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Magmatic water in the martian meteorite Nakhla
Earth and Planetary Science Letters 359-360 (2012) 84–92
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Magmatic water in the martian meteorite Nakhla L.J. Hallis a,b,c,n, G.J. Taylor a,c, K. Nagashima a, G.R. Huss a,c a
Hawai’i Institute of Geophysics and Planetology, Pacific Ocean Science and Technology (POST) Building, University of Hawai’i, 1680 East-West Road, Honolulu, HI 96822, United States Institute for Astronomy, University of Hawai’i, 2680 Woodlawn Drive, Honolulu, HI 96822-1839, United States c UH, NAI Astrobiology Program, United States b
a r t i c l e i n f o
a b s t r a c t
Article history: Received 7 March 2012 Received in revised form 8 September 2012 Accepted 27 September 2012 Editor: T. Elliot Available online 8 November 2012
Mars does not recycle crustal materials via plate tectonics. For this reason the magmatic water reservoir of the martian mantle has not been affected by surface processes, and the deuterium/hydrogen (D/H) ratio of this water should represent the original primordial martian value. Following this logic, hydrous primary igneous minerals on the martian surface should also carry this primordial D/H ratio, assuming no assimilation of martian atmospheric water during crystallization and no major hydrogen fractionation during melt degassing. Hydrous primary igneous minerals, such as apatite and amphibole, are present in martian meteorites here on Earth. Providing these minerals have not been affected by terrestrial weathering, martian atmospheric water, or shock processes after crystallization, they should contain a good approximation of the primordial martian D/H ratio. As Nakhla was seen to fall in the Egyptian desert in 1911, terrestrial contamination is minimized in this meteorite. The nakhlites are also among the least shocked of the martian meteorites. Therefore, apatite within Nakhla could contain primordial martian hydrogen isotope ratios. We produced in-situ measurements of the D/H ratios in Nakhla apatite grains, using a Cameca ims 1280 ion-microprobe. Our measurements produced D/H values in Nakhla apatite similar to terrestrial values, despite strong evidence that our samples were not significantly contaminated by terrestrial hydrogen. These results suggest that water trapped in the martian mantle has a similar D/H to that of the Earth. Therefore, the water of these two planets may have originated from the same source material. The D/H ratios of the carbonaceous chondrite meteorites, and the Jupiter-family comet 103P/ Hartley 2, are similar to the D/H of the two planets, making both these primitive inner solar system materials strong candidates for the source of the terrestrial planets water. These results support recent dynamical models of the formation of the solar system, which suggest material in the inner solar system was homogenized by the migration of Jupiter. & 2012 Elsevier B.V. All rights reserved.
Keywords: martian meteorites hydrogen isotopes terrestrial planet formation
1. Introduction Understanding the sources and delivery mechanisms of water to the Earth and the other terrestrial planets allows for the validation of planetary accretion models. A key parameter in determining the source(s) of terrestrial planetary water is hydrogen isotope composition. This is reasonably well established for the Earth’s water reservoirs. Here we report new measurements that attempt to establish the hydrogen isotopic composition of the martian interior. There are currently 48 martian meteorites available for study. Although these rocks were not sampled directly from the surface n Corresponding author at: Hawai’i Institute of Geophysics and Planetology, Pacific Ocean Science and Technology (POST) Building, University of Hawai’i, 1680 East-West Road, Honolulu, HI 96822, United States. E-mail address: [email protected] (L.J. Hallis).
0012-821X/$ - see front matter & 2012 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.epsl.2012.09.049
of Mars, their martian origin can be deduced by comparing the chemical composition of trapped gas bubbles with the composition of the martian atmosphere, as measured by the Viking landers (Bogard and Johnson, 1983). Despite the fact that there are various different types of martian meteorite, all are primarily igneous rocks from basaltic lava flows or shallow-level subsurface intrusions. Certain igneous primary minerals within these rocks, for example amphibole (potassic-chlorohastingsite) and apatite (Cl-rich), have water locked into their crystal structure. As the planetary accretion process does not appear to fractionate hydrogen isotopes (Le´cuyer et al., 1998), and as Mars does not recycle crustal materials via plate tectonics, the D/H ratio of water in these minerals could represent the primordial ratio of Mars’ water. If so, this ratio could be used to assess the origin of martian water. However, there are various processes that may have affected the D/H ratio of water within these minerals, including contamination via terrestrial weathering and/or martian atmospheric
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water, hydrogen fractionation during magmatic degassing, and shock processes after crystallization. The ratio of deuterium (2H or D) to hydrogen (1H) in water is commonly quoted relative to Vienna Standard Mean Ocean Water (VSMOW), and is given as dD¼ [((D/H)sample/(D/H)VSMOW) 1] 1000, with units of per ml (%). Hence, VSMOW has a dD value of 0%. On Earth, ice and rocks have negative dD values, with meteoric water varying from þ130% (some rainfall) to 480% (ice sheets), and the oceans at 0% (Le´cuyer et al., 1998 and references therein; Hoefs, 2004). Measurements of the Earth’s mantle have produced a range of values, from 140 to þ60% (e.g., Boettcher and O’Neil 1980; Michael, 1988; Ahrens, 1989; Deloule et al., 1991; Bell and Rossman, 1992; Thompson, 1992; Graham et al., 1994; Jambon, 1994; Wagner et al., 1996; Xia et al., 2002). In contrast, Mars’ atmospheric hydrogen has a positive dD value. Approximately 4 billion years ago Mars’ liquid core is believed to have solidified, causing the planet to lose its global magnetic field. This, in turn, initiated the loss of the martian atmosphere to space. As the lighter hydrogen isotope was preferentially lost, the atmosphere became deuterium enriched—the current martian atmospheric dD value is approximately þ4200% (Bjoraker et al., 1989). This enrichment was subsequently incorporated into secondary alteration minerals formed via weathering, such as the Fe-rich clay-like deposits and carbonates found in certain martian meteorites (Boctor et al., 2003). Amphibole and apatite are the primary igneous minerals, and hence they contain water from the martian interior. Providing there has been no subsequent contamination, their dD should reflect that of primordial Mars, without any atmospheric component. However, as reflected by previous hydrogen isotopic studies of both martian primary igneous and alteration phases (e.g., Watson et al., 1994; Leshin, 2000; Deloule, 2002; Gillet et al., 2002; Boctor et al., 2003; Greenwood et al., 2008) both martian and terrestrial surface fluids are potential sources of contamination.
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measured with a 4 nA primary ion beam (Feb 2012). As Nakhla apatites are rare and mostly very small, we were unable to collect a second set of results for these apatites with the higher beam current. However, although the hydrogen background is higher for the 2 nA analyses, it is not close to the level of counts obtained from the apatite grains [SOM]. Therefore, analyzing Nakhla apatite grains with a 4 nA primary beam current was not necessary. For Fe-rich clay-like veins, H, D, and 30Si were measured with 80 pA primary beam. The secondary ion mass spectrometer was operated at 10 keV with a 50 eV energy window. Isotopes were detected using an electron multiplier. The mass resolving power was 1900, sufficient to separate any interfering molecules. A normal-incidence electron flood gun was used for charge compensation. Prior to each measurement, the 2 nA primary beam (4 nA for Shergotty apatites) was rastered over a 25 25 mm2 area for 200 s, to remove the carbon coat and any surface contamination. During this rastering process the distribution of H ions in the region was displayed (although no image was permanently saved), providing the opportunity to avoid any H-rich cracks which may contain terrestrial contamination. The raster was then reduced to a 15 15 mm2 analysis area. To further reduce the contribution of unwanted hydrogen to our signal, we used the electronic gate to exclude the ions from the outer 50% of the 15 15 mm2 sputtered area, thus eliminating contributions from the crater walls and from hydrogen creeping along the surface of the collector. Therefore, our final analysis spot size was 8 8 mm2. The data were collected for 40 cycles, with H measured for 3 s, D for 40 s, and 18O (or 30Si for Fe-rich clay-like veins) for 2 s in each cycle. The beam was blanked for the first 10 cycles of the measurement, as well as for the last five cycles, to measure the background hydrogen and deuterium counts per second.
2.2. Data analysis 2. Methods 2.1. Measurement protocol Our sample set included eight UH polished thin-sections of Nakhla (20a–h), newly prepared from the same cm-sized rock chip (parent 20) without water (using polishing oil and petroleum ether to clean), and one thin-section of Shergotty (UNM 410), prepared with water prior to 1986 at the University of New Mexico [see SOM in the Appendix for petrographic descriptions]. The Nakhla samples were placed in a vacuum oven at 60 1C as soon as they were prepared to minimize atmospheric contamination, and were only removed for analysis. Shergotty UNM 410 was cleaned with methanol, placed in the vacuum oven for three days, re-carbon coated, and replaced in the vacuum oven one week prior to analysis. We utilized the JEOL JSM-5900LV scanning electron microscope at the University of Hawai’i (UH) to produce backscatter electron images and elemental X-ray images at 65 magnification (20 mm pixel size) for each thin-section, in order to locate areas of interest. These areas were subsequently imaged at higher resolutions, with various pixel sizes (approximately 1–5 mm), to pick out the individual apatite grains and Fe-rich clay-like veins for ion microprobe analysis. The areas analyzed for D/H were also imaged after analysis, to verify the positions of the sputtered regions. Deuterium and hydrogen isotopic compositions were analyzed in-situ with the UH Cameca ims 1280 ion microprobe. A 2 nA focused Cs þ primary ion beam was used to produce negative ions of H, D, and 18O for Nakhla apatites (June 2011), following the protocol of Boctor et al. (2003). However, the background counts per second (cps) were subsequently found to slightly decrease with a higher beam current, so the Shergotty apatite grains were
The data from the ion probe are reported in counts per second (cps). However, when we use the electronic gate, the beam is deflected away from the detector for a significant portion of each second, so the reported count rates are too low to make an appropriate deadtime correction. Therefore, we first correct the measured count rates to account for the fraction of time that the beam is out of the detector. The data are then corrected for deadtime. The isotope data were subsequently corrected for background, which was measured during the first 10 and last five cycles of each measurement as described above. The background signal was routinely 3000–7000 cps for Nakhla apatites, and 400–600 cps for Shergotty apatites (because of the increased beam current). The total hydrogen signal with the beam on was 200,000–1,200,000 cps for Nakhla apatites, and 800,000–1,000,000 cps for Shergotty apatites [Appendix (SOM)]. The background counts were subtracted from the measured isotope signal, collected between cycles 11 and 35 while the beam was on. The counting-statistical uncertainty of the background signal was propagated. Apatite data were corrected for instrumental fractionation using two terrestrial apatite standards (Crystal Lode ap005 and Russia ap018). The water contents and dD values of these apatites were previously determined by McCubbin et al. (2010, 2012) [SOM]. While these previous publications did not determine the dD value of Durango ap003, the water content of this standard was established. Therefore, as these three apatite standards (Durango ap003, Crystal Lode ap005 and Russia ap018) contain varying water contents, calibration curves were produced to estimate the water content of the unknowns (Fig. 1). The H2O content of each unknown apatite was estimated based on its 1H /18O ratio, and the slope of the
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Fig. 1. 1H/18O vs. H2O wt% calibration lines for the Nakhla apatite (a) and Shergotty apatite (b) analyses. These curves are based on the measurements of three terrestrial apatite standards with known water contents at 2 nA (a) and 4 nA (b).
