factors controlling sulfur concentrations in volcanic apatite · factors controlling sulfur...

12100003–004X/97/1112–1210$05.00

American Mineralogist, Volume 82, pages 1210–1224, 1997

Factors controlling sulfur concentrations in volcanic apatite

GENYONG PENG,1 JAMES F. LUHR,1 AND JAMES J. MCGEE21Department of Mineral Sciences, NHB-119, Smithsonian Institution, Washington, DC 20560, U.S.A.

2U.S. Geological Survey, Reston, Virginia 22092, U.S.A.

ABSTRACTApatite crystals from two types of samples were analyzed by electron microprobe for

15 major and trace elements: (1) apatite in H2O- and S-saturated experimental charges ofthe 1982 El Chichon trachyandesite and (2) apatite in volcanic rocks erupted from 20volcanoes. The SO3 contents of the experimental apatite increase with increasing oxygenfugacity (ƒ ), from #0.04 wt% in reduced charges buffered by fayalite-magnetite-quartzO2(FMQ), to 1.0–2.6 wt% in oxidized charges buffered by manganosite-hausmanite (MNH)or magnetite-hematite (MTH). The SO3 contents of MNH- and MTH-buffered apatite alsogenerally increase with increasing pressure from 2 to 4 kbar and decreasing temperaturefrom 950 to 800 8C. The partition coefficient for SO3 between apatite and oxidized meltincreases with decreasing temperature but appears to be independent of pressure.Apatites in volcanic rocks show a wide range of SO3 contents (,0.04 to 0.63 wt%).

Our sample set includes one group known to contain primary anhydrite and a secondgroup inferred to have been free of primary anhydrite. No systematic differences in apatiteS contents are observed between these two groups. Our study was initiated to define thefactors controlling S contents in apatite and to evaluate the hypothesis that high S contentsin apatite could be characteristic of S-rich anhydrite-bearing magmas such as those eruptedfrom El Chichon in 1982 and Pinatubo in 1991. This hypothesis is shown to be invalid,probably chiefly a consequence of the slow intra-crystalline diffusion that limits re-equil-ibration between early formed apatite and the evolving silicate melt. Contributing factorsinclude early crystallization of most apatite over a relatively small temperature interval,common late-stage magmatic enrichment of S, progressive oxidation during magmaticevolution, and strong controls on S contents in apatite exerted by ƒ , temperature, andO2pressure.

TABLE 1. Evaluation of analytical accuracy and precision based on analyses of Durango apatite (wt%)

SiO2 Al2O3 La2O3 Ce2O3 FeO MnO MgO CaO

Recommended values*Average of 113 analyses (SI)Average of 38 analyses (USGS)Detectability limit (SI)Sensitivity (SI)

0.340.33(4)0.36(4)0.05

0.070.00(0)0.00(1)0.01

0.490.44(4)0.44(7)0.03

0.550.59(5)0.53(12)0.03

0.050.02(1)0.04(4)0.01

0.010.02(2)0.01(2)0.01

0.010.00(1)0.00(0)0.01

54.0254.02(28)54.02(41)

0.14

Note: Numbers in parentheses are one standard deviation values.* From Young et al. (1969).

INTRODUCTIONRecent eruptions of anhydrite-bearing tephra at El Chi-

chon (1982) and Pinatubo (1991) volcanoes were accom-panied by large injections of S gases into the atmosphere:7 and 20 megatons of SO2, respectively (Bluth et al.1992). These observations stimulated research into thepossible relationship between S-rich volcanic eruptionsand short-term global climatic changes (Luhr 1991; Bluthet al. 1992, 1993; Hansen et al. 1992). Anhydrite-satu-rated magmas of the type erupted at El Chichon and Pin-atubo appear to have the following common features:highly-oxidized, crystal-rich, fluid-saturated, rich in

amphibole or biotite, and erupted in subduction-zones(Luhr et al. 1984; Luhr 1991; Imai et al. 1993; Pallisteret al. 1995, 1996; Luhr and Melson 1996). Primary vol-canic anhydrite is readily erased from the geological rec-ord because it is highly soluble in water. Thus, S-richanhydrite-bearing magmas may be much more commonthan presently realized (Luhr et al. 1984). One approachto the estimation of volatiles released to the atmosphereduring volcanic eruptions is the ‘‘petrologic method’’ ofDevine et al. (1984), which bases the estimate on themass of erupted material and the difference in S contentsbetween glass inclusions and matrix glass. This method,

1211PENG ET AL.: S IN APATITE

TABLE 1—Extended.

SrO Na2O K2O P2O5 SO3 Cl F

0.070.06(4)0.06(3)0.03

0.230.24(3)0.22(3)0.02

0.010.00(1)0.01(1)0.01

40.7840.78(45)40.78(28)

0.22

0.370.37(5)0.34(3)0.04

0.410.41(3)0.41(3)0.01

3.533.53(23)3.53(37)

0.10

however, has been shown to underestimate significantlyS releases for many eruptions for which independent es-timates are available, including both anhydrite-bearing(El Chichon: Luhr et al. 1984; Pinatubo: Gerlach et al.1996) and anhydrite-free (St. Helens: Gerlach and McGee1994; Redoubt: Gerlach et al. 1994) magmatic systems.Apatite is a S-bearing phosphate mineral that is highly

resistant to weathering, alteration, and diffusion processes(Elliott and Young 1973; Roegge et al. 1974; Watson etal. 1985; Piccoli 1992). Apatite in hydrothermal ore de-posits and hydrothermally altered rocks can contain up to52.6% SO3 (see review in Shiga and Urashima 1987).However, apatite in volcanic and plutonic rocks, com-monly found as an accessory mineral, is generally poorin S (Nash 1983; Shiga and Urashima 1987). Promptedby observations of relatively high SO3 contents in apatitesfrom El Chichon (0.34 wt%: Luhr et al. 1984) and Pina-tubo (0.89 wt%: Matthews et al. 1992), our study wasinitiated to define the factors controlling S contents inapatite and to determine whether S in apatite could beused to identify ancient eruptions of S-rich anhydrite-bearing magmas that released exceptionally largeamounts of S gases to the atmosphere. To achieve thesegoals, apatite crystals from two types of samples havebeen analyzed for S and 14 other major and trace ele-ments by electron microprobe: (1) H2O- and S-saturatedexperimental charges of the trachyandesite erupted fromEl Chichon volcano in 1982 (Luhr 1990) and (2) rockserupted from 20 volcanoes (Appendix 1). Among the nat-ural samples is a subset known to be anhydrite-saturated(from El Chichon, Pinatubo, Lascar, Julcani, SutterButtes, and Cerro Lanza).

