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Mineralogy and noble gas isotopes of micrometeorites collected from Antarctic snow
© Okazaki et al. 2015
- Received: 10 December 2014
- Accepted: 3 June 2015
- Published: 17 June 2015
We have investigated seven micrometeorites (MMs) from Antarctic snow collected in 2003 and 2010 by means of electron microscopy, X-ray diffraction, micro-Raman spectroscopy, transmission electron microscopy (TEM) observation, and noble-gas isotope analysis. Isotopic ratios of He and Ne indicate that the noble gases in these MMs are mostly of solar wind (SW). Based on the release patterns of SW 4He, which should reflect the degree of heating during atmospheric entry, the seven MMs were classified into three types including two least heated, three moderately heated, and two severely heated MMs. The heating degrees are well correlated to their mineralogical features determined by TEM observation. One of the least heated MMs is composed of phyllosilicates, whereas the other consists of anhydrous minerals within which solar flare tracks were observed. The two severely heated MMs show clear evidence of atmospheric heating such as partial melt of the uppermost surface layer in one and abundant patches of dendritic magnetite and Si-rich glass within an olivine grain in the other. It is noteworthy that a moderately heated MM composed of a single crystal of olivine has a 3He/4He ratio of 8.44 × 10−4, which is higher than the SW value of 4.64 × 10−4, but does not show a cosmogenic 21Ne signature such as 20Ne/21Ne/22Ne = 12.83/0.0284/1. The isotopic compositions of He and Ne in this sample cannot be explained by mixing of a galactic cosmic ray (GCR)-produced component and SW gases. The high 3He/4He ratio without cosmogenic 21Ne signature likely indicates the presence of a 3He-enriched component derived from solar energetic particles.
- Antarctic micrometeorites
- Noble gas
- Transmission electron microscopy
- Atmospheric entry heating
- Solar energetic particles
Our understanding of the evolution history of our early solar system materials stands on studies of extraterrestrial materials originating from asteroids and comets. Asteroidal materials occurring as meteorites have been studied in detail. In particular, those recovered from the asteroid Itokawa by the Hayabusa spacecraft have been subjected to comprehensive studies (e.g., Nakamura et al. 2011; Noguchi et al. 2011, 2014a, 2014b; Nagao et al. 2011; Tsuchiyama et al. 2011; Yurimoto et al. 2011; Nakamura et al. 2012; Harries and Langenhorst 2014; Keller and Berger 2014; Langenhorst et al. 2014; Mikouchi et al. 2014; Thompson et al. 2014; Takeda et al. 2015). In contrast, it is difficult to obtain samples identified to be cometary in origin, with the exception of direct sampling by the Stardust mission that returned the dust samples from the short-period comet 81P/Wild 2 (Zolensky et al. 2006). A possible approach for investigating cometary samples is the study of cosmic dust. This method is likely the best and easiest technique because cosmic dust is accreted to the Earth in large quantities of ~40,000 tons annually (e.g., Love and Brownlee 1993; Taylor et al. 1998; Yada et al. 2004) and can be collected in Antarctica and Greenland as micrometeorites (MMs) and in the stratosphere as interplanetary dust particles (IDPs). Although IDPs are suggested to be mainly of cometary origin (e.g., Joswiak et al. 2005), the abundance ratios, or relative proportions, of asteroidal and cometary particles in MMs are not well known and thus need to be clarified.
Previous studies have focused on distinguishing between asteroidal and cometary particles based on their overall properties such as the combination of mineralogy, chemistry, isotopic compositions, and orbital parameters inferred from the degree of heating during atmospheric entry (e.g., Sandford and Bradley 1989; Love and Brownlee 1991, 1994; Nier and Schlutter 1992, 1993). Thermal alteration during atmospheric entry causes various effects on IDPs and MMs including a loss of volatile elements (Kehm et al. 2002) and noble gases (e.g., Nier and Schlutter 1992, 1993; Füri et al. 2013), decomposition and formation of secondary minerals (e.g., Greshake et al. 1998; Toppani et al. 2001; Nozaki et al. 2006) such as magnetite shells or envelopes (Keller et al. 1992; Kurat et al. 1994), and modification of infrared spectra (Sandford and Bradley 1989). The release pattern of noble gases is one of the most sensitive parameters in atmospheric entry heating because solar wind (SW) noble gases are implanted onto the surfaces of IDPs and MMs in interplanetary space. A comparison of noble gas release patterns and mineralogical characteristics enable discussion of the time in which the dust particles were heated, whether in the parent body, in space, or during atmospheric entry, and in addition the manner of heating revealed by peak temperature, heating duration, and oxygen fugacity.
