Triple oxygen isotope exchange between chondrule melt and water vapor: An experimental study

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ScienceDirect Geochimica et Cosmochimica Acta 164 (2015) 17–34 www.elsevier.com/locate/gca

Triple oxygen isotope exchange between chondrule melt and water vapor: An experimental study Tommaso Di Rocco ⇑, Andreas Pack Geowissenschaftliches Zentrum, Abteilung Isotopengeologie, Georg-August-Universita¨t, Goldschmidtstraße 1, D-37077 Go¨ttingen, Germany Received 17 August 2014; accepted in revised form 22 April 2015; available online 30 April 2015

Abstract We have conducted time and f O2-dependent oxygen isotope exchange experiments between chondrule analogue melts and H2O in the phase. The aim of our study is to address the question whether the oxygen isotope composition of chondrules is the result of exchange with the ambient nebular gas or has been inherited from the precursor material. The silicate melt-H2O vapor exchange experiments were carried out in a vertical gas-mixing furnace using the metal loop technique at 1500 °C. The duration ranged from 5 to 1440 min and f O2 was set between IW  3.8 and IW  1.3 using the H2O/H2 buffer. Our experiments show that 50% exchange between H2O gas and silicate melt occurs in 4 h at f O2 = IW  3.8 and in 1 h at f O2 = IW  1.3. At solar nebula conditions, significant exchange occurs only if chondrule-melting times were several hours. Ó 2015 Elsevier Ltd. All rights reserved.

1. INTRODUCTION Chondrules are 0.1–1 mm-sized roundish objects that are frequent constituents (up to 80 vol%) of most chondrites (e.g. Zanda, 2004). Chondrules are chemically classified into low-FeO type-I and high-FeO type-II chondrules. Type-I chondrules contain <10 wt.% FeO in olivine, are generally poor in moderately volatile elements, and often contain Fe, Ni-metal and/or troilite (e.g. Hewins, 1997; Jones et al., 2005). Formation of type-I chondrules occurred at f O2 between IW  4 and IW  2 (Zanda et al., 1994; Lauretta et al., 2001; Schrader et al., 2013). Type-II chondrules contain >10 wt.% FeO in olivine, formed under more oxidizing conditions, and are richer in the moderately volatile elements. They are metal poor or metal free and formed at f O2 between IW  2 and IW (Lofgren, 1989; Schrader et al., 2013).

⇑ Corresponding author.

E-mail address: (T. Di Rocco).

[email protected]

http://dx.doi.org/10.1016/j.gca.2015.04.038 0016-7037/Ó 2015 Elsevier Ltd. All rights reserved.

One of the most distinctive features of chondrules is their O-isotope heterogeneity at different scales (reflected in bulk and single grain data; e.g. Clayton, 1993; Weinbruch et al., 1993; Leshin et al., 1997; Russell et al., 2000; Maruyama and Yurimoto, 2003; Jones et al., 2004; Pack et al., 2004; Krot et al., 2006a; Chaussidon et al., 2008). Chondrules were melted to near or super liquidus temperatures while freely floating in space. If chondrules were melted in a gas (e.g. of solar composition), exchange between the chondrule liquid and the gas may have occurred. The main carriers of O in a gas of solar composition are CO and H2O (Lodders and Fegley, 1998). It has been shown by Boesenberg et al. (2005) that exchange between CO and silicate melt does not occur. Instead, H2O was shown to be a reactive exchange partner (Yu et al., 1995). The isotope composition of solar nebula H2O may have been modified by CO self-shielding to high D17O (Clayton, 2002; Yurimoto and Kuramoto, 2002, 2004; Lyons and Young, 2005). Exchange with such 16O-poor H2O gas could be responsible for the mass independent variations in O isotopic composition of chondrules (Clayton et al., 1977, 1983; Jones et al., 2004). In such a scenario, chondrules would have behaved as open system with

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respect to oxygen. If, instead, melting times were too short or the partial pressure of H2O too low to permit exchange, chondrules would have behaved as closed system with respect to oxygen. In that scenario, the isotopic variability of chondrules would have been inherited from the precursor material. Arguments in favor of chondrule formation as chemically and isotopically open or closed system are presented in the following literature (e.g. Huang et al., 1996; Sears et al., 1996; Tissandier et al., 2002; Hezel et al., 2003, 2010; Libourel et al., 2006; Libourel and Krot, 2007 for open system; Grossman, 1988; Jones, 1990; Humayun and Clayton, 1995; Alexander et al., 2008; Borisov et al., 2008 for closed system). Apart of the inter-chondrule heterogeneity, also isotopic heterogeneity within chondrules has been observed. Chaussidon et al. (2008) determined the O-isotope composition of olivine, orthopyroxene, silica and mesostasis glass in a number of chondrules from unequilibrated chondrites. They show that D17O typically increases from olivine to low-Ca pyroxene to mesostasis (see also intra-chondrule heterogeneity reported by Maruyama et al., 1999, and Maruyama and Yurimoto, 2003; Pack et al., 2004). Chaussidon et al. (2008) conclude that this heterogeneity is due to a non-cogenetic origin of olivine (low D17O) and mesostasis (high D17O), that some chondrule olivines may represent xenocrysts sourcing from the mantles of early differentiated asteroids (Libourel and Krot, 2007; Chaussidon et al., 2008). An alternative explanation for intra-chondrule D17O heterogeneity is that chondrule melts exchanged O during the crystallization sequence (e.g. Pack et al., 2004; their Fig. 20). In such a scenario, minerals (ol, px) freeze in the D17O of the changing melt at time when they crystallized. The last phase (e.g. mesostasis glass) then is expected to show the highest D17O by protracted exchange with the 16 O-poor H2O from the ambient gas. Melting of chondrules was a transient process. The question is if melting times were long enough to allow O isotope exchange between chondrules and ambient H2O gas. How long chondrule melting lasted is difficult to assess from observation on chondrules. For cooling rates, mineral zoning (e.g. Weinbruch and Mu¨ller, 1995) and crystallization experiments (e.g. Hewins et al., 2005) post-liquidus cooling rates of 10–1000 K h1 were inferred. These cooling rates are, however, associated with a large uncertainty. Nagashima et al., 2006, 2008 conducted containerless crystallization experiments and suggested lower chondrule cooling rates of <10 K h1. The duration at which chondrules remained at super-solidus T can be assessed from the physical model of the heating event. The shock wave model (e.g. Boss, 1996; Jones et al., 2000; Iida et al., 2001; Ciesla and Hood, 2002; Miura et al., 2002; Connolly and Desch, 2004; Miura and Nakamoto, 2005; Connolly et al., 2006; Morris et al., 2009; Morris and Desch, 2010) predicts that chondrules were partially molten for a few hours. The exchange rate between chondrule melts and ambient H2O is not only function of T and the duration of melting, but also of the number density of H2O molecules. At 1 bar and f O2 = IW  0.5, Yu et al. (1995) demonstrated that exchange occurs in minutes. If O isotope exchange is a 1st

