1. Top of page
  2. Abstract
  3. Introduction
  4. Experimental
  5. Experimental Results
  6. Discussion and Conclusions
  7. References

Abstract– We used a combination of different analytical techniques to study particle W7190-D12 using microinfrared spectroscopy, micro-Raman spectroscopy, and field emission scanning electron microscopy (FESEM) energy dispersive X-ray spectroscopy (EDS). The particle consists mainly of hematite (α-Fe2O3) with considerable variations in structural disorder. It further contains amorphous (Na,K)-bearing Ca,Al-silicate and organic carbon. Iron-bearing spherules (<150 nm in diameter) cover the surface of this particle. At local sites of structural disorder at the hematite surface, the hematite spheres were reduced to FeO in the presence of organic carbons forming FeO-spheres. However, metallic Fe spheres cannot be excluded based on the available data. To the best of our knowledge, this particle is the first detection of such spherules at the surface of a stratospheric dust particle. Although there is no definitive evidence for an extraterrestrial origin of particle W7190-D12, we suggest that it could be an IDP that had moved away from the asteroid-forming region of the early solar system into the outer solar system of the accreting Kuiper Belt objects. After it was released from a Jupiter family comet, this particle became part of the zodiacal cloud. Atmospheric entry flash-heating caused (1) the formation of microenvironments of reduced iron oxide when indigenous carbon materials reacted with hematite covering its surface resulting in the formation of FeO-spheres and (2) Na-loss from Na,Al-plagioclase. The particle of this study, and other similar particles on this collector, may represent a potentially new type of nonchondritic IDPs associated with Jupiter family comets, although an origin in the asteroid belt cannot be ignored.


  1. Top of page
  2. Abstract
  3. Introduction
  4. Experimental
  5. Experimental Results
  6. Discussion and Conclusions
  7. References

For almost four decades, dust particles have been collected in the stratosphere between 17 and 19 km altitudes on flat-plate collectors that are carried aloft underneath the wings of high-flying aircraft (Brownlee 1978, 1985; Mackinnon et al. 1982; Zolensky et al. 1994; Rietmeijer 1998). The collected particles include interplanetary dust particles (IDPs), up to a few to hundred microns in size, that survived burning up in the mesosphere. With no exceptions, all of the collected IDPs experienced flash-heating when decelerating in the mesosphere. Specific rapid, thermally induced, modifications to some of the chemical and/or mineral properties of these IDPs are recognizable, and they are indicators of peak-heating temperatures (Flynn 1994; Rietmeijer 1998, 2002). About fifteen individuals, approximately 10–15 μm IDPs in size, will be collectable during an approximately 40 h collection period using a 30 cm2 flat-plate collector (Brownlee 1978, 1979, 1985). The main driving force behind these collections was, and still is, to have multiple low-cost sampling missions of comet and/or asteroid dust. Particles with a chondritic (solar) composition for the major elements (Mg, Al, Si, S, Ca, Fe, and Ni) and an aggregate structure are considered remnants of the least modified solar system materials. The collected chondritic IDPs showed systematic mineralogical differences that led to an infrared spectroscope classification scheme (Sandford and Walker 1985) and later to the acquisition of IR-reflectance spectra of individual chondritic porous IDPs (Bradley et al. 1992). Large nonchondritic IDPs ranging from 10 μm to approximately 60 μm in size are frequently collected, but they are infrequently analyzed. They are mostly forsterite and Fe-Ni sulfide IDPs (Christoffersen and Buseck 1986; Zolensky 1987; Schramm et al. 1989; Steele 1990; Zolensky and Barrett 1994; Rietmeijer 1996). Nonchondritic fragments of the much larger chondritic aggregate IDPs resemble the nonchondritic IDP grain sizes. Some of these particles are non-CI fragments of a cluster IDP (Thomas et al. 1995). They include Na-bearing, high-silica, and plagioclase-like particles as inferred from the reported compositions (Rietmeijer 1998). The data for the IDP-size fragments of the cluster IDP L008#5 (Thomas et al. 1995) lack many important details such as information about their amorphous or crystalline structure and if they are mixtures of minerals and glass, but the chemical and mineralogical variations among these fragments are quite large. Table 1 summarizes the types of nonchondritic IDPs that have been found.

Table 1.   Chondritic and nonchondritic IDPs, on average 10–15 μm in size, in the Earth’s lower stratosphere (modified after Rietmeijer 2002).
Chondritic aggregate IDPsNonchondritic IDPs (Chondritic aggregate material often adheres to surface)
A matrix of nanometer-scale amorphous silicates with variable amounts of embedded approximately 5 μm MG,FE- and Ca,Mg,Fe-silicates, Ni-free and low-Ni pyrrhotite, iron oxides 
 (1) Silicate IDPs, mostly Mg,Fe-silicates, and Mg(Fe),Ca,Al-Silicates
 (2) Sulfide IDPs, mostly NI-FREE and low-Ni pyrrhotite; rare pentlandite
 (3) Refractory, Ca,Ti,Al-rich IDPs
 (4) (Rare) plagioclase-like IDPs
 (5) IDPs that are admixtures of variable amounts of IDPs types 1–4
Cluster IDPs are mixtures of variable amounts of aggregate IDPs and nonchondritic IDPs (1)–(5) 

Iron oxides and Fe-oxyhydroxides, viz. magnetite, maghémite, hematite, and goethite, were found in several IDPs (Fraundorf 1981; Fraundorf et al. 1981; Rietmeijer et al. 1999; Rietmeijer 2002). Rotundi et al. (2007) reported micro-IR and micro-Raman measurements for five IDPs with evidence for the presence of maghémite and/or magnetite, including a very porous aggregate with maghémite and hematite that also showed the infrared silicate bands of olivine and pyroxene. Several Fe-oxide rich fragments ranging from 5 × 7 μm to 10 × 15 μm are also part of cluster IDP L0008#5 (Thomas et al. 1995). Weak Raman bands attributed to hematite and pyroxene-like infrared absorption features were previously detected in IDP 3U2 (Fraundorf et al. 1982). The pyroxene assignment was confirmed by EDX analysis. Electron diffraction of this particle supported the presence of magnetite (Fe3O4) or maghémite (α-Fe2O3) mixed with olivine. The D and G Raman bands associated with amorphous carbon were also detected in this IDP (Fraundorf et al. 1982).

Although not listed in Table 1, chondritic aggregate and cluster IDPs contain a wide range of organic compounds. Indeed, the 3.4 μm infrared feature allows the determination of the CH2/CH3 ratio in aliphatics (Matrajt et al. 2005). The D and G Raman bands allow the determination of the aromatic domain size. Thus, it was found that carbon in IDPs was mostly present as amorphous carbon (a-C), or hydrogenated amorphous carbon (a-C:H) for particles with a significant 3.4 μm infrared feature (Muñoz Caro et al. 2006). Such material corresponds to poorly graphitized carbon with an aromatic domain size between 1.1 and 1.6 nm (approximately 20–40 rings) that are either linked by aliphatic chains with CH2/CH3 ratios varying from 2.8 to 5.5 in IDPs containing a-C:H, or a carbon sp3-skeleton in IDPs containing a-C amorphous carbon (Muñoz Caro et al. 2006).

Particles in the same size range as IDPs on dust collectors can also be to linked to terrestrial sources such as volcanic dust injected into the stratosphere. These particles have nonchondritic compositions and lack the telltale signs of flash-heating, for example, a (partial) magnetite and/or maghémite rim and/or disseminated crystals of these oxides. However, from the lack of such signs, an extraterrestrial source cannot be excluded. A case in point would be the hydrated, low-Ni, nonchondritic stratospheric dust particles (Rietmeijer 1992) that could be a new type of IDPs of lunar, Martian, or differentiated asteroid origin (Flynn and Sutton 1990, 1991).

