Journal of Petrology Advance Access originally published online on July 13, 2005
Journal of Petrology 2005 46(11):2337-2366; doi:10.1093/petrology/egi058
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Volcán Popocatépetl, Mexico. Petrology, Magma Mixing, and Immediate Sources of Volatiles for the 1994–Present Eruption
1 DEPARTMENT OF EARTH AND SPACE SCIENCES, UNIVERSITY OF WASHINGTON, SEATTLE, WA 98195, USA
2 USGS-DEPARTMENT OF EARTH AND SPACE SCIENCES, UNIVERSITY OF WASHINGTON, SEATTLE, WA 98195, USA
RECEIVED APRIL 17, 2005; ACCEPTED MAY 23, 2005
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONANALYTICAL METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAAPPENDIXREFERENCESREFERENCES |
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Volcán Popocatépetl has been the site of voluminous degassing accompanied by minor eruptive activity from late 1994 until the time of writing (August 2002). This contribution presents petrological investigations of magma erupted in 1997 and 1998, including major-element and volatile (S, Cl, F, and H2O) data from glass inclusions and matrix glasses. Magma erupted from Popocatépetl is a mixture of dacite (65 wt % SiO2, two-pyroxenes + plagioclase + Fe–Ti oxides + apatite,
3 wt % H2O, P = 1·5 kbar, fO2 = 
NNO + 0·5 log units) and basaltic andesite (53 wt % SiO2, olivine + two-pyroxenes, 3 wt % H2O, P = 1–4 kbar). Magma mixed at 4–6 km depth in proportions between 45:55 and 85:15 wt % silicic:mafic magma. The pre-eruptive volatile content of the basaltic andesite is 1980 ppm S, 1060 ppm Cl, 950 ppm F, and 3·3 wt % H2O. The pre-eruptive volatile content of the dacite is 130 ± 50 ppm S, 880 ± 70 ppm Cl, 570 ± 100 ppm F, and 2·9 ± 0·2 wt % H2O. Degassing from 0·031 km3 of erupted magma accounts for only 0·7 wt % of the observed SO2 emission. Circulation of magma in the volcanic conduit in the presence of a modest bubble phase is a possible mechanism to explain the high rates of degassing and limited magma production at Popocatépetl. KEY WORDS: glass inclusions; igneous petrology; Mexico; Popocatépetl; volatiles
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONANALYTICAL METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAAPPENDIXREFERENCESREFERENCES |
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From December 1994 until the time of writing (August 2002), Volcán Popocatépetl (Mexico) has non-explosively degassed an average of 7000 tons/day SO2 into the troposphere [Delgado & Cárdenas-G, 1997
; Delgado-Granados et al., 2001; National Center for Disaster Prevention, Mexico City (CENAPRED), unpublished data, 2000]. The current eruptive period is similar to other eruptions that have occurred in historical time. This is the first time, however, that such massive degassing could be measured quantitatively. Popocatépetl has degassed a total of 20 Mt SO2 since 1994, accompanied by relatively minor eruptive activity, compared with the 17 Mt SO2 (Gerlach et al., 1996) injected into the stratosphere in a single day during the climactic, highly explosive eruption of Pinatubo (Philippines) on June 15, 1991. In this study, we examine various aspects of the process of passive degassing at Popocatépetl. We describe the petrology of recent magmas, and assess the volatile budget during recent volcanic activity. These data are used to estimate magmatic conditions and volatile sources immediately prior to eruption. Based on these constraints we discuss possible processes to liberate the large mass of gases non-explosively emitted by Popocatépetl during recent activity.
Background
Popocatépetl (5452 m) is an andesitic, subduction zone stratovolcano that lies 60 km SE of Mexico City (Fig. 1). Holocene eruptive activity at Popocatépetl included at least three Plinian eruptions (Siebe et al., 1996a, 1996b; Panfil et al., 1999). Since AD 1354, Popocatépetl has experienced at least 12 periods of significant unrest, all of which exhibited volcanic activity similar to the 1994–present eruptive period (De la Cruz-Reyna et al., 1995).
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, Holocene volcanoes; 
, Mexico City. Map made with GMT (Wessel & Smith, 1991After nearly 70 years of dormancy, Volcán Popocatépetl began a new period of activity in December 1994 (De la Cruz-Reyna & Siebe, 1997
). Elevated levels of degassing, seismicity, and phreatic explosions characterized the initial portion of volcanic unrest at Popocatépetl through early 1996. In March 1996, a dacitic lava dome extruded into the summit crater (GVN, 1996; S. De la Cruz-Reyna et al., unpublished data, 1998; Stimac et al., 1997). Since the onset of the current eruptive period, Popocatépetl has continuously degassed at high rates accompanied by relatively minor eruptions. Eruptive activity includes: (1) vulcanian explosions that eject lithic and juvenile ash and rock debris up to 8 km above the vent; (2) ash-poor ‘exhalations’ of volcanic gas; (3) periodic lava dome growth, subsidence, and explosive destruction (S. De la Cruz-Reyna et al., unpublished data, 1998). Eruptions at Popocatépetl are commonly characterized by mixing and mingling of distinct mafic and silicic magmas (Cantagrel et al., 1984; Robin, 1984; Boudal & Robin, 1988; Kolisnik, 1990). This includes the most recent active phase extending from 1994 to the present (Athanasopoulos, 1997; Straub & Martin-Del Pozzo, 2001).
From March 1996 until December 2000, 3·1 x 107 m3 of magma was extruded, based on aerial photogrammetry (S. De la Cruz-Reyna et al., unpublished data, 2000). During this period the average SO2 flux was 8400 tons/day (CENAPRED, unpublished data, 2000), yielding a total of 14·8 ± 3·0 Mt SO2 (average ±20% error because of uncertainties in measuring windspeed; Delgado-Granados et al., 2001). Assuming 20 vol. % crystals with melt density of 2300 kg/m3, about 5·7 x 1010 kg silicate melt has been erupted during the March 1996 to December 2000 period of activity. For this mass of silicate melt alone to supply the 14·8 Mt observed SO2 emission, the silicate melt phase would have to have initially contained 13 wt % S. This exceeds the experimentally derived sulfur solubility limits in silicic and intermediate magmas by two or more orders of magnitude (Wendlandt, 1982; Carroll & Rutherford, 1985; Luhr, 1990; Scaillet et al., 1998). This apparent paradox has been encountered by other workers and is commonly referred to as the ‘excess sulfur’ problem (Luhr et al., 1984; Sigurdsson et al., 1990; Andres et al., 1991; Westrich & Gerlach, 1992; Gerlach & McGee, 1994; Gerlach et al., 1994, 1996; Kazahaya et al., 1994, 2002; Wallace & Gerlach, 1994; Mandeville et al., 1996; Witter, 1997; Wallace, 2001).
Previous work
During the 1980s various workers pointed out the importance of magma mixing as a petrogenetic process at Popocatépetl and documented giant sector collapse events at the volcano (Cantagrel et al., 1984; Robin, 1984; Robin & Boudal, 1987; Boudal & Robin, 1988, 1989; Kolisnik, 1990). Renewal of eruptive activity at Popocatépetl in 1994 stimulated extensive new research on the eruptive history, petrology, and degassing behavior of the volcano (CENAPRED-UNAM, 1995; Athanasopoulos, 1997; Straub & Martin-Del Pozzo, 2001).
Gas measurements at Popocatépetl have focused on SO2, using correlation spectrometry (COSPEC). Carbon dioxide emissions have also been measured using LI-COR and passive infrared spectroscopy (FTIR). COSPEC measurements began in January 1994 and showed emission rates of 1500 to 12 500 tons/day SO2 (Delgado & Cárdenas-G, 1997; Delgado-Granados et al., 2001). On rare occasions, SO2 emissions have reached 30 000–50 000 tons/day SO2 (Goff et al., 1998). LI-COR measurements of CO2 in June 1995 indicated a flux of between 6400 and 9000 tons/day CO2 (Gerlach et al., 1997). Using the same technique, Delgado et al. (1998) reported average fluxes of >40 000 tons/day CO2 between 1996 and 1998. In February 1998, passive FTIR measured 38 000 tons/day CO2 with excursions that exceeded 100 000 tons/day CO2 (Goff et al., 2001). These measurements establish Popocatépetl as one of the greatest known natural producers of SO2 and CO2.
Relative HCl and HF emission rates were measured using volatile traps from February 1994 until the onset of activity in December 1994 (Goff et al., 1998). Some volatile trap samples were also collected in early 1996. Volatile trap studies, combined with SO2 flux measurements, suggest that between the onset of activity and November 1996, Popocatépetl had emitted 0·75 Mt HCl and 0·075 Mt HF (Goff et al., 1998). This represents an average emission rate of 740 tons/day HCl and 74 tons/day HF, respectively. Emission rates measured by passive FTIR in 1997 and 1998 have the range 190–1700 tons/day HCl and 35–250 tons/day HF (Goff et al., 2001). The uncertainty in these halogen measurements is estimated to be of the order of ±20% (Goff et al., 1998, 2001).
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ANALYTICAL METHODS |
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TOPABSTRACTINTRODUCTIONANALYTICAL METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAAPPENDIXREFERENCESREFERENCES |
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All analyses by electron microprobe (EPMA) were conducted on a four-spectrometer JEOL 733 Superprobe at the University of Washington. Mineral analyses were conducted at 15 kV accelerating voltage,
1–3 µm beam diameter, and 10–25 nA beam current with a maximum count time of 40 s for each element. Elements were calibrated using mineral standards, which were also analyzed at the end of each analytical session to check for instrumental drift. Matrix effects were corrected using the ZAF model of Armstrong (1984).
Major element glass analyses were performed at 15 kV accelerating voltage and 10 nA with a 5 µm beam diameter and maximum count time of 40 s for each element. The combination of analyzing sodium first and a defocused beam minimized Na loss during analysis. Plots of intensity vs time suggest Na loss of <0·2 wt % Na2O over a 40 s count time using an anhydrous glass standard. Na loss during analysis of glass inclusions, however, may be significant because of enhanced Na mobility in hydrous glasses (Devine et al., 1995). No correction was applied for potential alkali loss in hydrous inclusions. Mean standard errors as well as errors based on counting statistics for the major element mineral and glass analyses are shown in the Electronic Appendix (Tables A1 and A2), which may be downloaded from http://www.petrology.oupjournals.org/.
Volatile elements S, Cl, and F were analyzed in glasses, hornblende, and apatite using an improved analytical method originally developed by Thordarson (Thordarson et al., 1996). Analytical conditions were 15 kV accelerating voltage, 100 nA beam current, 5 µm beam diameter, and 400 s total peak count time. To minimize volatile loss, an iterative scheme was employed in which the beam was blanked for 10 s after every 10 s of counting. This procedure minimized heating of the sample and consequent volatile loss. We conducted a test for volatile mobility under the electron beam and found that the X-ray count rates for S, Cl, and F remained stable during the entire period of analysis. Sulfur and chlorine were analyzed using PET diffraction crystals with the spectrometer aperture window set at its widest setting to maximize the number of X-ray counts. Fluorine was analyzed using an LDE1 diffraction crystal and a new empirical correction method that overcomes peak overlap problems with the Fe K
line (Witter & Kuehner, 2004). Because the volatile elements were analyzed in a separate session, we reprocessed the volatile data with major element data (collected previously; see Witter & Kuehner, 2004) to correct for matrix effects using the ZAF model of Armstrong (1984). Prior to each analytical session, volatile elements were calibrated with NiS (sulfide sulfur), celestite (sulfate sulfur), scapolite 17 120 (chlorine), and synthetic F-phlogopite (fluorine). Analytical precision and accuracy were checked using basalt glass standards VG-2 and A-99 (for S, Cl, and F) and NBS soda-lime glasses 620 and 621 (for S and Cl). The 1
S, Cl and F variations for more than 30 analyses are 1·5% at 1500 ppm and 13% at 140 ppm for sulfur, 9% at 200 ppm for chlorine, and 5% at 750 ppm and 12% at 300 ppm for fluorine (Electronic Appendix, Table A3).
H2O, CO2, S, Cl, and F in glass inclusions were analyzed using secondary ion mass spectrometry (SIMS) at the Department of Terrestrial Magnetism, Carnegie Institution of Washington. The analytical methods for this technique have been described by Hauri et al. (2002). SIMS analysis could only be performed on inclusions that were larger than about 40 µm in the smallest dimension.
Major and trace element compositions of four pumice samples and dome rocks from the current eruptive period were analyzed by X-ray fluorescence at the WSU Geoanalytical Laboratory or by fusion inductively coupled plasma mass spectrometry (ICP-MS) at Activation Laboratories, Ontario, Canada. To augment our dataset of whole-rock analyses we flash-melted selected portions of eight more pumice fragments and dome rocks from the same sample suite. The resulting glass beads were analyzed for major elements by electron microprobe. This method enabled us to obtain a bulk chemical analysis of small, uniform portions of compositionally banded pumice and dome rock. All samples were ground using a mortar and pestle, and then melted in a platinum crucible at 1350°C for 10 min. The resulting glass was reground using a mortar and pestle, remelted at 1350°C for a further 10 min, and then quenched in water.
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RESULTS |
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TOPABSTRACTINTRODUCTIONANALYTICAL METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAAPPENDIXREFERENCESREFERENCES |
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Petrology
Rock samples used in this study were obtained from small explosive events that occurred in June and December 1997 and January, March, and April 1998. All samples represent juvenile magma and were initially studied under the petrographic microscope. We focused our analytical efforts on pumice lapilli from the eruption of June 30, 1997 and dense blocks of dome material from the eruption of January 1, 1998. The mineralogy of the dome rocks from explosive events in December 1997 and January, March, and April 1998 is identical to the mineralogy of the pumiceous samples of the June 30, 1997 eruption. The rapid quenching of lapilli from the June 30, 1997 eruption made these samples more suitable for analysis by electron microprobe.
Whole-rock compositions of the January 1998 dome samples (Table 1, Fig. 2) are silicic andesite to dacite—similar to samples from the March 1996 dome (Athanasopoulos, 1997). Brown and white pumice, erupted in June 1997, are compositionally distinct. Brown pumice is andesitic, whereas white pumice is intermediate in composition between brown pumice and the 1996 and 1998 dome rocks. Whole-rock data suggest no systematic change in major or trace element composition with time.
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, white pumice from the June 30, 1997 eruption; 
, gray dome rock from the January 1, 1998 eruption; +, whole-rock analyses of the March 1996 dome (Athanasopoulos, 1997Recent Popocatépetl eruptive products contain phenocrysts and microphenocrysts of olivine, clinopyroxene, orthopyroxene, plagioclase and hornblende. Microphenocrysts of titanomagnetite, ilmenite, and rare apatite are distributed in a microcrystalline groundmass. Chromian spinel commonly occurs as inclusions in olivine. Titanomagnetite and ilmenite occur as inclusions in pyroxene and rarely in plagioclase. Apatite and rare globular sulfide are also found as inclusions, most commonly in pyroxene. Anhedral xenocrysts of quartz are rare. The groundmass varies from light to dark in color. Dark groundmass consists of a dense network of plagioclase and Fe–Ti oxide microlites in brown glass. Light-colored groundmass contains sparse plagioclase microlites in colorless glass. The relative abundance of mineral phases in Popocatépetl's recent eruptives is: plagioclase > orthopyroxene
clinopyroxene > olivine > ilmenite titanomagnetite hornblende > chromian spinel apatite > sulfide. Modal analyses are shown in Table 2.
