Icarus 212 (2011) 24–41
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Icarus
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On the chemistry of mantle and magmatic volatiles on Mercury ⇑ Mikhail Yu. Zolotov
School of Earth and Space Exploration, Arizona State University, Tempe, AZ 85287-1404, USA article info abstract
Article history: The surface of Mercury contains ancient volcanic features and signs of pyroclastic activity related to Received 21 November 2009 unknown magmatic volatiles. Here, the nature of possible magmatic volatiles (H, S, C, Cl, and N) is dis- Revised 8 November 2010 cussed in the contexts of formation and evolution of the planet, composition and redox state of its mantle, Accepted 19 December 2010 solubility in silicate melts, chemical mechanisms of magma degassing, and thermochemical equilibria in Available online 25 December 2010 magma and volcanic gases. The low-FeO content in surface materials (<6 wt%) evaluated with remote
sensing methods corresponds to less than 2.3 fO2 log units below the iron–wüstite buffer. These redox Keywords: conditions imply restricted involvement of hydrous species in nebular and accretion processes, and a lim- Mercury, Interior ited loss of S, Cl, and N during formation and evolution of the planet. Reduced conditions correspond to Mercury, Surface Volcanism high solubilities of these elements in magma and do not favor degassing. Major degassing and pyroclastic Cosmochemistry activity would require oxidation of melts in near-surface conditions. Low-pressure oxidation of graphite in moderately oxidized magmas causes formation of low-solubility CO. Decompression of reduced N-sat-
urated melts involves oxidation of nitride melt complexes and could cause N2 degassing. Putative assim-
ilation of oxide (FeO, TiO2, and SiO2) rich rocks in magma chambers could have caused major degassing through oxidation of graphite and S-, Cl- and N-bearing melt complexes. However, crustal rocks may never have been oxidized, and sulfides, graphite, chlorides, and nitrides could remain in crystallized
rocks. Chemical equilibrium models show that N2, CO, S2,CS2,S2Cl, Cl, Cl2, and COS could be among the most abundant volcanic gases on Mercury. Though, these speciation models are restricted by uncer- tain redox conditions, unknown volatile content in magma, and the adopted chemical degassing mechanism. Ó 2010 Elsevier Inc. All rights reserved.
1. Introduction: Volcanic features and composition of surface (Robinson and Lucey, 1997; Murchie et al., 2008; Robinson et al., materials 2008; Ernst et al., 2010). The ground mid-infrared emission spectra
suggest silicate surface materials with spatially variable SiO2 con- Morphological and color analyses of Mariner 10 images suggest tent and dominant mafic to ultramafic compositions (Boynton a volcanic origin of some surface features on Mercury (Strom et al., et al., 2007). These interpretations are consistent with high heights 1975; Basaltic Volcanism, 1981; Spudis and Guest, 1988; Robinson of lobate flow fronts indicating high-volume eruptions of mafic and Lucey, 1997). The presence of volcanic rocks is confirmed by lava similar to terrestrial flood-basalts (Wilson and Head, 2008). the initial MESSENGER results (Head et al., 2008, 2009a; Robinson Some regions (e.g., the inner part of Caloris basin) could be repre- et al., 2008; Murchie et al., 2008; Prockter et al., 2010). The volca- sented by evolved igneous rocks with intermediate compositions, nic features demonstrate mostly effusive lunar-like activity with- as suggested from mid-infrared spectra (Sprague et al., 2002, out stratovolcanoes or large sheet structures (Head et al., 2008, 2009; Boynton et al., 2007). Intermediate compositions may repre- 2009a). However, several features could be interpreted as pyroclas- sent rocks formed through upward fractional crystallization of a tic deposits (Head et al., 2008, 2009a; Murchie et al., 2008; Robin- putative magma ocean (Brown and Elkins-Tanton, 2009)or son et al., 2008; Kerber et al., 2009; Prockter et al., 2010). through differentiation of impact melts and basaltic magma in Spectroscopic data in visible, near- and mid-infrared ranges chambers. Although magma chambers are suggested from some indicate the compositional diversity of the Mercury’s surface images (Head et al., 2009b), there are no morphological signs of (e.g., Sprague et al., 2009; Denevi et al., 2009; Vernazza et al., volcanic activity associated with evolved magmas (Head et al., 2010). Suggested volcanic plains and pyroclastic deposits are 2008, 2009a). brighter and redder than underlying rocks and impact ejectae Visible, near- and mid-infrared spectra of Mercury’s surface are consistent with the presence of Mg-rich olivine and pyroxene, Mg- ⇑ Fax: +1 480 965 8102. and Ca-clinopyroxene, plagioclase feldspars, silicate glasses, and E-mail address: [email protected] sulfides (Sprague et al., 1995, 2002, 2009; Warell et al., 2006,
0019-1035/$ - see front matter Ó 2010 Elsevier Inc. All rights reserved. doi:10.1016/j.icarus.2010.12.014 M.Yu. Zolotov / Icarus 212 (2011) 24–41 25
2010; Boynton et al., 2007; Vernazza et al., 2010), though interpre- The identity of the volatiles that caused the pyroclastic erup- tations are model-dependent and surface mineralogy remains tions is enigmatic. Kerber et al. (2009) showed that abundant mag- uncertain. The lack of spectral features in visible and near-infrared matic gases (up to 1.3 wt% depending on molecular mass) are ranges can be accounted for by the presence of low-albedo opaque needed to form the large pyroclastic deposit at Caloris basin. They phase(s), which could represent sulfides and/or Ti-bearing oxides suggested Earth-like volatiles (H2O, CO2, CO, H2S) along with abun- (Rava and Hapke, 1987; Denevi and Robinson, 2008; Denevi dant S gases to explain the eruptions. This paper discusses the ele- et al., 2009; Vernazza et al., 2010). The detection of Mg in the exo- mental abundances of H, S, C, Cl, and N in Mercury’s silicate sphere of Mercury (McClintock et al., 2009) agrees with the identi- interior, the composition of possible magmatic volatiles, and fication of Mg-silicates by ground-based spectroscopy, although chemical aspects of degassing. Then, thermochemical calculations Mg sulfides could be present as well (Sprague et al., 1995). The are used to estimate volcanic gas speciation at appropriate bulk presence of Ca and alkalis in the exosphere (e.g., McClintock compositions, redox conditions (fO2), temperatures (T), and pres- et al., 2009) also agrees with the identification of Ca-pyroxenes sures (P). and feldspars in surface materials. Surface spectra of Mercury in near- and mid-infrared are distinct from lunar spectra, probably 2. Volatiles in the interior indicating a fundamental difference in abundance and composition of ferrous silicates (Vernazza et al., 2010). Observational data on Mercury’s volatiles are limited to exo- Although the abundance and speciation of iron in surface mate- spheric Na, K, and possibly, S- and H-bearing ionized molecules rials remains uncertain, the majority of remote-sensing data indi- formed through sputtering of surface materials (Potter and Mor- cate a low-FeO content. The lack of 1 lm absorption in most of gan, 1985, 1986; Zurbuchen et al., 2008). The abundance and spe- Mercury’s spectra, the color of the surface, and microwave emissiv- ciation of volatiles in the interior would reflect the composition ity data indicate an FeO-depleted composition (<3–6 wt% FeO) and redox state of parent planetesimals, and specifics of accretion, (e.g., McCord and Clark, 1979; Vilas, 1988; Blewett et al., 1997; impact re-processing, differentiation, and initial degassing. The Robinson and Taylor, 2001; Boynton et al., 2007). Mid-infrared pathways of formation and evolution of Mercury remain poorly ground emission spectra do not reveal FeO-rich compositions constrained, although diverse scenarios have been suggested, as (Sprague et al., 2002, 2009; Boynton et al., 2007; Vernazza et al., reviewed by Solomon (2003), Taylor and Scott (2003), Benz et al. 2010). These ground observations are consistent with the interpre- (2008), and Breuer et al. (2008). Most of these models are aimed tations of Mariner 10 color data (Robinson and Lucey, 1997; Denevi at explaining a high core/mantle mass ratio and the observed and Robinson, 2008). Initial MESSENGER results also do not indi- FeO deficiency in surface materials. In this section, we discuss vol- cate a presence of abundant surface ferrosilicates (Solomon et al., atiles in the context of formation and evolution pathways. 