relevant 1H/18O vs. H2O wt% standard calibration line (Fig. 1), using the following equation: H2O wt%¼(1H /18O ) line slope
(1)
The H2O detection limit was estimated in a similar way, via the measurement of 1H /18O ratios in a nominally anhydrous San Carlos olivine standard. Detection limits are consistently 30 ppm H2O for Nakhla (2 nA), and 10 ppm for Shergotty (4 nA). The difference in these limits is caused by a reduction in instrumental hydrogen background during the Shergotty analyses, because of the higher beam current. As these detection limits are so low, the calibration lines for both primary beam currents were forced through the origin, following the methodology of McCubbin et al. (2010, 2012). The reproducibility of the apatite water contents, and D/H ratios, were calculated based on the subtraction of the true H2O or D/H value from the estimated H2O or D/H value of each apatite standard (error¼2 standard deviation (estimated-true)). Apatite water content reproducibility is 0.02 wt% H2O for Nakhla, and 0.01 wt% H2O for Shergotty. The 2s uncertainties of each apatite H2O content and D/H ratio were subsequently calculated using the following equation: ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi r ffi 2 2s error ¼ 2 measurement error2 þ standard reproducibility ð2Þ The H2O contents of the Nakhla Fe-rich clay-like veins cannot accurately be estimated in this way, as these veins are made up of more than one mineralogical phase. Fe-rich clay-like vein data were corrected for instrumental mass fractionation using a serpentine standard (UH 180C 18) from Conical Seamont, Mariana Arc, south of Japan (Alt and Shanks, 2006).
3. Results Apatite grains in the martian meteorite Nakhla (Fig. 2) show
dD values of between 111 and þ155% (Table 1). These values are similar to those of terrestrial meteoric water and the Earth’s mantle, excepting that the range stretches to slightly more positive values. The samples studied were sourced from a group of Nakhla specimens cataloged at the British Natural History Museum in 1913—2 yr after the Nakhla meteorite shower over Egypt [see SOM for petrographic descriptions in the Appendix]. It is unclear whether the specimens spent this 2 yr period in the museum waiting to be cataloged, or in the field exposed to weathering. Therefore, to test the extent of terrestrial hydrogen
contamination in our samples, we measured the dD value of coexisting pre-terrestrial Fe-rich clay-like veins. These clay-like vein assemblages, or ‘iddingsite’ veins formed as a result of martian surficial aqueous alteration of the crystallized nakhlite basalt (e.g., Ashworth and Hutchison, 1975; Reid and Bunch, 1975; Gooding et al., 1991; Bridges et al., 2001; Gillet et al., 2001, 2002; Treiman, 2005). They are reported to consist of mixtures of hydrous amorphous silicate gel, siderite, gypsum and carbonate (Changela and Bridges, 2011), possibly with some smectite clay (Gooding et al., 1991) [see also SOM in the Appendix]. The phases within these veins, especially smectite and hydrous amorphous silicate gel, contain only loosely bound water that would be susceptible to hydrogen exchange with the surrounding atmosphere. The original dD value of these alteration phases on Mars would depend on the dD value of the alterationforming fluid. However, even if this fluid had a low dD value (which may possibly be the case if some martian hydrothermal fluids exist/existed having had no contact with atmospheric hydrogen), subsequent exposure of the ‘iddingsite’ veins to the martian atmosphere would have resulted in relatively rapid hydrogen exchange. Therefore, water within these veins should have been enriched in deuterium at the time the nakhlites were blasted from Mars. Hydrogen isotope fractionation between the martian atmosphere and the specific alteration vein phases would affect the level of this enrichment. However, the fractionation factor between montmorillonite clay and water is 0.94 at sedimentary temperatures on Earth (Savin and Epstein, 1970), resulting in only a slight enrichment of the lighter hydrogen isotope in the solid phase. This same effect is seen in other clays, amphiboles, and micas (see Hallis et al., 2012 for further discussion). As explained above, the layered and commonly amorphous structure of the clay-like ‘iddingsite’ vein assemblages results in non-structural and/or loosely bound water molecules. Loosely bound water leads to enhanced hydrogen exchange with the surrounding atmosphere, compared to purely crystalline minerals (such as apatite) where water is locked into the crystal structure. It therefore follows that if these alteration assemblages still retain their martian D enrichment, their hydrogen isotopes have not completely exchanged with terrestrial atmospheric hydrogen— terrestrial contamination in the samples is minimal. If this very easily exchangeable martian hydrogen has not been completely replaced, neither has the more structurally bound martian hydrogen in the apatite grains. Our measurements from Nakhla samples 20c and 20e show dD values between þ1170 and 105% (Table 1) within these assemblages, indicating incomplete hydrogen isotope exchange with the terrestrial atmosphere. The abundance of 1H (cps) in these veins gives an indication of their
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Fig. 2. Backscattered electron (A, C, E, G), and secondary electron (B, D, F, H), images of Nakhla apatites 20b1 (A and B) 20b2 (C and D), 20b3 (E and F) and 20c1 (G and H). Backscattered electron images show apatite grains (ap) within the mesostasis (meso) of Nakhla, associated with clinopyroxene (cpx) and olivine (olv) phenocrysts, along with other more minor igneous phases, including the silica polymorph cristobalite (crist), the iron sulfide pyrrhotite (FeS), feldspar (feld) and Fe-rich clinopyroxene (Fecpx). Secondary electron images highlight the extent of the SIMS pre-sputtered raster pits on each apatite grain (the extent each grain is approximated by the white shape). The raster pits are 25 25 mm2, but D/H measurements only include counts from the central 8 8 mm2 (approximated by the black squares).
relative H2O contents, and these values also vary widely—from 150,000 to 880,000. This wide variation in H2O content reflects the heterogeneous mixture of hydrous and anhydrous silicates and non-silicates in these veins. Compositional variation is also suggested by the variation in 1H/30Si ratios (Table 1). When these values are compared to dD for individual analyses, there is generally an inversely proportional relationship—as 1H cps or 1H/30Si ratio increases, dD decreases. This relationship is indicative of different amounts of terrestrial atmospheric contamination, with the most
H-rich and deuterium depleted veins having absorbed the most terrestrial hydrogen. A wide variation in ‘iddingsite’ vein dD values has also been reported in other nakhlite samples, and appears to be strongly dependant on the sample and parent meteorite history (Gillet et al., 2002; Hallis et al., 2012). To further test our methodology, and the reliability of our Nakhla apatite results, we measured the D/H ratio of apatite grains in Shergotty (Fig. 3). Our measurements show dD values of 2953%–3931% (Table 1). The lowest of these values was
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Table 1 dD (%) and H2O contents of Nakhla and Shergotty apatite grains and standards, and dD (%) of Nakhla Fe-rich claylike veins and standards. The calculated H2O content of anhydrous minerals indicates the detection limits. 1
Sample
dD (%)
2r
Apatites (2 nA) Nakhla 20b apatite 1 Nakhla 20b apatite 1 (2) Nakhla 20b apatite 2 Nakhla 20c apatite 1 Durango apatite std 1 Durango apatite std 2 Crystal Lode apatite std 1 Crystal Lode apatite std 2 Crystal Lode apatite std 4 Russia apatite std 1 Russia apatite std 2 Russia apatite std 3
188 78 19 151 4 71 60 36 47 107 91 128
62 62 62 62 3 3 4 6 6 8 8 8
0.43 0.42 0.09 0.58 0.04 0.05 0.33 0.33 0.33 0.19 0.19 0.19
Apatites (4 nA) Shergotty apatite 1 Shergotty apatite 1 (2) Shergotty apatite 2 Shergotty apatite 8 Shergotty apatite 9 Durango apatite std 1 Durango apatite std 2 Crystal Lode apatite std 1 Crystal Lode apatite std 2 Crystal Lode apatite std 3 Russia apatite std 1 Russia apatite std 2 Russia apatite std 3
3563 3931 3267 2953 3516 141 108 99 113 101 145 104 109
12 13 11 10 11 10 11 4 3 3 4 5 5
0.63 0.55 0.66 0.71 0.67 0.04 0.04 0.37 0.39 0.38 0.21 0.20 0.20
Fe-rich clay-like veins (80 pA) Nakhla 20c Fe-vein 1 852 Nakhla 20c Fe-vein 2 6 Nakhla 20c Fe-vein 3 1165 Nakhla 20c Fe-vein 4 702 Nakhla 20c Fe-vein 5 16 Nakhla 20c Fe-vein 6 66 Nakhla 20c Fe-vein 7 98 Nakhla 20c Fe-vein 8 42 Nakhla 20e Fe-vein 1 110 Nakhla 20e Fe-vein 2 108 Nakhla 20e Fe-vein 3 229 Nakhla 20e Fe-vein 4 237 Nakhla 20e Fe-vein 5 35 Serpentine std 1 53 Serpentine std 2 36 Serpentine std 3 28 Serpentine std 4 0 Serpentine std 5 54 Serpentine std 6 42 Serpentine std 7 53
85 77 86 96 80 75 79 76 77 78 82 77 78 77 75 75 77 76 75 76
3.18 21.43 91.74 7.58 22.83 42.41 38.95 49.38 25.88 27.29 24.93 13.08 3.25 38.25 33.46 44.44 51.81 51.24 50.43 37.54
H/monitor isotopea
2r
H2O (wt%)
2r
0.0001 0.0001 0.00003 0.0001 0.00002 0.00002 0.00006 0.00006 0.00006 0.00004 0.00004 0.00004
0.47 0.46 0.10 0.64 0.03 0.03 0.37 0.37 0.37 0.20 0.20 0.20
0.02 0.02 0.02 0.02 0.01 0.01 0.02 0.02 0.02 0.02 0.02 0.02
0.0002 0.0002 0.0002 0.0002 0.0002 0.00005 0.00004 0.00018 0.00016 0.00015 0.00011 0.00010 0.00005
0.62 0.54 0.65 0.70 0.66 0.04 0.04 0.36 0.38 0.37 0.21 0.19 0.20
0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01
38.46 38.47 38.99 38.46 38.47 38.48 38.51 38.48 38.50 38.47 38.48 38.47 38.46 38.47 38.47 38.50 38.52 38.49 38.48 38.48
Anhydrous min (2 nA) Nakhla 20b augite San Carlos olivine 1 San Carlos olivine 2 San Carlos olivine 3
0.0088 0.0029 0.0031 0.0027
0.00002 0.00016 0.00017 0.00013
0.0097 0.0032 0.0034 0.0030
0.04076 0.04076 0.04076 0.04076
Anhydrous min (4 nA) Shergotty augite 1 Shergotty olivine 1 Shergotty olivine 2 San Carlos olivine 1 San Carlos olivine 2
0.0329 0.0013 0.0088 0.0012 0.0011
0.00002 0.00001 0.00002 0.00001 0.00001
0.0323 0.0012 0.0086 0.0012 0.0011
0.01414 0.01414 0.01414 0.01414 0.01414
a Monitor isotope¼ 18O for apatite and anhydrous mineral measurements, and 30Si for Fe-rich clay-like vein measurements.