order reaction, reduction of total pressure from 1 bar to 103 bar as suggested for a nebula of solar composition (e.g. Wood and Morfill, 1988) would results in exchange only if melting occurred longer than 10 h. This implies that O isotope exchange between molten chondrules and ambient H2O gas could have occurred during shock wave heating of chondrules. We have extended the data set of Yu et al. (1995) by conducting H2O gas–melt exchange experiments at lower f O2. Here, we report new results of times series experiments of the O isotope exchange between two isotopically distinct components, i.e. chondrule-analogue melt and water vapor at f O2 between IW  3.8 and IW  1.3 in a (H2–H2O)0.1– (N2)0.9 atmosphere. Aim of the experiments is to better constrain the kinetics of O isotope exchange in order to obtain information on the extent the O isotope exchange of chondrules with ambient gas. If exchange occurred, chondrules may also provide informations regarding the composition of the nebular H2O. 2. EXPERIMENTAL SETUP AND ANALYTICAL METHODS 2.1. Experimental details 2.1.1. Preparation of starting material The chondrule-analogue starting material used in this study is a synthetic glass in the system CMAS + TiO2 with 0.46 wt.% MnO (Table 1). The starting material was prepared from reagent grade oxides and carbonates. The mixture was ground in an agate mortar, slowly heated at 1100 °C to decompose the calcium carbonate in CaO and CO2 and then fused in a Pt crucible at 1600 °C in air before quenched to glass. The resulting glass was finely ground to powder to ensure homogeneity before used in the oxygen isotope exchange experiments. The starting material has chondritic Al/Ca ratio. The Al/Si and Ca/Si ratios are about four times the CI chondritic ratio (Lodders, 2003). Higher Al/Si and Ca/Si ratios were chosen to bring the liquidus temperature below the limit of the furnace (Tmax 6 1600 °C). The calculated liquidus temperature is 1480 °C using MELTS (Ghiorso and Sack, 1995) and the liquidus phase is olivine. The starting composition resembles that of low-FeO aluminum-rich chondrules (Bischoff and Keil, 1984). The CaO content is similar to

Table 1 Starting composition for all experiments (EMP data). Oxide

Concentration (wt.%)

SiO2 TiO2 Al2O3 FeO MnO MgO CaO P

45.00 0.76 15.70 bdl 0.46 25.10 13.40 100.40

bdl = below detection limit.

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that reported by Hezel and Palme (2010) for a PO chondrule in Kainsaz. 2.1.2. Gas mixing furnace setup The silicate melt-H2O vapor exchange experiments were conducted using the metal loop technique (Presnall and Brenner, 1974) at 1500 °C, different f O2 and durations (Table 2) in a GERO vertical gas-mixing furnace (Fig. 1). The f O2 was controlled by an H2O–H2–N2 gas mixture using the H2O/H2 buffer. A stream of a non-explosive mixture of 10 vol.% H2 and 90 vol.% N2 was split into two streams; the “water line” and the “dilution line” (Fig. 1). Flows trough both lines were controlled using calibrated (for a 10 vol.% H2, 90 vol.% N2 gas mixture) Vo¨gtlin Instruments (Switzerland) gas flow controllers. The gas of the “water line” was saturated with H2O at a flow rate between 18 and 300 mL min1 through two bottles, each containing two liters of deionized tap water and equipped with a diffusion stone. The fraction of H2O vapor in the H2–N2 gas stream is a function of the water temperature. For the first two sets of experiment (higher f O2, Series A and B, see Table 2), the water bottles were held at room temperature (26 ± 3 °C). The corresponding H2O saturation vapor pressure was 30 mbar giving a H2O/H2 volume

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ratio of 0.34 for the “water line”. At 1500 °C this corresponds to a f O2 of IW  0.94 ± 0.16 for Series A where no dry gas from the “dilution line” was added. For Series B, the “water line” (88 mL min1) was diluted with 220 mL min1 dry gas from the “dilution line” giving a f O2 of IW  2.04 ± 0.16. For the lower f O2 experiments (Series C and D, Table 2), we kept the bottles at a constant temperature of 15 ± 0.1 °C inside a LabVIEW-controlled thermo-electric chiller box. The H2O/H2 ratio in the “water line” was 0.17 giving a f O2 of IW  1.54 ± 0.16. Chilling of the water bottles had the positive effect that they were the coolest point of the line ensuring that no condensation of tap water occurred in any other part of the system. The f O2 was adjusted to IW  3.1 ± 0.2 (Series C) and IW  4.0 ± 0.2 (Series D) by appropriate addition of H2O-free H2–N2 gas. The uncertainty of about ±0.16 is dominated by the precision of the flow controllers. The time-zero experiment (sample E5 in Table 2) was run in dry H2–N2 gas, to avoid any possible O isotope exchange and get the isotopic composition of the starting material. The total flow rate through the tube furnace (“water line” + “dilution line”) was maintained at 300 ml min1 in all experiments.

Table 2 List of experiment numbers and experimental conditions. All experiments were run in Fe loops at a temperature of 1500 °C. Sample

Time (min)

f O2

“Water line” (mlmin1)

“Dilution line” (mlmin1)

A-5 A-10 A-20 A-40 A-90 A-180 A-720 A-1440 B-5 B-20 B-40 B-90 B-180 B-360 B-720 C-5 C-10 C-20 C-40 C-90 C-180 C-360 C-720 C-1440 D-5 D-10 D-20 D-40 D-90 D-180 D-360 D-720 E-5

5 10 20 40 90 180 720 1440 5 20 40 90 180 360 720 5 10 20 40 90 180 360 720 1440 5 10 20 40 90 180 360 720 5

IW  1.3 IW  1.3 IW  1.3 IW  1.3 IW  1.3 IW  1.3 IW  1.3 IW  1.3 IW  1.9 IW  1.9 IW  1.9 IW  1.9 IW  1.9 IW  1.9 IW  1.9 IW  3.1 IW  3.1 IW  3.1 IW  3.1 IW  3.1 IW  3.1 IW  3.1 IW  3.1 IW  3.1 IW  3.8 IW  3.8 IW  3.8 IW  3.8 IW  3.8 IW  3.8 IW  3.8 IW  3.8 H2 + N2

300 300 300 300 300 300 300 300 88 88 88 88 88 88 88 50 50 50 50 50 50 50 50 50 18 18 18 18 18 18 18 18 0

0 0 0 0 0 0 0 0 220 220 220 220 220 220 220 250 250 250 250 250 250 250 250 250 280 280 280 280 280 280 280 280 300

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Fig. 1. Sketch illustrating the experimental set up. The H2O/H2 ratio is controlled by bubbling the gas stream of the “water line” through two water reservoirs kept at constant temperature.