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Olivine (0·1–1·5 mm) comes in two textures. Larger-sized olivine grains (>0·5 mm) are the more abundant, are generally subhedral, and locally glomerocrystic. They contain abundant chromian spinel inclusions (
15 mm diameter) but no glass inclusions and consist of broad, compositionally uniform cores (Fo87–91) with narrow (20–70 µm) normally zoned rims (Fo73–88). NiO content is high (2300–6700 ppm) and varies sympathetically with MgO content in the cores of these olivine grains (Fig. 3 and Electronic Appendix, Table A4). Some of these glomerocrysts as well as some individual grains of olivine are anhedral and display embayed margins that have reacted extensively with the surrounding silicate liquid. We suggest that these large, Cr-spinel-rich olivines are probably xenocrystic.
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The other type of olivine, microphenocrystic (<0·3 mm), is much less abundant. Some grains have a skeletal (arrow-shaped) habit suggesting rapid growth from the surrounding silicate melt. Olivine microphenocrysts contain few, if any, chromian spinel inclusions but do contain 5–40 µm diameter inclusions of clear brown glass. Glass inclusion-bearing olivine microphenocrysts have uniform cores but have a lower Fo content and more limited compositional range (Fo84–86) than the large-sized Cr-spinel-rich olivines. Glass inclusion-bearing olivine microphenocrysts, however, have much lower NiO contents (1100–1800 ppm) for Fo84–86 (Fig. 3).
Orthopyroxene and clinopyroxene are common as phenocrysts in all samples and reach 1 mm in length. Both clinopyroxene and orthopyroxene phenocrysts contain abundant glass inclusions 5–50 µm in diameter. Rare, globular, sulfide inclusions 5–15 µm in size and lath-shaped apatite 10–50 µm long also occur as inclusions in pyroxene. Microphenocrysts of clinopyroxene and orthopyroxene (30–300 µm) are common, frequently have octagonal or lath-shaped cross-section, and rarely contain melt or mineral inclusions. Most pyroxene core compositions occur as high-Mg-number and low-Mg-number phases [Mg-number = Mg/(Mg + Fe2+); Fig. 4]. Plots of SiO2, Al2O3, and Cr2O3 vs Mg-number help delineate the two groups of clino- and orthopyroxene (Electronic Appendix, Fig. A1).
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Pyroxenes are characterized by broad, compositionally uniform cores with or without narrow, zoned rims (Electronic Appendix, Table A7). High- and low-Mg-number clinopyroxene cores are augite with compositions En47–49Fs8–10Wo42–44 (Mg-number 83–86 and Cr2O3 0·2–0·7 wt %) and En40–44Fs13–18Wo40–45 (Mg-number 70–78 and Cr2O3 < 0·3 wt %), respectively. High-Mg-number orthopyroxene cores are bronzite with composition En82–83Fs14–15Wo2–4 (Mg-number 84–85 and Cr2O3 0·1–0·4 wt %). Low-Mg-number orthopyroxene cores are hypersthene with composition En61–69Fs29–35Wo2–3 (Mg-number 63–72 and Cr2O3 < 0·15 wt %). Core-to-rim zonation in pyroxene is highly variable. High-Mg-number pyroxenes are generally normally zoned (up to
En = 21 mol %) whereas low-Mg-number pyroxenes are predominantly reversely zoned (up to En = 15 mol %). Many pyroxene grains, however, are unzoned or have weak normal to reverse zoning (En <3 mol %).
Hornblende phenocrysts commonly have an acicular habit and are up to 1 mm in length. More equant hornblende microphenocrysts are generally 0·2 mm x 0·3 mm in size. All hornblendes have dark ‘gabbroic-type’ (Garcia & Jacobson, 1979) breakdown rims 30–100 µm across made of tiny plagioclase, pyroxene, Fe–Ti oxides, and rare sulfides. Rare hornblende phenocrysts and microphenocrysts possess small plagioclase or pyroxene cores. Some hornblendes form partial overgrowths on olivine and pyroxene.
Popocatépetl hornblende has a restricted compositional range (Electronic Appendix, Table A9) with Mg-number = 63–72 [Mg-number = Mg/(Mg + FeT)], 9–13 wt % Al2O3, low concentrations of Cl (200 ppm) and high F (2100–5400 ppm). This contrasts with some subduction zone volcanoes, such as Unzen and Montserrat, that erupt magmas containing two compositionally distinct amphiboles. Popocatépetl hornblende compositions (Electronic Appendix, Fig. A2) are similar to groundmass amphibole (magnesiohastingsite) in mafic inclusions in Montserrat andesite (Murphy et al., 2000) and groundmass amphibole (pargasite) from Unzen volcano (Sato et al., 1999). They plot well outside of the field for amphibole phenocrysts (magnesiohornblende) in Montserrat andesite and Unzen dacite.
Complete descriptions of all samples and analyses of all mineral phases are given in the Electronic Appendix, which can be downloaded from http://www.petrology.oupjournals.org/.
Estimation of intensive variables: T, P, and fO2
Two-pyroxene geothermometry
Magmatic temperatures were calculated using orthopyroxene–clinopyroxene pair compositions and the QUILF program (Andersen et al., 1993). Results are listed in the Electronic Appendix (Table A12). Most pyroxene pairs used in the geothermometry calculation were in direct contact; however, a few pyroxene pairs were separated by up to a few millimeters of groundmass.
Low-Mg-number pyroxene core pairs yield temperatures of 959 ± 17°C (n = 12). High-Mg-number pyroxene core pairs are out of equilibrium and do not yield reliable temperature estimates. Temperatures derived from pyroxene rims occur in low- and high-temperature groups. The low-temperature rims yield 949 ± 32°C (n = 11) and involve unzoned to weakly zoned low-Mg-number pyroxenes or high-Mg-number pyroxenes with strong normal zonation. The temperatures derived from the pyroxene cores and low-Mg-number pyroxene rims are in good agreement. The high-temperature pyroxene rims yield 1055 ± 9°C (n = 3) and involve either low-Mg-number pyroxene with reversely zoned rims or high-Mg-number pyroxenes with weak normally zoned rims. Straub & Martin-Del Pozzo (2001) found one pair of low-Mg-number pyroxenes and determined a temperature of 930–950°C using the geothermometer of Lindsley & Andersen (1983).
Hornblende–plagioclase geothermometry
The composition of Popocatépetl hornblende allows use of the hornblende–plagioclase (Hbl–Plag) geothermometer of Holland & Blundy (1994). However, accurate application of the Hbl–Plag geothermometer is difficult, as the composition of plagioclase in equilibrium with hornblende at the time of crystallization is unknown. Taking a conservative approach, we assume that at hornblende saturation, the coexisting plagioclase had a composition that lies in the range of the measured plagioclase compositions (An28–61). Utilizing the edenite–tremolite geothermometer (for mineral assemblages lacking quartz) over a pressure range of 0–5 kbar, the entire measured plagioclase and hornblende compositional range yields temperatures from 775 to 900°C.
Apatite saturation temperature
Apatite saturation temperatures were calculated using the model equation from Piccoli & Candela (1994):
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is the P2O5 concentration in the initial melt (whole-rock P2O5 content), 
is the SiO2 concentration in the melt at apatite saturation (whole-rock SiO2 content), and X is the crystallinity (in wt %). Using whole-rock compositions of the most silicic clasts, and 35 vol. % (40 wt %) crystallinity (from modal analysis of June 30, 1997 white pumice samples), these calculations suggest apatite saturation at between 945°C and 959°C (n = 4). This is in good agreement with temperatures derived from cores of low-Mg-number pyroxene.
Fe–Ti oxide geothermometry and oxygen barometry
The thermodynamic formulation of Ghiorso & Sack (1991) and QUILF (Andersen et al., 1993) were both used to estimate temperature and fO2 from Fe–Ti oxide pairs (Electronic Appendix, Table A13). Oxide pairs that did not pass the Mg/Mn partition test of Bacon & Hirschmann (1988) were not used. Most of the oxide pairs used in the T–fO2 calculation were in contact; pairs not in contact were less than 100 µm apart.
Fe–Ti oxide compositions cover a wide range (XIlm = 0·63–0·79 and XUsp = 0·30–0·47). Similarly, temperature estimates from both geothermometers cover a broad range [910–1030°C using QUILF and 1040–1190°C using Ghiorso & Sack (1991); n = 27 pairs]. Oxygen fugacity results are comparable using the two methods and have the range NNO + 0·5 to NNO + 1·5 (where NNO is the nickel–nickel-oxide buffer). Athanasopoulos (1997) analyzed several Fe–Ti oxide pairs from the March 1996 lava dome that have a very similar compositional range. Variable degrees of magma mixing and/or incomplete thermal re-equilibration are possible explanations for the wide variation in T and fO2 obtained using each method.
The differing results from the two methods, however, must be reconciled. We suggest that the values presented here using the thermodynamic formulation of Ghiorso & Sack (1991) may significantly overestimate the temperature of Popocatépetl magma. Evans et al. (2001) found that for hydrous, silicic and intermediate volcanic rocks at high oxygen fugacity (>NNO + 1) and 700–900°C, the model of Ghiorso & Sack (1991) overestimates T and fO2. This error may be associated with inadequate characterization of the 
transition in the rhombohedral phase (B. W. Evans & M. S. Ghiorso, personal communication, 2001). New experimental results for the Fe–Ti–O system at 1 bar and 1000–1300°C are consistent with this hypothesis (Lattard et al., 2001). The new data suggest that Popocatépetl Fe–Ti oxides equilibrated at temperatures 
1000°C at an oxygen fugacity between NNO and NNO + 1. This temperature and fugacity estimate is consistent with QUILF calculations. We conclude from this that Fe–Ti oxides record temperatures that probably span the range 900–1000°C and fO2 = NNO to NNO + 1. This range in Fe–Ti oxide results suggests a cooling array that defines a line in T– fO2 space. Where this line intersects the temperature derived from pyroxene thermometry (950°C), oxygen fugacity has the value NNO + 0·5.
Glass inclusions
Glass inclusions occur in all mineral phases. We focused analytical efforts on glass inclusions in olivine and pyroxene with the aim of obtaining chemical data on the volatile contents of the pre-eruptive magmatic end-members. Only glass inclusions that were 
10 µm in diameter and without intersecting cracks were considered. Despite taking care in choosing glass inclusions, some analyses displayed elevated Ca and Al contents. These anomalous analyses were interpreted to result from microscopic post-entrapment plagioclase crystallization creating inhomogeneity of the quench products contained in the glass inclusion. Such analyses were excluded from further consideration. The concentration of H2O in glass inclusions was estimated using both SIMS and the difference technique (100 wt % – analytical total of major and volatile elements by EPMA; Anderson, 1974).
Forty-two glass inclusions in olivine were analyzed for major elements (Fig. 5; Electronic Appendix, Table A14). Twenty of these glass inclusions in olivine were analyzed for S, Cl, and F (Fig. 6; Electronic Appendix, Table A16). Sixteen of these were recalculated for post-entrapment crystallization (PEC) of olivine using the thermodynamic model of Kress & Ghiorso (2004). This correction calculation is performed iteratively by incrementally adding back into the liquid (corresponding to the glass composition) instantaneous equilibrium solid in 0·01 wt % increments. Initial glass composition is assumed to have been reached when the inclusion composition most closely approaches equilibrium with the bulk host phenocryst. The method assumes that the inclusions were sealed throughout the crystallization process and that crystallization was rapid relative to diffusive processes in the crystal. S, Cl, and F concentrations are adjusted by assuming that they are totally incompatible. The magnitude of PEC in olivine-hosted inclusions varies between 1 and 10 wt %. Thus, recalculated volatile contents can be up to 10% lower than the measured values. PEC calculations for these samples using PETROLOG v2.1.3 (Danyushevsky, 2001) yielded similar results.
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Twenty-one glass inclusions in orthopyroxene and 30 glass inclusions in clinopyroxene were analyzed for major elements (Fig. 5; Electronic Appendix, Table A14). Of these, 18 glass inclusions in orthopyroxene and 23 glass inclusions in clinopyroxene were analyzed for S, Cl, and F (Fig. 6; Electronic Appendix, Tables A17 and A18). PEC correction calculations were performed on 13 of the orthopyroxene-hosted glass inclusions using the thermodynamic model of Kress & Ghiorso (2004)
. Calculated PEC in orthopyroxene-hosted glass inclusions ranges from 0 to 5 wt %. PEC calculations using PETROLOG v2.1.3 (Danyushevsky, 2001) suggest 1–5 wt % post-entrapment crystallization for the clinopyroxene-hosted glass inclusions.
According to the classification scheme of Le Bas et al. (1986), PEC-corrected glass inclusions in olivine are mostly basaltic andesite. Glass inclusions in clinopyroxene span the range basaltic andesite to rhyolite. Glass inclusions in orthopyroxene are dacite to rhyolite (Fig. 5). With increasing silica, Al2O3, TiO2, MgO, FeO, CaO, and P2O5 decrease and K2O increases. There are no obvious trends for MnO, Cr2O3, and Na2O.
Glass inclusions were chosen for analysis by observation of two-dimensional thin sections. Unseen cracks or re-entrants intersecting a glass inclusion in the third dimension may have escaped detection. These potentially allowed for volatile leakage from the glass inclusion during ascent and eruption. Additionally, diffusive loss of H2O or OH may have influenced the water content preserved in glass inclusions. In an attempt to filter out a possible volatile leakage and/or diffusive bias in our data we used the composition of glass inclusions with the highest sulfur and water concentrations to represent pre-eruptive volatile contents (Fig. 7). Glass inclusions with the highest sulfur and water contents generally have the highest Cl contents.
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Volatile-rich glass inclusions hosted by olivine are basaltic andesite in composition. The olivine-hosted glass inclusion with the highest volatile content (GI1OL6tsB) contains 1980 ppm S, 1060 ppm Cl, 950 ppm F, and 3·3 wt % H2O. The five glass inclusions in orthopyroxene with the highest volatile contents (MI1opx8tsSCa, MI2opx8tsSCa, MI3opx8tsSCa, MI4opx8tsSCa, MI2opx1tsPumc) are rhyolitic and average 130 ± 50 ppm S, 880 ± 70 ppm Cl, 570 ± 100 ppm F, and 2·9 ± 0·2 wt % H2O (±1
). The five glass inclusions in clinopyroxene with the highest volatile contents (MI1cpx25ts1, MI2cpx25ts1, MI1cpx12ts2, MI3cpx12ts2, MI1cpx14tsPumc) are also rhyolitic and average 80 ± 100 ppm S, 760 ± 110 ppm Cl, 830 ± 270 ppm F, and 3·2 ± 0·6 wt % H2O (±1). The volatile contents of intermediate composition (andesite to dacite) glass inclusions hosted by olivine, orthopyroxene, and clinopyroxene are highly variable.
Matrix glasses
Fifty-seven major element analyses and 13 analyses of S, Cl, and F in matrix glass were performed by EPMA (Figs 5 and 6; Electronic Appendix, Tables A15 and A20). The concentration of H2O in the matrix was calculated using the difference technique (Anderson, 1974). Each analysis represents a single point in the matrix glass. Care was taken to avoid the solid phases in microlite-rich glass by inspecting the region of interest with high-magnification backscattered electron (BSE) imaging. Analyses with unusually high Ca, Al, or Fe contents were presumed to be due to subsurface plagioclase and/or magnetite microlites and were excluded. Our major and volatile element analyses are comparable with those of light and dark matrix glasses found in the March 1996 dome rocks of Popocatépetl reported by Athanasopoulos (1997).
