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Does planetary differentiation really fractionate iron isotopes?
Earth and Planetary Science Letters 256 (2007) 484 – 492 www.elsevier.com/locate/epsl
Discussion
Does planetary differentiation really fractionate iron isotopes? ☆ Franck Poitrasson Laboratoire d'étude des Mécanismes de Transfert en Géologie, Centre National de la Recherche Scientifique-UPS-IRD, 14–16, avenue Edouard Belin, 31400 Toulouse, France Received 10 June 2006; received in revised form 31 October 2006; accepted 31 January 2007 Available online 11 February 2007 Editor: R.W. Carlson
Abstract The difference in the mean Fe isotope composition of samples from the Earth, Moon, Mars and Vesta has been recently interpreted as tracking contrasted planetary accretion mechanisms [F. Poitrasson, A.N. Halliday, D.C. Lee, S. Levasseur, N. Teutsch, Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms, Earth Planet. Sci. Lett. 223 (2004) 253–266]. Using newly produced Fe isotopic data on terrestrial and lunar samples, pallasites, eucrites and Martian meteorites, Weyer et al. [S. Weyer, A.D. Anbar, G.P. Brey, C. Munker, K. Mezger, A.B. Woodland, Iron isotope fractionation during planetary differentiation, Earth Planet. Sci. Lett. 240 (2005) 251–264] reinterpreted these data as fingerprinting planetary differentiation. In particular, these authors suggested that partial melting in the terrestrial and lunar mantles produced melts isotopically heavy. It is shown here that the inference of Weyer et al. [S. Weyer, A.D. Anbar, G.P. Brey, C. Munker, K. Mezger, A.B. Woodland, Iron isotope fractionation during planetary differentiation, Earth Planet. Sci. Lett. 240 (2005) 251–264] is strongly biased by the sampling approach taken. Notably, these authors used olivine in place of the host bulk peridotites δ57Fe signatures despite this mineral has been shown to be frequently isotopically lighter than coexisting phases, and they analyzed lunar samples heavily affected chemically by the meteoritic bombardment, a process known to alter Fe isotope signatures. Their pallasite metal–silicate fractionation data are also likely biased by the approach adopted to estimate the iron isotope composition of the different mineral phases. In fact, their conclusion of Fe isotopic fractionation during basalt extraction from planetary mantles is invalidated by the observation that basaltic shergottites and eucrites have δ57Fe indistinguishable from those of chondrites. Therefore, the heavier Fe isotopic composition of the Moon relative to the Earth, itself heavier than most chondrites and achondrites remains best explained by loss of light iron isotopes during the high temperature event accompanying the interplanetary impact that led to the formation of the Moon [F. Poitrasson, A.N. Halliday, D.C. Lee, S. Levasseur, N. Teutsch, Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms, Earth Planet. Sci. Lett. 223 (2004) 253–266., F. Poitrasson, S. Levasseur, N. Teutsch, Significance of iron isotope mineral fractionation in pallasites and iron meteorites for the core–mantle differentiation of terrestrial planets, Earth Planet. Sci. Lett. 234 (2005) 151–164]. © 2007 Elsevier B.V. All rights reserved. Keywords: planet accretion; planet differentiation; mantle; core; crust; Earth; Moon; Mars; meteorites
DOI of original article: 10.1016/j.epsl.2007.01.038. A comment of “Iron isotope fractionation during planetary differentiation” by S. Weyer, A.D. Anbar, G.P. Brey, C. Münker, K. Mezger and A.B. Woodland, EPSL, 240: 251–264. E-mail address: [email protected].
1. Introduction
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0012-821X/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2007.01.039
In a recent study [1], we interpreted the differences of the mean iron isotope composition of bulk rocks from different planetary bodies as witnesses of their contrasted
F. Poitrasson / Earth and Planetary Science Letters 256 (2007) 484–492
accretion mechanisms. The iron isotope composition of lunar samples was on average heavier by 0.1‰ than mantle-derived rocks from the Earth, themselves heavier by 0.1‰ than SNC (Shergotty–Nakla–Chassigny) type and HED (Howardite–Eucrite–Diogenite) type meteorites. Our interpretation taking into account all available data is that the heavy isotopic composition of the Moon fingerprints the giant impact that led to its formation. It is likely that this high-energy event resulted in melting to partial vaporisation of portions of the early Earth and its impactor. It was shown quantitatively that during this process, light iron must have been lost to space, leaving planetary residues enriched in heavy Fe isotopes [1]. Subsequently, we studied the Fe isotope systematics of mineral separates from pallasites and iron meteorites [3]. It was found that the metal fraction of pallasites has a heavier iron isotope composition than the coexisting olivines, as are iron meteorites relative to chondrites. Extrapolating these results to the planetary scale, this would suggest that the silicate portion of planets – mantles – should be isotopically lighter than the chondritic starting material. That was not observed, however, since all planetary mantle-derived rocks analyzed have a Fe isotope composition similar, or heavier than those of chondrites. It was thus inferred that these mantle signatures could reflect equilibrium fractionation with metallic cores at high temperature, thus resulting in indiscernible metal–silicate fractionation. Alternatively, these results could reveal the lack of a global core–mantle equilibration at the planetary scale [3], thus leaving any possible Fe isotopic fractionation between the mean isotopic composition of the silicate mantle and the metallic core of planets. More recently, Weyer et al. [2] repeated these measurements using high resolution MC-ICP-MS with improved analytical reproducibilities over previous Fe isotopic analyses. While they were able to reproduce many of our previous measurements (see also the recent results of Schoenberg et al. [4], and the review of Dauphas and Rouxel [5]), they concluded that, because of their improved analytical techniques, they were able to find subtle isotopic differences between planetary reservoirs (i.e., between crust and mantle) that were previously unnoticed. On this basis, they reinterpreted the planetary differences we observed [1] as resulting not from contrasted accretion mechanisms, but merely as planetary differentiation processes. Although the analytical quality of the isotopic measurements of Weyer et al. [2] is beyond doubt, it is shown in this contribution that their conclusions were biased by inappropriate sampling for the question studied, plus an incomplete or oriented use of the literature. This
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study also suffers from a number of contradictions. From this, it is concluded that the new dataset of Weyer et al. [2] does not permit to assert that planetary differentiation induce Fe isotope fractionation. It is therefore maintained that given all data currently available, the most likely interpretation of the Fe isotope planetary differences remains that they trace contrasting planetary accretion mechanisms. 2. Mantle–core differentiation Given the lack of experimental data, this aspect has so far essentially been studied on the basis of Fe isotope measurements of pallasite mineral components and by comparing isotopic signatures obtained on chondrites, iron meteorites and other achondrites [3]. Both our previous study [3] and that of Weyer et al. [2] therefore rely on the assumption that pallasites are a valid proxy of the mantle–core interface of terrestrial planets for Fe isotopes. A quick comparison of Weyer et al. [2] data and ours [3], or the previous values of Zhu et al. [6], rapidly shows that there are significant differences between the Fe isotopic fractionation observed by Weyer et al. [2] compared to previous studies, even when the same meteorites have been analyzed (e.g., Eagle Station). Whereas Weyer et al. [2] found metal fractions either lighter or heavier than the olivines of various pallasites, we only measured metals with δ57Fe higher than the coexisting olivines. Where Weyer at al. [2] observed metal isotopically heavier than olivine, their magnitude was significantly different from those measured previously on the same meteorites (e.g., Marjalahti; [3]). Finally, there are also disagreements between the two studies on the isotopic fractionations inferred for troilite and schreibersite. These contradictory results cannot be due to an analytical bias because the analyses of well homogenised powders of 5 international rock standards (BIR1, BCR-1, BHVO-1, JP-1, PCC-1) reported by the two groups are in good agreement since they match within 0.03‰ in δ57Fe [1,2]. Weyer et al. [2] also obtained values that agree within this range with our own measurements of our in house hematite standard from Milhas, Pyrénées, used in Zürich and Toulouse. If the analytical procedures give similar results, the discrepancy between the two studies must be related to the different sampling methods used. Because the aim of our study was to monitor Fe isotope fractionation between coexisting minerals and to check whether they were in equilibrium (and not assume an a priori equilibrium, as incorrectly quoted by Weyer et al. [2]),
