Degassing of volatiles (H2O, CO2, S, Cl) during ascent, crystallization, and eruption at mafic monogenetic volcanoes in central Mexico

Journal of Volcanology and Geothermal Research 197 (2010) 225–238

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Journal of Volcanology and Geothermal Research j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / j vo l g e o r e s

Degassing of volatiles (H2O, CO2, S, Cl) during ascent, crystallization, and eruption at mafic monogenetic volcanoes in central Mexico Emily R. Johnson a,⁎, Paul J. Wallace a, Katharine V. Cashman a, Hugo Delgado Granados b a b

Department of Geological Sciences, University of Oregon, Eugene, OR 97403, USA Departamento de Vulcanología, Instituto de Geofísica, UNAM, Mexico

a r t i c l e

i n f o

Article history: Received 13 November 2008 Accepted 28 February 2010 Available online 11 March 2010 Keywords: degassing volatiles melt inclusions basalt crystallization

a b s t r a c t Mafic monogenetic volcanoes (cinder cones, maars) have eruption styles that include highly explosive, mildly explosive, and effusive regimes. Here we investigate the degassing and vapor-melt partitioning of volatiles (H2O, CO2, S, Cl) in monogenetic volcanoes from the subduction-related Michoacán–Guanajuato Volcanic Field (MGVF) in central Mexico. Olivine-hosted melt inclusions from these volcanoes contain variably degassed melts that were trapped over a wide range of pressures from b50 MPa to ∼ 300 MPa. Variations in melt compositions and volatile contents provide evidence that crystallization and differentiation were driven by degassing of H2O. Melt CO2 and H2O concentrations are highly variable, and much of the variation does not conform to equilibrium open- or closed-system degassing paths. Instead, we suggest that gas-fluxing – partial re-equilibration of magmas with CO2-rich gases rising from depth – can explain the variable CO2 and H2O concentrations in the melts. Such fluxing may be common in basaltic systems, and it increases the extent of crystallization during magma ascent by removing dissolved H2O from vaporsaturated (but H2O-undersaturated) melts. Strong degassing of S and Cl during magma ascent and crystallization begins at pressures of approximately 50 MPa. Using the relationship between degassing and crystallization, we calculate apparent vapor-melt partition coefficients for S and Cl. Our results show that, overall, S partitions more strongly into the vapor phase than Cl, consistent with published experimental data and thermodynamic models, and that vapor-melt partitioning of S increases more strongly with decreasing pressure than Cl. The S and Cl partitioning behavior inferred from the melt inclusion data are consistent with the gas fluxing model suggested by the H2O and CO2 data. © 2010 Elsevier B.V. All rights reserved.

1. Introduction Exsolution of volatiles from magma during ascent is the driving force for explosive eruptions. Until recently, studies of explosive volcanism have focused mainly on silicic systems. However, measurements of the pre-eruptive volatile concentrations in basaltic melts have demonstrated that mafic magmas can have high volatile contents (b6 wt.% H2O, b3000 ppm CO2, b4000 ppm S, b3000 ppm Cl; Sisson and Layne, 1993; Roggensack et al., 1997; Cervantes and Wallace, 2003; Spilliaert et al., 2006a,b; Wade et al., 2006; Benjamin et al., 2007; Johnson et al., 2008, 2009). These high volatile contents affect the explosivity of basaltic eruptions (Roggensack et al., 1997; Spilliaert et al., 2006b; Métrich and Wallace, 2008), leading to Strombolian, violent Strombolian, and even sub-Plinian eruption styles (Pioli et al., 2008, 2009). Furthermore, an increasing number of studies have emphasized the role of degassing in causing crystalliza⁎ Corresponding author. Present address: CODES ARC Centre of Excellence in Ore Deposits, University of Tasmania, Private Bag 126, Hobart, TAS 7001, Australia. Tel.: +61 3 6226 7210; fax: +61 3 6226 7662. E-mail address: [email protected] (E.R. Johnson). 0377-0273/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2010.02.017

tion of ascending melts (e.g., Cashman, 1992; Sisson and Layne, 1993; Blundy and Cashman, 2001;2005; Métrich et al., 2001; Roggensack, 2001; Atlas et al., 2006; Blundy et al., 2006; Johnson et al., 2008). Degassing of volatiles is complex because the solubilities and vapor-melt partition coefficients of the different volatile components vary with melt and gas composition, temperature, pressure, and in some cases, oxygen fugacity (e.g., Wallace and Carmichael, 1992; Dixon and Stolper, 1995; Webster et al., 1999; Newman and Lowenstern, 2002; Jugo et al., 2005; Alletti et al., 2009). Melt inclusion H2O and CO2 concentrations (which can be used to calculate inclusion entrapment pressures) can be combined with S and Cl contents of melt inclusions and matrix glasses to provide estimates of the depth and extent of degassing of different volatiles (Métrich et al., 1993, 2001; Gurenko et al., 2005; Spilliaert et al., 2006a,b). However, these studies have focused on large or persistently active basaltic volcanoes (e.g., Etna), where volcanic conduits are well established. Similar studies have not been done on monogenetic volcanoes, the most abundant type of volcano on land (Vespermann and Schmincke, 2002), where conduits to the surface must be created during early phases of activity and may evolve over the course of the eruption (e.g., Johnson et al., 2008; Erlund et al., 2009; Pioli et al., 2009).

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Here we present data for the pre-eruptive volatile concentrations (H2O, CO2, S, Cl) and major element compositions of olivine-hosted melt inclusions and matrix glasses from several monogenetic volcanoes in central Mexico. Using these data, we assess: (1) the pressures at which inclusions were trapped, (2) the causes of H2O and CO2 variations, (3) the effects of H2O loss on crystallization, (4) the vapor-melt partition coefficients for S and Cl, and (5) the ascent rates and gas outputs during violent Strombolian eruptions. 2. Samples and methods We sampled tephra from nine monogenetic volcanoes (eight cinder cones and one maar) across the Michoacán–Guanajuato Volcanic Field (MGVF), Mexico (Fig. 1). When possible, samples were taken from a single stratigraphic tephra section that allowed sampling from the earliest (basal) to latest phases of the eruption. Loose olivine crystals from the tephra samples were separated, washed, and cleaned of adhering glass using HBF4, and those with suitable melt inclusions (fully enclosed, glassy, with either a single small vapor bubble or no bubble) were prepared for analysis. Samples of the bulk tephra were analyzed by XRF at Washington State University for major and trace elements, and the analyses are presented in Johnson et al. (2008, 2009). Water and CO2 concentrations in melt inclusions were analyzed by Fourier Transform Infrared Spectroscopy (FTIR) at the University of Oregon. Concentrations of H2O and CO2 were calculated using Beer's law: c = MA/ρdε, where M is the molecular weight of H2O or CO2, A is the measured absorbance of the band of interest, ρ is the density of the hydrous basaltic glass (calculated following Luhr, 2001), d is the thickness of the melt inclusion and ε is the molar absorption coefficient. In most samples, water concentrations were calculated using the total OH peak at 3550 cm− 1 and an absorption coefficient of 63± 3 L/mol cm