2008; McClintock et al., 2008; Robinson et al., 2008). The interpretation of MESSENGER neutron spectrometer data obtained during three flybys indicates elevated (lunar-mare type) 2.1. Bulk composition in the framework of formation and evolution concentrations of Fe and Ti, though phase composition cannot be scenarios obtained (Lawrence et al., 2010). These authors suggest an ilmenite content of 7–18 wt% in surface materials. Lawrence et al. (2010) The high density of Mercury (5.43 g/cm3) implies a high Fe/Si also proposed early crystallization of ilmenite in oxidized lunar- ratio and a significant fractionation of materials that originated type magmas (cf., Usselman and Lofgren, 1976) that may lead to with solar Fe/Si ratio. The fractionation could have been associated late crystallization of FeO-poor pyroxenes. Alternatively, a pres- with nebular condensation under specific conditions, chemical gra- ence of Fe in surface sulfides (e.g., FeS, Sprague et al., 2009; Vern- dients in the nebula, differential re-condensation related to chon- azza et al., 2010) could also be consistent with the low-FeO content drule-forming and impact events, aerodynamic separation of in silicates (Blewett et al., 2009). In the latter case, Ti could be pres- particles, and preferential evaporation and/or impact removal of ent in less oxidized minerals than ilmenite. silicates. The latter two processes could have involved km-size The presence of Fe-bearing oxides (magnetite, ulvospinel, planetesimals, Moon- to Mars-sized planetary embryos, and Mer- ilmenite) in Mercury’s rocks seems unlikely because their forma- cury itself. tion requires oxidized conditions that are inconsistent with FeO- Thermodynamic models for condensation of the solar nebular poor silicates (Riner et al., 2010, this paper). In reduced (FeO-poor) gas (e.g., Grossman, 1972; Grossman and Larimer, 1974; Ebel, magmatic conditions, Ti could be present in Mg-rich geikielite– 2006) demonstrate that high temperature condensates ( 1300– ilmenite solid solution (MgTiO3–FeTiO3)(Riner et al., 2010), rutile 1450 K) have an elevated metal/silicate ratio compared to low-T (TiO2), perovskite (CaTiO3), or even sulfides. Although ilmenite condensates. Preferential accumulation these condensates in has been suggested as an opaque spectrally neutral phase to ac- high-T inner nebula could have accounted for a high Fe/Si ratio count for low reflectance in visible and near-infrared ranges (Rava of planetesimals that formed Mercury. High gas (H2) pressure in and Hapke, 1987; Robinson et al., 2008; Denevi et al., 2009; Vern- the innermost nebula would also facilitate early high-T condensa- azza et al., 2010), it does not provide a good match with near-infra- tion of Fe-metal compared to Mg-silicates (Grossman, 1972; Lewis, red emission spectra (Sprague et al., 2009). In contrast, rutile could 1972). Subsequent chondrule-forming events caused by gas shock be present in surface materials based on interpretation of mid- waves (Desch and Connoly, 2002) could have favored metal con- infrared emission spectra (Sprague et al., 2009) and near-infrared densation at high gas densities over formation of silicates (includ- reflectance spectra (Vernazza et al., 2010). Titanium-rich sulfides ing Fe-silicates) through re-condensation. This process could have are present in enstatite meteorites (Keil, 1969) that are depleted been more efficient in the innermost nebula because of high pres- in FeO-bearing minerals (Keil, 1968). The presence of spectral fea- sure. High-T removal of silicates through evaporation (cf. Fegley tures of enstatite in Mercury’s mid-infrared spectra (e.g., Sprague and Cameron, 1987) from separate chondrule melts and/or nebular et al., 2009) does not exclude the occurrence of Ti sulfides in dust could have also contributed to a high Fe/silicate ratio on Mer- FeO-depleted surface materials. This paper (Section 5) shows that cury. Extremely reduced conditions, which may reflect elevated C/ rutile, Ti2O3, and some Ti sulfides could form in FeO-poor mag- O and H2/H2O ratios in the inner nebula (e.g., Ciesla and Cuzzi, matic conditions. It is possible that a majority of Fe and Ti is in re- 2006), also support preferential formation of Feo-rich condensates duced forms that reflect redox conditions during magmatic (Larimer and Batholomay, 1979), and account for the low-FeO con- crystallization. tent in accreted materials. Note that condensation of nebular gas 26 M.Yu. Zolotov / Icarus 212 (2011) 24–41 with solar abundances also leads to reduced condensates with 1989; Keil et al., 1994) demonstrate that extremely reduced plane- minute FeO content in silicates (e.g., Ebel, 2006; Fedkin and Gross- tesimals were also involved in these processes. Independent of the man, 2006; Grossman et al., 2008). origin of metal-rich planetesimals, an aerodynamic separation of High T–P conditions in the inner solar nebula may imply a low denser and/or larger planetesimals could have led to an elevated amount of condensed volatiles (at least H, S, and alkalis). However, Feo/silicate ratio in the high-P inner nebula and may have contrib- elevated C/O and H2/H2O ratios in the inner nebula could have led uted to the formation of metal-rich Mercury (Weidenschilling, to accumulation of S, C, and N in reduced solids (e.g., Larimer and 1978). Batholomay, 1979). If Mercury’s primary materials were formed at Our working hypothesis is that Mercury formed from reduced, C/O > 0.8, the volatiles could have been condensed in phases high-T, high-P, anhydrous nebular condensates that were affected which occur in enstatite chondrites (Mg, Ca, Cr and Ti sulfides, C by removal of silicates in major collisions occurring before and/or species, and nitrides; Keil, 1968, 1989; Brearley and Jones, 1998) after formation of the planet. Although enstatite chondrites and (e.g., Larimer, 1975; Larimer and Batholomay, 1979; Grossman CB-like chondrites provide insights about reduced nebular solids et al., 2008). This notion and the low FeO in enstatite meteorites and metal–silicate fractionation through impacts, it is possible that led to the suggestion of significant contribution from enstatite none of these groups of meteorites represent the composition of chondrite (EC) material to the bulk composition of Mercury (Was- Mercury. In particular, FeO-poor and Feo-rich condensates of the son, 1988; Sprague et al., 1995). Enstatite chondrites are isotopi- solar gas composition could be among Mercury’s protoliths, though cally identical to Earth and could be similar to the dominant they do not contain minerals found in enstatite chondrites. Like- constituent of our planet (Javoy et al., 2010). A deficiency of H2O wise, the presence of hydrated carbonaceous clasts in CB chon- gas in the inner nebula caused by diffusion of water vapor toward drites (Bonal et al., 2010) and their oxygen isotopic composition the water condensation front (Stevenson and Lunine, 1988; Cyr (Clayton and Mayeda, 1999; Rubin et al., 2003; Ivanova et al., et al., 1998; Ciesla and Cuzzi, 2006) and/or evaporation of con- 2008) indicate their formation in the main asteroid belt. densed C-bearing materials that migrated toward the Sun (Larimer and Batholomay, 1979; Lodders and Fegley, 1993) could account 2.2. Hydrogen and water for even greater contribution of the EC-like material to the compo- ÀÁ þ þ sition of Mercury. Condensation models show that a 85% depletion Hydrogen-bearing ionized molecules H2O ; H2O; OH and + of nebular H2O gas content would lead to formation of minerals H2S have been tentatively identified in Mercury’s exosphere by found in enstatite chondrites (Hutson and Ruzicka, 2000; Pasek MESSENGER (Zurbuchen et al., 2008). These molecules could have et al., 2005). Mn–Cr isotopic systematics (Shukolyukov and Lugm- resulted from sputtering of surface materials, which may include air, 2004) and xenon isotopes (Jee-Yon et al., 2009) in enstatite water ice (Zurbuchen et al., 2008) and/or protons implanted with meteorites also indicate their formation in the hot inner solar neb- the solar wind. Water ice (Slade et al., 1992) may be associated ula sunward of the asteroid belt. However, the Fe/Si ratio in EC is with radar-bright craters in the north polar region (Harmon and similar to solar abundances and a metal–silicate separation is re- Slade, 1992; Harmon et al., 2001) and could represent molecules quired to form Mercury from reduced condensates, which may or trapped after cometary/meteoritic impacts (Moses et al., 1999). may not contain minerals found in enstatite chondrites. Although endogenic sources of Mercury’s H are not excluded,
Disruptive collisions of planetesimals and differentiated plane- the interior of Mercury is likely to be depleted in H and H2O. The tary embryos (e.g., Wetherill, 1986; Asphaug et al., 2006) would low-FeO content in surface silicates and suggested FeO-poor man- provide an efficient mechanism for metal–silicate separation. tle (Robinson and Taylor, 2001) are indicators of water scarcity that Numerical simulations of Benz et al. (1988, 2007) demonstrate that also implies reduced conditions. Water is the major oxidizing agent the density of Mercury could be accounted for by stripping of its in the solar nebula and inside asteroids and planetary bodies (Le- silicate mantle by major impact(s). Dynamical models show that wis and Prinn, 1984; Dreibus and Wänke, 1989; McSween and Lab- major collisions were more frequent within 2 AU from the Sun, otka, 1993; Arculus and Delano, 1987; Kasting et al., 1993; Ciesla and that iron meteorites could have formed through impact dis- and Cuzzi, 2006), and reduced rocks often reflect processes that oc- ruptions of differentiated bodies in the terrestrial planet region curred under deficiency of water. The low-FeO content in surface (Bottke et al., 2006). High relative velocities of planetesimals and silicates suggests a small (if any) contribution from processes such planetary embryos at low AU should have led to strong collisions as aqueous alteration of parent planetesimals, accretion of icy and o (e.g., Wetherill, 1986). Rapid accretion of planetesimals at low hydrated planetesimals, and Fe –H2O interaction in the interior. AU also favored their fast differentiation (through decay of 26Al, Outward diffusion of nebular water vapor and a limited pene-