2.1.3. Silicate-vapor oxygen exchange experiments The silicate melt-vapor O exchange experiments have been conducted in Fe loops. About 5 mg of powdered starting material, mixed with organic glue, were inserted into metal-loops with an average radius of 1.2 mm, prepared by cutting Fe foil (0.1 mm thick, 99.99% Fe, Alfa Aesar) into strips. All the experiments have been run at the temperature of 1500 °C. The furnace temperature was controlled to ±1.5 °C using a type B (Pt-30%Rh–Pt-6%Rh) thermocouple calibrated against the melting points of Au and Ni. For each experiment, two loops were hung into the furnace and kept for 30–45 min in the upper part of the furnace at the temperature of about 400 °C in an Ar atmosphere (flow rate of 1000 mL min1). The aim was to remove all water and organic glue from the charges and to provide a neutral atmosphere by pushing the air out of the furnace before moving into the hot zone of the furnace. We then flushed the furnace with the H2O–H2–N2 mixture and moved the samples into the hot zone of the furnace. There they were kept for 5 min to 24 h. Five minutes are required to move the charges to the hot zone of the furnace, establish thermal equilibrium between the sample and the furnace atmosphere and get a stable temperature reading. The samples were quenched by melting the Fe “holding wire” by means of an electric current and dropped into a glass cup at the bottom of the furnace (see Fig. 1, inset). One sample each was used for O isotope analysis and the second sample was mounted in resin and polished for scanning electron microscopy (SEM) and electron microprobe (EMP) analyses. 2.1.4. Determination of f O2 from FeO concentrations in the run products The f O2 was checked by equilibration of the Fe loop with the silicate melt. The oxidation of metallic iron is described by: FeðsÞ þ 1=2O2  FeOðlÞ

ð1Þ

Equilibrium between metallic iron and FeO dissolved in the melt is given by:

log10 f O2 ¼ 2 

Df G0T melt þ 2  log10 ðcmelt FeO Þ þ 2  log10 ðX FeO Þ 2:303  RT

ð2Þ

where Df G0 [J/mol] is the Gibbs free energy of formation of FeO(l) for the elements in their reference state and is equal to 244,118 + 115.559T–8.474TlnT (Holzheid et al., 1997 and references therein) and cFeO is the activity coefficient for FeO (=1.70 ± 0.22; Holzheid et al., 1997). The mole fraction X of FeO in the melt was calculated from EMP. 2.2. Analytical methods 2.2.1. Scanning electron microscopy Back scattered electron (BSE) images of the experimental run products were obtained with a LEO 1450VP SEM using a 15 kV accelerating voltage and 1–2 nA beam current. 2.2.2. Electron microprobe Electron microprobe analyses of the experimental glasses were performed with a JEOL 8900R electron microprobe equipped with five wavelength-dispersive spectrometers. An accelerating voltage of 15 kV, a beam current of 10 nA, a defocused electron beam of 2 lm and counting times of 5 s on background and 15 s on peak were used. The following standards (with detection limits) were analyzed for calibration: San Carlos olivine for Si (180 ppm) and Mg (190 ppm), anorthite for Al (160 ppm), hematite for Fe (320 ppm), rutile for Ti (330 ppm), rhodonite for Mn (300 ppm), wollastonite for Ca (210 ppm). The built-in JEOL ZAF correction routines were used to obtain the element concentrations from count rates. 2.2.3. Oxygen (d18O) and hydrogen (dD) isotope analysis of tap water and vapor The d18O and dD values of the tap water and water vapor of our experiments were measured by laser absorption spectrometry (Baer et al., 2002), using a Los Gatos Research water analyzer (IWA-35EP) equipped with a CTC LC-PAL liquid auto sampler for automated injection

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of 1 lL water. Working standards calibrated against VSMOW were analyzed together with the samples. For each sample and standard 6 injections were performed, the first 3 measurements were discarded due to memory effects and the average of the last 3 measurements were used. The analytical uncertainty in d18O and dD is in the range of ±0.3 and ±1.5& (1r), respectively. The vapor was collected at the outlet of the electric chiller using a U-trap immersed in liquid nitrogen (Fig. 1). A collecting time of 2–6 h was necessary to obtain 0.3 mL of water. After collection the water was sealed in glass ampoules before analyzing with the laser spectrometer.

The chosen slope of 0.5305 coincides with the high-T approximation for equilibrium fractionation (Matsuhisa et al., 1978; Young et al., 2002). The choice is somewhat arbitrary and it has to be noted that D17O is not a measured quantity, but was introduced to better display small variations in the relation between d17O and d18O (see e.g. Pack and Herwartz, 2014, and references therein for more details). Accordingly all the reported D17O values have been recalculated with respect to the reference line (RL) chosen here. The analytical uncertainty in d18O and D17O (see Eq. (3) for definition) are ±0.15& and ±0.05& (1r), respectively.

2.2.4. Oxygen isotope analyses (d17O, d18O) of experimental glasses by laser fluorination Oxygen isotope analyses of experimental glasses were carried out by means of infrared laser fluorination (Sharp, 1990) in combination with gas chromatography isotope ratio monitoring gas mass spectrometry (CF-GC-IRMS; for details, see Pack et al., 2013). About 0.3 mg glass was loaded along with NBS-28 quartz, into an 18-hole Ni metal sample holder. After evacuation and heating of the sample chamber to 70 °C overnight, materials were reacted in a 20–30 mbar of purified F2 gas (Asprey, 1976) by means of heating with a SYNRAD 50 W CO2-laser. Excess F2 was removed from the run products by reacting with heated NaCl (180 °C) and cryotrapping of Cl2. The released O2 was first cryofocused in a 6 mm stainless steel molecular sieve trap at 196 °C (liquid N2). The sample O2 was then released by heating to 120 °C and transported with He carrier gas through a second molecular sieve trap, where a fraction of the sample gas was again cryofocused at 196 °C. The sample O2 was released at 92 ± 2 °C (tempered hot water bath) back into the He carrier gas stream and transported through ˚ molecular sieve capillary gas chromatograph column. a5A The column was operated at room temperature and separated NF3 from O2, which is critical when analyzing d17O (e.g. Pack et al., 2007, and references therein). Purified sample O2 was then injected via an open split valve of the GasBench-II into the source of a THERMO MAT 253 gas mass spectrometer. The signals of 16O–16O, 17O–16O, and 18 O–16O were simultaneously monitored on 3 Faraday cups. Bell-shaped sample peaks had an amplitude of 25–30 V (m/z = 32, where m/z represents the mass divided by charge number) and a full width at half maximum of 20 s. Reference O2 was injected for 40 s two times before the sample through a second open split valve of the GasBench-II. The amplitudes (m/z = 32) of the reference gas peaks were 25 V. Variations in triple oxygen isotope ratios (17O/16O, 18 O/16O) are expressed in the d notation relative to VSMOW (McKinney et al., 1950) and the D17O (Eq. (3)). For mass dependent processes, variations in d17O are about half of variations in d18O. Deviations from such a mass dependent fractionation line are expressed in form of the D17O notation with:  17  d OVSMOW D17 O ¼ 1000  ln ¼ þ1 1000  18  d OVSMOW ð3Þ  0:5305  1000  ln þ1 1000