However, combined research on noble gas analysis and mineralogical investigation has rarely been performed despite its usefulness because of the small size of samples, which creates difficulties in their handling. In addition, the resin used to fix samples could cause a false increase in the background levels of a noble-gas mass spectrometer as reactive organic interference and can create an atmospheric gas trap. As exceptions, Osawa et al. (2003a) and Bajo et al. (2011) performed scanning electron microscopy (SEM)/energy dispersive spectrometry (EDS) and synchrotron radiation X-ray diffraction (SR-XRD) analyses to characterize the appearance, bulk chemical composition, and bulk mineralogy of the particles prior to conducting noble gas measurements. Okazaki and Nakamura (2006) conducted SEM, electron probe microanalysis (EPMA), and secondary ion mass spectrometry (SIMS) analyses on the cross sections of individual Antarctic MMs in addition to observations of their appearances and bulk chemistry measurements by SEM/EDS prior to the noble-gas analysis. In this study, we used an acetone-soluble resin and changed the configuration of the focused ion beam (FIB) sectioning, which enabled us to perform multidiscipline analyses of MMs, particularly transmission electron microscopy (TEM) observation and noble-gas analysis. Here, we present the results of a combined study of the mineralogy and noble gas isotopes of MMs collected from Antarctic snow.
Detection of micrometeorites and SEM observation
Sample list with the degree of atmospheric entry heating, apparent sizes, and mineralogical features
Degree of atmospheric entry
Apparent size (μm)
Ol, LPx, Kam, Po, Mt
Ol, Mt, Sap, Serp
Ol, LPx, Po, Mt
Ol, Low-Ca cpx
Ol, LPx, Kam, Po, Pl, Mt
Low-Ca cpx, Ol, Chr
Powder diffraction patterns of the individual MM samples were determined by SR-XRD at the beam line 3A at the Photon Factory Institute of Material Science, High-Energy Accelerator Research Organization, Tsukuba, Japan. Each sample was mounted on a thin glass fiber of 5 μm in diameter by using acetone-soluble resin and was exposed to synchrotron X-rays in a Gandolfi camera that was evacuated to ~7 Pa to decrease the scattering of X-rays by the atmosphere. The X-rays were monochromated to 2.1655 ± 0.0009 Å and were concentrated by a short-gap undulator. The X-ray diffraction pattern was recorded on imaging plates read by an IP reader (Typhoon FLA7000) with a resolution of diffraction angle of 0.025°.
FIB microsectioning and TEM observation
Following the SR-XRD analysis, the samples were mounted on Mo strips using acetone-soluble resin and were cut off parallel to the wafers as a 100-nm-thick section by using the FIB-SEM (JEOL JIB-4501) at Ibaraki University for TEM. Different from normal FIB sample preparation procedures, the FIB sections were prepared parallel to the surface of substrate Mo plates, upon which each of the MMs was attached, for further micro-Raman and EPMA analyses. The proportion consumed by the FIB cutting was small, corresponding to only 5–10 and 10–20 % of each particle in volume and in surface area, respectively, assuming that a portion 6 μm thick from the top surface of a spherical body was consumed. TEM observation was performed by using JEOL JEM-2100 TEM equipped with Oxford INCA EDS at Ibaraki University and FEI Tecnai 20 F TEM equipped with EDAX Exact Genesis EDS at Kyushu University.
Micro-Raman spectroscopy and EPMA analysis
Because the surface of the remaining sample after FIB sample preparation was parallel to the Mo substrate, we were able to use Raman spectroscopy to identify the minerals on the surface by using a micro-Raman spectrometer (JASCO NRS-3100) at Ibaraki University. The excitation laser has a wavelength of 532 nm and a beam diameter of 2 μm. The exposure time on the samples and the power of the excitation laser were 80 s (40 s × 2) and 10 mW, respectively. The peak position of the Raman spectra was calibrated by using the strongest Raman shift peak of metallic Si at 520 cm−1.
After the micro-Raman analysis, the MMs were coated with carbon to determine their bulk and mineral chemistries. The chemical analyses were performed with a JEOL JXA-8530 F FE-EPMA at JEOL Co. The acceleration voltage and the probe current were 15 kV and 9 nA, respectively. ZAF-oxide correction was applied to calculate the chemical compositions of the minerals.