(P. Dobson et al., unpublished data, cited by Dixon et al., 1995). In some instances, however, total H2O was calculated using an average of the molecular H2O peaks at 1630 cm− 1 and 5200 cm− 1 and the OH− peak at 4500 cm− 1. In these cases absorption coefficients were calculated based on major element compositions (Dixon et al., 1995). CO2 was calculated using the carbonate peaks at 1515 and 1435 cm− 1; an absorption coefficient was calculated (typically 290–300 L/mol cm) based on the major element composition of each sample (Dixon and Pan, 1995). The background around the carbonate peaks is complex, and thus it is necessary to subtract a carbonate-free spectrum from each sample spectrum to obtain a flat background (Dixon et al., 1995). We measured the absorbance of the carbonate doublet peaks using a peakfitting program (unpublished program by S. Newman). Major and minor element (including S and Cl) compositions of melt inclusions, their olivine hosts, and tephra groundmass glasses were analyzed on the Cameca SX-100 electron microprobe at the University of Oregon using a 15 kV accelerating voltage, 10 nA beam current (20 nA for olivine analyses), and a beam diameter of 10 μm (1 μm for olivine). The beam current was increased to 40 nA when analyzing S and Cl, and count times were increased to 80 s for S and 100 s for Cl. Using published analyses of FeO and Fe2O3 in whole rock samples of MGVF lavas (Hasenaka and Carmichael, 1985), we calculated relative magmatic oxygen fugacities (Kress and Carmichael, 1991) that vary from ΔNNO+ 0.5 to +1.0. These values allow us to predict a mean oxidation state of S in MGVF magmas that corresponds to an S Kα peak position in glass that is intermediate between those for anhydrite and pyrite (∼30% of the full shift between pyrite and anhydrite; Wallace and Carmichael, 1994). To correct for decreasing counting rates with time caused by alkali (Na, K) migration, and consequent increases for Si and Al, we used a correction routine that fits an exponential function to the varying count rates for these elements, allowing extrapolation to time zero. A combination of glass and mineral standards was used in the microprobe analyses.

Fig. 1. Digital elevation model of the Michoacán–Guanajuato Volcanic Field (MGVF) showing locations of volcanic centers sampled in this study. White contours indicate distances from the Middle America Trench (MAT). Inset shows map of Mexico and plate tectonic boundaries, with location of the MGVF denoted by a box (modified from Luhr and Carmichael, 1985). Dots indicate locations of volcanoes in Mexico and Central America.

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To determine the speciation of dissolved S in the melt inclusions, wavelength dispersive S Kα scans were performed on 2–4 inclusions per sample. Exposure to the electron beam causes an increase in sulfur oxidation state for glasses in which most S is present as S2−(Wallace and Carmichael, 1994) and a reduction of S6+ to S4+ in more oxidized glasses (Wilke et al., 2008). To minimize these effects, the sample was moved relative to the electron beam every 20 s. A peak-fitting program was used to locate the position of the S Kα peak, and the oxygen fugacity of the melt was calculated using the calibration of Wallace and Carmichael (1994), which has recently been confirmed using a larger dataset (Jugo et al., 2007). All melt inclusion data were corrected for the effects of postentrapment crystallization of olivine (Sobolev and Chaussidon, 1996) and diffusive loss of Fe (Danyushevsky et al., 2000). We correct for post-entrapment crystallization (PEC) by adding equilibrium olivine into the melt inclusion composition, in incremental fractions of 0.1 wt.%, until it is in equilibrium with its host olivine (as analyzed by electron microprobe). There are two variables used in calculating the equilibrium olivine composition: the KD value and the FeO/FeOT ratio. We used a KD of 0.3 ± 0.01 (Toplis, 2005) and FeO/FeOT values of 0.7–0.9, based on whole rock lava data (Hasenaka and Carmichael, 1985) for each cone. Following the procedure of Danyushevsky et al. (2000) we also corrected the inclusions, if necessary, for postentrapment Fe loss. To make this correction, we plotted the melt inclusion FeOT vs. MgO data and the bulk tephra XRF data and whole rock data from Hasenaka and Carmichael (1985). Inclusions with low FeOT compared to the bulk tephra and whole-rock data had FeO added back into their compositions until they matched the whole rock trend. Many inclusions required little to no correction for PEC and Fe loss because their analyzed compositions were close to equilibrium with the olivine host: the correction changed the Mg# by ≥0.05 for only 36% of all of the inclusions analyzed (Fig. 2a). The amount of postentrapment modification reflects the temperature difference between entrapment and eruption of an inclusion. A comparison of the corrected and uncorrected melt inclusion Mg#s and their host olivine compositions is shown in Fig. 2b. Inclusions requiring the largest correction were typically in the most Fo-rich olivine, probably because high-Fo olivines sometimes erupt in lower-Mg melts, such that the inclusions may experience a large cooling interval between entrapment and eruption. Both corrected and uncorrected (analyzed) values for major elements and volatiles in all melt inclusions are presented in Johnson et al. (2009). All melt inclusion major element and volatile data discussed in the text and shown in subsequent figures in this paper are corrected values. It should be noted that the H2O and CO2 values for Jorullo melt inclusions in Johnson et al. (2009) and used in this paper are slightly higher than the values originally reported in Johnson et al. (2008) because of changes made to the glass densities used in the Beer–Lambert Law. The higher values are due to use of compositionally dependent densities (primarily the large effect of H2O in lowering the glass density) in the Beer–Lambert Law whereas the lower values in Johnson et al. (2008) used a constant density value for all inclusions.

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Fig. 2. Effects of post-entrapment crystallization (PEC) and Fe-loss corrections on melt inclusion compositions. (a) Mg # of corrected melt inclusion composition minus the Mg# of the analyzed inclusion vs. olivine host forsterite content. The corrected Mg# values are for inclusions that were corrected for PEC and Fe-loss (if necessary) as described in the text. (b) Mg# of melt inclusions vs. olivine host forsterite content showing a comparison between corrected (symbols as in a) and uncorrected (x) melt inclusion compositions. Solid line is equilibrium line assuming a KD of 0.3, with the long dashed lines representing ±0.01 uncertainty in the KD value. Light dashed lines indicate the amount of olivine added to inclusions for the PEC correction, assuming no Fe loss. Values of olivine added would be greater than depicted by the lines for inclusions that were corrected for Fe loss.

Bulk compositions of the tephra samples are reflected in the associated phenocryst populations, which are dominated by Fo-rich olivine (Fo88–91) at many of the cones (Fig. 2b). In contrast, the more evolved magmas crystallized olivine (Fo75–87)±plagioclase±clinopyroxene phenocrysts. Some suites of melt inclusions from individual volcanoes show decreasing MgO and increasing K2O over time or within an individual eruption unit, suggesting evolution of melt compositions by fractional crystallization

3. Results 3.1. Melt compositions The melt inclusions from the MGVF trapped mostly basaltic to basaltic andesitic melt compositions (Fig. 3). Melt inclusions from the cinder cones are medium-K calc-alkaline in composition, whereas the maar locality (Hoya Alvarez) erupted more alkalic compositions. Although Hoya Alvarez plots in the shoshonitic field, its major and trace element characteristics are typical of intraplate alkaline or ocean island basalt (OIB) magmas (Johnson et al., 2009) rather than arc shoshonites (e.g., Vigouroux et al., 2008).

