This page provides an overview of the state of the art in volcanic gas methods, with sections on Direct sampling, In situ plume measurement, Remote sensing and more to come (soon) by prominent members in CCVG
Contributed by Franco Tassi
Continuous measurements of gases emitted from volcanic system would provide information of fundamental importance for surveillance purposes. However, the hostile environmental conditions experienced by electronic devices in the presence of corrosive gases, makes the ‘survival’ of the analytical instruments difficult. Attempts to quantify fumarolic discharges in the field were successfully performed on specific gas species by using Dragger tubes (Tonani 1971), field gas-chromatographs (Le Guern et al. 1982, Zimmer and Erzinger 2003, Tedesco et al. 2005) and mass spectrometers (Naumann et al. 2001), MultiGAS. However, the periodic ‘direct’ gas sampling still represents a fundamental approach for extensive geochemical surveys of active volcanoes, spanning from main, minor, trace and ultra-trace compounds to isotopic investigation (Cioni and Symond et al. 1994; Giggenbach 1992, 1996; Panichi and La Ruffa 2001; Jordan 2003; Schwandner et al. 2004). Moreover, a “direct” connection to the fumarolic vent aimed to minimize air contamination allows the study of chemical and isotopic compositions of inert/noble gases in volcanic gas discharges. The presence of steam and acidic (highly reactive) gases, the latter being typically enriched in volcanic fumaroles, poses one of the most challenging problems of gas sampling and analysis (Giggenbach et al. 2001, Taran et al. 2006). In order to avoid secondary chemical reactions among the different species present in the gas phase, as well as to enhance the “analytically available” concentrations of minor and trace compounds, a pre-concentration procedure must be adopted in the field.
The basis for most “direct” sampling methods and analytical procedures currently in use was developed by Giggenbach (1975) and Giggenbach and Goguel (1989), which significantly improved the previous techniques developed using KOH (De Fiore 1926) or NH3 (Sicardi 1955). Gases from the fumarolic vent are transferred through “sampling lines” composed of non-reactive material into a pre-evacuated and pre-weighed glass bottle (known as Giggenbach’s bottle) equipped with Teflon stopcocks and filled with a certain amount of a highly alkaline (generally, 4 to 6M NaOH) solution, to promote acidic gases and steam dissolution. The gases that do not react with the alkaline solution will be enriched in the headspace of the gas vial. Some modifications have then been proposed for different purposes, e.g. analysis of metals in trace amounts (Sortino et al. 2006) and speciation of S-bearing gases (Piccardi, 1982; Montegrossi et al. 2001).
Gas-chromatography (GC) plays a dominant role for the analysis of the composition of the gases stored in the sample flask headspace. The commercial availability of different GC detectors and columns allows the analysis of mayor and minor gas low-solubility inorganic and organic gases. This technique, coupled with analytical methods (e.g., titration, ion-chromatography, potentiometry, spectrophotometry) for the determination of reactive, soluble gas species dissolved in the soda solution and in condensate samples, is thus able to provide the complete chemical composition of any fumarolic gas. Mass spectrometry, using a variety of instruments set at different conditions, offers the opportunity to analyze the isotopic ratios of a number of gases of interest (e.g. H2O, CO2, S-bearing compounds, halogens, noble gases, light hydrocarbons) for volcanological and geothermal purposes.
Cioni R. and Corazza E. (1981). Medium-temperature fumarolic gas sampling. Bull. Volcanol., 44, 23-29
De Fiore O. (1926). Studi sull’esalazione vulcanica. Nuove ricerche sui metodi d’analisi quantitative dei gas vulcanici. Bull. Vulcanol., 7-8, 271-301.
Giggenbach W. F. (1975). A simple method for the collection and analysis of volcanic gas samples. Bull. Volcanol., 39, 132-145.
Giggenbach W. F. (1992). Isotopic shifts in waters from geothermal and volcanic systems along convergent plate boundaries and their origin. Earth Planet. Sci. Lett., 113, 495-510.