Noble gas analysis
Each MM sample mounted on a silicon wafer was removed by using acetone and was then washed separately by acetone, ethanol, and pure water. After weighing with an ultra-microbalance (Sartorius MSU-2.7S) at Kyushu University, individual MMs were wrapped with Al foil 10 μm in thickness. To remove adsorbed atmospheric gases, the samples were heated at 150 °C for 24 h in a sample holder connected to a gas purification line of the mass spectrometer. The heating temperature was determined considering that MMs and IDPs had been heated at a minimum of several hundred °C for several seconds (e.g., Love and Brownlee 1994). In addition, our samples consisted mainly of olivine with only small amounts of glass and quartz, in which He can diffuse faster (e.g., Kurz and Jenkins 1981) than that in olivine, from which He can hardly escape during the baking (e.g., Trull and Kurz 1993). This was supported by our noble gas results. For example, one MM consisting of hydrous minerals (Table 1), D03IB068, has a higher 4He/20Ne ratio and lower 4He release temperature than that of the other MMs, suggesting that preferential release of He by baking did not occur even in the least heated MM. Hence, the baking procedure did not affect the noble gas compositions of all of the MMs studied.
Each sample was dropped into a Mo crucible and was heated stepwise at 400, 650, and 1800 °C to extract noble gases in a “pot-pie” furnace designed for noble gas analysis of microgram samples. Isotopic ratios and amounts of noble gases were determined with the modified Micromass MM5400 noble-gas mass spectrometer at Kyushu University. Sensitivities and mass fractionation factors of the spectrometer were determined by measuring the known amounts of atmospheric gas and a He-standard gas with a 3He/4He ratio of 1.71 × 10−4. Repeated measurements of standard gases confirmed the stabilities of sensitivities of the mass spectrometer to be <10 %. The blank gas abundances were as follows: 4He = 5 × 10−12, 20Ne = 5–7 × 10−13, 36Ar = 1–2 × 10−12, 40Ar = 2–4 × 10−10, 84Kr = 1–3 × 10−14, 132Xe = 2–6 × 10−15 cm3 standard temperature and pressure (STP). The abundances of 20Ne and 22Ne were corrected for divalent interferences of 40Ar++ and 12C16O2 ++, which were around 5 × 10−14 and 1 × 10−14 cm3 STP during sample measurements, respectively.
Noble gas compositions and release patterns
Concentrations and isotopic ratios of micrometeorites collected from Antarctic snow
Sample ID (weight)
0.00023 ± 0.00012
13 ± 11
0.184 ± 0.062
215 ± 44
0.000240 ± 0.000017
11.77 ± 0.40
0.0258 ± 0.0094
0.204 ± 0.044
125 ± 40
0.000227 ± 0.000022
11.45 ± 0.33
0.0307 ± 0.0087
0.201 ± 0.013
214.5 ± 9.4
0.000236 ± 0.000014
11.54 ± 0.27
0.0294 ± 0.0068
0.200 ± 0.014
194 ± 12
0.193 ± 0.066
223 ± 31
0.00040 ± 0.00015
12.33 ± 1.18
0.028 ± 0.012
0.221 ± 0.064
216 ± 40
0.000450 ± 0.000060
11.78 ± 0.79
0.040 ± 0.020
0.187 ± 0.011
227.4 ± 9.7
0.000429 ± 0.000068
12.04 ± 0.70
0.035 ± 0.012
0.194 ± 0.016
225 ± 11
0.229 ± 0.084
195 ± 58
0.00026 ± 0.00016
9.9 ± 4.0
0.186 ± 0.054
153 ± 60
0.000249 ± 0.000015
11.41 ± 0.37
0.0255 ± 0.0058
0.1926 ± 0.0056
236.3 ± 5.9
0.000250 ± 0.000018
11.37 ± 0.37
0.0255 ± 0.0058
0.1940 ± 0.0079
227.4 ± 7.6
0.154 ± 0.071
294 ± 43
14.2 ± 2.8
0.173 ± 0.059
150 ± 47
0.000385 ± 0.000077
11.63 ± 0.53
0.0250 ± 0.0031
0.190 ± 0.010
188.0 ± 7.6
0.000385 ± 0.000077
11.73 ± 0.52
0.0250 ± 0.0031
0.185 ± 0.012
189.8 ± 8.9
0.173 ± 0.051
242 ± 28
0.00068 ± 0.00014
12.79 ± 0.71
0.0236 ± 0.0081
0.190 ± 0.027
93 ± 49
0.000900 ± 0.000035
12.84 ± 0.32
0.0291 ± 0.0036
0.187 ± 0.013
140 ± 12
0.000844 ± 0.000046
12.83 ± 0.29
0.0284 ± 0.0033
0.186 ± 0.011
141 ± 12
0.16 ± 0.11
238 ± 39
0.22 ± 0.10
137 ± 101
10.14 ± 0.95
0.029 ± 0.023
0.182 ± 0.015
266.3 ± 8.8
10.14 ± 0.95
0.029 ± 0.023
0.184 ± 0.019
251 ± 13
0.202 ± 0.073
236 ± 36
10.4 ± 7.2
0.218 ± 0.057
190 ± 50
10.88 ± 0.77
0.025 ± 0.014
0.196 ± 0.016
241 ± 11
10.85 ± 0.81
0.025 ± 0.014
0.201 ± 0.018
231 ± 13