EXPERIMENTAL TECHNIQUESDetails concerning the phase-equilibrium experiments

are given in Luhr (1990) and are discussed only brieflyhere. Only apatites from charges of the 1982 El Chichontrachyandesite composition were analyzed in this study.The experiments were performed at the U.S. GeologicalSurvey (USGS) in Reston, Virginia, using an internallyheated pressure vessel with argon gas as the pressure me-dium. Most experiments were conducted using a gridlikearray of temperature (800, 850, 900, 950, and 1000 8C)and pressure (2 and 4 kbar) conditions. Fewer experi-ments were conducted at pressures of 1 and 2.5 kbar.Oxygen fugacity was buffered using one of three differentH2O-saturated solid oxygen buffers in a double-gold-cap-

sule arrangement: FMQ (fayalite-magnetite-quartz),MNH (manganosite-hausmanite), and MTH (magnetite-hematite). No reversal experiments were conducted todemonstrate equilibrium. An approach to equilibrium inthese experiments is indicated by systematic changes inglass and mineral compositions and abundances withchanging T, P, and ƒ (Luhr 1990; Housh and LuhrO21991). As discussed below, apatite S contents and apatiteand glass S partition coefficients also vary systematicallywith T and ƒ .O2Apatite crystals in the experimental charges were iden-

tified by energy-dispersive X-ray mapping. Because oftheir tiny crystal sizes (mostly ,10 mm), apatites in theexperimental charges were analyzed using the JEOL 8900electron superprobe at the USGS in Reston. Electron im-aging capabilities on this instrument are far superior tothose on the Smithsonian ARL-SEMQ microprobe. An-alytical conditions for the USGS microprobe were: 15 kVaccelerating voltage, 10 nA sample current, 20 s countingtime for the peak of each element, 10 s counting time foroff-peak backgrounds of each element, a focused probebeam (1–2 mm), and ZAF corrections.Apatite crystals in volcanic rocks were identified using

an optical microscope or luminoscope. For some pumicesamples containing relatively rare apatites, POLY-GEEbrand sodium polytungstate liquid with a density of 2.89g/cm3 and a Frantz isodynamic separator were used toseparate apatite crystals. All analyses of apatite from vol-canic rocks were performed using the Smithsonian’sARL-SEMQ electron microprobe. Analytical conditionswere: 15 kV accelerating voltage, 25 nA sample current,2–5 mm beam diameter, and on-peak mean atomic num-ber background corrections. For nine elements (Si, Al,Fe, Mg, Ca, K, Na, P, and S), we used 40 s peak countingtimes. For the other six elements (La, Ce, Mn, Sr, Cl, andF), we used 20 s peak counting times.The same standards were used on both electron micro-

probes: Kakanui hornblende (Jarosewich et al. 1980) forSi, Al, Fe, Mg, K, and Na; Durango apatite (Jarosewichet al. 1980) for P, Ca, and F; manganite for Mn; Brazilianscapolite (Jarosewich et al. 1980) for Cl; celestine for Sand Sr; and REE3 glass (Drake and Weill 1972) for Laand Ce. The calibration was checked and analytical driftwas corrected by repeated analyses of Kakanui horn-blende, Durango apatite, Brazilian scapolite, celestine,REE3 glass, and glass standards V, W, and X from theMicrobeam Analysis Society. S was analyzed on an ADPcrystal at the Smithsonian and on a PETH crystal at theUSGS.Analytical accuracy and precision were evaluated (Ta-

ble 1) on the basis of repeated analyses of Durango ap-atite as an unknown during each operating session at theSmithsonian and the USGS. According to the methods ofGoldstein et al. (1992), analytical sensitivity of major el-ements and detectability limits of trace elements (95%confidence) were evaluated using 20 analyses of Durangoapatite performed on the Smithsonian microprobe (Table1); similar values are expected for the USGS microprobe.

1212 PENG ET AL.: S IN APATITE

FIGURE 1. Backscattered-electron images of apatite crystalsin the experimental charges: (A) an isolated apatite crystal en-closed in glass; (B) an apatite crystal included in plagioclase.Abbreviations of phases: ap 5 apatite; a 5 anhydrite; g 5 glass;h 5 hornblende; p 5 plagioclase; c 5 clinopyroxene; m 5magnetite.

FIGURE 2. Compositional zoning in single apatite crystalsfrom the MNH- and MTH-buffered experimental charges: (a)SO3 vs. SiO2; (b) SO3 vs. Na2O. Lines connect the cores and rimsof individual crystals.

Stormer et al. (1993) reported variations of F and ClX-ray intensities during electron microprobe analysis be-cause of anisotropic diffusion in apatite. A 3 min time-diffusion study similar to that of Stormer et al. (1993)was performed on oriented sections (c, \ c) of Durangoapatite using the Smithsonian’s ARL-SEMQ electron mi-croprobe. Similar to the results of Piccoli (1992), our testsshowed relatively constant F and Cl count rates over a 3min period. Nonetheless, we ran F and Cl first during themicroprobe analyses but do not believe that diffusion sig-nificantly affected our analytical results.

APATITE FROM EXPERIMENTAL CHARGES

Apatites in the experimental charges occur mainly asisolated grains in glass (or vesicles) (Fig. 1a) and areoccasionally included in or clustered with plagioclase, cli-nopyroxene, or hornblende (Fig. 1b). Among the 133 ex-perimental apatite crystals that were analyzed, only 12

grains are mostly or entirely included in plagioclase orclinopyroxene. The short dimension of each analyzedgrain section was recorded. Most of the apatite crystalsare 5–6 mm wide, and only eight grains have a shortdimension larger than 10 mm. The largest apatite grainfound in these experimental charges is 14 3 48 mm. Thecharges run at 900 8C and 4 kbar contain more abundantand larger apatite crystals than other charges.The tiny sizes of apatite in the experimental charges

and the realities of beam excitation volumes caused someanalyses to represent mixtures of apatite and adjacentglass (or silicate minerals). The following procedureswere used to correct these analytical results. BecauseAl2O3 contents of apatite in volcanic rocks are typicallyclose to zero (see Table 2), we deleted all ten analyseswith .0.30 wt% Al2O3 (arbitrarily chosen), which left 172apatite point analyses with Al2O3 above the detectabilitylimit (0.01%). We assumed that all measured Al2O3 re-sulted from excitation of the enclosing phase (mostlyglass). Accordingly, the composition of that phase, deter-mined for that particular charge (Luhr 1990; Housh andLuhr 1991; Luhr, unpublished data) was subtracted fromthe apatite analysis until Al2O3, MgO, Na2O, or K2Oreached zero.Zoning of trace elements (e.g., REE, Y, Sr, U, and Mn)

in apatite was reported by various works (e.g., Knutsonet al. 1985; Joliff et al. 1989; Rakovan and Reeder 1994).We also observed core-to-rim elemental zoning in apa-tites during this study. The largest apatite grains (;10 mmin width) from the experimental charges were evaluatedfor zoning. Apatites from the FMQ-buffered charges con-tain SO3 at or below the detectability limit, and thus showno SO3 zoning. In the MNH- and MTH-buffered charges,apatites show core-to-rim decreases of SO3, SiO2, andNa2O (Figs. 2a and 2b). The compositional zoning of thelargest apatite crystals in the experimental charges indi-cates that kinetic effects probably inhibited equilibriumexchange between apatite and melt during rapid experi-mental cooling from 1000 8C to experimental temperature(Luhr 1990). As discussed above, however, many lines ofevidence indicate that the experimental charges closelyapproach equilibrium of phase assemblages and compo-


TABLE 2. Average compositions of apatites from volcanic rocks (wt%)

VolcanoGrains

El Chichon25

Pinatubo11

Lascar55

Julcani19

Sutter Buttes16

Cerro Lanza4

Etna8

SiO2Al2O3La2O3Ce2O3FeOMnOMgOCaOSrONa2OK2OP2O5SO3CIF

0.17(7)0.01(1)0.08(2)0.19(3)0.19(9)0.14(1)0.01(1)55.18(41)0.10(2)0.09(2)0.01(1)41.93(54)0.23(9)0.59(8)1.87(11)

0.01(2)0.00(0)0.09(2)0.22(2)0.32(4)0.20(3)0.10(1)54.12(44)0.06(1)0.11(1)0.01(1)41.28(90)0.13(2)1.17(15)1.46(22)

0.15(5)0.00(0)0.08(4)0.23(8)0.46(21)0.09(2)0.20(6)54.63(51)0.07(3)0.13(3)0.01(1)41.81(46)0.24(11)0.67(14)2.51(25)