Matrix glasses have compositions ranging from andesite to rhyolite (Fig. 5; Electronic Appendix, Tables A15 and A20) according to the classification scheme of Le Bas et al. (1986). Brown, microlite-rich matrix glasses are more mafic than colorless, microlite-poor matrix glasses. The matrix glass compositional variations mimic that of the glass inclusions: with increasing silica, Al2O3, TiO2, MgO, FeO, CaO, and P2O5 decrease and K2O increases. Within error, no trend exists for MnO, Cr2O3, and Na2O. Matrix glasses from the brown andesite and white dacite pumices of the June 30, 1997 eruption contain 60–68 wt % SiO2 and 70–72 wt % SiO2, respectively. Matrix glass in the dacite dome rock erupted January 1, 1998 is the most felsic with 76–77 wt % SiO2.
The averages of all matrix glass volatile analyses are 30 ± 30 ppm S, 600 ± 110 ppm Cl, 720 ± 200 ppm F, and 1·3 ± 1·0 wt % H2O by difference (Fig. 7). Our data suggest a relatively high retention of H2O in the matrix glasses. Typically, degassed unaltered glasses at the surface will contain <0·3 wt % H2O (Newman et al., 1988). The estimated minimum detection limit using the H2O by difference technique is 1 wt % H2O. We believe the anomalously high H2O contents reported here for matrix glass reflect cumulative errors in the ‘H2O by difference’ technique. There appears to be no systematic variation in volatile content of the matrix glass, regardless of the matrix glass color, microlite content, or silica content.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONANALYTICAL METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAAPPENDIXREFERENCESREFERENCES |
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Evidence for magma mixing
Evidence for magma mixing is abundant in the pumice and dome rocks from Popocatépetl on the macroscopic, petrographic, and chemical level. Hand samples of pumice and dome rocks contain bands and lenses of light and dark material suggesting the mechanical mingling of magmas of differing composition. Disequilibrium textures and mineralogy include olivine (Fo90) in a dacitic to rhyolitic matrix glass (some with orthopyroxene reaction rims), sieve-textured and resorbed plagioclase, reverse zonation in the rims of low-Mg-number pyroxene, and normal zonation in the rims of olivine and high-Mg-number pyroxene. Our observations are consistent with other studies of Popocatépetl that suggest magma mixing has been a common process during the 40 kyr history of the modern Popocatépetl cone (Cantagrel et al., 1984
; Robin, 1984; Boudal & Robin, 1988; Kolisnik, 1990; Athanasopoulos, 1997).
Definition of the end-member magmas: MELTS modeling
It is likely that one or both of the presumed magmatic end-members may themselves be the product of mixing processes. Nevertheless, variation in the most recent eruption products can be well described by mixing of two distinct and relatively homogeneous compositions immediately prior to eruption, and this final mixing event is likely to be the most relevant in terms of unraveling details of the eruption process. We used the MELTS algorithm (Ghiorso & Sack, 1995) to help define the compositions and intensive parameters of the mixing end-members in the following manner.
We first assumed that end-member compositions lay along the linearly regressed mixing line (Fig. 2). Compositions along extrapolations of this mixing line were explored to find potential end-member compositions that could have been in equilibrium with phenocryst phases observed in the rock samples. The mineral compositions determined by MELTS must be consistent with phase compositions determined by electron microprobe. Reasonable pressure conditions were iteratively tested using MELTS. The magmatic conditions were constrained with temperatures, H2O contents, and fO2 determined from the methods described previously. This approach allows simultaneous application of a large number of thermodynamic constraints on potential conditions and compositions. In practice, this places tight constraints on permissible end-member compositions. Estimated bulk compositions and magmatic conditions of the silicic and mafic end-members are shown in Table 3.
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It is generally recognized that MELTS is better calibrated and more accurately represents phase equilibria in mafic compositions. Nevertheless, there is currently no tool that allows simultaneous application of as many phase equilibria constraints in compositions relevant to nature.
The MELTS-modeled silicic end-member is a dacite (65 wt % SiO2) that crystallized predominantly at low pressures (<1·5 kbar or <6 km) with 0·5–1·5 wt % H2O and fO2 = NNO + 0·3 log units. A crystallization interval of 1000–900°C produces a phase assemblage of plagioclase + clinopyroxene + orthopyroxene + titanomagnetite ± ilmenite. MELTS predicts 30–40 wt % feldspar, 6–8 wt % pyroxene, and 0·5–1 wt % titanomagnetite (at 0·5–1·5 wt % H2O, 0·1–1 kbar, 950°C, and fO2 = NNO + 0·3 log units), which is consistent with modal observations (Table 2). The MELTS-predicted silicate liquid compositions at these conditions also match the microprobe analyses for matrix glasses in the samples of June 30, 1997 white dacite pumice (Fig. 5).
The MELTS-modeled mafic end-member is a basaltic andesite (53 wt % SiO2) with 3 wt % H2O and fO2 = NNO + 1 log units (Table 3). The pressure–temperature evolution of the mafic end-member consists of olivine fractionation at moderate pressures (2–4 kbar, 1150–1300°C) followed by low-pressure (1–2 kbar, 1050–1150°C) pyroxene fractionation.
MELTS modeling and experimental determinations of the olivine/silicate melt distribution coefficient for Ni (Seifert et al., 1988; Snyder & Carmichael, 1992; Gaetani & Grove, 1997) support a xenocrystic origin of the large, Fo87–91, Ni-rich and Cr-spinel-bearing olivine (Fig. 3). Using experimental values for the Ni distribution coefficient, the high Ni content in these olivines required large degrees of fractionation to produce the observed Ni contents in the mafic liquid. Such high degrees of fractionation cannot be reconciled with the lower degrees of Fe/Mg fractionation implied in the same mafic liquid. The low-Ni, Fo84–86 olivines that contain glass inclusions probably did crystallize from the mafic end-member. We speculate that the xenocrystic Fo87–91, Ni-rich olivines and their included Cr-spinel were plucked from the mantle and/or crystallized from a primitive Mg-, Ni-, and Cr-rich mantle melt (Clynne & Borg, 1997) which were subsequently incorporated into Popocatépetl's mafic end-member magma.
Hornblende: a product of mixing
Hornblende phenocrysts and microphenocrysts are found evenly disseminated, surrounded by both the dark and light matrix glasses, in all pumices and lava blocks. Petrographic observations are inconsistent with hornblende originating from only the silicic or mafic end-members because hornblende is found as partial overgrowths on both low-Mg-number pyroxene and olivine. If Popocatépetl hornblende is derived neither from the mafic nor the silicic end-member, it must have crystallized during mixing. Plagioclase and pyroxene microphenocrysts form cores of some hornblende, suggesting that hornblende saturated very late in the crystallization sequence. In addition, the acicular habit of hornblende suggests that crystallization was rapid. Compositionally, Popocatépetl hornblende is very similar to groundmass pargasite in Unzen dacite (Sato et al., 1999) and groundmass amphibole (magnesiohastingsite) found in mafic inclusions in Montserrat andesite (Murphy et al., 2000; Electronic Appendix, Fig. A2).
The concentration of Cl in Popocatépetl hornblende is consistent with a low-pressure origin. Sato et al. (1999) found that magnesiohornblende phenocrysts in Unzen lavas contained much higher concentrations of Cl than did the groundmass pargasite. They concluded that the low Cl content of the Unzen groundmass pargasite (100–300 ppm Cl) is a result of crystallization from a partially degassed magma. Popocatépetl hornblende also has low Cl contents (<300 ppm Cl). We infer that Popocatépetl hornblende also crystallized from a partially degassed magma and that partial degassing may have served to enhance hornblende saturation (Ghiorso, 1999).
Proportions and depth of magma mixing
With the compositions of the mixing end-members estimated, we can calculate the proportions of mixing needed to derive the compositions of individual pumice or lava clasts (Table 3). Compositions of the most mafic samples (58 wt % SiO2), represented by brown andesite pumice erupted on June 30, 1997, are consistent with mixing of 55 wt % mafic end-member and 45 wt % silicic end-member. The most silicic clasts (63 wt % SiO2) are represented by white dacite pumice erupted on June 30, 1997 and gray dacite dome rocks erupted on January 1, 1998. These are inferred to be made up of 85 wt % silicic end-member and 15 wt % mafic end-member.
Based on the minimum pressure estimates for the mafic end-member (1 kbar) and the maximum pressure estimate for the silicic end-member (1·5 kbar) derived from the MELTS simulations presented above, we infer that magma mixing occurred at pressures between 1 and 1·5 kbar (4–6 km depth). This result is toward the shallow end of the 4–13 km range of mixing depths proposed by Straub & Martin-Del Pozzo (2001) for the recent eruptives of Popocatépetl.
Immediate sources of volatiles
There are several models that could explain the observed gas emissions at Popocatépetl. These include: (1) degassing from erupted silicate melt; (2) degassing from large bodies of unerupted magma intruded at shallow levels into the volcanic edifice; (3) fumarolic remobilization; (4) breakdown of sulfur-bearing magmatic phases; (5) assimilation of carbonate and evaporite country rock; (6) a separate volatile phase in the erupted magma. We also present convection of magma in the volcanic conduit supplemented by a separate volatile phase in the convecting magma as a viable scenario to explain the degassing process.
Degassing from erupted silicate melt
Volatile contents of glass inclusions and coexisting degassed matrix glasses can be compared to estimate the extent of SO2 degassed from silicate melt during eruption. This calculation is performed using the equation (Gerlach et al., 1996)
 | (2) |
melt is the density of the silicate melt phase, S is the difference between sulfur concentration in glass inclusions and matrix glass, and Fmelt is the volume fraction of melt (i.e. 1 – crystallinity). The amount of degassing from the silicate melt portion of Popocatépetl's erupted magma is governed by the relative contribution of volatiles from the melts of the two end-member magmas. As the bulk proportion of mixing between the end-member magmas prior to eruption is unknown and/or may vary with time, we investigate the entire range of mixing proportions represented by individual clasts, i.e. 45:55 wt % and 85:15 wt % silicic: mafic magma. The amount of sulfur each magmatic end-member contributes to degassing can be estimated by taking the difference in sulfur content between glass inclusions and the matrix glass for mafic and silicic end-members. Equation (2) yields 0·13 Mt SO2 for the 45:55 mixture and 0·04 Mt SO2 for the 85:15 mixture. Regardless of mixing proportion, sulfur dioxide emitted from the melt portion of 3·1 x 107 m3 of erupted magma is over two orders of magnitude lower than the 14·8 Mt SO2 emission measured by COSPEC during the magmatic period under study. Even if 100% of the exsolved gas came from the melt phase of an aphyric mafic end-member with a volume of 3·1 x 107 m3, SO2 emission increases to only 0·27 Mt.
HCl and HF emissions can be estimated using expressions analogous to equation (2). Chlorine emission calculations utilizing petrological data yield a total emission of 0·02 Mt HCl for the recent 4 year 10 month magmatic period. Similar to the case for SO2 described above, the calculated HCl emission is approximately two orders of magnitude less than the 1·8 Mt HCl total emission measured by passive FTIR (Goff et al., 2001). The amount of fluorine that has exsolved from the melt portion of erupted magma is difficult to assess because the matrix glass F content significantly overlaps the F contents in glass inclusions in phenocrysts. As a result, the mass of degassed F cannot be reliably calculated.
Degassing from large bodies of unerupted magma intruded to shallow levels
Degassing of magma emplaced at shallow levels within the volcanic edifice has been proposed as a source of ‘excess’ sulfur and other volatiles (Rose et al., 1982; Andres et al., 1991). To produce 14·8 Mt SO2, melt degassing from a hypothetical silicic end-member composition (S = 100 ppm, melt = 2340 kg/m3, Fmelt = 0·65) would require complete degassing of 50 km3 magma. Melt degassing from a mafic end-member composition (S = 1950 ppm, melt = 2470 kg/m3, Fmelt = 0·9) would require a volume of 2 km3 magma to completely degas.
Efficient degassing of silicate melt requires either high degrees of crystallization or low pressure. Our analyses of glass inclusions suggest Popocatépetl mafic and silicic melts contain 3 wt % H2O prior to eruption. According to the H2O solubility model of Moore et al. (1998), the mafic and silicic magma compositions of Popocatépetl are saturated with 3 wt % H2O at a pressure of 700 bars (3 km depth). CO2 will probably saturate at greater depth, but the volume fraction of this CO2-rich vapor will be small until depths at which appreciable H2O comes out of solution (Newman & Lowenstern, 2002).
If we assume complete degassing of a 2 km3 magma body intruded at 3 km depth as a result of crystallization, we can employ modified Mogi modeling (Okada, 1992) to estimate the surface deformation associated with intrusion of such a magma body. Calculations using the Mogi.m code (P. Cervelli, personal communication, 2000) predict 57 m of vertical displacement. Incomplete degassing would require more magma, leading to greater surface deformation. Likewise, emplacement at depths shallower than 3 km would also lead to greater vertical deformation than the predicted 57 m. Continuous GPS deformation monitoring of Popocatépetl has so far shown negligible deformation (Cabral-Cano et al., 1999). In conclusion, degassing from a shallow magma body recently intruded into the volcanic edifice is an unlikely source of ‘excess’ sulfur at Popocatépetl.
Fumarolic remobilization
Some sulfur emissions may be derived by fumarolic remobilization of pre-existing sulfur deposits within the volcanic edifice (M. O. Garcia, personal communication, 2000). Such sulfur deposits may take the form of fumarolic precipitates of native sulfur and a variety of other sulfur-bearing minerals filling cracks in wall rocks. Intrusion of fresh magma could provide the heat necessary to liberate sulfur from these deposits.
To produce the 14·8 Mt SO2 would require remobilization of 7·4 Mt of native sulfur. Such a mass of native sulfur would occupy 3·6 x 106 m3 ( = 2070 kg/m3 for native sulfur). This volume is equivalent to a 15 m wide, 1000 m long annulus of pure sulfur around a hypothetical 30 m radius cylindrical conduit. The existence of a pure sulfur deposit of this size seems unlikely. Additionally, remobilization of much smaller volume sulfur deposits that are widely disseminated in narrow cracks would require heating of an unreasonably large volume of wall rock to produce the observed sulfur emissions. The flux of sulfur derived from fumarolic remobilization is probably negligible compared with the total sulfur output.
Breakdown of sulfur-bearing magmatic phases: anhydrite and sulfide
The volume of silicate melt required to produce the observed SO2 emission can be reduced if melt degassing is supplemented by significant breakdown of sulfur-bearing phases such as anhydrite or sulfide. Primary anhydrite has not been observed in Popocatépetl magmas. This is consistent with our estimate of redox state for these lavas (NNO 0·5), which is not sufficiently oxidized to stabilize anhydrite (Carroll & Rutherford, 1987; Luhr, 1990).
Fe-rich sulfide globules are found as inclusions in pyroxene. In addition, Fe–Ni- and Fe–Cu-rich sulfide globules are found among the breakdown products in hornblende reaction rims and on the edges of phenocrysts in contact with silicate melt (Witter, 2003). These sulfides could contribute to sulfur output at Popocatépetl. Larocque et al. (2000) proposed that immiscible sulfide liquids break down during depressurization and degassing to form spongy Fe-oxides and sulfur gas. Rare spongy Fe-oxide grains are found on vesicle walls of recent pumices and lava blocks of Popocatépetl (Witter, 2003).