Fig. 3. Melt inclusion K2O and SiO2 concentrations. Most MGVF melt inclusions plot in the medium-K basalt and basaltic andesite fields. Hoya Alvarez inclusions are more alkalic, with lower SiO2 and higher alkali contents.

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and/or assimilation (Fig. 4). In general, the melt inclusions are less evolved, with lower incompatible element (K2O, TiO2, P2O5) concentrations, than the groundmass glasses (Fig. 4). 3.2. Volatile concentrations in melt inclusions MGVF melt inclusions trapped variably degassed melts with a wide range of dissolved volatile concentrations. The range in H2O measured for inclusions at a given cinder cone is typically between 1 and 4 wt.% H2O, with higher concentrations (up to nearly 6 wt.%) found at six localities (Fig. 5). Melt CO2 concentrations are highly variable, from levels below detection (b ∼ 50 ppm) to 2500 ppm. Volatile contents of alkali basaltic melt inclusions from the maar locality, Hoya Alvarez, are consistently different from those of the cinder cones, with low H2O (0.4–1.4 wt.%) and variable but elevated CO2 (up to 3000 ppm, with one measurement of ∼6000 ppm; Fig. 5c). Sulfur concentrations in melt inclusions from all volcanoes vary between 1000 and 2000 ppm (Fig. 6). Measured Cl contents in melt inclusions from the calc-alkaline cinder cones are mostly between 600 and 1400 ppm; Hoya Alvarez melt inclusions have lower Cl values (400–700 ppm) (Fig. 6). 4. Discussion To assess degassing of the MGVF basaltic melts, we must evaluate all of the magmatic processes that can affect volatile concentrations during magma storage, ascent and eruption. In the following sections, we first investigate the effects of fractional crystallization, assimilation, and boundary layer trapping on the melt inclusion compositions. We then calculate the melt inclusion trapping (olivine crystallization) pressures and use these pressures to address conditions of H2O and CO2 degassing in the MGVF melts. Our results suggest an important role for open-system gas fluxing during crystallization and ascent at many cones. Using robust sample suites that show evidence for degassing and fractional crystallization, we then assess the interplay amongst gas fluxing, exsolution of individual volatile species (H2O, S and Cl), vapor-melt partitioning of S and Cl, and crystallization. 4.1. Processes affecting melt inclusion compositions To interpret the record of magma degassing preserved by melt inclusions, we must first assess compositional variations related to conditions of magma storage. Most important are compositional controls exerted by crystallization and assimilation. Fractional crystallization of vapor-undersaturated magma can increase the

Fig. 4. Melt inclusion and matrix glass K2O and MgO concentrations. Melt inclusion compositions are less evolved than their respective matrix glasses. Error bars shown for melt inclusions (average value) and matrix glasses (shown on individual data points) represent one standard deviation uncertainties based on replicate electron microprobe analyses.

Fig. 5. Melt inclusion CO2 vs. H2O. Vapor saturation isobars (light gray), closed-system degassing paths (solid black), and isopleths of constant vapor composition (dashed lines) were calculated using the model of Papale et al. (2006). Most melt inclusions were trapped at pressures b 300 MPa. Closed-system degassing paths are shown for 5.8 wt.% initial H2O and 3000 ppm initial CO2 (in panel a), 4.6 wt.% initial H2O and 3000 ppm CO2 (a and b), and 3 wt.% initial H2O and 3000 ppm CO2 (c). The modeled degassing paths do not satisfactorily reproduce the variations in melt inclusion CO2 and H2O data. Note that a large number of the melt inclusions plot between the isopleths representing vapor with 40 and 75 mol% CO2. Average standard deviations for all melt inclusion data shown in a).

concentrations of volatile elements in the melt because such elements typically are incompatible. Even if the magma is vapor saturated, volatile components with low vapor-melt partition coefficients will increase during fractional crystallization because they do not partition strongly into the vapor phase. Conversely, fractionation of phases that contain volatile elements such as apatite (Cl, F) and sulfide (S) can deplete melt volatile contents even if the melt is not otherwise saturated with a vapor phase. Although quenched sulfide globules have not been observed in any MGVF samples, olivine crystals from Hoya Alvarez do contain small crystals of apatite, which suggests that the Cl concentrations in Hoya Alvarez melt inclusions have been affected by apatite fractionation. Assimilation of crustal wallrocks by basaltic melts is also common, as seen in studies of cinder cone

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Fig. 6. Melt inclusion and matrix glass S (a) and Cl (b) vs. H2O. Only a few inclusions show decreases in S with decreasing H2O (a), but matrix glass S contents are generally low. Melt Cl concentrations show no change with decreasing H2O, even in some matrix glasses.

eruptions like Parícutin (McBirney et al., 1987; Luhr, 2001) and Jorullo (Luhr and Carmichael, 1985; Johnson et al., 2008). Assimilation can affect both major and trace element concentrations of magmas, and the extent of this process can be determined if the assimilant composition is well constrained (e.g., Erlund et al., 2009). An additional complication is the possibility that melt inclusions do not faithfully record the compositions of the bulk melt from which they formed because of boundary layer effects during inclusion trapping. Melt inclusions form when the growing crystal traps small volumes of melt at the crystal-melt interface. Because melt inclusions can be trapped during relatively fast growth, elements that are slow to diffuse in the melt (e.g., Al2O3, P2O5) may be preferentially enriched in the boundary layer surrounding the growing crystal. Such elements may have higher concentrations in melt inclusions than in the bulk melt, as illustrated by recent crystallization experiments involving forsterite in the CMAS system (Faure and Schiano, 2005) and plagioclase and pyroxene in basaltic melt (Baker, 2008). To investigate processes that may have affected the compositions of trapped melt inclusions, we plot data for melt inclusions and bulk rock compositions as ratios of elements that are incompatible in olivine. We use P2O5 in the denominator because it is a slowly diffusing element in silicate melts (Fig. 7). Differences in these ratios between melt inclusion and bulk rock samples can be caused by boundary layer enrichment of the melts during inclusion formation, assimilation of crustal rocks, mixing of crystals and inclusions that formed in compositionally distinct magmas, and fractional crystallization involving titanomagnetite or apatite. Figs. 7 and 8a show that there is generally good agreement between melt inclusion and whole rock data at Jorullo, Astillero, San Juan, and Hungaro. In contrast, some Hoya Alvarez and Pelon inclusions have lower TiO2/P2O5 concentrations than bulk rock values, which suggests crystallization of