Giggenbach W. F. (1996). Chemical composition of volcanic gases. In: R. Scarpa, R. I. Tilling (eds.), Monitoring and Mitigation of Volcano Hazards, Berlin, Springer Verlag, 221-256
Giggenbach W. F. and Goguel R. L. (1989). Collection and analysis of geothermal and volcanic water and gas discharges. New Zealand, Inst. Geol. Nucl. Sci., report n. CD2401, 36-53.
Giggenbach W. F., Tedesco D., Sulistiyo Y., Caprai A., Cioni R., Favara R., Fischer T. P., Hirabayashi J.-I., Korzhinsky M., Martini M., Menyailov I. and Shinohara H. (2001). Evaluation of results from the fourth and fifth iavcei field workshops on volcanic gases, Vulcano Island, Italy and Java, Indonesia. J. Volcanol. Geotherm. Res., 108, 157-172.
Le Guern F., Gerlach T. M. and Nohl A. (1982). Field gas chromatograph analyses of gases from a glowing dome at Merapi volcano, Java, Indonesia, 1977, 1978, 1979. J. Volcanol. Geotherm. Res., 14, 223-245.
Montegrossi G., Tassi F., Vaselli O., Buccianti A. and Garofalo K. (2001). Sulphur species in volcanic gases. Anal. Chem., 73, 3, 709-3,715.
Naumann D., Zimmer M., Erzinger J. and Wiersberg T. (2001). Gas monitoring,fluid flux and fluid sampling at well GPK-2 (Soultz-sous-Forets, France) – First results from the 5000 m production test. Geo-ForschungsZentrum Potsdam, Section 4.2, STR00/23, Geothermie Report 00-1 71, 14.
Panichi C. and La Ruffa G. (2001). Stable isotope geochemistry of fumaroles: an insight into volcanic surveillance. J. Geodynamics, 32, 519-542.
Piccardi G. (1982). Fumarole gas collection and analysis. Bull. Volcanol., 45, 257-260.
Schwandner F. M., Seward T. M., Gize A. P., Hall P. A. and Dietrich V. J. (2004). Diffuse emission of organic trace gases from the flank and crater of a quiescent active volcano (Vulcano, Aeolian Islands, Italy). J. Geophys. Res. D -Atmospheres, 109, D04301, doi:10.1029/2003JD003890.
Sicardi L. (1955). Captazione ed analisi chimica dei gas della esalazione solfidrico-solforosa dei vulcani in fase solfatarica. Bull. Volcanol., 18, 107-112.
Sortino F., Nonell A., Toutain J. P., Munoz M., Valladon M. and Volpicelli G. (2006). A new method for sampling fumarolic gases: analysis of major, minor and metallic trace elements with ammonia solution. J. Volcanol. Geotherm. Res., 158, 244-256.
Taran Y., Inguaggiato S. and Fisher T. (2006). Evaluation of results from the 8th workshop. Commission on the Chemistry of Volcanic Gases, Newsletter, 19, 6-13.
Tedesco D., Castrillo A., Vaselli O., Gianfrani L. (2005). Method allows for continuous monitoring of volcanic gases. EOS, 86, 510-511.
Tonani F. (1971). Concepts and techniques for the geochemical forecasting of volcanic eruptions. In: The surveillance and prediction of volcanic activity. A review of methods and techniques, Paris, unesco, 145-166.
Zimmer M. and Erzinger J. (2003). Continuous H2O, CO2, 222Rn and temperature measurements on Merapi Volcano, Indonesia. J. Volcanol. Geotherm. Res., 125, 25-38.
Remote Sensing of Volcanic Gases
Contributed by Nicole Bobrowski
Although today it seems hard to imagine volcanic gas measurements without thinking of routine SO2 gas flux measurements, yet remote sensing of volcanic gases has a history of less than half a century. Only in 1971 at Mt Mihara (Japan), the first Correlation Spectrometer (COSPEC) measurements were carried out to determine the volcanic SO2 output (Moffat et al., 1972), which has since become a very broadly applied technique, including for monitoring. The 1982 eruption of El Chichón (Mexico) was the first volcanic eruption during which SO2 was detected from space (Krueger 1983). Remote sensing combines several advantages – it allows (a) near continuous measurements, (b) real-time evaluations, (c) non – contact measurements, (d) direct identification of trace gas molecules and overall (e) it allows measurements to be conducted from a safe distance from the volcano, also allowing the possibility to observe explosive eruptions. Remote sensing of volcanic gases can be carried out from different platforms, from the ground up to space, the latter that provides the opportunity to detect volcanic degassing at volcanoes around the planet on a daily basis.