We report here the results from combined micro-FTIR, micro-Raman, and FESEM-EDX analyses of the irregularly shaped, 10 × 9 μm, particle W7190-D12 collected over North America between July 17 and October 3, 1996 during accumulated period of 40 h. In Cosmic Dust Catalog volume 17, it is described as being opaque with a dull black luster. Those are properties commonly ascribed to stratospheric dust particles of “cosmic (C)” origin. From this catalog we quote, “particle type ‘Cosmic’ is used to conveniently group together all particles, which are judged to be of extraterrestrial origin, including those that have apparently experienced strong ablation heating or melting.” The smooth structures along the collected particle perimeter (Fig. 1) could be a macroscopic ablation feature. Particles with this particular morphology in other Cosmic Dust Catalogs were called chondritic rough IDPs (Rietmeijer and Warren 1994) that included IDPs with the distinctive petrological properties of hydrated CM-meteorite matrix (Bradley and Brownlee 1991; Rietmeijer 1996). On this stratospheric dust collector, 23 particles were identified as C-type particles, and 14 of them contained Si and Ca as the main element components, including particle W7190-D12. A majority of C-type IDPs listed in this particular catalog have an EDX spectrum that is qualitatively similar to that of particle W7190-D12. None of these particles are listed as possible fragments of cluster IDPs that might indicate an extraterrestrial origin. The high abundance of these particles on the collector in the lower stratosphere at the time of collection might suggest that these nonchondritic particles were part of a dust swarm, e.g., a dust cloud linked a bolide event (Klekociuk et al. 2005) or a volcanic dust cloud. However, the SEAN bulletins report no volcanic events during the time period between January and October 1, 1996, that ejected debris that could have reached the lower stratosphere across the North American airspace at the time of particle collection. Still, it is common to find volcanic ash in the stratosphere with no apparent recent volcanic eruption (Zolensky et al. 1989). However, this observation is limited to aerosol particles, i.e., sulfuric acid aerosol droplets, during the arctic winter period. It does not include approximately 10 μm sized solid dust particles. There is no record whether these particles were collected from a single catastrophic event or were gradually collected during the entire 40 h of exposure time. Although there is no definitive evidence for an extraterrestrial origin of these particles, we suggest that particle W7190-D12 and other similar particles on this collector were from the Zodiacal dust cloud and may be nonchondritic debris from a Jupiter family comet. If so, particle W7190-D12 represents another new type of nonchondritic IDPs.


Figure 1.  Scanning electron microscopy image of particle W7190-D12. This image was performed at the Center of Astrobiology using a JEOL JSM-5600LV instrument. This low vacuum image was obtained using the backscattered electron detection mode at an acceleration potential of 20 kV and 10 mm working distance.

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  1. Top of page
  2. Abstract
  3. Introduction
  4. Experimental
  5. Experimental Results
  6. Discussion and Conclusions
  7. References

Sample Preparation

Curatorial handling procedures can be found in MacKinnon et al. (1982) and Zolensky et al. (1994), and in the NASA-JSC Cosmic Dust Catalogs. There is no evidence supporting that new minerals or compounds were formed during processing and handling in the NASA/JSC Curatorial Facility. Subsequently, the particle was examined under an optical microscope in a laminar flow class-10000 clean room at the Institut d’Astrophysique Spatiale (Orsay, France). It was rinsed 5 times with hexane (Aldrich) to remove any residual silicon oil coating of the collector applied to entrap particles. It is conceivable that the hexane rinse may have removed some fraction of indigenous organics, but it is less likely that indigenous organics were modified during this process because hexane is quite stable, nonpolar, and nonreactive in most cases. It was then transferred to a KRS-5 infrared transparent window of thallium bromo-iodide using a hairbrush mounted on a micromanipulator. Once pressed onto the KRS-5 window, the particle covered an area of roughly 10 μm by 20 μm.

Sample Analysis

The analyses were performed in sequential order (1) nondestructive FTIR spectroscopy, (2) Raman spectroscopy at very low laser power less than 1 mW, and (3) FESEM-EDX analyses.

Micro-FTIR Analysis

Fourier Transform Infrared spectroscopy was conducted using a Nicolet Magna-IR 560 ESP [in full] spectrometer coupled to a Nicolet Nicplan infrared microscope, in line with a Synchrotron Radiation Source located on line SU5 at the Laboratoire pour l’Utilisation du Rayonnement Electromagnétique (LURE) at the University of Paris-Sud, Orsay, France. The spectrometer is equipped with a KBr beamsplitter and a nitrogen-cooled Mercury-Cadmium-Tellure (MCT) detector operating in the 4000–650 cm−1 (2.5–15 μm) spectral range. The microscope used a Schwarzchild-Cassegrain objective (32×) and a condenser (10×) in transmission mode. The analyses were performed using an internal black body-like source (Globar). To enhance the signal-to-noise ratio and minimize background fluctuations due to purge variations, one hundred acquisitions of 128 scans each at 2 cm−1 resolution were recorded and averaged. Integration of the absorbance area of the features was performed using an inhouse Interactive Data Language (IDL) code. The column densities were determined using the formula

  • image(1)

where N is the column density in cm−2, τ the optical depth of the band, dν the wave number differential in cm−1, and A the band strength in cm molecule−1.

We calculated the column density of aliphatic C atoms cm−2 as N(aliph. C) = N(CH2) + N(CH3), where the absorbance areas of the peaks at 2958 and 2920 cm−1 (asymmetric CH3 and CH2 stretching modes) are calculated using a Gaussian to fit the data. The adopted band strengths, A(CH3) = 1.25 × 10−17 cm (C atom)−1 and A(CH2) = 8.4 × 10−18 cm (C atom)−1, are those for hexane ice from Dartois et al. (2004). For evaluation of N(Si), i.e., the column density of the silicate band around approximately 1000 cm−1, we adopted A(Si) = 2.0 × 10−16 cm (Si atom)−1 (Matrajt et al. 2005). We obtained three infrared transmittance spectra, viz. (1) the entire particle, (2) the top half covering the area of the EDX spots 2, 5, and 6 (Fig. 1) and (3) the bottom half with EDX spots 1, 3, and 4 (see Figs. 2 and 3). Infrared beam scans across the particle and the particle absorbance was measured.


Figure 2.  Midinfrared spectra of particle W7190-D12 before (top panel) and after (bottom panel) baseline subtraction.

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Figure 3.  Top panel: Photograph of particle W7190-D12. Bottom panel: Infrared spectra corresponding to the top and bottom halves of the particle. D12 (top) spectrum was offset upward for clarity.

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Micro-Raman Analysis

Raman spectra were recorded using a Dilor XY confocal Raman microspectrometer at the Laboratoire de Sciences de la Terre (Labram HR800 vis Jobin Yvon) coupled to a Spectra Physics Argon+ laser tuned for these experiments at a wavelength of 514.5 nm (green). The laser beam was focused through microscope objectives (×50) to a 2 μm spot on the sample; the backscattered light was collected through the same objective. The resolution was λ/d(λ) ≈ 600. Wavelength calibration was performed using a silicon wafer sample. The wavelength deviations from the standard values were less than 1 cm−1 for all measurements. The laser power was kept well below 1 mW during 0.5–20 min of acquisition. We investigated the spectral region 200–3200 cm−1 that includes the first-order D (disorder line, approximately 1360 cm−1) and G (graphitic line, approximately 1580 cm−1) bands of carbonaceous materials, and the equivalent of the infrared 3.4 μm band at 2800–3000 cm−1 for hydrogenated carbon-bearing materials (Wopenka 1988; Ferini et al. 2004; Muñoz Caro et al. 2006). The Raman bands for minerals are also within the 200–3200 cm−1 spectral range. It is known that Fe-oxyhydroxides could be transformed to hematite during Raman analysis when the laser power is 7 mW or higher (de Faria et al. 1997). We chose our experimental conditions carefully to prevent this potential laboratory-induced artifact by using <1 mW laser power. Thus, any Fe-oxyhydroxides that might have been present could not be transformed to hematite.

