Geochemistry of CO2-Rich Gases Venting From Submarine Volcanism: The Case of Kolumbo (Hellenic Volcanic Arc, Greece)

Introduction

About 80% of Earth volcanism occurs on the ocean floor (Crisp, 1984), which has greatly hindered the understanding of the natural outgassing of volatiles from the Earth’s interior and its impact on the environment. Several submarine volcanoes have been discovered in recent decades, and geochemical studies have investigated the current status of submarine magmatic-hydrothermal systems. These studies have focused either on midocean ridges (e.g., Butterfield et al., 1990; Lilley et al., 1993; Von Damm, 1995; Von Damm et al., 1995; Lupton et al., 1999) or subduction-related settings (Taran et al., 1992; Tsunogai et al., 1994; Caracausi et al., 2005; Chiodini et al., 2006; Lupton et al., 2006, 2008; Lan et al., 2010; Rizzo et al., 2016; Stucker et al., 2017; and references therein). To the best of our knowledge, Panarea (Aeolian Islands, Italy) is the only volcanic system in the Mediterranean basin characterized by active submarine emissions for which geochemical studies have provided detailed and comprehensive reconstructions of the gas–water interaction process, the origin of gases and the magmatic-hydrothermal system (Caliro et al., 2004; Caracausi et al., 2005 and references therein; Chiodini et al., 2006; Capaccioni et al., 2007; Tassi et al., 2009, 2014; and references therein). The submarine emissions at Panarea are mainly located at relatively shallow water depths (down to 30 m), thereby providing easy access for gas sampling.

In 2006, an extensive hydrothermal vent field was discovered at a depth of 500 m on the floor of Kolumbo submarine crater (Sigurdsson et al., 2006), 7 km off the northeast coast of the Santorini Island (Figure 1) in the southern Aegean Sea, Greece (Nomikou et al., 2012). Kolumbo volcano lies along the Christianna-Santorini-Kolumbo volcanic line (CSK; Figure 1) that is in the central part of the Hellenic Volcanic Arc (HVA; Nomikou et al., 2016, 2018). The CSK consists of 23 submarine cones and craters (Nomikou et al., 2012; Hooft et al., 2017), of which Kolumbo is the largest and most-active, and last erupted in 1650 (Cantner et al., 2014). The crater is 1.7 km in diameter and up to 500 m deep, with the shallowest point nowadays being at 18 m b.s.l. (below sea level) (Nomikou et al., 2012). All of these volcanoes belong to the modern HVA formed by the subduction of the African plate beneath the European plate (McKenzie, 1972; Le Pichon and Angelier, 1979). HVA began to form 3–4 My ago (Pe-Piper and Piper, 2007), and it stretches from the Gulf of Saronikos in the northwest to the Kos-Nisyros-Yali Islands complex in the east (Figure 1).

FIGURE 1

Figure 1. Offshore and onshore topographic map of the Aegean Sea and the location of the main eruptive centers belonging to the modern HVA (Hellenic Volcanic Arc) (modified after Nomikou et al., 2013). Bathymetry is from the EMODNET database (www.emodnet.eu). CSK indicates the Christianna-Santorini-Kolumbo volcanic line. The inset map shows the main geodynamics of the Aegean Sea (Nomikou et al., 2018).

The Kolumbo hydrothermal vent field emits CO₂-dominated fluids at temperatures as high as 220°C (Sigurdsson et al., 2006; Carey et al., 2013) and with a clear mantle origin (Rizzo et al., 2016). However, the origin of CO₂ and CH₄, the extent of variability in the geochemistry of gases emitted from different chimneys, and the pressure and temperature conditions of the magmatic-hydrothermal system remain unclear.

This study aims to fill this gap in the knowledge by combining the chemical and isotopic data previously reported for noble gases by Rizzo et al. (2016) with new chemical data from the analysis of Hg(0) and the isotopic compositions of CO₂ and CH₄. We identified and modeled the main processes underlying the alterations in the isotopic composition of magmatic-hydrothermal fluids along the path of their ascent, in order to reconstruct the original chemistry of these gases at the earliest stages of their formation. Finally, we used geo-indicators to elucidate the pressure and temperature conditions of the hydrothermal system feeding the gas vents beneath Kolumbo volcano, and propose a conceptual scheme for the fluid circulation.

Present Knowledge About Kolumbo Submarine Volcano

The Kolumbo submarine volcano developed next to Santorini Island (Figure 1), and several multidisciplinary studies have been carried out over the last decade to shed light on its plumbing system and activity state. The findings of volcanological and geochemical investigations suggest that there are two distinct plumbing systems beneath the Santorini and Kolumbo volcanic systems (Francalanci et al., 2005; Dimitriadis et al., 2009; Kilias et al., 2013). The depth of the magma chamber beneath Kolumbo has been constrained by seismological and petrological observations to lie at depths of 5–7 km (Dimitriadis et al., 2009, 2010; Konstantinou and Yeh, 2012; Cantner et al., 2014). A more recent petrological study revealed striking geochemical differences (e.g., in Nb/Yb, Zr/Nb, ²⁰⁶Pb/²⁰⁴Pb, ⁸⁷Sr/⁸⁶Sr, and ³He/⁴He) between Kolumbo and Santorini magmas despite their close temporal and spatial associations, supporting the hypothesis that the two magmatic systems have distinct mantle sources (Klaver et al., 2016).

Modern-day microseismicity along the CSK is concentrated beneath Kolumbo at depths of 6–9 km (Bohnhoff et al., 2006; Dimitriadis et al., 2009), with the exception of the unrest at Santorini during 2011–2012 when the seismicity focus migrated within its caldera (e.g., Parks et al., 2012). This confirms that Kolumbo is the most-active volcanic system in the region at the present time (Francalanci et al., 2005; Dimitriadis et al., 2009, 2010; Nomikou et al., 2012; Hubscher et al., 2015). Similar indications come from the presence of intense degassing of hydrothermal vents on the floor of Kolumbo submarine crater and the geochemistry of these fluids (Sigurdsson et al., 2006; Carey et al., 2013; Rizzo et al., 2016). This intense degassing contrasts with the low-temperature fumaroles observed in the Santorini caldera (Sigurdsson et al., 2006; Rizzo et al., 2015; and references therein).

Sigurdsson et al. (2006) were the first to describe the presence of a widespread hydrothermal vent field on the floor of Kolumbo submarine crater, but only two subsequent geochemical studies have investigated the chemistry of these fluids (Carey et al., 2013; Rizzo et al., 2016). Sigurdsson et al. (2006) and Carey et al. (2013) reported that fluids emitted from the Kolumbo floor are at temperatures up to 220°C and pH ∼5 next to the vents, respectively. Carey et al. (2013) also reported the first data for the chemical composition of gas samples collected from two vents, which indicated that these gases are virtually pure CO₂. However, that study focused on the acidification-related hazards that may arise from the dissolution of CO₂ in seawater, which was found to occur within the first 10 m of the water column inside the bowl-shaped crater. This dissolution causes local increases in water density and favors the accumulation of CO₂-rich, acidic waters at the seafloor that do not permit the growth of macrofauna (Camilli et al., 2015), while additional hazards could arise from the abrupt release of gases at the surface.

Rizzo et al. (2016) reported on ³He/⁴He measurements of gas samples collected during May 2014 from seven chimneys. Briefly, they constrained the ³He/⁴He signature of Kolumbo gases and of the local mantle at 7.0 Ra (where Ra is the atmospheric ³He/⁴He equal to 1.39 × 10^-6), which is indicative of a MORB-like mantle. It was subsequently revealed that these values are significantly lower than the ³He/⁴He values measured in Santorini fluids and rocks (∼4 Ra; Rizzo et al., 2015 and references therein), whereas the observed decrease in the He-isotope signature was attributable to a crustal contamination below the Island. Finally, that study highlighted that ³He/⁴He values measured at Kolumbo were the highest ever measured across the entire HAV and indicative of the direct degassing through lithospheric faults.