A variety of remote sensing techniques for measuring volcanic plumes have been developed during the last few decades. These techniques can be grouped into dispersive (e.g. DOAS) and non-dispersive (e.g. SO2 Camera) approaches and into active (using their own artificial radiation source e.g. LIDAR) and passive (using natural sources of radiation e.g. sunlight or the thermal emission of the gas to be measured) techniques (Platt et al., 2015). Remote sensing measurements rely on the analysis of radiation emitted, absorbed, or scattered by gases and aerosols in volcanic plumes. Dispersive techniques determine a complete wavelength-dependent intensity spectrum for each pixel. The trace gas column density (TGCD) for that particular pixel is then derived from the measured spectrum. Non-dispersive approaches derive the TGCD from a small number (typically one) integrated intensity measurement at suitable wavelengths. Frequently, a reference intensity at one (or very few) different wavelength(s) is recorded and the TGCD is derived from the ratio of the two intensities (or as a function of a few intensity measurements).
Dispersive methods are usually more complex and slower than non-dispersive techniques. Their advantage lies in the fact that data collected with higher spectral resolution allow a much better identification of a particular trace gas and a more accurate removal of interferences from other gas species. Furthermore, analyses done using traceable, well-quantified absorption cross-sections increases the rigour of the technique. Many gas absorption features are naturally narrow, particularly for small molecules, and this limits the effectiveness of non-dispersive techniques. In fact, non-dispersive techniques have been only applied to major components (like SO2) of volcanic plumes, while trace species (like BrO) have been detected by dispersive approaches (Platt et al., 2015).
Remote sensing of volcanic gases has developed far beyond SO2 emission rate measurements. The first field portable FTIR instruments were developed in the beginning of the nineties for the detection and quantification of other volcanic gases including: HCl, H2O, SO2, HF, CO2, SiF4, OCS and CO (e.g. Notsu et al., 1993, Mori et al., 1993, Mori et al., 1997). In fact, an important aim of present research is to develop techniques to detect an even broader range of molecules and to image in two and three dimensions the distribution of volcanic gases within plumes. Despite impressive accomplishments, many techniques are still in an early stage of development and much progress can be expected in the near future.
The main disadvantages of remote sensing techniques for quantifying volcanic gas emissions are (1) the limited number of species determined, (2) the complex analysis for most species, and (3) the challenge of differentiating isotopes. In particular, two volcanic activity-driving gases, CO2 and H2O, are hard to measure due to their high abundance in ambient air and in the case of water its high variability over time and space. However, some successful remote sensing measurements for CO2 have been carried out (e.g. Goff et al., 2001).
A further challenge of measuring volcanic gases at a greater distance down-wind from the emission source is the possible chemical conversion of reactive gas species that can only be corrected when the complex mixing and reactions between volcanic gases and the surrounding atmosphere are well constrained.
In addition to modern technical advances in the hardware offering robust, relatively small and cheap spectrometers, much has been done to improve data analysis software, and work is currently underway to improve imaging processing. The latter, which will help improve our understanding of the heterogeneity inside volcanic plumes, the plume–atmospheric mixing and chemical transformation processes. Also long recognized problems, like radiative transfer in volcanic plumes and the atmosphere, remain to be solved, for example through improved radiative transfer modelling for the accurate determination of gas fluxes.
Goff, F., Love, S. P., Warren, R. G., Counce, D., Obenholzner, J., Siebe, C., & Schmidt, S. C. (2001). Passive infrared remote sensing evidence for large, intermittent CO 2 emissions at Popocatépetl volcano, Mexico. Chemical Geology, 177(1), 133-156.
