation is finished, the temperature increases back to what it was, and the leachate solution evaporates to form a mixture of Al and Ca chlorides and Na, K and Mg chloroferrates. Then, chloroferrates decay to produce molecular chlorine in the presence of hot air flux (Fig. 13c). Laboratory experiments showed that the oxidative decomposition of chloroferrates produces Cl2. The question is whether the above-described model really exists in the Tolbachik cones. No direct evidence was found to support the hypothesis that acid leaching and the subsequent decomposition of chloroferrates could contribute to observed chlorine emissions. However, some facts indicate that this process is possible. First, bleached basalts on the cones that cover more than 10,000 m2 of the cone summits (Fig. 2a) could provide a significant amount of chloroferrates through hydrochloric acid leaching. Second, powder of Fe2O3 and Mg and Al hydroxychlorides, which could be formed as a result of chloroferrate decomposition, is abundant within the subsurface layer at depths of 20– 40 cm (Fig. 2b). The possible contribution of chlorine from chloroferrate decay to the overall Tolbachik Cl2 emissions can be estimated from the total amount of leached iron from the cone summits, coupled with the “chlorine productivity of chloroferrates”. According to the average EDS profile of the outer layer of leached basalt (Fig. 11), approximately 75% Fe was extracted from 10,000 m2 of basalt within 35 years. With an average iron content of 8.1 wt.% in unaltered basalt (3 g/ cm3) and a 20 mm depth of alteration, about one ton of iron is extracted per year. On the basis of Eqs. (6a) and (6b), 1.5 mol Cl2 from 1 mol Fe, or 1.9 tons of Cl2 per year (in hypothetic case of 100% conversion) is expected. As chloroferrate-decay Cl2 is evolved only after heavy precipitations (according to the model of scoria leaching and flushing), this would create high chlorine concentrations. 4.4. Volcanic emissions of Cl2 Discharges of the chlorine-bearing gas at Tolbachik seem relatively insignificant on a global scale in terms of affecting the total tropospheric burden of chlorine (Graedel and Keene, 1995; Keene et al., 1999), possibly having some environmental effects only in close proximity to the cones. Winds disperse the gas and, even downwind from the cone summit, Cl2 concentrations rapidly decrease. Cl2 has a tropospheric lifetime of several hours (Khalil, 1999; Winterton, 2000) and during this time, much of the Cl2 reacts with H2 or VOC to form HCl or chlorinated organics. Within the chlorine – containing plume, tropospheric ozone loss can also occur (Platt and Ho¨nninger, 2003). Similarly,

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

223

Fig. 13. Illustration of the model of cyclic scoria flushing and chloroferrate formation/decomposition. (a) In wet weather, scoria particles are covered by a thin water layer that absorbs HCl from the percolating gas. (b) Heavy rain can flush the solution of HCl and leached cations down to the hot zone. (c) When the rain is finished, hydrochloric leachate in the hot zone dries to form a mixture of chlorides and chloroferrates; the latter decompose and evolve molecular chlorine Cl2.

emissions of Cl2 from volcanoes have little or no impact on the stratospheric ozone layer (WMO Report, 2006; von Glasow et al., 2009). Molecular chlorine is too reactive and has too short a lifetime to reach the upper horizons of the atmosphere from the boundary layer. The Tolbachik emissions of Cl2-rich gases represent an example of a new important natural Cl2 source in the continental environment, consisting of cooling and degassing

volcanic cones and lava flows and domes. Shortly after the eruption, the temperature inside a newly formed scoria cone or lava flow and the concentrations of volatiles can be high enough to form high amounts of the catalyst, leading to the intensive oxidation of HCl and the generation of molecular chlorine. In the case of passively degassing volcanoes, the considerable oxidation of HCl to Cl2 seems to be less likely.