Sampling and Analytical Techniques

During the 4-SeaBioTech survey on RV AEGAEO (Hellenic Centre for Marine Research) during May 2014, seven chimneys bubbling gas phases with variable sustained fluxes were sampled on the floor of Kolumbo submarine crater with the Greek Max Rover remotely operated vehicle (ROV) (Figure 2). These chimneys were selected based on a previous seafloor exploration of hydrothermal activity (Sigurdsson et al., 2006; Carey et al., 2013) and a oceanographic survey performed in 2014 (Figure 2). Although direct measurements of temperature were not possible during the 2014 survey, previous surveys found that the fluids discharged in the northern part of the crater floor had temperatures as high as 220°C, while those present along the northern and eastern margins were no hotter than 70°C, and characterized by ephemeral bubble fluxes (Sigurdsson et al., 2006; Carey et al., 2013).

FIGURE 2

Figure 2. Bathymetric map of the floor of Kolumbo submarine crater and the location of the sampled hydrothermal vents (modified after Kilias et al., 2013). Inset map shows the swathe map of Kolumbo volcano (Rizzo et al., 2016). Red rectangular indicates the location of the vent field.

The present study collected bubbling hydrothermal gases at the seafloor and stored them in titanium gas-tight bottles equipped with funnels, as described in detail by Rizzo et al. (2016). The obtained gas samples were analyzed in the laboratories of INGV (Istituto Nazionale di Geofisica e Vulcanologia), Sezione di Palermo for their chemistry and isotopic compositions of noble gases (He, Ne, and Ar), C of CO₂, and C and H of CH₄. The chemical composition of He, H₂, O₂, N₂, CO, CH₄, and CO₂ was measured by a gas chromatograph (Clarus 500, Perkin Elmer) equipped with a 3.5-m column (Carboxen 1000) and double detector (hot-wire detector and flame ionization detector [FID]), for which the analytical errors were < 3%. The concentrations of C₂H₆ and C₃H₈ were also measured in a few selected samples after hydrocarbon enrichment via bubbling pressurized gas in Giggenbach bottles filled with 4-M NaOH (Giggenbach, 1975). Higher hydrocarbons were analyzed using a gas chromatograph (Shimadzu, 2010) equipped with an FID and a capillary column (CP Poraplot) using He as the carrier gas. The analytical precision for these gas chromatography analyses was better than 5% for trace gases and 10% for alkanes.

The Hg(0) concentration in the gas samples was measured at the Geosciences Environment Toulouse laboratory in France. Titanium canisters filled with hydrothermal gases at a known pressure were connected to a 60-mL impinger (Savillex) via a stainless steel valve, 6 mm of PTFE tubing and a 6-mm-long Pyrex bubbler tube. The impinger contained 30 mL of a strongly oxidizing 40 vol% inverse aqua-regia solution (HNO₃:HCl = 2:1). The canister was partly opened, which induced the hydrothermal gases to bubble in the aqua-regia solution and the oxidation of Hg(0) to Hg(II). The volume of hydrothermal gas sampled was measured with a ball flowmeter and a chronometer. The flow rate was approximately 80 mL min^-1, and trapping continued until atmospheric pressure equilibration occurred between the canister and the oxidizing solution trap. The total Hg concentrations were measured in accordance with the USEPA 1631 method. Aliquots of 0.1–1 mL were analyzed in duplicate using semiautomatic cold vapor atomic fluorescence spectrometry (CV-AFS; Brooks Rand Model III, United States) with a single gold trap. The analysis accuracy of CV-AFS was evaluated according to the standard analysis method for the NRC ORMS-4 certified reference material (22.0 ± 1.6 ng L^-1, mean ± SD), and satisfactory results were obtained (21.3 ± 2.4 ng L^-1, n = 7).

The ³He/⁴He and ⁴He/²⁰Ne ratios addressed in the present study have been reported together with the respective analytical methods by Rizzo et al. (2016). The ⁴⁰Ar/³⁶Ar and ³⁸Ar/³⁶Ar ratios were measured using a mass spectrometer (Argus GVI) with analytical errors of < 0.2%.

The C-isotope composition of CO₂ [expressed as δ¹³C ‰ vs. V-PDB (Vienna-Pee Dee Belemnite)] was determined using a continuous-flow isotope-ratio mass spectrometer (Thermo Delta Plus XP, Finnigan), connected to a gas chromatograph (Trace GC) and interface (Thermo GC/C III, Finnigan). The gas chromatograph and its column (length = 30 m and i.d. = 0.32 mm; Poraplot-Q) were operated at a constant temperature of 50°C using He as the carrier gas. The analytical errors were <0.1‰. The C and H isotopes of CH₄ were analyzed using the same instrument. A combustion interface (Thermo GC III, Finnigan) was used to produce CO₂ from methane, while a gas-chromatograph/thermal-conversion interface provided on-line high-temperature conversion of CH₄ into H. The SDs for the δ¹³C and δD measurements of CH₄ were <0.2 and <2.5‰, respectively.

Results

Gas Chemistry

The chemical composition of submarine gases collected at Kolumbo is presented in Table 1. These gases are dominated by CO₂, which is present at concentrations up to 99.1% (Figure 3A). The concentration of CH₄ ranges from 1052 to 5521 ppm, while C₂H₆ and C₃H₈ exhibit much narrow ranges of 95–128 and 14–20 ppm, respectively (values measured in the Giggenbach bottle; Giggenbach, 1975). He varies between 9 and 40 ppm, while H₂ and CO range from 170 to 716 ppm and from 2 to 7 ppm, respectively. The O₂ and N₂ contents varied depending on the degree of sample contamination by ambient air, with maximum values of 5.1 and 21%, respectively (Figures 3B, 4A). ²⁰Ne ranges between 0.082 and 4.2 ppm, while ⁴⁰Ar ranges between 69 and 834 ppm (Figure 4B). It should be noted that ⁴⁰Ar was not measured in the sample with the greatest air contamination (O₂ = 5.1%, N₂ = 21%, and ²⁰Ne = 4.2 ppm). In order to evaluate the extent of air or the presence of air-saturated water (ASW), O₂ and Ar are plotted versus ²⁰Ne in Figures 4A,B, respectively. All of the analyzed samples fell along the line representing air rather than ASW, indicating that this contamination is probably due to sampling or extraction procedures. In order to determine the gas composition before air contamination (G_f), data were corrected based on the O₂ content measured in each analysis as follows:

$$\mathit{G}\mathit{f}\mathrm{=}\frac{\mathrm{(}\mathit{G}\mathit{s}\mathrm{-}\mathrm{(}\mathit{G}\mathit{a}\mathrm{\times }\mathit{F}\mathrm{)}\mathrm{)}}{\mathrm{(}\mathrm{1}\mathrm{-}\mathit{F}\mathrm{)}}$$

TABLE 1

Table 1. Chemical composition of major and minor gaseous components from Kolumbo hydrothermal vents. Ar, CO₂, and CH₄ isotope ratios are also reported.

FIGURE 3

Figure 3. (A) CO₂-CH₄-He ternary diagram showing the main components of Kolumbo gases. (B) He-Ar-N₂ ternary diagram displaying air and ASW (air-saturated water) contamination trends.

FIGURE 4

Figure 4. Binary plots of Ne versus (A) O₂ and (B) Ar, with the air line also shown. Data plotting along the air line represent evidence of atmospheric contamination due to sampling or storage conditions. See the main text for details.

where Gs is the concentration of the gaseous species measured in the sample (e.g., He), Ga is the concentration of the same gaseous species in air (e.g., He = 5.24 ppm) and F is the fraction of air calculated from the O₂ concentration measured in the sample divided by the atmospheric O₂ concentration (20.947%). We point out that even assuming that all of the ²⁰Ne or ⁴⁰Ar measured in our samples is of atmospheric origin (as for O₂), the recalculated percentages of air contamination would still be comparable. The back-corrected chemical composition of Kolumbo gases is reported in Table 2.

TABLE 2

Table 2. Restored chemical composition after correction for atmospheric contamination and pristine composition of gases after correction for CO₂ dissolution in water.

Furthermore, the concentration of gaseous Hg(0) differed markedly (and significantly) among the seven investigated chimneys, from 61 to 1301 ng m^-3 (Table 3).