Krueger, A. J. (1983). Sighting of El Chichon sulfur dioxide clouds with the Nimbus 7 total ozone mapping spectrometer. Science, 220(4604), 1377-1379.
Moffat, A. J., Kakara, T., Akimoto, T. and Langan, L. (1972) Air Note. Environmental Measurements, San Francisco.
Platt U, Lübcke P, Kuhn J, Bobrowski N, Prata F, Burton M, Kern C (2015) Quantitative imaging of volcanic plumes — Results, needs, and future trends. J Volcanol Geotherm Res 300:7–21. doi: 10.1016/j.jvolgeores.2014.10.006
Isotopic evaluation of volcanic gases
Contributed by Andrea Rizzo and Maarten de Moor
The term isotope derives from Greek (ἴσος “equal” and τόπος “place”) and means that atoms of a particular element, whose nuclei contain the same number of protons but a different number of neutrons, occupy the same position on the periodic table. In nature we can distinguish between radioactive and stable isotopes (Hoefs, 2004). The former provide powerful tracers for studying the ages and origins of Earth systems, and have yielded a revolution in the comprehension of geological processes (Dickin, 2005; Allègre, 2008); the latter are fundamental for geochemical investigations in the broad sense to assess the origin of fluids as well as to understand the modifying the pristine isotopic composition (i.e., magmatic degassing, crustal and atmospheric contamination, gas-water interaction, etc…).
The study of volcanic gases is based on the measurement and interpretation of concentrations and isotopic ratios of the main gaseous species (H2O, CO2, SO2, H2S, HCl, N2) and of a few key trace species such as noble gases (especially He, Ne, and Ar). The natural variations in the isotope ratios of these species are extremely slight, and for major gases are generally expressed in d unit permil (‰) (e.g., Hoefs, 2004). On the other hand, noble gases are expressed as raw ratios with 3He/4He of sample (R) commonly normalized to the same ratio in atmosphere (Ra) (Ozima and Podosek, 2002).
Since the middle of the previous century, a number of geochemical studies of volcanic and geothermal systems revealed that the isotope composition of steam H2O (d18O and dD) can be used to distinguish magma-derived water from meteoric, marine, hydrothermal and groundwater contributions, as well as their relative extents of mixing (e.g., Craig, 1961, 1963; Giggenbach and Stewart, 1982; Giggenbach, 1992; Panichi and La Ruffa, 2001). These early studies sparked more dedicated research on the isotopic compositions of volcanic gas emissions, and isotopes now form a fundamental source of information in volcanic gas chemistry.
The isotopic ratios of carbon (d13C), sulphur (d34S), nitrogen (d15N), chlorine (d37Cl), and noble gases (mostly 3He/4He, 4He/20Ne, 20Ne/22Ne, 21Ne/22Ne, 40Ar/36Ar) are principally used to identify the relative contributions from different volatiles sources, which are intricately related to geodynamic setting (Hoefs, 2004; Dickin, 2005; Allègre, 2008; Burnard, 2012). Considering that CO2 is generally the dominant component of volcanic dry gases, d13CCO2 has a long history of investigation (Naughton and Terada, 1954; Wasserburg et al., 1963; Allard et al., 1977) and is particularly useful to evaluate the extent of magmatic degassing. The isotopic fractionation along this process can be modeled assuming batch- and fractional-equilibrium conditions (Rayleigh, 1896; Taylor, 1986; Gerlach and Taylor, 1990; Albarède, 1995; Hoefs, 2004). The d13CCO2 is often coupled to CO2/3He to distinguish proportions of carbon originating from magma, sediments and limestones (Sano and Marty, 1995). Recently, d13CCO2 is studied together with d15N of N2 which is sensible to the presence of sediments especially in subduction-related environments (e.g., Fischer et al., 2002).