224

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

Thermodynamic modelling shows that in a plume within a hot air-gas mixing zone at 600–1200 °C, HCl can be partially oxidised to Cl, ClO, OClO and Cl2, with equilibrium concentrations that are 3–6 orders of magnitude lower than the initial concentrations of HCl (Gerlach, 2004; Martin et al., 2006; Roberts et al., 2009). Sulphur dioxide, at least at concentrations that occur within the plume, does not suppress the oxidation of HCl at the hot zone, nor does it reduce Cl, ClOX and Cl2 back to HCl to a noticeable degree at lower temperatures (Ryan et al., 2006 and references therein; our calculations). In fact, several direct measurements have confirmed the existence of ClO and OClO in volcanic plumes (Lee et al., 2005; Bobrowski et al., 2007), although several more recent measurements did not (Kern et al., 2008; von Glasow, 2010). The behaviour of Cl2 in a volcanic gas–air mixture differs from that of other chlorine species. Equilibrium concentrations of Cl and ClOX in gas–air mixture decrease with decreasing temperature, while the oxidation of HCl to Cl2 is thermodynamically favoured at low temperatures, where the reaction drastically slows down. It seems that a major limitation of the oxidation of HCl to Cl2 in the plume arises from kinetics. In the absence of a catalyst, significant concentrations of Cl2 in the plume below 600–700 °C cannot be achieved (see Section 4.2.1). Even at 600–700 °C, the oxidation of HCl to Cl2 occurs rather slowly. According to Eq. (4), for example, it required 9 s for the conversion of 0.1% of HCl to Cl2 at 700 °C, and approximately 3 min at 600 °C. At the same time, few quiescently degassing volcanoes have gas temperatures exceeding 700 °C that could maintain high rates of HCl oxidation, and the temperature of a hot air–gas mixture rapidly decreases because of further dilution by ambient air. However, some amount of aerosol particles always exists in volcanic plumes, even in fumarolic plumes, thus providing surface area for heterogeneous catalysis (Mather et al., 2003; Pfeffer et al., 2006; Bobrowski et al., 2007; von Glasow, 2010). It is a question for further studies, whether these particles can effectively catalyse the oxidation of HCl below 600–700 °C. Even if they can, the suppressive effect of SO2 can become apparent, possibly leading to reduced concentrations of Fe and Cu oxides/chlorides because of their conversion into sulphates (Ryan et al., 2006; Delmelle et al., 2007). Conditions inside eruptive ash-laden plumes and pyroclastic flows are potentially favourable for the formation of Cl2 because of the high temperature and long lifetime (up to minutes) of the hot air-gas mixture. What is important is that ash can possess substantial catalytic activity that is essential for the oxidation of HCl at moderate (400– 700 °C) temperatures. In some cases, the existence of a catalyst inside the hot zone of the eruptive plume is obvious. The evidence is the bright red colour of the basaltic scoria, which is impregnated with and covered by fine crystals of Fe2O3. The red colour of the scoria can often be seen at volcanic cones, including the New Tolbachik cones (Fig. 1). It is possible that eruptive gases in the hot zone of the plume can be oxidised in the presence of a sufficient quantity of air and the catalytic influence of volcanic ash covered by fine Fe2O3. In this case, a substantial part of HCl can be converted into Cl2. This assumption needs further studies,

involving either experimental investigations or requiring in situ measurements of chlorine species in eruptive plumes. If the oxidation of HCl occurs in the high-temperature inner zone of an eruptive plume heavily loaded with thin ash (a possible catalyst), detecting the resultant Cl2 is challenging. Direct observations of the chemistry in ash-laden plumes originating from strong eruptions are scarce because ash strongly restricts both remote and airborne measurements of the plume composition (Andres and Rose, 1994; Andres and Schmid, 2001). Another issue is that spectroscopic measurements use specific banding of SO2, BrO and ClOX absorption spectra (for example, Kern et al., 2008; and references therein). The absorption spectrum of molecular chlorine does not have strong narrow bands and is essentially structureless (Maric et al., 1993). This prevents the detection of this gas in volcanic plumes by spectroscopic methods. Common chemical methods are also problematic. SO2 and Cl2 can exist simultaneously in the gas phase, but they interact almost instantly in a sample that is commonly a water solution, where SO2 forms sulphurous acid. According to Fogelman et al. (1989), the reaction time between SO2 and Cl2 in the solution varies from several milliseconds at neutral or low pH to several seconds at high pH (>11). To explore eruptive plumes and detect molecular chlorine in such harsh conditions, new remote or sampling techniques are required.