TABLE 3

Table 3. Hydrothermal gas Hg(0) concentrations of the seven chimneys where it was possible to make measurements.

Isotopic Composition of Gases

Table 1 reports the isotopic compositions of CO₂, CH₄, and Ar in the analyzed gases. The C-isotope composition (δ¹³C_CO2) varied between -0.04 and 1.15‰ V-PDB, whereas increasing ratios were positively correlated with the concentrations of He, H₂, CO, and CH₄. The CH₄-isotope composition varied over a narrow range for both C (δ¹³C = -18.8 to -17.6‰ V-PDB) and H (δD = -118.6 to -115.4‰ V-SMOW).

The ³He/⁴He and ⁴He/²⁰Ne ratios used in the present study were previously measured and discussed by Rizzo et al. (2016). Here we further measured the Ar-isotope composition, which was found to be close to the values typically encountered in atmospheric air (⁴⁰Ar/³⁶Ar = 295.5 and ³⁸Ar/³⁶Ar = 0.188; Ozima and Podosek, 1983). ⁴⁰Ar/³⁶Ar in hydrothermal gases varies between 296 and 316, while ³⁸Ar/³⁶Ar is between 0.187 and 0.189. However, ⁴⁰Ar was corrected for atmospheric contamination (⁴⁰Ar^∗) in the samples having ⁴⁰Ar/³⁶Ar > 308 as follows:

$${{}^{\mathrm{40}}\mathrm{A}\mathrm{r}}^{\mathrm{*}}\mathrm{=}{}^{\mathrm{40}}{\mathrm{A}\mathrm{r}}_{\mathrm{s}\mathrm{a}\mathrm{m}\mathrm{p}\mathrm{l}\mathrm{e}}\mathrm{-}\mathrm{[}{}^{\mathrm{36}}{\mathrm{A}\mathrm{r}}_{\mathrm{s}\mathrm{a}\mathrm{m}\mathrm{p}\mathrm{l}\mathrm{e}}\mathrm{\cdot }{\mathrm{(}{}^{\mathrm{40}}\mathrm{A}\mathrm{r}\mathrm{/}{}^{\mathrm{36}}\mathrm{A}\mathrm{r}\mathrm{)}}_{\mathrm{a}\mathrm{i}\mathrm{r}}\mathrm{]}$$

This correction is useful for obtaining accurate estimates of the ⁴He/⁴⁰Ar^∗ ratio, which in Kolumbo gases vary between 4.8 and 5.5. These ratios are within the ratio range typical of the mantle (⁴He/⁴⁰Ar = 1–5; e.g., Ozima and Podosek, 1983; Marty, 2012), confirming that Kolumbo gases preserve the features that they have inherited from the magma.

Discussion

Gas–Water Interaction of Magmatic Fluids

The chemical composition of gases corrected for air contamination as well as the δ¹³C_CO2 show a variability clearly modulated by a process of selective dissolution of gases in water driven by their different solubilities (Figures 5, 6). In order to determine the original composition of the intact gases, which is necessary for evaluating the origin of CO₂ and the pressure and temperature conditions of the hydrothermal system, we initially defined the boundary conditions for modeling. Due to the lack of direct measures of temperature in the fluids discharged from the Kolumbo chimneys, we assumed a homogeneous emission temperature of 220°C, as measured by Sigurdsson et al. (2006) (see section “Sampling and Analytical Techniques”). This assumption is reasonable because it is based on direct measurements made in the same part of the floor of Kolumbo submarine crater and from vents showing a sustained flux of gas bubbles, which indicates the concomitant release of high-temperature fluids (Sigurdsson et al., 2006). Regardless of the accuracy of this estimate, it should be remembered that any slight difference in temperature between vents would mostly influence the extent of the CO₂ dissolution in water rather than the path of fractionation of less-soluble species, especially when considering elemental ratios. Since gases are emitted on the floor of Kolumbo submarine crater, which is at around 500 m b.s.l., we assumed a constant pressure of 50 bar. Based on the assumption of a temperature of 220°C, the saturated vapor pressure would be ∼23 bar. This means that the partial pressure of CO₂ would be ∼27 bar within a bubble of fluids exsolved from the vents located at the hydrostatic pressure of 50 bar.

FIGURE 5

Figure 5. Plots of He versus (A) CH₄ and (B) H₂, and CH₄ versus (C) H₂ and (D) CO. Legend symbols as in Figure 4. Lines show the path of selective dissolution of gases in water modeled as an open-system condensation process under equilibrium conditions. Arrows indicate the process direction. The starting gas composition for modeling was He = 8 ppm, H₂ = 150 ppm, N₂ = 0.2%, CH₄ = 1150 ppm, CO = 1.5 ppm and CO₂ = 99.8%. See the main text and Table 2 for further details.

FIGURE 6

Figure 6. Plot of CH₄/CO₂ versus He/CO₂. Legend symbols as in Figure 4. Line and arrow are as in Figure 5. The starting gas composition for modeling was CH₄/CO₂ = 1.15 × 10^-3 and He/CO₂ = 8.02 × 10^-6. See the main text for further details.

Considering that we are dealing with submarine emissions, it is reasonable to assume that hydrothermal water condenses in the seawater together with the highly soluble acidic gases (SO₂, HCl, and HBr). We cannot exclude that the CO₂ continued to dissolve in water inside the gas-tight bottles during the time that lapsed between sampling and gas extraction aboard the vessel, mostly because of cooling.

The residual gas phase that we studied comprised mainly CO₂ and secondarily N₂, H₂, CO and CH₄, as well as noble gases at low levels and trace gases. We ultimately assumed a geochemical system comprising CO₂, N₂, H₂, CO, CH₄, He, Ne and Ar, all of which dissolve in liquid water according to Henry’s law. After the condensation of water vapors and highly soluble acidic gases, a further progression of gas dissolution would induce the preferential dissolution of CO₂ in water and the consequent enrichment of less-soluble species such as N₂, H₂, CO, hydrocarbons and noble gases (Sander, 2015). Since our dry gas mixture comprised > 97% CO₂, we deduce that the extent of fractionation was low. Therefore, this process can be better evaluated using other components such as CH₄ versus He (Figures 5, 6) rather than the CO₂ concentration.

Following the thermodynamic approach proposed by Fernandez-Prini et al. (2003 and references therein) and based on our pressure and temperature boundary conditions, we calculated Henry’s constant k_H and the gas–liquid distribution constant K_D for each species included in our geochemical system. We simulated a condensation process under equilibrium conditions as expressed by the Rayleigh (1896) equation:

$$\frac{\mathit{R}\mathit{v}}{{\mathit{R}\mathit{v}}_{\mathrm{0}}}\mathrm{=}{\mathit{f}}^{\mathrm{\alpha }\mathrm{-}\mathrm{1}}$$

where Rv₀ is the initial ratio of the bulk composition (e.g., He/CO₂), Rv is the same instantaneous ratio in the residual gas phase (v), f is the fraction of the residual gas phase and α is the fractionation factor determined by the solubility ratio of the species under consideration (e.g., k_H-He/k_H-CO2). We also assumed open-system conditions considering the continuous removal of gas-saturated parcels of water.