To produce the entire 14·7 Mt of SO2 gas not accounted for by melt degassing would require breakdown of 15 vol. % sulfide liquid in the erupted volume. If this occurred, we would expect to find a sub-equal amount of spongy Fe-oxide breakdown products in the recent Popocatépetl eruptives. As such a large concentration of sulfide liquid is unlikely and we find only trace amounts of spongy Fe-oxides, breakdown of sulfide liquid is unlikely to have produced more than a small portion of the measured SO2 emissions.
Assimilation of carbonate and evaporite country rock
Cretaceous marine carbonate, shale, and evaporite form outcrops near Popocatépetl and are presumed to compose part of the basement rocks that underlie the volcano. Early Tertiary intrusive bodies cut these marine rocks and it seems plausible that dikes feeding Popocatépetl would also intersect these rocks (Goff et al., 2001). Volatile-bearing carbonate and evaporite country rocks could be assimilated into Popocatépetl magma and provide a rich non-magmatic source for the emitted volatiles.
Goff et al. (2001) hypothesized that episodic ingestion and assimilation of carbonate country rocks may explain huge excursions in the CO2 flux from Popocatépetl. Large, short-term increases in the CO2/SO2 mass ratio, the presence of metamorphosed carbonate xenoliths in eruptives from the last 14 kyr, and the lack of an adequate alternate source of CO2 were provided as support for their hypothesis.
Although we do not have CO2 data from glass inclusions in phenocrysts from the recent eruptives of Popocatépetl to address this hypothesis, S, Cl, and F emissions appear to be decoupled from CO2 emissions. During large increases in the CO2/SO2 mass ratio, SO2/HCl and SO2/HF ratios remain stable (Goff et al., 2001). If the ‘excess’ CO2 is caused by carbonate assimilation then this suggests that the assimilated country rock does not contain appreciable quantities of S-, Cl-, and F-bearing minerals.
Sulfur isotope data for Popocatépetl gases indicate 
34S = +1·5 to +6·5
with an average of +3·4 (Goff et al., 1998). Gypsum from a regional evaporite deposit yielded a 34S value of 16·4. The sulfur isotopic composition of Popocatépetl gases falls within the range for gases from continental margins and island arcs (34S = 0 to 12; Taylor, 1986). Sulfur originating from the primitive mantle has 34S 1. Goff et al. (1998) suggested that isotopic enrichment of the gas phase over mantle values reflects assimilation of evaporite country rocks. Alternatively, if a more typical arc 34S magmatic signature for Popocatépetl magmas is assumed, then little to no contamination is required to produce the observed 34S values. Isotopic fractionation of sulfur between various phases and oxidation states can be expected to increase 34S without requiring assimilation (Ohmoto & Rye, 1979). Continued degassing of the magma reservoir can be expected to increase 34S still further, as described by McKibben et al. (1996). Both acid gas ratios and sulfur isotope systematics provide evidence that the contribution of evaporite to the volatile budget at Popocatépetl is probably relatively minor.
Separate volatile phase in erupted magma
Recent work has suggested that arc magmas are commonly volatile-saturated at depth. Many of these magma bodies may also harbor a substantial coexisting volatile-rich fluid phase prior to eruption (Luhr et al., 1984; Anderson et al., 1989; Lowenstern et al., 1991; Westrich & Gerlach, 1992; Lowenstern, 1993; Gerlach & McGee, 1994; Gerlach et al., 1994, 1996; Wallace & Gerlach, 1994; Wallace et al., 1995; Wallace, 2001). A separate volatile phase that exists at depth and ascends with the silicate liquid could be a rich source of magmatic gases to account for the large-magnitude gas emissions at Popocatépetl. The low dissolved Cl content measured in glass inclusions (<1000 ppm) is insufficient for the formation of a hydrosaline brine in equilibrium with Popocatépetl silicate melt (Webster et al., 1999). A separate volatile phase at Popocatépetl is more likely to be dominated by H2O and CO2. The composition and relative proportion of such a volatile-rich fluid can be calculated (Table 4).
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Let us assume that the gas output at Popocatépetl is carried entirely as bubbles within the 3·1 x 107 m3 magma erupted from March 1996 to December 2000. The erupted magma has a mass of only 74 Mt. If we assume unidirectional, homogeneous flow of only the erupted silicate liquid and crystals and the separate volatile phase, the ascending mixture would have consisted of >60 wt % separate volatile phase. It is unlikely that such high volatile contents can be maintained unerupted at depth for sufficient time to allow their accumulation. For comparison, the Pinatubo magma is inferred to have contained only
1–6 wt % separate volatile phase (Gerlach et al., 1996). Apart from this unreasonably high gas–silicate ratio, the hypothesis that the entire gas output of the volcano is carried by and emitted only with the erupted magma is inconsistent with the observations of discontinuous magma effusion and continuous gas emission (CENAPRED, unpublished data, 2000).
Likely sources for Popocatépetl volatiles
The hypothesis that gas emissions are derived from the degassing of unerupted silicate melt is still possible; however, we must provide a mechanism for efficiently extracting volatile components from a large volume of melt without erupting more than a small portion of that melt or accumulating significant quantities of the degassed residue in the volcanic edifice. Circulation of magma in the volcanic conduit could explain shallow, non-explosive degassing of large volumes of magma without excessive surface deformation or implausibly large gas–silicate mass ratios. Kazahaya et al. (1994) proposed convection of magma in the volcanic conduit as a mechanism for passive degassing at Izu–Oshima (basaltic) and Sakurajima (andesitic) volcanoes. Allard (1997) suggested a similar process at Mount Etna. Shinohara et al. (1995) modeled convection of granitic magma in a 300 m diameter conduit to explain volatile transport and the development of porphyry Mo deposits. Stevenson & Blake (1998) refined the fluid dynamic equations presented by Kazahaya et al. (1994) with experimental data. They used this revised formulation to explain degassing at Stromboli (basaltic) and the June 1980 gas emissions at Mount St. Helens (dacitic) by convective overturn of dense degassed magma with buoyant volatile-rich magma in a conduit connected to a deep magma chamber.
The conceptual model for conduit convection at Popocatépetl is as follows. Magma ascends to very shallow levels through a cylindrical or dike-shaped conduit from a magma body located at deeper levels. At a level where the bubble content reaches 30–60 vol. % (Eichelberger et al., 1986; Klug & Cashman, 1996), collapse of magmatic foam would force exsolved magmatic gases to travel out of the magma and into wall rocks (Jaupart & Allègre, 1991). The relatively dense, degassed magma descends back down into the volcano's plumbing system, making room for influx of fresh, low-density, volatile-rich magma. Following magma–gas separation, gases ascend vertically through permeable wall rocks to be emitted from the crater walls or margins of the dome. This model is similar to that presented by Kazahaya et al. (2002) for the rhyolitic volcano Satsuma–Iwojima.
The required volume of convecting magma is greatly reduced if the majority of the observed volatile output at Popocatépetl resides in a separate volatile phase. Slow accumulation of exsolved gas could take place at the saturated cap of a large magma chamber (Wallace, 2001). This type of bubble accumulation by compositional convection in a stratified magma body has been modeled numerically by Simakin & Botcharnikov (2001) to explain passive degassing at volcanoes such as Popocatépetl. In the model of Simakin & Botcharnikov (2001) exsolved volatiles are transported from the cap of the magma chamber to the surface through fumarole pipes.
To understand the specific case of Popocatépetl, an estimate of the depth of gas-saturation of Popocatépetl magma is needed. If we use a CO2/SO2 mass ratio of four (Goff et al., 2001) and an H2O/SO2 mass ratio between 5·1 (Case 1: 95 mol % H2O + CO2) and 15·7 (Case 2: 98 mol % H2O + CO2) based on calculations of the total gas flux (Table 4) and assume that all of the volatiles were originally dissolved in silicate melt, we can calculate initial melt CO2 and H2O contents and use these estimates to place constraints on the depth of gas saturation. A magma mixture that consists of 35 wt % mafic end-member and 65 wt % silicic end-member (an average proportion based on erupted clasts) would have a bulk initial S content of 780 ppm. The gas mass ratios presented above imply 6200 ppm initial CO2 content and 0·8 wt % (Case 1) and 2·5 wt % (Case 2) initial H2O content. The initial H2O content calculated for Case 2 is consistent with the 3 wt % initial H2O estimates obtained from glass inclusions in phenocrysts. Using either 0·8 wt % or 3 wt % as the initial H2O and 6200 ppm as the initial CO2 content of the melt, the VolatileCalc 1.1 program (Newman & Lowenstern, 2002) calculates saturation pressures >5 kbar utilizing either the basalt or rhyolite solubility models. This suggests that Popocatépetl magma begins to exsolve gas bubbles at depths >20 km. The first bubbles to exsolve will be CO2-dominated; however, once the magma reaches low pressures (<1·5 kbar) the coexisting vapor will be H2O-rich (Newman & Lowenstern, 2002). From initial saturation to about 6 km depth, the CO2-dominated vapor phase will make up an insignificant portion of the magma (<2 vol. %). It is not until the magma reaches <2 km depth that volatiles can make up a sufficient proportion of the magma (30 vol. % or more) to facilitate efficient magma–gas segregation.
For the case of Popocatépetl, we propose that (1) silicate melt and exsolved gas are both sources of volatiles and (2) the mechanism for passive degassing of the large quantities of volatiles involves convection of volatile-saturated magma in the volcanic conduit (Fig. 8). We explored a range of volumes of silicate material (0·1–3 km3 for the range H2O + CO2 = 95–98 mol %) that consists of 65:35 wt % silicic:mafic end-members with 25 vol. % crystals (Table 5). The masses of SO2 and HCl derived from the silicate material are supplemented by a separate volatile phase until the observed masses of SO2 and HCl emissions are most closely matched. The calculated weight proportion of exsolved gas has the range 1–56 wt % (Table 5). HF emissions were not included in the calculation because of large uncertainties in the mass of HF emission from silicate melt. Volumes of silicate material <0·1 km3 require high proportions of exsolved gas (>38 wt %). Volumes of silicate material greater than 4 km3 would release quantities of HCl that are greater than the observed HCl emission. Although all of the mass balance calculations shown in Table 5 are consistent with the SO2 and HCl emission data, the scenarios that involve separate volatile phases smaller than 10 wt % are the most likely. An H2O-dominated separate volatile phase at 2 kbar and 1000°C would have a density of 500 kg/m3. At this density, a separate volatile phase that comprises >10 wt % of a magma would make up >30 vol. % of the magma body. Permeability of the magma is thought to begin at porosity values >30 vol. %, therefore, magmas with these porosities are unlikely to be stable storage reservoirs as they would probably readily degas into country rock or erupt as a result of their own buoyancy (Wallace, 2001). In summary, the best fit for matching the SO2 and HCl emission data, while also keeping the separate volatile phase a reasonable size, is achieved when degassing 1–3 km3 of silicate melt with an additional 1–9 wt % separate volatile phase that contains 95–98 mol % H2O + CO2.
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To illustrate an example of one possible scenario, our data suggest that complete degassing of
1 km3 of magma (consisting of 65 wt % dacite, 35 wt % basaltic andesite, with 25 vol. % crystals, and no exsolved gas) would release 2·5 Mt SO2 and 0·6 Mt HCl from the melt. Chemical breakdown of sulfide in the 1 km3 of Popocatépetl magma would supply an additional 0·3 Mt SO2. An accompanying separate volatile phase that comprises 5 wt % of the magma (Case 1: 95 mol % H2O + CO2), with volatile proportions as outlined in Table 5, would contain 12·0 Mt SO2 and 1·5 Mt HCl. This process would liberate 14·8 Mt SO2 and 2·0 Mt HCl. These values are within error of the measured gas emissions for the 4 year 10 month time period under study.
Our results imply that a hypothesized 1 km3 of silicate material degassed at an average rate of nearly 7 m3 magma per second from March 1996 to December 2000. If we assume an average magma ascent velocity of 1 cm/s, the ascending magma would occupy a cylinder with a 15 m radius to supply the calculated 7 m3 magma per second. These magma ascent rates and conduit diameters are comparable with those inferred for other andesitic and dacitic, dome-producing volcanoes (Swanson & Holcomb, 1990; Rutherford & Hill, 1993; Nakada & Motomura, 1999; Murphy et al., 2000).
Fluid dynamic modeling
Although circulation of magma in the volcanic conduit is easily imagined for low-viscosity basaltic systems such as Izu–Oshima (Kazahaya et al., 1994), Etna (Allard, 1997), and Stromboli (Stevenson & Blake, 1998), the higher viscosities found in the andesitic to dacitic system of Popocatépetl might inhibit magma convection within the volcanic conduit. As a test of the feasibility of convection of the Popocatépetl magma we use a simple fluid dynamic model. The following equation is based on the theory of laminar flow in a vertical cylindrical conduit and experiments by Stevenson & Blake (1998):
 | (3) |
), g = 9·8 m/s2, is the density difference between degassed and undegassed magma (kg/m3), and µd is the viscosity of the descending, degassed magma (Pa s). We can rearrange equation (3) to solve for conduit radius, Rc:
 | (4) |
). For density and viscosity we use the same 65 wt % dacite–35 wt % basaltic andesite mixture with 25 vol. % crystals as above. This mixture has a melt with 67 wt % SiO2, T = 1000°C, and 3 wt % H2O dissolved in the melt prior to degassing. We consider two cases: degassing of 2·0 wt % H2O with 1·0 wt % H2O remaining in the melt (equivalent to 150 bar or 600 m depth; Fig. 8) and degassing of 2·9 wt % H2O with 0·1 wt % H2O remaining in the melt (equivalent to 1 bar or surface conditions). These two cases result in calculated density differences of 59 kg/m3 and 92 kg/m3, respectively, using the model of Lange (1994). This density difference does not include the dramatic density decrease caused by bubble growth at low pressures. The viscosities of the undegassed and degassed magmas are estimated by taking the average viscosity calculated from the non-Arrhenian viscosity models for andesite melts (Richet et al., 1996) and rhyolite melts (Hess & Dingwell, 1996), coupled with the Einstein–Roscoe equation (Lejeune & Richet, 1995), which accounts for the effect of crystals on the bulk viscosity. The effect of bubbles on viscosity is expected to be small (less than a five-fold decrease; Stein & Spera, 2002) and is not included in the following calculations. The viscosity of undegassed magma is 104·0 Pa s. The viscosities of the descending magmas (µd) that have 1·0 wt % H2O and 0·1 wt % H2O remaining in the melt are 105·3 and 107·3 Pa s, respectively. Utilizing equation (4), the calculated conduit radius (Rc) is 14 m for the 1·0 wt % H2O remaining case and 38 m for the 0·1 wt % H2O remaining case. These values for conduit radii are entirely plausible for Popocatépetl. For comparison, Kazahaya et al. (2002) calculated a conduit radius Rc of 25 m for the rhyolitic magma of Satsuma–Iwojima with a degassing rate Q = 10 m3/s, = 50 kg/m3, and µd = 106·1 Pa s. Based on the dimensionless ratio R* = Ra/Rc = 0·6 (Stevenson & Blake, 1998) and magma flux of 7 m3/s, the calculated values of Rc for Popocatépetl correspond to ascending cylinders of magma with radii (Ra) of 8 m and ascent rates of 3·2 cm/s for the 1·0 wt % H2O remaining case and 23 m and 0·4 cm/s for the 0·1 wt % H2O remaining case. These calculated values of Ra bracket the radius of the ascending cylinder of magma (15 m) that is predicted by assuming a magma ascent velocity of 1 cm/s. It should be noted that if, instead, 3 km3 of magma were degassed, then Q = 21 m3/s and Ra = 11 m for the 1·0 wt % H2O remaining case and 30 m for the 0·1 wt % H2O remaining case. These values are also reasonable.