5. CONCLUSIONS

(1) Fumaroles at the Tolbachik scoria cones emit volcanic gas consisting primarily of atmospheric air. The gas contains less than 0.1 vol % acid species (CO2, HCl, HF) and molecular chlorine Cl2, species that have not been detected previously in volcanic gases. As much as 2–12 ppmv (6–36 mg/m3) Cl2 and 30– 80 ppmv HCl was measured in gas emissions from the 2nd North Cone, and up to 60 ppmv (180 mg/ m3) Cl2 was measured in emissions from the South Cone. These values are more than five orders of magnitude higher than concentrations of Cl2 previously measured in a marine boundary layer. (2) The main source of Cl2 in the Tolbachik gas is the catalytic oxidation of volcanic HCl by molecular oxygen in the air. Fine crystals of Fe and Cu oxides and chlorides on the altered basalt surface serve as catalysts. The catalytic oxidation of volcanic HCl in the presence of Fe and Cu oxides and chlorides within the scoria cones represents a case of abiogenic heterogeneous catalysis previously unknown in nature. This is an example of gas-rock interaction that affects volcanic gas composition. (3) Hydrochloric acid leaching of volcanic rocks is another type of gas-rock interaction producing unstable chemical compounds (Na, K and Mg chloroferrates) that may be considered as precursors of molecular chlorine. The oxidative decomposition of these chloroferrates can potentially create high concentrations of Cl2.

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

(4) The described emissions of Cl2 are too weak to have a significant influence on tropospheric and stratospheric composition, excluding the immediate vicinities of the cones. However, similar processes can potentially occur on other newly formed scoria cones, lava flows and domes. These volcanic manifestations should be considered as a potential new source of molecular chlorine in the continental environment.

ACKNOWLEDGEMENTS We are grateful to Vasilii Karandashev for ICP analyses, Tamara Sagalova and Salavat Khasanov for XRD studies, Slava Shapar for gas chromatography and Aleksey Nekrasov for electron micrographs and EDS analyses. We also thank Alessandro Aiuppa, Justin Filiberto, Tobias Fischer and anonymous reviewers, whose thoughtful and critical comments greatly improved the manuscript. This study was partially supported by RFBR Grant 07-05-00552 (Russia). We wish to dedicate this paper to the memory of Slava Shapar, a Russian volcanologist and friend, who left us in January, 2012.

REFERENCES Africano F. and Bernard A. (2000) Acid alteration in the fumarolic environment of Usu volcano, Hokkaido, Japan. J. Volcanol. Geoth. Res. 97, 475–495. Andres R. J. and Rose W. I. (1994) Remote sensing spectroscopy of volcanic plumes and clouds. In Monitoring Active Volcanoes: Strategies, Procedures and Techniques (ed. Bil McGuire). UCL Press, London, pp. 301–314. Andres R. J. and Schmid J. W. (2001) The effects of volcanic ash on COSPEC measurements. J. Volcanol. Geoth. Res. 108, 237–244. Bindeman I. N., Ponomareva V. V., Bailey J. C. and Valley J. W. (2004) Volcanic arc of Kamchatka: a province with high-d18O magma source and large-scale 18O/16O depletion of the upper crust. Geochim. Cosmochim. Acta 68, 841–865. Bobrowski N., Honninger G., Galle B. and Platt U. (2003) Detection of bromine monoxide in a volcanic plume. Nature 423, 273–276. Bobrowski N., von Glasow R., Aiuppa A., Inguaggiato S., Louban I., Ibrahim O. W. and Platt U. (2007) Reactive halogen chemistry in volcanic plumes. J. Geophys. Res. 112, 1–17. http:// dx.doi.org/10.1029/2006JD007206. Dai Y. D., Yu Z., Huang H. B., He Y., Shao T. and Hsia Y. F. (2003) Thermal decomposition of iron oxychloride as studied by thermal analysis, X-ray diffraction and Mo¨ssbauer spectroscopy. Mater. Chem. Phys. 79, 94–97. ´ skarssone Delmelle P., Lambert M., Dufreˆne Y., Gerind P. and O N. (2007) Gas/aerosol–ash interaction in volcanic plumes: new insights from surface analyses of fine ash particles. Earth Planet. Sci. Lett. 259, 159–170. Devine J. D., Sigurdsson H., Davis A. N. and Self S. (1984) Estimates of sulfur and chlorine yield to the atmosphere from volcanic eruptions and potential climatic effects. J. Geophys. Res. 89, 6309–6325. Fedotov S. A. and Markhinin Ye K. (Eds.), (1983) The Great Tolbachik Fissure Eruption: Geological and Geophysical Data, 1975–1976. Cambridge Univ. Press, Cambridge. 353p. Finley B. D. and Saltzman E. S. (2006) Measurement of Cl2 in coastal urban air. Geophys. Res. Lett. 33, L11809. http:// dx.doi.org/10.1029/2006GL025799.