0.21(6)0.00(0)0.15(5)0.36(7)0.25(19)0.16(3)0.01(1)54.36(32)0.09(3)0.12(4)0.01(2)41.64(22)0.17(9)0.98(38)2.14(23)

0.01(1)0.02(1)0.02(1)0.08(4)0.10(2)0.09(1)0.04(3)55.28(34)0.08(2)0.06(1)0.00(0)41.69(119)0.12(3)0.22(4)1.75(10)

0.09(1)0.00(0)0.08(1)0.15(3)0.13(1)0.19(1)0.00(0)55.36(36)0.06(1)0.08(1)0.00(0)41.97(50)0.20(0)0.76(8)2.26(4)

0.21(11)0.00(0)0.13(4)0.28(6)0.40(25)0.09(1)0.27(5)54.31(38)0.26(3)0.16(4)0.02(2)41.33(34)0.24(6)0.59(19)2.82(29)

O 5 CI, F 20.92 20.88 21.21 21.12 20.78 21.12 21.32Total 99.87 98.42 100.09 99.53 98.79 100.18 99.80

AlLaCeFeMnMgCaSrNaKSubtotal

0.00050.00250.00570.01320.01020.00144.94630.00510.01430.00085.0000

0.00020.00270.00690.02270.01440.01294.91800.00300.01830.00095.0000

0.00010.00250.00700.03240.00630.02464.90050.00350.02180.00135.0000

0.00000.00480.01120.01770.01150.00094.92800.00450.02000.00145.0000

0.00190.00080.00250.00720.00640.00554.96080.00410.01030.00055.0000

0.00000.00250.00470.00930.01350.00014.95360.00310.01260.00055.0000

0.00050.00400.00850.02830.00670.03344.87810.01260.02580.00215.0000

PSiSSubtotal

CIFSubtotal

2.96990.01450.01462.99890.08320.49600.5793

2.96410.00130.00812.97350.16870.39270.5615

2.96370.01280.01522.99170.09500.66490.7599

2.98310.01770.01103.01180.14040.57250.7129

2.95610.00070.00772.96450.03150.46240.4938

2.96960.00720.01292.98960.10750.59780.7053

2.93340.01790.01512.96630.08350.74880.8324

Note: Numbers in parentheses are one standard deviation values.

sitions. Because of the tiny sizes of the experimental ap-atite crystals and the systematic correlation of apatitecompositions with T, P, and ƒ , we believe that averageO2compositions of apatite in the experimental charges arein approximate equilibrium with the melt (glass)compositions.Average chemical compositions of the apatite from 19

experimental charges are listed in Table 3. Two additionalcharges, produced at 1000 8C and 2 kbar (no. 93 and 161:Luhr 1990), were mapped using the energy-dispersiveX-ray method, but no analyzable apatite crystals werefound. Apatite can be represented by the formulaCa5(PO4)3(F, Cl, OH). Apatite formulas in this paper werecalculated on the basis of five atoms of Al 1 La 1 Ce1 Fe 1 Mn 1 Mg 1 Ca 1 Sr 1 Na 1 K. We assumedthat Si and S substitute for P. No systematic composition-al differences were found between the apatites surroundedby glass and those included in silicate minerals from thesame charges. SO3 contents in the experimental apatitesbuffered by FMQ are at or below the detectability limitof 0.04 wt% (Table 3; Fig. 3a). In charges buffered byMNH and MTH, apatite SO3 concentrations are muchhigher (1.0–2.6 wt%: Table 3; Fig. 3a), with no consistentdifference between the two buffers.

The average SO3 contents of apatites in the MNH-and MTH-buffered experimental charges generally de-crease with increasing temperature at a given pressureand oxygen buffer (Fig. 3b), the opposite of SO3 be-havior in the experimental glasses (Luhr 1990). TheSO3 concentrations in MNH- and MTH-buffered apa-tites increase with increasing pressure at a given tem-perature and oxygen buffer (Fig. 3c), which is consis-tent with SO3 increases in the experimental glasses(Luhr 1990). Accordingly, the apatite-melt partition co-efficient for SO3 decreases with increasing temperature(Fig. 4) but appears to be independent of pressure. Theapatite-melt partition coefficient for SO3 is consistentlylower in the FMQ-buffered charges than in those buf-fered by MNH and MTH. This probably indicates thatat the lower ƒ of the FMQ buffer we are evaluatingO2

partitioning of sulfide rather than sulfate (Carroll andRutherford 1988). On the basis of linear regression ofthe experimental results in Figure 4, the apatite-meltpartition coefficient for SO3 buffered by MNH andMTH can be expressed as: lnKd 5 21 130/T 2 16.2.Two mechanisms have been proposed to explain the

substitution of S for P in apatite:


TABLE 2—Continued

VolcanoGrains

Santorini22

Krakatau13

Tambora5

Bogoslof11

Novarupta23

Lassen10

Mazama8

SiO2Al2O3La2O3Ce2O3FeOMnOMgOCaOSrONa2OK2OP2O5SO3CIF

0.23(15)0.00(0)0.08(3)0.22(9)0.75(23)0.14(2)0.15(6)54.18(84)0.04(2)0.08(2)0.02(2)41.65(44)0.03(3)1.04(10)2.38(30)

0.32(9)0.00(0)0.05(2)0.16(4)0.63(18)0.16(3)0.20(4)53.38(43)0.04(2)0.11(3)0.01(1)40.95(23)0.24(10)0.86(11)2.63(24)

0.09(2)0.00(0)0.07(3)0.19(3)0.29(10)0.08(2)0.20(1)54.53(27)0.29(2)0.08(1)0.01(1)42.01(31)0.18(2)0.44(2)2.13(7)

0.26(13)0.00(0)0.09(3)0.20(6)0.14(11)0.20(4)0.00(1)55.64(45)0.09(5)0.15(5)0.01(1)41.28(54)0.51(22)0.34(8)3.03(43)

0.30(10)0.00(0)0.05(3)0.19(9)0.74(19)0.13(3)0.12(5)53.70(48)0.04(2)0.09(3)0.01(1)41.39(34)0.11(7)0.94(15)2.78(25)

0.18(5)0.00(0)0.20(3)0.46(4)0.36(10)0.23(2)0.06(2)53.97(59)0.05(1)0.17(6)0.02(2)41.33(51)0.23(16)1.24(10)1.66(19)

0.19(3)0.00(0)0.03(2)0.09(4)0.33(4)0.08(2)0.32(7)54.04(37)0.18(6)0.20(3)0.01(2)42.01(59)0.63(10)0.95(7)1.83(25)

O 5 CI, FTotal

21.2499.75

21.3098.46

21.0099.59

21.35100.59

21.3999.21

20.9899.20

20.9899.90


0.00010.00260.00680.05280.01010.01844.89260.00180.01250.00235.0000

0.00020.00160.00500.04510.01170.02564.88910.00210.01850.00115.0000

0.00000.00220.00580.02050.00560.02514.91220.01430.01320.00125.0000

0.00020.00280.00590.00970.01430.00034.93750.00450.02340.00145.0000

0.00000.00170.00590.05240.00900.01564.89770.00190.01490.00095.0000

0.00000.00640.01410.02540.01650.00734.89680.00250.02850.00255.0000

0.00030.00080.00260.02360.00590.04074.88310.00870.03280.00155.0000

PSiSSubtotal

CIFSubtotal

2.97170.01910.00172.99250.14840.63320.7816

2.96350.02740.01573.00660.12500.71090.8359

2.99030.00730.01113.00870.06270.56560.6283

2.89450.02170.03182.94800.04720.79360.8408

2.98280.02550.00693.01520.13610.74950.8856

2.96260.01540.01492.99290.17870.44490.6236

2.99940.01600.03963.05510.13600.48690.6229

61 41 51S 1 Si 5 2P(e.g., Rouse and Dunn 1982) (1)