3. RESULTS 3.1. Texture of the run products The BSE images of the experimental run-products show that all of them were glassy Rarely, fine-grained quenched skeletal olivine crystals were observed (Fig. EA-1a and b in the Electronic Annex). 3.2. Chemical composition Representative chemical compositions of the experimental run products are given in Table 3 (the complete set of EMP data is available as Electronic Annex, therein Table EA-1). It is clear from Fig. 2 that the compositions of the charges become enriched in FeO from oxidation of the Fe loops by H2O from the gas mixture. Some experimental run shows a decrease in their SiO2 + MgO values (Fig. 2, Series C and D). This is more evident from Table EA1 where the MgO content (calculated on FeO free basis) slightly decrease from 25 to 23 wt% for the Series D and from 25 to 24 wt% for the Series C. This could be due to olivine crystallization at some part of the loop (Fig. EA1). Significant evaporative loss was observed for Mn (Fig. 3) of which 50% was lost after 75 min (series D) to 300 min (series A) suggesting that the evaporative loss is greater at lower f O2. The MnO evaporation rates of our charges are similar to those observed by the experiments of Cohen and Hewins (2004). 3.3. Partitioning of Fe between metal and silicate The FeO content in the silicate melt increased with increasing f O2 from initially zero up to 20 wt.% (Table 3). The oxygen fugacity of each experimental series was calculated from the iron–wu¨stite buffer curve using the FeO content of the equilibrated run products. The four experimental series give equilibrium XFeO of 0.010 ± 0.001, 0.019 ± 0.001, 0.079 ± 0.003, and 0.156 ± 0.006 (1r) after about 20 min. We used these values for the last term of Eq. (2) to get the log10f O2. These correspond to oxygen fugacities of IW  3.8, IW  3.1, IW  1.9, and IW  1.3 (Fig. 4). An uncertainty of about ±0.1 in the calculated f O2 value is mainly related to the uncertainty of the activity coefficient of FeO in the melt (Holzheid et al., 1997). The oxygen fugacity that is obtained from the FeO

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Table 3 Representative chemical composition of the experimental glasses (EMP data). The oxides Na2O and K2O were <0.05 wt.% in all samples. Experiment Time (min)

A-5 5

A-1440 1440

B-5 5

B-720 720

C-20 20

C-1440 1440

D-20 20

D-720 720

SiO2 (wt.%) TiO2 Al2O3 FeO MnO MgO CaO R

40.03 0.67 13.72 11.10 0.41 22.14 11.46 99.53

37.69 0.62 12.62 19.85 bdl 18.53 10.41 99.72

42.43 0.70 14.58 5.88 0.45 23.47 12.29 99.80

40.61 0.71 15.64 7.79 0.02 22.30 13.01 100.07

45.06 0.72 15.46 1.86 0.38 23.87 12.54 99.89

43.81 0.82 16.11 2.02 0.09 23.85 13.30 99.99

43.40 0.75 16.12 1.21 0.31 24.81 13.47 100.07

42.04 0.91 18.45 1.16 bdl 22.34 15.16 100.06

bdl = below detection limit.

Fig. 2. Plot showing the main element chemical composition of the run products (time series A: points, B: squares, C: diamonds, D: triangles; EMP data). The experimental series A reached about 20 wt.% FeO (Table 3).

concentrations in the equilibrated run products (Fig. 4) is in excellent agreement with the oxygen fugacity as defined by the H2O/H2 ratio (Fig. EA-2 in the Electronic Annex).

Although the comparison between the f O2 derived from FeO concentrations and that predicted from the H2O/H2 ratio support that the gas was saturated in vapor, we checked for saturation by analyzing the d18O and dD in the tap water and the vapor. Especially at high flow rates through the bottles, incomplete saturation could occur, which would result in erroneously low H2O/H2 ratios and hence too low f O2. The isotopic composition of the tap water is d18O = 8.5 ± 0.05& and dD = 58.6 ± 0.2& (1rm, Table 4). The composition of the tap water did not change beyond the analytical uncertainty during the course of the experiments (Fig. EA-3 in the Electronic Annex). This is due to the small amount of water removed as vapor during the experiments. The water is local tap water and falls on the meteoric water line (Hofmann et al., 2012) that is defined by:  17   18  d OVSMOW d OVSMOW ln þ 1 ¼ 0:528  ln þ 1 þ 0:033 1000 1000 ð4Þ in the d17O vs. d18O space (Barkan and Luz, 2005; Luz and Barkan, 2010). This gives a d17O and D17O for the tap water

Fig. 3. Plot of MnO (calculated on FeO free basis) vs. time for experimental series A (solid circles) and D (triangles). The curves are exponential fit functions of type y = a * (ebx) + c. The experiments of Cohen and Hewins (2004) (diamonds) are plot for comparison.

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Fig. 4. Plot of XFeO of the experimental melts vs. heating duration in minutes. Time series A: points, B: squares, C: diamonds, D: triangles. The dashed lines represent XFeO values at infinite times calculated using an exponential fit function of type y = a * (1  ebx). Error bars are not shown, as they are smaller than the symbols.

of 4.5 and 0.054&, respectively. Vapor was collected at flow rates of 300, 88 and 50 ml/min and water bath temperature of 15 °C for the four different experimental series A, B, and C. For the IW  4 experiment (series D), we could not collect vapor in a reasonable time because of the low H2O/H2 ratio in this series. The mean d18O of the vapor was 18.7 ± 1.7& and the dD was 144 ± 13& giving D18Ovapor-water = 10 ± 1.7& and DDvapor-water = 86 ± 13&. Within uncertainty, the measured vapor-water fractionation is in agreement with the equilibrium fractionation in both, dD and d18O (Horita and Wesolowski, 1994; Fig. EA-4 in the Electronic Annex)

and further confirms the conclusion from the chemical analysis that the “water line” was saturated in vapor. The d17O of the vapor is calculated from the triple isotope equilibrium fractionation data by Meijer and Li (1998) and Luz and Barkan (2010) to 9.9& giving a D17O = 0.069& (±0.02). The measurements (FeO, oxygen [d18O] and hydrogen isotopes [dD]) show that our experimental setup is suitable for controlling f O2 in a range between IW  1 and IW  4 using the H2O/H2 buffer with an accuracy of <±0.16 log units. 3.4. Oxygen isotope composition (d17O, d18O) of the silicates

Table 4 Hydrogen (dD) and oxygen isotope (d18O) composition of water and vapor. The water was collected during the course of the experiments. The analytical uncertainties in d18O and dD are ±0.3& and ±1.5&, respectively (1r, SD). Sample

dD (&)

d18O (&)

Water #1 Water #2 Water #3 Water #4 Water #5 Water #6 Water #7 Water #8 Water #9 Water #10 Water #11 Water #12 Water #13 Water #14 Water #15 Vapor @ 300 ml/min Vapor @ 88 ml/min Vapor @ 50 ml/min