Grimm and McSween, 1993) and made metal–silicate separation tration of H2O-bearing planetesimals into the inner solar nebula more efficient. In addition to silicates, some volatiles could have could have created H2O-deficient conditions (e.g., Stevenson and been lost through decompression and boiling of silicate fluids in Lunine, 1988; Ciesla and Cuzzi, 2006). Together with high temper- collisions (Melosh et al., 2004; Asphaug et al., 2006) followed by ature, these conditions would prevent gas–solid type hydration rapid quenching before re-condensation of volatiles. (Lewis, 1972, 1988). A metal–silicate separation through major col- The metal-rich (60–70 vol.%) CB carbonaceous chondrites (e.g., lisions (Benz et al., 1988, 2007) implies a loss of H and water Weisberg et al., 2001; Krot et al., 2003) could have formed through (Asphaug et al., 2006). Even if some water had accreted on Mer- collisions of planetary embryos (Campbell et al., 2002; Rubin et al., cury, it would have been consumed in oxidation reactions with 2003; Krot et al., 2005). CB chondrites are strongly depleted in Feo-metal (e.g., Dreibus and Wänke, 1989), elemental C (Holloway moderately volatile lithophile elements (Na, K, Mn, Zn, Se, etc.) and Jakobsson, 1986; Kadik et al., 2006) and other reduced solids and have high abundances of lithophile refractory elements (Weis- (Fe3P, MgS, etc.) leading to the formation and escape of H2 and berg et al., 2001; Krot et al., 2003). They have low-FeO contents in CO. If Mercury’s mantle formed through upward crystallization of silicates (<4 mol.%), and the high Feo/silicate ratio makes them a magma ocean (Brown and Elkins-Tanton, 2009), water might plausible analogs for Mercury’s bulk composition (Taylor and Scott, have been expelled toward the surface and partially degassed at 2003). Several other metal-rich meteorites (CH chondrites, palla- low-P conditions. Low solubility of water in Feo-metal saturated sites, ureilites, mesosiderites, and IVA iron meteorites) with signs magmas (Kadik et al., 2004 and refs. therein) would also favored of rapid cooling could have formed through major collisions as well water loss. Finally, possible impact removal of the upper part of (Keil et al., 1994; Yang et al., 2007; Scott, 2007). Signs of a major, the mantle (Benz et al., 1988, 2007) would have led to major loss early impact in the Shallowater enstatite achondrite (aubrite) (Keil, of water and H. M.Yu. Zolotov / Icarus 212 (2011) 24–41 27
The discovery of H-bearing glasses (4–46 ppm H O; Saal et al., 2 -8 o 2008) and phosphates in lunar basalts (Boyce et al., 2010) implies IW: 2Fe0.947O = 1.894Fe + O2 IRI: 2FeTiO = 2TiO + 2Feo + O IW that some water either survived a Moon-forming collision or was 3 2 2 -10 o trapped after the event. The estimated pre-eruptive water content ITOI: 2FeTiO3 = Ti2O3 + 2Fe + 1.5O2 RTO: 4TiO = 2Ti O + O in lunar glasses ( 700 ppm; Saal et al., 2008) could be considered 2 2 3 2 -12 M-A as an upper limit for water content in Mercury’s rocks because IRI , bar eO melt is enriched in water compared to mantle source rocks, and be- 2 F wt% CB/CH O 6 cause of the possible water-bearing nature of proto-Earth. f -14 eO % F log t 2 w EC ITOI 2.3. Mantle oxidation state -3 -16 IW RTO Oxygen fugacity (fO2) of lunar mare basalts (18–21 wt% FeO; -6 IW Basaltic Volcanism, 1981)is0to 2 log units below the iron–wüs- -18 tite buffer (IW, Eq. (1) in Table 1)(Wadhwa, 2008). The lesser FeO 1200 1300 1400 1500 1600 1700 1800 content on Mercury compared to lunar values implies more re- Temperature, K duced conditions. For 2–6 wt% FeO in silicate rocks, Malavergne Fig. 1. Oxygen fugacity at reduced conditions that corresponds to some mineral et al. (2010) estimated Mercury’s mantle fO2 as 3–6 log units below buffers and suggested analogs for Mercury’s mantle melts. The dotted curves refer IW (IW 3toIW 6) (Fig. 1). They assume ideal melt in which FeO to conditions of 3 and 6 log fO units below the iron–wüstite buffer (IW, Eq. (1) in o 2 activity is equal to concentration in equilibrium with Fe -metal. Table 1). IRI, ilmenite–rutile–iron buffer (Eq. (2)); ITOI, ilmenite–titanium oxide– Consideration of the non-ideality of silicate melts saturated iron buffer (Eq. (3)); RTO, rutile–titanium oxide buffer (Eq. (4)). The gray curves with Feo-metal using empirical equations from Ariskin et al. represent melts with 2 wt% and 6 wt% FeO based on the Ariskin et al. (1993) melt (1993) leads to more definitive fO evaluations. These equations model for 1500–1700 K. The curves marked CB/CH and M-A correspond to partial 2 melts of CB/CH chondrites and Mercury’s mantle composition from Morgan and describe relationships between fO2, FeO molar content, and con- Anders (1980), respectively. The circle symbol refers to a partial melt of an enstatite centrations of other major oxides in melts. For plausible magma chondrite. formed through partial melting of Mercury’s mantle compositions
Table 1 Chemical equilibria and their equilibrium constants.
Reaction Log K at 1 bar 298 K 1500 K 1700 K
o (1) 2Fe0.947O (wüstite) = 1.894Fe +O2(g); IW buffer 85.8 11.6 9.53 o (2) 2FeTiO3 (ilmenite) = 2TiO2 (rutile) + 2Fe +O2(g); IRI buffer 93.6 12.9 10.7 o (3) 4/3FeTiO3 = 2/3Ti2O3 + 4/3Fe +O2(g); ITOI buffer 103 14.1 11.6
(4) 4TiO2 = 2Ti2O3 +O2(g); RTO buffer 121 16.4 13.2
(5) 2MgS + O2(g) = 2MgO + S2(g) 65.8 11.9 10.3
(6) 2CaS + O2(g) = 2CaO + S2(g) 32.6 5.90 5.1
(7) 2MnS + O2(g) = 2MnO + S2(g) 36.6 6.42 –
(8) TiS2 +O2(g) = TiO2 +S2(g) 72.5 14.1 –
(9) TiS2 + 0.75O2(g) = 0.5Ti2O3 +S2(g) 42.2 10.0 –
(10) 2TiS + 1.5O2(g) = Ti2O3 +S2(g) 139 23.6 –
(11) 2MgS + 2SiO2 +O2(g) = 2MgSiO3 +S2(g) 77.3 14.1 12.4
(12) 2MgS + SiO2 +O2(g) = Mg2SiO4 +S2(g) 76.2 13.8 12.0
(13) 2CaS + 2SiO2 +O2(g) = 2CaSiO3 +S2(g) 62.9 12.1 10.6
(14) 2CaS + SiO2 +O2(g) = Ca2SiO4 +S2(g) 55.4 10.8 9.57
(15) 2MnS + 2SiO2 +O2(g) = 2MnSiO3 +S2(g) 45.5 8.03 6.93
(16) 2MnS + SiO2 +O2(g) = Mn2SiO4 +S2(g) 45.1 7.18 6.66
(17) 2TiS2 = 2TiS + S2(g) 54.3 3.69 – o a (18) 2FeS = 2Fe +S2(g) 50.0 4.8 – o (19) MgS + FeO = MgO + Fe + 0.5S2(g) 11.1 0.08 0.26 o (20) MgS + FeTiO3 = MgO + TiO2 +Fe + 0.5S2(g) 13.9 0.52 0.23 o (21) TiS2 + 2FeTiO3 = 3TiO2 + 2Fe +S2(g) 21.1 1.14 – o (22) 1.5MgS + FeTiO3 = 0.5Ti2O3 + 1.5MgO + Fe + 0.75S2(g) 27.7 1.65 0.96
(23) MgS + 2TiO2 =Ti2O3 + MgO + 0.5S2(g) 27.7 2.73 1.46
(24) Mg2SiO4 + SiO2 = 2MgSiO3 1.16 0.31 0.36 o (25) SiO2 =Si +O2(g) 150 22.4 18.7 (26) Co + FeO = Feo + CO(g) 20.0 2.45 3.09 o (27) C + 0.5O2(g) = CO(g) 24.0 8.47 7.99 o o (28) C + FeTiO3 =Fe + TiO2 + CO(g) 22.8 2.0 2.62 o o (29) C + 0.667FeTiO3 = 0.667Fe + 0.333Ti2O3 +CO 27.4 1.43 2.21 o (30) C + 2TiO2 =Ti2O3 +CO 39.6 0.27 1.39 o o (31) C + SiO2 =Si + 0.5O2(g) + CO(g) 126 14.0 10.7
(32) TiN + O2(g) = TiO2 + 0.5N2(g) 102 16.6 14.2
(33) 2TiN + 1.5O2(g) = Ti2O3 +N2(g) 143 25.1 21.7
(34) 2NH3(g) = N2(g) + 3H2(g) 5.75 8.41 8.87
(35) H2O(g) = 0.5O2(g) + H2(g) 40.0 5.73 4.67
Notes: The equilibrium constants are calculated using thermodynamic properties of corresponding solids, liquids, and gases mainly from Gurvich et al. (1989–1994), Pankratz
(1987), Robie and Hemingway (1995), and Barin (2004). Although O2 gas is not present in magmas, it is used to formally quantify fO2 at chemical equilibria. In reality, some of o o the species (FeO, MgS, SiO2, etc.) could be components of magma, a phase or a solid solution. In all equations, C ,Fe , TiO2, FeTiO3, and MgSiO3 represent graphite, iron metal, rutile, ilmenite, and enstatite, respectively. a Extrapolated from T < 1400 K. 28 M.Yu. Zolotov / Icarus 212 (2011) 24–41
(Taylor and Scott, 2003), concentrations of 2–6 wt% FeO result in the redox conditions of IW 0.5 imply a significant oxidation of fO2 of IW 2.3 to IW 3.4 at 1600–1700 K (Figs. 1 and 2). Specifi- the innermost solar nebula through incorporation of water-bearing cally, the redox state of a partial melt of the Indarch (EH4) enstatite planetesimals migrated from larger heliocentric distances (Ciesla chondrite at 1700 K with 0.25 wt% FeO (McCoy et al., 1999; Taylor and Cuzzi, 2006; Grossman et al., 2008). Although this migration and Scott, 2003) corresponds to IW 5.3 (log fO2 = 14.9). Melt de- could explain elevated oxidation states of some materials (e.g., rived from a mantle composition consisting of the silicate compo- Type II chondrules with high FeO content) formed in the asteroid nents of metal-rich CB/CH-like chondrites (Krot et al., 2001) with belt, these planetesimals are unlikely to have survived high-T con- 3.2 wt% FeO (Taylor and Scott, 2003) corresponds to IW 3 ditions in the innermost nebula. Considering limited contribution