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we used in situ sampling tools (a microdiamond-coated dentist saw) to sample, within less than a cm, the minerals in contact after optical identification [3]. Instead, Weyer et al. [2] analyzed mineral grain separates without textural control. Given the level of Fe isotope heterogeneity measured in metal of iron meteorites, (taenite can be isotopically heavier than kamacite by up to 0.3‰ [3]), which may also occur in the pallasite metals, it is not surprising that handpicked grains, not related texturally, may give so variable results. Instead, sampling of the metal close to the olivine led us to sample kamacite, not taenite [7]. Another reason for the scatter in the Fe isotope composition of “metal fractions” from pallasites measured by Weyer at al. [2] is that there is in fact a mixture between metal and variable amount of S- and P-rich minerals, which show contrasting Fe isotopic signatures relative to the pallasitic metal, as discussed below. Moreover, because of the important Fe isotopic variation potentially associated with weathering and low temperature aqueous alteration [8,9], we only sampled and analyzed fresh pallasites. Instead, Weyer et al. [2] report analyses of metals that are sometimes “altered”, corresponding presumably to low temperature terrestrial weathering, which can be an additional cause of modification of the Fe isotope compositions. The contradictory isotopic results obtained by Weyer et al. [2] and Poitrasson et al. [3] on troilite and schreibersite led us to check by scanning electron microscopy (SEM) the sample fractions left that we initially characterized optically using reflected light under binocular lenses before in situ sampling. It turned out that the Energy Dispersive Spectra (EDS) of what we called “troilite” in Table 2 and Fig. 2 of [3] are in fact those of schreibersite, and what we took as “schreibersite” may in fact be a complicated mixture of iron metal and iron oxide (although for this last sample, there was only one grain left for SEM characterization). There are therefore probably no troilite Fe isotope data in [3], and the two schreibersite data (incorrectly labelled troilite in Table 2 and Fig. 2 of [3]) are isotopically heavier than the coexisting metal and olivine. This is more consistent with the findings of Weyer et al. [2]. On the other hand, and as detailed previously, results of Weyer et al. [2] are in contradiction with the metal– olivine fractionation measured in two previous Fe isotopic studies of pallasites [3,6]. These studies were in good agreement with the theoretical predictions of Polyakov and Mineev [10], notably for those pallasites where olivine and metal reached equilibrium. As noted above, a central aspect of our interpretation [3] was to evaluate whether metal and olivine from the pallasites
analyzed reached equilibrium or not. We did not “assume isotopic equilibrium between silicate and adjacent metal” as stated by Weyer et al. [2]. We evaluated the possibility of equilibrium in combining our Fe isotope measurements with a theoretical framework developed for oxygen isotopes [11], previous textural studies of pallasites [12], diffusion data in olivine [13] and theoretical iron isotope fractionation calculations [10]. On this basis, it was concluded that only two (Springwater and Marjalahti) of the set of five pallasites studied showed equilibrium Fe isotope fractionation between metal and olivine. Instead, Weyer et al. [2] initially stated that metal and silicate fractions of pallasites were not in isotopic equilibrium, and then interpreted their Fe isotope mean analyses on these phases using a thermodynamic framework with the dependence of Fe isotope fractionation as a function of temperature to infer the effect of planetary mantle–core differentiation on isotopic signatures, therefore contradicting their first assertion. If they conclude that the metal and silicate in pallasites were not in isotopic equilibrium, then these results do not bear any relevance regarding the mantle–core fractionation of Fe isotopes if this planetary process occurred at equilibrium. These inconsistencies make the conclusions of Weyer et al. [2] derived from their data on the lack of metal–silicate Fe isotope fractionation in planets an a priori statement, but not a deduction, since these conclusions cannot be derived from their data if their interpretation of isotopic disequilibrium is correct. In such a situation, metal and silicate may take eminently variable Fe isotope fractionation factors (e.g. [14]) and the results obtained on pallasites will bring no clues to understand planetary core–mantle differentiation using Fe isotopes. Our previous conclusions therefore appear to be the simplest interpretation of our data [3]. If the metal–silicate Fe isotopic equilibration observed in two pallasites can be transposed to planetary core–mantle differentiation, this process will likely generate a smaller Fe isotopic fractionation because of the higher temperatures involved. This statement was repeated by Weyer et al. [2], but as explained above, it cannot be logically deduced from their data. Furthermore, it is unlikely that the mixtures analyzed by [2], containing Fe metal and up to ∼ 60% troilite and up to ∼ 90% schreibersite will have any relevance to the matter that forms planetary cores given that they certainly incorporate much lesser amounts of S and P. Our alternative interpretation was that if the final core–mantle differentiation of planets did not involve a final metal–silicate equilibration, then the pallasite data add very little in the way of constraints for this process [3]. Given the importance of this issue
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on our understanding of the chemical composition of planetary mantles or on Hf–W chronometry, HP–HT experiments of metal–silicate fractionation are required to clarify this topic in the future. 3. Crust–mantle differentiation It follows from the previous discussion that there is currently no reason to consider that core–mantle differentiation generates sizeable Fe isotope fractionation. The next step is to evaluate whether crust–mantle differentiation of planets can lead to contrasted planetary silicate reservoirs. 3.1. The Earth case First, it should be stressed that there is at least a consensus on the iron isotopic composition of the Earth's igneous crustal rocks. In agreement with the initial study of Beard et al. [15], we also found that mantle-derived rocks are very homogeneous and yield a mean δ57Fe value of 0.102 ± 0.032‰ [1] (delta values are computed relative to IRMM-14 and uncertainties are expressed as two standard errors throughout this comment; see [16] for details). Only the most fractionated, silica-rich granites (with SiO2 N 71 wt.%) begin to show Fe isotopic composition heavier than this, possibly as a result of the exsolution of a fluid phase [16]. Weyer et al. [2] analyzed four basalts and one komatiite, mostly international rock standards previously analyzed by Beard et al. [15] and Poitrasson et al. [1], and concluded that the mean composition derived in these previous studies was correct. The central question then revolves around the iron isotopic composition of the terrestrial mantle. In their initial survey studies, Beard et al. [15], and Poitrasson et al. [1] analyzed a wide range of terrestrial igneous rocks and found no reason to conclude that peridotites pointed to a different bulk Fe isotope composition for the mantle. For example, from the limited set of peridotites we analyzed [1], two yielded lighter δ57Fe values than the terrestrial igneous mafic rocks (0.034 ± 0.038‰ for PCC-1 and 0.006 ± 0.034% for JP-1) whereas one was heavier (DTS-1: 0.159 ± 0.044‰). The lighter values for PCC-1 and JP-1 were attributed to the fluid–rock interaction previously documented in these peridotites [17,18] whereas the Twin Sister dunite (DTS-1) only shows a low degree of serpentinization [19]. It is therefore more likely to have kept a pristine Fe isotope signature. I purified a second time and reanalyzed the iron solution of DTS-1 using high resolution MC-ICP-MS following the procedure described
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in [16], and I found an iron isotope composition of δ57Fe = 0.148 ± 0.050‰, thus indistinguishable from the value previously published [1]. Unfortunately, Weyer et al. [2] did not reanalyze this international rock standard, nor they considered it in their discussion. Instead, they reported the iron isotope composition of 10 new samples, and reanalyzed the peridotite rock standards PCC-1 and JP-1. From these results, they concluded that the bulk terrestrial mantle has a δ57Fe of 0.023 ± 0.027‰, thus significantly lighter than the mean igneous rocks previously reported by Beard et al. [15] and Poitrasson et al. [1], that Weyer et al. [2] considered to be essentially crustal values only, despite the inclusion of mantle peridotites in these means. A major problem with the mantle mean from Weyer et al. [2] is that besides PCC-1 and JP-1, it only adds 4 bulk peridotite analyses to the database, the remaining 6 values being those of olivines from other peridotites. Their underlying assumption is that olivine iron isotope composition is not significantly different from bulk peridotites and therefore from other minerals in the same rocks. However, previous studies have shown that this is a wrong assumption since olivine is frequently isotopically lighter than other coexisting minerals, such as pyroxenes (Fig. 1). Moreover, olivines do not “contain most of the Fe in peridotites” as stated by Weyer et al. [2]. Pyroxenes also hold a notable fraction of the iron budget (e.g., [20]). As a result, the terrestrial mantle mean of Weyer et al. [2] is biased toward the light Fe isotopic signatures of olivine. For example, using one of the most comprehensive iron isotope study of the terrestrial mantle published so far [20], the mean of the δ57Fe olivine values yields − 0.05 ± 0.10‰ (n = 15) (Fig. 2), that is significantly lighter than our mean of mafic igneous rocks from the Earth of 0.102 ± 0.032‰ [1]. In contrast, the mean of the Fe isotopic composition of analyzed bulk mantle rocks (thus excluding recalculated values from mineral separate analyses, likely to be slightly less accurate) gives 0.08 ± 0.24‰ (n = 8) (Fig. 2). Although the large uncertainty shows that it is difficult to define a mean mantle value on the basis of the analysis of a few bulk mantle rocks, the mean value (0.08‰) is very close to our “mean mafic Earth” estimate (Fig. 2). In fact, the available data lend no support to consider that the Earth's mantle is isotopically different from the igneous crust (see also [21]). Interestingly, Beard and Johnson [21] also took the Fe isotopic composition of mineral separates as representative of bulk-rock composition for differentiated crustal rocks, and this led them to overlook the conspicuous increase in δ57Fe found in various granites and rhyolites beyond 71 wt.% SiO2 [16,22,23].