The isotope fractionation is a mass-dependent process in which the partitioning of isotopes occurs differently for kinetic and equilibrium reactions (e.g., Hoefs, 2004). This process can occur during magmatic degassing, differentiation processes, interaction with hydrothermal systems, and precipitation of phases from volcanic gases. In all these processes, the fractionation mostly affects isotope systematics with greater mass difference between the light and heavy isotope, and those elements with multiple oxidation states. Thus, light elements such as D/H in water (Giggenbach, 1992), and multi-valent elements such as S in volcanic-hydrothermal systems can show large isotopic variation due to fractionation (Marini et al., 2013). “Seeing through” fractionation processes to understand variations in source compositions can thus require extensive characterization and modelling of the magmatic/hydrothermal system in question (e.g., Liotta et al., 2012; de Moor et al., 2013).
Most of the geochemical studies aimed at initial characterization of a volcanic system involve the use of noble gases isotopes, which are among the most powerful tracers in fluid geochemistry according to their chemical and physical properties (e.g., Porcelli et al., 2002; Burnard, 2012). In particular, studies from the 1960’s (Clarke et al.,1969; Mamyrin et al.,1969; Tolstikhin et al.,1974; Kurz and Jenkins, 1981) showed that 3He/4He allows unambiguous assessments of gas origin because of the different provenance of the two isotopes (Ozima and Podosek, 2002). The study of 3He/4He in volcanic gases also permits the definition of the magmatic/mantle source features, the evaluation of contaminations at crustal levels and gives clear signals of magma ascent in volcanic plumbing systems when monitored over time (e.g., Sano et al., 1995, 1997, 2015; Martelli et al., 2008; Rizzo et al., 2015a). On the other hand, 4He/40Ar* is a diagnostic tracer of magmatic degassing (e.g., Moreira and Sarda, 2000), while 4He/20Ne and 40Ar/36Ar are used to evaluate the extent of atmospheric contamination.
At present, the most common technique used to make isotopic measurements is still conventional isotope-ratio mass spectrometry (IRMS), which requires that gases are directly sampled in the field and then brought to the laboratory for processing and analysis. Even if this method provides a good understanding of isotopes behavior in active volcanic systems, it suffers the limitation that only a restricted number of samples can be analyzed over time resulting in a medium- to long-term information for volcanic surveillance. New techniques aimed at increasing the frequency of isotopic measurements of carbon dioxide, methane and water vapor at atmospheric concentrations have been developed in the last decades involving the use of laser-based isotope-ratio infrared spectrometer (IRIS) and cavity ring-down spectroscopy (CRDS)(e.g., O’Keefe and Deacon, 1988; Murnick and Peer, 1994). These techniques allow direct sampling and in-situ analysis of gas that is pumped inside an analyzer. The most recent and successful attempts of these applications to active volcanoes regarded the real-time measurements of the CO2 concentration and δ13C in fumarolic and plume gases emitted at Mount Etna (Italy; Rizzo et al., 2014, 2015b), at Turrialba (Costa Rica; Malowany et al., 2014) and at Long Valley (USA; Lucic et al., 2015). Although IRIS and CRDS techniques represent a significant advancement toward real-time volcano monitoring, development of permanent installations at the top of volcanoes are needed in future work.
When this next step forward becomes possible, our understanding of magmatic degassing and volcanic plumbing systems will most likely increase dramatically, thus yielding enhanced capacity for eruption forecasting.
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de Moor J. M., Fischer T. P., Sharp Z. D., King P. L., Wilke M., Botcharnikov R. E., Cottrell E., Zelenski M., Marty B., Klimm K., Rivard C., Ayalew D., Ramirez C., Kelley K.A, 2013. Sulfur degassing at Erta Ale (Ethiopia) and Masaya (Nicaragua) volcanoes: Implications for degassing processes and oxygen fugacities of basaltic systems. Geochemistry, Geophysics, Geosystems 14, 1525-2027. doi: 10.1002/ggge.20255.
Dickin A.P., 2005. Radiogenic Isotope Geology (Cambridge University Press).
Fischer, T. P., Hilton, D. R., Zimmer, M. M., Shaw, A. M., Sharp, Z. D., Walker, J. A., 2002. Subduction and recycling of nitrogen along the Central American margin. Science 297, 1154–1157.
Gerlach, T. M. and Taylor, B.E., 1990. Carbon isotope constraints on degassing of carbon dioxide from Kilauea Volcano, Geochim. Cosmochim. Ac., 54, 2051–2058.