The initial composition applied in our model is reported in Table 2 and in the captions of Figures 5–7. Similarly to gas levels, the isotopic composition of gaseous CO₂ (δ¹³C_CO2) changes as a result of its dissolution in water (Figure 7) and the isotopic fractionation between gaseous and dissolved inorganic CO₂ [CO₂(aq)]. CO₂(aq) is referred to as dissolved inorganic C (DIC) and is equal to the sum of the aqueous species H₂CO₃, ${\mathrm{H}\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{-}}$, and ${\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{2}\mathrm{-}}$. The fractionation process is modeled using the Rayleigh equation as follows (Clark and Fritz, 1997):

$${\mathrm{\delta }}^{\mathrm{13}}{{\mathrm{C}}_{\mathrm{C}\mathrm{O}}}_{\mathrm{2}}\mathrm{=}{\mathrm{\left(}{\mathrm{\delta }}^{\mathrm{13}}{{\mathrm{C}}_{\mathrm{C}\mathrm{O}}}_{\mathrm{2}}\mathrm{\right)}}_{\mathrm{0}}\mathrm{+}\mathrm{\epsilon }\mathrm{\text{In}}\mathrm{\left(}\mathit{f}\mathrm{\right)}$$

FIGURE 7

Figure 7. Plot of He/CO₂ versus δ¹³C_CO2. Legend symbols as in Figure 4. Line and arrow are as in Figure 5. The starting gas composition for modeling was δ¹³C_CO2 = –0.4‰ and He/CO₂ = 8.02 × 10^-6. ε_DIC-CO2g = –0.79 at 220°C, which implies H₂CO₃ and ${\mathrm{H}\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{-}}$ molar fractions of 0.96 and 0.04, respectively, equivalent to pH = 5. See the main text for further details.

where (δ¹³C_CO2)₀ is the initial CO₂-isotope composition, f is the fraction of the residual gas phase, and ε is the fractionation factor between DIC and gaseous CO₂ (CO_2(g)). This fractionation factor is obtained by summing up the fractionation factors of dissolved C species and CO_2(g) weighted for their molar fraction with respect DIC (Zhang et al., 1995; Allègre, 2008):

$${\mathrm{\epsilon }}_{\mathrm{D}\mathrm{I}\mathrm{C}\mathrm{-}{\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{\left(}\mathrm{g}\mathrm{\right)}}\mathrm{=}\frac{\mathrm{\left[}{\mathrm{H}}_{\mathrm{2}}{\mathrm{C}\mathrm{O}}_{\mathrm{3}}\mathrm{\right]}{\mathrm{\epsilon }}_{{\mathrm{H}}_{\mathrm{2}}{\mathrm{C}\mathrm{O}}_{\mathrm{3}}\mathrm{-}{\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{\left(}\mathrm{g}\mathrm{\right)}}\mathrm{+}\mathrm{\left[}\mathrm{H}{\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{-}}\mathrm{\right]}{\mathrm{\epsilon }}_{\mathrm{H}{\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{-}}\mathrm{-}{\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{\left(}\mathrm{g}\mathrm{\right)}}\mathrm{+}\mathrm{\left[}{\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{2}\mathrm{-}}\mathrm{\right]}{\mathrm{\epsilon }}_{{\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{2}\mathrm{-}}\mathrm{-}{\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{\left(}\mathrm{g}\mathrm{\right)}}}{\mathrm{\left[}{\mathrm{H}}_{\mathrm{2}}{\mathrm{C}\mathrm{O}}_{\mathrm{3}}\mathrm{\right]}\mathrm{+}\mathrm{\left[}\mathrm{H}{\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{-}}\mathrm{\right]}\mathrm{+}\mathrm{\left[}{\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{2}\mathrm{-}}\mathrm{\right]}}$$

The molar fraction of each C species depends on the water temperature and pH. We calculated each fractionation factor at 220°C based on the approach of Zhang et al. (1995), while the molar fractions of H₂CO₃ and ${\mathrm{H}\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{-}}$ were set to 0.96 and 0.04, respectively. We also assumed that the molar fraction of ${\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{2}\mathrm{-}}$ was 0 given that the pH of water on the floor of Kolumbo submarine crater is typically lower than 8 (Mandalakis et al., 2019), which means that the terms involving [${\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{2}\mathrm{-}}$] can be neglected in the above equation. Under these conditions, ε_DIC-CO2(g) was calculated to be -0.79. The molar fractions were also used in the following equation describing the dissociation of H₂CO₃ in order to estimate the pH of water in which CO₂ was dissolved:

$${\mathrm{H}}_{\mathrm{2}}{\mathrm{C}\mathrm{O}}_{\mathrm{3}\mathrm{\left(}\mathrm{a}\mathrm{q}\mathrm{\right)}}\mathrm{\leftrightarrow }{\mathrm{H}}_{\mathrm{a}\mathrm{q}}^{\mathrm{+}}\mathrm{+}{\mathrm{H}\mathrm{C}\mathrm{O}}_{\mathrm{3}\mathrm{\left(}\mathrm{a}\mathrm{q}\mathrm{\right)}}^{\mathrm{-}}$$

$${\mathrm{K}}_{\mathrm{A}\mathrm{1}}\mathrm{=}\frac{\mathrm{\left[}{\mathrm{H}\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{-}}\mathrm{\right]}\mathrm{\left[}{\mathrm{H}}^{\mathrm{+}}\mathrm{\right]}}{\mathrm{\left[}{\mathrm{H}}_{\mathrm{2}}{\mathrm{C}\mathrm{O}}_{\mathrm{3}}\mathrm{\right]}}\mathrm{=}\mathrm{4.3}\mathrm{\cdot }{\mathrm{10}}^{\mathrm{-}\mathrm{7}}$$

The chemical and δ¹³C_CO2 variability of gases emitted from Kolumbo (Figures 7, 8) is well-modeled by a trend of CO₂ dissolution in water having H₂CO₃ and ${\mathrm{H}\mathrm{C}\mathrm{O}}_{\mathrm{3}}^{\mathrm{-}}$ molar fractions of 0.96 and 0.04, respectively, which at 220°C corresponds to pH ∼5. It is particularly interesting that this value is identical to that measured by Carey et al. (2013) in fluids emitted by a neighboring vent on the floor of Kolumbo submarine crater thereby further validating the boundary conditions selected for our model.

FIGURE 8

Figure 8. Plot of δ¹³C_CO2 versus CO₂/³He. ³He data are taken from Rizzo et al. (2016), who analyzed the same suite of gas samples. Data for Santorini, Nisyros, Vulcano, and Panarea are from Brombach et al. (2003), Paonita et al. (2013), Tassi et al. (2014), and Rizzo et al. (2015), respectively. The violet dotted curve represents binary mixing between a MORB-like upper mantle having δ¹³C = –4‰ and CO₂/³He = 2 × 10⁹, and limestone having δ¹³C = 0‰ and CO₂/³He = 1 × 10¹³ (Marty and Jambon, 1987; Javoy and Pineau, 1991; Sano and Marty, 1995). The black dotted line represents the path of selective dissolution of gases in water modeled by an open-system condensation process under equilibrium conditions. The starting gas composition for modeling was δ¹³C_CO2 = –0.4‰ and CO₂/³He = 1.26 × 10¹⁰. The symbols are as in Figure 5.

Origin of Gases Emitted at Kolumbo

Origin of CO₂

Back-corrected calculations of the concentration and isotopic composition of gaseous CO₂ before its selective dissolution in water allow evaluation of its origin in the Kolumbo magmatic system and the making of inferences about the local mantle. To our knowledge, this is the first study to provide data on δ¹³C_CO2 emitted by Kolumbo submarine volcano.