61 1 51 21S 1 Na 5 P 1 Ca(e.g., Liu and Comodi 1993) (2)

The Si and Na contents of apatite in the more oxidizingMNH- and MTH-buffered charges increase with S (Figs.5a and 5c). However, the correlation coefficient betweenSi and S (r 5 0.93) is higher than that between Na andS (r 5 0.67). In addition, the Si/S ratios in the apatitesare closer to unity than Na/S ratios (Figs. 5a and 5c).Therefore, the substitution mechanism of Equation 1 isconsidered to be more significant than that of Equation 2for the apatites in the experimental charges buffered byMNH and MTH. Because of the positive correlation be-tween S and Si contents in apatite, the dependence of SO3

partitioning between apatite and melt on temperature(Fig. 4), and the poor correlation between the apatite-meltSiO2 partition coefficient and temperature (r 5 20.46),we conclude that the concentration of Si in apatites buf-fered by MNH and MTH is controlled mainly by the ap-atite S content. Sulfur content in apatite, in turn, is gov-erned chiefly by melt S content, T, P, and ƒ , asO2discussed above.Besides the substitution mechanism in Equation 1, Si

can also enter the apatite structure by the followingmechanisms:

42 22 32SiO 1 CO 5 2PO (Sommerauer and4 3 4

Katz-Lehnert 1985) (3)41 31 51 21Si 1 REE 5 P 1 Ca

(e.g., Roeder et al. 1987) (4)

The FMQ-buffered apatite contains 0.69–1.6 wt%SiO2. Their low SO3 contents (#0.04 wt%) preclude op-eration of the substitution mechanism in Equation 1. Themechanism of Equation 3 can also be ruled out becausethe sum of Si 1 P in many FMQ-buffered apatites is closeto 3 pfu (Fig. 6), and the presence of C in amounts equalto Si would cause Si 1 C 1 P to exceed three consid-erably. As is common for apatites (Nash 1983), those inEl Chichon trachyandesite are relatively enriched inLREE (Luhr et al. 1984). Apatites in the FMQ-bufferedexperimental charges have La2O3 1 Ce2O3 contents rang-ing from 0.07 to 0.46 wt%. We estimated the total REEcontents for these experimental apatites by assuming thatthey have the same relative REE abundances as the nat-ural apatite from 1982 El Chichon pumices, which Luhret al (1984) analyzed in bulk by instrumental neutron ac-tivation. Even so, total REE concentrations in the FMQ-


TABLE 2—Continued

VolcanoGrains

St. Helens10

Bishop Tuff5

Mesa Falls Tuff11

Ceboruco16

Popocatepetl15

Santa Maria13

SiO2Al2O3La2O3Ce2O3FeOMnOMgOCaOSrONa2OK2OP2O5SO3CI

0.04(6)0.00(0)0.04(2)0.10(3)0.44(20)0.08(2)0.14(2)54.51(36)0.07(3)0.04(2)0.01(1)42.08(79)0.04(3)0.72(5)

0.96(16)0.01(2)0.33(4)1.00(14)0.63(23)0.22(2)0.00(0)53.57(32)0.01(2)0.11(1)0.01(1)39.93(26)0.00(1)0.21(3)

2.10(115)0.02(1)0.77(36)2.11(99)0.85(32)0.12(2)0.00(0)51.09(223)0.01(2)0.13(2)0.02(2)37.62(206)0.01(1)0.24(4)

0.07(4)0.01(1)0.08(3)0.19(5)0.53(13)0.15(4)0.21(8)54.62(58)0.09(2)0.11(3)0.01(2)41.86(77)0.10(5)0.50(10)

0.16(9)0.00(0)0.07(2)0.19(4)0.61(25)0.07(1)0.26(5)54.73(44)0.10(2)0.15(5)0.02(3)41.86(68)0.29(13)0.45(9)

0.04(5)0.01(1)0.04(1)0.16(2)0.61(36)0.29(4)0.17(1)54.09(62)0.07(2)0.19(4)0.01(1)41.38(50)0.36(14)1.11(11)

F 1.88(15) 2.65(50) 3.27(39) 2.83(35) 2.67(31) 1.45(9)O 5 CI, FTotal

20.9599.26

21.1698.47

21.4396.95

21.30100.06

21.23100.39

20.8699.13


0.00020.00120.00320.03090.00600.01824.92860.00360.00700.00105.0000

0.00110.01030.03100.04480.01550.00004.87830.00070.01730.00095.0000

0.00240.02490.06780.06260.00890.00004.80730.00050.02280.00275.0000

0.00060.00250.00590.03730.01030.02634.89370.00460.01760.00135.0000

0.00000.00210.00580.04260.00500.03254.88140.00470.02380.00225.0000

0.00060.00120.00480.04300.02070.02174.87300.00350.03070.00075.0000

PSiSSubtotal

CIFSubtotal

3.00670.00370.00283.01310.10260.50230.6048

2.87340.08130.00022.95490.02970.71190.7416

2.79730.18460.00082.98270.03610.90700.9431

2.96370.00610.00612.97590.07060.74760.8183

2.95020.01330.01802.98150.06290.70400.7669

2.94570.00380.02302.97250.15890.38540.5443

buffered apatite are still far insufficient to balance chargesfor the substitution mechanism of Equation 4 because ofthe high SiO2 concentrations (0.69–1.6 wt%) in these ap-atites. Besides, the Si 1 Na remaining in FMQ-bufferedapatite after accounting for the coupled substitutions ofEquations 1 and 2 correlates poorly with La 1 Ce (Fig.7a). In the light of these observations, some additionalsubstitution mechanism for Si must operate in apatite ofthe FMQ-buffered experimental charges, and the follow-ing mechanism is speculatively proposed:

Si41 1 M 5 P51 1 (F, Cl, OH)2 (5)(where M represents a vacancy on the hydroxyl site).The apatite-melt partition coefficient for SiO2 in the

FMQ-buffered charges decreases with increasing temper-ature and is independent of pressure. Therefore, both meltSiO2 content and temperature may control SiO2 concen-trations in the FMQ-buffered apatite.

APATITE FROM NATURAL VOLCANIC ROCKS

Apatite from rocks erupted at 20 volcanoes was alsoanalyzed, including pumices from many of the largest ex-plosive eruptions of historical time (Appendix 1). Thesevolcanoes, which mainly occur in subduction-related arcs,can be divided into two groups according to the knownpresence or inferred absence of primary anhydrite. Sam-ples with primary anhydrite are from El Chichon (Luhr

et al. 1984), Pinatubo (Bernard et al. 1991), Lascar (Mat-thews et al. 1994), Julcani (Drexler and Munoz 1985),Sutter Buttes, and Cerro Lanza. Anhydrite is present indrill core samples from Sutter Buttes, but is absent inexposed volcanic rocks (Brian Hausback, personal com-munication). At Cerro Lanza, a Pliocene-Quaternarydome complex approximately 130 km south-southeast ofEl Chichon, anhydrite was observed as a single micro-scopic inclusion in plagioclase. All the primary anhydritecrystals occur in samples that are relatively oxidized(high Fe2O3/FeO), porphyritic, and rich in phenocrysts ofthe hydrous minerals amphibole or biotite. The secondgroup of samples lacks evidence of primary anhydrite,although anhydrite might actually have been present insome of them originally. These samples are from Etna,Santorini, Krakatau, Tambora, Bogoslof, Novarupta, Las-sen, Mazama, St. Helens, Long Valley (Bishop Tuff ),Yellowstone (Mesa Falls Tuff ), Ceboruco, Popocatepetl,and Santa Marıa.Apatite is a common accessory mineral in most of

these rocks, with the exceptions of rhyodacites and rhy-olites, which are typically impoverished in P2O5 (Greenand Watson 1982). Apatite in volcanic rocks occurs assingle crystals entirely or mostly surrounded by: (1) ma-trix; (2) anhydrous mineral phases (plagioclase, clinopy-roxene, or orthopyroxene); (3) hydrous mineral phases


FIGURE 6. P vs. Si for apatites from the experimental charges(atoms per formula unit). Line shows condition Si 1 P 5 3.Many apatites from FMQ-buffered charges fall along this trend.