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Fig. 1. Compilation of the average iron isotope fractionation observed between clinopyroxene and olivine in mafic rocks. Except for the Beard and Johnson [21] study, all others conclude that clinopyroxenes are significantly heavier isotopically than coexisting olivine [1,6,20].
As revealed by the large uncertainty of the mean bulk mantle rock value (0.08 ± 0.24‰) reported in Fig. 2, the mantle studies of Beard and Johnson [21] and Williams et al. [20] have shown that mantle rocks display eminently variable Fe isotopic signatures. From these investigations, it is therefore becoming clear that using only 6 bulk peridotite analyses, besides olivine separates, to define the iron isotopic signature of the mantle [2] is not sufficient. Future studies conducted to assess whether the Fe isotopic variations in the mantle are the consequences of a varying redox state, partial melting [20] or metasomatism [21], will help to determine the geometry and size of the various Fe isotopic reservoirs, and ultimately to estimate the bulk mantle composition. A similar issue on mantle heterogeneity occurred more than a decade ago for oxygen isotopes and this was
resolved through a careful study of oxygen isotopic fractionation of main minerals from a large variety of representative samples from the mantle [24,25]. There are thus no reasons to conclude at this stage that the mean δ57Fe of the mantle is different from that of crustal igneous rocks and therefore that mantle partial melting generates magmas isotopically heavy, as advocated by Weyer et al. [2]. Instead, it is emphasized that basaltic shergottites and eucrites, interpreted to result from the partial melting of the mantle of their parent bodies, have δ57Fe values indistinguishable from those of chondrites (Fig. 3) [26]. This shows that on at least two different planetary bodies, Fe isotope fractionation did not occur during partial melting and crust formation since the basalts have δ57Fe indistinguishable from the chondritic starting material. Iron isotope fractionation during partial
Fig. 2. Mean iron isotope composition of whole mantle rock powders measured by Williams et al. [20], compared to their mean values obtained on olivine separates, quoted as δ57Fe relative to IRMM-14. The mean δ57Fe of mantle olivines is significantly lighter from the mean mafic Earth baseline previously defined [1], whereas the bulk mantle rock mean is very close to this reference value. However, as a result of the large iron isotopic scatter of bulk mantle rocks, the uncertainty is large and further study is required to define precisely the bulk mantle δ57Fe value (see text for details).
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Fig. 3. Comparison of the mean δ57Fe values relative to IRMM-14 for the mafic Earth, Moon and chondrites [1,3] with the mean of extraterrestrial basalts (from 6 basaltic shergottites and 9 eucrites). The chondritic average (− 0.042 ± 0.043‰) is indistinguishable from basalts from Mars and Vesta, thus illustrating that partial mantle melting and crust extraction did not generate Fe isotope fractionation. Data for eucrites and basaltic shergottites are from [1], except for Millbillillie, Stanern and Zagami from [2], and Los Angeles and Sayh al Uhaymir 005 from [26].
melting and basaltic melt extraction cannot therefore be an explanation for the 0.1‰ heavier than chondrites δ57Fe signature of terrestrial mantle-derived rocks [1]. The same conclusion applies for the 0.2‰ heavier than chondrites δ57Fe signature measured in Apollo lunar samples [1], as discussed below. 3.2. The Moon case Evaluating whether lunar crust formation induced Fe isotope fractionation is not an easy task given the lack of samples from the lunar mantle (e.g., [27]). We are thus left with petrological considerations, comparisons with other planetary bodies and with theoretical estimates of Fe isotope fractionation. Weyer et al. [2] analyzed a series of lunar rocks and obtained a more scattered set of bulk-rock Fe isotopic data (from 0.042‰ in a “highland breccia” to 0.314‰ in a high-Ti basalt, with a mean δ57Fe of 0.156 ± 0.063‰, n = 12) than what we previously reported (from 0.111‰ for a quartz-normative basalt to 0.274‰ for a high-Ti basalt, with a mean δ57Fe of 0.206 ± 0.029‰, n = 14; [1]). This is despite their petrological range of igneous rock types was more restricted since they did not consider the norites, picritic basalts and anorthosites we analyzed [1]. Despite this isotopic scatter, they separated their samples in two groups, one made of high-Ti basalt with a heavy Fe isotope composition, and the rest, essentially consisting of “highland rocks”, low-Ti basalt and “lunar glasses” sharing an iron isotope composition similar to the five terrestrial basalt they analyzed, according to these authors [2].
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Again, the Weyer et al. [2] interpretations are biased by their sampling strategy. One first interesting feature is that although there is a good agreement between the Fe isotopic determinations from the two groups on well homogenised powders from terrestrial rock standards (see above), this is much less the case on the lunar samples analyzed by both Weyer et al. [2] and Poitrasson et al. [1]. The agreement is good for the KREEP basalt 15386 and quartz-normative basalt 15475, but it is much less so for olivine normative basalt 15555 and the two lunar ultramafic glasses 15426 and 74220. Lunar samples are precious and it is therefore not possible to work on powders produced out of large chunks of rocks. In our study, we worked on powders made of at least 1 g of rock [1], whereas Weyer et al. [2] made their powders from rock samples ten times smaller. Given the importance of Fe isotope mineral fractionation (Fig. 1) relative to the bulk-rock variations [1,23], one possible explanation of the scatter of Weyer et al. [2] data results from powders not representative enough of the bulk rocks. Samples 15426 and 74220 are made of glass beads of volcanic ultramafic glasses mixed with soil containing ejecta due to the meteoric bombardment of the lunar surface [28], potentially heavier isotopically than the volcanic glass beads [29]. We purified these glass beads under binocular lenses before chemical treatment and analysis, but it is not explained in the Weyer et al. [2] study whether they did so as well. If not, this can be an explanation of the higher δ57Fe values of Weyer et al. [2] on lunar glasses 15426 and 74220 compared to ours. Another critical aspect of our sampling is that we avoided rock samples chemically affected by the meteoritic bombardment as much as possible [1]. Instead, the lunar highland samples analyzed by Weyer et al. [2] were all impact melts mixed with clasts, contaminated chemically by meteoritic bombardment [28,30–33]. This is attested by their high siderophile contents. Rocks like 62235, for example, have Fe metal inclusions that have Co and Ni compositions within the meteoritic range [34]. We avoided analyzing these samples because besides the ejectas produced by the meteoritic bombardment, which are potentially heavy in their δ57Fe composition [29], the complementary fractions are likely to be isotopically lighter. Furthermore, the meteoritic iron contamination of these samples will shift the isotopic values of pristine lunar igneous rocks towards the lower Fe isotopic signatures of chondrites [3]. All these would therefore favour an isotopic scatter among samples and an overall lowering of the Fe isotope composition due to meteorite impacts (Fig. 4). We also found [1], like Weyer et al. [2], that high-Ti basalts tend to show a heavy Fe isotope composition.
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Fig. 4. Comparison of the lunar mean whole-rock δ57Fe from the data of Weyer et al. [2] and Poitrasson et al. [1], excluding the high-Ti basalts which tend to be isotopically heavy. This plot nevertheless illustrates that the mean δ57Fe lunar value computed using bulk rocks analyzed by [1], without high-Ti basalts, is indistinguishable from the lunar mean previously defined including high-Ti basalts [1]. This shows that these specific kinds of samples do not affect significantly the lunar mean bulk-rock Fe isotopic composition. Instead, the Weyer et al. [2] value (δ57Fe = 0.108 ± 0.042‰; n = 9) is lower than the mean Moon δ57Fe previously defined [1] because it comprises a large number of samples heavily affected by the meteoritic bombardment.