Giggenbach W. F. (1992). Isotopic shifts in waters from geothermal and volcanic systems along convergent plate boundaries and their origin. Earth Planet. Sci. Lett., 113, 495-510.
Giggenbach, W. F. and Stewart, M. K. (1982) Processes controlling the isotopic composition of steam and water discharges from steam vents and steam-heated pools in geothermal areas. Geothermics 11, 71–80.
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Liotta M., A. Rizzo, A. Paonita, A. Caracausi, and M. Martelli (2012), Sulfur isotopic compositions of fumarolic and plume gases at Mount Etna (Italy) and inferences on their magmatic source, Geochem. Geophys. Geosyst., 13, Q05015, doi:10.1029/2012GC004118.
Lucic, G., Stix, J., and Wing, B. , 2015. Structural controls on the emission of magmatic cabon dioxide gas, Long Valley caldera, USA., J. Geophys. Res.-Sol. Ea., 120, 2262–2278, 25 doi:10.1002/2014JB011760.
Malowany, K., Stix, J., and de Moor, J. M., 2014. Field measurements of the isotopic composition of carbon dioxide in a volcanic plume and its applications for characterizing an active volcanic system, Turrialba volcano, Costa Rica, in: CCVG-IAVCEI 12th Field Workshop on Volcanic Gases, Northern Chile, 17–25 November.
Mamyrin, B. A., Tolstikhin, I. N., Anufriyev, G. S. and Kamenskiy, I. L., 1969. Anomalous isotopic composition of helium in volcanic gases. Dokl. Akad. Nauka SSSR 184, 1197-1199.
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Martelli, M., Caracausi, A., Paonita, A. & Rizzo, A., 2008. Geochemical variations of air-free crater fumaroles at Mt Etna: New inferences for forecasting shallow volcanic activity. Geophys. Res. Lett. 35, L21302; doi: 10.1029/2008GL035118.
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Fluxes from soil diffuse degassing
Contributed by Franco Tassi and Carlo Cardellini
Investigations aimed to the estimation of gas fluxes, especially CO2, released to the atmosphere through diffuse emission from the soil are of fundamental importance for the evaluation of the gas budget in volcanic and hydrothermal systems. At Campi Flegrei (southern Italy), for example, diffuse degassing represents the main gas release and significantly contributes to the heat transfer from the hydrothermal system (e.g. Chiodini et al. 2005, 2012).
Indirect methods for the determination of soil gas fluxes are based on the analysis of gas concentrations at different depths in the soil. However, this approach assumes steady state diffusive fluxes and a correct estimation of soil porosity. Hence, direct methods, generally consisting of dynamic and static procedures, are to be preferred. The accumulation chamber (AC) method (e.g. Sorey et al. 1998; Chiodini et al. 1996, 1998, 2001; Gerlach et al. 2001) is a dynamic procedure commonly used to measure fCO2. The AC apparatus consists of: 1) a metal cylindrical chamber placed on the ground, 2) a CO2 detector, usually consisting of a Infra-Red (IR) spectrophotometer, 3) an analogue-digital (AD) converter, and 4) a palmtop computer. A low-flux pump (in the order of 20 mL s-1) continuously transfers the soil gas from the chamber to the IR for the continuous measurement of the CO2 concentrations. To minimize the disturbance effects due to changes of barometric conditions, the soil gas continuously circulated from the chamber to the detector.
The CO2 flux (fCO2) values are computed on the basis of the measured CO2 concentration inside the chamber over time (dCCO2/dt), according to the following equation:
fCO2 = cf × dCCO2/dt. (1)
The proportionality factor (cf) between dCCO2/dt and the fCO2 is theoretically (e.g., Chiodini et al. 1998) proportional to height of the chamber, but has to be determined for any specific measuring apparatus by laboratory tests. The laboratory tests consist in repeated measurements of different known gas fluxes (in a proper range of values), that allow to compute the cf of a specific apparatus as the slope of the best-fit line to the imposed CO2 flux versus measured dCCO2/dt.