The high temperature of the mixed magma (950–1050°C) enhances the potential for conduit circulation at Popocatépetl. An elevated magmatic temperature reduces magma viscosity and eases silicate–gas separation. The presence of a low-density, separate volatile phase could serve to initiate magma ascent and would enhance circulation by adding a potent buoyancy force. Circulation of magma in the volcanic conduit and degassing from both the silicate melt and a modest separate volatile phase is a plausible mechanism to explain the observed gas emissions at Popocatépetl.
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TOPABSTRACTINTRODUCTIONANALYTICAL METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAAPPENDIXREFERENCESREFERENCES |
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The current manifestation of unrest at Popocatépetl involves the ascent of a basaltic andesite magma with 53 wt % SiO2 and
3 wt % H2O. This mafic end-member crystallized over a pressure range of 1–4 kbar (4–16 km) and then mixed with silicic magma at a depth of 4–6 km. Hornblende crystallized during or after magma mixing. The silicic end-member magma, presumed to currently reside beneath Popocatépetl, is a dacite with 65 wt % SiO2 containing 3 wt % H2O that probably crystallized at pressures of 1·5 kbar. The silicic and mafic end-member magmas mixed in proportions in the range 45:55 wt % to 85:15 wt % silicic:mafic magma to produce the erupted clasts. Degassing from erupted silicate melt in addition to sulfide breakdown in that same magma accounts for only 0·7 wt % of the observed SO2 emission. Although assimilation of carbonate country rocks appears to be a plausible source for large excursions in CO2 emission, evaporite assimilation as a source of S, Cl and F gases seems unlikely. If the source of volatiles at Popocatépetl is a separate volatile phase formerly contained only in the erupted magma then the silicate–volatile mixture would consist of >60 wt % low-density volatile phase. This is implausible. If there is no excess volatile phase at depth, then 2 km3 of the mafic end-member magma or 50 km3 of the silicic end-member magma must completely degas to account for the observed gas emissions. Complete degassing of 1–3 km3 of mixed magma with 1–9 wt % separate pre-existing volatile phase can account for the observed gas emissions. Circulation of magma in the volcanic conduit may be an effective mechanism for non-explosive degassing of large quantities of volatiles at Popocatépetl. Any volcano that (1) emits large quantities of gases, (2) does not erupt sufficient quantities of magma to produce the observed gas emissions, (3) has no alternate source of volatiles other than magma, and where (4) intrusion of magma into the volcanic edifice is minimal, is a potential candidate for degassing by convection of magma in the conduit. For a volcano that exhibits these qualities, one may collect data on the flux of gas emissions, determine pre-eruptive gas content, magma temperature, density and viscosity from samples of coeval magma, then use analytical pipe flow equations to estimate conduit size and determine whether or not magma convection is feasible for the volcano in question. In our view, conduit convection is probably a common mechanism to explain long-term, non-explosive, ‘excess’ degassing behavior observed at many volcanoes.
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SUPPLEMENTARY DATA |
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TOPABSTRACTINTRODUCTIONANALYTICAL METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAAPPENDIXREFERENCESREFERENCES |
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Supplementary data are available at Journal of Petrology online.
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APPENDIX |
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TOPABSTRACTINTRODUCTIONANALYTICAL METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAAPPENDIXREFERENCESREFERENCES |
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Rock descriptions
Rock samples used in this study were obtained from various small explosive events in 1997 and 1998. All samples are vesicular or were emplaced hot as evidenced by nearby burnt vegetation or melted equipment. All samples represent juvenile magma. We focused our efforts on pumiceous samples from the eruption of June 30, 1997 and dome blocks from the eruption of January 1, 1998. Petrographic studies were also undertaken on samples from other small explosive events in December 1997 and January, March, and April 1998.
One suite of samples was derived from the moderately explosive eruption of June 30, 1997. The sampled eruption products include volumetrically dominant, brown andesite pumice and subordinate, white dacite pumice. Some clasts are banded. The eruption had two pulses and lasted a total of 3
h. The eruption column rose to a height of at least 13 km above sea level and showered ash on Mexico City. It was one of the most powerful eruptive events at Popocatépetl since 1994 (GVN, 1997a).
A second suite of samples was derived from a smaller eruption on January 1, 1998. This eruption produced a crystal-rich, poorly vesicular, gray dacite containing volumetrically minor, deformed blebs of darker material. The maximum intensity of the eruption lasted 1 min and the eruption plume reached 5 km above the crater (GVN, 1997b).
The various samples of ballistic ejecta from dome explosions in December 1997 to April 1998 are all light to dark gray, dense, crystal-rich lava. Some samples are micro-vesicular whereas others show subtle bands of lighter and darker material.
Petrography and igneous textures
The 1997 to 1998 Popocatépetl eruptive products contain phenocrysts and microphenocrysts of olivine, clinopyroxene, orthopyroxene, plagioclase and hornblende. Microphenocrysts of titanomagnetite, ilmenite, and rare apatite are distributed in a microcrystalline groundmass. Chromian spinel commonly occurs as inclusions in olivine. Titanomagnetite and ilmenite occur as inclusions in pyroxene and rarely in plagioclase. Apatite and rare globular sulfide are also found as inclusions, most commonly in pyroxene. Anhedral xenocrysts of quartz are rare. The groundmass varies from light to dark in color. Dark groundmass consists of a dense network of plagioclase and Fe–Ti oxide microlites in brown glass. Light-colored groundmass contains sparse plagioclase microlites in colorless glass.
Olivine (0·1–1·5 mm) comes in two textures. Larger-sized olivine (>0·5 mm) are the more abundant and are generally subhedral. They contain abundant chromian spinel inclusions (15 µm diameter) but no glass inclusions. Some of the large-sized, Cr-spinel-rich olivines occur as glomerocrysts of two or more crystals. Some of these glomerocrysts as well as some individual grains of olivine are anhedral and display embayed margins that have reacted extensively with the surrounding silicate liquid. These large, Cr-spinel-rich olivines are probably xenocrystic.
The other type of olivine, microphenocrystic (<0·3 mm), is much less abundant. Some have a skeletal (arrow-shaped) habit suggesting rapid growth from the surrounding silicate melt. Olivine microphenocrysts contain few, if any, chromian spinel inclusions but do contain 5–40 µm diameter inclusions of clear brown glass.
Hornblende phenocrysts commonly have an acicular habit and are up to 1 mm in length. More equant hornblende microphenocrysts are generally 0·2 mm x 0·3 mm in size. All hornblendes have dark ‘gabbroic-type’ (Garcia & Jacobson, 1979) breakdown rims 30–100 µm across made of tiny plagioclase, pyroxene, Fe–Ti oxides, and rare sulfides. Rare hornblende phenocrysts and microphenocrysts possess small plagioclase or pyroxene cores. Some hornblende grains form partial overgrowths on olivine and pyroxene. Glass inclusions in hornblende are rare.
Orthopyroxene rims on Cr-spinel-bearing olivine are rare and narrow in our samples from the June 30, 1997 eruption. Orthopyroxene rims are more common and better-developed on olivines in samples from the January 1, 1998 dome explosion event and the December 1997 to April 1998 dome rocks. Straub & Martin-Del Pozzo (2001) also reported orthopyroxene rims on olivine whereas Kolisnik (1990) found abundant and well-developed orthopyroxene reaction rims on olivine in Quaternary Popocatépetl lavas. The higher abundance of orthopyroxene rims in the dense lavas and dome rocks probably reflects longer cooling times, compared with the more rapidly quenched pumices erupted on June 30, 1997. Orthopyroxene rims are not found on olivine microphenocrysts.
Anhedral Fe–Ti oxides commonly occur as 5–100 µm inclusions in clinopyroxene, orthopyroxene and, less commonly, plagioclase phenocrysts. Titanomagnetite and ilmenite occur as microphenocrysts (up to 250 µm in diameter) and as microlites (<10 µm in diameter) in the groundmass. Rare Fe–Ti oxide grains contain glass inclusions up to 30 µm in size. Titanomagnetite and ilmenite occur separately as well as in intimate contact. Exsolution of titanomagnetite in ilmenite or vice versa has not been observed.
Orthopyroxene and clinopyroxene are common as phenocrysts in all samples and are up to 1 mm in length. Clinopyroxene phenocrysts are generally euhedral to subhedral. Most orthopyroxene phenocrysts are subhedral; however, some appear reacted and embayed. Both clinopyroxene and orthopyroxene phenocrysts contain abundant glass inclusions 5–50 µm in diameter. Rare, globular, sulfide inclusions 5–15 µm in size and lath-shaped apatite 10–50 µm long also occur as inclusions in pyroxene. Microphenocrysts of clinopyroxene and orthopyroxene (30–300 µm) are common, frequently have octagonal or lath-shaped cross-section, and rarely contain melt or mineral inclusions.
Plagioclase is by far the most abundant phenocryst in the recent Popocatépetl samples, reaching 2 mm in size. Many are rich in glass inclusions and concentric bands of glass inclusions in plagioclase are common. Complex, oscillatory zonation as well as subhedral or embayed plagioclase crystals are common. Many plagioclase grains have sieve-textured cores and/or evidence of mineral dissolution followed by regrowth. In contrast, other plagioclase grains are relatively pristine, euhedral and/or inclusion free. Microphenocrysts of plagioclase (30–300 µm) are common and frequently occur as euhedral laths.
The felty textured groundmass in dark-colored pumiceous hand samples is dominated by randomly oriented plagioclase and Fe–Ti oxide microlites. Some of the plagioclase microlites are lath-shaped whereas others are hopper-shaped. Very little dark-colored, quenched glass is present. Some areas of the matrix in the dark-colored pumiceous hand samples have partially aligned lath-shaped plagioclase microphenocrysts and microlites that show a pilotaxitic texture. In light-colored pumiceous hand samples, groundmass is composed of abundant, colorless, microlite-poor glass with a vesicular hyalopilitic texture.
Nearly all samples contain crystal clots of clinopyroxene + orthopyroxene + Fe–Ti oxides or plagioclase + clinopyroxene + orthopyroxene + Fe–Ti oxides. They usually occur as one glomerocryst per thin section. Pyroxene-dominated glomerocrysts often have a granular texture. However, glomerocrysts that contain plagioclase exhibit hypidiomorphic granular texture.
The mineralogy of the dome explosion samples of December 1997 and January, March, and April 1998 is identical to the mineralogy of the more pumiceous samples of the June 30, 1997 eruption. However, texturally, the matrix in the dome samples is generally coarser grained, with a smaller proportion of interstitial glass. In addition, textural variations, such as color banding (reflecting regions of microlite-rich vs microlite-free matrix), are less pronounced and subtler in the dome samples compared with the more dramatically banded pumice samples.
Modes and mineral chemistry
Modal analyses were performed by point counting 1300 points per thin section (Table 2). Hand samples that are darker in color generally have a higher proportion of olivine. White pumice from the June 30, 1997 eruption is rich in plagioclase and pyroxene phenocrysts, but is poor in olivine phenocrysts and plagioclase microlites. In contrast, brown pumice ejected in the June 30, 1997 eruption is richer in olivine, has abundant plagioclase microlites (35 vol. %) and is poorer in pyroxene and plagioclase phenocrysts than their white pumice counterparts. Vesicular dome rocks ejected on January 1, 1998 have modal proportions that are similar to the white pumice of June 30, 1997 (rich in plagioclase and pyroxene phenocrysts and poor in olivine phenocrysts and plagioclase microlites). However, the January 1, 1998 dome rocks are richer in total phenocrysts (60 vol. %) than the June 30, 1997 white pumices (30–40 vol. %). The brown pumice from June 30, 1997 has the lowest proportion of phenocrysts (<20 vol. %). Modal analysis of juvenile rocks by Athanasopoulos (1997) from the first magmatic material effused in March–April 1996 is shown in Table 2 for comparison.
Olivine
Cores and rims of 22 olivine xenocrysts were analyzed by EPMA along with an additional 11 core analyses of olivine microphenocrysts (Electronic Appendix Table A4). Large, Cr-spinel-rich olivine xenocrysts contain broad, compositionally uniform cores (Fo87–91) with narrow (20–70 µm) normally zoned rims (Fo73–88). Glass inclusion-bearing olivine microphenocrysts also have uniform cores but have a lower Fo content and more limited compositional range (Fo84–86). High NiO content (2300–6700 ppm) varies sympathetically with Fo87–91 in the olivine xenocrysts. Glass inclusion-bearing olivine microphenocrysts, however, have much lower NiO (1100–1800 ppm) for Fo84–86. The total range of olivine compositions presented in this study is very similar to that reported for Popocatépetl's Quaternary lavas (Kolisnik, 1990). Our results show a slightly broader range in Fo than the recent eruptives studied by Straub & Martin-Del Pozzo (2001) and are comparable with the data for olivine from the March 1996 dome (Athanasopoulos, 1997).
Clynne & Borg (1997, fig. 2B) distinguished between high-alumina olivine tholeiites, calc-alkaline basalts, and mafic andesites from the Lassen Peak region on a plot of Fo vs CaO content in olivine. Glass inclusion-free olivine xenocrysts from Popocatépetl plot within the field of calc-alkaline mafic andesites but outside the fields for calc-alkaline basalts and high-alumina olivine tholeiites. In contrast, glass inclusion-bearing olivine microphenocrysts from Popocatépetl plot outside the three fields in a region of lower Fo-content.
Chromian spinel
Twenty-two analyses of chromian spinel were conducted by EPMA (Electronic Appendix Table A5). Chromian spinel core compositions are rich in Cr, Al, and Mg. Chromian spinel rims, however, that are in contact with silicate melt have a much higher content of Fe2+, Fe3+ and Ti [all analyses recalculated according to Stormer (1983)].
Clynne & Borg (1997, fig. 6) distinguished rock types in the Lassen Peak area using chromian spinel compositions on the plot Mg-number [= Mg/(Mg + Fe2+)] vs Cr-number [= Cr/(Cr + Al)]. With this diagram, chromian spinel in mafic volcanic rocks from the Lassen region, Jorullo basalt to basaltic andesites (Luhr & Carmichael, 1985), and MORB lavas from the Lamont seamount chain (Allan et al., 1989) can be compared with those of Popocatépetl. Popocatépetl chromian spinel plots primarily within the fields of calc-alkaline mafic andesite from the Lassen region and Jorullo basalt to basaltic andesite. The relatively high Cr content may reflect a rather primitive source region for Popocatépetl's mafic input. Our analyses of Cr-spinel are similar to those presented in Straub & Martin-Del Pozzo (2001) for recent Popocatépetl eruptives.