However, our larger petrological range of rock types, chemically unaffected by the meteoric bombardment, also shows that bulk anorthosites, quartz-normative low-Ti basalts and KREEP basalts can be isotopically as heavy [1]. There are therefore no reasons to put the high-Ti basalt into a separate group on the basis of their Fe isotope signatures as done by [2]. Indeed, the mean of bulk igneous lunar rocks from [1], excluding high-Ti basalts (δ57Fe of 0.185 ± 0.038‰, n = 9) is indistinguishable with the mean originally published by [1] that includes these Ti-rich rocks (Fig. 4). Thus, taking all the lunar samples analyzed by Weyer et al. [2] representative of bulk rocks (i.e., excluding the soil sample 10084 and “lunar volcanic glass” clods 74220 and 15426), the computed lunar mean (δ57Fe of 0.156 ± 0.063‰, n = 12) is undistinguishable from that published earlier (δ57Fe of 0.206 ± 0.029‰, n = 14; [1]). Their lower value and larger uncertainty probably result largely from the effect of meteoritic bombardment on many of the samples selected by these authors for their study. Weyer et al. [2] attempted to explain a heavier Fe isotope composition of high-Ti basalts in terms of the involvement of ilmenite in the petrogenesis of these rocks. However, as previously discussed [1], the Polyakov and Mineev [10] theoretical estimates suggest that ilmenite is one of the high temperature Fe-bearing igneous mineral that fractionates the less iron isotopes. There is therefore no justification to assume that this phase played a specific role explaining the Fe isotope composition of high-Ti
basalt that in fact is, as shown above, not as different as implied by Weyer et al. [2] compared to other bulk lunar igneous samples. If, for example, massive ilmenite fractional crystallization occurred during magma differentiation, this would be accompanied by a lowering of Ti and little impact on, or in an extreme case a slight increase in δ57Fe of the magma. Conversely, assimilation of ilmenite in the melt would result in a minimal effect to slight lowering of δ57Fe values of the magma, but a significant increase in its Ti content. In both cases, the resulting hypothetic melt would not correspond to an isotopically heavy, Ti-rich basalt. Hence, there is currently no evidence to hypothesize that the production of a lunar crust fractionated Fe isotopes and that the heavy isotopic composition of lunar samples result from planetary differentiation. 4. Conclusions Knowing whether planetary differentiation generates Fe isotope fractionation requires first good estimates of the main planetary reservoir compositions. The Weyer et al. [2] study shows that inappropriate sampling may lead to exaggerated isotopic variability (e.g., in their lunar rocks) and biased estimates of bulk planetary reservoirs (i.e., the Earth's mantle). It is also important to take fully into account data from all previously published studies (e.g., [6]) and the available theoretical estimates of high temperature Fe isotope fractionation between minerals [10].
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The large variability of the Fe isotope composition of bulk mantle rocks [20] remains puzzling, particularly when compared to the isotopic homogeneity of mantlederived rocks [1,15,21]. Therefore, it is first necessary to understand the cause of this variability before a bulk mantle Fe isotope composition can be precisely established. Experimental studies of peridotite partial melting are required to monitor the effect of this process on the Fe isotopic signature on magmas. So far, there are no reasons to consider that the Earth's mantle has a Fe isotope composition different from that of igneous crustal rocks (see also [21]). The current Fe isotopic database, including notably the fact that Martian basalts and eucrites have a Fe isotope composition undistinguishable from that of the building block of planets, chondrites [1,3], makes untenable the suggestion of Weyer et al. [2] that planetary differentiation, and particularly the production of a crust, can generate Fe isotope fractionation. The same conclusion is also valid for the mantle–core differentiation if metal and silicate remained in equilibrium [3]. It is therefore re-emphasized that the ∼ 0.1‰ heavier Fe isotope composition of the mafic silicate Earth, and especially the ∼ 0.2‰ heavier iron isotope signature of samples from the Moon, relative to chondrites, SNC and HED meteorites remain best explained by contrasted planetary accretion mechanisms [1,3]. This interpretation remains the most logical when all current available data are taken into account. It is also promising since it implies that Fe isotopes can trace specific processes not easily tracked down by other geochemical tools. Acknowledgements Thierry Aigouy is acknowledged for his help with the SEM characterization of pallasite separates. Guillaume Delpech is thanked for a careful reading of the manuscript, an anonymous referee for a critical and helpful review and Rick Carlson for his editorial handling. References [1] F. Poitrasson, A.N. Halliday, D.C. Lee, S. Levasseur, N. Teutsch, Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms, Earth Planet. Sci. Lett. 223 (2004) 253–266. [2] S. Weyer, A.D. Anbar, G.P. Brey, C. Munker, K. Mezger, A.B. Woodland, Iron isotope fractionation during planetary differentiation, Earth Planet. Sci. Lett. 240 (2005) 251–264. [3] F. Poitrasson, S. Levasseur, N. Teutsch, Significance of iron isotope mineral fractionation in pallasites and iron meteorites for the core–mantle differentiation of terrestrial planets, Earth Planet. Sci. Lett. 234 (2005) 151–164.
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[4] R. Schoenberg, B.S. Kamber, F. von Blanckenburg, Comparative stable Fe isotope systematics of terrestrial and meteoritic materials, Geochim. Cosmochim. Acta 69 (2005) A398. [5] N. Dauphas, O. Rouxel, Mass spectrometry and natural variations of iron isotopes, Mass Spectrom. Rev. 25 (2006) 515–550. [6] X.K. Zhu, Y. Guo, R.J.P. Williams, R.K. O'Nions, A. Matthews, N.S. Belshaw, G.W. Canters, E.C. de Waal, U. Weser, B.K. Burgess, B. Salvato, Mass fractionation processes of transition metal isotopes, Earth Planet. Sci. Lett. 200 (2002) 47–62. [7] P.R. Buseck, Pallasite meteorites—mineralogy, petrology and geochemistry, Geochim. Cosmochim. Acta 41 (1977) 711–740. [8] O. Rouxel, N. Dobbek, J. Ludden, Y. Fouquet, Iron isotope fractionation during oceanic crust alteration, Chem. Geol. 202 (2003) 155–182. [9] M.S. Fantle, D.J. DePaolo, Iron isotopic fractionation during continental weathering, Earth Planet. Sci. Lett. 228 (2004) 547–562. [10] V.B. Polyakov, S.D. Mineev, The use of Mössbauer spectroscopy in stable isotope geochemistry, Geochim. Cosmochim. Acta 64 (2000) 849–865. [11] R.T. Gregory, R.E. Criss, Isotopic exchange in open and closed systems, in: J.W. Valley, H.P. Taylor, J.R. O'Neil (Eds.), Stable Isotopes in High Temperature Geological Processes, Reviews in Mineralogy, vol. 16, Mineralogical Society of America, Washington, 1986, pp. 91–127. [12] E. Ohtani, Formation of olivine textures in pallasites and thermal history of pallasites in their parent body, Phys Earth Planet. Inter. 32 (1983) 182–192. [13] O. Jaoul, Y. Bertran-Alvarez, R.C. Liebermann, G.D. Price, Fe– Mg interdiffusion in olivine up to 9 GPa at T = 600–900 °C; experimental data and comparison with defect calculations, Phys Earth Planet. Inter. 89 (1995) 199–218. [14] M. Roskosz, B. Luais, H.C. Watson, M.J. Toplis, C.M.O. Alexander, B.O. Mysen, Experimental quantification of the fractionation of Fe isotopes during metal segregation from a silicate melt, Earth Planet. Sci. Lett. 248 (2006) 851–867. [15] B.L. Beard, C.M. Johnson, J.L. Skulan, K.H. Nealson, L. Cox, H. Sun, Application of Fe isotopes to tracing the geochemical and biological cycling of Fe, Chem. Geol. 195 (2003) 87–117. [16] F. Poitrasson, R. Freydier, Heavy iron isotope composition of granites determined by high resolution MC-ICP-MS, Chem. Geol. 222 (2005) 132–147. [17] F.J. Flanagan, U.S. Geological Survey silicate rock standards, Geochim. Cosmochim. Acta 31 (1967) 289–308. [18] T. Morishita, S. Arai, Evolution of spinel-pyroxene symplectite in spinel-lherzolites from the Horoman Complex, Japan, Contrib. Mineral. Petrol. 144 (2003) 509–522. [19] E.C. Ferré, B. Tikoff, M. Jackson, The magnetic anisotropy of mantle peridotites: example from the Twin Sister dunite, Washington, Tectonophysics 398 (2005) 141–166. [20] H.M. Williams, A.H. Peslier, C. McCammon, A.N. Halliday, S. Levasseur, N. Teutsch, J.P. Burg, Systematic iron isotope fractionation variations in mantle rocks and minerals: the effect of partial melting and oxygen fugacity, Earth Planet. Sci. Lett. 235 (2005) 435–452. [21] B.L. Beard, C.M. Johnson, Inter-mineral Fe isotope variations in mantle-derived rocks and implications for the Fe geochemical cycle, Geochim. Cosmochim. Acta 68 (2004) 4727–4743. [22] B.L. Beard, C.M. Johnson, Comment on “Heavy iron isotope composition of granites determined by high resolution MC-ICPMS” by F. Poitrasson and R. Freydier, Chemical Geology, volume 222, pages 132–147, Chem. Geol. 235 (2006) 201–204.