225

Fischer T. P. (2008) Fluxes of volatiles (H2O, CO2, N2, Cl, F) from arc volcanoes. Geochem. J. 42, 21–38. Fogelman K. D., Walker D. M. and Margerum D. W. (1989) Nonmetal redox kinetics: hypochlorite and hypochlorous acid reactions with sulfite. Inorg. Chem. 28, 986–993. Francis P., Chaffin C., Maciejewski A. and Oppenheimer C. (1996) Remote determination of SiF4 in volcanic plumes: a new tool for volcano monitoring. Geophys. Res. Lett. 23, 249–252. Frumina N. S., Lisenko N. F. and Chernova M. A. (1983) Analytical Chemistry o Chlorine. Nauka, Moscow, 198p. (in Russian). Gavrilov A. P., Kochubei V. F. and Moin F. B. (1975) Kinetics of the oxidation of hydrogen chloride by oxygen. Kinet. Catal. 16, 778–780 (in Russian). Gerlach T. M. (1993) Oxygen buffering of Kilauea volcanic gases and the oxygen fugacity of Kilauea basalt. Geochim. Cosmochim. Acta 57, 795–814. Gerlach T. M. (2004) Volcanic sources of tropospheric ozonedepleting trace gases. Geochem. Geophys. Geosyst. 5, 1–16. http://dx.doi.org/10.1029/2004GC000747. Getahun A., Reed M. H. and Symonds R. B. (1996) Mount St. Augustine volcano fumarole wall rock alteration: mineralogy, zoning, composition and numerical models of its formation process. J. Volcanol. Geoth. Res. 71, 73–107. Giggenbach W. F. (1987) Redox processes governing the chemistry of fumarolic gas discharges from White Island, New Zealand. Appl. Geochem. 2, 143–161. Graedel T. E. and Keene W. C. (1995) Tropospheric budget of reactive chlorine. Global Biogeochem. Cy. 9, 47–77. Hisham M. W. M. and Benson S. W. (1995) Thermochemistry of the deacon process. J. Phys. Chem. 99, 6194–6198. Keene W. C., Khalil M. A. K., Erickson D. J., McCulloch A., Graedel T. E., Lobert J. M., Aucott M. L., Gong S.-L., Harper D. B., Kleiman G., Midgley P., Moore R. M., Seuzaret C., Sturges W. T., Benkovitz C. M., Koropalov V., Barrie L. A. and Li Y.-F. (1999) Composite global emissions of reactive chlorine from anthropogenic and natural sources: reactive chlorine emissions inventory. J. Geophys. Res. 104, 8429–8440. Keene W. C., Stutz J., Pszenny A. A. P., Maben J. R., Fischer E. V., Smith A. M., von Glasow R., Pechtl S., Sive B. C. and Varner R. K. (2007) Inorganic chlorine and bromine in coastal New England air during summer, J. Geophys. Res. 112, D10S12. http://dx.doi.org/10.1029/2006JD007689. Kern C., Sihler H., Vogel L., Rivera C., Herrera M. and Platt U. (2008) Halogen oxide measurements at Masaya Volcano, Nicaragua using active long path differential optical absorption spectroscopy. Bull. Volcanol. 71, 659–670. Khalil M. A. K. (1999) Reactive chlorine in the atmosphere. In Reactive Halogen Compounds in the Atmosphere (eds. P. Fabian and O. N. Singh). Springer-Verlag, Berlin Heidelberg, New York. pp. 45–79. Lee C., Kim Y. J., Tanimoto H., Bobrowski N., Platt U., Mori T., Yamamoto K. and Hong C. S. (2005) High ClO and ozone depletion observed in the plume of Sakurajima volcano, Japan. Geophys. Res. Lett. 32, 1–4. http://dx.doi.org/10.1029/ 2005GL023785. Maric D., Burrows J. P., Meller R. and Moortgat G. K. (1993) A study of the UV–visible absorption spectrum of molecular chlorine. J. Photoch. Photobio. A 70, 205–214. Martin R. S., Mather T. A. and Pyle D. M. (2006) Hightemperature mixtures of magmatic and atmospheric gases. Geochem. Geophys. Geosyst. 7, 1–14. http://dx.doi.org/10.1029/ 2005GC001186. Mather T. A., Pyle D. M. and Oppenheimer C. (2003) Tropospheric volcanic aerosol. In Volcanism and the Earth’s Atmosphere, Geophysical Monograph 139 (eds. A. Robock and C.