produced from an analysis which may have included some terrestrial contamination from a crack within the apatite grain (Fig. 3). All other analyses appear to be free from this problem, at least at the visible scale. These measurements can be directly
compared with those of Greenwood et al. (2008), who measured the dD of apatites in Shergotty and Los Angeles (another shergottite meteorite). Our Shergotty apatite dD values are comparable with the apatite dD values of Los Angeles, but are 650% lower
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Fig. 3. Secondary electron images of Shergotty UNM 410 SIMS analysis pits. Apatite 1 contains two analysis pits (A and B), whereas apatites 2 (C and D), 8 (E and F) and 9 (G and H) contain only one. A test pit was made in the olivine grain next to apatite 9 (G). Lower magnification images (A, C, E, G) outline the general extent of each apatite grain (within the white shapes), and higher magnification images (B, D, F, H) show the pits in detail. As in Fig. 2, the central 8 8 mm2 measurement areas are approximated by the black squares. Ap¼ apatite, olv ¼olivine.
than the two previously measured apatite dD values of Shergotty (Greenwood et al., 2008). This difference may be caused by our analyses of a larger area on each apatite grain (8 8 mm2 rather than a 2 2 mm2), causing an increased risk of terrestrial
contamination from micro-cracks. Previously measured apatite and amphibole dD measurements from other martian meteorites are comparable to, or significantly lower than, our Shergotty values (Watson et al., 1994; Leshin, 2000; Boctor et al., 2003).
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4. Discussion 4.1. Possibility and effect of different sources of hydrogen isotope contamination The relationship of our measured apatite dD values with the original primordial dD value of Mars is dependent on three factors. The first factor is the amount of martian atmospheric contamination and degassing the apatite grains were exposed to during their crystallization and subsequent 1.3 billion year residence on Mars’ surface (Stauffer, 1962). The fact that Nakhla’s apatite dD values are much lower than some of its clay-like alteration assemblage dD values indicates that the majority of the water in these apatites is not sourced from martian atmospheric water, either via assimilation of crustal material or interaction with martian groundwater during crystallization. In addition, it is clear that Nakhla’s apatite grains did not completely exchange their hydrogen with the martian atmosphere after crystallization—in this case, again, more positive dD values would have been measured. However, some minor martian atmospheric contamination remains a possibility, as both Nakhla 20b apatite 1 and 20c apatite 1 show slightly positive dD values, above the normal terrestrial range (Table 1; Fig. 4). There is also evidence that the nakhlite melt probably degassed as it cooled, and that hydrogen isotope fractionation occurred during this process. McCubbin et al. (2012) attempted to determine the water content of the martian mantle by measuring that of the shergottite apatites. These measurements indicate that the shergottite magmas underwent magmatic degassing, because their water content as calculated from apatites is greater than the measured bulk-rock water content. Prior analysis of an apatite grain within the shergottite meteorite Los Angeles established the presence of isotopic zoning, from a relatively deuterium-poor core (dD¼4120%) to a deuterium-rich rim (dD¼4348%) (Greenwood et al., 2007, 2008). This zoning suggests that as the magma degassed, apatites became D-enriched, hence dD increased. Deuterium enrichment during degassing agrees with the general rule that during isotopic diffusion the lighter isotope
Fig. 4. Comparison of the dD values of our data (black symbols) with relevant data from previous studies (black and gray lines) (Watson et al., 1994; Leshin, 2000; Boctor et al., 2003; Greenwood et al., 2008). The high dD values of amphibole and apatite grains within Shergotty and the other previously measured meteorites could be explained by the greater shock pressures these rocks were exposed to relative to Nakhla, or by greater assimilation with martian atmospheric groundwater. Fe-rich clay-like veins in Nakhla show dD similar to ALH 84001 carbonates, proving that the low apatite dD in this meteorite cannot be the result of terrestrial hydrogen isotope exchange.
becomes enriched in the gas phase (as it has a weaker bond than the heavier isotope). Therefore, degassing of the melt, and of the apatites after crystallization, should result in D enrichment in the nakhlite rocks/apatites. Hence, in addition to possible martian atmospheric contamination, magmatic degassing may have also increased the deuterium content in Nakhla’s apatites relative to the original melt value. The second factor is the effect of shock on the apatites as they were blasted from the martian surface via meteorite impact. The high shock pressures involved in impact events can volatilize hydrogen in hydrated minerals, with 1H being more susceptible to this process than D. Volatilization has been reported to increase dD by up to 87% in lab-based experiments (Minitti et al., 2008). In addition, volatilization is accompanied by shock-induced implantation of water from the surrounding atmosphere. Therefore, meteorites that were heavily shocked on the surface of Mars might be expected to contain hydrated primary minerals enriched in D compared to the equivalent minerals in less shocked meteorites. These experimental results are consistent with our measurements of dD in Shergotty apatite grains, as well as previously reported dD values from the igneous hydrated minerals of Shergotty and other martian meteorites (Watson et al., 1994; Leshin, 2000; Boctor et al., 2003; Greenwood et al., 2008) (Fig. 4). These previous studies involved meteorites that were heavily shocked (31–49 GPa, possibly as high as 90 GPa for Shergotty—Greshake and Langenhorst, 1997; Sharp et al., 1999; El Goresy et al., 2000; Bell, 2007), and reported apatite dD values between þ 751 and þ4606% (Watson et al., 1994; Leshin, 2000; Boctor et al., 2003; Greenwood et al., 2008). In contrast, Nakhla was probably less severely shocked (20 GPa—Greshake, 1998), therefore is less likely to have acquired D-rich water from shock. Based on their measurements of various martian meteorites, Boctor et al. (2003) also reached the conclusion that increased shock increases the dD value in primary igneous phases. All of the above processes—martian atmospheric contamination, magmatic degassing, and the twin processes of shockinduced hydrogen volatilization and martian atmospheric water implantation during shock—would potentially increase dD in the hydrated minerals of martian meteorites. Therefore, we could assume a maximum value of 111 762% for the hydrogen isotopic composition of the martian mantle, as this is the lowest dD value in our apatite dataset, hence the least affected by these processes. However, before assuming this value is generally correct, we need to take into account the third factor—terrestrial contamination. Despite the fact that Nakhla is a fall meteorite, and that our samples were prepared without the use of water, terrestrial contamination may still be an issue in the apatite grains. As discussed above, Fe-rich clay-like veins within these samples do show some high dD values, indicating that this easily altered material has not completely exchanged its hydrogen with the terrestrial atmosphere. Micro-cracks in the apatite grains have proved to be a source of terrestrial contamination, as illustrated by the low dD value produced when a visible crack was included in one of our Shergotty analyses (Table 1). Greenwood et al. (2008) also recorded the effect of these cracks. The 650% difference between our Shergotty apatite dD values and those of Greenwood et al. (2008) may be the result of increased microcrack contamination, caused by our larger analyses areas. Therefore, it may be that our maximum martian mantle value of 111% is too low. However, it should be expected that terrestrial micro-crack contamination in both our Shergotty data, and that of Greenwood et al. (2008), would be greater than in our Nakhla apatite data, as both Shergotty samples were prepared using water. Previous dD measurements of Shergotty hydrous igneous phases show a wide variation (Watson et al., 1994; Boctor et al., 2003; Greenwood et al., 2008).
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Table 2 Previously measured major terrestrial, martian and cometary reservoir dD reservoirs. Known dD reservoirs
dD (%)
Terrestrial Oceansa Terrestrial Ice Sheetsa Terrestrial Meteoric Waterb Terrestrial Mantlec Comet Hartley 2dn Comet P/Halleye Comet Hyakutakef Comet Hale-Boppg Martian Atmosphereh
0 300 to 400 40 to þ130 140 to þ 60 6 888 813 1063 4200
2s(%)
300 92 625 1000
a
Le´cuyer et al. (1998). Hoefs (2004). c Boettcher and O’Neil (1980), Michael (1988), Ahrens (1989), Deloule et al. (1991), Bell and Rossman (1992), Thompson (1992), Graham et al. (1994), Jambon (1994), Wagner et al. (1996) and Xia et al. (2002). d Hartogh et al. (2011). e Eberhardt et al. (1995). f Bockelee-Morvan et al. (1998). g Meier et al. (1998). h Bjoraker et al. (1989). n Comet Hartley 2 is a Jupiter family comet, whereas P/Halley, Hyakutake and Hale-Bopp are all Oort cloud comets. b
These differences may indicate heterogeneity between different apatite grains in Shergotty, possibly explaining the difference between our dD values for this meteorite and those of Greenwood et al. (2008). Taking into account the above information, the amount of Nakhla apatite dD alteration caused by terrestrial and martian atmospheric hydrogen contamination, shock, and magmatic degassing is apparently small. We can therefore say with some certainty that our Nakhla apatite data reflect the best estimate of the true martian mantle dD value to date, and that this value is similar to that of the Earth. 4.2. The source of martian water The similarity of our proposed martian mantle dD value to that of the terrestrial oceans, rocks and mantle could suggest that Earth and Mars, and possibly the other terrestrial planets, accreted water from the same source, or from more than one source with similar D/H ratios. Carbonaceous chondrites have been considered a likely source because of the similarity in their D/H to terrestrial ocean water (Robert et al., 2000). Jupiter-family comets may also be a source; comet 103P/Hartley 2 has a D/H ratio indistinguishable from that of Earth (Hartogh et al., 2011). A significant contribution from the long-period comets of the Oort Cloud seems to be ruled out by the elevated D/H ratios of their water (dD4 þ800%, Table 2; Eberhardt et al., 1995; BockeleeMorvan et al., 1998; Meier et al., 1998). Therefore, the inner solar system materials may all have similar D/H ratios, whereas materials from further out in the solar system appear to be D enriched. Recent dynamical models suggest that the inward and outward migration of Jupiter caused extensive mixing of material in the inner solar system, while leaving the outer solar system relatively unmixed (Walsh et al., 2011). These models could explain the difference between inner and outer solar system hydrogen isotope reservoirs, and the similarity between the D/H ratios of Mars and Earth.