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[23] F. Poitrasson, On the iron isotope homogeneity level of the continental crust, Chem. Geol. 235 (2006) 195–200. [24] D.P. Mattey, D. Lowry, C. MacPherson, Oxygen isotope composition of mantle peridotites, Earth Planet. Sci. Lett. 128 (1994) 231–241. [25] G. Chazot, D. Lowry, M. Menzies, D.P. Mattey, Oxygen isotopic composition of hydrous and anhydrous mantle peridotites, Geochim. Cosmochim. Acta 61 (1997) 161–169. [26] M. Anand, Fe isotopic composition of Martian meteorites, LPSC, vol. XXXVI, 2005, p. 1859, pdf, Houston. [27] P.H. Warren, A concise compilation of petrologic information on possibly pristine nonmare Moon rocks, Am. Mineral. 78 (1993) 360–376. [28] G. Ryder, Catalog of Apollo 15 Rocks. Part 2, vol. 15306–15468, NASA, Houston, 1985, pp. 339–777.
[29] R.A. Wiesli, B.L. Beard, L.A. Taylor, C.M. Johnson, Space weathering processes on airless bodies: Fe isotope fractionation in the lunar regolith, Earth Planet. Sci. Lett. 216 (2003) 457–465. [30] I.C. Carlson, J.A. Wayne, Apollo 14 Rock Samples, NASA, Houston, 1978, p. 413. [31] G. Ryder, M.D. Norman, Catalog of Apollo 16 Rocks. Part 1, vol. 60015–62315, NASA, Houston, 1980, p. 349. [32] G. Ryder, M.D. Norman, Catalog of Apollo 16 Rocks. Part 2, vol. 63335–66095, NASA, Houston, 1980, pp. 351–773. [33] C. Meyer, Catalog of Apollo 17 Rocks. Volume 4 — North Massif, NASA, Houston, 1994, p. 644. [34] G.W. Pearce, G.S. Hoye, D.W. Strangway, B.M. Walker, L.A. Taylor, Some complexities in the determination of lunar paleointensities, Proc. Lunar Planet. Sci. Conf. 7th, 1976, pp. 3271–3297.
Discussion
Does planetary differentiation really fractionate iron isotopes? ☆ Franck Poitrasson Laboratoire d'étude des Mécanismes de Transfert en Géologie, Centre National de la Recherche Scientifique-UPS-IRD, 14–16, avenue Edouard Belin, 31400 Toulouse, France Received 10 June 2006; received in revised form 31 October 2006; accepted 31 January 2007 Available online 11 February 2007 Editor: R.W. Carlson
Abstract The difference in the mean Fe isotope composition of samples from the Earth, Moon, Mars and Vesta has been recently interpreted as tracking contrasted planetary accretion mechanisms [F. Poitrasson, A.N. Halliday, D.C. Lee, S. Levasseur, N. Teutsch, Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms, Earth Planet. Sci. Lett. 223 (2004) 253–266]. Using newly produced Fe isotopic data on terrestrial and lunar samples, pallasites, eucrites and Martian meteorites, Weyer et al. [S. Weyer, A.D. Anbar, G.P. Brey, C. Munker, K. Mezger, A.B. Woodland, Iron isotope fractionation during planetary differentiation, Earth Planet. Sci. Lett. 240 (2005) 251–264] reinterpreted these data as fingerprinting planetary differentiation. In particular, these authors suggested that partial melting in the terrestrial and lunar mantles produced melts isotopically heavy. It is shown here that the inference of Weyer et al. [S. Weyer, A.D. Anbar, G.P. Brey, C. Munker, K. Mezger, A.B. Woodland, Iron isotope fractionation during planetary differentiation, Earth Planet. Sci. Lett. 240 (2005) 251–264] is strongly biased by the sampling approach taken. Notably, these authors used olivine in place of the host bulk peridotites δ57Fe signatures despite this mineral has been shown to be frequently isotopically lighter than coexisting phases, and they analyzed lunar samples heavily affected chemically by the meteoritic bombardment, a process known to alter Fe isotope signatures. Their pallasite metal–silicate fractionation data are also likely biased by the approach adopted to estimate the iron isotope composition of the different mineral phases. In fact, their conclusion of Fe isotopic fractionation during basalt extraction from planetary mantles is invalidated by the observation that basaltic shergottites and eucrites have δ57Fe indistinguishable from those of chondrites. Therefore, the heavier Fe isotopic composition of the Moon relative to the Earth, itself heavier than most chondrites and achondrites remains best explained by loss of light iron isotopes during the high temperature event accompanying the interplanetary impact that led to the formation of the Moon [F. Poitrasson, A.N. Halliday, D.C. Lee, S. Levasseur, N. Teutsch, Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms, Earth Planet. Sci. Lett. 223 (2004) 253–266., F. Poitrasson, S. Levasseur, N. Teutsch, Significance of iron isotope mineral fractionation in pallasites and iron meteorites for the core–mantle differentiation of terrestrial planets, Earth Planet. Sci. Lett. 234 (2005) 151–164]. © 2007 Elsevier B.V. All rights reserved. Keywords: planet accretion; planet differentiation; mantle; core; crust; Earth; Moon; Mars; meteorites
DOI of original article: 10.1016/j.epsl.2007.01.038. A comment of “Iron isotope fractionation during planetary differentiation” by S. Weyer, A.D. Anbar, G.P. Brey, C. Münker, K. Mezger and A.B. Woodland, EPSL, 240: 251–264. E-mail address: [email protected].