226

M. Zelenski, Y. Taran / Geochimica et Cosmochimica Acta 87 (2012) 210–226

Oppenheimer). Am. Geophys. Union, Washington, DC, pp. 189–212. Menyailov I. A., Nikitina L. P., Shapar V. N. (1980) Geochemical peculiarities of exhalations from the Great Tolbachik Fissure Eruption. Nauka, Moscow. 236 p. (in Russian). Naboko S.I. and Glavatskikh S.F. (1984) Post Eruptive Metasomatism and Ore-Forming Process. Nauka, Moscow. 165 p. (in Russian). Palin A. T. (1957) The determination of free and combined chlorine in water by use of diethyl-p-phenylene diamine. J. Am. Water Works Assoc. 49, 873–880. Pekov I.V. (2007) New Minerals from Former Soviet Union Countries, 1998–2006. Mineralogical Almanac, vol. 11. Moscow. 112 pp. Pfeffer M. A., Rietmeijer F. J. M., Brearley A. J. and Fischer T. P. (2006) Electron microbeam analyses of aerosol particles from the plume of Poas Volcano, Costa Rica and comparison with equilibrium plume chemistry modelling. J. Volcanol. Geoth. Res. 152, 174–188. Platt U. and Ho¨nninger G. (2003) The role of halogen species in the troposphere. Chemosphere 52, 325–338. Pszenny A. A. P., Keene W. C., Jacob D. J., Fan S., Maben J. R., Zetwo M. P., Springer-Young M. and Galloway J. N. (1999) Evidence of inorganic chlorine gases other than hydrogen chloride in marine surface air. Geophys. Res. Lett. 20, 699–702. Reed M. H. (1982) Calculation of multicomponent chemical equilibria and reaction processes in systems involving minerals, gases and an aqueous phase. Geochim. Cosmochim. Acta 46, 513–528. Roberts T. J., Braban C. F., Martin R. S., Oppenheimer C., Adams J. W., Cox R. A., Jones R. L. and Griffiths P. T. (2009) Modelling reactive halogen formation and ozone depletion in volcanic plumes. Chem. Geol. 263, 151–163. Roine A. (2007) HSC chemistry 6.1. Tech. rep. Outotec Research Oy. Pori, Finland. Ryan S. P., Li X. D., Gullett B. K., Lee C. W., Clayton M. and Touati A. (2006) Experimental study on the effect of SO2 on PCDD/F emissions: determination of the importance of gasphase versus solid-phase reactions in PCDD/F formation. Environ. Sci. Technol. 40, 7040–7047. Simpson W. R., von Glasow R., Riedel K., Anderson P., Ariya P., Bottenheim J., Burrows J., Carpenter L. J., Frieb U., Goodsite M. E., Heard D., Hutterli M., Jacobi H.-W., Kaleschke L., Neff B., Plane J., Platt U., Richter A., Roscoe H., Sander R., Shepson P., Sodeau J., Steffen A., Wagner T. and Wolff E. (2007) Halogens and their role in polar boundary-layer ozone depletion. Atmos. Chem. Phys. 7, 4375–4418. Spicer C. W., Chapman E. G., Finlayson-Pitts B. J., Plastridge R. A., Hubbe J. M., Fast J. D. and Berkowitz C. M. (1998) Unexpectedly high concentrations of molecular chlorine in coastal air. Nature 394, 353–356.