58.3 58.4 59.3 58.3 57.9 57.5 56.9 58.6 58.5 58.0 59.9 57.8 58.7 60.5 59.9 169.6 138.3 125.2

8.5 8.5 8.4 8.5 8.5 8.7 8.2 8.5 8.6 8.4 8.6 8.3 8.2 8.8 8.7 22.3 16.7 17.2

The oxygen isotope composition of the starting material (sample E5 in Table 5) was d18O = 11.3& with a D17O of 0.28& (Table 5). The d18O and the D17O of the silicate melt systematically change during the experiments and approach the composition of the vapor (Figs. 5 and 6). The oxygen isotope composition of the run products follows a mixing trend between the starting material and the vapor in a D17O vs. d18O diagram (Fig. 5). 4. DISCUSSION 4.1. Kinetic analysis of the exchange process and applications to the origin of chondrules Our experiments show that, depending on f O2, isotope exchange between vapor and the silicate melt in the 2 mm metal loops occurs on a timescale of minutes to hours (Table 5 and Figs. 5 and 6). Higher f O2 results in higher exchange rates. This behavior suggests that the rate-limiting mechanism of O isotope exchange is the reaction occurring at gas/melt interface, which is dependent on the partial pressure of water vapor (Yu et al., 1995).

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The isotope exchange reaction is written as: 16

18

16

16

18

Si O O þ H2 O  Si O2 þ H2 O

ð5Þ

Here, the silicate melt is represented by “SiO2”. The O isotope exchange rate constants k for the reaction (Eq. (5)) were calculated by linear regression from Ln(1  f ) versus time plots (Fig. 7a) with the assumption that the exchange between vapor and silicate melt is a first order reaction. The variable f represents the fractional approach of the system to equilibrium (0 6 f 6 1) and has the following form (Cole and Chakraborty, 2001 and references therein):   d18 Ot  d18 Ot¼0 ¼ 1  eðktÞ f ¼ ð6Þ d18 Oequilibrium  d18 Ot¼0 Rearranging and taking the logarithm of both sides yields:  18  d Ot  d18 Oequilibrium ¼ kt ð7Þ Ln 18 d Ot0  d18 Oequilibrium The d18Ot = 0 value is the starting composition which we took as the composition of the 20 min experiments instead of the composition of the starting material. This choice is

Table 5 Triple oxygen isotope composition of the experimental products. The analytical uncertainties in d18O and D17O are ±0.3& and ±0.05&, respectively (1r, SD). Sample

Time (min)

d17O (&)

d18O (&)

D17O (&)

A-5 A-10 A-20 A-40 A-90 A-180 A-720 A-1440 B-5 B-20 B-40 B-90 B-180 B-360 B-720 C-5 C-10 C-20 C-40 C-90 C-180 C-360 C-720 C-1440 D-5 D-10 D-20 D-40 D-90 D-180 D-360 D-720 E-5

5 10 20 40 90 180 720 1440 5 20 40 90 180 360 720 5 10 20 40 90 180 360 720 1440 5 10 20 40 90 180 360 720 5

6.4 6.1 6.5 7.5 9.1 9.6 9.9 10.2 5.3 6.9 7.5 8.2 8.6 9.2 8.9 5.3 5.1 3.8 1.1 1.8 4.9 7.5 9.2 8.9 5.2 5.5 5.2 4.2 2.1 1.1 6.3 8.0 5.7

11.9 11.4 12.2 14.1 17.1 18.0 18.7 19.1 9.7 12.9 14.0 15.3 16.2 17.4 18.5 10.3 10.0 7.4 2.3 3.1 8.9 14.1 17.3 16.7 10.3 10.9 10.2 8.3 4.3 1.8 11.6 14.9 11.3

0.02 0.01 0.00 0.00 0.01 0.05 0.07 0.05 0.10 0.04 0.04 0.01 0.04 0.05 0.03 0.19 0.17 0.19 0.16 0.19 0.12 0.01 0.02 0.02 0.22 0.25 0.21 0.21 0.22 0.17 0.10 0.08 0.28

based on the observation that the melts becomes saturated in FeO only after 20 min (Fig. 4). The oxidation of Fe from the loop leads to a net transfer of O into the melt (accompanied by a rapid shift in d18O) that is more rapid than the exchange with the gas. The d18Oequilibrium value in Eq. (7) refers to the compositions of the silicate melt that is in equilibrium with steam in the furnace atmosphere. Since our data do not reveal any isotope fractionation at equilibrium between the silicate melt and vapor at 1500 °C (i.e. D18Osilicate–vapor = 0 ± 0.15&; Fig. 6a), given a sufficient amount of time, the oxygen isotope composition of the fully equilibrated experimental charges will coincide with the composition of the surrounding gas phase. Accordingly, the equilibrium d18O value (d18Oequilibrium) for the silicate melt is equal to that of the vapor (18.7&). The d18Ot in Eq. (7) refers to the composition of the silicate melt after time t. The d18Ot = 0 represents the composition of the starting material (i.e. d18Ot = 0 = 11.3&). Fig. 7a shows our data in a Ln(1  f ) vs. t diagram. The slopes give the respective reaction rate constants k (Eq. (7)). The results of the experiments by Yu et al. (1995) that were conducted at 1450 °C and IW  0.5 are also displayed in Fig. 7a and compared with the data from this study. The data show that k decreases with decreasing f O2, i.e. H2O partial pressures. For comparison between the different experiments, we introduce the exchange half-lives (t1/2). These are the times that are necessary to reach 50% exchange, i.e. 1  f in Eq. (7) equals 0.5. The t1/2 values vary from 70 min for the IW  1.3 (A series) to 230 min for the IW  3.8 (D series). The experiments by Yu et al. (1995) were conducted at IW  0.5 and yield a t1/2  17 min. The exchange reaction (Eq. (5)) suggests that t1/2 inversely correlates with the H2O partial pressure. Since collisions are necessary for an exchange, we display t1/2 vs. H2O molecule density in the ambient gas phase (Fig. 7b). This allows extrapolation toward lower pressures and lower f O2. As expected for a first order reaction, t1/2 negatively correlates with the number density of H2O molecules in the gas. Extrapolated to a canonical solar nebula pressure (solar gas, 106 < Ptot < 103 bar; Wood and Morfill, 1988), 50% exchange takes between 610 (Ptot = 103 bar) and 940 (Ptot = 106 bar) minutes (10–16 h) (Fig. 7b), with the assumption of a linear relationship between t1/2 and nH2O. The FeO contents of both, type-I (<10 wt.%; Jones et al., 2005) and type-II chondrules (>10 wt.%; Jones et al., 2005) are much higher than what is buffered by a nebula of solar composition (f O2 = IW  6.8; Grossman et al., 2008). Higher f O2 (required to stabilize FeO in chondrule silicates) can be achieved by local enrichment of a nebula in dust relative to gas followed by high temperature interaction between silicate dust and ambient gas (Ebel and Grossman, 2000). A 50 dust enrichment has been suggested for type-I and a 800 dust enrichment for type-II chondrules (Ebel and Grossman, 2000). The corresponding number density of H2O molecules in dust-enriched systems is 5.6  1019 and 9  1020 m3, respectively (Ptot = 103 bar). This