We examined the origin of CO₂ by combining the ³He/⁴He values reported by Rizzo et al. (2016) with the CO₂ concentration and δ¹³C_CO2 values measured in the present study. The plot of δ¹³C_CO2 versus CO₂/³He is a convenient diagnostic for this purpose (Sano and Marty, 1995), although post-magmatic processes may strongly modify the original gas composition and thus compromise the accuracy of this approach (e.g., Oppenheimer et al., 2014). As discussed in Section “Gas–Water Interaction of Magmatic Fluids,” the trend of the variation of Kolumbo gases is modulated by the process of CO₂ dissolution in water, which can be observed in Figure 8. However, the back-corrected calculated values for the gas composition produced δ¹³C_CO2 and CO₂/³He values of -0.4‰ and 1.28 × 10¹⁰, respectively. This CO₂/³He ratio falls within the range of values reported for gases emitted from arc volcanoes worldwide (≥10¹⁰; Hilton et al., 2002), while δ¹³C_CO2 is within the range proposed for limestone (CO₂/³He ∼1 × 10¹³, δ¹³C = -1 to +1‰; Sano and Marty, 1995). Figure 8 presents the binary mixing line between MORB (CO₂/³He = 2 × 10⁹, δ¹³C = -4‰) and limestone (assuming CO₂/³He = 1 × 10¹³, δ¹³C = 0‰), which highlights that gases emitted at Kolumbo have a MORB source contaminated by limestone. The main question arising from this evaluation is whether the contamination by carbonates occurs in the mantle (by subduction of limestone-bearing sediments) or in the crust (from carbonates in the basement). Unfortunately we have not performed any direct measurements of mantle fluids that could shed light on this question for Kolumbo, in contrast to the values measured for the mantle cumulates at Stromboli arc volcano (Gennaro et al., 2017). We therefore attempted to address this issue indirectly by comparing our data with those for fumarole gases emitted at Santorini (Nea Kameni Island) and Nisyros (Figure 8), which are the most-active volcanoes along HVA. Recent measurements of fumarole samples from Santorini (Rizzo et al., 2015) produced δ¹³C_CO2 and CO₂/³He values that are comparable to those for Kolumbo gases, despite the considerable variability that may have been caused by gas–water interaction, similarly to our case study (Figure 8). In detail, δ¹³C_CO2 as measured previously at Nea Kameni fumaroles was -0.2 ± 2.7‰ (Dotsika et al., 2009; Parks et al., 2013; Tassi et al., 2013), which is consistent with the present results for Kolumbo (Figure 8). Similar δ¹³C_CO2 and CO₂/³He values were also found by Brombach et al. (2003) for Nisyros fumaroles, although those data exhibited substantial variability. However, the reported range of values fell mainly within the binary mixing line between MORB and limestone, as observed for Santorini and Kolumbo gases.

Parks et al. (2013) proposed that the C-isotope signature of CO₂ emitted at Santorini and Nisyros is consistent with mixing between magmatic fluids and crustal basement limestone, the presence of which has been inferred by Nicholls (1971) and Spandler et al. (2012). This interpretation is plausible for two main reasons: (1) similar δ¹³C_CO2 signatures characterize most of the Mediterranean volcanoes where there is evidence of the presence of a carbonate basement, and (2) the δ¹³C_CO2 variability observed in fumaroles at Santorini during the 2011–2012 unrest (Tassi et al., 2013) is not compatible with a mantle signature modified by subducted carbonates (Parks et al., 2013). In further support of the hypothesis of Parks et al. (2013), we highlight that there is strong evidence for the presence of Mesozoic carbonates within the crust beneath Kolumbo (Kilias et al., 2013). However, Rizzo et al. (2016) reported that the ³He/⁴He signature of Kolumbo gases is indicative of the direct degassing of a MORB-like mantle, while those at Santorini are subsequently modified by crustal contamination. Those authors therefore argued that the mantle beneath Kolumbo and Santorini is homogeneous in terms of the He-isotope signature.

We consider that the consistency of the δ¹³C_CO2 signatures at Santorini, Kolumbo, Nisyros, and most Mediterranean volcanoes with MORB and limestone mixing represents evidence of mantle metasomatism induced by decarbonation of subducting limestone-bearing sediments, rather than mixing of comparable proportions of magmatic and crustal fluids originating from local basement carbonates. Indeed, based on δ¹³C_CO2 in fluid inclusions of mantle cumulates from Stromboli, Gennaro et al. (2017) revealed that the isotopic signature of local mantle reflected CO₂ contamination arising from the decarbonation of sediments carried by the subducting Ionian slab. We therefore argue that the mantle beneath Kolumbo and Santorini is reasonably homogeneous also in terms of δ¹³C_CO2 signatures, and we cannot exclude that CO₂ is already contaminated by the decarbonation of slab sediments. Local crustal contamination may eventually further modify the isotopic composition of metasomatized mantle C, as observed for He isotopes beneath Santorini (Rizzo et al., 2015, 2016).

Origin of CH₄

Some studies have focused on the origin of CH₄ in fluids emitted from submarine hydrothermal systems (Welhan, 1988; McCollom and Seewald, 2007; Proskurowski et al., 2008; Keir, 2010; McDermott et al., 2015; Wen et al., 2016; Xue-Gang et al., 2016; Wang et al., 2018; and references therein). Besides the difficulty of sampling at considerable water depths, the availability of CH₄ data is further impaired by the gases emitted from hydrothermal or magmatic systems generally being dominated by CO₂ (e.g., Yang et al., 2005; Lupton et al., 2006, 2008), with only trace levels of CH₄. This compromises the ability to perform isotope analyses, especially those of H.

In this study we measured the stable C and H isotopes of CH₄ in Kolumbo hydrothermal gases and performed evaluations to ascertain the origin of CH₄. In this context we plotted δ¹³C_CH4 versus CH₄/(C₂H₆+C₃H₈), which is the most-common classification approach that was introduced by Bernard et al. (1978), and δD_CH4 versus δ¹³C_CH4, as introduced by Schoell (1980) and later modified by Mazzini et al. (2011), to distinguish CH₄ originating from thermogenic and microbial processes as well as sediment-free midocean ridges (Figure 9). It should be stressed that the CH₄-isotope composition of Kolumbo gases (Table 1) varied within a narrow range for both C (δ¹³C = -18.8 to -17.6‰ V-PDB) and H (δD = -118.6 to -115.4‰ V-SMOW). The data points in the plot of δ¹³C_CH4 versus CH₄/(C₂H₆+C₃H₈) for Kolumbo samples (Figure 9A) fall within an origin area of the sediment-free midocean ridge, similar to other Italian geothermal and hydrothermal systems (e.g., Panarea and Pantelleria; Tassi et al., 2012). Such data are often classified as abiogenic (McCollom and Seewald, 2007 and references therein), meaning that methanogenesis does not involve a biogenic organic precursor (Welhan, 1988). Nevertheless, the accuracy of this approach can be significantly compromised since the concentrations of light hydrocarbons and the isotopic composition of CH₄ can both be modified by oxidation and migration processes (Welhan, 1988). More specifically, δD_CH4 and δ¹³C_CH4 measured in Kolumbo gases may have been altered by thermogenic gas oxidation, as in the case of CH₄ from Salton Sea Geothermal Field (Mazzini et al., 2011). Alternatively, it is possible that abiogenic CH₄ is either formed in basalts and extracted into the circulating fluids (Welhan, 1988) or produced by Fischer-Tropsch-type chemical reactions (e.g., Proskurowski et al., 2008; Keir, 2010 and references therein; Etiope and Sherwood-Lollar, 2013). However, the latter process was recently questioned by Taran et al. (2010a), who attributed the isotopic trends in the δ¹³C and δD values of light hydrocarbons as mixing between two or more endmembers.

FIGURE 9

Figure 9. (A) Plot of δ¹³C_CH4 versus CH₄/(C₂H₆+C₃H₈) modified from Bernard et al. (1978). Field data for microbial and thermogenic gases and from sediment-free midocean ridges and for igneous rocks are from McCollom and Seewald (2007 and references therein). Data for Italian geothermal systems and hydrothermal systems from Campi Flegrei, Vesuvio and Panarea are from Tassi et al. (2012 and references therein). (B) Plot of δD_CH4 versus δ¹³C_CH4 as introduced by Schoell (1980) and modified by Mazzini et al. (2011). Field data are from Mazzini et al. (2011 and references therein). The symbols are as in Figure 5.

On the other hand, the plot of δD_CH4 versus δ¹³C_CH4 for Kolumbo samples (Figure 9B) falls within the area of Socorro gases (Mexico, Taran et al., 2010b), between the range of abiotic CH₄ (McCollom and Seewald, 2007 and references therein) and a thermogenic field, which is how most geothermal and hydrothermal systems worldwide are generally characterized (Whiticar, 1999; Mango, 2000; Taran et al., 2010b; Tassi et al., 2012). By combining He-isotope data (i.e., ³He/⁴He ratio of ∼7 Ra for Kolumbo gases; Rizzo et al., 2016) with those of CH₄ (e.g., Hsin-Yi et al., 2016), we obtained an average CH₄/³He ratio of ∼1.4 × 10⁷ (Table 1). This ratio falls within the range of values measured in fluids from the East Pacific Rise (e.g., Proskurowski et al., 2008; Keir, 2010; and references therein).