Figure 2 shows the relationship between the 4He concentrations and the elemental ratios of 4He/20Ne. The elemental ratios of 4He/20Ne in IDPs, MMs (our data and unmelted MMs reported in Osawa and Nagao 2002, Osawa et al. 2003a), and cosmic spherules are lower than those of SW (Heber et al. 2009) and implantation-fractionated SW (IFSW), previously interpreted as solar energetic particles (SEPs), represented in lunar ilmenite I71501/5 (Benkert et al. 1993; the “Formation of secondary magnetite and diffusive loss of He during atmospheric entry heating” section). However, the ratios were higher than that of the terrestrial atmosphere (complied by Ozima and Podosek 2002). It is noteworthy that a positive correlation is present between 4He concentrations and 4He/20Ne ratios of the MMs in this study, as reported in previous studies (e.g., Nier and Schlutter 1990; Nagao et al. 2011). The severely heated MMs studied here and cosmic spherules (Osawa et al. 2003b) have lower 4He concentrations and 4He/20Ne ratios than those in the least heated MMs and IDPs (Nier and Schlutter 1990), which tend to have higher values (Fig. 2). The 4He/20Ne ratio of one severely heated MM, D10IB170, is almost identical to the atmospheric value (Fig. 2), although the Ne and Ar isotopic ratios in the 1800 °C fraction are different from the atmospheric values (Table 2). Another severely heated MM, D03IB057, also has a 4He/20Ne ratio similar to the atmospheric value (Fig. 2), although the 40Ar/36Ar ratios are distinct from the terrestrial value (Table 2).
Figure 3b shows 3He/4He versus 20Ne/22Ne ratios for the least and moderately heated MMs; the 3He/4He ratios could not be determined for the severely heated MMs (Table 2). The mixing lines shown in Fig. 3b were calculated by assuming the following values: (3He/4He)Air = 1.40 × 10−6, (20Ne/22Ne)Air = 9.8, and (4He/20Ne)Air = 0.319 for the terrestrial atmosphere (compiled by Ozima and Podosek 2002); (3He/4He)SW = 4.57 × 10−4, (20Ne/22Ne)SW = 13.78, and (4He/20Ne)SW = 656 for the present SW (Heber et al. 2009); (3He/4He)IFSW = 2.29 × 10−4, (20Ne/22Ne)IFSW = 11.2, and (4He/20Ne)IFSW = 390 for the IFSW in lunar ilmenite grains (Benkert et al. 1993); (3He/4He)Itokawa = 3.52 × 10−4, (20Ne/22Ne)Itokawa = 13.6, and (4He/20Ne)Itokawa = 110 for an Itokawa particle RA-QD02-0065 (Nagao et al. 2011); and (3He/4He)GCR = 0.30, (20Ne/22Ne)GCR = 0.80, and (4He/20Ne)GCR = 27 for a galactic cosmic ray (GCR)-produced component calculated for a spherical L-chondrite body with a 10-cm radius irradiated at a 0–1 cm depth (Leya and Masarik 2009). The He and Ne in all of the MMs studied are mixtures of normal SW and IFSW. Moreover, variable but minor contributions of the GCR-produced component are noted in D03IB068 and D10IB049 but not in D10IB130. The origin of the excess 3He in D10IB130 is discussed in the “Formation of secondary magnetite and diffusive loss of He during atmospheric entry heating” section.
Surface appearance and bulk mineralogy
Interior texture, mineral chemistry, and TEM observations
After identification and observation of the samples, cross sections were prepared from each MM by FIB and were observed by using FE-SEM. Three types of the textural features were visible in the cross sections of the MMs studied here including fine-grained polycrystalline in anhydrous D03IB067 and hydrous D03IB068, coarse-grained crystalline in D10IB049, D10IB130, and D10IB170, and chondrule-like cryptocrystalline in D10IB020 and D03IB057 (Table 1). Low-Ca pyroxenes in D10IB049 and D03IB057 were identified as low-Ca clinopyroxene (low-Ca cpx) based on the micro-Raman spectrum (Huang et al. 2000). Chromite was identified in D03IB057 by using micro-Raman spectroscopy (Wang et al. 2004). Some minerals identified by SR-XRD could not be identified by Raman spectroscopy because they are Raman inactive or were not found in the cross sections. Hereafter, we will describe the mineralogical features as compared with the degrees of heating during atmospheric entry inferred from the release patterns of noble gases (the “Noble gas compositions and release patterns” section).