1. Introduction
☆
0012-821X/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2007.01.039
In a recent study [1], we interpreted the differences of the mean iron isotope composition of bulk rocks from different planetary bodies as witnesses of their contrasted
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accretion mechanisms. The iron isotope composition of lunar samples was on average heavier by 0.1‰ than mantle-derived rocks from the Earth, themselves heavier by 0.1‰ than SNC (Shergotty–Nakla–Chassigny) type and HED (Howardite–Eucrite–Diogenite) type meteorites. Our interpretation taking into account all available data is that the heavy isotopic composition of the Moon fingerprints the giant impact that led to its formation. It is likely that this high-energy event resulted in melting to partial vaporisation of portions of the early Earth and its impactor. It was shown quantitatively that during this process, light iron must have been lost to space, leaving planetary residues enriched in heavy Fe isotopes [1]. Subsequently, we studied the Fe isotope systematics of mineral separates from pallasites and iron meteorites [3]. It was found that the metal fraction of pallasites has a heavier iron isotope composition than the coexisting olivines, as are iron meteorites relative to chondrites. Extrapolating these results to the planetary scale, this would suggest that the silicate portion of planets – mantles – should be isotopically lighter than the chondritic starting material. That was not observed, however, since all planetary mantle-derived rocks analyzed have a Fe isotope composition similar, or heavier than those of chondrites. It was thus inferred that these mantle signatures could reflect equilibrium fractionation with metallic cores at high temperature, thus resulting in indiscernible metal–silicate fractionation. Alternatively, these results could reveal the lack of a global core–mantle equilibration at the planetary scale [3], thus leaving any possible Fe isotopic fractionation between the mean isotopic composition of the silicate mantle and the metallic core of planets. More recently, Weyer et al. [2] repeated these measurements using high resolution MC-ICP-MS with improved analytical reproducibilities over previous Fe isotopic analyses. While they were able to reproduce many of our previous measurements (see also the recent results of Schoenberg et al. [4], and the review of Dauphas and Rouxel [5]), they concluded that, because of their improved analytical techniques, they were able to find subtle isotopic differences between planetary reservoirs (i.e., between crust and mantle) that were previously unnoticed. On this basis, they reinterpreted the planetary differences we observed [1] as resulting not from contrasted accretion mechanisms, but merely as planetary differentiation processes. Although the analytical quality of the isotopic measurements of Weyer et al. [2] is beyond doubt, it is shown in this contribution that their conclusions were biased by inappropriate sampling for the question studied, plus an incomplete or oriented use of the literature. This
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study also suffers from a number of contradictions. From this, it is concluded that the new dataset of Weyer et al. [2] does not permit to assert that planetary differentiation induce Fe isotope fractionation. It is therefore maintained that given all data currently available, the most likely interpretation of the Fe isotope planetary differences remains that they trace contrasting planetary accretion mechanisms. 2. Mantle–core differentiation Given the lack of experimental data, this aspect has so far essentially been studied on the basis of Fe isotope measurements of pallasite mineral components and by comparing isotopic signatures obtained on chondrites, iron meteorites and other achondrites [3]. Both our previous study [3] and that of Weyer et al. [2] therefore rely on the assumption that pallasites are a valid proxy of the mantle–core interface of terrestrial planets for Fe isotopes. A quick comparison of Weyer et al. [2] data and ours [3], or the previous values of Zhu et al. [6], rapidly shows that there are significant differences between the Fe isotopic fractionation observed by Weyer et al. [2] compared to previous studies, even when the same meteorites have been analyzed (e.g., Eagle Station). Whereas Weyer et al. [2] found metal fractions either lighter or heavier than the olivines of various pallasites, we only measured metals with δ57Fe higher than the coexisting olivines. Where Weyer at al. [2] observed metal isotopically heavier than olivine, their magnitude was significantly different from those measured previously on the same meteorites (e.g., Marjalahti; [3]). Finally, there are also disagreements between the two studies on the isotopic fractionations inferred for troilite and schreibersite. These contradictory results cannot be due to an analytical bias because the analyses of well homogenised powders of 5 international rock standards (BIR1, BCR-1, BHVO-1, JP-1, PCC-1) reported by the two groups are in good agreement since they match within 0.03‰ in δ57Fe [1,2]. Weyer et al. [2] also obtained values that agree within this range with our own measurements of our in house hematite standard from Milhas, Pyrénées, used in Zürich and Toulouse. If the analytical procedures give similar results, the discrepancy between the two studies must be related to the different sampling methods used. Because the aim of our study was to monitor Fe isotope fractionation between coexisting minerals and to check whether they were in equilibrium (and not assume an a priori equilibrium, as incorrectly quoted by Weyer et al. [2]),
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we used in situ sampling tools (a microdiamond-coated dentist saw) to sample, within less than a cm, the minerals in contact after optical identification [3]. Instead, Weyer et al. [2] analyzed mineral grain separates without textural control. Given the level of Fe isotope heterogeneity measured in metal of iron meteorites, (taenite can be isotopically heavier than kamacite by up to 0.3‰ [3]), which may also occur in the pallasite metals, it is not surprising that handpicked grains, not related texturally, may give so variable results. Instead, sampling of the metal close to the olivine led us to sample kamacite, not taenite [7]. Another reason for the scatter in the Fe isotope composition of “metal fractions” from pallasites measured by Weyer at al. [2] is that there is in fact a mixture between metal and variable amount of S- and P-rich minerals, which show contrasting Fe isotopic signatures relative to the pallasitic metal, as discussed below. Moreover, because of the important Fe isotopic variation potentially associated with weathering and low temperature aqueous alteration [8,9], we only sampled and analyzed fresh pallasites. Instead, Weyer et al. [2] report analyses of metals that are sometimes “altered”, corresponding presumably to low temperature terrestrial weathering, which can be an additional cause of modification of the Fe isotope compositions. The contradictory isotopic results obtained by Weyer et al. [2] and Poitrasson et al. [3] on troilite and schreibersite led us to check by scanning electron microscopy (SEM) the sample fractions left that we initially characterized optically using reflected light under binocular lenses before in situ sampling. It turned out that the Energy Dispersive Spectra (EDS) of what we called “troilite” in Table 2 and Fig. 2 of [3] are in fact those of schreibersite, and what we took as “schreibersite” may in fact be a complicated mixture of iron metal and iron oxide (although for this last sample, there was only one grain left for SEM characterization). There are therefore probably no troilite Fe isotope data in [3], and the two schreibersite data (incorrectly labelled troilite in Table 2 and Fig. 2 of [3]) are isotopically heavier than the coexisting metal and olivine. This is more consistent with the findings of Weyer et al. [2]. On the other hand, and as detailed previously, results of Weyer et al. [2] are in contradiction with the metal– olivine fractionation measured in two previous Fe isotopic studies of pallasites [3,6]. These studies were in good agreement with the theoretical predictions of Polyakov and Mineev [10], notably for those pallasites where olivine and metal reached equilibrium. As noted above, a central aspect of our interpretation [3] was to evaluate whether metal and olivine from the pallasites
analyzed reached equilibrium or not. We did not “assume isotopic equilibrium between silicate and adjacent metal” as stated by Weyer et al. [2]. We evaluated the possibility of equilibrium in combining our Fe isotope measurements with a theoretical framework developed for oxygen isotopes [11], previous textural studies of pallasites [12], diffusion data in olivine [13] and theoretical iron isotope fractionation calculations [10]. On this basis, it was concluded that only two (Springwater and Marjalahti) of the set of five pallasites studied showed equilibrium Fe isotope fractionation between metal and olivine. Instead, Weyer et al. [2] initially stated that metal and silicate fractions of pallasites were not in isotopic equilibrium, and then interpreted their Fe isotope mean analyses on these phases using a thermodynamic framework with the dependence of Fe isotope fractionation as a function of temperature to infer the effect of planetary mantle–core differentiation on isotopic signatures, therefore contradicting their first assertion. If they conclude that the metal and silicate in pallasites were not in isotopic equilibrium, then these results do not bear any relevance regarding the mantle–core fractionation of Fe isotopes if this planetary process occurred at equilibrium. These inconsistencies make the conclusions of Weyer et al. [2] derived from their data on the lack of metal–silicate Fe isotope fractionation in planets an a priori statement, but not a deduction, since these conclusions cannot be derived from their data if their interpretation of isotopic disequilibrium is correct. In such a situation, metal and silicate may take eminently variable Fe isotope fractionation factors (e.g. [14]) and the results obtained on pallasites will bring no clues to understand planetary core–mantle differentiation using Fe isotopes. Our previous conclusions therefore appear to be the simplest interpretation of our data [3]. If the metal–silicate Fe isotopic equilibration observed in two pallasites can be transposed to planetary core–mantle differentiation, this process will likely generate a smaller Fe isotopic fractionation because of the higher temperatures involved. This statement was repeated by Weyer et al. [2], but as explained above, it cannot be logically deduced from their data. Furthermore, it is unlikely that the mixtures analyzed by [2], containing Fe metal and up to ∼ 60% troilite and up to ∼ 90% schreibersite will have any relevance to the matter that forms planetary cores given that they certainly incorporate much lesser amounts of S and P. Our alternative interpretation was that if the final core–mantle differentiation of planets did not involve a final metal–silicate equilibration, then the pallasite data add very little in the way of constraints for this process [3]. Given the importance of this issue