FESEM Analysis

Field emission scanning electron microscopy was performed using a FESEM JSM-6500F equipped with an Oxford INCA EDS microanalyzer located at the Department of Materials and Structures (INTA). It was operated in the backscattered electron (BSE) detection mode at an acceleration potential of 20 kV and working distance of 9.4 mm for EDS analysis. The backscattered signal is strongly dependent on the mean atomic number of the sample (Joy 1991). As heavy elements (high atomic number) backscatter electrons more strongly than light elements (low atomic number), they will appear brighter in BSE images used to detect areas of different chemical compositions. Internal standards for the EDS spot analysis were used including a standard hematite sample from the Rio Tinto site (Spain) that was pure hematite based on X-ray diffraction analysis. The accuracy of the FESEM-EDS analyses is potentially affected by surface roughness and intrinsic submicron scale chemical heterogeneity of the particle. To test this, we repeated the measurements in different spots located close to the original spot. For all the elements, variations in the abundance values for close spots were less than 5%. The locations of these analyses are shown in a SEM image of the particle (Fig. 1).

Experimental Results

  1. Top of page
  2. Abstract
  3. Introduction
  4. Experimental
  5. Experimental Results
  6. Discussion and Conclusions
  7. References

Micro-FTIR Analysis

Figure 2 shows the infrared spectrum corresponding to the absorbance of the whole particle before (top panel) and after (bottom panel) baseline correction. The column densities of the relevant absorption bands in the spectrum (Fig. 2) are given in Table 2. The absorption band between approximately 3600 and 3100 cm−1 can be attributed either to a trace of adsorbed water to the sample or to constitutional OH groups present in mineral and/or carbon materials. The 3.4 μm feature of the CH-stretch of aliphatics with subfeatures at 2953, 2915, and 2848 cm−1 is present. A similar band is commonly observed in chondritic IDP infrared spectra (e.g., Flynn et al. 2003; Matrajt et al. 2005) and was associated with hydrogenated amorphous carbon, abbreviated to a-C:H (Muñoz Caro et al. 2006). The value N(CH2)/N(CH3) = 3.5 falls within the range of chondritic IDPs (Muñoz Caro et al. 2006). The prominent broad band peaking at approximately 983 cm−1 is due to the Si-O silicate stretching mode in amorphous silicates. The ratio N(aliph. C)/N(Si) approximately 0.7 is high compared with chondritic IDP values reported in the literature (Muñoz Caro et al. 2006). Unfortunately, we could not obtain a quantitative determination of the hematite content of this particle because the diagnostic infrared bands for hematite are located below 700 cm−1, which falls outside the spectral range of the instrument that was used. The band around 1415 cm−1 falls in the spectral region where weaker absorptions associated with the CH bending modes of aliphatics are expected, but it could also show a contribution from carbonates (Bradley et al. 1992), although the EDS data do not confirm carbonates in this particle.

Table 2.   Column densities for particle W7190-D12 of the CH2 and CH3 groups, total aliphatic carbon, and silicate Si atoms [respectively N(CH2), N(CH3), N(aliph. C), and N(Si)], calculated after baseline correction.
N(CH2)aN(CH3)bN(aliph. C)cN(Si)d
  1. aA = 8.4 × 10−18 cm (CH2 group)−1, Dartois et al. (2004).

  2. bA = 1.25 × 10−17 cm (CH3 group)−1, Dartois et al. (2004).

  3. cN(aliph. C) = N(CH2) + N(CH3).

  4. dA = 2.0 × 10−16 cm (Si atom)−1, Matrajt et al. (2005) and ref. therein.

CH2 cm−2CH3 cm−2aliph. C cm−2Si cm−2
4.9 × 10161.4 × 10166.3 × 10168.6 × 1016

The absorption feature around 1612 cm−1 could be due in part to an aromatic component (Pouchert 1997), but we found no indication for aromatics using micro-Raman spectroscopy, see the Micro-Raman Analysis section. The C=O stretching of carbonyl groups in organic materials could also contribute to the band peaking at 1612 cm−1. Finally, the apparent absorption features observed in the middle of the spectrum with no assigned frequency are probably artifacts due to reflections and other optical effects. The top panel of Fig. 3 is a photograph made with the optical microscope of the Raman spectrometer; the bottom panel of Fig. 3 shows the infrared spectra collected on the top and bottom halves of the particle. The top half covers FESEM-EDX spots 2, 5, and 6 of Fig. 1 and the bottom half covers spots 1, 3, and 4 (Fig. 1). There are some minor differences between both spectra. They are mainly due to changes in band profiles and the absence of the 1415 cm−1 band in the bottom spectrum. Still, they indicate a rather uniform bulk-infrared silicate composition of the particle.

Micro-Raman Analysis

The expected phonon lines in the Raman spectrum of hematite are the A1g modes at 225 and 498 cm−1, and the Eg modes at 247, 293, 299, 412, and 613 cm−1 (de Faria et al. 1997). The broad band that appears around 1320 cm−1 is assigned to a 2nd harmonic vibration (e.g., de Faria et al. 1997). The spectrum of pure, crystalline hematite does not show a band around 660 cm−1, but in reality, this band is observed in many hematite spectra (Zoppi et al. 2005). Recent studies assigned this band to either a lack of long-range order in hematite or the presence of impurities (Zoppi et al. 2005), although it is commonly referred to as the disorder band. The intensity ratio of the 615 and 660 cm−1 peaks can be an indicator of disorder or impurities in hematite (Zoppi et al. 2005). The top panel of Fig. 4 shows the Raman spectrum of a well-crystallized hematite standard with quite sharp bands except that the presence of the 660 cm−1 band indicates a small amount of disorder. The other three spectra correspond to three different regions located at the center, top, and bottom of this particle; they are labeled R(center), R(top), and R(bottom) (Fig. 1). Compared with the standard hematite spectrum, all three spectra show low relative intensities of the approximately 225 and 292 cm−1 bands. The (center) spectrum has the largest 611 cm−1/655 cm−1 intensity ratio, which indicates that hematite at this location is less disordered and/or has fewer impurities compared with the top and especially the bottom regions of the particle. The absence of a peak around 500 cm−1 (Fig. 4, top panel, top spectrum) was noted in hematite samples containing (unspecified) feldspars (Zoppi et al. 2005). It suggests that (unspecified) feldspars are present with hematite in this part of the particle.


Figure 4.  Raman spectroscopy of particle D12. Top panel: Top spectrum corresponds to a standard of well-crystallized hematite (sharp bands) in the 200–1500 cm−1 range. Bottom spectra correspond to three regions of particle W7190-D12 for 30s integration. One spectrum was taken at the center of the particle, where the particle displayed a darker color; the other two spectra were taken respectively at the top and bottom regions of the particle where the color was lighter. The color contrast in the particle is best appreciated in the SEM image of Fig. 1. Spectra were offset for clarity. Bottom panel: Spectra of the same three regions of particle D12 in the 800–3200 cm−1 range. Each spectrum corresponds to 20 min integration. The hematite band around 1320 cm−1 is observed. The small band in the region 2800–3000 cm−1 is due to aliphatic species. Spectra were offset for clarity.

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As mentioned in the Micro-Raman Analysis section, previous work showed that with laser power below 1 mW hematite and other Fe-mineral are not modified (de Faria et al. 1997). The positions of the Raman peaks for particle W7190-D12 displayed in Fig. 4 match well with the spectrum of hematite measured at 0.7 mW (see table 1 of de Faria et al. 1997), while the hematite spectrum at 7 mW in de Faria et al. (1997) (the minimum laser power required to convert other iron oxides and oxyhydroxides into hematite according to these authors) displays shifted peak positions and significantly broader FWHM. For instance, the peak at about 292 cm−1 is shifted to 283 cm−1 when acquired at 7 mW laser power. In addition, peaks associated with other iron oxides or oxy-hydroxides are absent in the Raman spectra of the particle; only hematite is observed. We conclude that (1) other Fe-minerals in the Raman spectrum of this particle that could have led to hematite formation are absent, and that (2) the peak positions in the Raman spectrum of this particle are similar to those of hematite measured at laser power below 0.7 mW (no laser alteration), and clearly differ from the peak positions of altered hematite measured at a laser power of 7 mW. If the hematite detected in the particle was the result of alteration during Raman analysis of a different Fe-mineral, the observation (1) above would require a large laser alteration leading to a nearly complete transformation of the original Fe-mineral into hematite. Observation (2) indicates that such laser alteration leading to a shift in the Raman peaks did not occur. Therefore, the hematite detected in the particle is not a product of laser-induced alteration as was expected for laser power below 1 mW. It is indigenous to the particle.

