Divergent T–ƒO2 paths during crystallisation of H2O-rich and H2O-poor magmas as recorded by Ce and U in zircon, with implications for TitaniQ and TitaniZ geothermometry

  • Robert R. LoucksEmail author
  • Marco L. Fiorentini
  • Bruce D. Rohrlach
Open Access
Original Paper


During solidification of magma chambers as systems closed to chemical exchange with environs, the residual siliceous melt may follow a trend of rising, constant, or decreasing oxidation state, relative to reference buffers such as nickel + nickel oxide (NNO) or fayalite + magnetite + quartz. Titanomagnetite–hemoilmenite thermometry and oxybarometry on quenched volcanic suites yield temperature versus oxygen fugacity arrays of varied positive and negative slopes, the validity of which has been disputed for several years. We resolve the controversy by introducing a new recorder of magmatic redox evolution employing temperature- and redox-sensitive trace-element abundances in zircon. The zircon/melt partition coefficients of cerium and uranium vary oppositely in response to variation of magma redox state, but vary in tandem as temperature varies. Plots of U/Pr versus Ce4+/Ce3+ in zircon provide a robust test for change in oxidation state of the melt during zircon crystallisation from cooling magma, and the plots discriminate thermally induced from redox-induced variation of Ce4+/Ce3+ in zircon. Temperature-dependent lattice strain causes Ce4+/Ce3+ in zircon to increase strongly as zircon crystallises from cooling magma at constant Ce4+/Ce3+ ratio in the melt. We examine 19 zircon populations from igneous complexes in varied tectonic settings. Variation of zircon Ce4+/Ce3+ due to minor variation in melt oxidation state during crystallisation is resolvable in 11 cases but very subordinate to temperature dependence. In many zircon populations described in published literature, there is no resolvable change in redox state of the melt during tenfold variation of Ce4+/Ce3+ in zircons. Varied magmatic redox trends indicated by different slopes on plots of zircon U/Pr versus Ce4+/Ce3+ are corroborated by Fe–Ti-oxide-based T–ƒO2 trends of correspondingly varied slopes. Zircon and Fe–Ti-oxide compositions agree that exceptionally, H2O-rich arc magmas tend to follow a trend of rising oxidation state of the melt during late stages of fluid-saturated magmatic differentiation at upper-crustal pressures. We suggest that H2 and/or SO3 and/or Fe2+ loss from the melt to segregating fluid is largely responsible. Conversely, zircon and Fe–Ti-oxide compositions agree in indicating that H2O-poor magmas tend to follow a T–ƒO2 trend of decreasing oxidation state of the melt during late stages of magmatic differentiation at upper-crustal pressures, because the precipitating mineral assemblage has higher Fe3+/Fe2+ than coexisting rhyolitic melt. We present new evidence showing that the Fe–Ti-oxide oxybarometer calibration by Ghiorso and Evans (Am J Sci 308(9):957–1039, 2008) retrieves experimentally imposed values of ƒO2 in laboratory syntheses of Fe–Ti-oxide pairs to a precision of ± 0.2 log unit, over a large experimental temperature range, without systematic bias up to at least log ƒO2 ≈ NNO + 4.4. Their titanomagnetite–hemoilmenite geothermometer calibration has large systematic errors in application to Ti-poor oxides that precipitate from very oxidised magmas. A key outcome is validation of Fe–Ti-oxide-based values of melt TiO2 activity for use in Ti-in-zircon thermometry and Ti-in-quartz thermobarometry.


Magmatic oxidation state Ce4+/Ce3+ in zircon Fe–Ti-oxide geothermometry Fe–Ti-oxide oxybarometry Ti-in-zircon geothermometer Ti-in-quartz geothermometer 


Accurate geothermometry is a key that often unlocks access to other geologically useful information from chemically analysed glass and/or mineral assemblages—e.g., rock or magma viscosity, pressure of crystallisation and depth of subsequent erosional exhumation, oxygen or sulphur fugacity, or the dissolved H2O content in a magma. These pieces of information may be required by earth scientists investigating problems as diverse as rock mechanics, crustal deformation dynamics, ore-forming processes, and volcanic eruption dynamics. Reliable characterisation of the redox state of magmas is important, because about 20% of the elements in the periodic table vary in ionic charge within the range of oxidation states of terrestrial and lunar silicate melts. In many cases, whether an element is sequestered in early minerals or accumulates in residual melts depends on the oxidation state of the melt, which enhances or thwarts a magma’s ability to produce ore deposits of many elements. The oxidation state of sulphur in silicate melts strongly influences its availability for degassing during volcanic eruptions, with climatic implications (Scaillet et al. 1998; Burgisser and Scaillet 2007). Evolution of the oxidation state of silicate melts during generation and crystallisation is normally dominated by the Fe3+/Fe2+ couple, with the S6+/S2− couple usually playing a subordinate but occasionally dominant role (Carmichael and Ghiorso 1986), whereas the other redox-sensitive elements are present in natural silicate magmas in relatively low concentrations and respond passively but promptly to the lead by iron and sulphur.

The widely used Ti-in-zircon geothermometer and Ti-in-quartz geothermobarometer are sensitive to errors in estimates of melt TiO2 activity, which are required as input data (Yakymchuk et al. 2017). Different methods of evaluating \(a_{{{\text{Ti}}{{\text{O}}_{\text{2}}}}}^{{{\text{melt}}}}\) give substantially discrepant results. As an example, for the Bishop Tuff melt \(a_{{{\text{Ti}}{{\text{O}}_{\text{2}}}}}^{{{\text{melt}}}}\) has been estimated at 0.45–0.65 using the Fe–Ti-oxide method (Ghiorso and Gualda 2013), but Thomas and Watson (2012) obtained \(a_{{{\text{Ti}}{{\text{O}}_{\text{2}}}}}^{{{\text{melt}}}}\) = 0.15 using the Rhyolite-MELTS algorithm of Gualda et al. (2012). At an actual crystallisation temperature of, say, 750 °C, these discrepancies in \(a_{{{\text{Ti}}{{\text{O}}_{\text{2}}}}}^{{{\text{melt}}}}\) correspond to temperature discrepancies of ~ 60 to 100 °C, and many kilobars discrepancy in pressures calculated by the Ti-in-quartz geobarometer (Kularatne and Audétat 2014). Kularatne and Audétat report experimentally based values of \(a_{{{\text{Ti}}{{\text{O}}_{\text{2}}}}}^{{{\text{melt}}}}\) in the Bandelier Tuff, Oruanui Tuff, Fish Canyon Tuff, and Bishop Tuff which are consistently ~ 50 to 200% larger than thermodynamic estimates by Rhyolite-MELTS for the same analysed glass samples.

Disagreement about evaluation of \(a_{{{\text{Ti}}{{\text{O}}_{\text{2}}}}}^{{{\text{melt}}}}\) may underlie a dispute in the literature (e.g., Ghiorso and Gualda 2013; Ghiorso et al. 2016; Gualda et al. 2017; Evans and Bachmann 2013; Chamberlain et al. 2015; Evans et al. 2016) as to whether the reducing–cooling trend in the Bishop Tuff (Fig. 1) and similar T–ΔNNOƒO2 trends of Yellowstone caldera tuffs, Cougar Point Tuff, Mt. Katmai, and Toba Tuff are a valid representation of T–ƒO2 conditions in the melt in their pre-eruption magma chambers.

Fig. 1

Divergent trends of magmatic temperature and oxidation state relative to the nickel metal + nickel-oxide reference buffer, ∆NNO, during late, felsic stages of magma evolution, as recorded by titanomagnetite–hemoilmenite phenocryst pairs in volcanic suites and evaluated by the Ghiorso and Evans (2008) calibration of the Fe–Ti-oxide oxybarometer–thermometer. They are shown in two panels to avoid obfuscating overlap. These trends are to be compared with redox trends recorded by analyses of zircon populations which are available for Bishop, Yellowstone, St Helens Yn, South Sister, Toba, and Tampakan. Fe–Ti oxide compositions in the White River Ash are from Preece et al. (2014); Mt St Helens Yn tephra, Gardner et al. (1995); Mt St Helens May 1980 and Shiveluch from Blundy et al. (2006, 2008); South Sister Devil’s Hills rhyolitic domes, Stelten and Cooper (2012); Cougar Point Tuff, Cathey and Nash (2004); Toba Tuff, Chesner (1998). Data for Daisen, Fish Canyon Tuff, Katmai, Bishop Tuff, and Yellowstone Lake Creek and Huckleberry Ridge tuffs are from the compilation by Ghiorso and Evans (2008), which cites literature sources. The NNO and FMQ buffer curves are from Pownceby et al. (1994) and from O’Neill (1987), respectively

In this paper, we proceed by examining the reliability of the Fe–Ti-oxide geothermometer–oxybarometer calibration that is the basis of the contentious T-ΔΝΝΟ trends, as plotted in Fig. 1. This is followed by introduction of zircon trace-element indicators of temperature variation and redox variation in magmas during crystallisation of zircon populations from many volcanic and intrusive complexes. Some of the zircon populations are from volcanic suites in which Fe–Ti oxides have also been analysed, providing a test for consistency of results. These tests are followed by interpretation of the causes of variation among temperature-redox trends during crystallisation–differentiation of varied magma types, and by suggestions for improved thermodynamic modelling of TiO2 activity in silicate melts.

Conflicting opinions on validity of varied Fe–Ti-oxide-based T–ƒO2 trends

Titanomagnetite–hemoilmenite geothermometry and oxybarometry, as thermodynamically calibrated by Anderson et al. (1993) and by Ghiorso and Evans (2008), are widely applied to volcanic rocks with the aim of recovering information about intra-telluric conditions preceding eruption. The validity of the variation among T–ƒO2 trends obtained from Fe–Ti oxides in Fig. 1 has been questioned by Ghiorso and Gualda (2013), who note that trends having positive slopes are inconsistent with predictions by the MELTS and Rhyolite-MELTS free-energy-minimisation algorithms (Gualda et al. 2012; Ghiorso and Gualda 2013). Rhyolite-MELTS modelling of Bishop Tuff compositions has been used to propose that the crystallisation temperature for Bishop Tuff phenocrysts lies within a 10 °C range from top to bottom of the eruptive sequence (Gualda et al. 2012, 2017), whereas Bindeman and Valley (2002), Hildreth and Wilson (2007), Evans and Bachmann (2013), Chamberlain et al. (2015), and Evans et al. (2016) present mutually consistent results from zircon-saturation thermometry, two-feldspar thermometry, Ti-in-quartz thermometry, \(\delta\)18Omagnetite–quartz thermometry and Fe–Ti-oxide thermometry documenting a vertical gradient of ~ 100 °C in the pre-eruptive Bishop Tuff magma chamber tapped by the eruptive sequence. Bindeman and Valley (2002) also report a vertical gradient of ~ 112 °C indicated by \(\delta\)18Omagnetite–quartz thermometry in pre-eruptive magma chamber of the Lava Creek Tuff (Yellowstone), which corroborates the range, as shown in Fig. 1.

Ghiorso and Evans (2008) and Ghiorso and Gualda (2013) and Gardner et al. (2014) show plots of temperature versus oxygen fugacity in the Bishop Tuff derived by applying the Ghiorso and Evans (2008) calibration of the Fe–Ti-oxide oxybarometer–thermometer to compositions of Titanomagnetite–hemoilmenite phenocryst pairs reported by Hildreth and Wilson (2007), as shown in Fig. 1. Ghiorso and Gualda (2013) suggest that the discrepancy between the Rhyolite-MELTS model and the Fe–Ti-oxide-based T–ƒO2 in the Bishop Tuff eruptive suite implies that the Fe–Ti-oxide phenocrysts (or ilmenite, in particular) were not in equilibrium with the melt that is represented by enclosing volcanic glass. Ghiorso and Gualda (2013) contend that, in that case, the crystal-melt disequilibrium would propagate to erroneous estimates of the thermodynamic activity of TiO2 in the melt, which would propagate, in turn, to erroneous estimates of magma temperature based on Ti-in-zircon (TitaniZ) geothermometry (Ferry and Watson 2007) and Ti-in-quartz (TitaniQ) geothermobarometry (Wark and Watson 2006; Huang and Audétat 2012; Kularatne and Audétat 2014), both of which require input of TiO2 activity for temperature calculations.