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titanomagnetite occurred before melt inclusion entrapment. In this case, TiO2 decreases in the melt with progressive fractionation, but other elements, including volatiles, should not be affected. The Hoya Alvarez magma also crystallized apatite, which likely would lower Cl concentrations. Parícutin magma compositions evolved over time (Wilcox, 1954; Luhr, 2001; Erlund et al., 2009). Melt inclusion compositions from Parícutin reflect this change (Fig. 7c). Early melt inclusions plot near the early lavas, but as the eruption progressed, the K2O/P2O5 ratio of the melts increased. This progressive evolution is likely the result of substantial assimilation of granitic bedrock, which has high K2O/P2O5 and low TiO2/P2O5 values (McBirney et al., 1987; Erlund et al., 2009). More complex magmatic processes are suggested by melt inclusions that show significant variation from whole rock values (Fig. 7d). Here the data cluster into two (San Miguel) and possibly three (La Loma) distinct melt inclusion populations, only some of which plot near the bulk rock values. Variations in TiO2/P2O5 at San Miguel and La Loma could be the result of titanomagnetite fractionation, but the variations in K2O/P2O5 must reflect either mixing of compositionally distinct magmas or boundary layer enrichment of slowly diffusing P2O5. For San Miguel, the lack of correlation between K2O/P2O5 and TiO2/P2O5 is not consistent with boundary layer enrichment of P2O5 and thus seems more likely to be caused by magma mixing. For La Loma, the two ratios are correlated, but one group of inclusions has higher ratios than the whole rock, which is also inconsistent with boundary layer enrichment of P2O5. We conclude that the compositional variations at La Loma are related to combined magma mixing and titanomagnetite fractionation. To further test for possible boundary layer enrichment effects, we compared CaO/Al2O3 and K2O/P2O5 as a function of inclusion size and morphology (Fig. 8). In general, we find no evidence of anomalous enrichment in slowly diffusing elements, even in small inclusions or rapid growth forms like skeletal or hopper crystals. Kent (2008) examined major element compositional data for olivine-hosted melt inclusions from a number of localities and similarly found little evidence for boundary layer effects (see also Métrich and Wallace, 2008). Data for Icelandic picrites suggest that fractionation of CaO/ Al2O3 attributable to boundary layer effects is observed only in small inclusions (≤15 µm in diameter; Kuzmin and Sobolev, 2004). 4.2. Vapor saturation pressures The considerable range of H2O and CO2 in melt inclusions from each of the volcanic centers that we studied (Fig. 5) indicates melt inclusion entrapment (and olivine crystallization) over a wide range of pressures. The pressure at which a melt inclusion was trapped can be estimated by using the dissolved H2O and CO2 to calculate a vapor saturation pressure (Newman and Lowenstern, 2002; Papale et al., 2006; Moore, 2008). A recent comparison of solubility models (Moore, 2008) has shown that the model of Papale et al. (2006) provides the best estimates of vapor saturation pressure and vapor composition for typical mafic arc magma compositions. Calculated vapor saturation pressures for the H2O and CO2 contents of the MGVF melt inclusions using the Papale et al. (2006) model indicate entrapment pressures of b50 to 300 MPa. These pressures are substantially lower (by ∼40%, on average) than those calculated with VolatileCalc (Newman and Lowenstern, 2002). The reason for this is that the solubility of CO2 in mafic arc melt compositions is higher than is predicted with VolatileCalc (Moore, 2008), which is calibrated for tholeiitic to silica-undersaturated compositions (Newman and Lowenstern, 2002). Mafic magmas are typically vapor saturated at middle to upper crustal pressures because of the relatively low solubility of CO2 in silicate melts (Métrich and Wallace, 2008). A complication, however, to using melt inclusion H2O and CO2 contents to calculate inclusion trapping pressure is the formation of shrinkage bubbles. Shrinkage bubbles form during post-entrapment cooling because of the much greater thermal contraction of the melt relative to the crystal host. The

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Fig. 7. Melt inclusion (points) and whole rock (shaded fields) values of K2O/P2O5 vs. TiO2/P2O5. (a) Melt inclusions from Jorullo and Hoya Alvarez plot near the whole rock compositions. (b) Melt inclusions from San Juan and Hungaro have compositions that overlap with whole rock values. (c) Melt inclusions from Astillero generally plot near whole rock values, but Parícutin inclusions show evolution toward higher K2O/P2O5, which likely represents assimilation of granitic bedrock (McBirney et al., 1987). However, melt inclusions from this study and the early inclusions from Luhr (2001) plot near the early lava flows (McBirney et al., 1987), suggesting that these early melts were not affected by assimilation. (d) Melt inclusions from San Miguel, Pelon, and La Loma are variable; some San Miguel and La Loma inclusions plot with the bulk rock compositions, whereas others are compositionally different. Pelon melt inclusions overlap slightly with whole-rock values but show variable TiO2/P2O5, suggesting minor fractionation of titanomagnetite.

formation of a shrinkage bubble lowers the initial dissolved CO2 content of a melt inclusion because CO2 has such low solubility (Anderson and Brown, 1993; Cervantes et al., 2002), but the effect is greatest in H2O-poor melts because they exsolve vapor that is very CO2 rich (see vapor isopleths in Fig. 5). Given the large uncertainties in measuring vapor bubble volumes (because of irregular inclusion shapes) and estimating the pre-eruption (before eruptive expansion) bubble size (see Appendix A), we have chosen to use the analyzed dissolved CO2 contents in all figures throughout this paper and to use vapor saturation pressures based on the analyzed rather than restored CO2 values (Fig. 5). It should be noted that these are minimum values and that the real trapping pressures for inclusions at the higher pressure end of our dataset could have been as much as 100 MPa higher (Appendix A). However, because many inclusions trapped at relatively low pressures contain either no vapor bubble or have CO2 below detection, the pressures shown in Fig. 5 and subsequent figures should be close approximations to the true trapping pressures for the inclusions at the lower pressure end of our dataset. Thus the effect of accounting for shrinkage bubbles would be to spread out the data shown in Fig. 5 (i.e., higher pressure inclusions probably trapped up to 100 MPa higher than shown in Fig. 5, but lower pressure inclusions plotting close to true trapping pressures). 4.3. Sulfur solubility and oxygen fugacity The oxygen fugacity of basaltic magmas influences the speciation of S, which in turn affects S solubility and vapor-melt partitioning. Experimental studies have demonstrated that sulfate is the dominant S species in basaltic melts at high fO2 (NNNO + 1) and that sulfide is the dominant species in low-fO2 melts (Carroll and Rutherford, 1985,

1987; Luhr, 1990; Jugo et al., 2005). Furthermore, S solubility in mafic magmas increases dramatically at higher oxygen fugacities (Carroll and Rutherford, 1985, 1987; Luhr, 1990; Jugo et al., 2005). Oxygen fugacities and S contents of the MGVF melt inclusions are shown in Fig. 9. Based on measured S Kα peak positions, oxygen fugacities of MGVF magmas range from NNO + 0.5 to NNO + 1.4. There is substantial scatter in the data, but melts with high S contents (1500–2000 ppm) have higher fO2 (NNO + 1.1 to + 1.4). Our data generally agree with the experimental differences in sulfide and sulfate solubilities (Jugo et al., 2005), which result in a strong increase in total dissolved S in basaltic melts at NNO + 1.1. The MGVF melt inclusions plot around this transition: Hoya Alvarez melts, with lower fO2, were likely saturated in sulfide, Jorullo melts plot around the transition in stability from sulfide to sulfate, and the other melt inclusions generally plot in the region of much higher S solubility where solubility is controlled by sulfate saturation. 4.4. Degassing of major volatile elements (H2O and CO2) Calculated vapor saturation isobars (Fig. 5) show that melt inclusions at each volcanic center were trapped over a wide range of pressures (typically from b50 MPa to 300 MPa). The H2O and CO2 variations probably represent trapping of variably degassed melts as magma ascended and crystallized within the upper crust. In detail, however, the data show significant scatter (beyond analytical error) compared to equilibrium degassing models, and closed-system degassing models cannot successfully predict most of the H2O and CO2 variations (Fig. 5). This misfit is illustrated in Fig. 10a, which shows changes in melt inclusion H2O content for four different cinder cones as a function of entrapment pressure (calculated using