Symonds R. B., Rose W. I. and Reed M. H. (1988) Contribution of CI- and F-bearing gases to the atmosphere by volcanoes. Nature 334, 415–418. Symonds R. B., Reed M. H. and Rose W. I. (1992) Origin, speciation and fluxes of trace-element gases at Augustine volcano, Alaska: insights into magma degassing and fumarolic processes. Geochim. Cosmochim. Acta 56, 633–657. Symonds R. B., Rose W. I., Bluth, G. J. S. and Gerlach, T. M. (1994) Volcanic gas studies: methods, results, and applications. In Volatiles in Magmas (eds. M. R. Carroll and J. R. Holloway), Rev. Mineral. 30, 1–66. Tarabanko V. E., Tarabanko N. V., Zhyzhaev A. M. and Koropachinskaya N. V. (2009) A novel vanadium catalyst for oxidation of hydrogen chloride with dioxygen. J. Siber. Fed. Univ. (Chemistry) 1, 11–18. Taran Y. A. (1988) Geothermal Gas Geochemistry. Nauka, Moscow. 169 p. (in Russian). Vergasova L. P., Filatov S. K., Semenova T. F., Krivovichev S. V., Shuvalov R. R., Starova G. L. and Berlepsch P. (2000) Neue mineralien aus Kamchatka. Fumarolenmineralien vom vulkan Tolbachik, Rusland, 1996–1999; vorkommen, paragenese und hinwaise zur aufbewahrung. Lapis 1, 37–40 (in German). von Glasow R. (2010) Atmospheric chemistry in volcanic plumes. PNAS 107, 6594–6599. von Glasow R. and Crutzen P. J. (2003) Tropospheric Halogen Chemistry. In The Atmosphere (ed. R. F. Keeling). Treatise on Geochemistry, vol. 4 (eds. H. D. Holland and K. K. Turekian). Elsevier-Pergamon, Oxford. pp. 21–64. von Glasow R., Bobrowski N. and Kern C. (2009) The effects of volcanic eruptions on atmospheric chemistry. Chem. Geol. 263, 131–142. Washburn E. W. (1926–1930; 2003) International Critical Tables of Numerical Data, Physics, Chemistry and Technology, first electronic ed. Knovel. Winterton N. (2000) Chlorine: the only green element – towards a wider acceptance of its role in natural cycles. Green Chem. 2, 173–225. WMO, Scientific assessment of ozone depletion (2006) World Meteorological Organization, Global Ozone Research and Monitoring Project – Report 50. Zelenski M., Taran Yu., Shapar V., (2008) Emission of molecular chlorine from Earth’s volcanoes: a case of catalytic oxidation. in 10th IAVCEI Gas, Workshop, 10–20 November, 2008, Mexico (abstr). Zelenski M. E., Zubkova N. V., Pekov I. V., Boldyreva M. M., Pushcharovsky D. Yu. and Nekrasov A. N. (2011) Pseudolyonsite, Cu3(VO4)2, a new mineral species from the Tolbachik volcano, Kamchatka Peninsula, Russia. Eur. J. Mineral. 23, 475–481. Associate editor: Jun-ichiro Ishibashi