Fe–Ti oxides
Twenty grains of titanomagnetite and 20 grains of ilmenite were analyzed by EPMA (Electronic Appendix Table A6). Titanomagnetite and ilmenite compositions have the range XUsp = 0·29–0·48 and XIlm = 0·65–0·85, respectively [all analyses recalculated according to Stormer (1983)]. The compositional range of Popocatépetl Fe–Ti oxides is broadly similar to those found in pumices from the 2·2 ka eruption of Mount Rainier (Venezky & Rutherford, 1997). Compared with 1991 Unzen Fe–Ti oxides, Popocatépetl titanomagnetites have somewhat higher XUsp whereas ilmenites have lower XIlm (Venezky & Rutherford, 1999).
Pyroxene
Twenty-four core and 23 rim analyses of orthopyroxene were performed by EPMA along with 11 additional analyses of orthopyroxene host mineral near glass inclusions. Three analyses were obtained of an orthopyroxene host near a sulfide inclusion. Twenty-three core and 21 rim analyses of clinopyroxene were performed by EPMA along with 23 additional analyses of clinopyroxene host mineral near glass inclusions. Analyses of pyroxene cores and rims are shown in Electronic Appendix Table A7. Pyroxene core compositions occur as high-Mg-number and low-Mg-number phases [Mg-number = Mg/(Mg + Fe2+)]. A few pyroxene grains do not fit into the high- or low-Mg-number groupings. Two pyroxenes have intermediate compositions. One clinopyroxene is interpreted to be a xenocryst because of anomalously high Al2O3 content. One orthopyroxene is interpreted to be a xenocryst because of anomalously high Cr2O3 content and low Mg-number. Plots of SiO2, Al2O3, and Cr2O3 vs Mg-number help delineate the two groups of clino- and orthopyroxene (Electronic Appendix Fig. A1).
Pyroxenes are characterized by broad, compositionally uniform cores with or without narrow, zoned rims. We interpret the pyroxene core compositions as an indication of the origin of the pyroxene. Compositional variations in the pyroxene rims are interpreted to result from magma mixing. High- and low-Mg-number clinopyroxene cores are augite with compositions En47–49Fs8–10Wo42–44 (Mg-number 83–86 and Cr2O3 0·2–0·7 wt %) and En40–44Fs13–18Wo40–45 (Mg-number 70–78 and Cr2O3 < 0·3 wt %), respectively. High-Mg-number orthopyroxene cores are bronzite with composition En82–83Fs14–15Wo2–4 (Mg-number 84–85 and Cr2O3 0·1–0·4 wt %). Low-Mg-number orthopyroxene cores are hypersthene with composition En61–69Fs29–35Wo2–3 (Mg-number 63–72 and Cr2O3 < 0·15 wt %). EPMA data for pyroxene in recent Popocatépetl eruptives by Athanasopoulous (1997) and Straub & Martin-Del Pozzo (2001) cover the same compositional range as the analyses presented here. However, the high-Mg-number orthopyroxenes in their studies occur as mantles around olivine. This is in contrast to the high-Mg-number orthopyroxene phenocrysts presented in this study. Kolisnik (1990) also found high- and low-Mg-number populations of clino- and orthopyroxene phenocrysts in her study of Quaternary lavas of the Popocatépetl edifice that cover the same compositional range as those presented here.
Core-to-rim zonation in pyroxene is highly variable. High-Mg-number pyroxenes are generally normally zoned (up to En = 21 mol %) whereas low-Mg-number pyroxenes are predominantly reversely zoned (up to En = 15 mol %). Many pyroxene grains, however, are unzoned or have weak normal to reverse zoning (En < 3 mol %). Zonation variations in pyroxene are similar in magnitude to that found in studies by Kolisnik (1990), Athanasopoulous (1997), and Straub & Martin-Del Pozzo (2001).
Plagioclase
Twenty-four plagioclase core and 20 plagioclase rim analyses were performed by EPMA (Table A8). Plagioclase phenocryst core compositions cover the range An28–61. Rim compositions cover a similar range. Zonation is dominantly reverse (up to An = 20 mol %) but some phenocrysts are normally zoned (up to An = 5 mol %). Petrographic observations discussed previously (oscillatory zonation and dissolution surfaces) suggest that zoning patterns in plagioclase are much more complex than the zoning patterns revealed by individual core and rim EPMA data. It is apparent that there is no correlation between plagioclase composition and phenocryst texture. In addition, plagioclase compositions do not fall into two clearly defined compositional groups, as do the pyroxenes.
Our plagioclase data are similar to those of Straub & Martin-Del Pozzo (2001) for other recent Popocatépetl eruptives. In an exhaustive study of plagioclase zonation in Quaternary lavas from Popocatépetl, Kolisnik (1990) found plagioclase phenocrysts with a broader compositional range (An21–71) characterized by complex discontinuous zoning patterns. She also inferred repetitive dissolution and renewed growth. The extremely heterogeneous plagioclase crystal populations had little correlation of zoning between coexisting grains with calcic excursions in An content of up to 37 mol %.
Hornblende
Seventeen major element analyses of hornblende were performed by EPMA. Of these, seven were also analyzed for volatile elements S, Cl, and F (Electronic Appendix Table A9). Popocatépetl hornblende has a restricted compositional range with Mg-number 63–72 [Mg-number = Mg/(Mg + FeT)], 9–13 wt % Al2O3, low concentrations of Cl (200 ppm) and high F (2100–5400 ppm). This contrasts with some subduction zone volcanoes such as Unzen and Montserrat that erupt magmas containing two compositionally distinct amphiboles. Popocatépetl hornblende compositions (Electronic Appendix Fig. A2) are similar to groundmass amphibole (magnesiohastingsite) in mafic inclusions in Montserrat andesite (Murphy et al., 2000) and groundmass amphibole (pargasite) from Unzen volcano (Sato et al., 1999). They plot well outside the field for amphibole phenocrysts (magnesiohornblende) in Montserrat andesite and Unzen dacite.
Utilizing the amphibole nomenclature scheme of Leake et al. (1997) and depending on how they are recalculated, Popocatépetl hornblende can be classified as the calcic amphiboles tschermakite, pargasite, or magnesiohastingsite. The low Na content and high Ca content of Popocatépetl hornblende places them in the calcic group like many amphiboles found in subduction zone magmas. The (Na + K) content in the A-site differentiates the nomenclature between tschermakite and pargasite–magnesiohastingsite at (Na + K)A = 0·5. Popocatépetl hornblende has (Na + K)A 0·5. The relative proportions of AlVI and Fe3+ distinguish pargasite and magnesiohastingsite. Ferric–ferrous recalculation schemes for amphibole generally produce large errors, making exact classification difficult. As a result of these considerations we choose to refer to the Popocatépetl amphiboles under scrutiny with the more generic name ‘hornblende’. What is important to emphasize, however, is that despite the difficulty in nomenclature, Popocatépetl hornblende has a restricted compositional range. EPMA of hornblende rims in contact with silicate melt was not possible because of the ubiquitous presence of breakdown rims on hornblende.
Recalculating hornblende EPMA data on the basis of 23 oxygens and utilizing the ferric–ferrous scheme of Holland & Blundy (1994) gives the following average chemical formula:
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from Popocatépetl's March 1996 dome.
Apatite
Popocatépetl apatite is classified as fluorapatite (Table A10). Systematic zonation of F, Cl, S, Si, Na, Mn, Fe, Mg and Sr in apatite was not found. Recalculating apatite EPMA data on the basis of 13 (O, OH, F, Cl) gives the following chemical formula:
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Our analyses are similar to those of other volcanic apatites (Peng et al., 1997).
Sulfide
Sulfide globules (Electronic Appendix Table A11) contain 33–40 wt % S and are rich in Fe (30–58 wt %). Sulfide inclusions in pyroxene are the most Fe-rich. Other sulfide globules found in contact with silicate melt, between two differing minerals, or in hornblende reaction rims have variable concentrations of Cu (up to 27 wt %) and Ni (up to 33 wt %).
Sulfide–sulfate wavelength shift analysis
As another measure of oxygen fugacity, two glass inclusions in olivine were analyzed for sulfide–sulfate using the wavelength shift of sulfur K X-rays (Carroll & Rutherford, 1988). Analytical conditions were 7 kV accelerating voltage, 80 nA beam current, and 10 µm beam diameter. We collected X-ray counts simultaneously from two PET diffraction crystals with the spectrometer aperture window set at its narrowest setting to maximize the spatial resolution of the spectrometer's position. The separation between S6+ and S2– is 0·00032 sin
units. We performed wavelength scans across 0·00072 sin units in 0·000018 sin unit steps counting 30 s at each step to fully define the curve of the sulfur peak. The electron beam was blanked for 30 s between each step to allow the sample to cool. The low accelerating voltage combined with blanking the beam minimized the effects of element mobility and oxidation during analysis (Métrich & Clocchiatti, 1996). Analytical accuracy of the method was checked on basalt glass standard VG-2 and NBS soda-lime glass 620. The minimum sulfur content necessary to adequately define the sulfur peak at these analytical conditions is 500 ppm.
The results were converted to oxygen fugacity using the method of Wallace & Carmichael (1994). This method assumes that sulfur is dissolved as S6+ and S2– in the glass inclusion and that there is a linear relationship between the S6+/Stotal ratio and the wavelength shift. Recent X-ray absorption spectroscopy measurements (Paris et al., 2001) support the validity of these assumptions. Measurements suggest that the silicate melt in olivine-hosted glass inclusions has an oxygen fugacity of NNO +1·2 log units and the sulfur dissolved in the melt is dominated by sulfate (87 mol %). Métrich & Clocchiatti (1996) commented that oxidation of sulfur in glasses can be caused by post-entrapment changes in the Fe2+/Fe3+ ratio of the melt, H2 loss (by diffusion), and heating as a result of long time periods under the electron beam. In our case, limited post-entrapment crystallization of the host olivine probably minimizes the Fe2+/Fe3+ evolution of the melt. Furthermore, redox calculations using the scheme of Kress & Ghiorso (2004) suggest that the ferric–ferrous ratio of the melt will not change substantially during limited post-entrapment crystallization. Hydrogen loss is difficult to assess but may be significant considering the high beam currents, long count times, and hydrous nature of the glass inclusions compared with the nominally anhydrous standard glasses.
Data tables
The data tables found in the Electronic Appendix contain the following information:
- Table A1. Mean error associated with electron microprobe analysis of mineral and glass standards.
- Table A2. Relative errors based on counting statistics for electron microprobe analyses of minerals and glasses.
- Table A3. Comparison of analyses of S, Cl, and F in various glass standards by EPMA and wet chemistry.
- Table A4. Electron microprobe analyses of olivine.
- Table A5. Electron microprobe analyses of Cr-spinel.
- Table A6. Electron microprobe analyses of titanomagnetite and ilmenite.
- Table A7. Electron microprobe analyses of clinopyroxene and orthopyroxene (Sheets 1 and 2).
- Table A8. Electron microprobe analyses of plagioclase.
- Table A9. Electron microprobe analyses of hornblende.
- Table A10. Electron microprobe analyses of apatite.
- Table A11. Electron microprobe analyses of sulfide globules.
- Table A12. Pyroxene geothermometry.
- Table A13. Oxygen thermobarometry results from Fe–Ti oxide pairs.
- Table A14. Electron microprobe analyses of glass inclusions.
- Table A15. Electron microprobe analyses of matrix glasses.
- Table A16. Compositions of olivine-hosted glass inclusions and host olivines.
- Table A17. Compositions of orthopyroxene-hosted glass inclusions and host orthopyroxenes.
- Table A18. Compositions of clinopyroxene-hosted glass inclusions and host clinopyroxenes.
- Table A19. Composition of hornblende-hosted glass inclusion.
- Table A20. Electron microprobe analyses of matrix glasses.
- Table A21. Popocatépetl whole-rock analyses.
- Table A2. Relative errors based on counting statistics for electron microprobe analyses of minerals and glasses.
Fig. A1. Pyroxene core compositions shown on plots of (a) SiO2, (b) Al2O3, and (c) Cr2O3 vs Mg-number [= Mg/(Mg + Fe2+)]. Ortho- and clinopyroxenes that are inferred to derive from the mafic and silicic end-members are grouped. Pyroxenes of intermediate composition and xenocrysts are labeled. Pyroxene rim compositions not shown for clarity.
Fig. A2. Hornblende compositions shown on a plot of Mg-number [= Mg/(Mg + FeT), where FeT is total iron] vs Al2O3. Hornblende data from this study (x) fall within the field for groundmass pargasite in dacite from the 1991–1995 eruption of Unzen volcano, Japan (Sato et al., 1999) and the field for groundmass magnesiohastingsite in andesite from the continuing eruption of Soufrière Hills volcano, Montserrat (Murphy et al., 2000). Hornblende from this study falls outside the field for amphibole (magnesiohornblende) phenocrysts found in rocks from both the 1991–1995 eruption of Unzen and the current eruption of Soufrière Hills (Sato et al., 1999; Murphy et al., 2000).
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ACKNOWLEDGEMENTS |
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We would like to thank J.-L. Macías and S. De la Cruz-Reyna (Instituto de Geofísica, UNAM) and E. Ramos (CENAPRED) for kindly providing us with many of the rock samples used in this study. We also thank S. De la Cruz-Reyna for generously providing data on magma volumes and eruption chronology for Popocatépetl. Special thanks go to S. Kuehner for his patient assistance in the UW electron microprobe lab. I appreciate stimulating discussions on conduit convection with Y. Taran, S. Shinohara, K. Kazahaya, G. Bergantz, and R. Breidenthal. Reviews by M. Coombs, J. Lowenstern, C. Bacon, T. Sisson, W. Rose, P. Wallace, W. Leeman, P. Gauthier, B. R. Frost, and two anonymous reviewers significantly improved the manuscript. This investigation was financially supported by NSF Grant EAR-9980566 to V.C.K. and C.G.N. and a NASA ESS Graduate Fellowship to J.B.W.
* Corresponding author. E-mail: witt_98103{at}yahoo.com
|
REFERENCES |
|---|
|
TOPABSTRACTINTRODUCTIONANALYTICAL METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAAPPENDIXREFERENCESREFERENCES |
|---|
Allard, P. (1997). Endogenous magma degassing and storage at Mount Etna. Geophysical Research Letters 24(17), 2219–2222.[CrossRef][Web of Science]
Andersen, D. J., Lindsley, D. H. & Davidson, P. M. (1993). QUILF: a Pascal program to assess equilibria among Fe–Mg–Mn–Ti oxides, pyroxenes, olivine, and quartz. Computers and Geosciences 19(9), 1333–1350.
Anderson, A. T. (1974). The before-eruption water contents of some high-alumina magmas. Bulletin of Volcanology 37(4), 530–552.
Anderson, A. T., Jr, Newman, S., Williams, S. N., Druitt, T. H., Skirius, C. & Stolper, E. (1989). H2O, CO2, Cl, and gas in Plinian and ash-flow Bishop rhyolite. Geology 17(3), 221–225.
Andres, R. J., Rose, W. I., Kyle, P. R., deSilva, S., Francis, P., Gardeweg, M. & Moreno Roa, H. (1991). Excessive sulfur dioxide emissions from Chilean volcanoes. Journal of Volcanology and Geothermal Research 46(3–4), 323–329.[CrossRef][Web of Science]
Armstrong, J. T. (1984). Quantitative analysis of silicate and oxide minerals: a reevaluation of ZAF corrections and proposal for new Bence–Albee coefficients. In: Romig, A. D., Jr & Goldstein, J. I. (eds) Microbeam Analysis. San Francisco, CA: San Francisco Press, pp. 208–212.