The bottom panel of Fig. 4 extends the spectral range to 3200 cm−1. The hematite peak around 1320 cm−1 (Fig. 4, top panel) is shown for comparison. The Raman equivalent of the infrared 3.4 μm feature is observed in the range 2800–3000 cm−1. The absence of the D band in the spectra indicates that there are no aromatic rings. Only a very weak G band around 1600 cm−1 is observed (Fig. 4, bottom panel, D12 [top] spectrum), which indicates the presence of sp2 paired carbon atoms (C=C stretching mode) in olefinic chains (Ferrari 2002).

FESEM Analysis

Main Particle

The particle shows a rectilinear core of a compact but thin-sheeted material with irregular outlines with small fragments of the same covering the more continuous sheets. It is embedded in material that has a very smooth surface, but appears to be coarsely granular. This material becomes more platy when in direct contact with the core material. The EDS analyses (Table 3) and the Fe- and CI-normalized element abundances (Fig. 5) show (1) the nonchondritic composition of this particle; (2) different compositions of the core and smooth embedding material; (3) carbon, oxygen, and iron that are present across the entire particle; and (4) Na, Al, Si, and Ca that are limited to the coarsely granular and smooth material. Trace amounts (≤1 atom%) of Mg are randomly present across the particle. Trace amounts of K and Ti were found in the coarsely granular smooth material represented by analysis spot 5 (Fig. 1). Comparing the EDS data measured across the particle with the pure hematite composition (Table 3) supports that the core (spots 1, 3, and 4) is pure hematite, which is consistent with the micro-Raman identification of hematite in this particle. This identification is also consistent with the micaceous (i.e., sheeted) habit of hematite iron rose. These three analyses and analysis #6 (Table 3) indicate the presence of a C,O-compound with C/O ≈ 1. A closer look at the data in Table 3 shows that each analysis represents a mixture of hematite, a C,O-compound, and an Na,Al-aluminosilica compound that is the embedding material. This particular mixture describes the entire particle and is consistent with the spectral data that organic carbon covers the entire surface, and that the only variations correspond to the relative proportions of hematite and an Na,Ca-aluminosilica compound. The data for spot 4 located at the interface of the core and embedding material are consistent with FeO instead of hematite (Fe2O3). This area also has spherules at the surface.

Table 3.   Renormalized element abundances (atom%) obtained at an accelerating voltage of 20 keV for six selected spots on particle W7190-D12 that are identified in Fig. 1 after removal of I and Tl contributed by the KRS-5 window with an approximate Tl50Br20-I30 formula.
  1. Notes: nd = not detected; tr = trace only.

  2. aHematite sample from the Rio Tinto region in Huelva (Spain).


Figure 5.  Fe- and CI-normalized abundances measured in particle W7190-D12 (Fig. 1), viz. in an area of hematite with a rough surface (spots 1 and 3) (gray symbols) and areas with a smooth surface (spots 2, 5, 6) (black symbols) corresponding to the silicate component. Spot 4 is the smooth periphery that developed during atmospheric entry. It is probably a mixture of silicates and hematite. The CI data are from Anders and Grevesse (1989).

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The C,O-compound with C/O ≈ 1 was inferred from the measured element abundances only. The micro-FTIR data showed spectral signatures that match the CH-stretch and CH-bending modes of aliphatic carbons, bands that might be due to constitutional OH groups present in mineral and/or carbon materials and to C=O (carbonyl) groups that would be likely attached to the aliphatic chains. Micro-Raman data found a very weak G band indicating the presence of sp2 paired carbon atoms (C=C stretching mode) in olefinic chains. The measured oxygen may be associated with these organic carbons and with the adsorption of water by the sample.

The presence of Na, Ca, Al, Si (Table 3), and oxygen suggests the presence of a Na,Ca-aluminosilica compound. It is not a stoichiometric chemical mixture of albite (NaAlSi3O8)––plagioclase (CaAl2Si2O8). The micro-FTIR data support an amorphous silicate, which removes the constraints of chemical stoichiometry, viz. Na:Al:Si = 1–1.3:1–4:1–8. For example, the composition of spot 4 (Table 3) suggests that a structural formula would be consistent with Na0.5AlSi3O8, i.e., Na-deficient albite. The amorphous Na-aluminosilicates seem to show increasing Na-loss from albite while keeping the silicate framework intact. This is not the case for spot 6 where the amorphous Ca,Al-silicate composition is reminiscent of a deep metastable eutectic calciosilica compound with approximately 85 wt% CaO and approximately 15 wt% SiO2 (Rietmeijer et al. 2008). Such deep metastable compositions are nonequilibrium features that form after rapid quenching from high-temperature vapor or melt. The low amount of Al (Table 3) could easily be incorporated into this amorphous high-Ca calciosilica compound.

The element abundances presented in this paper match qualitatively with those shown in the EDS spectrum for particle W7190-D12 published in Cosmic Dust Catalog 17. The published spectrum shows major peaks for Si and Ca and only a very small Fe peak. The latter seems to be inconsistent with the data in Table 3 that show a considerable quantity of iron across the particle. The strong Si and Ca peaks in the published spectrum are also at odds with the findings in this study. Analysis #6 would be the “best match” of the catalog spectrum and the data in Table 3. If that were correct, the inferred amorphous Ca,Al-silicate composition for the smooth material in this location would be incorrect. Instead, it could suggest amorphous Ca-silicate with a larnite (Ca2SiO4) or wollastonite (CaSiO3) composition, but without presumption of mineral stoichiometry. There is still the possibility that the catalog spectrum and the data in Table 3 are for different sides of this massive particle. The SEM image of this particle in Cosmic Dust Catalog 17 shows a cluster of three granular features on top of a platy grain with straight edges with a fourth grain attached along the edge. There are also many smaller grains scattered on top of the platy grain. It is not clear how this particle morphology matches the image shown in Fig. 1 in this article. The platy grain is probably the hematite grain.

Assuming Si is in the form of Na,Al-silicates, the abundance of Si in the particle can be roughly estimated from the N(Si) = 8.6 × 1016 Si cm−2 value measured by FTIR spectroscopy (see the Micro-FTIR Analysis section). For simplification, we use the elemental abundances Na:Al:Si:O = 1:2:4:8 (Table 3), assuming two O atoms per Si atom in the mineral fraction that corresponds to 317.4 g per mol of material, and an amorphous silicate density of 3 g cm−3, which gives an average thickness of the amorphous silicate component around 0.04 μm. Unfortunately, the N(Fe) value for hematite could not be measured by FTIR. On the basis of the silicate column density and the elemental Fe/Si ratio around 30 measured by FESEM at the center of the particle on spots 1, 3, and 4, we obtain N(Fe) = 2.6 × 1018 Fe cm−2, or 1.3 × 1018 Fe2O3 molecules cm−2 if all the iron is in hematite. For a hematite density of 5.3 g cm−3, and a molecular weight of 159.7 g, the average thickness of hematite in the particle is roughly 0.7 μm. The aliphatic carbon component for which N(aliph. C) = 6.3 × 1016 C cm−2 with a density of roughly 1 g cm−3 corresponds to an average thickness of 0.02 μm.


A FESEM-BSE image of the particle (Fig. 6, top panel) reveals the presence of spherules (Fig. 6, bottom panel). The spherules are distributed randomly across the surface of this particle. They display variable sizes in the range 30–150 nm. There might be spherules smaller than 30 nm, but they would not stand out at ×19,000 magnifying power. Some small areas around the center of the particle contained the largest number of iron spherules, including the location of EDS spot 1. One of these spots is shown on the bottom panel of Fig. 6. Their brighter appearance and a higher iron density than in the hematite standard suggest that they could be FeO or pure iron. Magnetite was not considered because its backscatter coefficient, and therefore its “brightness” in the BSE image, is very similar to that of hematite (Andersen et al. 2009). From our data, we can only assert that the spherules are made either of FeO or pure Fe, or a combination of both. No FeO was directly detected by Raman spectroscopy, but the EDS data for spot 4 are consistent with the presence of FeO in this particle.