FIGURE 7. Compositional plot to test the importance of thesubstitution mechanisms in Equations 4 and 6: Si 1 Na - S vs.La 1 Ce in (a) experimental charges and (b) volcanic rocks.Lines show 1:1 correlations. Symbols as in Figure 5. The termSi 1 Na - S accounts for remaining Si and Na atoms after in-corporating them into the apatite structure through the combinedactions of the mechanisms in Equations 1 and 2.

8a), (2) rimward increases in SO3 (Fig. 8b), and (3) os-cillatory variations (Fig. 8c). A single sample may con-tain crystals with all three types of zoning. Our experi-mental results indicate that under quasi-equilibriumconditions, apatite S contents are controlled mainly bymelt S content and by T, P, and ƒ of the magma system.O2Compositional zoning and grain-to-grain compositionaldifferences in volcanic apatite from single thin sectionsimply that other factors are important in real magma sys-tems. Probably most important are slow intracrystallinediffusion rates for apatite, which inhibit its continuedre-equilibration, coupled with time-dependent variationsin melt composition and T, P, and ƒ in real magmaticO2systems (Elliott and Young 1973; Roegge et al. 1974;Knutson et al. 1985; Watson et al. 1985; Joliff et al. 1989;Piccoli 1992; Rakovan and Reeder 1994).Apatite compositions in volcanic rocks fall into two

groups on the S vs. Si 1 Na diagram (Fig. 5f ). Apatitesof the first group, from high-silica rhyolites of the BishopTuff and Mesa Falls Tuff, plot parallel to the Si 1 Naaxis along a similar trend as shown by the experimentalapatites buffered by FMQ (Fig. 5e); these apatites containlow SO3 but considerable SiO2 contents. The first-groupvolcanic apatites are different from the FMQ-buffered ex-perimental apatites, however, in that the Si 1 Na remain-ing after accounting for the coupled substitutions in Equa-tions 1 and 2 are positively correlated with La 1 Ce (Fig.7b), suggesting that the substitution mechanism given byEquation 4 is important in apatite from volcanic rocks,as found earlier by Piccoli and Candela (1988).Apatites of the second group, from the other 18 vol-

canoes, are characterized by a positive relationship be-tween S and Si 1 Na (Fig. 5f ). However, the substitutionmechanism of Equation 2 appears to be more important

than that of Equation 1 for most samples because S andNa (r 5 0.82) have better correlation than S and Si (r 50.38) (Figs. 5b and 5d). Besides the Na substitutionmechanism of Equation 2, another mechanism proposedfor Na in apatite is:

Na1 1 REE31 5 2Ca21 (e.g., Roeder et al. 1987) (6)However, it seems this is not significant for the apatite involcanic rocks because Na and La 1 Ce are poorly cor-related (r 5 0.25). The (La 1 Ce)/(Si 1 Na 2 S) valuesof the apatites in volcanic rocks are significantly less thanunity (Fig. 7b), because other REE not analyzed in thisstudy may be present in relatively high abundance (Luhret al. 1984; Fleet and Pan 1995). Among the studied ap-atites, those from the Bishop Tuff and Mesa Falls Tuffcontain the highest La and Ce concentrations (Table 2),which are possibly due to high apatite-melt REE partitioncoefficients and high melt-phase concentrations of REEin high-silica rhyolites (Hildreth 1977; Watson and Green1981; Nash 1983). For melts of similar SiO2 content, ap-atite-liquid REE partition coefficients increase with de-creasing temperature (Watson and Green 1981). AlthoughSiO2 contents of rhyolitic pumices from the early BishopTuff (SiO2, 74.5%) and Mesa Falls Tuff (SiO2, 74.6%) arevery close and the crystallization temperature of the MesaFalls Tuff (;880 8C) is higher than that in the early Bish-op Tuff (;725 8C), LREE concentrations in apatites ofthe former are much higher, consistent with whole-rockREE concentration differences (Hildreth 1977, 1981; Hil-dreth et al. 1991). Therefore, according to the coupled-substitution mechanism of Equation 4, it appears thatREE concentrations in the melt may control the silicacontent in apatite rather than silica activity in melt havingcontrol over the REE composition of apatite (Rønsbo1989).

DISCUSSIONS in magmatic systems and apatiteS in silicate melts can exist as both reduced (sulfide)

and oxidized (sulfate) species. The experiments of Carroll


TABLE 3. Average compositions of apatites from experimental charges (wt%)

Charge No.GrainsT (8C)P (kbar)O2 Buffer

15328002

FMQ

14928002

MNH

10538002

MTH

26318002.5MNH

26838002.5MTH

21888004

FMQ

20388004

MNH

193138004

MTH

14438502

FMQ

SiO2Al2O3La2O3Ce2O3FeOMnOMgOCaOSrONa2OK2OP2O5SO3CIFO 5 CI, FTotal

1.58(36)0.01(2)0.03(4)0.04(5)0.58(0)0.10(3)0.00(0)54.53(22)0.25(3)0.00(1)0.09(2)40.08(38)0.02(2)0.29(9)1.67(8)

20.7798.51

1.66(48)0.01(1)0.07(6)0.08(10)0.16(7)0.17(3)0.02(1)55.46(114)0.08(6)0.22(1)0.03(3)37.23(44)2.30(48)0.29(1)2.12(27)

20.9698.94

0.88(18)0.00(0)0.07(8)0.33(21)0.23(9)0.20(5)0.03(4)55.47(93)0.08(5)0.24(1)0.08(3)39.49(43)1.43(48)0.29(7)2.03(29)

20.9299.94

0.950.000.020.400.210.170.0055.110.140.400.0338.562.340.252.12

20.9599.73

1.39(4)0.00(0)0.05(1)0.16(14)0.31(12)0.12(4)0.05(3)53.93(49)0.13(2)0.21(1)0.09(4)37.82(32)2.14(18)0.21(4)1.84(8)

20.8297.62

1.61(62)0.00(0)0.07(7)0.11(11)0.52(16)0.11(5)0.00(1)54.63(70)0.26(7)0.03(4)0.08(4)39.97(63)0.03(5)0.12(3)1.52(10)20.6798.41

1.59(71)0.00(0)0.04(4)0.10(10)0.21(10)0.12(6)0.05(4)55.62(76)0.12(4)0.31(10)0.07(3)37.89(190)2.37(96)0.09(5)1.79(47)

20.7799.58

1.78(36)0.00(0)0.04(4)0.13(12)0.23(11)0.10(5)0.03(4)55.89(55)0.13(4)0.35(16)0.08(4)37.64(89)2.60(55)0.10(3)1.61(22)