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Fig. 8. Jorullo melt inclusion Al2O3 vs. CaO (a) and CaO/Al2O3 vs. melt inclusion size (b). In (a), the Jorullo melt inclusion data plot near the whole rock data, or along the trend representing olivine fractionation, suggesting the melt inclusion compositions accurately represent the bulk melt compositions. In (b), the ratio of CaO/Al2O3 in the melt inclusions overlaps with the whole rock values, and shows no variation with melt inclusion size, suggesting boundary layer enrichment did not occur.

degassing curve. Thus although we cannot rule out some postentrapment loss of H, we conclude that such loss is not the main process responsible for variations in melt inclusion H2O contents. The implications of this interpretation for the ascent rates of mafic magma during violent Strombolian eruption are discussed later in Section 5.1. A possible explanation for the wide range of H2O and CO2 values is that ascending magmas were affected by gas fluxing, in which CO2-rich vapor percolates through the system from below, where it is released by magma degassing deeper in the system (Anderson et al., 1989; Rust et al., 2004; Spilliaert et al., 2006a; Johnson et al., 2008; Vigouroux et al., 2008; Blundy et al., 2010). The effect of this fluxing by CO2-rich vapor is to shift melt H2O contents to lower values along vapor saturation isobars as the melts attempt to re-equilibrate with the introduced vapor phase. To illustrate the effects of gas fluxing, we have shown isopleths corresponding to vapors with 40 and 75 mol% CO2 in Fig. 5. In the majority of sample suites there is a subset of inclusions for which the H2O and CO2 contents roughly co-vary, forming an array that parallels the isopleths representing 40 to 75 mol% CO2. This range of vapor compositions is similar to those estimated for mafic magmas at Mount Etna (Spilliaert et al., 2006a), Jorullo (Johnson et al., 2008), Popocatépetl (Roberge et al., 2009), and high-K minettes and basanites in the Colima Graben to the west of the MGVF (Vigouroux et al., 2008). For other cones such as Parícutin, the evidence for gas fluxing is less clear, though this may be an artifact of the small number of inclusions analyzed. Based on the relations summarized above, we suggest that the wide range of H2O contents for inclusions with measurable CO2 (Fig. 5) can be explained by variations in the extent of buffering provided by gas fluxing during magma ascent. In contrast, inclusions shown in Fig. 5 with CO2 below detection can be explained by equilibrium closed- or open-system degassing. Evidence for fluxing of a vapor phase with 40–75 mol% CO2 in basaltic melts in the MGVF, the Colima graben, Popocatépetl, and Mount Etna suggests that the CO2 contents of the parental basaltic magmas were similar. A compilation of volatile contents from arc basalts shows that many melts are in equilibrium with vapors with 40–75 mol% CO2 (Fig. 11). This suggests that much of the variation in H2O and CO2 in basaltic magmas can be explained by primitive mafic magmas with ≥ 0.7 wt.% CO2, as suggested by Wallace (2005). The evidence for the pervasive effects of gas fluxing in volcanic conduit systems suggests that such magmas pond and degas CO2-rich vapor at lower crustal depths, probably as a result of storage and recharge in lower crustal sills (e.g., Annen et al., 2006). Additional gas could be

measured H2O and CO2). All four show generally higher CO2 relative to predicted degassing curves at intermediate pressures. Similar scatter in H2O and CO2 values have been documented for other magmatic systems (Rust et al., 2004; Atlas et al., 2006; Spilliaert et al., 2006a; Vigouroux et al., 2008; Roberge et al., 2009; Blundy et al., 2010), and the possible causes of this in the Jorullo melt inclusion data have been discussed in Johnson et al. (2008). Of particular concern is the possibility that post-entrapment diffusive loss of either H (Danyushevsky et al., 2002; Massare et al., 2002) or molecular H2O (Portnyagin et al., 2008) through the olivine host crystals has shifted the inclusions towards lower H2O values than when they were originally trapped. However, the loss of 2.5 to 4 wt.% H2O that would be required to account for our data (assuming that all inclusions with measurable CO2 initally plotted along the closed-system degassing curve in Fig. 5) would cause considerable crystallization inside the inclusions (e.g., Danyushevsky et al., 2002; Massare et al., 2002) and formation of a substantial shrinkage vapor bubble. These features are not observed in the high CO2, low H2O inclusions, leading us to conclude that significant H2O diffusive loss through the host olivine has not occurred. Additionally, we see no correlation of either melt inclusion Fe loss or extent of post-entrapment crystallization with the deviation of melt inclusion H2O values from the closed-system

Fig. 9. Concentration of S in MGVF melt inclusions vs. relative oxygen fugacity of the melts (ΔNNO) based on measurements of S Kα wavelength shifts as described in the text. The dashed line shows the sulfide solubility limit at lower oxygen fugacities and the transition from sulfide to sulfate saturation at ∼NNO + 1.1, which causes an increase in the concentration of S dissolved in basaltic melts (Jugo et al., 2005). The experimentally determined sulfate solubility limit is 1.3 wt.% S at 1 GPa pressure (Jugo et al., 2005).

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supplied by magma as it ascends through the lower to middle crust provided that the magma transport system is ‘open’ to gas fluxing. The main requirement for any particular batch of magma to be able to provide the gas is that it have sufficient CO2, and lose gas at sufficiently high pressure, to create vapor with ≥40 mol% CO2. Other basaltic volcanoes where gas fluxing may play a role include Arenal and Irazú, where degassing models can reproduce trends in melt inclusion H2O and CO2 data only if additional buffering gas is present (Wade et al., 2006; Benjamin et al., 2007; see also Métrich and Wallace, 2008). However, in these cases, smaller mass fractions and less CO2-rich vapor are needed to explain the degassing trends. 4.5. Degassing-induced crystallization and its relationship to gas fluxing Fluxing of the melts with a CO2-rich vapor phase should also affect melt crystallization behavior. Consideration of H2O-undersaturated phase equilibria for basaltic melts shows that gas fluxing causes

Fig. 11. Compilation of CO2 and H2O from basaltic melt inclusions worldwide (modified from Wallace, 2005). Dashed lines are vapor composition isopleths, thin grey lines are vapor saturation isobars, and the thick solid line represents the closed-system degassing path for a basalt with 5 wt.% H2O and 7000 ppm CO2 (calculated using an average calc-alkaline MGVF composition at 1125 °C and the model of Papale et al., 2006). Most of the melt CO2 and H2O concentrations fall between the 40 and 75 mol% CO2 isopleths.