Acknowledgments This material is based upon work supported by the National Aeronautics and Space Administration through the NASA
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Astrobiology Institute under Cooperative Agreement no. NNA09DA77A issued through the Office of Space Science. We also thank the NASA Johnson Space Center for allocation of Nakhla (20), and JoAnn Sinton at the University of Hawaii for producing the thinsections. Eric Hellebrand is thanked for his assistance with EMP analyses. Prof. F. McCubbin is thanked for supplying the terrestrial apatite standards. The comments and suggestions of Prof. S. Krot, Dr. Shoichi Itoh and Prof. Tim Elliott, as well as two anonymous reviewers, were invaluable in improving this manuscript.
Appendix A. Supporting information Supplementary data associated with this article can be found in the online version at http://dx.doi.org/10.1016/j.epsl.2012.09.049.
References Ashworth, J.R., Hutchison, R., 1975. Water in non-carbonaceous stony meteorites. Nature 256, 714–715. Ahrens, T.J., 1989. Water storage in the upper mantle. Nature (London) 342, 122–123. Alt, J.C., Shanks, W.C., 2006. Stable isotope compositions of serpentinite seamounts in the Mariana forearc: serpentinization processes, fluid sources and sulfur metasomatism. Earth Planet. Sci. Lett. 242, 272–285. Bell, M.S., 2007. Experimental shock decomposition of siderite and the origin of magnetite in Martian meteorite ALH 84001. Meteorit. Planet. Sci. 42, 935–949. Bell, D.R., Rossman, G.R., 1992. Water in Earth’s mantle: the role of nominally anhydrous minerals. Science 255, 1391–1396. Bjoraker, G.L., Mumma, M.J., Larson, H.P., 1989. Isotopic abundance ratios for hydrogen and oxygen in the Martian atmosphere. Bull. Am. Astron. Soc. 21, 991. Bockelee-Morvan, D., Gautier, D., Lis, D.C., Young, K., Keene, J., Phillips, T., 1998. Deuterated water in comet C/1996 B2 (Hyakutake) and its implications for the origin of comets. Icarus 133, 147–162. Boctor, N.Z., Alexander, C.M. O’D., Wang, J., Hauri, E., 2003. The sources of water in martian meteorites: clues from hydrogen isotopes. Geochim. Cosmochim. Acta 67, 3971–3989. Boettcher, A.L., O’Neil, J.R., 1980. Stable isotope, chemical and petrographic studies of high pressure amphiboles and micas: evidence for metasomatism in the mantle source regions of alkali basalts sand kimberlites. Am. J. Sci. 280A, 594–621. Bogard, D.D., Johnson, P., 1983. Martian gases in an Antarctic meteorite. Science 221, 651. Bridges, J.C., Catling, D.C., Saxton, J.M., Swindle, T.D., Lyon, I.C., Grady, M.M., 2001. Alteration assemblages in martian meteorites: implications for near-surface processes. Space Sci. Rev. 96, 365–392. Changela, H.G., Bridges, J.C., 2011. Alteration assemblages in the nakhlites: variation with depth on Mars. Meteorit. Planet. Sci. 45, 1847–1867. Deloule, E., 2002. D/H ratio ion probe measurements on magmatic minerals from martian meteorites: Implications for degassing of the martian mantle (abs). Lunar Planet. Sci. Conf. 33, 1607. Deloule, E., Albare de, F., Sheppard, S.M.F., 1991. Hydrogen isotope heterogeneities in the mantle from ion probe analysis of amphiboles from ultramafic rocks. Earth Planet. Sci. Lett. 105, 543–553. Eberhardt, P., Reber, M., Krankowsky, D., Hodges, R.R., 1995. The D/H and 18O/16O ratios in water from comet P/Halley. Astron. Astrophys. 302, 301–316. El Goresy, A., Dubrovinsky, L., Sharp, T.G., Saxena, S.K., Chen, M., 2000. A monoclinic post-stishovite polymorph of silica in the Shergotty meteorite. Science 288, 632–634. Gillet, P., Barrat, J.A., Crozaz, G., Deloule, E., Jambon, A., Neuville, D., Sautter, V., Wadhwa, M., 2001. Aqueous alterations in the NWA817 Martian meteorite (abs). Meteorit. Planet. Sci. 36, A66. Gillet, Ph., Barrat, J.A., Deloule, E., Wadhwa, M., Jambon, A., Sautter, V., Devouard, B., Neuville, D., Benzerara, K., Lesourd, M., 2002. Aqueous alteration in the Northwest Africa 817 (NWA 817) martian meteorite. Earth Planet. Sci. Lett. 203, 431–444. Gooding, J.L., Wentworth, S.J., Zolensky, M.E., 1991. Aqueous alteration of the Nakhla meteorite. Meteoritics 26, 135–143. Graham, C., Kinny, P., Harte, B., Valley, J., 1994. The nature and scale of stable isotope disequilibrium in the mantle: ion and laser microprobe evidence. Mineral. Mag. 58A, 345–346. Greenwood, J.P., Itoh, S., Sakamoto, N., Vicenzi, E.P., Yurimoto, H., 2008. Hydrogen isotope evidence for loss of water from Mars through time. Geophys. Res. Lett. 35, L05203. Greshake, A., 1998. Transmission electron microscopy characterization of shock defects in minerals from the Nakhla SNC meteorite (abs). Meteorit. Planet. Sci. 33, A63.
92
L.J. Hallis et al. / Earth and Planetary Science Letters 359-360 (2012) 84–92
Greshake, A., Langenhorst, F., 1997. Transmission electron microscope characterization of shock defects in minerals of the Martian meteorite Chassigny (abs). Meteorit. Planet. Sci. 32, A52. ¨ Hartogh, P., Lis, D.C., Bockelee-Morvan, D., de Val-Borro, M., Biver, N., Kuppers, M., Emprechtinger, M., Bergin, E.A., Crovisier, J., Renge, M., Moreno, R., Szutowicz, S., Blake, G.A., 2011. Ocean-like water in the Jupiter-family comet 103P/ Hartley 2. Nature 478, 218–220. Hallis, L.J., Taylor, G.J., Nagashima, K., Huss, G.R., Needham, A.W., Grady, M.M., Franchi, I.A., 2012. Hydrogen isotope analyses of alteration phases in the nakhlite martian meteorites. Geochim. Cosmochim. Acta 97, 105–119. Hoefs, J., 2004. Stable Isotope Geochemistry, 5th ed. Springer, Berlin (244 pp.). Jambon, A., 1994. Earth degassing and large-scale geochemical cycling of volatile elements. In: Carroll, M.R., Holloway, J.R. (Eds.), Volatiles in Magmas, vol. 30; 1994, pp. 479–517 (Min. Soc. Am., Rev. Min). Le´cuyer, C., Gillet, P., Robert, F., 1998. The hydrogen isotope composition of seawater and the global water cycle. Chem. Geol. 145, 249–261. Leshin, L.A., 2000. Insights into martian water reservoirs from analyses of martian meteorite QUE 94201. Geophys. Res. Lett. 27, 2017–2020. McCubbin, F.M., Steele, A., Haurib, E.H., Nekvasilc, H., Yamashitad, S., Hemleya, R.J., 2010. Nominally hydrous magmatism on the Moon. Proc. Natl. Acad. Sci. 107, 11223–11228. McCubbin, F.M., Hauri, E.H., Elardo, S.M., Vander Kaaden, K.E., Wang, J., Shearer Jr., C.K., 2012. Hydrous melting of the martian mantle produced both depleted and enriched shergottites. Geology 40, 683–686. Meier, R., Owen, T.C., Matthews, H.E., Jewitt, D.C., Bockelee-Morvan, D., Biver, N., Crovisie, J., Gautier, D., 1998. A determination of the HDO/H20 ratio in comet C/1995 01 (Hale-Bopp). Science 279, 842–844. Michael, P.J., 1988. The concentration, behavior and strorage of H2O in the suboceanic upper mantle: implications for mantle metasomatism. Geochim. Cosmochim. Acta 52, 555–566.
Minitti, M.E., Leshin, L.A., Darby Dyar, M., Ahrens, T.J., Guan, Y., Luo, S-N., 2008. Assessment of shock effects on amphibole water contents and hydrogen isotope compositions: 2. Kaersutitic amphibole experiments. Earth Planet. Sci. Lett. 266, 288–302. Reid, A.M., Bunch, T.E., 1975. The nakhlites, part II. Where, when and how? Meteoritics 10, 317–324. Robert, F., Gautier, D., Dubrulle, B., 2000. The solar system D/H ratio: observations and theories. Space Sci. Rev. 92, 201–224. Savin, S.M., Epstein, S., 1970. The oxygen and hydrogen isotope geochemistry of ocean sediments and shales. Geochim. Cosmochim. Acta 34, 43–63. Sharp, T.G., El Goresy, A., Wopenka, B., Chen, M., 1999. A post-stishovite SiO2 polymorph in the meteorite Shergotty: implications for impact events. Science 284, 1511–1514. Stauffer, H., 1962. On the production rates of rare gas isotopes in stone meteorites. J. Geophys. Res. 67, 2023–2028. Thompson, A.B., 1992. Water in the Earth’s upper mantle. Nature (London) 358, 295–302. Treiman, A.H., 2005. The nakhlite meteorites: augite-rich igneous rocks from Mars. Chem. Erde 65, 203–270. (invited review). Wagner, C., Deloule, E., Mokhtari, A., 1996. Richterite-bearing peridotites and MARID-type inclusions in lavas from North Eastern Morocco: Mineralogy and D/H isotopic studies. Cont. Min. Pet. 124, 406–421. Walsh, K.J., Morbidelli, A., Raymond, S.N., O’Brien, D.P., Mandell, A.M., 2011. A low mass for Mars from Jupiter’s early gas-driven migration. Nature 475, 206–209. Watson, L.L., Hutcheon, I.D., Epstein, S., Stolper, E.M., 1994. Water on Mars: clues from D/H and water contents of hydrous phases in SNC meteorites. Science 265, 85–90. Xia, Q-K., Deloule, E., Wu, Y-B., Chen, D-G., Cheng, H., 2002. Anomalously high dD values in the mantle. Geophys. Res. Lett. 29, 2008–2011.