In contrast, Evans and Bachmann (2013) present several lines of evidence that eruption of the Bishop Tuff magma preserved equilibrium of exchange components and element concentrations among magnetite, ilmenite, biotite, apatite, zircon, and liquid. Evans et al. (2016) and earlier papers cited therein offer numerous lines of independent thermometric and other evidence that the Fe–Ti-oxide phenocrysts probably also maintained equilibrium with feldspars and quartz in the Bishop Tuff magma. Evans et al. (2016) contend that the T–ƒO2 trend recovered from the analytical data of Hildreth and Wilson (2007) and thermodynamic calibration by Ghiorso and Evans (2008) are a valid representation of the T–ƒO2 conditions in the zoned Bishop Tuff magma chamber. In that case, questions arise regarding the adequacy of the calibration data or thermodynamic procedures used by the Rhyolite-MELTS free-energy-minimization algorithm (Gualda et al. 2012; Ghiorso and Gualda 2013) for calculating \(a_{{{\text{Ti}}{{\text{O}}_{\text{2}}}}}^{{{\text{melt}}}}\), with implications for its reliability in thermodynamic modelling of silicic melts more generally.

Trail et al. (2015) present a T–ƒO2 trend for the Bishop Tuff magma, based on X-ray absorption near-edge structure (XANES) measurements of the Ce4+/Ce3+ ratio in zoned zircons and Ti-in-zircon geothermometry in one pumice sample, showing a rise in Ce4+/Ce3+ ratio toward the lower temperature rims of the zircon crystals. They interpret their rim-ward rise of Ce4+/Ce3+ as supporting the arguments by Ghiorso and Gualda (2013) that the cooling–reducing T–ƒO2 trend indicated by Fe–Ti oxides in the Bishop Tuff is not valid. In that case, cooling–reducing trends indicated by Fe–Ti oxides in rhyolites at other eruptive centres (Fig. 1), such as Yellowstone caldera tuffs, Cougar Point Tuff, Katmai/Novarupta tuff, and the Toba Tuff are dubious too, as they also violate melt evolution trends derived by the Rhyolite-MELTS algorithm (Ghiorso and Gualda 2013).

Test of the Ghiorso and Evans (2008) Fe–Ti-oxide thermometer–oxybarometer

To resolve these conflicting propositions, we begin by testing the reliability of the Ghiorso and Evans (2008) thermodynamic calibration of the Fe–Ti-oxide oxybarometer by using it to calculate ƒO2 values from titanomagnetite–hemoilmenite compositions synthesised at controlled ƒO2 and temperature in seven experimental studies (listed in Electronic Supplementary Material 1). Most of them are represented in a plot of ΔΝΝΟ vs Τ shown as Fig. 27 of Ghiorso and Evans (2008), but we use here different plotting variables, and we include some additional experimental data (Blatter et al. 2013) that were not included in their calibration dataset.

Excluding only the highly deviant results of the study by Prouteau and Scaillet (2003) at 9.2 kbar in extremely hydrous felsic melts, we show in Fig. 2a that the Ghiorso and Evans (2008) oxybarometer retrieves the reported experimental ƒO2 values of all 45 experiments in the other seven studies with a standard error of just ± 0.2 log unit ƒO2, which does not exceed the 0.2–0.3 log unit uncertainty in experimental control of ƒO2 (Scaillet and Evans 1999; Pichavant et al. 2002). Figure 2b shows large deviations of calculated temperatures from the reported experimental temperatures for Ti-poor oxide pairs synthesised at high ƒO2. Under those conditions, experimental calibration data available to Ghiorso and Evans were scant; in a Roozeboom plot of Xilm versus Xulv (Ghiorso and Evans 2008, their Fig. 14), isotherms converge tightly, causing poor temperature resolution in Ti-poor compositions. The data trend at low Xilmenite in Fig. 2b may be used to estimate empirical corrections to calculated temperatures of Ti-poor oxide pairs until an improved thermodynamic model becomes available. However, Fig. 2b also indicates that the Ghiorso and Evans (2008) Fe–Ti-oxide geothermometer is accurate in the composition range of Fe–Ti oxides in the Bishop Tuff and other suites plotted in Fig. 1 (Xilmenite > 0.45, ΔΝΝΟ ≤ 2). Therefore, we conclude that mis-calibration of the titanomagnetite–hemoilmenite oxybarometer–thermometer is not responsible for the disagreement of the positive T–ƒO2 trend of the Bishop Tuff Fe–Ti-oxide data in Fig. 1 with the negative slope of the magmatic T–ƒO2 trend predicted by Rhyolite-MELTS or negative T–ƒO2 slope inferred by Trail et al. (2015) from zircon Ce4+/Ce3+ measurements and Ti-in-zircon temperatures.

Fig. 2

a Reliability of the Ghiorso and Evans (2008) thermodynamic calibration of the Fe–Ti-oxide oxybarometer is tested by using it to calculate ƒO2 values (at experiment temperatures) from titanomagnetite–hemoilmenite compositions reported in 45 experiments by seven investigative groups over an experimentally controlled (buffered) log ƒO2 range from NNO to NNO + 4.4, 0.5 to 7 kbar (mostly 2–3 kbar), and 750–1050 °C in hydrous, metaluminous bulk compositions. Colours represent experimental temperature intervals; symbols of different shape identify different investigations. The abscissa variable is the mole fraction of FeTiO3 in Mg–Fe–Mn–Al–Ti hemoilmenite projected onto the FeTiO3–Fe2O3 binary, calculated according to the procedure of Ghiorso and Evans (2008). b There are large systematic deviations of temperatures calculated by the Ghiorso and Evans (2008) geothermometer from Fe–Ti-oxide pairs synthesised at high ƒO2 (low mole fraction FeTiO3), evident in this plot of the same experiments as in a. But over the composition range of Fe–Ti-oxide pairs in the Bishop Tuff (grey bar; Hildreth 1977; Hildreth and Wilson 2007), the geothermometer is accurate. The plotted data are listed in the Electronic Supplementary Material 1

Independent test of magmatic T–ƒO2 trends using Ce and U redox couples

Zircon takes in cerium and uranium in quantities that are easy to measure quantitatively. Zircon is very resistant to adjustment of its initial composition by ion diffusion within the crystal lattice (Cherniak and Watson 2003). Ce and U vary in ionic charge within the ƒO2 range of terrestrial silicate melts; zircon/melt partition coefficients of Ce and U vary sensitively with ƒO2 and can, in principle, be used as magmatic oxybarometers. We examine first the variation in ionic charge of Ce and U in the melt as its T and ƒO2 vary and follow with an examination of the role of temperature-sensitive lattice strain on zircon/melt partition coefficients of U4+, U5+, Ce4+, Ce3+, and trivalent lanthanides. That is followed by introduction of plots of U/Pr versus Ce4+/Ce3+ in zircon as robust indicators of T–ƒO2 trends as zircon crystallises during magmatic differentiation. Those zircon-based magmatic redox trends are to be compared with Fe–Ti-oxide trends in Fig. 1 to test for consistency.

Relative sensitivity of Ce, U, and Fe redox couples in the melt to variation of T and ƒO2

The effect of varying temperature on Ce4+/Ce3+ and U5+/U4+ in silicate melts is illustrated in Fig. 3. These show that the T–ƒO2 slopes of contours of Ce4+/Ce3+ and U5+/U4+ in the melt are nearly indistinguishable from the slope of Fe3+/Fe2+ contours in the melt and from the NNO and FMQ reference buffers. Schreiber (1987) and Schreiber et al. (1987) use redox micro-titrations and optical spectroscopy to measure redox ratios in silicate glasses quenched at varied ƒO2. They report Ce4+/Ce3+ = 1 at log ƒO2 = − 0.25 at 1150 °C and at log ƒO2 = + 3.13 at 1500 °C (uncertainties are not given). They report that U5+/U4+ = 1 at log ƒO2 = − 8.8 at 1150 °C and log ƒO2 = − 5.1 at 1500 °C. Those results are in satisfactory accord with ƒO2 dependence of mineral/melt partition coefficients of Ce and U, as experimentally constrained by Burnham and Berry (2012) and of U by Fonseca et al. (2014), and in accord with range brackets by Farges et al. (1992) on U5+/U4+ measured by optical spectroscopy in aluminosilicate glasses synthesised at varied ƒO2.

Fig. 3

Position and slope of the fayalite–magnetite–quartz equilibrium and nickel metal–nickel oxide equilibrium, as calibrated by O’Neill (1987) and Pownceby and O’Neill (1994), are compared with experimentally calibrated contours along which Ce4+/Ce3+ and U5+/U4+ and Fe3+/Fe2+ in the silicate melt are unity. The cerium and uranium redox ratios in synthetic aluminosilicate glasses at 1150 and 1500 °C are from Schreiber (1987) and Schreiber et al. (1987). The curve for Fe3+/Fe2+ is according to the experimental calibration by Kress and Carmichael (1991). Near parallelism of the curves implies very similar enthalpies for these redox reactions

We can calculate the average enthalpy of the Ce and U redox reactions in the melt over that temperature interval using the Van’t Hoff relation (T in kelvins):
$$\log \;f{{\text{O}}_{2(T2)}}-{\text{ }}\log {\text{ }}f{{\text{O}}_{2(T1)}}=~~\frac{H}{{2.303~R}} \times \left( {~\frac{1}{{T2}}-\frac{1}{{T1}}} \right).$$

We obtain ΔΗ = − 466.55 kJ for the homogenous reaction 2 Ce2O3 + O2 = 4 CeO2 in the melt. That value compares closely with − 463.44 kJ for the NNO buffer reaction 2Ni + O2 = 2 NiO (Pownceby and O’Neill (1994). For the uranium oxidation reaction 4 UO2 + O2 = 4 UO2.5 in the melt, we obtain Δ H = − 510.72 kJ. That result is close to ∆= − 499.18 kJ for the FMQ buffer reaction, 3 Fe2SiO4 + O2 = 2 Fe3O4 + 3 SiO2 in the same temperature interval (O’Neill 1987). Our result for the cerium oxidation reaction is in accord with the value − 436.8 ± 38.4 kJ (2σ) reported by Smythe and Brenan (2015) in experiments at 1200–1500 °C and 1 bar ƒΟ2 on anhydrous, Fe-free, metaluminous haplo-basalt, haplo-andesite, and haplo-rhyolite compositions in which Ce4+/Ce3+ was determined potentiometrically and by XANES spectroscopy. Our result for Ce is also compatible with a value of − 397 ± 292 kJ (2σ) reported by Burnham and Berry (2014) in experiments at 1300–1500 °C in melt of anorthite–diopside eutectic composition in which they measured Ce4+/Ce3+ by XANES in the quenched glass. We are unaware of any independent constraints on the enthalpy of the U4+ → U5+ reaction.

The results calculated above and shown in Fig. 3 indicate that T–ƒO2 slopes of the Ce and U redox reactions in aluminosilicate melts are statistically indistinguishable from the T–ƒO2 slopes of the NNO and FMQ reference buffers. We expect that magmatic differentiation trends at essentially constant Ce4+/Ce3+ and U5+/U4+ in the melt are the usual case over most of the course of magmatic differentiation in upper-crustal magma chambers. Carmichael (1991) reports that experimentally constrained T–ƒO2 slopes of contours of Fe3+/Fe2+ in metaluminous aluminosilicate melts are also indistinguishable from those of the NNO and FMQ buffers, as we illustrate in Fig. 3. As he shows, natural metaluminous magmatic differentiation series that are dominated by fractional crystallisation in convergent and divergent plate-margin and intra-plate tectonic settings typically evolve from basaltic–andesitic through dacitic compositions along T–ƒO2 trends subparallel to NNO or FMQ:

  • Hildreth (1983) uses Fe–Ti oxide phenocryst pairs to document a T–ƒO2 trend from andesitic to felsic melts subparallel to the FMQ buffer at Mt. Katmai, Alaska.

  • Blundy et al. (2006, 2008) demonstrate Fe–Ti-oxide-based T–ƒO2 trends parallel to NNO in arc andesite-to-dacite differentiation series at Shiveluch (Kamchatka) and Mt St Helens (Washington).

  • Rohrlach (2002) and Rohrlach and Loucks (2005) show Fe–Ti-oxide-based T–ƒO2 trends parallel to NNO in arc andesite-to-dacite differentiation series at Tampakan volcano (Mindanao, Philippines).

  • Perfit and Fornari (1983a, b) report that along the Eastern Galapagos Rift spreading ridge and Inca Transform, the magmatic differentiation series from olivine tholeiitic melts to Fe–Ti basalt to icelandite and ferro-dacite (glass compositions 51–64 wt% SiO2 and 7.5 to 1.0 wt% MgO) evolved subparallel to the FMQ buffer along the liquid line of descent from ~ 1200 °C to < 910 °C.

  • Righter et al. (1998) demonstrate a T–ƒO2 trajectory parallel to FMQ over the course of magmatic differentiation from olivine basalt to rhyolite at Volcán Alcedo, Galápagos.