crystallization because CO2-rich gas removes dissolved H2O from melt even though the melts are H2O undersaturated (Johnson et al., 2008). This causes crystallization of olivine in ascending or ponded magmas in the absence of any temperature decrease. In contrast, melt inclusions with very low CO2 (e.g., values below detection in Fig. 5) are likely to represent melts that degassed along equilibrium closedor open-system trends and were relatively unaffected by gas fluxing. Along such paths, ascending melts crystallize olivine only once they have lost nearly all of their CO2 and hence become saturated with H2O-rich vapor, which occurs at ∼ 150–200 MPa for melts similar in composition to many of the MGVF inclusions (Moore and Carmichael, 1998; Johnson et al., 2008). To examine crystallization and its relationship to H2O loss, we have plotted melt inclusion H2O contents as a function of the extent of crystallization, as estimated from melt inclusion K2O for selected cinder cones (Fig. 10b). Because K2O is incompatible during crystallization, the concentration of K2O in a melt inclusion relative to the K2O concentration of the least evolved (parental) melt from a given volcano gives an estimate of how much crystallization differentiation has occurred. Although the data show considerable scatter, the most evolved melt inclusion compositions (requiring the most crystallization differentiation) were generally trapped at lower pressures and with lower H2O contents. This supports the interpretation that H2O loss from melts provides a driving force for crystallization. The H2O contents and inferred extents of crystallization for the melt inclusions are also consistent with phase equilibrium results (Fig. 10b). Because many of the melt inclusions deviated from the closed system equilibrium degassing trend (Fig. 10a), our results further suggest that gas fluxing plays a role in driving crystallization. Fig. 10. Melt inclusion entrapment pressure vs. H2O (a) and melt inclusion H2O vs. % fractional crystallization inferred from K2O contents (b). Panel (a) shows closed-system degassing paths (calculated using the model of Papale et al., 2006) for melts with 5.8, 4.6, and 3 wt.% H2O and sufficient CO2 to be vapor saturated at 300 MPa. Many melt inclusions plot at lower H2O contents than would be expected under closed-system degassing, an observation that could be explained by fluxing of CO2-rich vapor through the system. In panel (b), extents of fractional crystallization required to produce a given inclusion from the least evolved melt composition at each volcano were calculated based on variations in the K2O contents. Also shown are curves showing the relationship between H2O and % crystallization during adiabatic ascent for a calc-alkaline basaltic andesite with 6.7 wt.% MgO (Moore and Carmichael, 1998) and a basaltic melt from Jorullo with 10.5 wt.% MgO (Johnson et al., 2008; based on pMELTS calculations, Ghiorso et al., 2002). These compositions bracket the range of likely starting melt compositions for the MGVF melt inclusions.

4.6. Degassing of volatiles (H2O, S, Cl) during crystallization Crystallization increases the concentrations of dissolved volatiles in the remaining melt and thus can partially to fully offset the decrease of volatiles caused by degassing. To quantify the degassing behavior of volatiles, we focus on four MGVF localities — Jorullo, Parícutin, Pelon and San Juan. The melt inclusions and matrix glasses from these volcanoes have the greatest ranges in K2O and volatile concentrations, which makes robust interpretations about degassing behavior possible. We correct for the effect of crystallization by normalizing the concentrations of H2O, S, and Cl to K2O (Fig. 12). Because assimilation at Parícutin affected the melt compositions later

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Fig. 13. S/K2O and Cl/K2O for melt inclusions and matrix glasses. Covariation of these ratios indicates concurrent degassing of S and Cl during crystallization. The datasets with substantial variation show that S degasses to a greater extent than Cl because the data arrays indicate negative intercepts on the S/K2O axis. High matrix glass Cl contents supports this interpretation.

S/K2O ratios than even the most degassed melt inclusions (Fig. 12), indicating loss of H2O and S during eruptive degassing. In contrast, at all three volcanic centers, Cl/K2O ratios in groundmass glasses are similar to or very slightly lower than the values of the most degassed melt inclusions, indicating little loss of Cl occurred during eruption. Variations in S/K2O and Cl/K2O in the melt inclusions (Fig. 13) suggest that as much as 70–80% of initial S and 17–53% of initial Cl were degassed during crystallization. Further low pressure degassing occurred during eruption, and matrix glass values indicate that 98–99% of initial S and only 50–66% of initial Cl had degassed by the time of eruptive quenching. 4.7. Vapor-melt partitioning of S and Cl

Fig. 12. Ratios of volatiles/K2O vs. K2O in melt inclusions and matrix glasses from Jorullo, Parícutin (light triangles from this study, dark triangles from Luhr, 2001), Pelon and San Juan. The H2O contents of all matrix glasses are estimated to be ∼0 wt.%, based on the electron microprobe H2O-by-difference method. Vectors in panel (c) indicate general paths for melts undergoing crystallization, degassing, and degassing + crystallization. Decreases in H2O/K2O (a), S/K2O (b), and Cl/K2O (c) with increasing K2O indicate degassing of these volatiles during fractional crystallization.

in the eruption, we included only data from tephra from the earliest stages of the Parícutin eruption, which show no effects of assimilation (Erlund et al., 2009). Decreases in volatile/K2O ratios with increasing K2O in all melt inclusions indicate degassing of H2O, S, and Cl during crystallization (Fig. 12). However, the amount of degassing of these volatiles varies. In all samples, S appears to degas to a greater extent than Cl, as illustrated by the co-variation of S/K2O and Cl/K2O (Fig. 13). Although both of these ratios decrease, S/K2O decreases more rapidly indicating more extensive degassing of S. We also analyzed S and Cl in matrix glasses to assess syn-eruptive degassing. For all samples groundmass glasses have lower H2O/K2O and

Using the melt inclusion and matrix glass S, Cl, H2O and K2O concentrations, we modeled the vapor-melt partitioning of S and Cl during differentiation (see Appendix B for a complete description). Using similar methods as applied to Mount Etna melt inclusions by Spilliaert et al. (2006b), we calculated vapor-melt DS and DCl values assuming three different physical degassing processes: (1) open-system degassing, in which exsolved vapor is lost as it forms, (2) closed-system degassing, in which exsolved vapor is retained and remains in equilibrium with the dissolved volatiles in the melt, and (3) closedsystem degassing in which the melt contains an initial amount of exsolved vapor. In case (3), the initial vapor is assumed to contain no Cl or S. This might be the case where magma and entrained vapor ascended rapidly from lower crustal or upper mantle depths, because at high pressures, the high solubilities and low vapor-melt partition coefficients of Cl and S cause the vapor phase to have very low Cl and S. Importantly, the equations for a closed system with initial exsolved vapor also apply to an open system fluxed with a fixed mass of vapor because the equations describe mass balance for a system that contains more gas than it can generate internally. The equations are valid for an open system fluxed with vapor provided the fluxing vapor initially contains no Cl or S. In both of these physical scenarios (closed system with initial exsolved vapor, open system fluxed with vapor), the effect of the initial exsolved vapor during subsequent cooling, crystallization and decompression is to draw more of the dissolved S and Cl out of the melt and into the vapor phase. Fig. 14a shows vapor-melt DS values calculated from the Jorullo melt inclusion data. For our calculations, D-values equal to one indicate equal partitioning of the volatile between the vapor and the

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melt, and D-values greater than one indicate preferential partitioning into the vapor phase. Results are shown for each of four models (open system, closed system, closed system plus 1% and 2% initial exsolved vapor). For most inclusions, the closed-system models yield the highest calculated DS values. Values for closed-systems with initial exsolved vapor are generally lower because the presence of the additional vapor, which is assumed to contain no S, draws more dissolved S out of the melt. In this case, a lower D-value is required to account for a given set of initial and final dissolved S concentrations. However, the differences between models are generally relatively small compared to the uncertainties in calculated D values for inclusions in any given model. As described above, closed-system degassing with initial exsolved vapor is quantitatively the same as the open-system gas fluxing process that is suggested by the H2O and CO2 variations. Therefore we have chosen to use the D-values modeled for a closed system with 1% exsolved vapor for the MGVF data. It should be noted that the variation in D-value between this and the other models is relatively small (±20% of the D-value).