20.7099.99

0.77(10)0.00(0)0.08(9)0.16(4)0.55(16)0.15(5)0.04(2)55.57(74)0.22(8)0.05(3)0.11(4)41.35(42)0.02(2)0.30(3)1.84(17)

20.84100.36


0.00130.00100.00120.04090.00730.00004.92540.01220.00070.01015.0000

0.00050.00210.00250.01100.01190.00224.92730.00380.03570.00305.0000

0.00000.00230.01010.01600.01390.00374.90300.00370.03860.00885.0000

0.00000.00060.01200.01470.01220.00004.88780.00650.06360.00295.0000

0.00010.00160.00500.02200.00850.00664.90510.00630.03510.00985.0000

0.00020.00210.00350.03660.00820.00034.92240.01260.00520.00895.0000

0.00000.00110.00300.01440.00800.00654.90370.00560.05010.00765.0000

0.00000.00110.00390.01590.00710.00394.89870.00610.05550.00795.0000

0.00000.00230.00490.03820.01050.00514.90930.01030.00780.01155.0000

PSiSSubtotal

CIFSubtotal

2.86070.13320.00142.99530.04150.44660.4881

2.61510.13750.14262.89510.04100.55800.5990

2.75880.07270.08872.92010.04130.53070.5720

2.70210.07840.14522.92560.03490.55530.5903

2.71830.11760.13632.97220.03020.49460.5248

2.84580.13530.00202.98300.01650.40360.4201

2.63950.13050.14632.91630.01260.46530.4779

2.60680.14530.15952.91160.01360.41640.4300

2.88680.06370.00122.95170.04210.47870.5208

Note: Numbers in parentheses are one standard deviation values.

and Rutherford (1988) indicated that the sulfate-sulfideratio in silicate melts increases with increasing ƒ . Sul-O2fate solubility in silicate melts generally increases withincreasing ƒ , temperature, and pressure (Haughton et al.O21974; Katsura and Nagashima 1974; Carroll and Ruth-erford 1987; Luhr 1990).The eruptions of certain magmas, especially those con-

taining phenocrystic anhydrite, have been shown to re-lease to the atmosphere one to two orders of magnitudemore S gases than could have been dissolved in the meltphase of the erupted magmas at the pre-eruptive temper-ature and pressure (Luhr et al. 1984; Carroll and Ruth-erford 1987; Luhr 1990; Andres et al. 1991; Westrich andGerlach 1992; Gerlach and McGee 1994; Gerlach et al.1994, 1996). The S in addition to that dissolved in thepre-eruption melt has been referred to as ‘‘excess sulfur’’and is commonly envisioned to reside as a separate gasphase in these vesiculated vapor-saturated magmas. Var-ious origins for the excess sulfur have been evaluated: (1)fractionation of H2O- and S-rich basaltic magma (Luhr1990); (2) interaction between magma and S-enrichedsubvolcanic rocks, such as altered pyrite- and anhydrite-bearing lithic fragments (Carroll and Rutherford 1987;

Oppenheimer 1996); (3) gas-phase contributions fromunerupted magmas or intrusions (Andres et al. 1991; Pal-lister et al. 1992; Hattori 1993, 1996; Wallace and Ger-lach 1994); (4) mixing of S-poor, oxidized, evolved mag-ma with S- and volatile-rich, reduced, mafic magma(Matthews et al. 1992); and (5) direct accumulation of avapor phase within the magma body before eruption(Gerlach et al. 1996).Anhydrite can occur as a stable phenocrystic phase in

S-saturated silicate melts at ƒ greater than or equal toO21.0–1.5 log ƒ units above NNO (DNNO $ 1.0–1.5)O2(Carroll and Rutherford 1987). This ƒ of anhydrite sta-O2bility derived from the experiments is consistent with theƒ values estimated from natural anhydrite-bearing mag-O2mas at El Chichon (Luhr et al. 1984), Pinatubo (Ruther-ford and Devine 1996), Lascar (Matthews et al. 1994),and Julcani (Drexler and Munoz 1985). Many subduc-tion-related calc-alkaline volcanic rocks, particularlythose rich in hornblende or biotite, appear to have crys-tallized at ƒ high enough to stabilize anhydrite (Fig. 9a).O2The rarity of anhydrite reports for historical calc-alkalineeruptive products suggests that S contents in most of themagmas are at levels below significant anhydrite satura-


TABLE 3—Extended

Charge No.GrainsT (8C)P (kbar)O2 buffer

16568502

MNH

13988502

MTH

18579002

FMQ

16999002

MNH

11439002

MTH

223149004

FMQ

189159004

MNH

181119004

MTH

104129502

MNH

11959502

MTH

SiO2Al2O3La2O3Ce2O3FeOMnOMgOCaOSrONa2OK2OP2O5SO3CIFO 5 CI, FTotal

0.74(20)0.00(0)0.05(4)0.18(14)0.33(11)0.17(9)0.09(9)55.26(32)0.12(4)0.31(6)0.09(2)39.85(78)1.39(32)0.30(10)1.69(19)

20.7899.80

1.08(42)0.00(0)0.05(4)0.16(12)0.24(7)0.18(3)0.07(2)55.02(47)0.13(5)0.22(11)0.14(3)39.41(94)1.52(51)0.35(4)1.82(16)

20.8599.95

0.72(28)0.00(0)0.11(8)0.35(11)0.61(13)0.12(4)0.07(2)54.74(76)0.17(4)0.04(3)0.11(3)41.02(83)0.04(5)0.24(4)1.51(22)

20.6999.16

0.73(16)0.00(0)0.02(5)0.06(6)0.32(8)0.13(4)0.14(2)54.95(42)0.15(4)0.25(7)0.10(3)40.08(51)1.24(28)0.27(2)1.53(15)

20.7099.27

0.56(14)0.00(0)0.06(5)0.21(11)0.31(5)0.15(4)0.16(2)55.06(55)0.13(4)0.23(5)0.12(2)40.99(25)1.01(16)0.31(6)1.70(11)

20.79100.21

0.69(20)0.01(1)0.11(6)0.28(18)0.43(9)0.08(5)0.07(2)55.28(48)0.19(4)0.01(2)0.08(3)41.15(51)0.02(2)0.11(2)1.16(12)

20.5199.16

1.13(28)0.00(0)0.03(4)0.08(6)0.26(12)0.11(4)0.12(2)54.89(48)0.15(3)0.29(7)0.09(4)38.99(68)1.84(43)0.12(2)1.22(16)

20.5498.79

1.24(37)0.00(0)0.03(3)0.09(11)0.28(24)0.09(5)0.13(2)55.61(45)0.14(3)0.29(7)0.08(3)38.95(99)2.09(51)0.12(2)1.26(10)

20.5699.86

0.80(15)0.00(0)0.03(3)0.08(8)0.29(8)0.11(4)0.17(4)55.36(41)0.14(3)0.20(5)0.10(3)39.99(49)1.34(32)0.25(5)1.39(11)

20.6499.60

0.90(10)0.00(0)0.04(3)0.11(11)0.25(4)0.11(2)0.18(2)55.52(34)0.16(3)0.22(6)0.11(2)39.80(29)1.47(23)0.30(10)1.42(4)