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Fig. 5. Oxygen isotope composition of starting material, vapor, and experimental run products. The silicate melts change isotope composition and approach the composition of the vapor. A mixing line with weight fractions is displayed. Error bars for d18O have the size of the symbols.

corresponds to a Log(f O2) = IW  3.8 for type-I and Log(f O2) = IW  1.1 for type-II chondrules, respectively. For dust-enriched systems (Ptot = 103 bar), 50% exchange takes between 300 (800 dust enrichment, type-II chondrules) and 430 min (50 dust enrichment, type-I chondrules) (Fig. 7b). The time spans necessary to obtain 50% exchange between a molten chondrule and ambient gas at different conditions are listed in Table 6. Tenner et al. (2015) suggest, that for CR3.0 chondrites type-I and type-II chondrule formed at dust enrichments of about 100–200 and 2500, respectively. At these conditions 50% exchange takes between 370 min for type-I chondrules and 250 min for type II chondrules (Fig. 7b). Chondrules were melted in a transient heating event. Heating of chondrules in nebular shock waves fulfills a number of boundary conditions necessary to explain the properties of chondrules (Boss and Durisen, 2005; Morris and Desch, 2010). The shock wave model predicts that chondrules were heated to super liquidus temperatures for 67 h and were partially molten for up to some tens of hours (Fig. 8, the liquidus and solidus temperatures of chondrules were adopted from Hewins and Connolly, 1996). This implies that for type-I chondrules, that have formed by melting in a gas with 50 dust enrichment (Ptot = 103 bar) about 50% of the chondrules’ O would have exchanged with the ambient H2O. For type-II chondrules (800 dust enrichment, Ptot = 103 bar), given an exchange half-life of 300 min (Fig. 7b), a melting time of 7 h would result in 70% oxygen exchange between chondrules and ambient gaseous H2O. This confirms that isotopically heterogeneous chondrules may be the result of incomplete isotope exchange with ambient H2O gas (Jones et al., 2004) and demonstrates that both, type-I and type-II chondrules are expected to inherit (more for type-II, less for type-I) information about the oxygen isotope composition of the surrounding gas (e.g. Tenner et al., 2012; Schrader et al., 2013, 2014a).

The relationship between ambient redox conditions and degree of O exchange highlighted by our experiments and depicted in Fig. 7b may explain the O isotope variability in the composition of olivines from type-I and type-II chondrules observed in carbonaceous chondrite chondrules (Krot et al., 2006b; Connolly and Huss, 2010; Libourel and Chaussidon, 2011; Rudraswami et al., 2011; Ushikubo et al., 2012; Schrader et al., 2013, 2014a; Tenner et al., 2013, 2015). Fig. EA-5 shows that the D17O values of olivines in type-II chondrules occurring in CV, CO, Acfer 094 and CR chondrite (Fig. EA-5a and EA-5b in the Electronic Annex) is higher that the values of olivines in type-I. If chondrules exchange with 16O-poor H2O that formed in the self shielding process (Clayton, 2002), higher D17O is expected for type-II chondrules than for type-I chondrules. This is, indeed, observed. Our experiments show that in dust enriched systems, as suggested for type-I and type-II chondrules, approximately 50% of the O in chondrules exchanged with 16O-poor H2O in the ambient gas. Such a scenario does not contradict the formation of all chondrules (and likely also matrix), in the same nebula region (Hezel and Palme, 2008, 2010). Higher D17O values of chondrules can be explained by higher exchange rates in a gas with higher H2O molecular number density. 4.2. Oxygen isotope composition of the gaseous reservoir Our experimental data show that considerable exchange between chondrule melts and the nebular gas is expected. By using mass balance we calculate the O isotopic composition of the nebular gas in the formation region of type-I and type-II chondrules, respectively. For our calculations we apply the degrees of exchange of 50% and 70%, constrained by our experiments and use the average O isotope data of olivines from relict-free type-I and type-II chondrules studied by secondary ion mass spectrometry (final compositions in Table 7; data from Krot et al., 2006b; Connolly and Huss,

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Fig. 6. Plot of d18O (a) and D17O (b) values of the quenched glasses vs. time. The dashed lines represent the d18O and D17O values of the water vapor (gray: 1r uncertainty envelope). Solid lines are exponential regressions passing through the oxygen isotope composition of the starting material. Error bars in (a) have the size of the symbols.

2010; Libourel and Chaussidon, 2011; Rudraswami et al., 2011; Ushikubo et al., 2012; Schrader et al., 2013; Tenner et al., 2013). The O isotope composition of chondrule precursor is poorly constrained. We approximate the composition of chondrule precursors from the following considerations: (i) the composition of refractory condensates demonstrates that the gas reservoir in the solar nebula was originally 16O-rich. Ale´on et al. (2002) and Makide et al. (2009) presents O isotope analyses of calcium– aluminum-rich inclusions (CAIs) and amoeboid olivine aggregates (AOAs) from CR chondrites and show that primary minerals are 16O rich (26& < D17O < 20&), while objects presenting textural evidences of melting are 16O depleted (D17O grater than 18&). (ii) Olivines of both type-I and type-II chondrules, are 16O-depleted compared to AOA and most CAI, demonstrating that the gas reservoir in the solar nebula evolved from being 16O-rich to 16O-poor. (iii) Relict olivines occur in both type-I and –II chondrules

and can be distinguished on the basis of texture, Mg# and O isotope compositions (e.g. Connolly and Huss, 2010; Tenner et al., 2013; Schrader et al., 2013 and references therein). Rudraswami et al. (2011) and Ushikubo et al. (2012) show that type-I and type-II chondrules from Acfer 094 and CV chondrite contain relict olivine grains that are either 16O-enriched or 16O-depleted compared to the other minerals of the same chondrule. Here, for the sake of simplicity we consider only those relicts having lower D17O and take the average values of relict grains from both type-I and type-II chondrules reported in CV chondrites (Rudraswami et al., 2011), CO chondrites (Kunihiro et al., 2004; Tenner et al., 2013), Acfer 094 (Ushikubo et al., 2012) and CR chondrites (Krot et al., 2006b; Chaussidon et al., 2008; Schrader et al., 2013). Accordingly, type-I chonand drules relict grains have d18O = 11.2& 17 d O = 14.3&. Relicts of type-II chondrules have d18O = 1.5& and d17O = 5.1& (Table 7).