Athanasopoulos, P. (1997). The origin and ascent history of the 1996 dacitic dome, Volcán Popocatépetl, Mexico. B.Sc. thesis, University of Manitoba, Winnipeg.
Bacon, C. R. & Hirschmann, M. M. (1988). Mg/Mn partitioning as a test for equilibrium between coexisting Fe–Ti oxides. American Mineralogist 73(1–2), 57–61.[Abstract]
Boudal, C. & Robin, C. (1988). Relations entre dynamismes éruptifs et réalimentations magmatiques d'origine profonde au Popocatépetl. Canadian Journal of Earth Sciences 25, 955–971.[Web of Science]
Boudal, C. & Robin, C. (1989). Volcán Popocatépetl: Recent eruptive history and potential hazards and risks in future eruptions. In: Latter, J. H. (ed.) Volcanic Hazards, Vol. 1. Berlin: Springer, pp. 110–128.
Cabral-Cano, E., Meertens, C. & Correa-Mora, F. (1999). Surface deformation on active volcanoes in central Mexico. (Abstract.) American Geophysical Union Fall Meeting, December 13–17, San Francisco, CA.
Cantagrel, J. M., Gourgaud, A. & Robin, C. (1984). Repetitive mixing events and Holocene pyroclastic activity at Pico de Orizaba and Popocatépetl (Mexico). Bulletin of Volcanology 47–4(1), 735–748.
Carroll, M. R. & Rutherford, M. J. (1985). Sulfide and sulfate saturation in hydrous silicate melts. Journal of Geophysical Research 90(Supplement), C601–C612.
Carroll, M. R. & Rutherford, M. J. (1987). The stability of igneous anhydrite: experimental results and implications for sulfur behavior in the 1982 El Chichon trachyandesite and other evolved magmas. Journal of Petrology 28, 781–801.
CENAPRED-UNAM (1995). Volcán Popocatépetl: estudios realizados durante la crisis de 1994–1995. Mexico City: Centro Nacional de Prevencion de Desastres.
Clynne, M. A. & Borg, L. E. (1997). Olivine and chromian spinel in primitive calc-alkaline and tholeiitic lavas from the southernmost Cascade Range, California; a reflection of relative fertility of the source. Canadian Mineralogist 35, 453–472.[Web of Science]
Danyushevsky, L. V. (2001). The effect of small amounts of H2O on crystallization of mid-ocean ridge and backarc basin magmas. Journal of Volcanology and Geothermal Research 110(3–4), 265–280.[CrossRef][Web of Science]
De la Cruz-Reyna, S. & Siebe, C. (1997). The giant Popocatépetl stirs. Nature 388, 227.[CrossRef]
De la Cruz-Reyna, S., Quezada, J. L., Peña, C., Zepeda, O. & Sánchez, T. (eds) (1995). Historia de la Actividad Reciente del Popocatépetl (1354–1995). Volcán Popocatépetl estudios realizados durante la crisis de 1994–1995. Mexico City: CENAPRED–UNAM, pp. 3–22.
Delgado, H. & Cárdenas-G, L. (1997). Passive degassing at Volcán Popocatépetl (México): 2·6 x 106 tons of SO2 released in 617 days of activity. (Abstract.) IAVCEI General Assembly, Puerto Vallarta, Mexico.
Delgado, H., Piedad-Sánchez, N., Galvan, L., Julio, P., Alvarez, J. M. & Cárdenas, L. (1998). CO2 flux measurements at Popocatépetl volcano: II. Magnitude of emissions and significance. (Abstract.) American Geophysical Union Fall Meeting, December 6–10, San Francisco, CA, F926.
Delgado-Granados, H., Cárdenas-Gonzáles, L. & Piedad-Sánchez, N. (2001). Sulfur dioxide emissions from Popocatépetl volcano (Mexico): case study of a high-emission rate, passively degassing erupting volcano. Journal of Volcanology and Geothermal Research 108(1–4), 107–120.[CrossRef][Web of Science]
Devine, J. D., Gardner, J. E., Brack, H. P., Layne, G. D. & Rutherford, M. J. (1995). Comparison of microanalytical methods for estimating H2O contents of silicic volcanic glasses. American Mineralogist 80, 319–328.[Abstract]
Eichelberger, J. C., Carrigan, C. R., Westrich, H. R. & Price, R. H. (1986). Non-explosive silicic volcanism. Nature 323(6089), 598–602.[CrossRef]
Evans, B. W., Scaillet, B. & Kuehner, S. M. (2001). Iron–titanium oxide thermobarometry at high oxygen fugacity. (Abstract.) Geological Society of America Meeting, November 1–10, Boston, MA.
Gaetani, G. A. & Grove, T. L. (1997). Partitioning of moderately siderophile elements among olivine, silicate melt, and sulfide melt: constraints on core formation in the Earth and Mars. Geochimica et Cosmochimica Acta 61, 1829–1846.[CrossRef][Web of Science]
Garcia, M. O. & Jacobson, S. S. (1979). Crystal clots, amphibole fractionation, and the evolution of calc-alkaline magmas. Contributions to Mineralogy and Petrology 69, 319–327.[CrossRef][Web of Science]
Gerlach, T. M. & McGee, K. A. (1994). Total sulfur dioxide emissions and pre-eruption vapor-saturated magma at Mount St. Helens. Geophysical Research Letters 21(25), 2833–2836.[CrossRef][Web of Science]
Gerlach, T. M., Westrich, H. R., Casadevall, T. J., Finnegan, D. L., Miller, T. P. & Chouet, B. A. (1994). Vapor saturation and accumulation in magmas of the 1989–1990 eruption of Redoubt Volcano, Alaska. Journal of Volcanology and Geothermal Research 62(1–4), 317–337.[CrossRef][Web of Science]
Gerlach, T. M., Westrich, H. R. & Symonds, R. B. (1996). Preeruption vapor in magma of the climactic Mount Pinatubo eruption: source of the giant stratospheric sulfur dioxide cloud. In: Newhall, C. G. & Punongbayan, R. S. (eds) Fire and Mud: Eruptions and Lahars of Mount Pinatubo, Philippines. Quezon City: PHIVOLCS; Seattle, WA: University of Washington Press, pp. 415–433.
Gerlach, T. M., Delgado, H., McGee, K. A., Doukas, M. P., Venegas, J. J. & Cardenas, L. (1997). Application of the LI-COR CO2 analyzer to volcanic plumes; a case study, Volcan Popocatépetl, Mexico, June 7 and 10, 1995. Journal of Geophysical Research 102(4), 8005–8019.[CrossRef]
Ghiorso, M. S. (1999). On the stability relations of hydrous minerals in water-undersaturated magmas. American Mineralogist 84, 1506–1511.[Abstract]
Ghiorso, M. S. & Sack, R. O. (1991). Fe–Ti oxide geothermometry; thermodynamic formulation and the estimation of intensive variables in silicic magmas. Contributions to Mineralogy and Petrology 108(4), 485–510.[CrossRef][Web of Science]
Ghiorso, M. S. & Sack, R. O. (1995). Chemical mass transfer in magmatic processes. IV. A revised and internally consistent thermodynamic model for the interpolation and extrapolation of liquid–solid equilibria in magmatic systems at elevated temperatures and pressures. Contributions to Mineralogy and Petrology 119, 197–212.[Web of Science]
Goff, F., Janik, C. J., Delgado, H., Werner, C., Counce, D., Stimac, J. A., Siebe, C., Love, S. P., Williams, S. N., Fischer, T. & Johnson, L. (1998). Geochemical surveillance of magmatic volatiles at Popocatépetl Volcano, Mexico. Geological Society of America Bulletin 110(6), 695–710.
Goff, F., Love, S. P., Warren, R. G., Counce, D., Obenholzner, J., Siebe, C. & Schmidt, S. C. (2001). Passive infrared remote sensing evidence for large, intermittent CO2 emissions at Popocatépetl volcano, Mexico. Chemical Geology 177, 133–156.[CrossRef][Web of Science]
GVN (1996). Popocatépetl (México): continued ash emissions; new lava dome and lava flows in summit crater. Global Volcanism Network Bulletin 21(3).
Hauri, E., Wang, J., Dixon, J. E., King, P. L., Mandeville, C. & Newman, S. (2002). SIMS analysis of volatiles in silicate glasses 1. Calibration, matrix effects and comparisons with FTIR. Chemical Geology 183, 99–114.[CrossRef][Web of Science]
Hess, K.-U. & Dingwell, D. B. (1996). Viscosities of hydrous leucogranitic melts: a non-Arrhenian model. American Mineralogist 81, 1297–1300.[Web of Science]
Holland, T. & Blundy, J. (1994). Non-ideal interactions in calcic amphiboles and their bearing on amphibole–plagioclase thermometry. Contributions to Mineralogy and Petrology 116, 433–447.
Jaupart, C. & Allègre, C. J. (1991). Gas content, eruption rate and instabilities of eruption regime in silicic volcanoes. Earth and Planetary Science Letters 102(3–4), 413–429.[CrossRef][Web of Science]
Kazahaya, K., Shinohara, H. & Saito, G. (1994). Excessive degassing of Izu–Oshima volcano; magma convection in a conduit. Bulletin of Volcanology 56(3), 207–216.[CrossRef][Web of Science]
Kazahaya, K., Shinohara, H. & Saito, G. (2002). Degassing process of Satsuma–Iwojima volcano, Japan: supply of volatile components from a deep magma chamber. Earth, Planets, Space 54, 327–335.
Klug, C. & Cashman, K. V. (1996). Permeability development in vesiculating magmas; implications for fragmentation. Bulletin of Volcanology 58(2–3), 87–100.[CrossRef][Web of Science]
Kolisnik, A. M. E. (1990). Phenocryst zoning and heterogeneity in andesites and dacites of Volcán Popocatépetl, Mexico. M.Sc. thesis, Queen's University, Kingston, Ont.
Koyaguchi, T. & Blake, S. (1989). The dynamics of magma mixing in a rising magma batch. Bulletin of Volcanology 52, 127–137.[CrossRef][Web of Science]
Kress, V. C. & Ghiorso, M. S. (2004). Thermodynamic modeling of post-entrapment crystallization in igneous phases. Journal of Volcanology and Geothermal Research 137(4), 247–260.[CrossRef][Web of Science]
Lange, R. A. (1994). The effect of H2O, CO2 and F on the density and viscosity of silicate melts. In: Carroll, M. R. & Holloway, J. R. (eds) Volatiles in Magmas. Mineralogical Society of America, Reviews in Mineralogy 30, 331–369.
Larocque, A. C. L., Stimac, J. A., Keith, J. D. & Huminicki, M. A. E. (2000). Evidence for open-system behavior in immiscible Fe–S–O liquids in silicate magmas: implications for contributions of metals and sulfur to ore-forming fluids. Canadian Mineralogist 38, 1233–1249.[CrossRef][Web of Science]
Lattard, D., Käsmann, M., Fleischhammer, M. & Höltzen, H. (2001). Towards a better calibration of the titanomagnetite–ilmenite thermo-oxybarometer: new experimental results in the Fe–Ti–O system at 1 bar and 1000–1300°C. (Abstract.) European Union of Geosciences XI Meeting, p. 585.
Le Bas, M. J., LeMaitre, R. W., Streckeisen, A. L. & Zanettin, B. (1986). A chemical classification of volcanic rocks based on the total alkali–silica diagram. Journal of Petrology 27, 745–750.
Lejeune, A.-M. & Richet, P. (1995). Rheology of crystal-bearing silicate melts; an experimental study at high viscosities. Journal of Geophysical Research 100(3), 4215–4229.[CrossRef]
Lindsley, D. H. & Andersen, D. J. (1983). A two pyroxene thermometer. Journal of Geophysical Research 88(Supplement), A887–A906.[CrossRef]
Love, S. P., Goff, F., Counce, D., Siebe, C. & Delgado, H. (1998). Passive infrared spectroscopy of the eruption plume at Popocatépetl volcano, Mexico. Nature 396, 563–567.[CrossRef]
Lowenstern, J. B. (1993). Evidence for a copper-bearing fluid in magma erupted at the Valley of Ten Thousand Smokes, Alaska. Contributions to Mineralogy and Petrology 114(3), 409–421.[CrossRef][Web of Science]
Lowenstern, J. B., Mahood, G. A., Rivers, M. L. & Sutton, S. R. (1991). Evidence for extreme partitioning of copper into a magmatic vapor phase. Science 252, 1405–1409.
Luhr, J. F. (1990). Experimental phase relations of water- and sulfur-saturated arc magmas and the 1982 eruptions of El Chichon volcano. Journal of Petrology 31(5), 1071–1114.
Luhr, J. F., Carmichael, I. S. E. & Varekamp, J. C. (1984). The 1982 eruptions of El Chichon volcano, Chiapas, Mexico: mineralogy and petrology of the anhydrite-bearing pumices. Journal of Volcanology and Geothermal Research 23, 69–108.[CrossRef][Web of Science]
Mandeville, C. W., Carey, S. & Sigurdsson, H. (1996). Magma mixing, fractional crystallization and volatile degassing during the 1883 eruption of Krakatau Volcano, Indonesia. Journal of Volcanology and Geothermal Research 74(3–4), 243–274.[CrossRef][Web of Science]
McKibben, M. A., Eldridge, C. S. & Reyes, A. G. (1996). Sulfur isotopic systematics of the June 1991 Mount Pinatubo eruptions: a SHRIMP ion microprobe study. In: Newhall, C. G. & Punongbayan, R. S. (eds) Fire and Mud: Eruptions and Lahars of Mount Pinatubo, Philippines. Quezon City: PHIVOLCS; Seattle, WA: University of Washington Press, pp. 825–843.
Moore, G., Vennemann, T. & Carmichael, I. S. E. (1998). An empirical model for the solubility of H2O in magmas to 3 kilobars. American Mineralogist 83, 36–42.[Abstract]
Murphy, M. D., Sparks, R. S. J., Barclay, J., Carroll, M. R. & Brewer, T. S. (2000). Remobilization of andesite magma by intrusion of mafic magma at the Soufrière Hills Volcano, Montserrat, West Indies. Journal of Petrology 41(1), 21–42.
Nakada, S. & Motomura, Y. (1999). Petrology of the 1991–1995 eruption at Unzen: effusion pulsation and groundmass crystallization. Journal of Volcanology and Geothermal Research 89(1–4), 173–196.[CrossRef][Web of Science]
Newman, S. & Lowenstern, J. B. (2002). VolatileCalc: a silicate melt–H2O–CO2 solution model written in Visual Basic for Excel. Computers and Geosciences 28, 597–604.[CrossRef]
Newman, S., Epstein, S. & Stolper, E. (1988). Water, carbon dioxide, and hydrogen isotopes in glasses from the ca. 1340 A.D. eruption of the Mono Craters, California; constraints on degassing phenomena and initial volatile content. Journal of Volcanology and Geothermal Research 35(1–2), 75–96.[CrossRef][Web of Science]
Ohmoto, H. & Rye, R. O. (1979). Isotopes of sulfur and carbon. In: Barnes, H. L. (ed.) Geochemistry of Hydrothermal Ore Deposits. New York: John Wiley, pp. 509–567.
Okada, Y. (1992). Internal deformation due to shear and tensile faults in a half-space. Bulletin of the Seismological Society of America 82, 1018–1049.