20.6699.91


0.00000.00160.00530.02270.01180.01144.88250.00580.04930.00965.0000

0.00000.00150.00490.01690.01260.00874.89930.00650.03490.01485.0000

0.00000.00350.01080.04290.00850.00864.89940.00840.00590.01215.0000

0.00000.00070.00170.02200.00950.01754.89080.00740.04030.01015.0000

0.00000.00180.00630.02160.01080.01934.88500.00640.03660.01235.0000

0.00080.00350.00860.02960.00560.00864.92300.00920.00200.00905.0000

0.00000.00080.00240.01840.00790.01524.89130.00730.04670.00995.0000

0.00000.00080.00260.01950.00660.01614.89300.00680.04630.00845.0000

0.00000.00090.00240.02000.00780.02134.89950.00650.03150.01015.0000

0.00000.00110.00340.01690.00740.02194.89540.00750.03450.01205.0000

PSiSSubtotal

CIFSubtotal

2.78250.06070.08602.92920.04210.44110.4832

2.77260.08980.09522.95960.04940.47900.5284

2.90070.06000.00252.96320.03350.39740.4308

2.81900.06050.07702.95650.03840.40090.4393

2.87400.04620.06282.98310.04340.44530.4887

2.89520.05720.00092.95330.01530.30370.3190

2.74570.09400.11452.95420.01760.32190.3395

2.70840.10210.12902.93960.01610.32850.3446

2.79690.06630.08282.94610.03460.36430.3990

2.77300.07370.09092.93760.04200.36850.4105

tion, although the scarcity of anhydrite in volcanic rockscould be the result of post-eruption leaching of anhydriteby surface and ground waters (Luhr et al. 1984; Luhr1991). Basaltic magmas tend to be richer in dissolved Sthan more silicic magmas (Wendlandt 1982; Carroll andRutherford 1985), but their typically low oxidation states(DNNO 5 20.1 to 23.2: Wallace and Carmichael 1992)prohibit stability of anhydrite.Sulfur (as SO ) substitutes into apatite, replacing the22

4

PO group. As demonstrated in this study, high mag-324

matic oxidation states favor this substitution and result inthe formation of S-rich apatite (Fig. 3a). Our study alsoindicates that temperature and pressure can affect the Sconcentration in apatite, although they are less significantfactors than ƒ (Figs. 3b and 3c). Figure 9 shows SO3O2contents of apatite from both experimental charges andvolcanic rocks plotted against ƒ , T, and P. For volcanicO2rocks we have plotted available literature estimates forpre-eruptive ƒ , T, and P for magmas from the sameO2volcanoes (Figs. 9a–c). Although the experimental datashow a clear correlation between SO3 in apatite andDNNO, no such relationship is evident for the volcanicrocks. Likewise, apatites from natural volcanic rocks

show no correlation between apatite SO3 content and pre-eruptive T or P.

Apatite fails as an indicator of ancient S-rich eruptionsFrom Table 2, it is evident that the apatites from an-

hydrite-bearing volcanic rocks do not contain exception-ally high S concentrations and are indistinguishable inthis regard from those in the anhydrite-absent volcanicrocks. This suggests that no direct relationship occurs be-tween S concentration in volcanic apatite and anhydritesaturation, which itself seems to be related to the largesteruptive releases of S gases to the atmosphere. The aboveobservations may be explained as follows:(1) Apatite appears to precipitate early and rapidly in

magmatic systems (a large volume of apatite crystallizesover a small DT just below the apatite liquidus: Piccoliand Candela 1994). Slow intracrystalline diffusion ratesmay prevent apatite from re-equilibrating with the meltand other phases during continued magmatic evolution.Specifically, apatite in anhydrite-bearing volcanic rockslikely formed before late-stage enrichment of S, whichmay be a common trigger for anhydrite saturation. In thisscenario, the apatite and anhydrite are not in exchange


FIGURE 8. Three types of compositional zoning patterns inapatite from volcanic rocks: (a) rimward decreases (an apatitecrystal from a 1902 Santa Marıa pumice; sample 5 NMNH no.113100-10; point interval 5 4.1 mm); (b) rimward increases (anapatite crystal from a 1982 El Chichon pumice; sample 5CH70-1; point interval 5 10.2 mm); (c) oscillatory variations (anapatite crystal from a 1796 Bogoslof pumice; sample 5 BogI;point interval 5 9.9 mm).

FIGURE 9. Average SO3 contents of apatite in volcanic rocksplotted against: (a) oxygen fugacity as DNNO (anhydrite is stableat values of DNNO $ 1), (b) temperature, and (c) pressure. Sym-bols as in 5. Estimates for DNNO, T, and P of volcanic rocksare taken from literature sources (Appendix 1). For comparison,lines show SO3 variations of apatite in the experimental chargeswith oxygen fugacity, temperature, and pressure (see details inFig. 3)

equilibrium; the apatite mainly crystallized under S-poorconditions and did not record the late-stage S enrichment.Experimental work shows that anhydrite can be stablethroughout the liquidus-solidus crystallization intervalprovided that sufficient S is present to exceed the meltsaturation limit and that the magma is sufficiently oxi-dized (Carroll and Rutherford 1987; Luhr 1990). None-theless, anhydrite in volcanic rocks typically occurs asisolated microphenocrysts (,0.03 mm) in matrix, or it isincluded in the outmost 200 mm of plagioclase or otherphenocrystic minerals. Accordingly, we interpret primaryanhydrite to have crystallized generally late in the mag-matic history. Apatite can be found included in or clus-tered with anhydrite (Luhr et al. 1984; Bernard et al.1991; Matthews et al. 1992, Pallister et al. 1996), but noanhydrite is reported to be included within apatite. Thisalso indicates that apatite in S-rich volcanic rocks crys-tallized earlier than anhydrite although both apatite andanhydrite can in principle crystallize near liquidus tem-peratures (Watson 1980; Harrison and Watson 1984; Car-roll and Rutherford 1987; Luhr 1990). The crystallizationof apatite before S saturation and inhibited re-equilibra-tion with other evolving phases may also account for thedifference in S concentrations between natural apatite and

experimental apatite, the latter having a maximum valuefour times greater than the former (Tables 2 and 3).(2) One of the strongest controls on S content in apatite

is the S content of coexisting melt. If the melt has littleS, then so does the apatite. This study indicates that thepartition coefficient for S between apatite and melt de-pends strongly on temperature. The S in apatite and S inmelt have opposite relationships with temperature; that is,with increasing temperature, S in apatite tends to decreasewhereas S in the coexisting melt increases (Carroll andRutherford 1987; Luhr 1990). This opposite relationshipserves to complicate the correlation between apatite Scontents and magmatic S contents.(3) Late-stage oxidation might also occur during the


evolution of calc-alkaline magma systems. Oxygen fu-gacity strongly controls the S concentration in apatite.Magmas can contain S in various oxidation states, butapatite only incorporates significant S61. Therefore, evenif a magma is saturated with S, low ƒ prohibits S fromO2entering the apatite structure. During the evolution of amagma system, an increase in the activity of H2O resultsin the crystallization of hornblende or biotite, and an in-crease in ƒ (Frost and Lindsley 1991). The calculationsO2of Candela (1986) also demonstrated that the evolutionof vapor in magmas, especially in those with low Fe,commonly would be accompanied by an increase in mag-matic ƒ . Although pyrrhotite was reported in theO2groundmass of subduction-related calc-alkaline volcanicrocks (Ueda and Itaya 1981; Luhr el al. 1984), many suchrocks lack groundmass pyrrhotite in equilibrium withglass and contain pyrrhotite only as inclusions within sil-icate and oxide phenocrysts (Carroll and Rutherford1987). The restricted occurrence of phenocrystic pyrrho-tite in calc-alkaline volcanic rocks supports the conceptof late-stage oxidation during the evolution of subduc-tion-related magmatic systems.(4) The replacement of PO4 by SO4 in apatite involves

the coupled substitutions given by Equations 1 and 2.This implies that Na2O and SiO2 activities in the meltmight also affect the concentration of S in apatite. Stoppaand Liu (1995) proposed that silica activity in magmascontrols the S concentration in apatites from alkalinerocks. However, our results indicate that Si and Na in theMNH- and MTH-buffered apatites are controlled mainlyby the S content of apatite and that Si in the FMQ-buf-fered apatites is mainly controlled by the Si content andtemperature of the coexisting melt.