Christakis et al. (2018) analyzed the microbial community at Kolumbo sulfide chimneys using next-generation sequencing technologies, and did not find any microbial lineages closely related to CH₄ production processes. In addition, only a few phylotypes typically involved in CH₄ oxidation were identified. A further metagenomic investigation of the Kolumbo seafloor found negligible methanogenesis-related marker genes (Oulas et al., 2015), supporting the abiotic origin of CH₄.

It was beyond the scope of this study to resolve the scientific controversies of the different theories about abiotic CH₄. Overall, it is more likely that CH₄ in Kolumbo gases comprise a mixture of oxidized thermogenic and abiogenic CH₄ formed in high-temperature (>200°C) magmatic-hydrothermal systems.

Submarine Hydrothermal Hg(0) Emissions

The natural aerial volcanic emissions of Hg via passive degassing have been estimated at 76 ± 30 × 10⁶ g yr^-1 (Bagnato et al., 2014). However, the quantity and impact of Hg released by submarine volcanoes and hydrothermal systems is less well-known due to a lack of observations (Varekamp and Buseck, 1981; Bagnato et al., 2017). In particular, there are severe logistical constraints in sampling submarine volcanic gas emissions on the floor of Kolumbo submarine crater, since it is at 500 m b.s.l. (Fitzgerald and Lamborg, 2004). Moreover, previous studies of submarine hydrothermal Hg have investigated dissolved Hg(II) concentrations but not dissolved gaseous Hg(0) (Lamborg et al., 2006).

The present study deployed gas-tight titanium syringes connected to inverted funnels above hydrothermal vents and activated by an ROV, which provided the unique opportunity to collect pristine bubbling gas emissions discharged from an active submarine volcano. The levels of gaseous Hg(0) in the gas samples varied from 61 to 1300 ng m^-3, which are roughly 10 times higher than the levels previously reported for on-land Santorini fumaroles (9 to 121 ng m^-3; Bagnato et al., 2013) and also the worldwide aerial volcanic Hg(0) concentrations (4 to 125 ng m^-3; Bagnato et al., 2014). These data seem to further highlight that the level of magmatic activity is higher at Kolumbo volcano than at Santorini. It is also worth mentioning that the global average concentration of atmospheric Hg(0) in the northern hemisphere is 1.5 ng m^-3.

Hydrothermal Gas Equilibrium

The temperature and pressure conditions at which gaseous species equilibrate in the hydrothermal system beneath the floor of Kolumbo submarine crater can be estimated from concentration data of minor reactive species together with CO₂ and H₂O (Chiodini and Marini, 1998 and references therein). The most-important assumptions in this approach are that the gaseous species have attained chemical equilibrium in the hydrothermal system and are quenched during their ascent to the sea surface. We evaluated two stability diagrams based on the concentrations of H₂, CO, CH_4, and CO₂ in the dry gas phase (Figure 10), following the approach proposed by Chiodini and Cioni (1989) and Chiodini et al. (2001) and subsequently modified by Chiodini et al. (2006) for submarine gases at Panarea. Because H₂O represents the main component of hydrothermal fluids, we assumed that pure water coexisted with vapor, whereas the fugacity of water vapor (fH₂O) as a function of temperature (Giggenbach, 1987) can be expressed as follows:

$$\log \mathrm{}\mathrm{\left(}\mathit{f}{\mathrm{H}}_{\mathrm{2}}\mathrm{O}\mathrm{\right)}\mathrm{\left(}\mathrm{b}\mathrm{a}\mathrm{r}\mathrm{\right)}\mathrm{=}\mathrm{5.510}\mathrm{-}\mathrm{2048}\mathrm{/}\mathit{T}$$

FIGURE 10

Figure 10. Plots of (A) log(CO/CH₄) versus log(H₂/CH₄) and (B) log(CO/CO₂) versus log(H₂/CH₄). The theoretical grid for hydrothermal gases was calculated by assuming their coexistence with pure liquid water (Chiodini et al., 2006 and references therein). Data for Vulcano Island are from Paonita et al. (2013), and those for Panarea are from Caracausi et al. (2005), Chiodini et al. (2006), and Tassi et al. (2014). The symbols are as in Figure 5. See the main text for further details.

where T is the temperature in kelvin. It is also important to consider the fugacity of O₂ (fO₂), since this controls the redox conditions at which hydrothermal systems are formed. Since the H₂O concentration of hydrothermal vapors in our samples was not available and we have no information on the mineral assemblage that could fix fO₂, we considered the redox buffer typically proposed for hydrothermal systems by D’Amore and Panichi (1980):

$$\log \mathrm{}\mathrm{\left(}\mathit{f}{\mathrm{O}}_{\mathrm{2}}\mathrm{\right)}\mathrm{\left(}\mathrm{b}\mathrm{a}\mathrm{r}\mathrm{\right)}\mathrm{=}\mathrm{8.20}\mathrm{-}\mathrm{23643}\mathrm{/}\mathit{T}$$

Complementarily, we considered the following reactions and temperature-dependent equilibrium constants based on the thermodynamic data reported by Stull et al. (1969) and Giggenbach (1980, 1987):

$${\mathrm{H}}_{\mathrm{2}}\mathrm{O}\mathrm{\leftrightarrow }{\mathrm{H}}_{\mathrm{2}}\mathrm{+}\frac{\mathrm{1}}{\mathrm{2}}{\mathrm{O}}_{\mathrm{2}}\mathrm{\text{\hspace{1em}log}}\mathrm{\left(}{\mathit{K}}_{{\mathrm{H}}_{\mathrm{2}}\mathrm{O}}\mathrm{\right)}=\mathrm{2.548}\mathrm{-}\mathrm{12707}\mathrm{/}\mathrm{T}$$

$${\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{\leftrightarrow }\mathrm{C}\mathrm{O}\mathrm{+}\frac{\mathrm{1}}{\mathrm{2}}{\mathrm{O}}_{\mathrm{2}}\mathrm{\text{\hspace{1em}log}}\mathrm{\left(}{\mathit{K}}_{{\mathrm{C}\mathrm{O}}_{\mathrm{2}}}\mathrm{\right)}\mathrm{=}\mathrm{5.033}\mathrm{-}\mathrm{14955}\mathrm{/}\mathit{T}$$

$${\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{+}\mathrm{2}{\mathrm{H}}_{\mathrm{2}}\mathrm{\leftrightarrow }{\mathrm{C}\mathrm{H}}_{\mathrm{4}}\mathrm{+}{\mathrm{O}}_{\mathrm{2}}\mathrm{\text{\hspace{1em}log}}\mathrm{\left(}{\mathit{K}}_{{\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{+}{\mathrm{H}}_{\mathrm{2}}}\mathrm{\right)}\mathrm{=}\mathrm{-}\mathrm{4.569}\mathrm{-}\mathrm{16593}\mathrm{/}\mathit{T}$$