(log fO2 = 13.5) at 1600 K, and IW 3.1 (log fO2 = 12.6) at from water-bearing bodies to the composition of Mercury (O’Brien 1700 K. Similar results are obtained for melt with 3.7 wt.% FeO et al., 2006), a reduced mantle seems likely. Metal–silicate frac- (Taylor and Scott, 2003) derived from the mercurian mantle com- tionation through collisions, gas drag, and igneous differentiation position of Morgan and Anders (1980):IW 2.9 (log fO2 = 13.4) at (see Section 1) would not have changed the reduced nature of 1600 K, and IW 3 (log fO2 = 12.5) at 1700 K. materials that formed Mercury. These estimates show that fO2 in Mercury’s magmas could be close to the conditions determined by mineral equilibria that in- 2.4. Sulfur, carbon, chlorine, nitrogen, and alkalis clude Ti(IV) and Ti(III) oxides (ITOI and RTO buffers, Figs. 1 and 2). The conditions of the IRI buffer (Eq. (2) in Table 1) correspond 2.4.1. Sulfur to high FeO melt contents that are not consistent with remote Estimations of Mercury’s bulk S content strongly depend on the sensing observations of surface materials (Fig. 2) (except, possibly, formation scenario. Accretion of Mercury from EC-like materials Lawrence et al., 2010, interpretation). In other words, the presence would lead to solar abundances of S (Table 2). In addition to FeS, of ilmenite contradicts low FeO concentration in silicates, and pure Mg-, Ca- and Cr-sulfides are the likely products of nebular pro- ilmenite may not be present in FeO-poor Mercury’s rocks (Riner cesses in H2O-depleted inner nebula (Cyr et al., 1999; Pasek et al., 2010). et al., 2005; Grossman et al., 2008). Conversely, formation of Mer- The possibility of crystallization of ilmenite before pyroxene at cury from metal-rich products of planetary embryo collisions could IW 0.5 (Usselman and Lofgren, 1976) could be reconciled with lead to an S-depleted planet, considering the low S content in CB the low-FeO content in surface silicates, and the lunar-type oxi- chondrites (Table 2). However, independent of the formation sce- dized mantle cannot be excluded (Lawrence et al., 2010). However, nario and redox state of accreted solids, Mercury’s silicate mantle abFeO wt % 12 5 6 7 -1 IRI, 1700 K IRI, 1500 K K 2 0 - 600 K -12 170 1 IW -2 ITOI, 1700 K K 1500 ITOI, 1500 K
K 3 0 - -3 150 W K
, bar -14 I 1700 2 4 - RTO, 1700 K O W
I f
relative to IW -4 2 O log -16 f RTO, 1500 K -5 log
-18 -6 0 5 10 15 20 0123456 FeO mole % FeO mole %
Fig. 2. Oxygen fugacity as a function FeO concentration in silicate melts. The data are obtained with empirical equations for silicate systems saturated withFeo-metal (Ariskin et al., 1993). The curves refer to magmas formed through partial melting of plausible Mercury’s mantle compositions from Taylor and Scott (2003). Solid, dashed, and dotted curves represent partial melts of an enstatite chondrite (McCoy et al., 1999), averaged silicate components of metal-rich CB/CH like chondrites (Krot et al., 2001), and a
Mercury’s mantle (Morgan and Anders, 1980), respectively. In (a), thin dashed curves show fO2 values relative the IW buffer at specific temperatures and melt compositions.
In (b), the data are presented relative to the IW buffer for low-FeO contents. The dotted lines refer to fO2 buffers depicted in Fig. 1. The upper horizontal axis corresponds to wt% FeO in the melt derived from metal-rich chondrites.
Table 2 Abundances of elements in some chondrites and aubrites.
S, wt% C, wt% Cl, ppm N, ppm CI carbonaceous chondrites 5.41 3.45 700 3180 H ordinary chondrites 2.0 0.21 140 48 L ordinary chondrites 2.2 0.25 270 43 LL ordinary chondrites 2.1 0.31 200 70 EH enstatite chondrites 5.6 0.39 570 420 EL enstatite chondrites 3.1 0.43 230 240 Aubrites (enstatite achondrites) 0.24–0.86a 0.02–0.14b 1–31c 10–100b CB chondrites 0.52e 0.37e – 25–106d
Notes: Chondritic data are from Lodders and Fegley (1998). a Gibson et al. (1985). b Grady et al. (1986). c Garrison et al. (2000), without Shallowater aubrite which contains 400 ppm Cl. d Krot et al. (2003) and Ivanova et al. (2008). e Rubin et al. (2003) and Ivanova et al. (2008). M.Yu. Zolotov / Icarus 212 (2011) 24–41 29 may only contain accessory sulfides because they would likely a 10 accumulate in the core (cf., Riner et al., 2008). Besides FeS, this accumulation could have affected Mg-, and Ca-sulfides, if they EC gas 1 were present (McCoy et al., 1999). The effect of subsequent major collisions (Benz et al., 1988, 2007) on concentration of S in the mantle is not clear. 0.1 Mercury's mantle? Sun 0.01 EH L LL CI 2.4.2. Carbon Cl/S mole ratio EL H C/S Mole ratio Kilauea The solar abundance of C is an order of magnitude higher than 0.001 the abundance in CI carbonaceous chondrites and is about two or- 0.1 1 10 ders of magnitude higher than in ordinary, enstatite, and CB chon- b 10 drites (Table 2; Asplund et al., 2005; Lodders and Fegley, 1998). EC gas This corresponds to a decrease in C content in the main asteroid 1 Sun belt where carbonaceous bodies become less abundant toward the Sun (Gradie et al., 1989). The limited formation of C-bearing CI 0.1 Mercury's mantle? solids in the inner solar nebula could be related to formation of EH EL CB CO from condensed organic species transported from higher helio- 0.01 N/S mole ratio LL centric distances (AU > 5) (Cuzzi et al., 2003). Reduced nebula H L conditions have not cased accumulation of abundant graphite in 0.001 enstatite chondrites, though these meteorites are slightly enriched 0.1 1 10 20 in C compared to ordinary chondrites (Table 2). Nebular condensa- C/S mole ratio tion models of Lewis (1988) do not predict C solids formed at high T–P conditions at the heliocentric distance of Mercury. It follows Fig. 3. Normalized elemental abundances in chondritic materials, gases released that solids condensed in the vicinity of the Sun could be C depleted. from heated enstatite chondrites (EC gas, Muenow et al., 1992), and Kilauea 1918 Nebular condensates (and Mercury) could be C-rich only if con- summit eruption (Symonds et al., 1994). Data for carbonaceous (CI, CB), ordinary (H, L, LL) and enstatite chondrites (EH, EL) are from Table 2, and solar photosphere data densed at higher C/O ratios compared to enstatite chondrites, that are from (Asplund et al., 2005). The gray ovals show possible elemental ratios in is at C/O > 0.83 when high-T condensates become rich in graph- Mercury’s mantle and magmas. ite, TiC, and SiC (e.g., Larimer and Batholomay, 1979; Grossman et al., 2008). Although such enrichment could be caused by mas- sive infall of C-rich solids into the inner solar nebula (Lodders ula could have led to formation of metal halides, as exemplified by and Fegley, 1993) no meteorite rich in these phases is found. Dy- reactions namic models show that only a limited fraction of carbonaceous 2HClðgÞþNa2Oðin gas or solidsÞ!2NaCl þ H2OðgÞ planetesimals formed at AU > 2 could have accreted on Mercury (O’Brien et al., 2006). Major collisions of planetary embryos may HClðgÞþNaAlSi O ðsolidÞ!NaCl þ 0:5Al O þ 3SiO have no major effect on C abundance, as implied from similar C 3 8 2 3 2 abundance in CB and ordinary chondrites (Table 2). Although C/S þ 0:5H2OðgÞ ratio in Mercury could vary from that in enstatite and carbona- These interactions of HCl(g) could have led to initial Cl abundances ceous chondrites (Fig. 3), lower ratios similar to EC seem probable. on Mercury as high as in CI chondrites and the Sun. During differentiation, accreted organic carbon could have been Highly metamorphosed EL6 chondrites are depleted in Cl com- lost as CO and hydrocarbon gases formed through thermal process- pared to EH4 chondrites (Rubin and Choi, 2009). One can suggest ing, oxidation, and graphitization of organics. However, in anhy- some loss of Cl during thermal/impact metamorphism after accre- drous high-P igneous systems, the majority of C could have not tion of Mercury and/or planetary embryos. Chlorine and alkali ele- degassed because of precipitation of low-solubility graphite (see ments could have been accumulated in the upper horizons of Section 4). Independent of the origin of Mercury and bulk C con- Mercury through magmatic differentiation possibly involving up- tent, a major portion of accreted C could be sequestered in the me- ward crystallization of a magma ocean (Brown and Elkins-Tanton, tal core. Although reducing conditions in the mantle may imply a 2009). Subsequently, these elements could have been removed stable existence of native C species, Mercury’s silicate rocks and through impact stripping. Despite the uncertainties, a Cl/S mole ra- magmas may not be enriched in C compared to other terrestrial tio of 0.01 characterizes many solar system objects (Fig. 3a) and planets. could be used a proxy for Mercury’s mantle rocks and magmas.