ACKNOWLEDGMENTSThe experimental charges used in this study could not have been pro-

duced without the generosity and encouragement of Tren Haselton, Ros-alind Helz, Bob Rye, Gary Cygan, Steve Huebner, Phil Bethke, I.-M.Chou, and many other USGS scientists. Thanks are also due to manypeople in the Smithsonian Institution, especially Joe Nelen, Gene Jarose-wich, and Jim Collins for assistance with electron microprobe analysesand mineral separation; Vic Avery for instruction on use of the lumino-scope; Sara Russell for help with SEM analyses; Tim Gooding for prep-aration of polished sections; and Pete Dunn and Jeff Post for kindly pro-viding comments on an earlier version of the manuscript. Mike Brewer,Jim Carboy, and Jennifer Kling assisted with this project during theirtenures as summer interns at the Natural History Museum. We are alsograteful to many individuals who aided us in obtaining samples for thisstudy: Charlie Bacon (Mazama), Bob Christiansen (Yellowstone), MichaelClynn (Lassen), John Drexler (Julcani), John Foden (Tambora), BrianHausback and Peggy Genaro (Sutter Buttes), Jose-Luis Macıas (pre-1982El Chichon), Steve Matthews (Lascar), Floyd McCoy (Santorini), BillMelson (St. Helens and Santa Marıa), Marc Monaghan (Bogoslof ), SteveNelson (Ceboruco), Chris Newhall (Pinatubo), Paul Wallace, and Wes Hil-dreth (Novarupta and Long Valley). We also thank John Pallister for manyinsightful discussions on the significance of S in volcanic apatite, andPhillip Piccoli and Malcolm Rutherford, whose reviews significantly im-proved the manuscript. This study was supported by NASA’s Volcano-Climate Interaction Program (1438-VCIP-007) and by the Research Ini-tiatives Program of the Smithsonian’s National Museum of NaturalHistory.

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MANUSCRIPT RECEIVED JANUARY 8, 1997MANUSCRIPT ACCEPTED JULY 28, 1997

APPENDIX 1: SAMPLE DESCRIPTIONS ANDREFERENCES (IN PARENTHESES) USED FOR

ESTIMATING PRE-ERUPTION ƒ ,O2TEMPERATURE, AND PRESSURE

Samples containing anhydriteEl Chichon, Mexico (Luhr et al. 1984):CH70-1: trachyandesite pumice erupted on March 28–29, 1982.

CH70-2: trachyandesite pumice erupted on March 28–29, 1982.

EC600: pumice from a 600 year old pumice-falldeposit.

EC900: pumice from a 900 year old pyroclastic-flowdeposit.

EC1400: pumice from a 1400 year old block-and-ash-flow deposit.

Pinatubo, Philippines (Rutherford and Devine 1996):NMNH no. 116534-1 and NMNH no. 116534-2: pum-ices from a June 15, 1991 pyroclastic-flow deposit.

Lascar, Chile (Matthews et al. 1994):LA135: 2-pyroxene andesitic lava of Center I.Sm93/22: pyroxene-hornblende dacite of PiedrasGrandes.

LA121: 2-pyroxene dacite pumice of Soncor.LA122: 2-pyroxene dacite pumice of Soncor.LA147: 2-pyroxene dacite pumice of Soncor.LA140: 2-pyroxene dacite lava of Capricorn.LA141: hornblende-rich mafic inclusion.LA123: 2-pyroxene andesite pumice of Tumbres Flow.LA102: 2-pyroxene andesite lava erupted February,1990.

Sm94/5: 2-pyroxene andesite pumice erupted April,1993.

Sm94/6: banded pyroxene pumice erupted April, 1993.

Julcani, Peru (Drexler and Munoz 1985):J8-1B-25: dacitic dome, stage II.LT79.9-20: Tenta Dora dike, stage III.J8-10A-14: Bulolo dike, stage IV.

Sutter Buttes, U.S.A.:SB BH-1: hornblende-biotite andesite from surface.UCB-103-14-1 and UCB-103-14-2: anhydrite-bearinghornblende-biotite andesite drillcore taken fromsouth side of Sutter Buttes in 1949. Univ. Californiapetrology teaching collection, 103A, spec. no. 14.

Cerro Lanza, Mexico:M83-18: hornblende trachyandesite.

Samples lacking obvious anhydriteEtna, Italy:NMNH no. 115482-4: mugearite from Poggio Laca.NMNH no. 100021: andesite from Grotta DellePalombe.

NMNH no. 100074: andesite from Mt. Rosso.

Santorini, Greece (Huijsmans 1985):NMNH no. 98696: lava from Mikra Kameni dome.NMNH no. 98697: lava from Nea Kameni dome.NMNH no. 115999: dacitic pumice from ashflow unitB03 (about 1450) from Phira Quarry.

Krakatau, Indonesia (Mandeville et al. 1996):NMNH no. 35516: andesitic pumice collected floatingin water soon after 1883 eruption.

Tambora, Indonesia (Foden 1986):NMNH no. 116570: pumice from 1815 eruption.

Bogoslof, U.S.A.:BogI and Bog II: pumice from 1796 eruption.


Novarupta, U.S.A. (Hildreth 1983; Avery 1992):NMNH no. 116660-3: andesitic pumice from 1912eruption.

NMNH no. 116660-4: dacitic pumice from 1912eruption.

NMNH no. 116660-5: rhyolitic pumice from 1912eruption.

Lassen, U.S.A. (Carmichael 1967):NMNH no. 116619-53: banded pumice from May 22,1915 eruption.

NMNH no. 116619-54: dacite of Lassen Peak.Mazama, U.S.A. (Druitt and Bacon 1989):NMNH no. 116619-56: rhyodacitic pumice from py-roclastic-flow on north side of Rouge River alongHwy. 230.

NMNH no. 116619-59: andesitic scoria from the Pin-nacles pyroclastic-flow deposit.

St. Helens, U.S.A. (Rutherford et al. 1985):NMNH no. 115379-7: pumice from May 18, 1980eruption.

NMNH no. 115379-15: pumice from pyroclastic-flowdeposit erupted on July 22, 1980.

Long Valley, U.S.A. (Hildreth 1977):

BT: rhyolitic pumice from the early fall unit of theBishop Tuff.

Yellowstone, U.S.A. (Hildreth 1981; Hildreth et al. 1991):MFT: rhyolitic pumice from fall unit of the Mesa FallsTuff.

Ceboruco, Mexico (Nelson 1979):983-3: 1870 dacite from the toe of small bulbous daciteflow.

983-106: second-stage dacite from the Copales lavaflow.

983-117: post-caldera andesite from the Coapan lavaflow.

Popocatepetl, Mexico:Popo-2: massive andesite lava on upper NW flank.Popo-3: andesite lava from northern crater rim.Popo-4: andesite lava on upper E flank, W of Las Cru-ces hut.

Popo-5: andesite lava on upper NE flank, at base ofsteep hill below Las Cruces hut.

Santa Marıa, Guatemala:NMNH no. 113100-9 and NMNH no. 113100-10: pum-ices from the 1902 eruption.

factors controlling sulfur concentrations in volcanic apatite · factors controlling sulfur...

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