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Fig. 7. (a) Fraction of oxygen isotopic exchange given as Ln(1  f) vs. time in minutes. Lines are least square fits to the data. Rate constant (k) values are given by the slopes of the lines. (b) Plot of half-lives of exchange reaction vs. number density of water in the gas phase (nH2O in logarithmic scale) for the Series A, B, C, D and the experiments of Yu et al. (1995) performed at IW  0.5 (all plotted as points). The thick solid line is the least square regression fit to the data. The dashed line represents the extrapolation to conditions relevant to solar nebula (solar gas, Ptot = 103 and 106 bars) and chondrule forming events (50 and 800 dust enrichment, Ebel and Grossman, 2000 and this study; 200 and 2500 dust enrichment, Tenner et al., 2015). The gray shaded region (from 0 to 420 min) indicates the time chondrules are expected to remain above their liquidus temperature in a shock wave model (see text and Fig. 8 for further explanations).

Table 6 H2O/H2 ratios, f O2 and number density of water (nH2O) at conditions relevant to solar nebula (Lodders, 2003) and chondrules formation (dust-enriched systems; Ebel and Grossman, 2000). The half-lives are calculated from the dashed line of Fig. 12, representing the extrapolation of the experimental results in a t1/2 vs. nH2O plot. Solar Nebula Ptot = 106 bar H2O/H2 Log10f O2-IW nH2O (m3) t1/2 (min)

4

4.19  10 6.8 1.50  1019 954

Dust-enrichment Ptot = 103 bar 4

4.19  10 6.8 1.50  1018 620

50, Ptot = 103 bar 2

1.60  10 3.8 5.63  1019 434

800, Ptot = 103 bar 3.36  101 1.1 8.98  1020 304

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Fig. 8. Thermal history of chondrules predicted by shock wave model (Morris and Desch, 2010) (solid curve). The chondrule liquidus (solid lines) and solidus (dashed line) temperatures are from Hewins and Connolly (1996).

Table 7 Calculated triple oxygen isotope composition of the gas reservoir. d18O (&)

d17O (&)

D17O (&)

40.7 11.2 1.5

43.5 14.3 5.1

22.5 8.4 4.3

Final compositions Type-I olivines (CO, CV)(c) Type I olivines (CR)(d) Type-II olivines (CO; CV) (e) Type-II olivines (CR) (f)

5.4 1.0 1.4 6.7

8.6 1.9 1.6 2.6

5.7 2.5 2.4 1.0

Gas compositions Gas phase (50% exchange) Gas phase (50% exchange) Gas phase (50% exchange) Gas phase (50% exchange) Gas phase (70% exchange) Gas phase (70% exchange) Gas phase (70% exchange) Gas phase (70% exchange)

29.8 42.7 0.3 13.2 19.4 27.1 2.6 10.3

26.3 39.7 3.0 10.4 16.3 22.3 0.2 5.8

10.4 16.7 3.1 3.4 6.0 7.9 1.6 0.4

Precursor’s compositions Average CAI + AOA(a) Relict type-I olivines(b) Relict type-II olivines(b)

(a)

Data from Ale´on et al. (2002) and Makide et al. (2009). Data from Krot et al. (2006a,b), Chaussidon et al. (2008), Rudraswami et al. (2011), Ushikubo et al. (2012), Tenner et al. (2013) and Schrader et al. (2013). (c) From Chaussidon et al. (2008), Libourel and Chaussidon (2011), Rudraswami et al. (2011), Ushikubo et al. (2012) and Tenner et al. (2013). (d) From Krot et al. (2006a,b), Chaussidon et al. (2008), Libourel and Chaussidon (2011), Schrader et al. (2013, 2014a) and Tenner et al. (2015). (e) Data from Rudraswami et al. (2011), Ushikubo et al. (2012) and Tenner et al. (2013). (f) From Krot et al. (2006a,b), Connolly and Huss (2010), Schrader et al. (2013) and Schrader et al. (2014a). (b)

We consider that the chondrules precursor must have had an isotopic composition lying somewhere between the composition of refractory condensates and that of relict olivines (precursor’s compositions in Table 7), as relict grains can be considered remnants of the precursor that escaped efficient partial melting and exchange (e.g. Jones et al., 2005).

On the basis of these assumptions and applying an exchange of 50% as calculated above for type-I chondrules, the d18O and D17O values of the gas phase that has exchanged with type-I chondrules range from 0 to +43& and from 3 to +17&, respectively. Applying an exchange of 70%, the d18O and D17O of the gas reservoir that exchanged with type-II chondrules varies instead from

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Fig. 9. Oxygen three isotopes plot showing the compositions of the nebular gas reservoir that have exchanged with (a) type-I, (b) type-II chondrule melt. The composition of the gas phase was calculated by mass balance (data from Table 7), applying an exchange of 50% for type-I chondrules, and 70% for type-II chondrules. (c) Two distinct oxygen isotope trends (i.e. mass dependent and mixing lines) for FUN CAIs indicating evaporation during crystallization of 16O-rich melt, and remelting (possibly associated with chondrule formation event) accompanied by oxygen isotope exchange with a 16O poor gas phase, respectively (data from Wasserburg et al., 1977; Lee et al., 1980; Clayton et al., 1984; Davis et al., 2000; Thrane et al., 2008) (d) Enlargement of the enclosed area showing the overlap between the fields describing the composition of the nebular gas inferred from type-I and -II chondrules and the FUN CAIs mixing lines. The close match in both d18O and D17O between FUN CAIs, BO chondrules and terrestrial planets suggests equilibration with a common nebular reservoir. Bulk Earth (mantle) data are from Pack and Herwartz (2014), Mars and Moon meteorites from Clayton and Mayeda (1996). The composition of BO chondrules in CV and CR chondrites are from Clayton et al. (1983) and Schrader et al. (2014a). The reference line (RL) of slope 0.5305 (see text for further explanations), the CCAM line (Clayton et al., 1977) and the Y + R line (Young and Russell, 1998) are plotted for reference.

+3 to +27& and from 2 to +8&, respectively (Fig. 9a and b; Table 7). Clayton and Mayeda (1984) estimated the O isotope composition of the gas reservoir to vary from +25 to +30& for d18O and from +6 to +8.5& for D17O. Schrader et al. (2013) consider the degree of exchange between the gas phase and type-II chondrule liquid from CR chondrites to vary between 25 and 50%. This gives a gas compositions ranging from +13 to +27& for d18O and from +3 to +9&, for D17O. Type-I and type-II barred olivine (BO) chondrules from CR chondrites are suggested to have exchanged nearly completely with a gas reservoir having 3& < D17O < +3& (Schrader et al., 2014a). Differently from what pointed out by Schrader

et al. (2013) and more in agreement with Schrader et al. (2014a), our experiments demonstrate that, an exchange of about 70% can occur for type-II chondrules during the chondrule forming event. We believe that the major uncertainty in the calculation of the gas reservoir is related to the uncertainty in the composition of the precursor, as chondrule precursors are not unequivocally identified yet (Hezel et al., 2010). To overcome this issue, the composition of the gas phase may be further constrained using a class of CAIs named FUN CAIs in reference to their fractionation and unidentified nuclear effects (Lee et al., 1980; Davis et al., 2000; Krot et al., 2006b, 2008; Thrane et al., 2008). In addition to large O mass dependent isotope fractionation related to