A different approach (Lewicki et al. 2005) has been used to calculated fCO2 (g m−2 day−1) considering the theoretical relation:
fCO2 = k × V/A × T0/T × P/P0 × dCCO2/dt. (2)
where k is a constant (1,558,656 g s m−3 day−1), T and T0 are measured and standard temperature (K), respectively, P and P0 are measured and standard pressure (kPa), respectively, V is the system volume (m3), A is the accumulation chamber footprint area (m2), and dCCO2/dt is the initial rate of change of CCO2 in the chamber after the chamber is placed on the soil surface (vol % s−1).
The cf determined theoretically using Eq. (2) (cf = k × V/A × T0/T × P/P0) resulted higher for specific measuring apparatus when compared to that determined by experimental laboratory test (Lewicki et al. 2005), suggesting that real measurement deviates from the theoretical one.
Other gas compounds (e.g. CH4, Hg0) have been successfully measured using this method (Etiope 1997; Cardellini et al. 2003a; Bagnato et al. 2014).
Alternatively, diffuse fluxes from the soil of hydrothermal gas, including fCH4, fC6H6, Hg0, can be measured using a static approach, the so-called closed-chamber (CC) method (Hutchinson and Mosier 1981; Rolston 1986; Livingstone and Hutchinson 1995), which was developed in agricultural sciences to determine soil respiration (e.g. Parkinson 1981; Raich et al. 1990; Smith et al. 1995; Dueñas et al. 1996; Tsuyuzaki et al. 2001; Huttunen et al. 2003; Hirota et al. 2004), whereas it has been less frequently used in geothermal and volcanic environments (Klusman and LeRoy 1996, Etiope 1999; Klusman et al. 2000; Castaldi and Tedesco 2005; D’Alessandro et al. 2009; Tassi et al. 2011, 2013).
The CC equipment consists of a cylindrical chamber equipped with a three-way valve on its top. Once a chamber was positioned on the ground, an aliquot (5 to 10 cc) of gas is collected from the chamber at fixed time intervals using a syringe, and transferred into sampling vials equipped with a rubber septum for subsequent analyses. The variation in time of the CH4 and C6H6 concentrations in the chamber is proportional to fCH4 and fC6H6, respectively, as described by the following equation:
f(X) = dCx/dt × V/A (3)
where dCx/dt is the rate of concentration change of the X gas compound within the chamber positioned on the ground, whereas V and A are the volume and the basal area of the chamber, respectively. The dCx/dt value is calculated from the linear regression of the concentrations of the X compound in samples collected from the chamber starting from zero time.
Gas fluxes from soil can be effectively shown using contour maps constructed by using interpolation algorithms, generally kriging (e.g. Bergfeld et al. 2001; Chiodini et al. 2001; Rogie et al. 2001), which provide an estimate of a variable without specific regard to the resulting spatial statistics of all the estimates taken together (Deutsch and Journel 1998). The kriging interpolation is considered to smooth out the extreme extrapolated values, with small values being overestimated, whereas large values, i.e. those that define degassing structures, are underestimated.
The total soil gas emission can be calculated either by multiplying the arithmetic mean value of the gas fluxes by the investigated areas, or by applying volume or by a Graphical Statistical Approach (GSA; Chiodini et al. 1998). Using the kriging interpolation, the estimation of local accuracy is incomplete, except if a Gaussian model for errors is assumed, the GSA approach allows the definition of a confidence interval for the estimation, but it does not take into account the spatial correlation of the data, thus causing an overestimation of the uncertainty.
For mapping purposes and to estimate the total CO2 release from an investigated area, stochastic simulation algorithms and in particular Sequential Gaussian Simulation (SGS, Deutsch and Journel 1998; Cardellini et al. 2003b) are generally used. SGS allows the generation of a set of equi-probable representations of the spatial distribution of the gas flux values, able to reasonably reproduce their global statistic and spatial features. The uncertainty is thus related to the differences among many simulated maps (Goovaerts 2001). We suggest that the sequential Gaussian simulation method yields the most realistic representation of the spatial distribution of CO2 fluxes and allows the determination of a reliable uncertainty of the estimation of the total CO2 release from an investigated area.
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