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Magmatic water in the martian meteorite Nakhla L.J. Hallis a,b,c,n, G.J. Taylor a,c, K. Nagashima a, G.R. Huss a,c a
Hawai’i Institute of Geophysics and Planetology, Pacific Ocean Science and Technology (POST) Building, University of Hawai’i, 1680 East-West Road, Honolulu, HI 96822, United States Institute for Astronomy, University of Hawai’i, 2680 Woodlawn Drive, Honolulu, HI 96822-1839, United States c UH, NAI Astrobiology Program, United States b
a r t i c l e i n f o
a b s t r a c t
Article history: Received 7 March 2012 Received in revised form 8 September 2012 Accepted 27 September 2012 Editor: T. Elliot Available online 8 November 2012
Mars does not recycle crustal materials via plate tectonics. For this reason the magmatic water reservoir of the martian mantle has not been affected by surface processes, and the deuterium/hydrogen (D/H) ratio of this water should represent the original primordial martian value. Following this logic, hydrous primary igneous minerals on the martian surface should also carry this primordial D/H ratio, assuming no assimilation of martian atmospheric water during crystallization and no major hydrogen fractionation during melt degassing. Hydrous primary igneous minerals, such as apatite and amphibole, are present in martian meteorites here on Earth. Providing these minerals have not been affected by terrestrial weathering, martian atmospheric water, or shock processes after crystallization, they should contain a good approximation of the primordial martian D/H ratio. As Nakhla was seen to fall in the Egyptian desert in 1911, terrestrial contamination is minimized in this meteorite. The nakhlites are also among the least shocked of the martian meteorites. Therefore, apatite within Nakhla could contain primordial martian hydrogen isotope ratios. We produced in-situ measurements of the D/H ratios in Nakhla apatite grains, using a Cameca ims 1280 ion-microprobe. Our measurements produced D/H values in Nakhla apatite similar to terrestrial values, despite strong evidence that our samples were not significantly contaminated by terrestrial hydrogen. These results suggest that water trapped in the martian mantle has a similar D/H to that of the Earth. Therefore, the water of these two planets may have originated from the same source material. The D/H ratios of the carbonaceous chondrite meteorites, and the Jupiter-family comet 103P/ Hartley 2, are similar to the D/H of the two planets, making both these primitive inner solar system materials strong candidates for the source of the terrestrial planets water. These results support recent dynamical models of the formation of the solar system, which suggest material in the inner solar system was homogenized by the migration of Jupiter. & 2012 Elsevier B.V. All rights reserved.
Keywords: martian meteorites hydrogen isotopes terrestrial planet formation
1. Introduction Understanding the sources and delivery mechanisms of water to the Earth and the other terrestrial planets allows for the validation of planetary accretion models. A key parameter in determining the source(s) of terrestrial planetary water is hydrogen isotope composition. This is reasonably well established for the Earth’s water reservoirs. Here we report new measurements that attempt to establish the hydrogen isotopic composition of the martian interior. There are currently 48 martian meteorites available for study. Although these rocks were not sampled directly from the surface n Corresponding author at: Hawai’i Institute of Geophysics and Planetology, Pacific Ocean Science and Technology (POST) Building, University of Hawai’i, 1680 East-West Road, Honolulu, HI 96822, United States. E-mail address: [email protected] (L.J. Hallis).
0012-821X/$ - see front matter & 2012 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.epsl.2012.09.049
of Mars, their martian origin can be deduced by comparing the chemical composition of trapped gas bubbles with the composition of the martian atmosphere, as measured by the Viking landers (Bogard and Johnson, 1983). Despite the fact that there are various different types of martian meteorite, all are primarily igneous rocks from basaltic lava flows or shallow-level subsurface intrusions. Certain igneous primary minerals within these rocks, for example amphibole (potassic-chlorohastingsite) and apatite (Cl-rich), have water locked into their crystal structure. As the planetary accretion process does not appear to fractionate hydrogen isotopes (Le´cuyer et al., 1998), and as Mars does not recycle crustal materials via plate tectonics, the D/H ratio of water in these minerals could represent the primordial ratio of Mars’ water. If so, this ratio could be used to assess the origin of martian water. However, there are various processes that may have affected the D/H ratio of water within these minerals, including contamination via terrestrial weathering and/or martian atmospheric
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water, hydrogen fractionation during magmatic degassing, and shock processes after crystallization. The ratio of deuterium (2H or D) to hydrogen (1H) in water is commonly quoted relative to Vienna Standard Mean Ocean Water (VSMOW), and is given as dD¼ [((D/H)sample/(D/H)VSMOW) 1] 1000, with units of per ml (%). Hence, VSMOW has a dD value of 0%. On Earth, ice and rocks have negative dD values, with meteoric water varying from þ130% (some rainfall) to 480% (ice sheets), and the oceans at 0% (Le´cuyer et al., 1998 and references therein; Hoefs, 2004). Measurements of the Earth’s mantle have produced a range of values, from 140 to þ60% (e.g., Boettcher and O’Neil 1980; Michael, 1988; Ahrens, 1989; Deloule et al., 1991; Bell and Rossman, 1992; Thompson, 1992; Graham et al., 1994; Jambon, 1994; Wagner et al., 1996; Xia et al., 2002). In contrast, Mars’ atmospheric hydrogen has a positive dD value. Approximately 4 billion years ago Mars’ liquid core is believed to have solidified, causing the planet to lose its global magnetic field. This, in turn, initiated the loss of the martian atmosphere to space. As the lighter hydrogen isotope was preferentially lost, the atmosphere became deuterium enriched—the current martian atmospheric dD value is approximately þ4200% (Bjoraker et al., 1989). This enrichment was subsequently incorporated into secondary alteration minerals formed via weathering, such as the Fe-rich clay-like deposits and carbonates found in certain martian meteorites (Boctor et al., 2003). Amphibole and apatite are the primary igneous minerals, and hence they contain water from the martian interior. Providing there has been no subsequent contamination, their dD should reflect that of primordial Mars, without any atmospheric component. However, as reflected by previous hydrogen isotopic studies of both martian primary igneous and alteration phases (e.g., Watson et al., 1994; Leshin, 2000; Deloule, 2002; Gillet et al., 2002; Boctor et al., 2003; Greenwood et al., 2008) both martian and terrestrial surface fluids are potential sources of contamination.
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measured with a 4 nA primary ion beam (Feb 2012). As Nakhla apatites are rare and mostly very small, we were unable to collect a second set of results for these apatites with the higher beam current. However, although the hydrogen background is higher for the 2 nA analyses, it is not close to the level of counts obtained from the apatite grains [SOM]. Therefore, analyzing Nakhla apatite grains with a 4 nA primary beam current was not necessary. For Fe-rich clay-like veins, H, D, and 30Si were measured with 80 pA primary beam. The secondary ion mass spectrometer was operated at 10 keV with a 50 eV energy window. Isotopes were detected using an electron multiplier. The mass resolving power was 1900, sufficient to separate any interfering molecules. A normal-incidence electron flood gun was used for charge compensation. Prior to each measurement, the 2 nA primary beam (4 nA for Shergotty apatites) was rastered over a 25 25 mm2 area for 200 s, to remove the carbon coat and any surface contamination. During this rastering process the distribution of H ions in the region was displayed (although no image was permanently saved), providing the opportunity to avoid any H-rich cracks which may contain terrestrial contamination. The raster was then reduced to a 15 15 mm2 analysis area. To further reduce the contribution of unwanted hydrogen to our signal, we used the electronic gate to exclude the ions from the outer 50% of the 15 15 mm2 sputtered area, thus eliminating contributions from the crater walls and from hydrogen creeping along the surface of the collector. Therefore, our final analysis spot size was 8 8 mm2. The data were collected for 40 cycles, with H measured for 3 s, D for 40 s, and 18O (or 30Si for Fe-rich clay-like veins) for 2 s in each cycle. The beam was blanked for the first 10 cycles of the measurement, as well as for the last five cycles, to measure the background hydrogen and deuterium counts per second.
2.2. Data analysis 2. Methods 2.1. Measurement protocol Our sample set included eight UH polished thin-sections of Nakhla (20a–h), newly prepared from the same cm-sized rock chip (parent 20) without water (using polishing oil and petroleum ether to clean), and one thin-section of Shergotty (UNM 410), prepared with water prior to 1986 at the University of New Mexico [see SOM in the Appendix for petrographic descriptions]. The Nakhla samples were placed in a vacuum oven at 60 1C as soon as they were prepared to minimize atmospheric contamination, and were only removed for analysis. Shergotty UNM 410 was cleaned with methanol, placed in the vacuum oven for three days, re-carbon coated, and replaced in the vacuum oven one week prior to analysis. We utilized the JEOL JSM-5900LV scanning electron microscope at the University of Hawai’i (UH) to produce backscatter electron images and elemental X-ray images at 65 magnification (20 mm pixel size) for each thin-section, in order to locate areas of interest. These areas were subsequently imaged at higher resolutions, with various pixel sizes (approximately 1–5 mm), to pick out the individual apatite grains and Fe-rich clay-like veins for ion microprobe analysis. The areas analyzed for D/H were also imaged after analysis, to verify the positions of the sputtered regions. Deuterium and hydrogen isotopic compositions were analyzed in-situ with the UH Cameca ims 1280 ion microprobe. A 2 nA focused Cs þ primary ion beam was used to produce negative ions of H, D, and 18O for Nakhla apatites (June 2011), following the protocol of Boctor et al. (2003). However, the background counts per second (cps) were subsequently found to slightly decrease with a higher beam current, so the Shergotty apatite grains were
The data from the ion probe are reported in counts per second (cps). However, when we use the electronic gate, the beam is deflected away from the detector for a significant portion of each second, so the reported count rates are too low to make an appropriate deadtime correction. Therefore, we first correct the measured count rates to account for the fraction of time that the beam is out of the detector. The data are then corrected for deadtime. The isotope data were subsequently corrected for background, which was measured during the first 10 and last five cycles of each measurement as described above. The background signal was routinely 3000–7000 cps for Nakhla apatites, and 400–600 cps for Shergotty apatites (because of the increased beam current). The total hydrogen signal with the beam on was 200,000–1,200,000 cps for Nakhla apatites, and 800,000–1,000,000 cps for Shergotty apatites [Appendix (SOM)]. The background counts were subtracted from the measured isotope signal, collected between cycles 11 and 35 while the beam was on. The counting-statistical uncertainty of the background signal was propagated. Apatite data were corrected for instrumental fractionation using two terrestrial apatite standards (Crystal Lode ap005 and Russia ap018). The water contents and dD values of these apatites were previously determined by McCubbin et al. (2010, 2012) [SOM]. While these previous publications did not determine the dD value of Durango ap003, the water content of this standard was established. Therefore, as these three apatite standards (Durango ap003, Crystal Lode ap005 and Russia ap018) contain varying water contents, calibration curves were produced to estimate the water content of the unknowns (Fig. 1). The H2O content of each unknown apatite was estimated based on its 1H /18O ratio, and the slope of the
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Fig. 1. 1H/18O vs. H2O wt% calibration lines for the Nakhla apatite (a) and Shergotty apatite (b) analyses. These curves are based on the measurements of three terrestrial apatite standards with known water contents at 2 nA (a) and 4 nA (b).