Least heated MMs
Moderately heated MMs
D10IB049 consists mainly of low-Ca cpx. Other mineral phases identified with the SR-XRD are observed as fine-grained materials attached onto its original surface (indicated by arrows in Fig. 8b). Although this MM is classified to be coarse-grained based on FE-SEM observation (Table 1), it is an aggregate of fine-grained low-Ca cpx of <2 μm in size (Fig. 9b).
D10IB130 is a MM consisting of a single olivine crystal with uniform chemical composition (Fig. 7b). Arrays of ~1-μm-sized holes and some discrete 3–5-μm holes are visible in this MM (Fig. 8c). Dendrites of magnetite are visible at the original surface in abundance (Fig. 9c, d) just below the tungsten deposition (W depo. in Fig. 9c) and are developed on a large scale (typically ~100 nm) at the original surface and along cracks (Fig. 9d) and are embedded in glass (Fig. 9e). Each of dendritic magnetite grains appears to be ~10–20 nm in size (Fig. 9e).
Severely heated MMs
D10IB170 consists mainly of olivine containing numerous small magnetite inclusions, and glass patches occur within the interior (Figs. 10b and 11b); these results are highly similar to those reported by Wu and Kohlstedt (1988). The olivine in this MM is more ferroan (Fa28–31) than that in D10IB057 (Fig. 7b).
Peak temperature during atmospheric entry heating
Numerous studies conducted on the condition of heating during atmospheric entry have cited several factors such as particle size, entry angle, and entry speed (e.g., Love and Brownlee 1991, 1994). As discussed in previous studies (e.g., Nier and Schlutter 1992, 1993), the peak temperature can be estimated from the 50 % release temperature of 4He during stepped-heating extractions. In this study, we performed only three steps of extraction to obtain precise isotopic ratios, particularly of Ne; hence, the peak temperature cannot be determined from only the release profiles in most cases. Here, we estimate the peak temperature for the MMs studied based on the results of our noble-gas and mineralogical analyses, taking the results of previous numerical calculations and simulation experiments into consideration as well.
Least heated MMs
For the least heated MMs, the release patterns of 4He (Fig. 1) indicate peak temperatures of 400 to 650 °C because ~50 % of 4He is released by 650 °C, including 71 and 43 % for D03IB067 and D03IB068, respectively, as calculated from Table 2 (Nier and Schlutter 1992, 1993).
The hydrous polycrystalline MM (D03IB068) contains serpentine and saponite that would decompose to amorphous phases by brief heating at 600 and 700 °C, respectively (Nozaki et al. 2006). The decomposition of saponite could be preceded by shrinkage of interlayer spacing from 1.32 nm due to loss of water at 600 °C (Nozaki et al. 2006). Our TEM observation revealed that part of the saponite in this MM shrunk to 0.90 nm (Fig. 6b). Therefore, D03IB068 should have experienced atmospheric entry heating with a peak temperature of ~600 °C.
The other least heated MM, D03IB067, is composed of anhydrous minerals. Although the constituent minerals are less sensitive to thermal alteration than hydrous minerals, some constraint on the peak temperature is indicated by the presence of SF tracks in the olivine and pyroxene grains in this MM. Fraundorf et al. (1982) reported that SF tracks in olivine become difficult to detect after pulse heating in the range of 500–600 °C, whereas those in pyroxene can survive pulse heating at 625 °C but disappear after 700 °C. Considering these experimental results by Fraundorf et al. (1982), the peak temperature of D03IB067 is regarded as lower than ~600 °C.
Moderately and severely heated MMs
Based on the noble gas release profiles (Fig. 1), the moderately and severely heated MMs should have experienced atmospheric heating with peak temperatures >650 °C because the cumulative releases of 4He in the 400 and 650 °C fractions are much lower than 50 %. It is difficult to estimate the peak temperatures for each MM. If it is assumed that the moderately heated MMs have initially contained 4He at the maximum (saturated) concentration of 10−1–10−2 cm3 STP SW-4He/g in particles ~50 μm in diameter (e.g., Bogard 1977; Jull and Pillinger 1977; Signer et al. 1977), the present 4He concentrations correspond to 10−1–10−3 of the initial concentrations (Table 2). According to the results of stepped extraction experiments, most of noble gases are extracted below ~900 °C from lunar soil (Nier and Schlutter 1992; Fu et al. 2013) and synthetic olivine irradiated by 4He with 50-keV kinetic energy (Fu et al. 2013). Hence, the peak temperatures for the moderately heated MMs are likely between 650 and 900 °C. For the severely heated MMs, the observed concentrations of 4He are 10 to 100 times lower than those for the moderately heated MMs and correspond to 10−2–10−5 of the hypothetical initial concentrations. Hence, the peak temperatures of the severely heated MMs are >900 °C.