2.4.3. Chlorine 2.4.4. Nitrogen The abundance of Cl in Mercury’s mantle rocks and magmas is Chondrites are strongly depleted in N compared to cometary hard to constrain because of the unknown bulk planetary content materials (Jessberger, 1999; Bockelée-Morvan et al., 2004) and so- and limited solubility data for dry reduced melts. On the one hand, lar abundances (Asplund et al., 2005). This indicates a major devol- a low Cl content in ordinary chondrites compared to CI chondrites atilization of N-bearing compounds (organics, CHON particles, and
(Table 2) and solar abundances implies Cl loss in high-T processes NH3 hydrate) in the inner solar nebula. The position of Mercury in in the inner nebula and thermally processed planetesimals. The the Solar System implies a limited delivery of N in organic-bearing secondary nature of all Cl phases in ordinary and carbonaceous planetesimals and planetary embryos (O’Brien et al., 2006). How- chondrites also suggests a low-T (<200 K) condensation of Cl (as ever, Mercury’s rocks may not be depleted in N compared to other HCl hydrates, Zolotov and Mironenko, 2007) which may account terrestrial planets because of possible accretion of refractory for Cl depletion in the inner Solar System. On the other hand, re- nitrides (TiN, Si3N4) and N-bearing metal alloys condensed from duced conditions in the inner nebula could have favored high-T reduced nebular and/or impact-generated gases. In enstatite chon- formation of halides consistent with elevated Cl abundances in drites, the presence of reduced N solids (Keil, 1968, 1989; Rubin EH enstatite chondrites (Table 2; Rubin and Choi, 2009). A defi- and Choi, 2009) accounts for significantly larger N content com- ciency of water vapor and elevated P–T conditions in the inner neb- pared to ordinary and some carbonaceous (CB, CV, CO) chondrites 30 M.Yu. Zolotov / Icarus 212 (2011) 24–41
(Table 2; Lodders and Fegley, 1998). The association of N with the have high FeO/MgO molar ratios ( 0.3–2, Mayne et al., 2009; McS- metal phase and its high-T ( 1000 °C) release from a CB/CH mete- ween et al., 2010) and may not be considered as close analogs for orite (Ivanova et al., 2008) may indicate trapping of degassed N Mercury’s igneous melts and cumulates. through condensation of metal from impact gases (Sugiura et al., Low-FeO content in Mercury’s surface silicates also implies a 1999). It follows that major collisions involving differentiated FeO-poor mantle (Robinson and Taylor, 2001) that would produce planetary embryos and Mercury may not have resulted in a major high-T melts. A mantle rich in refractory oxides is consistent with N loss. Abundance and speciation of N in enstatite and CB/CH chon- the non-global appearance of volcanic features and spectral signs drites could be used as a proxy for Mercury’s initial materials (Ta- of Mg- and Ca-rich silicates in surface materials (see Section 1). ble 2 and Fig. 3b). The refractory mineralogy may represent accreted Fe-depleted sil- The high stability of nitrides during metamorphism at extre- icates (enstatite, forsterite, diopside, and feldspars) and remained mely reduced conditions (Petaev and Khodakovsky, 1986; Fogel in the mantle after segregation of Feo–FeS melts into the core. et al., 1989) could have prevented major loss of N before the igne- Alternatively, the mantle could have formed through upward crys- ous differentiation of Mercury and planetary embryos. The high tallization of a magma ocean leading to a FeO-rich upper layer stability of nitride complexes in reduced, H-free and high-P mag- though fractional crystallization (Brown and Elkins-Tanton, mas (see Section 4) could also have limited N degassing from a 2009). Later on, this layer could have been stripped by impact(s) putative magma ocean, as suggested for the terrestrial counterpart (Benz et al., 1988) or even sunk into the lower mantle (Brown (Libourel et al., 2003). The presence of N in meteoritic Feo alloy also and Elkins-Tanton, 2009). The putative evaporation of an early implies some sequestering of N into the metal cores on planetary mantle also results in refractory rocks (Cameron, 1985; Fegley embryos and Mercury (cf., Fegley, 1983). and Cameron, 1987). The interpretation of MESSENGER neutron spectrometer data 2.4.5. Alkalis suggest Mercury’s rocks similar to lunar mare basalts from Luna Alkalis could have not been abundant in high-T nebular conden- 24 and Luna 16 samples (Lawrence et al., 2010). These data are sates and alkali oxides may compose only a few tenths of weight inconsistent with the melt models for FeO-poor protoliths (Taylor percent of materials accreted on Mercury (Lewis, 1988). Alkalis and Scott, 2003; Burbine et al., 2002). Although neutron spectrom- could have been depleted through high-T evaporation in the inner eter data would revolutionize our views on Mercury’s magmatism, nebula (through chondrule formation and/or collisions), after they need to be confirmed with independent orbital or in situ accretion (Fegley and Cameron, 1987), or stripped by impacts to- measurements. gether with a feldspathic crust (Benz et al., 1988; Lewis, 1988). Although the presence of Na and K in the exosphere (Potter and 4. Solubility of S, C, Cl, and N in reduced anhydrous magmas Morgan, 1985, 1986; Domingue et al., 2007; McClintock et al., 2009) is consistent with feldspar-bearing surface rocks, it does The specifics of degassing during eruptions and the abundances not imply an alkali-rich mantle. Enhanced K emission over the Cal- of volatiles in magma are affected by their solubility, which is non- oris basin (Sprague et al., 1990) may indicate a low degree of melt- linearly dependent on T, P, and fO2. Reduced melts could be en- ing rather than alkali enrichment in the mantle. Note that the riched in S, Cl, and N because of the formation of non-volatile melt apparent lack of crustal subduction implies limited recycling of complexes without oxygen. However, some volatiles could be alkalis and other incompatible elements in igneous processes. A undersaturated because of limited abundances in mantle source deficiency of alkalis in the mantle also implies elevated liquidus regions. temperatures. Petrological and material science experiments performed in sil-
icate melts at fO2 below IW 2(Fincham and Richardson, 1954; Fo- 3. Temperature and composition of mantle magmas gel et al., 1996; McCoy et al., 1999; Siebert et al., 2004; Berthet et al., 2009) reveal S solubility over an order of magnitude higher Temperature and composition of magmas mainly depend on the than at more oxidizing conditions (Mavrogenes and O’Neill, uncertain mineralogy and volatile content in Mercury’s mantle. 1999; Holzheid and Grove, 2002)(Fig. 4a). The highest S solubility Petrologic experiments demonstrate that dry mantle rocks melt has been reported by Malavergne et al. (2007) (up to 10 wt% at at temperatures up to several hundred degrees higher than H2O- 10 kbar and IW 7) and Fogel et al. (1996) (7.8 wt% in CaS-satu- rich mantle substrates (e.g., Green, 1975; Righter et al., 2006). rated enstatite chondrite melts at 1 bar and 1673 K). In contrast Dry lavas on the Moon and Io are hotter than Earth’s counterparts to more oxidized systems, S solubility is higher in FeO-poor melts (Basaltic Volcanism, 1981; McEwen et al., 1998; Keszthelyi et al., and could be caused by formation of metal (Mg, Ca, Mn, Ti, etc.)