  • Carmichael (1967) demonstrates the same trajectory for the basalt-to-rhyodacite series at Thingmuli, Iceland.

  • Delano (2001) examines a large database for mid-ocean-ridge igneous suites and demonstrates that a T–ƒO2 trajectory parallel to FMQ is typical.

According to Carmichael (1991), this behavior is because (1) the magma chambers are effectively closed to O2 exchange with environs and (2) during cooling and crystallisation, the poly-phase cotectic mineral assemblage extracts Fe3+ and Fe2+ in essentially the same proportions as were present in the melt at the onset of Fe–Ti-oxide precipitation. To the degree that a magmatic differentiation series evolves along such a trend of constant Fe3+/Fe2+ in the melt, it also evolves along a trend of essentially constant Ce4+/Ce3+ and U5+/U4+ in the melt, as shown in Fig. 3. In a following section, we introduce an independent, robust demonstration that igneous intrusions usually follow a liquid line of descent from basaltic–andesitic through dacitic compositions controlled by fractional crystallisation in a system closed to oxygen and hydrogen.

Temperature-dependent lattice-strain influence on M4+ and M3+ partition coefficients

Blundy and Wood (1994) introduced a method for estimating mineral-melt trace-element partition coefficients, based on the elasticity of the crystal lattice and the degree to which substituent ions of various sizes misfit the lattice sites of the host mineral. Their method extends to crystal-melt ion-exchange reactions an expression derived by Brice (1975) that relates the mechanical strain energy, ΔGstrain, around a size-mis-fitting ion to the elasticity (Young’s Modulus, E) of the host crystal:
$$\Delta {G_{{\text{strain}}}}={\text{ 4}}\pi {\text{E}}{{\text{N}}_{\text{A}}}\left[ {\left( {{\text{ }}{r_{\text{o}}}/{\text{2}}} \right){\text{ }}{{\left( {{r_i}-{r_{\text{o}}}} \right)}^{\text{2}}}+{\text{ 1}}/{\text{3 }}{{\left( {{r_i}-{\text{ }}{r_{\text{o}}}} \right)}^{\text{3}}}} \right]{\text{ }} \approx {\text{ }}-{\text{ RT ln }}(D_{i}^{{{\text{z/m}}}}/D_{o}^{{{\text{z/m}}}})$$
wherein ion radius ri and partition coefficient Di refer to the substituent element i in zircon (z) and melt (m), and ro and Do refer to an element that is an ideal, strain-free fit in the lattice site. NA is Avogadro’s number, R is the gas constant, and T is in kelvins. This relation is of the same form as the familiar relation between the Gibbs energy and equilibrium constant for an ion-exchange reaction, ΔG° = − In Κexchg, but Nernst-type partition coefficients, Di (in units of concentration by weight) can be used, because conversion factors for ppmw to mole fractions of components in zircon and melt cancel in the ratio. Division of Eq. (2) by RT yields a relation that illustrates the reciprocal variation of partition coefficients with temperature.

We apply Eq. (2) to zircon/melt partition coefficients reported for a series of tetravalent ions and trivalent ions, to evaluate ro and Do and E by least-squares regression on each valence series. Ionic radii, ri, are listed in Table 1 (from Shannon 1976). Figure 4a illustrates the results for a series of tetravalent and trivalent ions in the Zr-dominated sublattice of zircon, as measured by ion probe on zircons and haplo-andesitic melt in experiments at 1300 °C and 1 atm and log ƒO2 ≤ FMQ by Burnham and Berry (2012). Figure 4a also shows a series of partition coefficients, measured by ion probe on unpolished surfaces of natural zircon crystals and on enclosing rhyolitic volcanic glass, reported by Claiborne et al. (2018). Figure 4a shows that parabolas fitted to the data are tighter for tetravalent ions than for trivalent ions at each temperature, and the parabolas are narrower at lower temperature, reflecting decrease in lattice-site elasticity with increasing ionic charge and with decreasing temperature.

Table 1

Effective ionic radii (Å) of cations in (A) eightfold coordination with oxygen that may substitute for Zr in zircon (Shannon 1976) and (B) fourfold coordination with oxygen that may substitute for Si in zircon (Shannon 1976)






(A) Cations in eightfold coordination with oxygen


































































(B) Cations in fourfold coordination with oxygen





























The similarities of the radius of U4+ with Ce4+ and of Pr3+ with Ce3+ are highlighted in boldface

Fig. 4

a Comparison of predicted (lines) and observed (symbols) zircon/melt partition coefficients for tetravalent and trivalent ions occupying the Zr-dominated sublattice in zircon. Blue symbols represent ion-probe measurements 1–2 µm deep on the natural surface of zircon crystals and enclosing volcanic glass from Mt. St. Helens (sample SHL-26Z), reported by Claiborne et al. (2018), using their zircon-saturation temperature (703 ± 20 °C; 1σ; calibration of Watson and Harrison 1983). Red symbols represent ion-probe analyses of zircon rims and enclosing haplo-andesitic glass produced in experiments at 1300 °C and 1 atm by Burnham and Berry (2012). Ionic radii are from Shannon (1976). b Temperature dependence of Ce4+/Ce3+ fractionation between zircon and melt due to lattice strain, as represented by variation of the zircon/melt partition coefficients of Ce4+ and Ce3+ with temperature, determined by least-squares fitting (as in a) of Eq. (2) to partition coefficients and zircon-saturation temperatures reported for natural phenocryst–glass pairs by Claiborne et al. (2018) (black and blue symbols; standard errors not reported) and fitted to experimental (1300 °C) zircon/melt partition coefficients analysed by SIMS and by LA-ICPMS by Burnham and Berry (2012). Least-squares regression yields the equation: log \(\left( {D_{{{\text{C}}{{\text{e}}_{\text{4}}}}}^{{{\text{z/m}}}}/D_{{{\text{C}}{{\text{e}}_{\text{3}}}}}^{{{\text{z/m}}}}} \right)\) = − 2.34 + 0.9639 × 104/T(K); R = 0.93

Least-squares fits of Eq. (2) to those and other volcanic samples described by Claiborne et al. (2018) yield the parameters listed in Table 2. The derived values in Table 2 for the “ideal fit” ion size, ro, and for the Young’s Modulus, EM+, for felsic volcanic samples are similar to the values ro4+ = 0.912 Å, E4+ ≈ 750 GPa, ro3+ = 0.962 Å, and E3+ ≈ 620 GPa derived by Blundy and Wood (2003) for a zircon–glass pair from a natural rhyolite. A strong trend of increasing Do values (apex of the parabolas) and \(D_{{{\text{C}}{{\text{e}}_{\text{4}}}}}^{{{\text{z/m}}}}\) values with decreasing temperature is evident; a weak tendency for E values and \(D_{{{\text{C}}{{\text{e}}_{\text{3}}}}}^{{{\text{z/m}}}}\) to increase with decreasing temperature is suggested. In the case of mis-fitting trivalent ions, there is a competition between the tendency of crystal/melt partition coefficients to increase with decreasing temperature, and the increasing inability of the lattice to deform to accommodate them as temperature decreases. Figure 4b shows a plot of zircon/melt partition coefficients of Ce4+ and Ce3+ (Table 2) at the experimental temperature 1300 °C or at the zircon-saturation temperature reported by Claiborne et al. for natural zircon–glass pairs. Dispersion of the data for volcanic suites is due mainly to uncertainty in experimental calibration of the zircon-saturation geothermometer (around ± 15 °C, 1σ; Hanchar and Watson 2003), and due to micrometer-scale compositional zoning in analysis spots 15 µm diameter on “rims” of natural crystals in samples IHB, ITHn, and HRL21. Natural surfaces of zircon crystals were analysed in SHL samples. Magma chamber replenishments and mixing preceding and accompanying volcanic eruptions may, in some cases, cause mis-match of zircon crystals with the melt from which they grew. Least-squares regression of the array in Fig. 4b yields the equation: log (\(D_{{{\text{C}}{{\text{e}}_{\text{4}}}}}^{{{\text{z/m}}}}\)/\(D_{{{\text{C}}{{\text{e}}_{\text{3}}}}}^{{{\text{z/m}}}}\)) = − 2.34 + 0.9639 × 104/T(K); R = 0.93.

Table 2

Best fit parameters for Eq. (2) applied to experimental partition data (B&B-SIMS, B&B-ICP) from Burnham and Berry (2012) and to chemical analyses of natural zircon/glass pairs (other samples) from Claiborne et al. (2018)


T (°C)




E3+ (GPa)

ro4+ (Å)

D o 4+

E4+ (GPa)






6.4 × 10−3










5.1 × 10−3










5.8 × 10−3










6.8 × 10−4










1.8 × 10−4










8.3 × 10−4










1.0 × 10−4










4.6 × 10−3






Also shown are zircon/melt partition coefficients of Ce3+ and Ce4+ in each sample, obtained from parabolic fits to other M3+ and M4+ ions, respectively

In Fig. 4a, the measured values of \(D_{{\text{U}}}^{{{\text{z/m}}}}\) plot below the parabolas fitted to other tetravalent ions. Blundy and Wood (2003), Ballard et al. (2002), and others have noted the same sense of deviation of measured of \(D_{{\text{U}}}^{{{\text{z/m}}}}\) from the value predicted for U4+ in natural zircons from peralkaline and metaluminous rocks. The deviation may be due to incorporation of a minor proportion of the uranium as U5+ in zircon via the charge-coupled substitution 2 Zr4+ ↔ REE3+ + U5+. Burnham and Berry (2017) report that zircons from peraluminous S-type granites have phosphorus contents near the amounts expected for charge balance of the measured abundances of Y + REE3+ according to the charge-coupled substitution Zr4+ + Si4+ ↔ REE3+ + P5+, but they report that zircons from metaluminous I-type granitoids have a systematic deficit of charge-balancing phosphorus. This is understandable in light of experiments by Pichavant et al. (1992) showing that, at a given temperature and pressure and wt% H2O in the melt, apatite solubility and phosphorus content of silicate melts increase dramatically with increasing Al/(Ca + Na + K), due to increase of Ca–Al complexing in the melt with increasing A/CNK. Therefore, in metaluminous melts with higher lime activity and lower phosphate activity at apatite saturation, other charge-balancing mechanisms are implicated for incorporation of REE3+ into zircon. Ionic radii in Table 1 suggest that U5+ should be a lower strain candidate for that charge-coupled substitution role than, say, Nb5+, Ta5+, or V5+. Variation of the charge-coupled substitution mechanism with melt A/CNK is expected to diminish a correlation of E3+ with temperature in Table 2.

Cooling-induced variation of Ce4+/Ce3+ and U5+/U4+ in zircon at constant ratios in the melt

In Fig. 5, we present new measurements illustrating strongly temperature-dependent variation of Ce4+/Ce3+ in a natural zircon assemblage from a calc–alkalic volcanic arc suite. A few drill–core and outcrop samples of eruptive andesites and dacites from the Tampakan volcanic complex are sufficiently free of hydrothermal alteration to permit application of the titanomagnetite–ilmenite geothermometer (calibration of Ghiorso and Evans 2008) and the plagioclase–hornblende geothermometer (calibration of Holland and Blundy 1994). New data for Tampakan phenocryst compositions that are represented in Figs. 1 and 5 are shown in Electronic Supplementary Materials 2 and 3. In a zircon-bearing subset of those samples, we use whole-rock trace-element analyses by ICPMS together with zircon compositions analysed by ArF ELA-ICPMS to calculate zircon Ce4+/Ce3+ ratios by the lattice-strain method of Blundy and Wood (1994, 2003). All available paired T–Ce4+/Ce3+ data are plotted in Fig. 5, which shows strong lattice-strain discrimination, as the zircon lattice thermally contracts during growth from cooling magma, in favour of smaller, better-fitting Ce4+ relative to larger, size- and charge-mis-fitting Ce3+ ions in Zr4+ lattice sites. The T–ƒO2 trend of Tampakan Fe–Ti oxide data in Fig. 1 and trace-element compositions of Tampakan zircons presented ahead (Fig. 9) demonstrate that the Tampakan andesite–dacite series developed at essentially constant oxidation state of Fe, Ce, and U in the melt.