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on our understanding of the chemical composition of planetary mantles or on Hf–W chronometry, HP–HT experiments of metal–silicate fractionation are required to clarify this topic in the future. 3. Crust–mantle differentiation It follows from the previous discussion that there is currently no reason to consider that core–mantle differentiation generates sizeable Fe isotope fractionation. The next step is to evaluate whether crust–mantle differentiation of planets can lead to contrasted planetary silicate reservoirs. 3.1. The Earth case First, it should be stressed that there is at least a consensus on the iron isotopic composition of the Earth's igneous crustal rocks. In agreement with the initial study of Beard et al. [15], we also found that mantle-derived rocks are very homogeneous and yield a mean δ57Fe value of 0.102 ± 0.032‰ [1] (delta values are computed relative to IRMM-14 and uncertainties are expressed as two standard errors throughout this comment; see [16] for details). Only the most fractionated, silica-rich granites (with SiO2 N 71 wt.%) begin to show Fe isotopic composition heavier than this, possibly as a result of the exsolution of a fluid phase [16]. Weyer et al. [2] analyzed four basalts and one komatiite, mostly international rock standards previously analyzed by Beard et al. [15] and Poitrasson et al. [1], and concluded that the mean composition derived in these previous studies was correct. The central question then revolves around the iron isotopic composition of the terrestrial mantle. In their initial survey studies, Beard et al. [15], and Poitrasson et al. [1] analyzed a wide range of terrestrial igneous rocks and found no reason to conclude that peridotites pointed to a different bulk Fe isotope composition for the mantle. For example, from the limited set of peridotites we analyzed [1], two yielded lighter δ57Fe values than the terrestrial igneous mafic rocks (0.034 ± 0.038‰ for PCC-1 and 0.006 ± 0.034% for JP-1) whereas one was heavier (DTS-1: 0.159 ± 0.044‰). The lighter values for PCC-1 and JP-1 were attributed to the fluid–rock interaction previously documented in these peridotites [17,18] whereas the Twin Sister dunite (DTS-1) only shows a low degree of serpentinization [19]. It is therefore more likely to have kept a pristine Fe isotope signature. I purified a second time and reanalyzed the iron solution of DTS-1 using high resolution MC-ICP-MS following the procedure described
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in [16], and I found an iron isotope composition of δ57Fe = 0.148 ± 0.050‰, thus indistinguishable from the value previously published [1]. Unfortunately, Weyer et al. [2] did not reanalyze this international rock standard, nor they considered it in their discussion. Instead, they reported the iron isotope composition of 10 new samples, and reanalyzed the peridotite rock standards PCC-1 and JP-1. From these results, they concluded that the bulk terrestrial mantle has a δ57Fe of 0.023 ± 0.027‰, thus significantly lighter than the mean igneous rocks previously reported by Beard et al. [15] and Poitrasson et al. [1], that Weyer et al. [2] considered to be essentially crustal values only, despite the inclusion of mantle peridotites in these means. A major problem with the mantle mean from Weyer et al. [2] is that besides PCC-1 and JP-1, it only adds 4 bulk peridotite analyses to the database, the remaining 6 values being those of olivines from other peridotites. Their underlying assumption is that olivine iron isotope composition is not significantly different from bulk peridotites and therefore from other minerals in the same rocks. However, previous studies have shown that this is a wrong assumption since olivine is frequently isotopically lighter than other coexisting minerals, such as pyroxenes (Fig. 1). Moreover, olivines do not “contain most of the Fe in peridotites” as stated by Weyer et al. [2]. Pyroxenes also hold a notable fraction of the iron budget (e.g., [20]). As a result, the terrestrial mantle mean of Weyer et al. [2] is biased toward the light Fe isotopic signatures of olivine. For example, using one of the most comprehensive iron isotope study of the terrestrial mantle published so far [20], the mean of the δ57Fe olivine values yields − 0.05 ± 0.10‰ (n = 15) (Fig. 2), that is significantly lighter than our mean of mafic igneous rocks from the Earth of 0.102 ± 0.032‰ [1]. In contrast, the mean of the Fe isotopic composition of analyzed bulk mantle rocks (thus excluding recalculated values from mineral separate analyses, likely to be slightly less accurate) gives 0.08 ± 0.24‰ (n = 8) (Fig. 2). Although the large uncertainty shows that it is difficult to define a mean mantle value on the basis of the analysis of a few bulk mantle rocks, the mean value (0.08‰) is very close to our “mean mafic Earth” estimate (Fig. 2). In fact, the available data lend no support to consider that the Earth's mantle is isotopically different from the igneous crust (see also [21]). Interestingly, Beard and Johnson [21] also took the Fe isotopic composition of mineral separates as representative of bulk-rock composition for differentiated crustal rocks, and this led them to overlook the conspicuous increase in δ57Fe found in various granites and rhyolites beyond 71 wt.% SiO2 [16,22,23].
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Fig. 1. Compilation of the average iron isotope fractionation observed between clinopyroxene and olivine in mafic rocks. Except for the Beard and Johnson [21] study, all others conclude that clinopyroxenes are significantly heavier isotopically than coexisting olivine [1,6,20].
As revealed by the large uncertainty of the mean bulk mantle rock value (0.08 ± 0.24‰) reported in Fig. 2, the mantle studies of Beard and Johnson [21] and Williams et al. [20] have shown that mantle rocks display eminently variable Fe isotopic signatures. From these investigations, it is therefore becoming clear that using only 6 bulk peridotite analyses, besides olivine separates, to define the iron isotopic signature of the mantle [2] is not sufficient. Future studies conducted to assess whether the Fe isotopic variations in the mantle are the consequences of a varying redox state, partial melting [20] or metasomatism [21], will help to determine the geometry and size of the various Fe isotopic reservoirs, and ultimately to estimate the bulk mantle composition. A similar issue on mantle heterogeneity occurred more than a decade ago for oxygen isotopes and this was
resolved through a careful study of oxygen isotopic fractionation of main minerals from a large variety of representative samples from the mantle [24,25]. There are thus no reasons to conclude at this stage that the mean δ57Fe of the mantle is different from that of crustal igneous rocks and therefore that mantle partial melting generates magmas isotopically heavy, as advocated by Weyer et al. [2]. Instead, it is emphasized that basaltic shergottites and eucrites, interpreted to result from the partial melting of the mantle of their parent bodies, have δ57Fe values indistinguishable from those of chondrites (Fig. 3) [26]. This shows that on at least two different planetary bodies, Fe isotope fractionation did not occur during partial melting and crust formation since the basalts have δ57Fe indistinguishable from the chondritic starting material. Iron isotope fractionation during partial
Fig. 2. Mean iron isotope composition of whole mantle rock powders measured by Williams et al. [20], compared to their mean values obtained on olivine separates, quoted as δ57Fe relative to IRMM-14. The mean δ57Fe of mantle olivines is significantly lighter from the mean mafic Earth baseline previously defined [1], whereas the bulk mantle rock mean is very close to this reference value. However, as a result of the large iron isotopic scatter of bulk mantle rocks, the uncertainty is large and further study is required to define precisely the bulk mantle δ57Fe value (see text for details).
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Fig. 3. Comparison of the mean δ57Fe values relative to IRMM-14 for the mafic Earth, Moon and chondrites [1,3] with the mean of extraterrestrial basalts (from 6 basaltic shergottites and 9 eucrites). The chondritic average (− 0.042 ± 0.043‰) is indistinguishable from basalts from Mars and Vesta, thus illustrating that partial mantle melting and crust extraction did not generate Fe isotope fractionation. Data for eucrites and basaltic shergottites are from [1], except for Millbillillie, Stanern and Zagami from [2], and Los Angeles and Sayh al Uhaymir 005 from [26].
melting and basaltic melt extraction cannot therefore be an explanation for the 0.1‰ heavier than chondrites δ57Fe signature of terrestrial mantle-derived rocks [1]. The same conclusion applies for the 0.2‰ heavier than chondrites δ57Fe signature measured in Apollo lunar samples [1], as discussed below. 3.2. The Moon case Evaluating whether lunar crust formation induced Fe isotope fractionation is not an easy task given the lack of samples from the lunar mantle (e.g., [27]). We are thus left with petrological considerations, comparisons with other planetary bodies and with theoretical estimates of Fe isotope fractionation. Weyer et al. [2] analyzed a series of lunar rocks and obtained a more scattered set of bulk-rock Fe isotopic data (from 0.042‰ in a “highland breccia” to 0.314‰ in a high-Ti basalt, with a mean δ57Fe of 0.156 ± 0.063‰, n = 12) than what we previously reported (from 0.111‰ for a quartz-normative basalt to 0.274‰ for a high-Ti basalt, with a mean δ57Fe of 0.206 ± 0.029‰, n = 14; [1]). This is despite their petrological range of igneous rock types was more restricted since they did not consider the norites, picritic basalts and anorthosites we analyzed [1]. Despite this isotopic scatter, they separated their samples in two groups, one made of high-Ti basalt with a heavy Fe isotope composition, and the rest, essentially consisting of “highland rocks”, low-Ti basalt and “lunar glasses” sharing an iron isotope composition similar to the five terrestrial basalt they analyzed, according to these authors [2].