In Fig. 14b, we compare the calculated vapor-melt DS values for the closed system plus 1% initial exsolved vapor to values predicted for melts of similar temperature, pressure, oxygen fugacity and melt composition using the empirical model of Scaillet and Pichavant (2005); (see also Roberge et al., 2009, for a description of procedures). The model calculations indicate that vapor-melt partitioning of S is dependent on pressure, with DS values b∼10 at pressures N200 MPa and D S rapidly increasing at pressures b200 MPa. This is consistent with the high calculated DS values based on the groundmass glasses and the lower values at higher pressures calculated from the melt inclusion data, but the values calculated for most of the lower pressure melt inclusions are systematically lower than predicted by the model curves. This could be the result of shrinkage bubbles causing us to underestimate the true trapping pressures. For comparison, melt inclusion data from Etna yield closed-system degassing DS values from 0–60 that increase with decreasing pressure (Spilliaert et al., 2006b), similar to the range of values for our data. However, our estimated Ds values are lower than have been inferred at some other arc volcanoes (e.g., Wade et al., 2006; Benjamin et al., 2007). Similar plots of vapor-melt DCl values are shown in Fig. 15. As for S, lower values of DCl generally result from the closed-system models with initial vapor. Values are shown in Fig. 15b for the case of 1% initial vapor as being representative of the gas fluxing case inferred from the H2O and CO2 data. Recent experimental work on an Etna basalt composition indicates DCl values of 6 at 200 MPa and 11–14 at 1– 25 MPa under NNO conditions (Alletti et al., 2009), thus demonstrating that Cl partitions more strongly into the vapor phase at lower pressures. Values of DCl inferred from the MGVF melt inclusions are in general agreement with these values. Melt inclusion data from Etna yield closed-system degassing DCl values of 0–0.65 (Spilliaert et al., 2006b), lower than the experimental values. We emphasize that for the MGVF melt inclusion data, the uncertainties in the calculated DS and DCl values for any given degassing model (open system, etc.) are large enough relative to the differences between the different models that we cannot rigorously test between models. In the case of Etna, the compositions of volcanic gases were shown to be incompatible with an open-system degassing model (Spilliaert et al., 2006b). In our case, we lack gas data to test the different degassing models. However, our results do show that open-system fluxing with CO2-rich gas, as inferred from the H2O and CO2 data, is not incompatible with the observed depletions in S and Cl that are caused by degassing. 5. Implications for monogenetic basaltic eruptions 5.1. Ascent rates of mafic magma during violent Strombolian eruptions

Fig. 14. Melt inclusion S versus calculated S vapor-melt partition coefficient (DS) for Jorullo (a) and pressure versus DS for Jorullo, Parícutin, Pelon, San Juan and matrix glasses(b)(see Appendix B for details of the D-value calculations). Panel (a) illustrates the variations in the calculated D value depending on the model conditions. The maximum variation between our chosen DS values (closed + 1% vapor model) and the other modeled D values is ±25%. Panel (b) illustrates the pressure dependence of DS. Melt inclusion DS values were calculated using the closed system +1% vapor model (see text for discussion). Matrix glass DS values were calculated using the open system model, which should more closely represent the conditions during eruption and extensive vesiculation of the matrix glasses. Also shown are curves illustrating the pressure dependence of S vapor-melt partitioning calculated with the model of Scaillet and Pichavant (2005). Calculations were made for basaltic melts at oxygen fugacities ranging from ΔNNO to ΔNNO +1.5 at 1150 °C. Increases in oxygen fugacity (at a constant pressure) result in higher values of DS.

As discussed above, we interpret much of the variation in melt inclusion H2O contents to reflect conditions at the time of inclusion entrapment rather than post-entrapment H loss as a result of diffusion through the olivine host crystals. Because it has been shown experimentally that such diffusion is relatively rapid (Danyushevsky et al., 2002; Massare et al., 2002; Portnyagin et al., 2008), the moderate to high H2O contents of many MGVF inclusions provides information about the ascent rates of mafic magma from depths of last crystallization during violent Strombolian eruptions. Recent modeling has shown that mass eruption rates typical of violent Strombolian eruptions range from 103–105 kg/s (Pioli et al., 2009). For bubbly magma flowing upwards through a conduit with a 5-m radius, this corresponds to ascent rates of ∼ 0.01 to 1 m/s. We use a simple linear parameterization of the experimental results of Massare et al. (2002; their Fig. 4) to constrain the fraction of the initial H2O content of a melt inclusion that would be lost as a function of time. We recognize, however, that actual loss rates will be much more variable because they depend on many factors such as inclusion

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Fig. 15. Jorullo melt inclusion Cl contents and calculated Cl vapor-melt partition coefficients (a) and pressure versus DCl for Jorullo, Parícutin, Pelon and San Juan melt inclusions and matrix glasses (b). As in Fig. 14, panel (a) illustrates the variation in DCl depending on the model conditions chosen. The variations between our chosen DCl (closed + 1% vapor model) and that of the other models is ∼±20%, which is within the error of the individual DCl values. Panel (b) illustrates the relationship between pressure and vapor-melt partitioning of Cl. As in Fig. 14, the melt inclusion DCl values are for the closed system + 1% vapor model, and the matrix glass DCl values were calculated for the open system model. Also shown are experimental data from Alletti et al. (2009), which indicate a slight increase in DCl values with decreasing pressure.

and host olivine size and temperature (Qin et al., 1992). The H2O solubility in mafic melt (Fig. 16a) provides an estimate of how much H2O might be lost from a melt inclusion with 4 wt.% initial H2O as it attempts to re-equilibrate with external melt during ascent from 5 km depth to the surface. As shown in Fig. 16a, at depths of ∼ 1 to 2 km during ascent, melt inclusions will begin to experience the strong drive for H loss that occurs as solubility in external melt drops eventually to 0.1 wt.%, the 0.1 MPa solubility value, when the inclusion reaches the surface. The results of these model calculations (Fig. 16b) suggest that melt inclusions in magmas rising at 0.1 to 1 m/s will lose ≤30% of their original H2O, consistent with our observations and interpretations. However, these calculations indicate that melt inclusions in magma rising at b0.1 m/s could be affected by large H2O losses. Because H2O solubility in melts changes very strongly from 0.1 to ∼ 50 MPa (Fig. 16a), the time spent in final ascent towards the surface and at high temperature in surface lava flows before quenching is when the strongest drive for diffusive loss will occur. This is why melt inclusions from lava flows, which cool slowly at the Earth's surface, are most commonly affected by diffusive H loss (e.g., Hauri, 2002). The fact that we find no CO2-bearing melt inclusions with

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Fig. 16. The pressure dependence of H2O solubility (a) and the dependence of H2O loss (via H diffusion) on magma rise speed (b) calculated as described in the text. The curve in (b) shows % loss of initial inclusion H2O during ascent from 5 km depth. As magma ascends from this depth, for an inclusion with 4 wt.% initial H2O external melt H2O drops very rapidly with decreasing pressure, and therefore the drive for diffusive loss becomes very large.