relevant 1H/18O vs. H2O wt% standard calibration line (Fig. 1), using the following equation: H2O wt%¼(1H /18O ) line slope
(1)
The H2O detection limit was estimated in a similar way, via the measurement of 1H /18O ratios in a nominally anhydrous San Carlos olivine standard. Detection limits are consistently 30 ppm H2O for Nakhla (2 nA), and 10 ppm for Shergotty (4 nA). The difference in these limits is caused by a reduction in instrumental hydrogen background during the Shergotty analyses, because of the higher beam current. As these detection limits are so low, the calibration lines for both primary beam currents were forced through the origin, following the methodology of McCubbin et al. (2010, 2012). The reproducibility of the apatite water contents, and D/H ratios, were calculated based on the subtraction of the true H2O or D/H value from the estimated H2O or D/H value of each apatite standard (error¼2 standard deviation (estimated-true)). Apatite water content reproducibility is 0.02 wt% H2O for Nakhla, and 0.01 wt% H2O for Shergotty. The 2s uncertainties of each apatite H2O content and D/H ratio were subsequently calculated using the following equation: ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi r ffi 2 2s error ¼ 2 measurement error2 þ standard reproducibility ð2Þ The H2O contents of the Nakhla Fe-rich clay-like veins cannot accurately be estimated in this way, as these veins are made up of more than one mineralogical phase. Fe-rich clay-like vein data were corrected for instrumental mass fractionation using a serpentine standard (UH 180C 18) from Conical Seamont, Mariana Arc, south of Japan (Alt and Shanks, 2006).
3. Results Apatite grains in the martian meteorite Nakhla (Fig. 2) show
dD values of between 111 and þ155% (Table 1). These values are similar to those of terrestrial meteoric water and the Earth’s mantle, excepting that the range stretches to slightly more positive values. The samples studied were sourced from a group of Nakhla specimens cataloged at the British Natural History Museum in 1913—2 yr after the Nakhla meteorite shower over Egypt [see SOM for petrographic descriptions in the Appendix]. It is unclear whether the specimens spent this 2 yr period in the museum waiting to be cataloged, or in the field exposed to weathering. Therefore, to test the extent of terrestrial hydrogen
contamination in our samples, we measured the dD value of coexisting pre-terrestrial Fe-rich clay-like veins. These clay-like vein assemblages, or ‘iddingsite’ veins formed as a result of martian surficial aqueous alteration of the crystallized nakhlite basalt (e.g., Ashworth and Hutchison, 1975; Reid and Bunch, 1975; Gooding et al., 1991; Bridges et al., 2001; Gillet et al., 2001, 2002; Treiman, 2005). They are reported to consist of mixtures of hydrous amorphous silicate gel, siderite, gypsum and carbonate (Changela and Bridges, 2011), possibly with some smectite clay (Gooding et al., 1991) [see also SOM in the Appendix]. The phases within these veins, especially smectite and hydrous amorphous silicate gel, contain only loosely bound water that would be susceptible to hydrogen exchange with the surrounding atmosphere. The original dD value of these alteration phases on Mars would depend on the dD value of the alterationforming fluid. However, even if this fluid had a low dD value (which may possibly be the case if some martian hydrothermal fluids exist/existed having had no contact with atmospheric hydrogen), subsequent exposure of the ‘iddingsite’ veins to the martian atmosphere would have resulted in relatively rapid hydrogen exchange. Therefore, water within these veins should have been enriched in deuterium at the time the nakhlites were blasted from Mars. Hydrogen isotope fractionation between the martian atmosphere and the specific alteration vein phases would affect the level of this enrichment. However, the fractionation factor between montmorillonite clay and water is 0.94 at sedimentary temperatures on Earth (Savin and Epstein, 1970), resulting in only a slight enrichment of the lighter hydrogen isotope in the solid phase. This same effect is seen in other clays, amphiboles, and micas (see Hallis et al., 2012 for further discussion). As explained above, the layered and commonly amorphous structure of the clay-like ‘iddingsite’ vein assemblages results in non-structural and/or loosely bound water molecules. Loosely bound water leads to enhanced hydrogen exchange with the surrounding atmosphere, compared to purely crystalline minerals (such as apatite) where water is locked into the crystal structure. It therefore follows that if these alteration assemblages still retain their martian D enrichment, their hydrogen isotopes have not completely exchanged with terrestrial atmospheric hydrogen— terrestrial contamination in the samples is minimal. If this very easily exchangeable martian hydrogen has not been completely replaced, neither has the more structurally bound martian hydrogen in the apatite grains. Our measurements from Nakhla samples 20c and 20e show dD values between þ1170 and 105% (Table 1) within these assemblages, indicating incomplete hydrogen isotope exchange with the terrestrial atmosphere. The abundance of 1H (cps) in these veins gives an indication of their
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Fig. 2. Backscattered electron (A, C, E, G), and secondary electron (B, D, F, H), images of Nakhla apatites 20b1 (A and B) 20b2 (C and D), 20b3 (E and F) and 20c1 (G and H). Backscattered electron images show apatite grains (ap) within the mesostasis (meso) of Nakhla, associated with clinopyroxene (cpx) and olivine (olv) phenocrysts, along with other more minor igneous phases, including the silica polymorph cristobalite (crist), the iron sulfide pyrrhotite (FeS), feldspar (feld) and Fe-rich clinopyroxene (Fecpx). Secondary electron images highlight the extent of the SIMS pre-sputtered raster pits on each apatite grain (the extent each grain is approximated by the white shape). The raster pits are 25 25 mm2, but D/H measurements only include counts from the central 8 8 mm2 (approximated by the black squares).
relative H2O contents, and these values also vary widely—from 150,000 to 880,000. This wide variation in H2O content reflects the heterogeneous mixture of hydrous and anhydrous silicates and non-silicates in these veins. Compositional variation is also suggested by the variation in 1H/30Si ratios (Table 1). When these values are compared to dD for individual analyses, there is generally an inversely proportional relationship—as 1H cps or 1H/30Si ratio increases, dD decreases. This relationship is indicative of different amounts of terrestrial atmospheric contamination, with the most
H-rich and deuterium depleted veins having absorbed the most terrestrial hydrogen. A wide variation in ‘iddingsite’ vein dD values has also been reported in other nakhlite samples, and appears to be strongly dependant on the sample and parent meteorite history (Gillet et al., 2002; Hallis et al., 2012). To further test our methodology, and the reliability of our Nakhla apatite results, we measured the D/H ratio of apatite grains in Shergotty (Fig. 3). Our measurements show dD values of 2953%–3931% (Table 1). The lowest of these values was
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Table 1 dD (%) and H2O contents of Nakhla and Shergotty apatite grains and standards, and dD (%) of Nakhla Fe-rich claylike veins and standards. The calculated H2O content of anhydrous minerals indicates the detection limits. 1
Sample
dD (%)
2r
Apatites (2 nA) Nakhla 20b apatite 1 Nakhla 20b apatite 1 (2) Nakhla 20b apatite 2 Nakhla 20c apatite 1 Durango apatite std 1 Durango apatite std 2 Crystal Lode apatite std 1 Crystal Lode apatite std 2 Crystal Lode apatite std 4 Russia apatite std 1 Russia apatite std 2 Russia apatite std 3
188 78 19 151 4 71 60 36 47 107 91 128
62 62 62 62 3 3 4 6 6 8 8 8
0.43 0.42 0.09 0.58 0.04 0.05 0.33 0.33 0.33 0.19 0.19 0.19
Apatites (4 nA) Shergotty apatite 1 Shergotty apatite 1 (2) Shergotty apatite 2 Shergotty apatite 8 Shergotty apatite 9 Durango apatite std 1 Durango apatite std 2 Crystal Lode apatite std 1 Crystal Lode apatite std 2 Crystal Lode apatite std 3 Russia apatite std 1 Russia apatite std 2 Russia apatite std 3
3563 3931 3267 2953 3516 141 108 99 113 101 145 104 109
12 13 11 10 11 10 11 4 3 3 4 5 5
0.63 0.55 0.66 0.71 0.67 0.04 0.04 0.37 0.39 0.38 0.21 0.20 0.20
Fe-rich clay-like veins (80 pA) Nakhla 20c Fe-vein 1 852 Nakhla 20c Fe-vein 2 6 Nakhla 20c Fe-vein 3 1165 Nakhla 20c Fe-vein 4 702 Nakhla 20c Fe-vein 5 16 Nakhla 20c Fe-vein 6 66 Nakhla 20c Fe-vein 7 98 Nakhla 20c Fe-vein 8 42 Nakhla 20e Fe-vein 1 110 Nakhla 20e Fe-vein 2 108 Nakhla 20e Fe-vein 3 229 Nakhla 20e Fe-vein 4 237 Nakhla 20e Fe-vein 5 35 Serpentine std 1 53 Serpentine std 2 36 Serpentine std 3 28 Serpentine std 4 0 Serpentine std 5 54 Serpentine std 6 42 Serpentine std 7 53
85 77 86 96 80 75 79 76 77 78 82 77 78 77 75 75 77 76 75 76
3.18 21.43 91.74 7.58 22.83 42.41 38.95 49.38 25.88 27.29 24.93 13.08 3.25 38.25 33.46 44.44 51.81 51.24 50.43 37.54
H/monitor isotopea
2r
H2O (wt%)
2r
0.0001 0.0001 0.00003 0.0001 0.00002 0.00002 0.00006 0.00006 0.00006 0.00004 0.00004 0.00004
0.47 0.46 0.10 0.64 0.03 0.03 0.37 0.37 0.37 0.20 0.20 0.20
0.02 0.02 0.02 0.02 0.01 0.01 0.02 0.02 0.02 0.02 0.02 0.02
0.0002 0.0002 0.0002 0.0002 0.0002 0.00005 0.00004 0.00018 0.00016 0.00015 0.00011 0.00010 0.00005
0.62 0.54 0.65 0.70 0.66 0.04 0.04 0.36 0.38 0.37 0.21 0.19 0.20
0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01
38.46 38.47 38.99 38.46 38.47 38.48 38.51 38.48 38.50 38.47 38.48 38.47 38.46 38.47 38.47 38.50 38.52 38.49 38.48 38.48
Anhydrous min (2 nA) Nakhla 20b augite San Carlos olivine 1 San Carlos olivine 2 San Carlos olivine 3
0.0088 0.0029 0.0031 0.0027
0.00002 0.00016 0.00017 0.00013
0.0097 0.0032 0.0034 0.0030
0.04076 0.04076 0.04076 0.04076
Anhydrous min (4 nA) Shergotty augite 1 Shergotty olivine 1 Shergotty olivine 2 San Carlos olivine 1 San Carlos olivine 2
0.0329 0.0013 0.0088 0.0012 0.0011
0.00002 0.00001 0.00002 0.00001 0.00001
0.0323 0.0012 0.0086 0.0012 0.0011
0.01414 0.01414 0.01414 0.01414 0.01414
a Monitor isotope¼ 18O for apatite and anhydrous mineral measurements, and 30Si for Fe-rich clay-like vein measurements.