The mineralogical features determined by our TEM observation put constraints on the peak temperatures for the severely heated MMs. The appearance of olivine and glass at the outermost edge of D03IB057 suggests that this MM had been partly melted (Fig. 10a). The twinned clinopyroxene in the interior likely converted from proto-pyroxene stable at temperatures of about 1000–1450 °C, which depends on the Fe content (Huebner 1980). The mineralogical features of D03IB057 suggest that this MM experienced strong heating and rapid cooling. The textures of the magnetite and glass patches observed in D10IB170 are quite similar to that produced by heating of olivine under oxidized conditions (Wu and Kohlstedt 1988), suggesting that they were formed by decomposition of olivine under oxidizing conditions at temperatures higher than 1000 °C (Wu and Kohlstedt 1988; Gualtieri et al. 2003).
Formation of secondary magnetite and diffusive loss of He during atmospheric entry heating
As observed in the severely heated MMs, the presence of secondary magnetite is indicative of strong heating during atmospheric entry. However, a moderately heated MM, D10IB130, also contains abundant wormy secondary magnetite crystals (Fig. 9d, e). This mineralogical feature of D10IB130 appears to contradict the noble-gas characteristics. Here, we determine whether the magnetite formation by the atmospheric heating is compatible with the retention of He in this MM.
According to Wu and Kohlstedt (1988), magnetite formation is controlled by self-diffusion of Fe2+, which is comparable to Mg self-diffusion. As discussed above, D10IB130 appears to have been heated to 650–900 °C. By assuming the Mg diffusion rate shown by Wu and Kohlstedt (1988), the heating duration necessary to form the observed magnetite, assumed to be ~10 nm (Fig. 9e), was calculated to be 6500 and 100 s at 650 and 900 °C, respectively, under the ground atmosphere pressure. Such prolonged heating is impossible because particles with ~50-μm diameter and initial velocities of 15 and 20 km/s can experience peak heating only for 10–20 s at altitudes of 85–90 km (Love and Brownlee 1991). To reproduce 10-nm-sized magnetite within a short heating duration such as 20 s, a heating temperature of around 1100–1200 °C is necessary (Wu and Kohlstedt 1988), although this temperature range is inconsistent with the noble-gas release pattern of D10IB130 (Fig. 1). Even if the peak temperature expected for this MM based on the release pattern is underestimated, heating at 1100 °C for 10 s would release most of the SW noble gases implanted in the MMs because of the ~0.04-μm typical penetration depth of SW-He (1 keV/nucleon) and the typical diffusion length of He in olivine (0.1 μm for heating at 1100 °C for 10 s; Trull and Kurz 1993). In addition to diffusion, the loss of noble gases could take place by abrasion or evaporation of the outermost layer of MMs during atmospheric entry.
Even worse, the estimated duration for the magnetite formation should be the lower limit because the formation rate of magnetite is dependent also on the oxygen pressure. The oxygen pressure at the altitude (P fly) at which extraterrestrial particles have experienced the maximum heating temperature is calculated by P fly = ρv 2, where ρ and v are the oxygen density and the particle entry velocity at the relevant altitude, respectively. The P fly at 85–90 km of altitudes and with 15–20 km/s of the entry velocity is about one to two orders lower than the oxygen pressure on the ground, according to the density data provided by the United States Standard Atmosphere of 1976 and by assuming that the oxygen partial pressure does not depend on altitude. This calculation indicates that the magnetite formation at a high altitude would be slower than that on the ground.
Hence, some mechanisms are necessary for accelerating the oxidation of olivine to compensate for the discrepancy between the formation rate of magnetite and the loss rate of noble gases in a peculiar situation of atmospheric entry heating caused by interaction between hypervelocity particles and atmospheric molecules. The flow of O2 molecules encountering the surface of dust particles or active oxygen atoms and compounds might accelerate the formation rate of magnetite during atmospheric entry.