Fig. 5

For zircon-bearing silicic andesites and dacites from the Tampakan volcano, Mindanao, that are sufficiently free of hydrothermal alteration to permit geothermometry on associated phenocrysts, we combine whole-rock ICPMS analyses of trace elements with analyses of zircons by ArF ELA-ICPMS (Rohrlach 2002) to calculate zircon Ce4+/Ce3+ ratios by the lattice-strain method of Blundy and Wood (1994). Those Ce4+/Ce3+ ratios are plotted against temperatures obtained for the same samples on touching titanomagnetite–ilmenite phenocryst pairs (calibration of Ghiorso and Evans 2008) and on hornblende–plagioclase phenocryst pairs (calibration by Holland and Blundy 1994). Estimated uncertainty bars are ± 1 SE

The effect of thermally modulated lattice strain on incorporation of U4+, U5+, Ce4+, Ce3+, and other trivalent lanthanides is illustrated in Fig. 6, wherein each panel shows a compilation of experimental results at varying temperature but constant pressure, constant bulk composition of the charge, and constant oxidation state of Ce and U in the melt. In each suite of experiments, those at the three highest temperatures are illustrated, as they are most likely to approximate equilibrium partitioning. Each panel shows experimentally determined zircon/melt partition coefficients of an ion, U4+ or Ce4+, that fits relatively well in the Zr4+ sublattice of zircon, and also shows partition coefficients of a larger ion that fits relatively poorly in the Zr4+ sublattice (see ionic radii in Table 1). In each panel of Fig. 6, the ratio of the better-fitting, smaller ion to poorer fitting, larger ion increases in zircon with decreasing temperature at constant redox state of the melt—including the ratio Ce4+/Ce3+ in zircon (Fig. 6d). All experiments were buffered at NNO, which is shown in Fig. 3 to correspond to effectively constant oxidation state of U and Ce in the melt. These experiments corroborate calculated results in Fig. 4 for temperature dependence of lattice-strain influence on zircon/melt partition coefficients of ions of various sizes.

Fig. 6

Experimental determinations of the effect of varied temperature on zircon/melt partition coefficients (Dizirc/melt) of U, Ce, La, Pr, Gd, and Dy in experiments in which ƒO2 was buffered by Ni + NiO, at constant pressure, and nearly constant major-element composition of the hydrous granitoid melts. In each panel, the element that is most compatible in zircon is red; the element least compatible in zircon is blue. In each panel the ratio of the better-fitting ion to the poorer fitting ion rises with decreasing temperature. a, b Experimental results by Rubatto and Hermann (2007) in hydrous peraluminous tonalite, 66.1–69.8 wt% SiO2 at 20 kbar and ƒO2 buffered at NNO. c Experimental results by Luo and Ayers (2009) in hydrous peralkaline granite, 68.0–70.3 wt% SiO2 at 15 kbar and ƒO2 on the NNO buffer; d experimental results by Trail et al. (2012) in peraluminous haplo-granitic melt at 10 kbar, showing an increase in the zircon’s Ce4+/Ce3+ ratio with decreasing temperature at constant redox state of Ce in the melt on the NNO buffer

In thoughtfully designed experimental studies aiming to calibrate Ce4+/Ce3+ in zircon as a magmatic oxybarometer, Trail et al. (2011, 2012) recognise substantial temperature dependence of zircon/melt partition coefficients of Ce3+ and Ce4+. However, zircon/melt disequilibrium partitioning is problematic in laboratory experiments (and in nature), as discussed by Luo and Ayers (2009), Hofmann et al. (2014), and Trail et al. (2015), so the temperature dependence has not been well characterised by laboratory experiments on granitoid magmas. However, Claiborne et al. (2018) address much of that deficiency by determination of zircon/glass partition coefficients in natural volcanic suites, coupled with zircon-saturation geothermometry to characterise the varying temperature dependence of zircon/melt partition coefficients for ions of charge 5 +, 4 +, and 3 +.

Effect of isothermally varying ƒO2 on relative incorporation of Ce and U by zircon

Figure 7 illustrates the effect of varying magmatic ƒO2 on zircon/melt partition coefficients of Ce and U at constant pressure (1 atm) and temperature (1300 °C) and constant elemental composition of the melt (Fe-free haplo-andesite). The experiments show that zircon/melt partition coefficients of U and Ce have opposite senses of variation in response to isothermally varying ƒO2. Zircon/melt partition coefficients of the charge-invariant elements La, Pr, Nd, Sm, Gd and heavier REEs do not vary over that range of ƒO2. Therefore, within the log ƒO2 range FMQ−2 to FMQ+5 that is relevant to most terrestrial magmas, nearly all Ce in zircon is Ce4+, and the element ratios U/Pr and Ce/Nd show opposite variation trends. Ce/Nd is a proxy for Ce4+/Ce3+ in these experiments, in which Ce valence was not measured (Fig. 7). Although the bulk composition of experimental charges includes an ample supply of Sc3+, Y3+ and REE3+ and of P5+, Nb5+ and Ta5+ ions that could team up for coupled substitutions of the type 2 Zr4+ \(\leftrightarrow\) M3+ + M5+ or Zr4+ + Si4+ \(\leftrightarrow\) M3+ + M5+, the trends of zircon/melt partition coefficients of Ce and U demonstrate that coupled substitutions of charge- and size-misfitting ions are strongly disfavoured by zircon. The variation trends can be described by a sigmoid function, as derived by Burnham and Berry (2012).

Fig. 7

Nernst partition coefficients of Ce and U and other trace elements between zircon and synthetic andesitic melt were measured by secondary-ion mass spectrometry (SIMS, ion probe) and by LA-ICPMS in Burnham and Berry’s (2012) experiments at 1300 °C and 14 log units variation of oxygen fugacity relative to the fayalite–magnetite–quartz reference buffer (FMQ). a Oxidation state of cerium varies from entirely Ce3+ in melt at the lowest ƒO2 to increased (but still small) amounts of Ce4+ in the melt at the highest ƒO2, over which range the Ce content of zircon rises ~ 100-fold as the proportion of Ce4+ in the melt increases. Uranium shows the opposite trend, being entirely U4+ in the melt at the lowest ƒO2, but with increasing amounts of U5+ the melt as ƒO2 increases. The data show a very strong preference for U4+ in Zr4+ lattice sites, and the zircon/melt partition coefficient of U declines approximately 20-fold as the melt’s supply of U4+ diminishes with rising ƒO2. Burnham and Berry report standard errors of \(D_{{{\text{Ce}}}}^{{{\text{z/m}}}}\) that average ± 74% of the measured value; average standard errors of \(D_{{\text{U}}}^{{{\text{z/m}}}}\) are ± 34% of the measured values. b Ce/Nd ratio in zircon (a proxy for ΣCe/Ce3+ in this experimental dataset in which Ce valence was not reported) rises approximately 100-fold over the experimental ƒO2 range, while U/Pr in zircon declines by a factor of about 20

Burnham and Berry (2012) question the validity of the U trend in Fig. 7 because of suspected (but unsubstantiated) volatilisation of UO3 from the melt in the 1-atm gas-flow furnace used for these experiments. However, Luo and Ayers (2009) document the same trend of \(D_{{\text{U}}}^{{{\text{zirc/melt}}}}\) in experiments in sealed capsules in cold-seal pressure vessels at 2 kbar and 800 °C spanning 9 log units ƒO2. They report \(D_{{\text{U}}}^{{{\text{zirc/melt}}}}\) = 32 on the wüstite + iron buffer (= FMQ−4.0); \(D_{{\text{U}}}^{{{\text{zirc/melt}}}}\) = 20 on the nickel + nickel-oxide buffer (= FMQ+0.84); and \(D_{{\text{U}}}^{{{\text{zirc/melt}}}}\) = 3.3 on the hematite + magnetite buffer (= FMQ+5.05). That 10-fold range in \(D_{{\text{U}}}^{{{\text{zirc/melt}}}}\) is the same as in the equivalent ∆FMQ range in Fig. 7. Various spectroscopic methods have been employed by Schreiber and Andrews (1980), Farges et al. (1992) and Berry et al. (2008) to demonstrate that the oxidation of U4+ to U6+ in the melt proceeds via a U5+ intermediate stage which has a predominance range a bit above FMQ at magmatic temperatures.

Introduction of zircon U/Pr versus Ce4+/Ce3+ plots

The experiments reported in Fig. 7 demonstrate that, at oxidation states of common terrestrial magmas, essentially all Ce in zircon occurs as Ce4+ and nearly all U as U4+. We introduce plots of U/Pr versus Ce4+/Ce3+ in zircon to discriminate whether Ce4+/Ce3+ variations in zircon are due to redox variations in the melt during zircon crystallisation and/or due to temperature-regulated fractionation of Ce4+ from Ce3+ by lattice strain. Table 1 shows that the ionic radius of U4+ in eightfold coordination with oxygen is 1.00 Å, very similar to that of Ce4+, 0.97 Å, and shows that the U4+/Pr3+ ionic radius ratio, 0.888, is very similar to the Ce4+/Ce3+ ionic radius ratio, 0.849. Therefore, the effect of temperature variation on the relative zircon/melt partition coefficients of these ratios must be nearly identical. The close similarity of ionic radii of Pr3+ and Ce3+ is the rationale for choosing Pr as a denominator element. However, those ratios in zircon vary oppositely in response to ƒO2 variation, as indicated in Fig. 7b. This test of T–ƒO2 trends during magmatic differentiation is applicable to slowly cooled intrusive rocks, and it can test the reliability of T–ƒO2 trends recovered from titanomagnetite–hemoilmenite phenocryst pairs in volcanic rocks. The Fe–Ti-oxide oxybarometer–thermometer is reliably applicable only to quickly quenched volcanic rocks, inasmuch as Fe–Ti oxides re-equilibrate on the time scale of days to weeks at near-solidus temperatures (Venezky and Rutherford 1997).

Magmatic redox trends from zircon U/Pr versus Ce4+/Ce3+plots

Analytical and computational methods and sources of data dispersion

Electronic Supplementary Material 3 lists previously unpublished analyses of zircons from the Tampakan volcanic complex, Mindanao, done by ArF ELA-ICPMS at the Research School of Earth Sciences, Australian National University. In following paragraphs, we show zircon U/Pr versus Ce4+/Ce3+ derived from ion-probe and LA-ICPMS analyses that were compiled from the literature and combined with our data for the Tampakan complex. Calculation of Ce4+/Ce3+ by the lattice-strain method of Blundy and Wood (1994, 2003) utilises zircon/melt partition coefficients obtained by combining analyses of zircon crystal rims or surfaces with analyses of associated glass (in volcanic suites) or with analyses of host whole rocks (in intrusive suites).

Zircons in all suites in the following data plots have complex oscillatory growth bands and sector zoning similar to that of Bishop Tuff zircon in Fig. 8. There is evidence from synthetic and natural zircons that very slow diffusion of high-electric-field-strength ions of charge 3+, 4 + and 5 + through a microns-thick melt boundary layer adjacent to a growing zircon crystal surface results in disequilibrium between the boundary-layer melt film and the bulk melt, and hence between zircon and bulk melt; that diffusional disequilibrium causes oscillatory compositional zoning of zircon at sub-micron to tens-of-microns scale (Luo and Ayers 2009). Such disequilibrium oscillations in composition are likely to be especially pronounced in zircons that grew from very viscous, H2O-poor, high-silica rhyolitic melts, as documented in the Younger Toba Tuff and Bishop Tuff by Hofmann et al. (2014). There is also persuasive evidence that the zircon/melt partition coefficients of high-field-strength elements vary between prism and various pyramidal crystal faces along the same growth band in zircon, as demonstrated by Tailby et al. (2011) and in Bishop Tuff zircons by Chamberlain et al. (2014). Although all utilised zircon analyses in Figs. 9, 10, and 11 are described by the respective data sources as “rims”, the compositional heterogeneity within outer growth bands can be substantial at the 15–30 μm scale of ion-probe and LA-ICPMS beam spots. This contributes scatter to the zircon composition arrays that exceeds instrumental analytical errors on homogeneous material. However, the data dispersion within zircon arrays typically does not substantially exceed the dispersion within Fe–Ti-oxide arrays for volcanic suites in Fig. 1. Dispersion of zircon and Fe–Ti-oxide compositions within volcanic suites is in many cases due in part to chamber replenishments and magma mixing immediately preceding the eruptions, as documented for the Bishop Tuff, Younger Toba Tuff, and Yellowstone supereruptions, among others (Evans and Bachmann 2013; Chesner and Luhr 2010; Wotzlaw et al. 2015). Eruption of inhomogeneous melt contributes some uncertainty in identifying equilibrium pairs of zircon rims and host glass in those volcanic suites.