Finally, the following equations derived by Chiodini et al. (2006) were taken into account:

$$\log \mathrm{}\mathrm{(}{\mathrm{H}}_{\mathrm{2}}\mathrm{/}{\mathrm{C}\mathrm{H}}_{\mathrm{4}}\mathrm{)}\mathrm{=}\mathrm{8.811}\mathrm{-}\mathrm{4121.5}\mathrm{/}\mathit{T}\mathrm{-}\log \mathrm{}\mathrm{\left(}{\mathit{P}}_{{\mathrm{C}\mathrm{O}}_{\mathrm{2}}}\mathrm{\right)}\mathrm{\text{log}}\mathrm{(}\mathrm{C}\mathrm{O}\mathrm{/}{\mathrm{C}\mathrm{H}}_{\mathrm{4}}\mathrm{)}\mathrm{=}\mathrm{5.786}\mathrm{-}\mathrm{4326.5}\mathrm{/}\mathit{T}$$

where P is the partial pressure. The plot of log(CO/CH₄) versus log(H₂/CH₄) in Figure 10A suggests that Kolumbo gases would have reached an equilibrium at 200–250°C and P_CO2 ∼ 50 bar. In comparison with the submarine hydrothermal system of Panarea and the high-temperature fumaroles of Vulcano (Aeolian Islands, Italy), Kolumbo gases fall close to or within the range for Panarea gases, confirming their hydrothermal nature. The same pressure and temperature conditions were derived by examining the plot of log(CO/CO₂) versus log(H₂/CH₄) (Figure 10B). Given the compositional changes that may arise in hydrothermal gases due to the partial dissolution of CO₂ in water (see section “Gas–Water Interaction of Magmatic Fluids”), we recalculated the pressure and temperature conditions by applying the pristine composition estimated for each gas vent. Based on the values of Henry’s law constant, CO₂ was found to differ considerably from H₂, CO and CH₄, with the former having a greater tendency to partition from the gas phase into the aqueous phase (Sander, 2015). Therefore, the correction for gas–water partitioning was deemed necessary only when considering ratios that involved CO₂ [e.g., log(CO/CO₂); Figure 10B]. By applying the pristine composition, we estimated a gas equilibrium temperature of ∼200°C and P_CO2 ∼ 10 bar. These values are slightly lower than the pressure and temperature conditions derived from the plot of log(CO/CH₄) versus log(H₂/CH₄), but they remain well within the typical values reported for hydrothermal gases. In general, the estimated temperature using CO/CH₄, CO/CO_2, and H₂/CH₄ was equal to or below the highest temperature measured on the floor of Kolumbo submarine crater vents by Sigurdsson et al. (2006) and Carey et al. (2013). This suggests that part or all of these gaseous species did not achieve equilibrium under the assumed redox conditions or that the magmatic activity has waned in recent years.

To investigate these two possibilities, we applied the approach proposed by Chiodini and Cioni (1989) and Chiodini et al. (2006), which is based on the reaction:

$${\mathrm{C}\mathrm{H}}_{\mathrm{4}}\mathrm{+}\mathrm{3}{\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{\leftrightarrow }\mathrm{4}\mathrm{C}\mathrm{O}\mathrm{+}\mathrm{2}{\mathrm{H}}_{\mathrm{2}}\mathrm{O}$$

Since this reaction does not involve gaseous O₂, Chiodini et al. (2006) suggested that this approach provides equilibrium pressure and temperature estimates that are independent of the redox conditions. The resulting equations that allow pressure and temperature to be determined are

$$\mathit{T}\mathrm{\left(}\mathit{K}\mathrm{\right)}\mathrm{\left[}\frac{\mathrm{13606}}{\mathrm{8065}\mathrm{-}\log \mathrm{}\mathrm{\left(}\frac{{\mathrm{C}\mathrm{O}}^{\mathrm{4}}}{{\mathrm{C}\mathrm{O}}_{\mathrm{2}}^{\mathrm{3}}\mathrm{\cdot }{\mathrm{C}\mathrm{H}}_{\mathrm{4}}}\mathrm{\right)}}\mathrm{\right]}\mathrm{\text{log}}{\mathrm{P}}_{{\mathrm{C}\mathrm{O}}_{\mathrm{2}}}\mathrm{\left(}\mathrm{b}\mathrm{a}\mathrm{r}\mathrm{\right)}\mathrm{=}\mathrm{3.573}\mathrm{-}\frac{\mathrm{46}}{\mathrm{T}\mathrm{\left(}\mathrm{K}\mathrm{\right)}}\mathrm{-}\log \mathrm{}\mathrm{\left(}\frac{{\mathrm{H}}_{\mathrm{2}}}{\mathrm{C}\mathrm{O}}\mathrm{\right)}$$

This specific approach revealed that Kolumbo gases equilibrated at 200–250°C and when P_CO2 was slightly below 50 bar (Figure 11A). The presence of CO₂ in one of the above equations prompted us to again consider the pristine composition of gases prior to the gas–water partitioning process. This yielded a slightly lower equilibrium temperature (189–216°C), whereas P_CO2 remained the same (Figure 11A). This temperature range is comparable to those estimated using CO/CH₄, CO/CO₂ and H₂/CH₄ ratios, suggesting that the assumed redox conditions are not responsible for the difference in the estimated equilibrium temperatures.

FIGURE 11

Figure 11. (A) Plot of log[CO⁴/(CO₂³⋅ CH₄)] versus log(H₂/CO). The theoretical grid refers to vapor–liquid equilibrium conditions (Chiodini et al., 2006 and references therein). (B) Plot of log(CO/CO₂) versus log(H₂/H₂O) following the approach proposed by Chiodini and Marini (1998). The composition of liquid, vapor and vapor separated from liquid (single-step) phases are reported assuming that the fO₂ buffer proposed for hydrothermal systems by D’Amore and Panichi (1980) controls the redox conditions. Red dotted lines indicate the range of log(CO/CO₂) for Kolumbo gases corrected for the selective dissolution of CO₂ in water. See the main text and Chiodini and Marini (1998) for further details.

The findings of previous investigations of crater fumaroles at White Island (Giggenbach, 1987) and Vulcano (Chiodini et al., 1993, 1995) suggest that the two most-reactive gaseous species in hydrothermal systems are H₂ and CO, while CH₄ is one of the least reactive (Giggenbach, 1991; Taran and Giggenbach, 2003). To evaluate if CH₄ measured in Kolumbo gases attained full equilibrium, we considered the following formation reaction and relative equilibrium constant:

$${\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{+}{\mathrm{H}}_{\mathrm{2}}\mathrm{\leftrightarrow }\mathrm{C}\mathrm{O}\mathrm{+}{\mathrm{H}}_{\mathrm{2}}\mathrm{O}\mathrm{\text{log}}\mathrm{\left(}{\mathrm{K}}_{{\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{+}{\mathrm{H}}_{\mathrm{2}}}\mathrm{\right)}\mathrm{=}\log \mathrm{}\mathrm{(}\mathrm{C}\mathrm{O}\mathrm{/}{\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{)}\mathrm{-}\log \mathrm{}\mathrm{(}{\mathrm{H}}_{\mathrm{2}}\mathrm{/}{\mathrm{H}}_{\mathrm{2}}\mathrm{O}\mathrm{)}$$

In the gas phase, log(K_CO2+H2) = -2248 / T + 2.485

In a single saturated liquid phase, as assumed in our case, gas equilibrium contents are computed using the vapor–liquid distribution coefficient (B) (Giggenbach, 1980; D’Amore and Truesdell, 1988; Chiodini and Marini, 1998):

$$\log \mathrm{}\mathrm{\left(}{\mathrm{B}}_{\mathrm{C}\mathrm{O}}\mathrm{\right)}\mathrm{=}\mathrm{6.3173}\mathrm{-}\mathrm{0.01388}\mathrm{(}\mathrm{T}\mathrm{-}\mathrm{273.15}\mathrm{)}$$

$$\log \mathrm{}\mathrm{\left(}{\mathrm{B}}_{{\mathrm{C}\mathrm{O}}_{\mathrm{2}}}\mathrm{\right)}\mathrm{=}\mathrm{4.7593}\mathrm{-}\mathrm{0.01092}\mathrm{(}\mathrm{T}\mathrm{-}\mathrm{273.15}\mathrm{)}$$

$$\log \mathrm{}\mathrm{\left(}{\mathrm{B}}_{{\mathrm{H}}_{\mathrm{2}}}\mathrm{\right)}\mathrm{=}\mathrm{6.2283}\mathrm{-}\mathrm{0.01403}\mathrm{(}\mathrm{T}\mathrm{-}\mathrm{273.15}\mathrm{)}$$

Therefore, the above reported reaction for the liquid phase becomes

$$\log \mathrm{}\mathrm{\left(}{\mathrm{K}}_{{\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{+}{\mathrm{H}}_{\mathrm{2}}}\mathrm{\right)}\mathrm{-}\log \mathrm{}\mathrm{(}{\mathrm{B}}_{\mathrm{C}\mathrm{O}}\mathrm{/}{\mathrm{B}}_{{\mathrm{C}\mathrm{O}}_{\mathrm{2}}}\mathrm{)}\mathrm{+}\log \mathrm{}\mathrm{\left(}{\mathrm{B}}_{{\mathrm{H}}_{\mathrm{2}}}\mathrm{\right)}\mathrm{=}\log \mathrm{}\mathrm{(}\mathrm{C}\mathrm{O}\mathrm{/}{\mathrm{C}\mathrm{O}}_{\mathrm{2}}\mathrm{)}\mathrm{-}\log \mathrm{}\mathrm{(}{\mathrm{H}}_{\mathrm{2}}\mathrm{/}{\mathrm{H}}_{\mathrm{2}}\mathrm{O}\mathrm{)}$$

An equilibrium temperature of 250–300°C and log(K_CO2+H2) – log(B_CO/B_CO2) + log(B_H2) ≈ 0.0 ± 0.1 implies log(H₂/H₂O) ≈ log(CO/CO₂). Thus, the log(CO/CO₂) values measured in our samples and corrected for CO₂ dissolution can be used to roughly estimate the equilibrium temperature in liquid water, which varies in the narrow range of 263–276°C (Figure 11B). These values are higher than those previously calculated including CH₄, and thus we deduce that the latter species was not in equilibrium with the other gases and led to underestimations of the equilibrium temperature.