Excess 3He in a moderately heated coarse-grained MM
The isotopic ratios of He in most of the MMs studied are essentially identical to the SW values, 3He/4He = 4.64 × 10−4 (Heber et al. 2009), although the 3He/4He ratio of D10IB130, 8.44 × 10−4, is clearly higher than the SW ratio (Table 2 and Fig. 3b). Although it might be expected that the high 3He/4He ratio is the result of long exposure to GCRs (i.e., large contribution of cosmogenic isotopes), it is impossible to explain the 21Ne/22Ne ratio that is not indicative of cosmogenic 21Ne excess. In the following discussion, we will show that the He and Ne isotopic ratios in D10IB130 are explained by the contribution of solar energetic particles rather than by mixing of SW and cosmogenic noble-gas components. In the mixing calculation explained below, the 4He/20Ne ratios observed in the MMs do not have to be reproduced because the measured 4He/20Ne ratios of the MMs could have been altered and lowered (i.e., depleted in 4He) during the atmospheric heating to varying degrees. Moreover, the fractionation effects would be considerable in elemental ratios, as expected from Fig. 2, but small in isotopic ratios.
As shown in Fig. 3b, the isotopic ratios of He and Ne in the least and moderately heated MMs can be explained by mixing of normal SW and implantation fractionation in regard to both elemental and isotopic ratios (Grimberg et al. 2006, 2008; Wieler et al. 2007) of SW represented in lunar ilmenite, designated as IFSW in Fig. 3, except for D10IB130. Some Antarctic MMs show 3He/4He ratios larger than 5 × 10−4 (Osawa and Nagao 2002, Osawa et al. 2003a), which can be explained by mixing of normal SW and GCR-produced components, as clearly presented in Fig. 3b. Although the data for D03IB068 and D10IB049 slightly deviate from the mixing line between SW and IFSW, the deviation can be explained by the addition of a very small amount of a GCR-produced component, corresponding to about 5 Myr and 0.1 Myr exposure for D03IB068 and D10IB049, respectively. When a more fractionated 4He/20Ne ratio of ~100 is assumed for IFSW, the GCR exposure is not necessary to compensate for the deviation between the data for D03IB068 and D10IB049 and the mixing line. Such a highly fractionated SW (i.e., low 4He/20Ne of ~100 but 20Ne/22Ne similar to the IFSW value) is possible for olivine (e.g., Kiko et al. 1978; Signer et al. 1977; Wieler et al. 1980).
However, the isotopic composition of He and Ne in D10IB130 cannot be explained by the mixing of SW, IFSW, and GCR-produced components (Fig. 1b). When an elementally fractionated (i.e., depleted in He) but isotopically unfractionated component is assumed instead of the normal SW, the isotopic ratios of D10IB130 can be reproduced to some degree. Such an elementally fractionated SW has been observed in olivine grains recovered from the Itokawa asteroid by the Hayabusa spacecraft (Nagao et al. 2011). Two Itokawa particles, RA-QD02-0015 and RA-QD02-0065, have solar-like 3He/4He and 20Ne/21Ne/22Ne ratios but are associated with elementally fractionated 4He/20Ne ratios (110–127). Mixing of the elementally fractionated SW and a GCR-produced component reduces the degree of the 3He excess, but an excess corresponding to 150 Myr exposure to GCRs remains. Moreover, the mixing of the Itokawa particle composition and a GCR-produced component results in 21Ne excess (21Ne/22Ne of 0.08 at 3He/4He of 8.44 × 10−4 or 0.05 at 20Ne/22Ne of 12.83), contradicting the observation (Table 2 and Fig. 3a). The contribution of a solar cosmic ray (SCR)-produced component is excluded because the SCR-produced 3He/4He ratio of 0.03 (calculated from Trappitsch and Leya 2013) is lower than the GCR-produced value, which makes the situation worse.
Therefore, we argue that the 3He excess in D10IB130 originated from a component having solar-like Ne isotopic ratios and a higher 3He/4He ratio than the SW value. Such a component is expected in solar energetic particles emitted by gradual coronal mass ejection (CME) or impulsive SF events that enhance the 3He/4He ratio by a factor of 5 in CME events (Mason et al. 1999) or by factors of 250 to 80,000 in SF events (Reames 1998; Miller 1998). In this case, long exposure to cosmic rays is not necessary, and the noble-gas composition in this MM can be attributed to the mixing of a solar energetic particle component and a normal or fractionated SW component. To evaluate our argument, it is necessary to determine the precise Ne isotopic ratios and the He/Ne elemental ratio for the solar energetic particle component by space craft/satellite observations or by theoretical calculations in the future. To avoid misunderstanding, it is important to confirm that SEP noble gas is not a component related to SF or CME events but is merely normal SW fractionated during implantation (e.g., Grimberg et al. 2006, 2008; Wieler et al. 2007). The isotopic ratios of the SEP noble gas are enriched in heavier isotopes (i.e., lower 3He/4He and 20Ne/22Ne ratios compared with those of normal SW, as shown in Fig. 3), which is opposite that observed in D10IB130.