Figure 6.  Top panel: Electron backscattered (FESEM-BSE) image (×4300) of the entire particle D12. This particle was crushed onto the infrared transparent window to allow spectroscopy, which probably affected the original distribution of the Fe-spheres on the particle surface. The Fe-spheres do not appear to be limited to hematite-rich areas and are also found on silicate-rich areas. The spots outside the particle are due to the texture of the infrared transparent window. Bottom panel: FESEM-BSE image with a larger magnifying power (×19,000) showing only a small section of particle D12, located at the center-right position in top panel (this region is labeled as “1” in Fig. 1, indicating the position at which elemental analysis was performed using the FESEM-EDS instrument, see results in Table 2, Spot 1). The dark background corresponds to hematite with randomly distributed 30–150 nm sized FeO and/or Fe spherules at its surface. The lighter color on the top right corner of this image corresponds to a silicate grain.

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Discussion and Conclusions

  1. Top of page
  2. Abstract
  3. Introduction
  4. Experimental
  5. Experimental Results
  6. Discussion and Conclusions
  7. References

Based on its optical and chemical attributes, particle W7190-D12 was classified as a C-type particle (p. 67 of the particle catalog at This is not a definitive classification. Especially in the case of a compact nonchondritic particle, additional data are required such as stable isotope analyses before a cosmic origin can be accepted. The provisional identifying criteria listed in the Cosmic Dust Catalogs make no allowance for Al- and/or Ca-rich particles such as (rare) refractory IDPs (Zolensky 1987), but stable oxygen isotopes of such a particle proved its extraterrestrial origin (McKeegan 1987). Nonchondritic IDPs can be compact mineral agglomerates and mono-mineralic particles often with fine-grained aggregate material that resembles the fine-grained matrix material of chondritic aggregate IDPs attached to the surface (Rietmeijer 2002). If such fine-grained aggregate matrix material had been attached to particle W7190-D12, it would have been melted during atmospheric entry forming the smooth material along its perimeter as was observed for sulfide IDPs (Rietmeijer 2004). No such material was present on this particle. Nor does this particle show a partial or complete Fe3O4 rim. In the absence of obvious flash-heating indicators (Rietmeijer 1998) and stable isotope data, proving an extraterrestrial origin for this particle will be challenging and in fact will ultimately rely on proof of flash-heating.

Consistency with an Extraterrestrial Origin

This study of combined micro-FTIR and micro-Raman spectroscopy, and FESEM-EDS of particle W7190-D12 that measures 10 μm × 9 μm shows an aggregate of a hematite grain partially embedded in amorphous Na,Ca-aluminosilica material and with probably similar grains also attached to its surface. The surface of this material is quite smooth. An organic carbon layer covers the particle surface that was analyzed in this study. It is undoubtedly a natural material.

  • 1
     Particle morphology: a platy (catalog image) or somewhat blocky (Fig. 1) hematite grain coated with smooth amorphous feldspathic material could be volcanic dust. The amorphous Na,Ca-aluminosilica material is then volcanic glass. The absence of recorded volcanic activity prior to its collection in the lower stratosphere seems to rule out, but does not conclusively prove, that it is not terrestrial volcanic dust. Other nonchondritic particles collected in the lower stratosphere have been linked to nonterrestrial sources, but those particles were not (partially) covered or associated with amorphous Na,Ca-aluminosilica.
  • 2
     Hematite: If this particle is in fact an IDP, it will be the first occurrence of this mineral among these collected materials, but it would not be the first nonchondritic IDP. Hematite-magnetite mixtures were found in Stardust samples from comet Wild 2 (Bridges et al. 2010), but they were much smaller than the grain in particle W7190-D12.
  • 3
     Amorphous Na,Ca-aluminosilicate: Amorphous albite-anorthite and other nonstoichiometric Na,Ca-aluminosilica materials that frequently contain minor amounts of Mg and Fe are present in chondritic aggregate IDPs (Rietmeijer 1991, 1998), while amorphous Ca-aluminosilicates are present both in chondritic aggregate IDPs and in samples of Stardust glass (Rietmeijer 2009). The Si-O Stardust glass contains variable amounts of Mg, Al and Ca (Leroux et al. 2008; Tomeoka et al. 2008; Rietmeijer 2009a) and Si-O-Al glass has variable minor amounts of Na, Mg, and K (Tomeoka et al. 2008). The prominent broad band peaking at approximately 983 cm−1 is due to the Si-O stretching mode of amorphous silicates in particle W7190-D12.
  • 4
     Organic carbons: The aliphatic component in particle W7190-D12 resembles this component found in most chondritic IDPs. The ratio N(CH2)/N(CH3) = 3.5 is within the range found in chondritic IDPs.

This tabulation shows that particle W7190-D12 shares common properties with chondritic aggregate IDPs and comet Wild 2, but they are not proof of an extraterrestrial origin of the particle.

Indicators of Atmospheric Entry Flash-Heating

The smooth structures along the particle perimeter could be a macroscopic ablation feature, or at least serve as an indicator of particle modification. The still angular feature of the hematite grain suggests that, if the particle was heated, the temperatures stayed well below its melting point at 1665 °C. It also means that crystalline albite-anorthite would not have melted. Thus, we accept that the amorphous Na,Ca-aluminosilicate was always amorphous. Its original compositions are unknown, but the proposed Na-loss would be consistent with heating of particle W7190-D12. However, it does not offer tight thermal constraints, but Na-loss from these framework silicates is fast and does not require high temperatures. The Raman data show that hematite in the peripheral areas of the particle are more disordered than in the central part in the general area of analysis spot 1. We submit that thermal annealing might be the cause of this observed decrease in disorder. A thermal annealing study tracking the annealing of structural irregularities in nanometer scale hematite found that temperatures between 420 and 510 °C sufficed (Pérez-Maqueda et al. 2002). The presence of a 3.4 μm feature in the infrared spectrum of the particle indicates that the annealing temperature was less than 600 °C (Muñoz Caro et al. 2006). While these studies are not a direct analog to the conditions during atmospheric entry flash-heating of IDPs, it does suggest that the reduction in hematite disorder was probably a response to a heating event at low temperatures. In IDPs, most iron is present as ferrous iron, excluding GEMS that contain metallic iron (Bradley 1994). Flash-heating modifications during atmospheric entry in an oxidizing atmosphere causes oxidation and formation of ferric iron-bearing minerals, viz. magnetite and laihunite (Rietmeijer 1998). As the thermal spike during atmospheric entry is short-lived (5–15 s), diffusion distances need to be short, which is amply demonstrated by the nanometer-scale sizes of magnetite and maghémite grains in the rims on heated chondritic and nonchondritic IDPs alike. Size-wise, the <150 nm in diameter FeO-spherules in this particle could have formed during atmospheric entry. However, there are also indicators that flash-heating of C-bearing IDPs occurred in reducing environments (Rietmeijer 1998).

The particle W7190-D12 is exceptional among the collected natural stratospheric dust particles because it contains hematite with a surface layer of organic carbons. Surface heating could initiate a simple reaction of hematite reduction, viz. Fe2O+ C = 2FeO + CO. In this reaction, the reagent was most likely organic carbon. The CO product of this reaction would most likely be desorbed from the particle and be lost as a gas phase molecule to the atmosphere. Although FeO spherules are found across the entire particle, there is a distinct concentration located at one side of the hematite grain (Fig. 6). This side shows the “micaceous” layering that would be favorable sites for spherule formation. We suggest that at such sites local structural disorder in hematite became the loci of Fe3+ reduction and the formation of FeO-spherules.