Fig. 8

Cathodoluminescence image showing oscillatory banding and sector zoning of zircons in the Bishop Tuff (after Crowley et al. 2007), which have been shown by Hofmann et al (2014) and Chamberlain et al. (2014) to correspond to significant variations of trace-element composition from band to band within a sector, and from sector to sector along individual growth bands. The outer oscillatory growth bands are narrower than the usual 10–30 µm diameter of ion-probe or LA-ICPMS beam spot analyses on zircon “rims”, so such heterogeneity contributes dispersion to arrays of precise analyses in Figs. 9, 10, and 11

Fig. 9

Ce4+/Ce3+ ratios in zircons (calculated by the lattice-strain method of Blundy and Wood 1994, 2003) from eight igneous units are plotted against U/Pr of the zircons, measured by ArF ELA-ICPMS. As discussed in the text, these ratios must have nearly identical temperature sensitivity of their zircon/melt partition coefficients. Note that the slopes of the arrays do not differ significantly from 1.0 in these eight examples. Standard error of the slope and intercept are ± 1σ. After considering the presence of non-negligible (but unquantifiable) analytical uncertainty in these datasets, it is apparent from the slopes and R2 values that there is no resolvable component of the variance in Ce4+/Ce3+ within each array that is attributable to variation of magmatic oxidation state. Data sources: Koksai, Erdenet, Nurkazghan, and Bozshakol porphyry copper deposits: Shen et al. (2015); Shiyaogou porphyry molybdenum deposit: Han et al. (2013); El Abra porphyry copper igneous complex: Ballard et al. (2002); Tampakan volcanic complex: Electronic Supplementary Material 3

Fig. 10

Slopes of zircon composition arrays test whether the melt was undergoing reduction or oxidation as the zircon suite crystallised. Zircon Ce4+/Ce3+ values reported by authors of the source publications (evaluated by the lattice-strain method of Blundy and Wood 1994; Ballard et al. 2002) are plotted against the U/Pr ratios in those zircons, except for South Sister and Mt St Helens suites, for which we calculated zircon Ce4+/Ce3+ by the same procedure. All arrays have regression slopes < 1, which implies oxidation of the melt during zircon crystallisation, as indicated in three panels on the right, in which the data trends are decomposed into vector components representing cooling (slope + 1.0) and oxidation (slope − 1.0). Standard error of the slope and intercept are ± 1σ. Similar vector diagrams apply to panels on the left. South Sister data are from Stelten and Cooper (2012); Mt St Helens data from Claiborne (2011) and Claiborne et al. (2018); El Teniente porphyry copper igneous complex (Chile) data from Muñoz et al. (2012); Lower Yangtze Valley data from Wang et al. (2013); Chuquicamata porphyry copper complex and Montecristo (Chile) data are from Ballard et al. (2002)

Fig. 11

Slopes of zircon composition arrays test whether the melt was undergoing reduction or oxidation as the zircon suite crystallised. The least-squares regression in each panel has a slope > 1.0, which is decomposed into vector components representing cooling (slope + 1) and decrease in oxidation state of the melt (slope − 1). Standard error of the slope and intercept are ± 1σ. In each panel the data trend shows that the redox state of the melt decreased as zircon crystallised from cooling melt. The trends of these zircon arrays corroborate the Fe–Ti-oxide T–ƒO2 trends for these volcanic suites in Fig. 1. a Bishop Tuff zircon crystal rims analysed by Smythe and Brenan (2016, sample BTZ) and by Reid et al. (2011; other sample numbers). The label following the comma is the eruptive stratigraphic unit from which the sample was collected (nomenclature of Hildreth and Wilson 2007). Calculation of zircon Ce4+/Ce3+ utilises analyses of glass inclusions in Bishop Tuff phenocrysts by Wallace et al. (1999), Anderson et al. (2000), and Roberge et al. (2013). b Zircon crystal–rim compositions in eruptive Unit B of the Yellowstone caldera-forming Lava Creek Tuff supereruption (0.62 Ma), as reported by Wotzlaw et al. (2015) are combined with Lava Creek Tuff Unit B pumice composition as reported by Christiansen (2001). c Toba Tuff zircons and host glass analyses from two pumice samples reported by Gaither (2011)

Suites showing no change in oxidation state during cooling

Electronic Supplementary Material 3 lists Tampakan zircon trace-element analyses by ArF ELA-ICPMS at the Research School of Earth Sciences, Australian National University. Our analyses of Tampakan zircons have been combined with our compilation from published literature of zircon compositions in all igneous suites for which Ce4+/Ce3+ in zircon was calculated by the respective authors using the lattice-strain method, and which have enough data points to perform a test of trend. No examples were found that show a negative correlation of U/Pr with Ce4+/Ce3+, as would represent variation of Ce4+/Ce3+ dominated by changing oxidation state of the melt. Figure 9 shows strong positive correlations of zircon Ce4+/Ce3+ with U/Pr in eight igneous complexes. In each of them, the regression slope is near 1.0, which means that the variations in both ratios are attributable entirely to cooling during crystallisation. Accordingly, there is no resolvable variation in redox state of the melt (as represented by redox indicators in Fig. 3) within any zircon suite shown in Fig. 9, including the Tampakan suite that is also illustrated in Fig. 5. Of the igneous suites represented in Fig. 9, the Tampakan volcanic suite is the only one for which analyses of Fe–Ti-oxide phenocrysts are available for independent corroboration of the T–ƒO2 trend (Electronic Supplementary Material 2). The Tampakan Fe–Ti oxide T–ƒO2 array in Fig. 1 shows no resolvable change in oxidation state of the melt over the 920–705 °C temperature interval, which corroborates the redox independence of the Tampakan zircon trend in Fig. 9 showing ~ 1:1 co-variation of U/Pr with Ce4+/Ce3+.

Suites showing oxidation during cooling

In Fig. 10, we show a collection of zircon analyses from rhyolitic flow domes at South Sister volcano (Cascade arc) and from the 3.5 ka Yn tephra erupted from Mt St Helens (Cascade arc) and from six subduction-related calc-alkalic intrusions. All except South Sister and Montecristo are “adakitic” (high Sr/Y and La/Yb, modest or no Eu anomaly in the whole-rock REE pattern). Montecristo is a pre-ore intrusion in the Chuquicamata district. The chondrite-normalised REE patterns of the South Sister’s Devil’s Hills rhyolite lava domes are “adakite-like”, steeply dipping, with a minimum at Ho and LaN/YbN = 14 and Eu/Eu* = 0.96 to 0.97 in the two rock analyses reported by Stelten and Cooper (2012).

All the Ce4+/Ce3+ values plotted in Fig. 10 are those reported in the cited publications (following the calculation procedure of Ballard et al. 2002, using whole-rock compositions), except South Sister and Mt St Helens Yn tephra, which we calculated using glass analyses reported by the respective authors (Electronic Supplementary Material 4). All data arrays in Fig. 10 have regression slopes significantly < 1. As indicated by the inset vector-decompositions of the data trends, we interpret these regression trends as signifying magma oxidation during crystallisation of the analysed zircon suites. The Devil’s Hills rhyolitic lava domes in the South Sister volcanic complex and the Mt St Helens Yn tephra are the only suites in Fig. 10 for which Fe–Ti oxide data are available for an independent test of redox trend. The negative slopes of the Mt St Helens Yn tephra and South Sister Fe–Ti oxide array in Fig. 1 corroborate the zircon-based U/Pr versus Ce4+/Ce3+ trends in Fig. 10 that indicates melt oxidation during cooling and crystallisation.

Prouteau and Scaillet (2003), Rohrlach and Loucks (2005), Loucks (2014) and Lu et al (2015), among others, demonstrate that adakitic magmas are exceptionally hydrous, commonly with 8–14 wt% H2O dissolved in the silicate melt over most of the course of high-pressure magmatic differentiation. As silicate melts cannot dissolve that much H2O at upper-crustal pressures, it is implied that fluid exsolved during ascent of andesitic and dacitic melts to subvolcanic magma chambers, with crystallisation at fluid-saturated conditions upon epizonal emplacement. Oxidation of iron and other redox couples in the melt accompanies dehydration of the melt (Czamanske and Wones 1973; Candela 1986; Gaillard et al. 2002; Rowe et al. 2006). Figure 1 shows that the adakitic suites (Pinatubo, Daisen, Fish Canyon Tuff, and Mt St Helens Yn tephra) have systematically higher ƒO2 and negative T–ƒO2 slopes in the dacitic to rhyolitic composition range. We suggest that exsolution of hydrothermal fluid during magma ascent and epizonal crystallisation is accompanied by partitioning of oxidised sulphur from melt into hydrothermal fluid according to the heterogeneous reaction:

$$\mathop {{\text{S}}{{\text{O}}_{\text{3}}}}\limits^{{{\text{melt}}}} +{\text{ 2}}\mathop {{\text{FeO}}}\limits_{}^{{{\text{melt}}}} \to \mathop {{\text{F}}{{\text{e}}_{\text{2}}}{{\text{O}}_{\text{3}}}}\limits^{{{\text{melt}}}} +{\text{ }}\mathop {{\text{S}}{{\text{O}}_{\text{2}}}}\limits^{{{\text{fluid}}}} ,$$
wherein exsolution of one mole of S6+ in the melt as S4+ in the fluid is accompanied by oxidation of two moles of Fe2+ in the melt. The effect on Fe3+/Fe2+ ratio in the residual melt would be greatest in oxidised, Fe-poor felsic melts in which the S:Fe ratio is relatively large at the onset of fluid saturation. Scaillet et al. (1998) demonstrate experimentally that oxidised sulphur is much more volatile as SO2 than as SO3, and that at log ƒO2 > NNO + 1.3, the fluid/melt partition coefficient of sulphur is ≥ 1000 in dacitic and more silicic magmas, so this provides a plausible explanation of the negative T–ƒO2 slopes at high ƒO2 in Fig. 1 and slopes < 1 in Fig. 10. Experiments by Scaillet et al. (1998) indicate that the fluid/melt partition coefficient of sulphur is about 300 times smaller at log ƒO2 ≤ NNO + 0.5, in the stability field of S2− predominance in the melt. Therefore, gravitational segregation of buoyant hydrothermal fluid as it exsolves should induce some oxidation of the Ce4+/Ce3+ and U5+/U4+ and Fe3+/Fe2+ redox couples in the residual silicate melt and in zircons that grow during decompression and loss of SO3 to exsolving fluid.

Baker and Rutherford (1996) find in experiments at controlled ƒO2 that change of melt structure with variation of dissolved H2O content of metaluminous rhyolite has no effect on the Fe3+/Fe2+ ratio in the melt at temperatures below 850 °C. However, at magmatic temperatures molecular H2O in the melt and in exsolved fluid partially dissociates according to the reaction H2O \(\to\) H2 + ½ O2, and loss of H2 to segregating fluid oxidises iron in the residual melt by the reaction 2FeO + H2O \(\to\) Fe2O3 + H2. As pointed out by Carmichael (1991), in relatively H2O-poor silicic melts, such as the Toba Tuffs (Chesner and Luhr 2010) and Bishop Tuff (Wallace et al. 1999; Anderson et al. 2000; Roberge et al. 2013) and Yellowstone rhyolites (Almeev et al. 2012; Befus and Gardner 2016), most of the dissolved H2O occurs as hydroxyl. Therefore, the H2O molecular dissociation reaction does not proceed to as significant a degree as in extremely hydrous magmas.

H2 transfer from silicate melt to exsolving hydrothermal fluid and gravitational segregation of buoyant bubbles from the conjugate melt (forming a fluid cap at the top of the magma column) may provide a supplementary mechanism for inducing the oxidising–cooling trend revealed by the zircon data in Fig. 10 and by the Fe–Ti-oxide T–ƒO2 trends (Fig. 1) of the low-temperature (< 800 °C) dacitic/rhyolitic segments of the Pinatubo and St Helens Yn arrays and similar cooling–oxidising trends of other adakitic suites (e.g., Daisen and Fish Canyon). Czamanske and Wones (1973) attribute to segregation of exsolving magmatic hydrothermal fluid and H2 loss from the melt an oxidising–cooling trend during crystallisation of the Finnmarka granitoid complex, Norway. Cornejo and Mahood (1997) attribute to H2 loss the oxidising–cooling T–ƒO2 trend in the adakitic La Gloria pluton of Miocene age in central Chile.

Bell and Simon (2011) have identified an additional mechanism for increasing the Fe3+/Fe2+ ratio in the residual melt as hydrothermal fluid exsolves and physically segregates. Their experiments and thermodynamic data and modelling studies show that over a large range of ƒO2 relevant to natural magmas, Fe2+ partitions into a chloride-bearing magmatic hydrothermal fluid much more strongly than Fe3+, with the result that fractional segregation of the brine raises the oxidation state of the residual melt. The effect is especially pronounced during exsolution of a relatively large mass of brine from relatively Fe-poor felsic melts.