b1 wt.% H2O (Fig. 5; except for Hoya Alvarez, which has inherently low H2O) means that either ascent rates for all of these eruptions were ≥0.1 m/s or that the rates of diffusive H loss in nature are much slower than those in the 1 bar experiments performed by Massare et al. (2002). Our results suggest that melt inclusions from tephra deposited by violent Strombolian eruptions should be minimally affected by diffusive H loss during their ascent towards the surface. Prior to ascent, however, melt inclusions are susceptible to diffusive H loss or gain, and thus may record conditions of storage in the middle to upper crust rather than their original H2O contents from the time of trapping. This underscores the need to analyze large populations of inclusions from a given volcano or deposit: the highest H2O contents can be used to infer the original magmatic concentration (e.g., Sisson and Layne, 1993; Bureau et al., 1998; Cervantes and Wallace, 2003), whereas the entire melt inclusion population provides detailed information on magma storage en route to the surface (e.g., Roggensack et al., 1997; Blundy and Cashman, 2005; Spilliaert et al., 2006a). Analysis of multiple volatile species permits more complete evaluation of magmatic degassing patterns (e.g., Spilliaert et al., 2006b; Métrich and Wallace, 2008).

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5.2. Gas emissions during monogenetic eruptions The MGVF melt inclusion and matrix glass data can be used to make estimates of the masses of H2O, S and Cl gas released during monogenetic cinder cone eruptions. The estimated extents of volatile degassing (98–99% S, 30–67% Cl) are very similar to estimates for Etna (N95% S, 22–55% Cl; Spilliaert et al., 2006b). To calculate gas emissions, we used the mass eruption rates for Parícutin summarized by Pioli et al. (2009), volumetric estimates of 0.5 km3 (lava), 1.25 km3 (tephra) and 0.2 km3 (cone) for Jorullo (Luhr and Carmichael, 1985), and densities of 2380 kg/m3 (lava), 1460 kg/m3 (tephra), and 1800 kg/m3 (cone material) (Fries, 1953; Pioli et al., 2009) (Table 1). Our data suggest that degassing during the eruption of Jorullo could have released 1.9 × 1011 kg H2O, 1.3 × 1010 kg SO2, and 3 × 109 kg HCl into the atmosphere, with similar results found at Parícutin (1.9 × 1011 kg H2O, 1.4 × 1010 kg SO2, 2.1 × 109 kg HCl; Table 1). Considering that these eruptions continued for 15 and 9 years, respectively, we can estimate average daily outputs of volcanic gases into the atmosphere. However, the time and volume constraints for the Jorullo eruption are poorly known and we know that the mass eruption rates changed over time during the eruption of Parícutin (e.g., Pioli et al., 2009), so these estimates do not reflect the maximum fluxes or the variability of flux during the eruptions. Assuming that the mass of SO2 degassed was constant over the course of the eruptions, we calculated daily fluxes of 2320 t/d SO2 for Jorullo and 4360 t/d SO2 for Parícutin (Table 1). The SO2 fluxes calculated here are higher than SO2 gas emissions measured at both Kilauea (60– 210 t/d; Hager et al., 2008) and the 2007 eruption of Stromboli (620 t/ d; Burton et al., 2009), but are similar to estimates made for the 2002 eruption of Mt. Etna (7340 t/d; Spilliaert et al., 2006b) and recent gas emissions at Mt. Etna (on average 3250–3530 t/d; Salerno et al., 2009). Although our calculated gas emissions are minimum estimates, as they neglect any SO2 derived from a pre-eruptive vapor phase (e.g., Wallace, 2005), they show that monogenetic eruptions are capable of releasing relatively large amounts of volcanic gases into the atmosphere.

6. Conclusions We analyzed the compositions and volatile contents of olivinehosted melt inclusions from monogenetic volcanoes in the MGVF region of Mexico. Using these data, we were able to assess degassing and melt compositional evolution within these systems. Most of the analyzed melt inclusions have incompatible element ratios that are similar to their host magmas, suggesting the inclusions are represen-

tative of the bulk melt compositions from which they formed and were not affected by boundary layer enrichment. The melt inclusion data reveal that degassing in basaltic systems is complex. The CO2 and H2O variations record a wide range of entrapment pressures (from near-surface to ∼ 300 MPa). Compositional variations in melt inclusions show that extensive crystallization accompanied magma ascent as a consequence of both decompression-degassing and fluxing with a CO2-rich (40–75 mol%) gas. Interestingly, this vapor composition is similar to those estimated for CO2-fluxing in other basaltic systems, suggesting that primary basaltic magmas in arcs have similar and high initial CO2 concentrations. The chemical evidence for gas fluxing in these monogenetic systems reinforces physical models of gas fluxing (Krauskopf, 1948; Pioli et al., 2008) and suggests that fluxing occurs not only in the shallow conduit system but also deeper where phenocryst growth is occurring in response to H2O loss. Extensive degassing of S and Cl during magma ascent and crystallization begins at pressures of approximately 50 MPa. Using the relationship between degassing and crystallization, we calculated apparent vapor-melt partition coefficients for S and Cl. Our results show that S partitions more strongly into the vapor phase than Cl with decreasing pressure, consistent with published experimental data and solubility models. The partitioning behavior inferred from the melt inclusion data are consistent with the open-system gas fluxing model suggested by the H2O and CO2 data. Acknowledgments We would like to dedicate this manuscript to the memory of our friend and colleague, Jim Luhr, whose scientific work in Mexico and elsewhere has inspired us in so many ways. P.W. in particular owes an immeasurable debt to Jim for his guidance and forbearance throughout many years of work in the Trans-Mexican Volcanic Belt. We would like to thank Nathalie Vigouroux for S Kα wavelength measurements and oxygen fugacity calculations, Colleen Donegan for assistance with fieldwork, sample preparation and analysis, Rory San Filippo for assistance with sample preparation and analysis, and John Donovan for assistance with the electron microprobe analyses. We would also like to thank Nicole Métrich, Hiroshi Shinohara, and Patrick Allard for helpful reviews and assistance in calculating vapor-melt partition coefficients. Finally, we are grateful to Mark Ghiorso for making calculations with the Papale et al. (2006) solubility model possible through the OFM Research website. This research was supported by NSF grants EAR-0309559 and EAR-0510493. Appendices A and B Supplementary material associated with this article can be found, in the online version, at doi: 10.1016/j.jvolgeores.2010.02.017.

Table 1 Volatile emissions during the eruptions of Jorullo and Paricutin.

Erupted mass (kg)a Initial S(wt.%) Initial CI (wt.%) Initial H20 (wt.%) %loss S % loss CI % loss H2O SO2 emission (kg) HCI emission (kg) H2O emission (kg) SO2 flux (t/d) HCI flux (t/d) H2O flux (t/d)

Jorullo

Paricutin

3.4 × 1012 0.19 0.13 5.75 99 67 99 1.3 × 1010 3.0 × 109 1.9 × 1011 2320 540 35090

3.8 × 1012 0.19 0.10 4.93 99 53 99 1.4 × 1010 2.1 × 109 1.9 × 1011 4360 630 56630

Fluxes were calculated based on the recorded lengths of the eruptions: 15 years (Jorullo), 9 years (Paricutin). a See text for calculation of the erupted mass.

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