Panfil, M. S., Gardner, T. W. & Hirth, K. G. (1999). Late Holocene stratigraphy of the Tetimpa archaeological sites, northeast flank of Popocatépetl Volcano, central Mexico. Geological Society of America Bulletin 111(2), 204–218.
Piccoli, P. & Candela, P. (1994). Apatite in felsic rocks: a model for the estimation of initial halogen concentrations in the Bishop Tuff (Long Valley) and Tuolomne Intrusive Suite (Sierra Nevada Batholith) magmas. American Journal of Science 294, 92–135.
Richet, P., Lejeune, A. M., Holtz, F. & Roux, J. (1996). Water and the viscosity of andesite melts. Chemical Geology 128(1–4), 185–197.[CrossRef][Web of Science]
Robin, C. (1984). Le Volcan Popocatépetl (Mexique): structure, evolution pétrologique et risques. Bulletin Volcanologique 47(1), 1–23.[CrossRef]
Robin, C. & Boudal, C. (1987). A gigantic Bezymianny-type event at the beginning of modern Volcán Popocatépetl. Journal of Volcanology and Geothermal Research 31, 115–130.[CrossRef][Web of Science]
Rose, W. I., Jr, Stoiber, R. E. & Malinconico, L. L. (1982). Eruptive gas compositions and fluxes of explosive volcanoes; budget of S and Cl emitted from Fuego Volcano, Guatemala. In: Thorpe, R. S. (ed.) Andesites; Orogenic Andesites and Related Rocks. New York: John Wiley, pp. 669–676.
Rutherford, M. J. & Hill, P. M. (1993). Magma ascent rates from amphibole breakdown; an experimental study applied to the 1980–1986 Mount St. Helens eruptions. Journal of Geophysical Research, B, Solid Earth and Planets 98(11), 19667–19685.[CrossRef]
Sato, H., Nakada, S., Fujii, T., Nakamura, M. & Suzuki, K. (1999). Groundmass pargasite in the 1991–1995 dacite of Unzen volcano: phase stability experiments and volcanological implications. Journal of Volcanology and Geothermal Research 89(1–4), 197–212.[CrossRef][Web of Science]
Scaillet, B., Clemente, B., Evans, B. W. & Pichavant, M. (1998). Redox control of sulfur degassing in silicic magmas. Journal of Geophysical Research, Solid Earth 103(B10), 23937–23949.[CrossRef][Web of Science]
Seifert, S., O'Neill, H. S. C. & Brey, G. (1988). The partitioning of Fe, Ni, and Co between olivine, metal, and basaltic liquid: an experimental and thermodynamic investigation, with application to the composition of the lunar core. Geochimica et Cosmochimica Acta 52, 603–616.[CrossRef][Web of Science]
Shinohara, H., Kazahaya, K. & Lowenstern, J. B. (1995). Volatile transport in a convecting magma column; implications for porphyry Mo mineralization. Geology 23(12), 1091–1094.
Siebe, C., Abrams, M., Macías, J. L. & Obenholzner, J. (1996a). Repeated volcanic disasters in Prehispanic time at Popocatépetl, central Mexico: past key to the future? Geology 24(5), 399–402.
Siebe, C., Macías Vázquez, J. L., Abrams, M. & Obenholzner, J. (1996b). La destrucción de Cacaxtla y Cholula: un suceso en la historia eruptiva del Popocatépetl. Ciencias 41, 36–45.
Sigurdsson, H., Carey, S., Palais, J. M., Devine, J. & Williams, S. N. (1990). Pre-eruption compositional gradients and mixing of andesite and dacite magma erupted from Nevado del Ruiz Volcano, Colombia in 1985. Journal of Volcanology and Geothermal Research 41(1–4), 127–151.[CrossRef][Web of Science]
Simakin, A. & Botcharnikov, R. (2001). Degassing of stratified magma by compositional convection. Journal of Volcanology and Geothermal Research 105, 207–224.[CrossRef][Web of Science]
Snyder, D. A. & Carmichael, I. S. E. (1992). Olivine–liquid equilibria and the chemical activities of FeO, NiO, Fe2O3, and MgO in natural basic melts. Geochimica et Cosmochimica Acta 56, 303–318.[CrossRef][Web of Science]
Stein, D. & Spera, F. (2002). Shear viscosity of rhyolite–vapor emulsions at magmatic temperatures by concentric cylinder rheometry. Journal of Volcanology and Geothermal Research 113, 243–258.[CrossRef][Web of Science]
Stevenson, D. S. & Blake, S. (1998). Modelling the dynamics and thermodynamics of volcanic degassing. Bulletin of Volcanology 60(4), 307–317.[CrossRef][Web of Science]
Stimac, J. A., Siebe, C. & Schaaf, P. (1997). Magmatic processes at Popocatépetl, Mexico. (Abstract.) IAVCEI General Assembly, Puerto Vallarta, Mexico.
Straub, S. M. & Martin-Del Pozzo, A. L. (2001). The significance of phenocryst diversity in tephra from recent eruptions at Popocatépetl volcano (central Mexico). Contributions to Mineralogy and Petrology 140, 487–510.[CrossRef][Web of Science]
Swanson, D. A. & Holcomb, R. T. (1990). Regularities in growth of the Mount St. Helens dacite dome 1980–1986. In: Fink, J. H. (ed.) Lava Flows and Domes; Emplacement Mechanisms and Hazard Implications, Vol. 2. Berlin: Springer, pp. 3–23.
Taylor, B. E. (1986). Magmatic volatiles: isotopic variation of C, H, and S. In: Valley, J. W., Taylor, H. P. & O'Neil, J. R. (eds) Stable Isotopes in High Temperature and Geological Processes. Mineralogical Society of America, Reviews in Mineralogy 16, 185–225.
Thordarson, T., Self, S., Oskarsson, N. & Hulsebosch, T. (1996). Sulfur, chlorine, and fluorine degassing and atmospheric loading by the 1783–1784 AD Laki (Skaftar Fires) eruption in Iceland. Bulletin of Volcanology 58(2–3), 205–225.[CrossRef][Web of Science]
Wallace, P. J. (2001). Volcanic SO2 emissions and the abundance and distribution of exsolved gas in magma bodies. Journal of Volcanology and Geothermal Research 108, 85–106.[CrossRef][Web of Science]
Wallace, P. J. & Gerlach, T. M. (1994). Magmatic vapor source for sulfur dioxide released during volcanic eruptions; evidence from Mount Pinatubo. Science 265(5171), 497–499.
Wallace, P. J., Anderson, A. T., Jr & Davis, A. M. (1995). Quantification of pre-eruptive exsolved gas contents in silicic magmas. Nature 377(6550), 612–616.[CrossRef]
Webster, J., Kinzler, R. & Mathez, E. (1999). Chloride and water solubility in basalt and andesite melts and implications for magmatic degassing. Geochimica et Cosmochimica Acta 63(5), 729–738.[CrossRef][Web of Science]
Wendlandt, R. (1982). Sulfide saturation of basalt and andesite melts at high pressures and temperatures. American Mineralogist 67(1–2), 170–174.[Abstract]
Wessel, P. & Smith, W. H. F. (1991). Free software helps map and display data. EOS Transactions, American Geophysical Union 72, 441.
Westrich, H. R. & Gerlach, T. M. (1992). Magmatic gas source for the stratospheric SO2 cloud from the June 15, 1991, eruption of Mount Pinatubo. Geology 20(10), 867–870.
Wilcox, R. E. (1954). Petrology of Paricutin Volcano, Mexico. US Geological Survey Bulletin 965-C, 281–349.
Witter, J. B. (1997). Volatile emissions and potential climatic impact of the Great Kuwae (Vanuatu) eruption of 1452–3 A.D. M.Sc. thesis, University of Hawaii at Manoa, Honolulu.
Witter, J. B. (2003). Convection of magma in volcanic conduits as a degassing mechanism at active volcanoes. Ph.D. thesis, University of Washington, Seattle.
Witter, J. B. & Kuehner, S. M. (2004). A simple empirical method for high-quality electron microprobe analysis of fluorine at trace levels in Fe-bearing minerals and glasses. American Mineralogist 89, 57–63.
|
REFERENCES |
|---|
|
TOPABSTRACTINTRODUCTIONANALYTICAL METHODSRESULTSDISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAAPPENDIXREFERENCESREFERENCES |
|---|
Allan, J. F., Sack, R. O. & Batiza, R. (1989). Cr-rich spinels as petrogenetic indicators; MORB-type lavas from the Lamont seamount chain, eastern Pacific. American Mineralogist 73(7–8), 741–753.[Web of Science]
Andersen, D. J., Lindsley, D. H. & Davidson, P. M. (1993). QUILF: a Pascal program to assess equilibria among Fe–Mg–Mn–Ti oxides, pyroxenes, olivine, and quartz. Computers and Geosciences 19(9), 1333–1350.
Athanasopoulos, P. (1997). The origin and ascent history of the 1996 dacitic dome, Volcán Popocatépetl, Mexico. B.Sc. thesis, University of Manitoba, Winnipeg.
Bacon, C. R. & Hirschmann, M. M. (1988). Mg/Mn partitioning as a test for equilibrium between coexisting Fe–Ti oxides. American Mineralogist 73(1–2), 57–61.[Abstract]
Carroll, M. R. & Rutherford, M. J. (1988). Sulfur speciation in hydrous experimental glasses of varying oxidation state; results from measured wavelength shifts of sulfur X-rays. American Mineralogist 73(7–8), 845–849.[Abstract]
Clynne, M. A. & Borg, L. E. (1997). Olivine and chromian spinel in primitive calc-alkaline and tholeiitic lavas from the southernmost Cascade Range, California; a reflection of relative fertility of the source. Canadian Mineralogist 35, 453–472.
Dixon, J. E., Clague, D. A. & Stolper, E. M. (1991). Degassing history of water, sulfur, and carbon in submarine lavas from Kilauea Volcano, Hawaii. Journal of Geology 99(3), 371–394.[Web of Science]
Frost, B. R. (1991). Introduction to oxygen fugacity and its petrologic importance. In: Lindsley, D. H. (ed.) Oxide Minerals: Petrologic and Magnetic Significance. Mineralogical Society of America, Reviews in Mineralogy 25, 1–9.
Garcia, M. O. & Jacobson, S. S. (1979). Crystal clots, amphibole fractionation, and the evolution of calc-alkaline magmas. Contributions to Mineralogy and Petrology 69, 319–327.[CrossRef][Web of Science]
Ghiorso, M. S. & Sack, R. O. (1991). Fe–Ti oxide geothermometry; thermodynamic formulation and the estimation of intensive variables in silicic magmas. Contributions to Mineralogy and Petrology 108(4), 485–510.[CrossRef][Web of Science]
GVN (1997a). Largest ash emission of the 1994–97 eruption on 30 June. Global Volcanism Network Bulletin 22(7).
GVN (1997b). Explosion on 1 January causes fires, ashfall. Global Volcanism Network Bulletin 22(12).
Holland, T. & Blundy, J. (1994). Non-ideal interactions in calcic amphiboles and their bearing on amphibole–plagioclase thermometry. Contributions to Mineralogy and Petrology 116, 433–447.
Kolisnik, A. M. E. (1990). Phenocryst zoning and heterogeneity in andesites and dacites of Volcán Popocatépetl, Mexico. M.Sc. thesis, Queen's University, Kingston, Ont.
Kress, V. C. & Ghiorso, M. S. (2004). Thermodynamic modeling of post-entrapment crystallization in igneous phases. Journal of Volcanology and Geothermal Research 137(4), 247–260.[CrossRef][Web of Science]
Leake, B. E., Woolley, A. R., Arps, C. E. S., et al. (1997). Nomenclature of amphiboles: Report of the Subcommittee on Amphiboles of the International Mineralogical Association, Commission on New Minerals and Mineral Names. American Mineralogist 82(9–10), 1019–1037.[Abstract]
Luhr, J. F. & Carmichael, I. S. E. (1985). Jorullo Volcano, Michoacan, Mexico (1759–1774); the earliest stages of fractionation in calc-alkaline magmas. Contributions to Mineralogy and Petrology 90(2–3), 142–161.[CrossRef][Web of Science]
Métrich, N. & Clocchiatti, R. (1996). Sulfur abundance and its speciation in oxidized alkaline melts. Geochimica et Cosmochimica Acta 60(21), 4151–4160.[CrossRef][Web of Science]
Murphy, M. D., Sparks, R. S. J., Barclay, J., Carroll, M. R. & Brewer, T. S. (2000). Remobilization of andesite magma by intrusion of mafic magma at the Soufrière Hills Volcano, Montserrat, West Indies. Journal of Petrology 41(1), 21–42.
Paris, E., Giuli, G., Carroll, M. R. & Davoli, I. (2001). The valence and speciation of sulfur in glasses by X-ray absorption spectroscopy. Canadian Mineralogist 39, 331–339.[CrossRef][Web of Science]
Peng, G., Luhr, J. F. & McGee, J. J. (1997). Factors controlling sulfur concentrations in volcanic apatite. American Mineralogist 82, 1210–1224.[Abstract]
Piccoli, P. & Candela, P. (1994). Apatite in felsic rocks: a model for the estimation of initial halogen concentrations in the Bishop Tuff (Long Valley) and Tuolomne Intrusive Suite (Sierra Nevada Batholith) magmas. American Journal of Science 294, 92–135.
Saito, G., Kazahaya, K., Shinohara, H., Stimac, J. & Kawanabe, Y. (2001). Variation of volatile concentration in a magma system of Satsuma–Iwojima volcano deduced from melt inclusion analyses. Journal of Volcanology and Geothermal Research 108(1–4), 11–31.[CrossRef][Web of Science]
Sato, H., Nakada, S., Fujii, T., Nakamura, M. & Suzuki, K. (1999). Groundmass pargasite in the 1991–1995 dacite of Unzen volcano: phase stability experiments and volcanological implications. Journal of Volcanology and Geothermal Research 89(1–4), 197–212.[CrossRef][Web of Science]
Stormer, J. C., Jr (1983). The effects of recalculation on estimates of temperature and oxygen fugacity from analyses of multicomponent iron–titanium oxides. American Mineralogist 68(5–6), 586–594.[Abstract]
Straub, S. M. & Martin-Del Pozzo, A. L. (2001). The significance of phenocryst diversity in tephra from recent eruptions at Popocatépetl volcano (central Mexico). Contributions to Mineralogy and Petrology 140, 487–510.[CrossRef][Web of Science]
Thordarson, T., Self, S., Oskarsson, N. & Hulsebosch, T. (1996). Sulfur, chlorine, and fluorine degassing and atmospheric loading by the 1783–1784 AD Laki (Skaftar Fires) eruption in Iceland. Bulletin of Volcanology 58(2–3), 205–225.[CrossRef][Web of Science]
Venezky, D. Y. & Rutherford, M. J. (1997). Preeruption conditions and timing of dacite–andesite magma mixing in the 2·2 ka eruption at Mount Rainier. Journal of Geophysical Research 102(B9), 20069–20086.[CrossRef][Web of Science]
Venezky, D. Y. & Rutherford, M. J. (1999). Petrology and Fe–Ti oxide reequilibration of the 1991 Mount Unzen mixed magma. Journal of Volcanology and Geothermal Research 89(1–4), 213–230.[CrossRef][Web of Science]
Wallace, P. J. & Carmichael, I. S. E. (1994). S speciation in submarine basaltic glasses as determined by measurements of S K-alpha X-ray wavelength shifts. American Mineralogist 79, 161–167.[Abstract]
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