Suites showing reduction during cooling

In contrast to the differentiation trends of wet magmas that have slopes < 1 in Fig. 10 and negative slopes in Fig. 1, we show in Fig. 11 log U/Pr versus log Ce4+/Ce3+ trends for zircons in H2O-poor, fayalite- and pyroxene-phyric, high-silica rhyolites that have positive slopes of the T–ƒO2 trends in Fig. 1. The data and procedures for evaluating zircon Ce4+/Ce3+ in the suites plotted in Fig. 11 are shown in Electronic Supplementary Material 4. Those trends having positive slopes are a subject of dispute, as described in the introduction of this paper. In Fig. 1, the T–ƒO2 trend of the Yellowstone Lake Creek fayalite rhyolite and Toba Tuff pyroxene dacite and Toba Tuff high-silica fayalite rhyolite show positive T–ƒO2 slopes like the Bishop Tuff pyroxene rhyolite. The zircon composition trends in all three panels of Fig. 11 have slopes substantially > 1, which corroborate the cooling–reducing trend from Fe–Ti oxides for those three suites shown in Fig. 1.

In each of these volcanic suites in Fig. 11, and in the Tampakan and South Sister and Mt St Helens Yn volcanic suites in Figs. 9 and 10, the data dispersion in each array is similar to the degree of dispersion within T–ƒO2 arrays of Fe–Ti oxide phenocrysts from the same volcanic suites in Fig. 1. The Bishop Tuff and Yellowstone zircon Ce4+/Ce3+ values in Fig. 11 were calculated by pairing the average zircon “rim” composition (outer 10–15 µm) in each pumice sample with glass analyses from the same eruptive unit and sampling area, reported by other analysts, to derive apparent zircon/melt partition coefficients of REEs, Th, Zr and Hf, which were then used to estimate Ce4+/Ce3+ in zircon according to the lattice-strain method of Blundy and Wood (1994, 2003) (see Electronic Supplementary Material 4). The assumption that zircon and glass represent equilibrium pairs may not be perfect in every case, because chamber replenishment and magma mixing that immediately preceded the eruptions have been well documented in the Bishop Tuff (Evans and Bachmann 2013), Toba Tuff (Chesner and Luhr 2010), and Yellowstone Lake Creek Tuff (Wotzlaw et al. 2015). Such possible heterogeneity in composition of the host glass may contribute analytical dispersion in addition to the compositional heterogeneity within zoned crystals (Fig. 8).

The Bishop Tuff data trend in Fig. 11 implies that successive eruptive units tapped successively deeper levels of a magma chamber that was zoned in oxidation state, with early-erupted, more reduced, cooler magma overlying hotter, less reduced magma that was tapped later in the eruption sequence, corroborating interpretations by Hildreth and Wilson (2007) based on independent evidence from Fe–Ti oxides (Fig. 1) and other phenocrysts. Their interpretation that the Bishop Tuff magma chamber had strong vertical gradients in temperature and oxidation state was disputed by Gualda et al. (2012, 2017) who claimed, on the basis of Rhyolite-MELTS modelling, that the crystallisation temperature for Bishop Tuff phenocrysts lies within a 10 °C range from top to bottom of the eruptive sequence.

The cause of the reducing trend during cooling of the Bishop Tuff magma is illuminated by consideration of Fe3+/Fe2+ ratios in crystals and coexisting melt. Hildreth and Wilson (2007) report that ilmenite phenocrysts are exceedingly sparse in all eruptive units of the Bishop Tuff, and very subordinate to titanomagnetite in modal abundance. Hildreth (1977) carries out wet chemical measurements of Fe3+/Fe2+ ratios in biotite in several eruptive units of the Bishop Tuff and reports that 40–80% of the iron in biotite is Fe3+. Multiple lines of evidence summarized by Evans et al. (2016) and references therein show that the pyroxene phenocrysts found in late-erupted units that tapped deeper, hotter parts of the magma chamber are xenocrysts out of chemical equilibrium with all other phenocrysts and glass in the samples. Therefore, the only modally significant Fe-bearing phenocrysts precipitating from the melts represented by erupted pumice are titanomagnetite with Fe3+/Fe2+ ≈ 1.2 on average (Hildreth and Wilson 2007) and biotite with a similar Fe3+/Fe2+ ratio. At 800–850 °C, the Bishop melt was at log ƒO2 ≈ NNO + 0.5 (Fig. 1), so Fe3+/Fe2+ in the melt was ~ 0.06 to 0.07 (Gaillard et al. 2001; Wilke et al. 2002). Precipitation of a phenocryst assemblage having a higher Fe3+/Fe2+ than the high-silica, Fe–Mg-poor melt can account for the evolution of the melt toward lower Fe3+/Fe2 with decreasing temperature, as illustrated in Figs. 1 and 11a. Zircon arrays from slowly cooled intrusive suites plotted in Figs. 9 and 10 tend to show less data dispersion.

Implications for formulation of TiO2 activity by the MELTS and Rhyolite-MELTS thermodynamic models

Rutile solubility experiments in hydrous granitoid melts performed and compiled by Kularatne and Audétat (2014) underpin their thermodynamic model for TiO2 activity in rhyolitic melts. They apply their TiO2 activity model to quartz-hosted glass inclusions in metaluminous and peraluminous rhyolitic volcanic rocks to calculate values of TiO2 activity that are in good agreement with TiO2 activities calculated from Fe–Ti-oxide phenocrysts in those samples using the method of Ghiorso and Gualda (2013). However, Kularatne and Audétat (2014) report that titania solubilities predicted by MELTS (Ghiorso and Sack 1995) and Rhyolite-MELTS (Gualda et al. 2012) free-energy-minimisation programs are 1.5–2.9 times higher than actually observed TiO2 solubilities in all subaluminous and peraluminous melt compositions investigated in their experimental dataset. Kularatne and Audétat’s experiments indicate that rutile solubility is insensitive to alumina saturation index in metaluminous melts having Al/Na + K ≥ 1, but rutile solubility in peralkaline melts increases strongly with decreasing Al/Na + K. The sensitivity of rutile solubility to alumina saturation index in peralkaline melts is consistent with the formation of alkali–titanate complexes as suggested by Dickinson and Hess (1985) and later confirmed by X-ray absorption spectroscopy (e.g., Farges et al. 1996; Farges and Brown 1997).

We infer that the errors in the formulation for \(a_{{{\text{Ti}}{{\text{O}}_{\text{2}}}}}^{{{\text{melt}}}}\) by the MELTS and Rhyolite-MELTS thermodynamic models, which underlie the dispute over T–ƒO2 trends described in the introduction, may be due to failure to adequately account for the varying structural role of Ti with varying alumina saturation index and hydration state of the melt, with Ti being largely a “network forming” component (tetrahedrally coordinated with oxygen) in low-pressure peralkaline melts and predominantly a “network-modifying” component (octahedrally coordinated, as in rutile) in peraluminous melts and high-pressure metaluminous melts. If so, then the enthalpy of fusion of rutile must be combined with variable enthalpy of solvation (mixing) of molten TiO2 in aluminosilicate melts of varied A/CNK and hydration state (taking Na–OH complexing into account; Kohn et al. 1998), to thermodynamically model \(a_{{{\text{Ti}}{{\text{O}}_{\text{2}}}}}^{{{\text{melt}}}}\) in thermally and compositionally varied aluminosilicate melts. Our literature survey indicates that rutile saturation at mid- to upper-crustal pressures in nature is largely limited to peraluminous granitoids. We expect that the configurational behavior of Ti in silicate melts resembles that of Al, which is entirely a tetrahedrally coordinated “network former” in low-pressure peralkaline melts (in which liquidus minerals rich in octahedral Al are unstable), but largely an octahedrally coordinated “network modifier” in peraluminous melts and high-pressure melts that stabilise liquidus minerals rich in octahedrally coordinated Al. By extension of the Burnham (1981) “quasi-crystalline model” for configurational properties of aluminosilicate liquids, MELTS-type thermodynamic models may benefit by transitioning among pressure- and bulk-composition-dependent alternative sets of thermodynamic components of the melt (e.g., using the compositions and fusion enthalpies of CaTS or grossular rather than anorthite, and of jadeite rather than albite) to simplify treatment of configurational entropy and enthalpic mixing properties of aluminosilicate melts over a range of mantle to crustal pressures and varied alumina saturation index. Thermodynamic properties of biotite solid solutions, with particular attention to effective biotite/melt partition coefficients of Fe3+ and Fe2+, may also warrant re-examination in the MELTS and Rhyolite-MELTS programs.


This work presents new evidence that shows that the Fe–Ti-oxide oxybarometer calibration by Ghiorso and Evans (2008) retrieves experimentally imposed values of ƒO2 in laboratory syntheses of Fe–Ti-oxide pairs to a precision of ± 0.2 log unit, over a large experimental temperature range, without systematic bias up to at least log ƒO2 ≈ NNO + 4.4. Proceeding from that result, this study addresses a recent dispute (Ghiorso and Gualda 2013; Ghiorso et al 2016; Evans and Bachmann 2013; Evans et al. 2016) regarding the reliability of T–fO2 trends in the Bishop Tuff magma chamber recorded by titanomagnetite–hemoilmenite phenocryst pairs and trends recovered by Trail et al. (2015) from zircon Ce4+/Ce3+ ratios and Ti-in-zircon temperatures in the Bishop Tuff. Results from our study show that lattice-strain effects cause Ce4+/Ce3+ in zircon to increase strongly as zircon crystallises from cooling magma at constant Ce4+/Ce3+ ratio in the melt. In many natural assemblages described in published literature and interpreted as redox indicators, there is no resolvable change in redox state of the melt over a large range of Ce4+/Ce3+ variation in zircons. We indicate that those magma chambers have crystallised at constant redox state of the melt in systems closed to oxygen (or hydrogen) exchange with the environs. Exceptionally H2O-rich arc magmas tend to follow a trend of rising oxidation state of the melt during late stages of fluid-saturated magmatic differentiation at upper-crustal pressures. We suggest that SO3 and/or H2 and/or Fe2+ loss from the melt are responsible. H2O-poor magmas tend to follow a T–ƒO2 trend of decreasing oxidation state of the melt during late stages of magmatic differentiation at upper-crustal pressures, because the precipitating mineral assemblage has higher Fe3+/Fe2+ than coexisting rhyolitic melt. In conclusion, this study corroborates the validity of the Ghiorso and Evans (2008) Fe–Ti-oxide oxybarometer, and supports the methodology of Ghiorso and Gualda (2013) for evaluating TiO2 activity from quenched Fe–Ti-oxide phenocrysts, and thereby supports use of those TiO2 activity values in application to Ti-in-zircon and Ti-in-quartz geothermometry to address a wide range of fundamental geological problems. We offer suggestions for improving the Rhyolite-MELTS program of Gualda et al. (2012).