2007) and a deficiency of H2O on Mercury would contribute to ele- sulfide melt complexes (Fogel et al., 1996; McCoy et al., 1999; Fo- vated melt temperatures affecting the composition of silicate melt. gel, 2005). The anomalous S solubility in reduced melts is illus- Numerical modeling of partial melting using plausible mantle trated by the presence of 0.8–2.5 wt% S in basaltic glasses in compositions of Mercury demonstrates the formation of mafic to aubrites, which are the most S-rich samples among achondrites ultramafic melts with low-FeO contents (Taylor and Scott, 2003; and silicate planetary rocks (Fogel, 2005). These data show that S Section 2.3). Experimental melting of the Indarch (EH4) enstatite content in Mercury’s igneous materials would indicate redox con- chondrite (McCoy et al., 1999) suggests magma temperatures of ditions in magma and mantle source regions. 1670–1720 K at 20–50% partial melting. The high melt tempera- In anhydrous reduced magmas, the solubility of C is mainly con- ture reflects low abundances of FeO, alkalis, and water. The basaltic trolled by the equilibria of C-bearing solids (e.g., graphite) with CO normative composition of those melts ( 34 vol% plagioclase, 42% (Eqs. (26) and (27)). Low fO2 and temperature and high pressure fa- enstatite, 15% diopside, and 6% sulfides) has been suggested as vor the stability of graphite. The low solubilities of CO, CO2 (Pan a proxy for Mercury’s lavas (Burbine et al., 2002). Note that the use et al., 1991) and graphite, which precipitates, account for the low of enstatite achondrites (aubrites, Watters and Prinz, 1979)as solubility of C in magma (< 0.1 wt%, Holloway and Jakobsson, compositional analogs for Mercury’s igneous rocks is limited by 1986; Kadik et al., 2004). Despite possible formation of SiC melt their residual nature. Aubrites do not reveal a clear relationship complexes (Kadik et al., 2004), high-P (1–5 GPa) solubility of C with enstatite chondrites (Keil, 1989) and could have formed after in graphite-saturated systems decreases with decrease of fO2 removal of basaltic melts (e.g., Burbine et al., 2002). Despite forma- owing to precipitation of C-bearing solids (graphite, diamond, tion of HED achondrites from anhydrous reduced magmas, they carbides, metal alloys) (Holloway et al., 1992; Kadik et al., 2004; M.Yu. Zolotov / Icarus 212 (2011) 24–41 31
þ a complexes (SiON, Roskosz et al., 2006;NH2 ; NH2 ; NH3, Mysen 10 9-27 kbar data: et al., 2008). These data reveal N solubility as high as 0.7–1.6 Holzheild and Grove, 2000 wt% at pressures of 20–30 kbar (Fig. 4b), which may represent Malavergne et al., 2007 Berthet et al., 2009 mantle source regions 190–280 km below Mercury’s surface. Although the effect of pressure on N solubility in reduced 1 bar data: 1 McCoy et al., 1999 ( S Solubility, wt% log fO relatively to IW 2 rocks, which may also be N-rich. 0.1 5. Chemical pathways of magma degassing -8 -6 -4 -2 0 2 b 25 kbar 0 20 kbar Changes in pressure, composition or redox state of magma may 15 kbar affect speciation of volatiles in the melt and cause their separation -1 10 kbar 1673K 10-30 kbar in the gas phase. Here we consider degassing mechanisms of indi- 1673-1973 K -2 3 kbar vidual volatiles and discuss devolatilization of plausible Mercury’s 1 kbar 1523 K magmas with different redox states. -3 -4 5.1. Sulfur 1 bar, 1700 K 1 bar, 1573 K log N Solubility, wt% -5 Pressure does not have a definitive effect on S solubility (Ber- -6 thet et al., 2009; Fig. 4a) and decompression during ascent of re- -8 -6 -4 -2 0 2 duced magma would not cause major degassing. A decrease in solubility would cause formation of sulfide liquid(s) (McCoy log fO2 relative to IW et al., 1999; Berthet et al., 2009) rather than degassing. Major S Fig. 4. Experimental data on solubilities of S (a) and N (b) in reduced silicate melts. degassing from synthetic silicate and enstatite chondrite melts The plot (a) is a modified version of a figure from Berthet et al. (2009) and has not been reported at 1 bar (Fogel et al., 1996; McCoy et al., represents data for temperatures 1500–2300 K and pressures from 1 bar to 1999; Fincham and Richardson, 1954). Partial melting of enstatite 15 kbar. In (b), the circle dotted symbols represent data for basaltic FeO-free chondrite samples in a vacuum at 1200–1600 K (Muenow et al., magmas at 1700 K and 1 bar from Libourel et al. (2003). The solubility data at 1992) did not produce much S gases compared to C-, N-, and Cl- IW+0.9 and 1 bar is from Miyazaki et al. (2004). The high-P data for Na2O–SiO2 melts from Shilobreeva et al. (1994), Mysen et al. (2008), and Roskosz et al. (2006) species (see Section 6). A high S content in glasses in aubrites (Fo- are presented by squares, triangles, and horizontal dashed lines, respectively. gel, 2005) also does not indicate S degassing at suggested low-P Oxidation conditions in (Roskosz et al., 2006) are not well justified. The conditions conditions on a parent body (cf., Wilson and Keil, 1991). Relatively of the IW buffer correspond to Eq. (1). The plots illustrate that oxidation of magma oxidized (IW to IW 2) lunar mare basalts do not reveal signs of at fO2 a 5 b TiS -TiO MgS-En MgS-Fo 2 2 MgS-MgO 0 CaS-Woll 2 MnS-Rhd S f MnS-MnO -5 FeS-FeO CaS-CaOl MnS-Tph FeS-Feo CaS-CaO log -10 o IW buffer FeS-Fe IW buffer -15 c 5 d MgS-Fo 0 TiS -TiO MgS-En 2 2 MgS-MgO 2 S CaS-Woll CaS-CaOl f -5 MnS-Rhd o FeS-Fe MnS-MnO MnS-Tph log -10 CaS-CaO IW-3 FeS-Feo IW-3 -15 e 5 f 0 2 MgS-Fo S MgS-En TiS2-Ti2O3 f -5 MgS-MgO CaS-Woll log MnS-Rhd CaS-CaOl -10 o MnS-MnO FeS-Fe FeS-Feo MnS-Tph CaS-CaO IW-6 IW-6 -15 1000 1200 1400 1600 1800 1000 1200 1400 1600 1800 Temperature, K Temperature, K Fig. 5. Fugacity of S2 in sulfide–oxide (a, c, e; Eqs. (5)–(10)) and sulfide–silicate (b, d, f; Eqs. (11)–(16) equilibria at 1 bar total pressure and different fO2 relative to the IW buffer. The activities of sulfides, oxides, and silicates are chosen to be unity. The dashed curves refer to Eq. (18) that may control fS2 in magma. En, Fo, Woll, CaOl, Rhd, and Tph are referred to enstatite, forsterite, wollastonite, rhodonite, and tephroite, correspondingly. These equilibria provide estimates for relative fS2 over different sulfides in silicate melts. Higher fS2 values correspond to lower stability of sulfides and sulfide complexes. Decompression and/or oxidation of melts would favor sulfide to oxide/silicate conversion and formation of S-bearing gases. oxidation, decomposition of Ti and Mg sulfide melt complexes Some decomposition of sulfide melt complexes may also be could be a major reason for S degassing from initially reduced facilitated by assimilation of SiO2-rich (e.g., intermediate) rocks. melts. If this occurred, an increase in SiO2 activity (a) would increase fS2 Oxidation of magma on Mercury could potentially be driven by in sulfide-silicate equilibria (Eqs. (11)–(16), Fig. 6c). However, if assimilation of oxides from crustal rocks. Several equilibrium oxide aSiO2 is controlled by the forsterite-enstatite equilibrium (Eq. assemblages correspond to the redox conditions from IW to IW 6 (24)) it may not increase beyond 0.4–0.5. Without forsterite, an (Fig. 1, Eqs. (1)–(4)) and the assimilation of corresponding O-bear- increase in aSiO2 from 0.4–0.5 to 1 increases fS2 by only an or- ing phases would increase the fO2 of magma. Although oxidation der of magnitude. The small increase in aSiO2 also limits the oxida- through melting of FeO-bearing silicates is limited by their possible tion of magma by SiO2 that leads to formation of elemental Si (Eq. scarcity, assimilation of putative ilmenite-rich rocks suggested in (25), Fig. 7). In addition, these processes could be limited by a low (Denevi et al., 2009; Lawrence et al., 2010) could oxidize magma. abundance of evolved rocks. The relatively oxidizing conditions of ilmenite-bearing assem- The discussed S degassing through oxidation could be presented blages are expressed by the fO2 values of the IRI buffer (Eq. (2), by net reaction Figs. 1 and 6). The pathways of sulfide oxidation are illustrated 2S2 ðmeltÞþO ! S ðgÞþ2O2 ðmeltÞ by reactions (20) and (21), were FeO is the oxidizing agent. The 2 2 corresponding sulfide-oxide equilibria at 1500 K give fS2 = 0.09 (Fogel, 2005). Removal of S2 gas would reduce magma, which will and 14 bar, respectively, providing unity activity of non-gaseous increase stability of sulfide melt complexes. The reduction may lim- species (Table 1, Fig. 6a and b). Magmas could also be oxidized it degassing of S-bearing species even from long-standing shallow by Ti(IV) assimilated from ilmenite and/or rutile leading to the for- magma chambers that are vulnerable for gas escape through a frac- mation of Ti(III) species (see Fig. 1 and Eqs. (3) and (4)). If oxidation tured crust. is driven by ilmenite’s Ti(IV) and Fe(II) (e.g., Eq. (22)), magma could be oxidized beyond IW 3. In this case, the oxidation limit is set 5.2. Carbon by the ITOI buffer (Eq. (3), Fig. 