Particle W7190-D12 probably experienced mild thermal heating. When it is not a volcanic dust particle, the alternative is an extraterrestrial origin and heating occurred during deceleration in the Earth’s atmosphere. The FeO-spherules formed during atmospheric entry. They can be considered the equivalent of a partial iron-oxide rim on IDPs that decelerated in under-oxidizing conditions. Ablation of its organic materials created the reducing conditions to effectuate limited reduction of hematite in particle W7190-D12. Fe-Ni spherules on the surface of silicate spheres were observed in a few rare IDPs and appear to have been produced by reduction processes related to strong heating of organic materials (Brownlee et al. 2001). We have not proven the extraterrestrial origin for particle W7190-D12. No search for isotopic anomalies or solar flare tracks could be conducted because the particle was first crushed onto the infrared transparent window. Clearly, more work is needed on these particles to prove their currently putative extraterrestrial origin such as determinations of the noble gas contents or searching for solar flare tracks in other Ca-rich particles on the same collector. Also, experimental work is needed on the formation of FeO spherules in chondritic and nonchondritic IDPs with organic materials in the presence of ferric-iron minerals during simulated conditions of flash-heating.

The petrological properties of nonchondritic particle W7190-D12 are consistent with platy hematite, which is a common mineral in terrestrial iron ore occurrences. It was recently also identified as present at the Martian surface (Lane et al. 2002), which provides a potential extraterrestrial source for such particles assuming that there is an ejection mechanism to place them in Earth-crossing orbits. It is noteworthy that it has been suggested that nonchondritic IDPs, including hydrated low-Ni nonchondritic IDPs, could be a new type of interplanetary dust particle that could have a lunar, Martian, or differentiated-asteroid origin (Flynn and Sutton 1990, 1991; Rietmeijer 1992).

Assuming that the particle is a nonchondritic IPD, we might then speculate on a possible cometary origin. The mass of particle W7190-D12 is contained in hematite (ρ = 5.3−3) and amorphous Na,Ca-aluminosilicate (ρ ≈ 2.7 g cm−3 for the mineral form, but less when amorphous estimated at ρ ≈ 2 g cm−3) in roughly equal proportions. The estimated density of this particle that is approximately 9 μm in diameter is then ρ ≈ 3.65 g cm−3. Love and Brownlee (1994) developed a matrix to estimate atmospheric entry peak heating temperatures for an “average” IDP (ρ = 2.0 g cm−3) striking the atmosphere at a 45° angle. An important invariable in this matrix is that the product of particle density and diameter is constant. Exploiting this constancy and using the diameter and estimated density for particle W7190-D12 its equivalent diameter is about 16 μm. Its corresponding peak heating temperature would be approximately 600 °C at an entry velocity of 10 km s−1 and approximately 1200 °C at an entry velocity of 20 km s−1 (i.e., a low-cometary velocity). There is no evidence that this particle was heated to approximately 1200 °C, but there is the possibility that this particle was heated to approximately 500 °C, which suggests that particle W7190-D12 is an IDP that entered the atmosphere at approximately 10 km s−1. The average atmospheric entry velocity of dust from the Zodiacal Cloud is 14.5 km s−1, but a sizeable fraction enters at velocities of approximately 12 km s−1. The estimated low entry velocity for this particle suggests that it was on an orbit of low eccentricity and low inclination to the ecliptic. Unfortunately, as its time of entry and direction (azimuth) of flight are unknown, there is no constraint on its source region. Interestingly, a recently developed dynamical model of the Zodiacal cloud showed that 85–95% of 100–200 μm particles in this cloud were ejected from Jupiter-family comets, <10% is contributed by particles from long-period comets, and that the contribution of asteroidal dust is no more than 10% (Nesvorný et al. 2010). Flynn (1996) had previously suggested that approximately 10 μm IDPs from the Kuiper Belt might be present among the Zodiacal dust cloud. We submit that particle W7190-D12 and other C-type IDPs on the same collector might be Jupiter-family comet debris. This particle W7190-D12 is the first detection of FeO spherules at the surface of a mildly heated nonchondritic IDP where it could form only in the presence of carbonaceous material. The present data cannot exclude the possibility that some fraction of the spherules could be metallic iron. The particle of this study may represent a potentially new type of nonchondritic IDP that could either be from the asteroid belt or was possibly associated with Jupiter-family comets.

Prior to the successful Stardust mission, this proposition would be without context. The results of this mission have shown that Kuiper Belt objects could harbor the same highly evolved minerals that were thought to be restricted to objects in the asteroid belt, such as those found in the unequilibrated ordinary chondrites (among others, Brownlee et al. 2006; Joswiak et al. 2009; Keller et al. 2006; Zolensky et al. 2006, 2008). This is an opportune time to search the NASA JSC Cosmic Dust Collection for nonchondritic IDPs that could be associated with Jupiter-family comets.

Acknowledgments–– We are grateful to M. E. Zolensky, G. J. Flynn, and D. E. Brownlee for very constructive reviews. We thank the curator and colleagues at the NASA JSC Curatorial Facility for providing particle W7190-D12. We acknowledge P. Dumas for assistance with the use of the micro-FTIR spectrometer, and G. Montagnac for the use of the micro-Raman spectrometer. We are grateful to E. Dartois for the Raman measurements. We thank G. Matrajt for requesting this particle. G. M. M. C. was supported by a Marie Curie Individual Fellowship from the European Union, a Ramón y Cajal research contract, and project AYA2008-06374 from the MCYT. F. J. M. R. was supported by NASA grant NNX07AM65G.