As discussed above, the high 3He/4He ratio of D10IB130 is likely indicative of the contribution of solar energetic particles in this MM. Similar or even higher 3He/4He ratios have been observed relatively often in IDPs (Pepin et al. 2000, 2001), although it is not conclusive because of the absence of Ne isotope data for these IDPs. In contrast, numerous studies on lunar soils have determined that there is no evidence for the contribution of solar energetic particles and that the isotopic ratios of solar wind have been quite constant (e.g., Heber et al. 2001). It is possible that the difference in duration of exposure to solar particles is responsible for the absence of 3He-rich signatures in lunar samples because shorter exposure durations of MMs and IDPs increase the odds of detecting such sudden events. The SW exposure duration of IDPs reported by Pepin et al. (2000) is calculated to be in the order of 100–10,000 years at 1 AU based on the 20Ne concentrations and the SW flux (Heber et al. 2009), which is shorter than those of lunar soils that have been exposed to SW typically for 104 − 105 years (Signer et al. 1977; Wieler et al. 1983). In addition to the duration of exposure, the magnetic field of the Earth influences the irradiation conditions on the lunar surface. The propagation ranges of protons and other ions in the impulsive SF events are narrow compared with those in the gradual CME events (e.g., Reames 1999); thus, more 3He-enriched SF particles were partly prevented from arriving at the Moon. Consequently, the 3He-rich signature observed in D10IB130 is more likely to originate from impulsive SF events rather than CME events.
Atmospheric entry velocity and origin of MMs
Based on the 50 % release temperature (Nier and Schlutter 1992, 1993), Joswiak et al. (2005) calculated the peak temperatures experienced by IDPs during atmospheric entry and divided IDPs in two categories as those having peak temperatures >700 and <500 °C. The IDPs studied by Joswiak et al. (2005) are within 5–15 μm in size; hence, they assumed that the peak temperatures are correlated mainly with atmospheric entry velocity. The particles with higher peak temperatures were believed to have velocities of ≥18 km/s, and those with lower temperature have ≤14 km/s. The differences in atmospheric entry velocities are considered to result from differences in their parent bodies, i.e., particles with higher velocities are from comets and those with lower velocities are from asteroids.
Considering our estimates for the peak temperatures discussed above, it is possible that the least heated MMs originated from asteroids and the severely heated MMs came from comets. Peak temperatures for the moderately MMs are intermediate between the criterion temperatures, i.e., between 500 and 700 °C; hence it is difficult to distinguish whether the moderately heated MMs originated from asteroids or comets based only on the peak temperature. Therefore, other criteria are needed to distinguish the origins of interplanetary particles. One possible source of information is the ratio of SW- and GCR-produced gases because the SW flux depends on the distance from the Sun whereas the GCR flux is virtually constant within the solar system. Differences in averaged distance from the Sun during interplanetary flights of particles could be determined as SW/GCR-produced abundance ratios (e.g., 20Ne-SW/21Ne-GCR ratios). Such a 21Ne-GCR signature can be observed only in IDPs and MMs with relatively long GCR exposure ages of > ~10 Myr (e.g., Pepin et al. 2000, 2001; Osawa et al. 2003b). As previously mentioned, longer GCR exposure ages are expected for cometary particles considering the Poynting–Robertson effect.
Seven MMs recovered from Antarctic snow are classified into three types based on their noble-gas release patterns reflecting the degree of heating during the atmospheric entry. This classification is consistent with the mineralogical and petrographic features of the MMs.
A moderately heated MM, D10IB130, shows a 3He/4He ratio of 8.44 × 10−4 but no cosmogenic 21Ne signature. This can be explained by the addition of a 3He-enrich component related to solar energetic particle events such as flare events. This result may be attributed to a short exposure to solar wind and/or a terrestrial magnetic field.
We are grateful to H. Motoyama, N. Imae, and the JARE teams for collection and transportation of the Antarctic snow. We thank A. Okubo for her assistance with the sample collection. Reviews by two anonymous reviewers resulted in significant improvements to this paper. This work was supported by Grant-in-Aid for Scientific Research (S) (No. 22224010, PI: H. Nagahara) and partly by a Grant-in-Aid for Young Scientists (A) (No. 23684046, PI: R. Okazaki).
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