We thank Paul Agnew and Alan Kobussen of Rio Tinto Exploration for financial support and for authorisation to publish research results. Additional financial support was provided by the ARC Centre of Excellence for Core to Crust Fluid Systems Grant CE110001017. MLF acknowledges the support of the Australian Research Council through the Future Fellowship Grant scheme (FT110100241). MLF also acknowledges support from the Australian Research Council through Linkage Project LP120100668. BDR acknowledges WMC Resources Ltd for financial support of his PhD research. RRL thanks the University of Bristol for a Benjamin Meaker Visiting Professorship which afforded an opportunity to clarify his understanding of this material through discussions with Professors Jon Blundy and Chris Hawkesworth. Journal reviews by Jon Blundy and an anonymous reviewer improved the clarity of presentation. This is contribution 1223 from the ARC Centre for Excellence for Core to Crust Fluid Systems (

Supplementary material

410_2018_1529_MOESM1_ESM.xls (905 kb)
Supplementary material 1 (xls 905 kb)
410_2018_1529_MOESM2_ESM.xls (68 kb)
Supplementary material 2 (xls 68 kb)
410_2018_1529_MOESM3_ESM.xls (60 kb)
Supplementary material 3 (xls 60 kb)
410_2018_1529_MOESM4_ESM.xls (224 kb)
Supplementary material 4 (xls 224 kb)


  1. Almeev RR, Bolte T, Nash BP, Holtz F, Erdmann M, Cathey HE (2012) High-temperature, low-H2O silicic magmas of the Yellowstone hotspot: an experimental study of rhyolite from the Bruneau–Jarbidge eruptive center, central Snake River Plain, USA. J Petrol 53(9):1837–1866Google Scholar
  2. Andersen DJ, Lindsley DH, Davidson PM (1993) QUILF: a Pascal program to assess equilibria among Fe–Mg–Mn–Ti oxides, pyroxenes, olivine and quartz. Comput Geosci 19(9):1333–1350Google Scholar
  3. Anderson AT, Davis AM, Lu F (2000) Evolution of the Bishop Tuff rhyolitic magma based on melt and magnetite inclusions and zoned phenocrysts. J Petrol 41(3):449–473Google Scholar
  4. Baker LL, Rutherford MJ (1996) The effect of dissolved water on the oxidation state of silicic melts. Geochim Cosmochim Acta 60(12):2179–2187Google Scholar
  5. Ballard JR, Palin JM, Campbell IH (2002) Relative oxidation states of magmas inferred from Ce(IV)/Ce(III) in zircon: application to porphyry copper deposits of northern Chile. Contrib Mineral Petrol 144(3):347–364Google Scholar
  6. Befus KS, Gardner JE (2016) Magma storage and evolution of the most recent effusive and explosive eruptions from Yellowstone Caldera. Contrib Mineral Petrol 171:30Google Scholar
  7. Bell AS, Simon A (2011) Experimental evidence for the alteration of the Fe3+/ΣFe of silicate melt caused by the degassing of chlorine-bearing aqueous volatiles. Geology 39(5):499–502Google Scholar
  8. Berndt J, Koepke J, Holtz F (2005) An experimental investigation of the influence of water and oxygen fugacity on differentiation of MORB at 200 MPa. J Petrol 46(1):135–167Google Scholar
  9. Berry AJ, O’Neill HSTC, Foran GJ (2008) The oxidation state of uranium in mantle melts. Geochim Cosmochim Acta 72(12):A79Google Scholar
  10. Bindeman IN, Valley JW (2002) Oxygen isotope study of the Long Valley magma system, California: isotope thermometry and convection in large silicic magma bodies. Contrib Mineral Petrol 144:185–205Google Scholar
  11. Blatter DL, Sisson TW, Hankins WB (2013) Crystallization of oxidized, moderately hydrous arc basalt at mid- to lower-crustal pressures: implications for andesite genesis. Contrib Mineral Petrol 166:861–886Google Scholar
  12. Blundy J, Wood B (1994) Prediction of crystal-melt partition coefficients from elastic moduli. Nature 372(6505):452–454Google Scholar
  13. Blundy J, Wood B (2003) Mineral-melt partitioning of uranium, thorium and their daughters. Revs Mineral Geochem 52:59–123Google Scholar
  14. Blundy J, Cashman K, Humphreys M (2006) Magma heating by decompression-driven crystallisation beneath andesite volcanoes. Nature 443:76–80Google Scholar
  15. Blundy J, Cashman K, Berlo K (2008) Evolving magma storage conditions beneath Mount St. Helens inferred from chemical variations in melt inclusions from the 18979–1986 and current (2004–2006) eruptions. USGS Prof Pap 1750:755–790Google Scholar
  16. Brice JC (1975) Some thermodynamic aspects of the growth of strained crystals. J Cryst Growth 28:249–253Google Scholar
  17. Burgisser A, Scaillet B (2007) Redox evolution of a degassing magma rising to the surface. Nature 445:194–197Google Scholar
  18. Burnham CW (1981) The nature of multicomponent aluminosilicate melts. Phys Chem Earth 13/14:197–229Google Scholar
  19. Burnham AD, Berry AJ (2012) An experimental study of trace element partitioning between zircon and melt as a function of oxygen fugacity. Geochim Cosmochim Acta 95:196–212Google Scholar
  20. Burnham AD, Berry AJ (2014) The effect of oxygen fugacity, melt composition, temperature and pressure on the oxidation state of cerium in silicate melts. Chem Geol 366:52–60Google Scholar
  21. Burnham AD, Berry AJ (2017) Formation of Hadean granites by melting of igneous crust. Nat Geosci 10:457–461Google Scholar
  22. Candela PA (1986) The evolution of aqueous vapor from silicate melts: effect on oxygen fugacity. Geochim Cosmochim Acta 50:1205–1211Google Scholar
  23. Carmichael ISE (1967) The mineralogy of Thingmuli, a Tertiary volcano in eastern Iceland. Amer Mineral 52:1815–1841Google Scholar
  24. Carmichael ISE (1991) The redox states of basic and silicic magmas: a reflection of their source regions? Contrib Mineral Petrol 106(2):129–141Google Scholar
  25. Carmichael ISE, Ghiorso MS (1986) Oxidation-reduction relations in basic magma: a case for homogeneous equilibria. Earth Planet Sci Lett 78(2–3):200–210Google Scholar
  26. Cathey HE, Nash BP (2004) The Cougar Point Tuff: implications for thermochemical zonation and longevity of high-temperature, large-volume silicic magmas of the Miocene Yellowstone hotspot. J Petrol 45(1):27–58Google Scholar
  27. Chamberlain KJ, Wilson CJN, Wooden JL, Charlier BLA, Ireland TR (2014) New perspectives on the Bishop Tuff from zircon textures, ages, and trace elements. J Pet 55:395–426Google Scholar
  28. Chamberlain KJ, Wilson CJN, Wallace PJ, Millet M-A (2015) Micro-analytical perspectives on the Bishop Tuff and its magma chamber. J Petrol 56:605–640Google Scholar
  29. Cherniak DJ, Watson EB (2003) Diffusion in zircon. Rev Mineral Geochem 53:113–143Google Scholar
  30. Chesner CA (1998) Petrogenesis of the toba tuffs. Sumatra Indonesia J Petrol 39:397–438Google Scholar
  31. Chesner CA, Luhr JF (2010) A melt inclusion study of the TobaTuffs, Sumatra, Indonesia. J Volc Geoth Res 197:259–278Google Scholar
  32. Christiansen RL (2001) The Quaternary and Pliocene Yellowstone Plateau volcanic field of Wyoming, Idaho, and Montana. US Geol Surv Prof Pap 729-G:156Google Scholar
  33. Claiborne LL (2011) Understanding upper crustal silicic magmatic systems using the temporal, compositional, and thermal record in zircon. Ph.D. Dissertation, Vanderbilt University, p 385Google Scholar
  34. Claiborne LL, Miller CF, Gualda GAR, Carley TL, Covey AK, Wooden JL, Fleming MA (2018) Zircon as magma monitor: robust, temperature-dependent partition coefficients from glass and zircon surface and rim measurements from natural systems. In: Moser DE, Corfu F, Darling JR, Reddy SM, Tait K (eds) Microstructural geochronology: planetary records down to atom scale. Geophysical Monograph 232, American Geophysical Union/Wiley, Oxford, p 3–33Google Scholar
  35. Cornejo PC, Mahood GA (1997) Seeing past the effects of re-equilibration to reconstruct magmatic gradients in plutons: La Gloria Pluton, central Chilean Andes. Contrib Mineral Petrol 127:159–175Google Scholar
  36. Crowley JL, Schoene B, Bowring SA (2007) U-Pb dating of zircon in the Bishop Tuff at the millennial scale. Geology 35:1123–1126Google Scholar
  37. Czamanske GK, Wones DR (1973) Oxidation during magmatic differentiation, Finnmarka Complex, Oslo area, Norway: Part 2, the mafic silicates. J Petrol 14:349–380Google Scholar
  38. Delano J (2001) Redox history of the Earth’s interior since ~ 3900 Ma: implications for prebiotic molecules. Orig Life Evol Biosph 31:311–341Google Scholar
  39. Dickinson JE, Hess PC (1985) Rutile solubility and titanium coordination in silicate melts. Geochim Cosmochim Acta 49:2289–2296Google Scholar
  40. Evans BW, Bachmann O (2013) Implications of equilibrium and disequilibrium among crystal phases in the Bishop Tuff. Am Mineral 98:271–274Google Scholar
  41. Evans BW, Hildreth W, Bachmann O, Scaillet B (2016) In defense of magnetite-ilmenite thermometry in the Bishop Tuff and its implication for gradients in silicic magma reservoirs. Am Mineral 101(2):469–482Google Scholar
  42. Farges F, Brown GE Jr, Rehr JJ (1996) Coordination chemistry of Ti(IV) in silicate glasses and melts: I. XAFS study of titanium coordination in oxide model compounds. Geochim Cosmochim Acta 60:3023–3038Google Scholar
  43. Farges F, Brown GE Jr (1997) Coordination chemistry of Ti(IV) in silicate glasses and melts: IV. XANES study of synthetic and natural volcanic glasses and tektites at ambient temperature and pressure. Geochim Cosmochim Acta 61:1863–1870Google Scholar
  44. Farges F, Ponader W, Calas G, Brown GE Jr (1992) Structural environments of incompatible elements in silicate glass/melt systems: II. UIV, UV, and UVI. Geochim Cosmochim Acta 56:4205–4220Google Scholar
  45. Ferry JM, Watson EB (2007) New thermodynamic models and revised calibrations for the Ti-in-zircon and Zr-in-rutile thermometers. Contrib Mineral Petrol 154:429–437Google Scholar
  46. Fonseca ROC, Mallmann G, Sprung P, Sommer JE, Heuser A, Speelmanns IM, Blanchard H (2014) Redox controls on tungsten and uranium crystal/silicate melt partitioning and implications for the U/W and Th/W ratio of the lunar mantle. Earth Planet Sci Lett 404:1–13Google Scholar
  47. Gaillard F, Scaillet B, Pichavant M, Bény J-M (2001) The effect of water and ƒO2 on ferric-ferrous ratio of silicic melts. Chem Geol 174:255–273Google Scholar
  48. Gaillard F, Scaillet B, Pichavant M (2002) Kinetics of iron oxidation-reduction in hydrous silicic melts. Amer Mineral 87:829–837Google Scholar
  49. Gaither TA (2011) Zircon geochemistry and geothermometry of the youngest Toba Tuff. Unpub. MSc. Thesis, Northern Arizona University, p 154Google Scholar
  50. Gardner JE, Carey S, Rutherford MJ, Sigurdsson H (1995) Petrologic diversity in Mount St. Helens dacites during the last 4,000 years: implications for magma mixing. Contrib Mineral Petrol 119:224–238Google Scholar
  51. Gardner JE, Befus K, Gualda GAR, Ghiorso MS (2014) Experimental constraints on rhyolite-MELTS and the Late Bishop Tuff magma body. Contrib Mineral Petrol 168(2):1–14Google Scholar
  52. Ghiorso MS, Sack RO (1995) Chemical mass transfer in magmatic processes IV. A revised and internally consistent thermodynamic model for the interpolation and extrapolation of liquid-solid equilibria in magmatic systems at elevated temperatures and pressures. Contrib Mineral Petrol 119(2–3):197–212Google Scholar
  53. Ghiorso MS, Evans BW (2008) Thermodynamics of rhombohedral oxide solid solutions and a revision of the Fe–Ti two oxide geothermometer and oxygen-barometer. Am J Sci 308(9):957–1039Google Scholar
  54. Ghiorso MS, Gualda GAR (2013) A method for estimating the activity of titania in magmatic liquids from the compositions of coexisting rhombohedral and cubic iron–titanium oxides. Contrib Mineral Petrol 165:73–81Google Scholar
  55. Ghiorso MS, Gualda GAR, Pamukcu AS, Gardner JE (2016) Apparent temperature-redox trends in the Bishop Tuff: a result of rapid growth of Fe–Ti oxides under constant temperature/redox state conditions. Am Geophys Union Fall Mtg, Abs V24C-01Google Scholar
  56. Gualda GAR, Ghiorso MS, Lemons RV, Carley TL (2012) Rhyolite-MELTS: a modified calibration of MELTS optimized for silica-rich, fluid-bearing magmatic systems. J Pet 53:875–890Google Scholar