Editorial Handling–– Dr. Donald Brownlee


  1. Top of page
  2. Abstract
  3. Introduction
  4. Experimental
  5. Experimental Results
  6. Discussion and Conclusions
  7. References
  • Anders E. and Grevesse N. 1989. Abundances of the elements: Meteoritic and solar. Geochimica et Cosmochimica Acta 53:197214.
  • Andersen J. C. Ø., Rollinson G. K., Snook B., Herrington R., and Fairhurst R. J. 2009. Use of QEMSCAN® for the characterization of Ni-rich and Ni-poor goethite in laterite ores. Mineral Engineering 22:11191129.
  • Bradley J. P. 1994. Chemically anomalous, preaccretionally irradiated grains in interplanetary dust from comets. Science 265:925929.
  • Bradley J. P. and Brownlee D. E. 1991. An interplanetary dust particle linked directly to type CM meteorites and an asteroidal origin. Science 251:549552.
  • Bradley J. P., Humecki H. J., and Germani M. S. 1992. Combined infrared and analytical electron microscope studies of interplanetary dust particles. The Astrophysical Journal 394:643651.
  • Bridges J. C., Burchell M. J., Changela H. C., Foster N. J., Creighton J. A., Carpenter J. D., Gurman S. J., Franchi I. A., and Busemann H. 2010. Iron oxides in comet 81P/Wild 2. Meteoritics & Planetary Science 45:5572.
  • Brownlee D. E. 1978. In Cosmic dust, edited by McDonnell J. A. M. New York: Wiley Interscience Publication. pp. 295336.
  • Brownlee D. E. 1979. Interplanetary dust. Reviews of Geophysics and Space Physics 17:17351743.
  • Brownlee D. E. 1985. Cosmic dust: Collection and research. Annual Review of Earth and Planetary Sciences 13:147173.
  • Brownlee D. E., Joswiak D. J., Bradley J., Kress M., Pepin R., Schlutter D., and Palma R. 2001. Carbonaceous meteor ash––A significant carrier of carbon, organic material and noble gas to the surfaces of terrestrial planets? (abstract #2170). 32nd Lunar and Planetary Science Conference. CD-ROM.
  • Brownlee D., Tsou P., Aléon J., Alexander C. M. O’D., Araki T., Bajt S., Baratta G. A., Bastien R., Bland Ph., Bleuet P., Borg J., Bradley J. P., Brearley A., Brenker F., Brennan S., Bridges J. C., Browning N. D., Brucato J. R., Bullock E., Burchell M. J., Busemann H., Butterworth A., Chaussidon M., Cheuvront A., Chi M., Cintala M. J., Clark B. C., Clemett S. J., Cody G., Colangeli L., Cooper G., Cordier P., Daghlian C., Dai Z., D’Hendecourt L., Djouadi Z., Dominguez G., Duxbury T., Dworkin J. P., Ebel D. S., Economou T. E., Fakra S., Fairey S. A. J., Fallon S., Ferrini G., Ferroir T., Fleckenstein H., Floss C., Flynn G., Franchi I. A., Fries M., Gainsforth Z., Gallien J.-P., Genge M., Gilles M. K., Gillet P., Gilmour J., Glavin D. P., Gounelle M., Grady M. M., Graham G. A., Grant P. G., Green S. F., Grossemy F., Grossman L., Grossman J. N., Guan Y., Hagiya H., Harvey R., Heck P., Herzog G. F., Hoppe P., Hörz F., Huth J., Hutcheon I. D., Ignatyev K., Ishii H., Ito M., Jacob D., Jacobsen C., Jacobsen S., Jones S., Joswiak D., Jurewicz A., Kearsley A. T., Keller L. P., Khodja H., Kilcoyne A. L. D., Kissel J., Krot A., Langenhorst F., Lanzirotti A., Le L., Leshin L. A., Leitner J., Lemelle L., Leroux H., Liu M-C., Luening K., Lyon I., MacPherson G., Marcus M. A., Marhas K., Marty B., Matrajt G., McKeegan K., Meibom A., Mennella V., Messenger K., Messenger S., Mikouchi T., Mostefaoui S., Nakamura T., Nakano T., Newville M., Nittler L. R., Ohnishi I., Ohsumi K., Okudaira K., Papanastassiou D. A., Palma R., Palumbo M. E., Pepin R. O, Perkins D., Perronnet M., Pianetta P., Rao W., Rietmeijer F. J. M., Robert F., Rost D., Rotundi A., Ryan R., Sandford S. A., Schwandt G. S., See T. H., Schlutter D., Sheffield-Parker J., Simionovici A., Simon S., Sitnitsky I., Snead C. J., Spencer M. K., Stadermann F., Steele A., Stephan T., Stroud R., Susini J., Sutton S. R., Suzuki Y., Taheri M., Taylor S., Teslich N., Tomeoka K., Tomioka N., Toppani A., Trigo-Rodríguez J. M., Troadec D., Tsuchiyama A., Tuzzolino A. J., Tyliszczak T., Uesugi K., Velbel M., Vellenga J., Vicenzi E., Vincze L., Warren J., Weber I., Weisberg M., Westphal A. J., Wirick S., Wooden D., Wopenka B., Wozniakiewicz P., Wright I., Yabuta H., Yano H., Young E. D., Zare R. N., Zega T., Ziegler K., Zimmerman L., Zinner E., and Zolensky M. 2006. Comet 81P/Wild 2 under a microscope. Science 314:17111716.
  • Christoffersen R. and Buseck P. R. 1986. Mineralogy of interplanetary dust particles from the “olivine” infrared class. Earth and Planetary Science Letters 78:5366.
  • Dartois E., Marco O., Muñoz Caro G. M., Brooks K., Deboffle D., and d’Hendecourt L. 2004. Organic matter in Seyfert 2 nuclei: Comparison with our Galactic center lines of sight. Astronomy & Astrophysics 423:549558.
  • de Faria D. L. A., Venâncio Silva S., and de Oliveira M. T. 1997. Raman microspectroscopy of some iron oxides and oxyhydroxides. Journal of Raman Spectroscopy 28:873878.
  • Ferini G., Baratta G. A., and Palumbo M. E. 2004. A Raman study of ion irradiated icy mixtures. Astronomy & Astrophysics 414:757766.
  • Ferrari A. C. 2002. Determination of bonding in diamond-like carbon by Raman spectroscopy. Diamond and Related Materials 11:10531061.
  • Flynn G. J. 1994. Interplanetary dust particles collected from the stratosphere: Physical, chemical, and mineralogical properties and implications for their sources. Planetary and Space Science 42:11511161.
  • Flynn G. J. 1996. Sources of 10 micron interplanetary dust: The contribution from the Kuiper belt. In Physics, chemistry and dynamics of interplanetary dust, edited by Gustafson B. Å. S. and Hanner M. S. Astronomical Society of the Pacific Conference Series 104:171175.
  • Flynn G. J. and Sutton S. R. 1990. Synchrotron X-ray fluorescence analyses of stratospheric cosmic dust: New results for chondritic and low-nickel particles. Proceedings, 20th Lunar and Planetary Science Conference. pp. 335342.
  • Flynn G. J. and Sutton S. R. 1991. Chemical characterization of seven Large Area Collector particles by SXRF. Proceedings, 21st Lunar and Planetary Science Conference. pp. 549556.
  • Flynn G. J., Keller L. P., Feser M., Wirick S., and Jacobsen C. 2003. The origin of organic matter in the solar system: Evidence from the interplanetary dust particles. Geochimica et Cosmochimica Acta 67:47914806.
  • Fraundorf P. 1981. Interplanetary dust in the transmission electron microscope: Diverse materials from the early solar system. Geochimica et Cosmochimica Acta 45:915943.
  • Fraundorf P., Patel R. I., and Freeman J. J. 1981. Infrared spectroscopy of interplanetary dust in the laboratory. Icarus 47:368380.
  • Fraundorf P., Patel R. I., Walker R. M., Freeman J. J., and Adar F. 1982. Raman spectroscopy of graphite and other phases in meteorites and interplanetary dust (abstract). 13th Lunar and Planetary Science Conference. pp. 231232.
  • Joswiak D. H., Brownlee D. E., Matrajt G., Westphal A. J., and Snead C. J. 2009. Kosmochloric Ca-rich pyroxenes and FeO-rich olivines (Kool grains) and associated phases in Stardust tracks and chondritic porous interplanetary dust particles: Possible precursors to FeO-rich type II chondrules in ordinary chondrites. Meteoritics & Planetary Science 43:15611588.
  • Joy D. C. 1991. An introduction to Monte Carlo simulations. Scanning Microscopy 5:329337.
  • Keller L. P., Bajt S., Baratta G. A., Borg J., Bradley J. P., Brownlee D. E., Busemann H., Brucato J. R., Burchell M., Colangeli L., d’Hendecourt L., Djouadi Z., Ferrini G., Flynn G., Franchi I. A., Fries M., Grady M. M., Graham G. A., Grossemy F., Kearsley A., Matrajt G., Nakamura-Messenger K., Mennella V., Nittler L., Palumbo M. E., Stadermann F. J., Tsou P., Rotundi A., Sandford S. A, Snead C., Steele A., Wooden D., and Zolensky M. 2006. Infrared spectroscopy of comet 81P/Wild 2 samples returned by Stardust. Science 314:17281731.
  • Klekociuk A. R., Brown P. G., Pack D. W., Revelle D. O., Edwards W. N., Spalding R. E., Tagliaferri E., Yoo B. B., and Zagari J. 2005. Meteoritic dust from the atmospheric disintegration of a large meteoroid. Nature 436:11321135.
  • Lane M. D., Morris R. V., Mertzman A., and Christensen P. R. 2002. Evidence for platy hematite grains in Sinus Meridiani, Mars. Journal of Geophysical Research 107:5126.