Based on the relationships proposed by Chiodini and Cioni (1989), we can finally calculate P_CO2 ∼ 30 bar and P_H2O ∼ 36 bar, which sum to a total pressure of ∼66 bar. Considering that Kolumbo gases are emitted at 500 m b.s.l., which corresponds to a pressure of ∼50 bar, the hydrothermal system must be located at a pressure of ∼116 bar (∼1000 m b.s.l.).

Conceptual Scheme of the Hydrothermal System Beneath Kolumbo

Based on this study and other previous findings (Sigurdsson et al., 2006; Carey et al., 2013; Kilias et al., 2013; Rizzo et al., 2016), in Figure 12 we propose a physical and geochemical model for the Kolumbo magmatic-hydrothermal system. We made some assumptions in order to simplify the system: (1) the permeability of the system was considered uniform, isotropic and sufficiently high; (2) the shallow magma chamber located 5–7 km beneath Kolumbo (Dimitriadis et al., 2010; Konstantinou and Yeh, 2012; Cantner et al., 2014) represents the source of magmatic gases feeding the hydrothermal system, (3) magmatic degassing does not or only weakly modifies these magmatic gases, and (4) seawater infiltrates from the sea bottom and favors condensation of magmatic gases both within and above the hydrothermal system.

FIGURE 12

Figure 12. Conceptual scheme of the hydrothermal system beneath Kolumbo. The location of the shallow magma chamber at 5–7 km beneath Kolumbo is from Dimitriadis et al. (2010), Konstantinou and Yeh (2012), and Cantner et al. (2014). The He-isotope composition is from Rizzo et al. (2016).

At Kolumbo, magmatic gases (H₂O, CO₂, S-bearing species, halogens and noble gases) having ³He/⁴He ∼ 7 Ra (Rizzo et al., 2016) and δ¹³C_CO2 ∼-0.4‰ ascend from the magma chamber and feed conduits up to the hydrothermal system (Figure 12). During the cooling of magmatic gases, H₂O and more-acidic species (S and halogens) condensate to form a hydrothermal system at ∼270°C in which water at pH ≤ 5 probably circulates. We cannot exclude that the hydrothermal system has multiple levels; however, we are able to reconstruct the upper level that feeds the bottom vents of the crater. Hydrothermal waters are probably fed by seawater infiltrating from the sea bottom and any brine formed from the adsorption of acidic gases in groundwater by rock dissolution.

The hydrothermal system has a total pressure of ∼66 bar, corresponding to ∼1,160 m b.s.l. and 650 m below the crater bottom, if the hydrostatic pressure is assumed at depth. Hydrothermal gases within the hydrothermal system undergo gas–water interactions along the fractures feeding the floor of Kolumbo submarine crater and from the shallow vents that favor the removal of most of the acidic gases (S and halogens) and the partial dissolution of CO₂ (see section “Gas–Water Interaction of Magmatic Fluids”). This produces an excess of N₂, noble gases and reactive gases via migration paths that permit their rapid ascent under advective degassing. At the crater bottom, several vents discharge fluids at temperatures up to 220°C, variable gas fluxes and pH down to ∼5 (Sigurdsson et al., 2006; Carey et al., 2013). Acoustic and visual imaging of the ascending bubbles suggests that CO₂ is being dissolved into seawater within ∼10 m above the crater floor (Carey et al., 2013).

Conclusions

We have investigated the geochemistry of CO₂-rich gases venting at 500 m b.s.l. from Kolumbo submarine volcano, which is located 7 km northeast of Santorini Island. The main findings are as follows:

- Gases are dominated by CO₂ (>97%), with a small air contamination probably related to sampling conditions, but are fractionated by a process of gas–water interaction reasonably related to variable fluxes of gases emitted from the different vents. This process induces a partial dissolution of gaseous CO₂ in water, leading to substantial enrichment of the residual gas in those species that are much less soluble in water (i.e., He, H₂, CO, CH_4, and N₂). This fractionation also affects the C-isotope composition of CO₂.

- We modeled the gas–water interaction process (∼220°C, ∼50 bar and pH ∼ 5) and reconstructed the chemistry and δ¹³C_CO2 of magmatic gases before interaction. We assess that the pristine CO₂ is characterized by δ¹³C ∼-0.4‰. Combining our data with ³He/⁴He measurements carried out in the same gas samples by Rizzo et al. (2016) yields CO₂/³He ∼1 × 10¹⁰. These data are in the same range as those obtained for the Santorini and Nisyros fumaroles. We argue that CO₂ emitted at Kolumbo could originate from a mantle contaminated by CO₂ via the decarbonation of subducting limestone.

- The CH₄-isotope composition falls within the range typical of hydrothermal gases, similar to other Mediterranean hydrothermal systems (Panarea and Campi Flegrei), suggesting that it originates from mixing between thermogenic and abiotic CH₄.

- We found that the Hg(0) concentration in Kolumbo gases ranges from ∼60 to 1300 ng m^-3. These levels are particularly high when compared to those of land-based fumaroles located on Santorini Island and worldwide aerial volcanic emissions, which suggests that the magmatic activity is higher at Kolumbo than at Santorini.

- Based on geo-indicators of pressure and temperature, we calculated that magmatic gases equilibrate within the Kolumbo hydrothermal system at about 270°C and 116 bar.

Author Contributions

AR and PN conceived the study. AR, VC, PN, PP, MM, GK, AM, and AlC participated in the collection of gas samples. AR, VC, and AnC analyzed the gas samples and elaborated the data. DL helped in figures preparation. All of the authors contributed to the preparation and editing of the final manuscript.

Funding

The authors acknowledge financial support from the SeaBioTech project (spider.science.strath.ac.uk/seabiotech/) funded by the European Commission within its FP7 Programme (Grant No. 311932), as well as from INGV, Sezione di Palermo, which permitted the access to analytical facilities.

Conflict of Interest Statement

The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

Acknowledgments

We thank Giuseppe Riccobono and Paolo Cosenza for providing technical support to the project and constructing the gas-tight sampler at INGV, Sezione di Palermo that was used to collect some of the submarine fluid samples. The officers and crew of RV AEGAEO are especially acknowledged for their valuable help during sampling. We thank INGV, Sezione di Palermo for providing analytical support, particularly Francesco Salerno and Mauro Martelli for performing analyses of gases chemistry, and Mariano Tantillo for supporting the laboratory activities involving noble-gas isotopes. We also thank Fausto Grassa, Giorgio Capasso, Ygor Oliveri, and Aldo Sollami for their help with the analyses of CO₂ and CH₄ isotopes in the stable-isotope laboratory. We further thank Cinzia Federico and Fausto Grassa for useful discussions that helped in the elaborative and interpretative framework. Data reported in Tables 1–3 are available by contacting the corresponding author. We thank the Chief Editor VA the Guest Associate Editor GG for handling and revising the manuscript, YT and an reviewer for suggestions that greatly improved the manuscript. English Science Editing revised the use of English in the manuscript.

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