  57. Gualda GAR, Ghiorso MS, Hurst AA, Allen MC, Bradshaw RW (2017) Patchy distribution of magma that fed the Bishop Tuff supereruption: evidence from matrix glass major and trace-element compositions. Am Geophys Union Fall Mtg, Abs V11C-0369Google Scholar
  58. Han Y, Zhang S, Pirajno F, Zhou X, Zhao G, Qu W, Liu S, Zhang J, Liang H, Yang K (2013) U–Pb and Re–Os isotopic systematics and zircon Ce4+/Ce3+ ratios in the Shiyaogou Mo deposit in eastern Qinling, central China: insights into the oxidation state of granitoids and Mo (Au) mineralization. Ore Geol Revs 55:29–47Google Scholar
  59. Hanchar JM, Watson EB (2003) Zircon saturation thermometry. Revs Mineral Geochem 53:89–112Google Scholar
  60. Hildreth W (1977) The magma chamber of the Bishop Tuff: gradients, temperature, pressure and composition. Ph.D. thesis, University of California, Berkeley, California, p 328Google Scholar
  61. Hildreth W (1983) The compositionally zoned eruption of 1912 in the valley of ten thousand smokes, Katmai National Park, Alaska. J Volc Geoth Res 18:1–56Google Scholar
  62. Hildreth W, Wilson CJN (2007) Compositional zoning of the Bishop Tuff. J Pet 48(5):951–999Google Scholar
  63. Hofmann AE, Baker MB, Eiler JM (2014) Sub-micron-scale trace-element distributions in natural zircons of known provenance: implications for Ti-in-zircon thermometry. Contrib Mineral Petrol 168:1057Google Scholar
  64. Holland T, Blundy J (1994) Non-ideal interactions in calcic amphiboles and their bearing on amphibole-plagioclase thermometry. Contrib Mineral Petrol 116(4):433–447Google Scholar
  65. Huang R, Audétat A (2012) The titanium-in-quartz (TitaniQ) thermobarometer: a critical examination and re-calibration. Geochim Cosmochim Acta 84:75–89Google Scholar
  66. Kohn SC, Smith ME, Dirken PJ, van Eck ERH, Kentgens APM, Dupree R (1998) Sodium environments in dry and hydrous albite glasses: improved 23Na solid state NMR data and their implications for water dissolution mechanisms. Geochim Cosmochim Acta 62:79–87Google Scholar
  67. Kress VC, Carmichael ISE (1991) The compressibility of silicate liquids containing Fe2O3 and the effect of composition, temperature, oxygen fugacity and pressure on their redox states. Contrib Mineral Petrol 108:82–92Google Scholar
  68. Kularatne K, Audétat A (2014) Rutile solubility in hydrous rhyolite melts at 750–900 °C and 2 kbar, with application to titanium-in-quartz (TitaniQ) thermobarometry. Geochim Cosmochim Acta 125:196–209Google Scholar
  69. Loucks RR (2014) Distinctive composition of copper-ore-forming arc magmas. Austral J Earth Sci 61:5–16Google Scholar
  70. Lu Y-J, Loucks RR, Fiorentini ML, Yang ZM, Hou Z-Q (2015) Fluid flux melting generated postcollisional high Sr/Y copper-ore-forming water-rich magmas in Tibet. Geology 43:583–586Google Scholar
  71. Luo Y, Ayers JC (2009) Experimental measurements of zircon/melt trace-element partition coefficients. Geochim Cosmochim Acta 73:3656–3679Google Scholar
  72. Muñoz M, Charrier R, Fanning CM, Maksaev V, Deckart K (2012) Zircon trace element and O–Hf isotope analyses of mineralized intrusions from El Teniente ore deposit, Chilean Andes: constraints on the source and magmatic evolution of porphyry Cu–Mo related magmas. J Pet 53(6):1091–1122Google Scholar
  73. O’Neill HSC (1987) Quartz-fayalite-iron and quartz-fayalite-magnetite equilibria and the free energy of formation of fayalite (Fe2SiO4) and magnetite (Fe3O4). Am Mineral 72:67–75Google Scholar
  74. Perfit MR, Fornari DJ (1983a) Geochemical studies of abyssal lavas recovered by DSRV Alvin from the eastern Galapagos Rift, Inca Transform, and Ecuador Rift 3. Trace element abundances and petrogenesis. J Geophys Res 88:10551–10572Google Scholar
  75. Perfit MR, Fornari DJ (1983b) Geochemical studies of abyssal lavas recovered by DSRV alvin from eastern galapagos rift, inca transform, and ecuador rift 2. phase chemistry and crystallisation history. J Geophys Res 88:10530–10550Google Scholar
  76. Pichavant M, Montel J-M, Richard LR (1992) Apatite solubility in peraluminous liquids: experimental data and an extension of the Harrison-Watson model. Geochim Cosmochim Acta 56(10):3855–3861Google Scholar
  77. Pichavant M, Martel C, Bourdier JL, Scaillet B (2002) Physical conditions, structure, and dynamics of a zoned magma chamber: mount Peleé (Martinique, Lesser Antilles Arc). J Geophys Res 107(B5):2093Google Scholar
  78. Pownceby MI, O’Neill HSC (1994) Thermodynamic data from redox reactions at high temperatures. IV. Calibration of the Re–ReO2 oxygen buffer from EMF and NiO + Ni-Pd redox sensor measurements. Contrib Mineral Petrol 118(2):130–137Google Scholar
  79. Preece SJ, McGimsey RG, Westgate JA, Pearce NJG, Hart WK, Perkins WT (2014) Chemical complexity and source of the White River Ash, Alaska and Yukon. Geosphere 10:1020–1042Google Scholar
  80. Prouteau G, Scaillet B (2003) Experimental constraints on the origin of the 1991 Pinatubo dacite. Jour Petrol 44(12):2203–2241Google Scholar
  81. Reid MR, Vazquez JA, Schmitt AK (2011) Zircon-scale insights into the history of a supervolcano, Bishop Tuff, Long Valley, California, with implications for the Ti-in-zircon geothermometer. Contrib Mineral Petrol 161:293–311Google Scholar
  82. Righter K, Chesley JT, Geist D, Ruiz J (1998) Behavior of Re during magma fractionation: an example from Volcán Alcedo, Galápagos. J Pet 39:785–795Google Scholar
  83. Roberge J, Wallace PJ, Kent AJR (2013) Magmatic processes in the Bishop Tuff rhyolitic magma based on trace elements in melt inclusions and pumice matrix glass. Contrib Mineral Petrol 165:237–257Google Scholar
  84. Rohrlach BD (2002), Tectonic evolution, petrochemistry, geochronology and palaeohydrology of the Tampakan porphyry and high sulphidation epithermal Cu–Au deposit Mindanao, Philippines. Unpub PhD thesis, Australian National University, Canberra, Australia, p 499Google Scholar
  85. Rohrlach BD, Loucks RR (2005) Multi-million-year cyclic ramp-up of volatiles in a lower-crustal magma reservoir trapped below the Tampakan copper–gold deposit by Mio–Pliocene crustal compression in the southern Philippines. In: Porter TM (ed) Super porphyry copper & gold deposits—a global perspective, vol 2. PCG, Adelaide, pp 369–407Google Scholar
  86. Rowe MC, Nielsen RL, Kent AJR (2006) Anomalously high Fe contents in rehomogenized olivine-hosted melt inclusions from oxidised magmas. Am Mineral 91:82–91Google Scholar
  87. Rubatto D, Hermann J (2007) Experimental zircon/melt and zircon/garnet trace element partitioning and implications for the geochronology of crustal rocks. Chem Geol 241:38–61Google Scholar
  88. Scaillet B, Evans BW (1999) The 15 June 1991 eruption of Mount Pinatubo. I. Phase equilibria and pre-eruption P–T–ƒO2–ƒH2O conditions of the dacite magma. J Pet 40:381–411Google Scholar
  89. Scaillet B, Clemente B, Evans B, Pichavant M (1998) Redox control of sulfur degassing in silicic magmas. J Geophys Res 103:23937–23949Google Scholar
  90. Schreiber HD (1987) An electrochemical series of redox couples in silicate melts: a review and applications to geochemistry. J Geophys Res 92:9225–9232Google Scholar
  91. Schreiber HD, Andrews SM (1980) The redox states of uranium in synthetic basaltic magmas. [Abstract]. In: Lunar and planetary science conference proceedings XI, pp 1000–1002Google Scholar
  92. Schreiber HD, Merkel RC Jr, Schreiber VL, Balazs GB (1987) Mutual interactions of redox couples via electron exchange in silicate melts: models for geochemical melt systems. J Geophys Res 92:9233–9245Google Scholar
  93. Shannon RD (1976) Revised effective ionic radii and systematic studies of interatomic distances in halides and chalcogenides. Acta Crystall A32:751–767Google Scholar
  94. Shen P, Hattori K, Pan H, Jackson S, Seitmuratova E (2015) Oxidation condition and metal fertility of granitic magmas: zircon trace-element data from porphyry Cu deposits in the Central Asian Orogenic Belt. Econ Geol 110(7):1861–1878Google Scholar
  95. Smythe DJ, Brenan JM (2015) Cerium oxidation state in silicate melts: combined fO2, temperature and compositional effects. Geochim Cosmochim Acta 170:173–187Google Scholar
  96. Smythe DJ, Brenan JM (2016) Magmatic oxygen fugacity estimated using zircon-melt partitioning of cerium. Earth Planet Sci Lett 453:260–266Google Scholar
  97. Stelten ME, Cooper KM (2012) Constraints on the nature of the subvolcanic reservoir at South Sister volcano, Oregon, from U-series dating combined with sub-crystal trace-element analysis of plagioclase and zircon. Earth Planet Sci Lett 313–314:1–11Google Scholar
  98. Tailby ND, Walker AM, Berry AJ, Hermann J, Evans KA, Mavrogenes JA, O’Neill HSC, Rodina IS, Soldatov AV, Rubatto D, Sutton SR (2011) Ti site occupancy in zircon. Geochim Cosmochim Acta 75:905–921Google Scholar
  99. Thomas JB, Watson BE (2012) Application of the Ti-in-quartz thermobarometer to rutile-free systems. Reply to: a comment on: “TitaniQ under pressure: the effect of pressure and temperature on the solubility of Ti in quartz” by Thomas et al. Contrib Mineral Petrol 164:369–374Google Scholar
  100. Trail D, Watson EB, Tailby ND (2011) The oxidation state of Hadean magmas and implications for early Earth’s atmosphere. Nature 480(7375):79–82Google Scholar
  101. Trail D, Watson EB, Tailby ND (2012) Ce and Eu anomalies in zircon as proxies for the oxidation state of magmas. Geochim Cosmochim Acta 97(1):70–87Google Scholar
  102. Trail D, Tailby ND, Lanzirotti A, Newville M, Thomas JB, Watson EB (2015) Redox evolution of silicic magmas: insights from XANES measurements of Ce valence in Bishop Tuff zircons. Chem Geol 402:77–88Google Scholar
  103. Venezky DY, Rutherford MJ (1997) Pre-eruption conditions and timing of dacite-andesite mixing in the 2.2 ka eruption of Mount Rainier. J Geophys Res 102(B9):20,068–020,086Google Scholar
  104. Wallace PJ, Anderson AT, Davis AM (1999) Gradients in H2O, CO2, and exsolved gas in a large-volume silicic magma system: interpreting the record preserved in melt inclusions from the Bishop Tuff. J Geophys Res, 140(B9):20,097–20,122Google Scholar
  105. Wang F, Liu S-A, Li S, He Y (2013) Contrasting zircon Hf–O isotopes and trace elements between ore-bearing and ore-barren adakitic rocks in central-eastern China: implications for genetic relation to Cu–Au mineralization. Lithos 156–159:97–111Google Scholar
  106. Wark DA, Watson EB (2006) TitaniQ: a titanium-in-quartz geothermometer. Contrib Mineral Petrol 152:743–754Google Scholar
  107. Watson EB, Harrison TM (1983) Zircon saturation revisited: temperature and composition effects in a variety of crustal magma types. Earth Planet Sci Lett 64:295–304Google Scholar
  108. Wilke M, Behrens H, Burkhard DJM, Rossano S (2002) The oxidation state of iron in silicic melt at 500 MPa water pressure. Chem Geol 189:55–67Google Scholar
  109. Wotzlaw J-F, Bindeman IN, Stern RA, D’Abzac F-X, Schaltegger U (2015) Rapid heterogeneous assembly of multiple magma reservoirs prior to Yellowstone supereruptions. Nat Sci Rep 5:14026Google Scholar
  110. Yakymchuk C, Clark C, White RW (2017) Phase relations, reaction sequences and petrochronology. Revs Mineral Geochem 83:13–53Google Scholar

Copyright information

© The Author(s) 2018

Open AccessThis article is distributed under the terms of the Creative Commons Attribution 4.0 International License (, which permits unrestricted use, distribution, and reproduction in any medium, provided you give appropriate credit to the original author(s) and the source, provide a link to the Creative Commons license, and indicate if changes were made.

Authors and Affiliations

  1. 1.Centre for Exploration Targeting, School of Earth SciencesUniversity of Western AustraliaCrawleyAustralia
  2. 2.ARC Centre of Excellence in Core to Crust Fluid SystemsUniversity of Western AustraliaCrawleyAustralia
  3. 3.Research School of Earth SciencesAustralian National UniversityActonAustralia
  4. 4.Sunstone Metals Ltd.MiltonAustralia

Personalised recommendations