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Again, the Weyer et al. [2] interpretations are biased by their sampling strategy. One first interesting feature is that although there is a good agreement between the Fe isotopic determinations from the two groups on well homogenised powders from terrestrial rock standards (see above), this is much less the case on the lunar samples analyzed by both Weyer et al. [2] and Poitrasson et al. [1]. The agreement is good for the KREEP basalt 15386 and quartz-normative basalt 15475, but it is much less so for olivine normative basalt 15555 and the two lunar ultramafic glasses 15426 and 74220. Lunar samples are precious and it is therefore not possible to work on powders produced out of large chunks of rocks. In our study, we worked on powders made of at least 1 g of rock [1], whereas Weyer et al. [2] made their powders from rock samples ten times smaller. Given the importance of Fe isotope mineral fractionation (Fig. 1) relative to the bulk-rock variations [1,23], one possible explanation of the scatter of Weyer et al. [2] data results from powders not representative enough of the bulk rocks. Samples 15426 and 74220 are made of glass beads of volcanic ultramafic glasses mixed with soil containing ejecta due to the meteoric bombardment of the lunar surface [28], potentially heavier isotopically than the volcanic glass beads [29]. We purified these glass beads under binocular lenses before chemical treatment and analysis, but it is not explained in the Weyer et al. [2] study whether they did so as well. If not, this can be an explanation of the higher δ57Fe values of Weyer et al. [2] on lunar glasses 15426 and 74220 compared to ours. Another critical aspect of our sampling is that we avoided rock samples chemically affected by the meteoritic bombardment as much as possible [1]. Instead, the lunar highland samples analyzed by Weyer et al. [2] were all impact melts mixed with clasts, contaminated chemically by meteoritic bombardment [28,30–33]. This is attested by their high siderophile contents. Rocks like 62235, for example, have Fe metal inclusions that have Co and Ni compositions within the meteoritic range [34]. We avoided analyzing these samples because besides the ejectas produced by the meteoritic bombardment, which are potentially heavy in their δ57Fe composition [29], the complementary fractions are likely to be isotopically lighter. Furthermore, the meteoritic iron contamination of these samples will shift the isotopic values of pristine lunar igneous rocks towards the lower Fe isotopic signatures of chondrites [3]. All these would therefore favour an isotopic scatter among samples and an overall lowering of the Fe isotope composition due to meteorite impacts (Fig. 4). We also found [1], like Weyer et al. [2], that high-Ti basalts tend to show a heavy Fe isotope composition.
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Fig. 4. Comparison of the lunar mean whole-rock δ57Fe from the data of Weyer et al. [2] and Poitrasson et al. [1], excluding the high-Ti basalts which tend to be isotopically heavy. This plot nevertheless illustrates that the mean δ57Fe lunar value computed using bulk rocks analyzed by [1], without high-Ti basalts, is indistinguishable from the lunar mean previously defined including high-Ti basalts [1]. This shows that these specific kinds of samples do not affect significantly the lunar mean bulk-rock Fe isotopic composition. Instead, the Weyer et al. [2] value (δ57Fe = 0.108 ± 0.042‰; n = 9) is lower than the mean Moon δ57Fe previously defined [1] because it comprises a large number of samples heavily affected by the meteoritic bombardment.
However, our larger petrological range of rock types, chemically unaffected by the meteoric bombardment, also shows that bulk anorthosites, quartz-normative low-Ti basalts and KREEP basalts can be isotopically as heavy [1]. There are therefore no reasons to put the high-Ti basalt into a separate group on the basis of their Fe isotope signatures as done by [2]. Indeed, the mean of bulk igneous lunar rocks from [1], excluding high-Ti basalts (δ57Fe of 0.185 ± 0.038‰, n = 9) is indistinguishable with the mean originally published by [1] that includes these Ti-rich rocks (Fig. 4). Thus, taking all the lunar samples analyzed by Weyer et al. [2] representative of bulk rocks (i.e., excluding the soil sample 10084 and “lunar volcanic glass” clods 74220 and 15426), the computed lunar mean (δ57Fe of 0.156 ± 0.063‰, n = 12) is undistinguishable from that published earlier (δ57Fe of 0.206 ± 0.029‰, n = 14; [1]). Their lower value and larger uncertainty probably result largely from the effect of meteoritic bombardment on many of the samples selected by these authors for their study. Weyer et al. [2] attempted to explain a heavier Fe isotope composition of high-Ti basalts in terms of the involvement of ilmenite in the petrogenesis of these rocks. However, as previously discussed [1], the Polyakov and Mineev [10] theoretical estimates suggest that ilmenite is one of the high temperature Fe-bearing igneous mineral that fractionates the less iron isotopes. There is therefore no justification to assume that this phase played a specific role explaining the Fe isotope composition of high-Ti
basalt that in fact is, as shown above, not as different as implied by Weyer et al. [2] compared to other bulk lunar igneous samples. If, for example, massive ilmenite fractional crystallization occurred during magma differentiation, this would be accompanied by a lowering of Ti and little impact on, or in an extreme case a slight increase in δ57Fe of the magma. Conversely, assimilation of ilmenite in the melt would result in a minimal effect to slight lowering of δ57Fe values of the magma, but a significant increase in its Ti content. In both cases, the resulting hypothetic melt would not correspond to an isotopically heavy, Ti-rich basalt. Hence, there is currently no evidence to hypothesize that the production of a lunar crust fractionated Fe isotopes and that the heavy isotopic composition of lunar samples result from planetary differentiation. 4. Conclusions Knowing whether planetary differentiation generates Fe isotope fractionation requires first good estimates of the main planetary reservoir compositions. The Weyer et al. [2] study shows that inappropriate sampling may lead to exaggerated isotopic variability (e.g., in their lunar rocks) and biased estimates of bulk planetary reservoirs (i.e., the Earth's mantle). It is also important to take fully into account data from all previously published studies (e.g., [6]) and the available theoretical estimates of high temperature Fe isotope fractionation between minerals [10].
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The large variability of the Fe isotope composition of bulk mantle rocks [20] remains puzzling, particularly when compared to the isotopic homogeneity of mantlederived rocks [1,15,21]. Therefore, it is first necessary to understand the cause of this variability before a bulk mantle Fe isotope composition can be precisely established. Experimental studies of peridotite partial melting are required to monitor the effect of this process on the Fe isotopic signature on magmas. So far, there are no reasons to consider that the Earth's mantle has a Fe isotope composition different from that of igneous crustal rocks (see also [21]). The current Fe isotopic database, including notably the fact that Martian basalts and eucrites have a Fe isotope composition undistinguishable from that of the building block of planets, chondrites [1,3], makes untenable the suggestion of Weyer et al. [2] that planetary differentiation, and particularly the production of a crust, can generate Fe isotope fractionation. The same conclusion is also valid for the mantle–core differentiation if metal and silicate remained in equilibrium [3]. It is therefore re-emphasized that the ∼ 0.1‰ heavier Fe isotope composition of the mafic silicate Earth, and especially the ∼ 0.2‰ heavier iron isotope signature of samples from the Moon, relative to chondrites, SNC and HED meteorites remain best explained by contrasted planetary accretion mechanisms [1,3]. This interpretation remains the most logical when all current available data are taken into account. It is also promising since it implies that Fe isotopes can trace specific processes not easily tracked down by other geochemical tools. Acknowledgements Thierry Aigouy is acknowledged for his help with the SEM characterization of pallasite separates. Guillaume Delpech is thanked for a careful reading of the manuscript, an anonymous referee for a critical and helpful review and Rick Carlson for his editorial handling. References [1] F. Poitrasson, A.N. Halliday, D.C. Lee, S. Levasseur, N. Teutsch, Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms, Earth Planet. Sci. Lett. 223 (2004) 253–266. [2] S. Weyer, A.D. Anbar, G.P. Brey, C. Munker, K. Mezger, A.B. Woodland, Iron isotope fractionation during planetary differentiation, Earth Planet. Sci. Lett. 240 (2005) 251–264. [3] F. Poitrasson, S. Levasseur, N. Teutsch, Significance of iron isotope mineral fractionation in pallasites and iron meteorites for the core–mantle differentiation of terrestrial planets, Earth Planet. Sci. Lett. 234 (2005) 151–164.
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