produced from an analysis which may have included some terrestrial contamination from a crack within the apatite grain (Fig. 3). All other analyses appear to be free from this problem, at least at the visible scale. These measurements can be directly
compared with those of Greenwood et al. (2008), who measured the dD of apatites in Shergotty and Los Angeles (another shergottite meteorite). Our Shergotty apatite dD values are comparable with the apatite dD values of Los Angeles, but are 650% lower
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Fig. 3. Secondary electron images of Shergotty UNM 410 SIMS analysis pits. Apatite 1 contains two analysis pits (A and B), whereas apatites 2 (C and D), 8 (E and F) and 9 (G and H) contain only one. A test pit was made in the olivine grain next to apatite 9 (G). Lower magnification images (A, C, E, G) outline the general extent of each apatite grain (within the white shapes), and higher magnification images (B, D, F, H) show the pits in detail. As in Fig. 2, the central 8 8 mm2 measurement areas are approximated by the black squares. Ap¼ apatite, olv ¼olivine.
than the two previously measured apatite dD values of Shergotty (Greenwood et al., 2008). This difference may be caused by our analyses of a larger area on each apatite grain (8 8 mm2 rather than a 2 2 mm2), causing an increased risk of terrestrial
contamination from micro-cracks. Previously measured apatite and amphibole dD measurements from other martian meteorites are comparable to, or significantly lower than, our Shergotty values (Watson et al., 1994; Leshin, 2000; Boctor et al., 2003).
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4. Discussion 4.1. Possibility and effect of different sources of hydrogen isotope contamination The relationship of our measured apatite dD values with the original primordial dD value of Mars is dependent on three factors. The first factor is the amount of martian atmospheric contamination and degassing the apatite grains were exposed to during their crystallization and subsequent 1.3 billion year residence on Mars’ surface (Stauffer, 1962). The fact that Nakhla’s apatite dD values are much lower than some of its clay-like alteration assemblage dD values indicates that the majority of the water in these apatites is not sourced from martian atmospheric water, either via assimilation of crustal material or interaction with martian groundwater during crystallization. In addition, it is clear that Nakhla’s apatite grains did not completely exchange their hydrogen with the martian atmosphere after crystallization—in this case, again, more positive dD values would have been measured. However, some minor martian atmospheric contamination remains a possibility, as both Nakhla 20b apatite 1 and 20c apatite 1 show slightly positive dD values, above the normal terrestrial range (Table 1; Fig. 4). There is also evidence that the nakhlite melt probably degassed as it cooled, and that hydrogen isotope fractionation occurred during this process. McCubbin et al. (2012) attempted to determine the water content of the martian mantle by measuring that of the shergottite apatites. These measurements indicate that the shergottite magmas underwent magmatic degassing, because their water content as calculated from apatites is greater than the measured bulk-rock water content. Prior analysis of an apatite grain within the shergottite meteorite Los Angeles established the presence of isotopic zoning, from a relatively deuterium-poor core (dD¼4120%) to a deuterium-rich rim (dD¼4348%) (Greenwood et al., 2007, 2008). This zoning suggests that as the magma degassed, apatites became D-enriched, hence dD increased. Deuterium enrichment during degassing agrees with the general rule that during isotopic diffusion the lighter isotope
Fig. 4. Comparison of the dD values of our data (black symbols) with relevant data from previous studies (black and gray lines) (Watson et al., 1994; Leshin, 2000; Boctor et al., 2003; Greenwood et al., 2008). The high dD values of amphibole and apatite grains within Shergotty and the other previously measured meteorites could be explained by the greater shock pressures these rocks were exposed to relative to Nakhla, or by greater assimilation with martian atmospheric groundwater. Fe-rich clay-like veins in Nakhla show dD similar to ALH 84001 carbonates, proving that the low apatite dD in this meteorite cannot be the result of terrestrial hydrogen isotope exchange.
becomes enriched in the gas phase (as it has a weaker bond than the heavier isotope). Therefore, degassing of the melt, and of the apatites after crystallization, should result in D enrichment in the nakhlite rocks/apatites. Hence, in addition to possible martian atmospheric contamination, magmatic degassing may have also increased the deuterium content in Nakhla’s apatites relative to the original melt value. The second factor is the effect of shock on the apatites as they were blasted from the martian surface via meteorite impact. The high shock pressures involved in impact events can volatilize hydrogen in hydrated minerals, with 1H being more susceptible to this process than D. Volatilization has been reported to increase dD by up to 87% in lab-based experiments (Minitti et al., 2008). In addition, volatilization is accompanied by shock-induced implantation of water from the surrounding atmosphere. Therefore, meteorites that were heavily shocked on the surface of Mars might be expected to contain hydrated primary minerals enriched in D compared to the equivalent minerals in less shocked meteorites. These experimental results are consistent with our measurements of dD in Shergotty apatite grains, as well as previously reported dD values from the igneous hydrated minerals of Shergotty and other martian meteorites (Watson et al., 1994; Leshin, 2000; Boctor et al., 2003; Greenwood et al., 2008) (Fig. 4). These previous studies involved meteorites that were heavily shocked (31–49 GPa, possibly as high as 90 GPa for Shergotty—Greshake and Langenhorst, 1997; Sharp et al., 1999; El Goresy et al., 2000; Bell, 2007), and reported apatite dD values between þ 751 and þ4606% (Watson et al., 1994; Leshin, 2000; Boctor et al., 2003; Greenwood et al., 2008). In contrast, Nakhla was probably less severely shocked (20 GPa—Greshake, 1998), therefore is less likely to have acquired D-rich water from shock. Based on their measurements of various martian meteorites, Boctor et al. (2003) also reached the conclusion that increased shock increases the dD value in primary igneous phases. All of the above processes—martian atmospheric contamination, magmatic degassing, and the twin processes of shockinduced hydrogen volatilization and martian atmospheric water implantation during shock—would potentially increase dD in the hydrated minerals of martian meteorites. Therefore, we could assume a maximum value of 111 762% for the hydrogen isotopic composition of the martian mantle, as this is the lowest dD value in our apatite dataset, hence the least affected by these processes. However, before assuming this value is generally correct, we need to take into account the third factor—terrestrial contamination. Despite the fact that Nakhla is a fall meteorite, and that our samples were prepared without the use of water, terrestrial contamination may still be an issue in the apatite grains. As discussed above, Fe-rich clay-like veins within these samples do show some high dD values, indicating that this easily altered material has not completely exchanged its hydrogen with the terrestrial atmosphere. Micro-cracks in the apatite grains have proved to be a source of terrestrial contamination, as illustrated by the low dD value produced when a visible crack was included in one of our Shergotty analyses (Table 1). Greenwood et al. (2008) also recorded the effect of these cracks. The 650% difference between our Shergotty apatite dD values and those of Greenwood et al. (2008) may be the result of increased microcrack contamination, caused by our larger analyses areas. Therefore, it may be that our maximum martian mantle value of 111% is too low. However, it should be expected that terrestrial micro-crack contamination in both our Shergotty data, and that of Greenwood et al. (2008), would be greater than in our Nakhla apatite data, as both Shergotty samples were prepared using water. Previous dD measurements of Shergotty hydrous igneous phases show a wide variation (Watson et al., 1994; Boctor et al., 2003; Greenwood et al., 2008).
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Table 2 Previously measured major terrestrial, martian and cometary reservoir dD reservoirs. Known dD reservoirs
dD (%)
Terrestrial Oceansa Terrestrial Ice Sheetsa Terrestrial Meteoric Waterb Terrestrial Mantlec Comet Hartley 2dn Comet P/Halleye Comet Hyakutakef Comet Hale-Boppg Martian Atmosphereh
0 300 to 400 40 to þ130 140 to þ 60 6 888 813 1063 4200
2s(%)
300 92 625 1000
a
Le´cuyer et al. (1998). Hoefs (2004). c Boettcher and O’Neil (1980), Michael (1988), Ahrens (1989), Deloule et al. (1991), Bell and Rossman (1992), Thompson (1992), Graham et al. (1994), Jambon (1994), Wagner et al. (1996) and Xia et al. (2002). d Hartogh et al. (2011). e Eberhardt et al. (1995). f Bockelee-Morvan et al. (1998). g Meier et al. (1998). h Bjoraker et al. (1989). n Comet Hartley 2 is a Jupiter family comet, whereas P/Halley, Hyakutake and Hale-Bopp are all Oort cloud comets. b
These differences may indicate heterogeneity between different apatite grains in Shergotty, possibly explaining the difference between our dD values for this meteorite and those of Greenwood et al. (2008). Taking into account the above information, the amount of Nakhla apatite dD alteration caused by terrestrial and martian atmospheric hydrogen contamination, shock, and magmatic degassing is apparently small. We can therefore say with some certainty that our Nakhla apatite data reflect the best estimate of the true martian mantle dD value to date, and that this value is similar to that of the Earth. 4.2. The source of martian water The similarity of our proposed martian mantle dD value to that of the terrestrial oceans, rocks and mantle could suggest that Earth and Mars, and possibly the other terrestrial planets, accreted water from the same source, or from more than one source with similar D/H ratios. Carbonaceous chondrites have been considered a likely source because of the similarity in their D/H to terrestrial ocean water (Robert et al., 2000). Jupiter-family comets may also be a source; comet 103P/Hartley 2 has a D/H ratio indistinguishable from that of Earth (Hartogh et al., 2011). A significant contribution from the long-period comets of the Oort Cloud seems to be ruled out by the elevated D/H ratios of their water (dD4 þ800%, Table 2; Eberhardt et al., 1995; BockeleeMorvan et al., 1998; Meier et al., 1998). Therefore, the inner solar system materials may all have similar D/H ratios, whereas materials from further out in the solar system appear to be D enriched. Recent dynamical models suggest that the inward and outward migration of Jupiter caused extensive mixing of material in the inner solar system, while leaving the outer solar system relatively unmixed (Walsh et al., 2011). These models could explain the difference between inner and outer solar system hydrogen isotope reservoirs, and the similarity between the D/H ratios of Mars and Earth.