1). Assimilation of only Ti(IV) phases (rutile) could only cause oxidation of very reduced magmas toward Formation of CO through graphite oxidation by FeO in melt (Eq. IW 3–IW 5, depending on temperature (Fig. 1), and the oxidation (26)) has been suggested to explain pyroclastic eruptions on the limit is placed by the RTO buffer (Eq. (4)). Oxidation of magma and Moon (e.g., Sato, 1976; Fogel and Rutherford, 1995; Nicholis and sulfide complexes by Fe(II) from ilmenite (Eqs. (20) and (21))is Rutherford, 2009), parent asteroid(s) of ureilites (Warren and Kal- more efficient than reactions involving Ti(IV) reduction to Ti(III) lemeyn, 1992), and was mentioned with respect to Mercury’s vol- (Eqs. (22) and (23)). This is illustrated by higher fS2 values in sul- canism (Kerber et al., 2009). Degassing could occur when partial fide-oxide equilibria that correspond to the IRI buffer (Eq. (2)) com- pressure of CO exceeds total pressure. At elevated FeO activities, pared to buffers that include Ti(IV) reduction (Eqs. (3) and (4)) graphite oxidation is favorable at lower pressure and could occur (Fig. 6a). However, the possible lack of FeO-rich ilmenite in Mer- during magma ascent and in shallow magma chambers. cury’s rocks (Riner et al., 2010) limits the oxidation potential of this Elevated fO2 and temperature correspond to higher fCO in the mineral. graphite-CO equilibrium (Eq. (27), Fig. 8) and favor degassing. M.Yu. Zolotov / Icarus 212 (2011) 24–41 33 -4 a 0 SiO = Si + O 1500 K 1700 K aSiO 2 2 -6 2 = 1 -1 IRI 0.1 ITOI TiS2 RTO -8 -2 0.44 aSi , bar 2 -10 S TiO f 2 log -3 TiS 1400 K aSiO -12 2 = 1 log 0.1 log fO relative to IW -4 Ti2O3 2 0.5 -14 FeS-Feo -5 -6 -5 -4 -3 -2 -1 0 -6 -4 -2 0 log f O relative to IW b 2 2 2MgS + O2 = 2MgO + S2 Fig. 7. Activity of Si as a function of fO2 for the conditions of Eq. (25) for 1400 K and 0 1700 K at 1 bar total pressure. The values at lines refer to activity (a) of SiO2. Solid lines correspond to aSiO2 controlled by Eq. (24). The plot shows that aSiO2 has much lower effect on aSi than fO2 and temperature. ITOI, 1400 K , bar 2 MgS -2 S f RTO, 1400 K 2000K The involvement of ilmenite’s Fe(II) in graphite oxidation (Eq. log IRI, 2000 K -4 IRI, 1400 K 1700K (28)) would lead to CO degassing at greater depths. At fO2 con- log fO relativeMgO to IW 1400K 2 trolled by the IRI buffer (Eq. (2)) at 1700 K, fCO in graphite-bearing RTO, 2000 K -6 ITOI, 2000 K magma could be as high 400 bar (Fig. 8). This means that CO may -6 -4 -2 0 start degassing 4 km below the surface (were partial P of CO be- c 2 comes equal to lithostatic P). Although possible, oxidation of 1700 K graphite by Ti(IV) (Fogel and Rutherford, 1995) is less efficient than MgS by FeO, as illustrated by lower fCO values in corresponding equilib- 0 2MgS + 2SiO2 + O2 = .44 ria (Eqs. (29) and (30); Fig. 8). Nevertheless, high magma temper- 2MgSiO3 + S2 = 0 , bar ature strongly favors oxidation by TiO (Eq. (30)), and a major SiO 2 2 2 a -2 1 S = assimilation of TiO2 could lead to fCO up to 25–70 bar at 1700– f SiO 2 .1 a =0 1 1800 K, which implies a possibility for degassing from the depth iO 2 0.0 log aS = -4 O 2 <230–650 m. aSi ITOI IRI RTO Oxidation of graphite by the SiO2 in magma (Eq. (31)) men- MgSiO3 tioned by Kerber et al. (2009) may also contribute to CO production -6 -6 -4 -2 0 at high-T and low-P conditions. This notion is consistent with signs o o of low-P reduction of SiO2 in C -, Fe -bearing enstatite chondritic log f O2 relative to IW melts leading to dissolution of Sio in Feo-metal (McCoy et al., 1999). Fig. 6. Phase equilibria among sulfides and oxides of Ti and Mg at a function of fO2 Degassing and escape of low-solubility CO would favor further at 1 bar total pressure. Vertical dotted lines show conditions of redox buffers at graphite oxidation by shifting Eq. (27) to the right hand side. corresponding temperatures. In (a), the lines represent Eqs. (4), (8)–(10), and (17) at Although CO removal makes magma more reduced and decreases unity activities of sulfides and oxides at 1500 K. The dashed line corresponds to the o the potential for oxidation, the extent of reduction could be limited FeS–Fe equilibrium (Eq. (18)) that may control fS2. In (b), fugacity of S2 is shown for Eq. (5) at unity activities of MgS and MgO. In (c), equilibrium conditions of Eq. (11) are shown for 1700 K at different activities of silica (aSiO2) and unity activities of MgS and MgSiO3.The value of aSiO2 = 0.44 corresponds to the forsterite–enstatite 4 equilibrium (Eq. (24)) that may control silica activity in the melt. Stability of MgS is C(graphite) + 0.5O2(g) = CO(g) lower at higher fO2 and aSiO2 values. IW 2 However, at fO2 Si3N4ðin meltÞþ3O2 ! 3SiO2ðin meltÞþ2N2ðgÞ magmas may emit more N through decompression. Although degassing of N may account for pyroclastic eruptions of extremely (Ito and Fruehan, 1988). Experimental data show that major degas- reduced melts, more experimental efforts are needed to support sing is also expected from decompression of N-saturated magma at this suggestion. In addition to N2 (see Section 6), corresponding fO > IW 1. At IW, up to 1.6 wt% N could degas through magma 2 2 volcanic gases would contain C, Cl, and S species formed through ascent from 25 kbar ( 230 km below Mercury’s surface) (Mysen low-P decomposition of graphite and melt complexes in the close et al., 2008). If N solubility in reduced melts is proportional to proximity of Mercury’s surface. Degassing experiments with EC P1=2 (Libourel et al., 2003), decompression of FeO-depleted basaltic N2 melts (Muenow et al., 1992) cold be considered as proxy for this magma (IW 6, 1700 K) from 1 kbar causes degassing of up to 0.7 case. wt% N . Similar amounts of degassed N are expected if reduced 2 2 The most efficient degassing is expected through putative oxi- melts (fO 2 RTO f 0 N some S-, C-, and Cl-species could separate into the N2-rich gas. log f IW-6 TiO Nitrogen degassing from extremely reduced magmas has been 2 -2 discussed to account for presumed pyroclastic activity on the par- log Ti2O3 ent body of aubrites (Wilson and Keil, 1991). The low N content in -4 aubrites compared to enstatite chondrites (Table 2) may suggest a ab loss of 150–400 ppm N in igneous processes. These numbers do -6 not contradict the amounts of N released through partial melting 1000 1200 1400 1600 -6 -5 -4 -3 -2 -1 0 of enstatite chondrites in a vacuum (Muenow et al., 1992), through Temperature, K log f O2 relative to IW degassing mechanism is unknown. In addition to decomposition of nitride melt complexes, N could be released through near-surface Fig. 9. Phase equilibria between osbornite (TiN) and Ti oxides. (a) Equilibrium fugacities of N for Eqs. (32) and (33) at fO controlled by redox buffers resented by oxidation of solid nitrides, if they were present in igneous systems 2 2 Eqs. (3) and (4).AtIW 6, TiN coexists with Ti2O3. (b) Stability fields at 1700 K. The on Mercury and elsewhere. Fig. 9 illustrates that osbornite (TiN) figure illustrates that decompression and/or oxidation of reduced magma may could decompose because of oxidation and/or decompression of cause TiN decomposition and nitrogen degassing. M.Yu. Zolotov / Icarus 212 (2011) 24–41 35 For example, reduced volatile-rich magmas delivered from core– in enstatite chondrite melts correspond to higher concentrations mantle boundary after the disrupting Caloris impact could be oxi- of O-bearing gases (CO2,SO2, SO, S2O) and lower contents of C spe- dized through assimilation of crustal rocks that may (Denevi et al., cies (CS2, CS, ClCN) (Fig. 10b). For redox conditions below IW the 2010; Lawrence et al., 2010) or may not contain ilmenite. This pro- most abundant volatiles are CO, N2, Cl, Cl2,S2, and CS2. These com- cess should strongly increase the amount of formed S–C–N–Cl positions are not similar to terrestrial volcanic gases that contain gases and may account for major pyroclastic activity. Below we H2O, CO2,SO2,H2S, HCl, and H2 (Symonds et al., 1994; Oppenhei- use thermochemical equilibrium calculations to evaluate the speci- mer, 2003). ation of corresponding volcanic gases. Another set of speciation models has been developed for bulk gas compositions typical for chondritic analogs of Mercury and 6. Speciation models for volcanic gases its mantle (Figs. 11–13). The calculations were performed at Cl/S mole ratio of 0.01 that represents enstatite, ordinary, and CI chon- Ideal gas chemical equilibria in the S–C–O–Cl–N system have drites, the solar photosphere, and the Kilauea summit 1918 volca- been calculated at variable T, P, fO2, and assumed bulk composi- nic gases (Fig. 3a). This value represents a low range of the Cl/S tions with the use of thermodynamic properties of gas species ratio in Earth’s volcanic gases (Aiuppa et al., 2009) and could also (Barin, 2004) and graphite. The calculations were performed by be considered as a lower limit for Mercury’s gases. The preferred the Gibbs free energy minimization method (Van Zeggeren