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evaporation, some FUN CAIs provide evidences of re-melting processes and O isotope exchange with a 16O poor gas phase. Melilite, anorthite and fassaite can be 16 O-depleted with respect to spinel and forsterite and plot off the mass fractionation line, but on a mixing trend (Fig. 9c). Thrane et al. (2008) suggest that partial melting and O exchange of a FUN CAI from a CV carbonaceous chondrite was associated to the formation of CV chondrules. If this was the case, we can speculate that FUN CAIs and chondrules (at least from CV chondrites) exchanged with the same gas reservoir. Extrapolation of the mixing lines intercept and overlap the fields describing the composition of the nebular gas inferred from type-I and -II chondrules at their lower ends (Fig. 9c and d), close to the composition of barred olivine (BO) chondrules from CV and CR chondrites (Clayton et al., 1983; Schrader et al., 2014a). BO chondrules are ideal for this purpose as their texture indicates that they have experienced a higher degree of melting than porphyritic chondrules (e.g. McSween, 1985; Hewins et al., 2005; Jones, 2012) and complete equilibration with the nebular gas (Clayton et al., 1983; Chaussidon et al., 2008; Rudraswami et al., 2011; Weisberg et al., 2011). Schrader et al. (2014a) suggest that type-I and -II BO chondrule melts have experienced 90% of exchanged with the surrounding gas phase, approaching the composition of the ambient gas. Accordingly, the gas phase composition obtained from BO chondrule and FUN CAIs data define a smaller range of values with d18O, d17O and D17O that vary between 1 and 6&, 3 and 3& and between 3 and 0&, respectively. As pointed out by Schrader et al. (2014a) such range of values is consistent with the water composition for the CR chondrite parent body (D17O = 0&; Tyra et al., 2011; Keller et al., 2012; Schrader et al., 2014b; or D17O = 0.9&, Clayton and Mayeda, 1999; Schrader et al., 2011) and incompatible with heavy water–ice reservoirs (D17O > 80&) inferred from Fe–S-bearing oxides (Sakamoto et al., 2007). Yurimoto and Kuramoto (2004), Kuramoto and Yurimoto (2005) and Krot et al. (2005) suggest that because of low temperature the outer portions of the protoplanetary disk keep the composition of the molecular cloud components (i.e. water ice is enriched in 17O and 18O as a result of UV photodissociation of CO). In the inner solar disk, instead, H2O–ice evaporates and the O isotope composition of the nebula gas shifts toward a 16O poor composition approaching planetary O isotope composition. This is confirmed by Fig. 9d showing that the composition of Mars, Moon and bulk Earth (Clayton and Mayeda, 1996; Wiechert et al., 2001; Herwartz et al., 2014), falls in the field described by the nebular gas calculated from FUN CAIs and BO chondrules. The observed close match in both d18O and D17O implies that the O isotope composition of BO chondrules, FUN CAIs, and terrestrial planets reflect the composition of the ambient gas they equilibrated with and suggests the existence of a common reservoir plausibly expected to be dominant even during the chondrule forming event.

4.3. Evaporative loss of MnO Type-I chondrules are generally characterized by lower concentrations of volatile elements (Mn, Na, K, S, and P) relative to type-II (Jones and Scott, 1989; Jones, 1990, 1994; Lauretta et al., 2006; Berlin et al., 2011; Schrader et al., 2013). This behavior could be caused by differences in the composition of the chondrule precursor material or by partial volatile loss as result of higher liquidus temperatures (Jones, 1990, 1994) and lower f O2 (Jones and Scott, 1989) of type-I chondrules. Our experiments show that MnO is readily lost from the charges (Table 3 and Table EA1). The loss rate increases with decreasing f O2 (Fig. 3). This is consistent with the partial volatile loss in reduced type-I chondrules with respect to more oxidized type-II chondrules and can be understood in terms of an evaporation reaction: melt MnOsilicate þ H2 ! Mng þ H2 O liquid

ð8Þ

Our experiments imply that in 7 h more than 90% of Mn should evaporate from molten type-I chondrules and more than 75% from molten type II chondrules, respectively. However, it has been shown for element similarly or more volatile than Mn that complete evaporation during chondrule melting did not occur (Humayun and Clayton, 1995, for K; Alexander et al., 2008, and Borisov et al., 2008, for Na). The presence of Mn in chondrules (Lauretta et al., 2006) hence indicates that evaporation was hampered, possibly by the same mechanism that prevented K and Na from evaporating (e.g. chondrule formation in high density region for Alexander et al., 2008; extremely short melting events for Borisov et al., 2008). Alternatively, on the basis on negative Fe/Mn vs. Fe/Mg trend described by olivines from type-II chondrules, Berlin et al. (2011) and Schrader et al. (2013) suggest that Mn could have recondensed from the surrounding gas phase during chondrule crystallization (Libourel et al., 2006; Nagahara et al., 2008). 5. CONCLUSIONS Experimental results from this study demonstrate that: (1) A non-toxic mixture of H2O–H2–N2 is a suitable alternative for toxic CO–CO2 in high temperature-low f O2 gas mixing furnace experiments. (2) The degree of O isotope exchange between a silicate melt and the H2O–H2–N2 gas is proportional to the f O2. Between IW  1.3 and IW  3.8 the exchange half-life varies from 70 to 230 min. (3) Extrapolation to conditions relevant to chondrule formation reveal that 50% exchange takes between 5 h (800 dust enrichment, corresponding to type-II chondrules) to 7 h (50 dust enrichment, corresponding type-I chondrules). (4) The shock wave model predicts that chondrules were heated to super liquidus temperatures for 67 h

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(Hewins and Connolly, 1996; Morris and Desch, 2010). This implies that type-I and type-II chondrules inherit about 50% and 70%, respectively, of their composition form the ambient gas they exchanged with and are proxies for the composition of the nebular reservoir. Protracted exchange of the partial melt with the gas during cooling can explain the observed disequilibrium in D17O between chondrule phases (e.g. olivine and glass). (5) If the hypothesis that re-melting of FUN CAIs is associated with chondrule formation and that they both exchanged oxygen isotopes with the same gas, a nebular composition similar to that of the large bodies Earth, Moon, Mars, is suggested.

ACKNOWLEDGMENTS We thank Reinhold Przybilla and Nina Albrecht for assistance in isotope measurements; Andreas Kronz for help in EMP analyses; Roman Mathieu for the preparation of the starting material; Axel Dierschke and Lothar Laake for technical assistance in the laboratory; Zach Sharp for reading an earlier version of the manuscript and for helpful comments. We thank Dominik Hezel, Devin Schrader and an anonymous reviewer for constructive reviews, Associate Editor Sarah S. Russell for editorial handling and useful comments. This work was supported by DFG-SPP 1385.

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