Acknowledgments This material is based upon work supported by the National Aeronautics and Space Administration through the NASA
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Astrobiology Institute under Cooperative Agreement no. NNA09DA77A issued through the Office of Space Science. We also thank the NASA Johnson Space Center for allocation of Nakhla (20), and JoAnn Sinton at the University of Hawaii for producing the thinsections. Eric Hellebrand is thanked for his assistance with EMP analyses. Prof. F. McCubbin is thanked for supplying the terrestrial apatite standards. The comments and suggestions of Prof. S. Krot, Dr. Shoichi Itoh and Prof. Tim Elliott, as well as two anonymous reviewers, were invaluable in improving this manuscript.
Appendix A. Supporting information Supplementary data associated with this article can be found in the online version at http://dx.doi.org/10.1016/j.epsl.2012.09.049.
References Ashworth, J.R., Hutchison, R., 1975. Water in non-carbonaceous stony meteorites. Nature 256, 714–715. Ahrens, T.J., 1989. Water storage in the upper mantle. Nature (London) 342, 122–123. Alt, J.C., Shanks, W.C., 2006. Stable isotope compositions of serpentinite seamounts in the Mariana forearc: serpentinization processes, fluid sources and sulfur metasomatism. Earth Planet. Sci. Lett. 242, 272–285. Bell, M.S., 2007. Experimental shock decomposition of siderite and the origin of magnetite in Martian meteorite ALH 84001. Meteorit. Planet. Sci. 42, 935–949. Bell, D.R., Rossman, G.R., 1992. Water in Earth’s mantle: the role of nominally anhydrous minerals. Science 255, 1391–1396. Bjoraker, G.L., Mumma, M.J., Larson, H.P., 1989. Isotopic abundance ratios for hydrogen and oxygen in the Martian atmosphere. Bull. Am. Astron. Soc. 21, 991. Bockelee-Morvan, D., Gautier, D., Lis, D.C., Young, K., Keene, J., Phillips, T., 1998. Deuterated water in comet C/1996 B2 (Hyakutake) and its implications for the origin of comets. Icarus 133, 147–162. Boctor, N.Z., Alexander, C.M. O’D., Wang, J., Hauri, E., 2003. The sources of water in martian meteorites: clues from hydrogen isotopes. Geochim. Cosmochim. Acta 67, 3971–3989. Boettcher, A.L., O’Neil, J.R., 1980. Stable isotope, chemical and petrographic studies of high pressure amphiboles and micas: evidence for metasomatism in the mantle source regions of alkali basalts sand kimberlites. Am. J. Sci. 280A, 594–621. Bogard, D.D., Johnson, P., 1983. Martian gases in an Antarctic meteorite. Science 221, 651. Bridges, J.C., Catling, D.C., Saxton, J.M., Swindle, T.D., Lyon, I.C., Grady, M.M., 2001. Alteration assemblages in martian meteorites: implications for near-surface processes. Space Sci. Rev. 96, 365–392. Changela, H.G., Bridges, J.C., 2011. Alteration assemblages in the nakhlites: variation with depth on Mars. Meteorit. Planet. Sci. 45, 1847–1867. Deloule, E., 2002. D/H ratio ion probe measurements on magmatic minerals from martian meteorites: Implications for degassing of the martian mantle (abs). Lunar Planet. Sci. Conf. 33, 1607. Deloule, E., Albare de, F., Sheppard, S.M.F., 1991. Hydrogen isotope heterogeneities in the mantle from ion probe analysis of amphiboles from ultramafic rocks. Earth Planet. Sci. Lett. 105, 543–553. Eberhardt, P., Reber, M., Krankowsky, D., Hodges, R.R., 1995. The D/H and 18O/16O ratios in water from comet P/Halley. Astron. Astrophys. 302, 301–316. El Goresy, A., Dubrovinsky, L., Sharp, T.G., Saxena, S.K., Chen, M., 2000. A monoclinic post-stishovite polymorph of silica in the Shergotty meteorite. Science 288, 632–634. Gillet, P., Barrat, J.A., Crozaz, G., Deloule, E., Jambon, A., Neuville, D., Sautter, V., Wadhwa, M., 2001. Aqueous alterations in the NWA817 Martian meteorite (abs). Meteorit. Planet. Sci. 36, A66. Gillet, Ph., Barrat, J.A., Deloule, E., Wadhwa, M., Jambon, A., Sautter, V., Devouard, B., Neuville, D., Benzerara, K., Lesourd, M., 2002. Aqueous alteration in the Northwest Africa 817 (NWA 817) martian meteorite. Earth Planet. Sci. Lett. 203, 431–444. Gooding, J.L., Wentworth, S.J., Zolensky, M.E., 1991. Aqueous alteration of the Nakhla meteorite. Meteoritics 26, 135–143. Graham, C., Kinny, P., Harte, B., Valley, J., 1994. The nature and scale of stable isotope disequilibrium in the mantle: ion and laser microprobe evidence. Mineral. Mag. 58A, 345–346. Greenwood, J.P., Itoh, S., Sakamoto, N., Vicenzi, E.P., Yurimoto, H., 2008. Hydrogen isotope evidence for loss of water from Mars through time. Geophys. Res. Lett. 35, L05203. Greshake, A., 1998. Transmission electron microscopy characterization of shock defects in minerals from the Nakhla SNC meteorite (abs). Meteorit. Planet. Sci. 33, A63.
92
L.J. Hallis et al. / Earth and Planetary Science Letters 359-360 (2012) 84–92
Greshake, A., Langenhorst, F., 1997. Transmission electron microscope characterization of shock defects in minerals of the Martian meteorite Chassigny (abs). Meteorit. Planet. Sci. 32, A52. ¨ Hartogh, P., Lis, D.C., Bockelee-Morvan, D., de Val-Borro, M., Biver, N., Kuppers, M., Emprechtinger, M., Bergin, E.A., Crovisier, J., Renge, M., Moreno, R., Szutowicz, S., Blake, G.A., 2011. Ocean-like water in the Jupiter-family comet 103P/ Hartley 2. Nature 478, 218–220. Hallis, L.J., Taylor, G.J., Nagashima, K., Huss, G.R., Needham, A.W., Grady, M.M., Franchi, I.A., 2012. Hydrogen isotope analyses of alteration phases in the nakhlite martian meteorites. Geochim. Cosmochim. Acta 97, 105–119. Hoefs, J., 2004. Stable Isotope Geochemistry, 5th ed. Springer, Berlin (244 pp.). Jambon, A., 1994. Earth degassing and large-scale geochemical cycling of volatile elements. In: Carroll, M.R., Holloway, J.R. (Eds.), Volatiles in Magmas, vol. 30; 1994, pp. 479–517 (Min. Soc. Am., Rev. Min). Le´cuyer, C., Gillet, P., Robert, F., 1998. The hydrogen isotope composition of seawater and the global water cycle. Chem. Geol. 145, 249–261. Leshin, L.A., 2000. Insights into martian water reservoirs from analyses of martian meteorite QUE 94201. Geophys. Res. Lett. 27, 2017–2020. McCubbin, F.M., Steele, A., Haurib, E.H., Nekvasilc, H., Yamashitad, S., Hemleya, R.J., 2010. Nominally hydrous magmatism on the Moon. Proc. Natl. Acad. Sci. 107, 11223–11228. McCubbin, F.M., Hauri, E.H., Elardo, S.M., Vander Kaaden, K.E., Wang, J., Shearer Jr., C.K., 2012. Hydrous melting of the martian mantle produced both depleted and enriched shergottites. Geology 40, 683–686. Meier, R., Owen, T.C., Matthews, H.E., Jewitt, D.C., Bockelee-Morvan, D., Biver, N., Crovisie, J., Gautier, D., 1998. A determination of the HDO/H20 ratio in comet C/1995 01 (Hale-Bopp). Science 279, 842–844. Michael, P.J., 1988. The concentration, behavior and strorage of H2O in the suboceanic upper mantle: implications for mantle metasomatism. Geochim. Cosmochim. Acta 52, 555–566.
Minitti, M.E., Leshin, L.A., Darby Dyar, M., Ahrens, T.J., Guan, Y., Luo, S-N., 2008. Assessment of shock effects on amphibole water contents and hydrogen isotope compositions: 2. Kaersutitic amphibole experiments. Earth Planet. Sci. Lett. 266, 288–302. Reid, A.M., Bunch, T.E., 1975. The nakhlites, part II. Where, when and how? Meteoritics 10, 317–324. Robert, F., Gautier, D., Dubrulle, B., 2000. The solar system D/H ratio: observations and theories. Space Sci. Rev. 92, 201–224. Savin, S.M., Epstein, S., 1970. The oxygen and hydrogen isotope geochemistry of ocean sediments and shales. Geochim. Cosmochim. Acta 34, 43–63. Sharp, T.G., El Goresy, A., Wopenka, B., Chen, M., 1999. A post-stishovite SiO2 polymorph in the meteorite Shergotty: implications for impact events. Science 284, 1511–1514. Stauffer, H., 1962. On the production rates of rare gas isotopes in stone meteorites. J. Geophys. Res. 67, 2023–2028. Thompson, A.B., 1992. Water in the Earth’s upper mantle. Nature (London) 358, 295–302. Treiman, A.H., 2005. The nakhlite meteorites: augite-rich igneous rocks from Mars. Chem. Erde 65, 203–270. (invited review). Wagner, C., Deloule, E., Mokhtari, A., 1996. Richterite-bearing peridotites and MARID-type inclusions in lavas from North Eastern Morocco: Mineralogy and D/H isotopic studies. Cont. Min. Pet. 124, 406–421. Walsh, K.J., Morbidelli, A., Raymond, S.N., O’Brien, D.P., Mandell, A.M., 2011. A low mass for Mars from Jupiter’s early gas-driven migration. Nature 475, 206–209. Watson, L.L., Hutcheon, I.D., Epstein, S., Stolper, E.M., 1994. Water on Mars: clues from D/H and water contents of hydrous phases in SNC meteorites. Science 265, 85–90. Xia, Q-K., Deloule, E., Wu, Y-B., Chen, D-G., Cheng, H., 2002. Anomalously high dD values in the mantle. Geophys. Res. Lett. 29, 2008–2011.