  • Leroux H., Rietmeijer F. J M., Velbel M. A., Brearley A. J., Jacob D., Langenhorst F., Bridges J. C., Zega T. J., Stroud R. M., Cordier P., Harvey R. P., Lee M., Gounelle M., and Zolensky M. E. 2008. A TEM study of thermally modified comet 81P/Wild 2 dust particles by interactions with the aerogel matrix during the Stardust capture process. Meteoritics & Planetary Science 43:97120.
  • Love S. G. and Brownlee D. E. 1994. Peak atmospheric entry temperature of micrometeorites. Meteoritics 29:6970.
  • Mackinnon I. D. R., McKay D. S., Nace G., and Isaacs A. M. 1982. Classification of the Johnson Space Center Stratospheric Dust Collection. Proceedings, 13th Lunar and Planetary Science Conference. Journal of Geophysical Research 87, Suppl.: A413A421.
  • Matrajt G., Muñoz Caro G. M., Dartois E., D’Hendecourt L., Deboffle D., and Borg J. 2005. FTIR analysis of the organics in IDPs: Comparison with the IR spectra of the diffuse interstellar medium. Astronomy & Astrophysics 433:979995.
  • McKeegan K. D. 1987. Oxygen isotopes in refractory stratospheric dust particles: Proof of extraterrestrial origin. Science 237:14681471.
  • Muñoz Caro G. M., Matrajt G., Dartois E., Nuevo M., D’Hendecourt L., Deboffle D., Montagnac G., Chauvin N., Boukari C., and Le Du D. 2006. Nature and evolution of the dominant carbonaceous matter in interplanetary dust particles: Effects of irradiation and identification with a type of carbon. Astronomy & Astrophysics 459:147159.
  • Nesvorný D., Jenniskens P., Levison H. F., Bottke W. F., Vokrouhlický D., and Gounelle M. 2010. Cometary origin of the Zodiacal cloud and carbonaceous micrometeorites. Implications for hot debris disks. The Astrophysical Journal 713:816836.
  • Pérez-Maqueda L. A., Criado J. M., Real C., Balek V., and Šubrt J. 2002. Study of thermal evolution of porous hematite by emanation thermal analysis. Journal of the European Ceramic Society 22:22772281.
  • Pouchert C. 1997. Aldrich library of FT-IR spectra, 2nd ed. Milwaukee, WI: Aldrich Chemical Company. 5100 pp.
  • Rietmeijer F. J. M. 1991. Aqueous alteration in five chondritic porous interplanetary dust particles. Earth and Planetary Science Letters 102:148157.
  • Rietmeijer F. J. M. 1992. A detailed petrological analysis of hydrated, low-nickel, nonchondritic stratospheric dust particles. Proceedings, 22nd Lunar and Planetary Science Conference. pp. 195201.
  • Rietmeijer F. J. M. 1996. CM-like interplanetary dust particles in the lower stratosphere during 1989 October and 1991 June/July. Meteoritics & Planetary Science 31:278288.
  • Rietmeijer F. J. M. 1998. Interplanetary dust particles. In Planetary materials, edited by Papike J. J. Reviews in Mineralogy, vol. 36. Chantilly, VA: Mineralogical Society of America. pp. 2-12-95.
  • Rietmeijer F. J. M. 2002. The earliest chemical dust evolution in the solar nebula. Chemie der Erde 62:145.
  • Rietmeijer F. J. M. 2004. Dynamic pyrometamorphism during atmospheric entry of large (∼10 micron) pyrrhotite fragments from cluster IDPs. Meteoritics & Planetary Science 39:18691887.
  • Rietmeijer F. J. M. 2009a. Stardust glass: Indigenous and modified comet Wild 2 particles. Meteoritics & Planetary Science 44:17071715.
  • Rietmeijer F. J. M. 2009b. A cometary aggregate interplanetary dust particle as an analog for comet Wild 2 grain chemistry preserved in silica-rich Stardust glass. Meteoritics & Planetary Science 44:15891608.
  • Rietmeijer F. J. M. and Warren J. L. 1994. Windows of opportunity in the NASA Johnson Space Center Cosmic Dust Collection. In Analysis of interplanetary dust, edited by Zolensky M. E., Wilson T. L., Rietmeijer F. J. M., and Flynn G. J. American Institute Physics Conference Proceedings 310. Woodbury, NY: American Institute Physics Press. pp. 255275.
  • Rietmeijer F. J. M., Nuth J. A. III., and Karner J. M. 1999. Metastable eutectic condensation in a Mg-Fe-SiO-H2-O2 Vapor: Analogs to circumstellar dust. The Astrophysical Journal 527:395404.
  • Rietmeijer F. J. M., Pun A., Kimura Y., and Nuth J. A. III 2008. A refractory Ca-SiO-H2-O2 vapor condensation experiment with implications for calciosilica dust transforming to silicate and carbonate minerals. Icarus 195:493503.
  • Rotundi A., Ferrini G., Baratta G. A., Palumbo M. E., Palomba E., and Colangeli L. 2007. Combined micro-infrared (IR) and micro-Raman measurements on stratospheric interplanetary dust particles. Proceedings, Dust in Planetary Systems, Kauai, Hawaii, USA (ESA SP-643, January 2007), p. 149153.
  • Sandford S. A. and Walker R. M. 1985. Laboratory infrared transmission spectra of individual interplanetary dust particles from 2.5 to 25 microns. The Astrophysical Journal 291:838851.
  • Schramm L. S., Brownlee D. E., and Wheelock M. M. 1989. Major element composition of stratospheric micrometeorites. Meteoritics 24:99112.
  • Steele I. M. 1990. Minor elements in forsterites of Orgueil (Cl), Alais (Cl) and two interplanetary dust particles compared to C2-C3-UOC forsterites. Meteoritics 25:301307.
  • Thomas K. L., Blanford G. E., Clemett S. J., Flynn G. J., Keller L. P., Klöck W., Maechling C. R., McKay D. S., Messenger S., Nier A. O., Schlutter D. J., Sutton S. R., Warren J. L., and Zare R. N. 1995. An asteroidal breccia: The anatomy of a cluster IDP. Geochimica et Cosmochimica Acta 59:27972815.
  • Tomeoka K., Tomioka N., and Ohnishi I. 2008. Silicate minerals and Si-O glass in comet Wild 2 samples: Transmission electron microscopy. Meteoritics & Planetary Science 43:273284.
  • Wopenka B. 1988. Raman observations on individual interplanetary dust particles. Earth and Planetary Science Letters 88:221231.
  • Zolensky M. E. 1987. Refractory interplanetary dust particles. Science 237:14661468.
  • Zolensky M. E. and Barrett R. 1994. Compositional variations of olivines and pyroxenes in chondritic interplanetary dust particles. Meteoritics 29:616620.
  • Zolensky M. E., McKay D. S., and Kaczor L. A. 1989. A tenfold increase in the abundance of large solid particles in the stratosphere, as measured over the period 1976–1984. Journal of Geophysical Research 94:10471056.
  • Zolensky M. E., Wilson T. L., Rietmeijer F. J. M., and Flynn G. J., eds. 1994. Analysis of interplanetary dust. American Institute Physics Conference Proceedings vol. 310. New York: American Institute Physics Press. 357 p.
  • Zolensky M. E., Zega T. J., Yano H., Wirick S., Westphal A. J., Weisberg M. K., Weber I., Warren J. L., Velbel M. A., Tsuchiyama A., Tsou P., Toppani A., Tomioka N., Tomeoka K., Teslich N., Taheri M., Susini J., Stroud R., Stephan T., Stadermann F. J., Snead C. J., Simon S. B., Simionovici A., See T. J., Robert F., Rietmeijer F. J. M., Rao W., Perronnet M. C., Papanastassiou D. A., Okudaira K., Ohsumi K., Ohnishi I., Nakamura-Messenger K., Nakamura T., Mostefaoui S., Takashi Mikouchi T., Meibom A., Matrajt G., Marcus M. A., Leroux H., Lemelle L., Le L., Lanzirotti A., Langenhorst F., Krot A. N., Keller L. P., Kearsley A. T., Joswiak D., Jacob D., Ishii H., Harvey R., Hagiya K., Grossman L., Grossman J. N., Graham G. A., Gounelle M., Gillet P., Genge M. J., Flynn G., Ferroir T., Fallon S., Ebel D. S., Dai Z., Cordier P., Clark B., Chi M., Butterworth A. L., Brownlee D. E., Bridges J. C., Brennan S., Brearley A., Bradley J. P., Bleuet P., Bland P. A., and Bastien R. 2006. Mineralogy and petrology of comet Wild 2 nucleus samples. Science 314:17351739.
  • Zolensky M., Nakamura-Messenger K., Sverdrup J., Rietmeijer F., Leroux H., Mikouchi T., Ohsumi K., Simon S., Grossman L., Stephan T., Weisberg M., Velbel M., Zega T., Stroud R., Tomeoka K., Ohnishi I., Tomioka N., Nakamura T., Matrajt G., Joswiak J., Brownlee D., Langenhorst F., Krot A., Kearsley A., Ishii H., Graham G., Dai Z. R., Chi M., Bradley J., Hagiya K., Gounelle M., and Bridges J. 2008. Comparing Wild 2 particles to chondrites and IDPs. Meteoritics & Planetary Science 43:261272.
  • Zoppi A., Lofrumento C., Castellucci E. M., and Migliorini M. G. 2005. The Raman Spectrum of hematite: Possible indicator for a compositional or firing distinction among Terra Sigillata Wares. Annali di Chimica 95:239246.