and N/S mole ratio of 0.017 characterizes EH enstatite chondrites and Storey, 1970) with the GEOCHEQ code (Mironenko et al., 2008) is similar to that in EL and CB chondrites (Fig. 3b). The C/S mole ra- modified by the author. The applicability of these models to Mer- tios of 0.19 and 1.9 were chosen to represent EH and CB chondrites, cury is limited by the uncertainties in the oxidation state of mag- respectively. The latter value also roughly represents CI chondrites mas, bulk compositions of volatiles, and their solubilities in and Kilauea gases (Fig. 3). melts. Therefore, we explore the effects of these parameters on Elevated temperature favors higher abundances of CO, Cl, S, CS, gas speciation rather than formulate definite predictions. SCl, SO, and S2O(Fig. 11). Concentrations of CS2,S3–S8,S2Cl, SCl2, First, we evaluated speciation of gases released from heated and and S2Cl2 are higher at lower temperatures. Condensation of melted enstatite chondrites (Muenow et al., 1992). The gases are graphite occurs at temperatures below those suggested for Mer- characterized with the following elemental ratios S:C:Cl:N = cury’s typical magmas. At 1700 K, extremely reduced gases 1:12.5:1.5:3.5. These ratios are significantly different from bulk (IW 5.3, see Section 2.3) mostly consist of S2,CS2, and CO compositions of chondritic materials (Fig. 3) and could reflect (Fig. 11a). Carbon monoxide dominates in more oxidized and C- limited volatility of S. The speciation of gases was modeled at rich gases (Fig. 11b). The latter gases could be associated with 1700 K, 1 bar total pressure, and log fO2 = 14.9 (IW 5.3) evalu- melts formed through partial melting of silicates from metal-rich ated for the partial melt of the Indarch enstatite chondrite (Sec- (CB/CH) chondrites. tion 2.3). The calculated volume (mole) fractions of major gases High-P gases, which characterize magma before its fragmenta- are as follows: CO, 0.788; N2, 0.113; Cl, 0.033; Cl2, 0.030; S2, tion, have elevated concentrations of CS2,S2Cl, and S2Cl2 0.015; CS2, 0.014. Concentrations of other gases can be seen in (Fig. 12). Low pressure corresponds to higher fractions of CO, S2, Fig. 10a at 1700 K. The results presented in Fig. 10 also explore S, Cl, SO, S2O, and SO2. In C-rich systems, concentrations of C gases effects of temperature and fO2 on the speciation of gases with that could be controlled by graphite which is stable at elevated pres- bulk composition. Elevated temperatures favor higher concentra- sures. Lowering pressure in moderately oxidized systems favors tions of smaller gas molecules: Cl, S2, CS, SCl and S. Condensation oxidation of graphite to CO (Fig. 12b and Section 5.2). Decompres- of graphite below 1460 K strongly decreases CO abundance and sion of ascending magma could be accompanied by increasing CO/ gas becomes rich in N2 and Cl2. More oxidizing conditions than CS2 and S2/CS2 ratios. a 0 b 0 CO CO 1 bar, IW-5.3 1700 K, 1 bar CO2 N2 N2 -1 -1 CS2 Cl Cl Cl 2 Cl 2 S2 S 2 S2 -2 -2 COS SO2 COS graphite SCl2 SCl2 COS log Gas Mole Fraction log Gas CS -3 -3 2 SO S Cl 2 2 CO2 S2Cl Gas Mole Fraction Cl S2Cl SCl S Cl ClCN CS 2 S2O CS -4 S SCl -4 3 SCl S S COCl S 3 COCl COCl2 ClCN S3 S2Cl2 ClCN IW -5 S3 -5 RTO IRI 1000 1200 1400 1600 1800 2000 -6 -4 -2ITOI 0 EC melt BC, M-A melts Temperature, K log f O2 relative to IW Fig. 10. Speciation of volcanic gases with bulk composition of volatiles released from enstatite chondrite samples (Muenow et al., 1992). (a), At log fO2 = 14.9 (IW 5.3) and 1 bar total pressure. (b), At T = 1700 K and 1 bar. In (a), the vertical dotted line shows upper temperature boundary of graphite stability, which condenses below 1460 K. In (b), the symbols at the X axis show conditions for redox buffers (Fig. 1) and fO2 of partial melts of enstatite (EC) and CB chondrites, and melts of Mercury’s mantle from Morgan and Anders (1980), M–A melts (see Section 2.3). 36 M.Yu. Zolotov / Icarus 212 (2011) 24–41 a 0 b 0 S 2 CO 1 bar, IW-5.3, C/S=0.19 1 bar, IW-3.1, C/S=1.9 CS2 CS2 S2 -1 CO -1 COS S3 graphite N2 -2 COS -2 CO2 S2Cl N2 Cl 2 Cl S2Cl Gas Mole Fraction SCl S2Cl S -3 Cl 2 -3 S3 S3 Gas Mole Fraction Cl S4 CS 2 SCl2 CS S2Cl2 S4 -4 CO2 -4 SCl SCl S Cl S Cl 2 2 5 SO S S S2O S6 SO 5 S4 S -5 -5 6 1000 1200 1400 1600 1800 2000 1000 1200 1400 1600 1800 2000 Temperature, K Temperature, K Fig. 11. Speciation of volcanic gases as a function of temperature at 1 bar total pressure. (a), At a generalized bulk composition and redox conditions of EH enstatite chondrites, S:C:Cl:N = 1:0.19:0.01:0.017, IW 5.3. (b), At bulk composition and redox conditions of CB chondrites, S:C:Cl:N = 1:1.9:0.01:0.017, IW 3.1. In (b), the vertical dotted line shows upper boundary of graphite stability, which condenses below 1203 K. a 0 b 0 CO 1700 K, IW-5.3, C/S = 0.19S2 1700 K, IW-3.1, C/S = 1.9 CO S2 -1 -1 CS 2 CS2 Cl N2 COS -2 -2 S S2Cl CO2 S COS N Cl2 2 Cl S2Cl Gas Mole Fraction -3 SCl2 -3 S S 3 3 SC SO SCl2 Gas Mole Fraction Cl2 CO2 SC -4 -4 SO2 S2Cl2 SO SCl SCl S5 S4 COCl S O 2 S4 S6 S2Cl2 -5 -5 -4 -3 -2 -1 0 1 2 -4 -3 -2 -1 0 1 2 log Pressure, bar log Pressure, bar Fig. 12. Speciation of volcanic gases as a function of total pressure at 1700 K. Bulk gas compositions and redox conditions are referred to EH (a) and CB (b) chondrites (see capture of Fig. 11). In (b), the vertical dotted line shows pressure (70 bar) of graphite condensation. Redox conditions below IW strongly affect concentrations of magma assent, or low-P degassing of enstatite chondrite melts CS2, CS, and O-bearing gases: CO2, CO, SO2,S2O, and SO (Fig. 13). (Muenow et al., 1992) would lead to C/S > 2 that results in CO dom- At fO2 a 0 b 0 S CO 1700 K, 1 bar, C/S = 0.19 2 1700 K, 1 bar, C/S = 1.9 CO S2 -1 -1 CO CO 2 CS2 2 CS2 SO2 COS COS N2 -2 Cl -2 SO2 N2 S3 S2Cl Cl Cl2 SO CS log Gas Mole Fraction S2Cl S -3 SCl2 S2O -3 3 Cl2 S log Gas Mole Fraction SCl CS S -4 S4 -4 SCl2 S Cl SCl 2 2 SO S O RTO IRI IW RTO 2 IRI IW -5 -5 ITOI -6EC melt -4 -2ITOI 0 -6EC melt -4 -2 0 BC, MA melts BC, M-A melts log f O relative to IW 2 log f O2 relative to IW Fig. 13. Speciation of volcanic gases as a function of fO2 at 1700 K and 1 bar total pressure. Bulk gas compositions are referred to EH (a) and CB (b) chondrites (see capture of Fig. 11). The symbols at the X axis show conditions for redox buffers (Fig. 1) and fO2 of partial melts of enstatite (EC) and CB chondrites, and melts of Mercury’s mantle from Morgan and Anders (1980), M–A melts (see Section 2.3). a 0 b 0 c 0 S S2CO 2 Cl S2 2 N2 CO CO -1 -1 Cl -1 CS CS2 CS2 2 SCl2 COS S Cl COS 2 COS -2 N -2 -2 2 N2 Cl Cl S S2Cl 3 SCl S Cl CO 2 Cl S3 Cl 2 log Gas Mole Fraction 2 log Gas Mole Fraction 2 -3 -3 -3 SCl2 S3 CO2 CO2 SCl2 log Gas Mole Fraction S S CS -4 S -4 S2Cl2 CS -4 S2Cl2 SCl S SCl 4 COCl S4 CS S2O S4 S2O SO S2O SO S2Cl2 SO CI COCl2 -5 -5 CI -5 -2 -1EH EL 0 1 -2Sun -1 0 CB CI Sun CB Kilauea EL EC gas -2 -1 0EC gas 1 EH, Kilauea EL, EH log C/S Mole Ratio log Cl/S Mole Ratio log N/S Mole Ratio Fig. 14. Speciation of volcanic gases as a function of C/S, Cl/S, and N/S ratios at 1700 K, log fO2 = 13.5 bar (IW 4) and 1 bar total pressure. In (a), S:Cl:N = 1:0.01:0.017. Graphite condenses at C/S > 10. In (b), S:C:N = 1:0.37:0.017. The C/S ratio represents EL chondrites. In (c), S:C:Cl = 1:0.37:0.01. The symbols at the X axis show elemental ratios in chondrites, the solar photosphere, and the Kilauea 1918 volcanic gas, and gases released from enstatite chondrites (EC gas, Muenow et al., 1992). S ratio corresponds to dilution of the gas with N2, which is the dom- have involved preferential high-T and high-P condensation in the inant N species (Fig. 14c). Volume fraction of ClCN may exceed 10 5 solar nebula, evaporation of silicates, and collision-driven losses. only in reduced C–Cl–N rich gas that could be associated with ensta- High temperature could have restricted accretion of H-bearing spe- tite chondrites (Fig. 10). The fraction of NS does not exceed 10 5, and cies and favored their lost early in the body’s history, if they ac- other considered N gases (CN, C2N2, NO, N, NOCl, etc.) are less abun- creted. However, by analogy with enstatite and CB/CH dant than 10 9. Even if H is present in Mercury’s magmatic gases, chondrites, Mercury may not be strongly depleted in S, Cl, and N. 5.9 11.9 1 5 NH3/N2 ratio is 10 –10 (at H/N = 10 –10 , 1700 K and These elements could have survived major losses because of stabil- 1 bar; see Eq. (34)). At 1700 K and fO2 from IW 3toIW 6, the ity of corresponding reduced solids. H2/H2O mole ratio is 37–1175, respectively (see Eq. (35)). H2S, H2, The FeO content in Mercury’ surface silicates below 6 wt% cor- o and HCl could be major H-bearing species in reduced gases. responds to fO2 < IW 2.3 in Fe -metal-saturated magmas. Re- duced conditions favor high solubility of S, N, and Cl in magmas 7. Concluding remarks and would cause partitioning of these elements into melts during magma formation. 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