Meteoritics & Planetary Science 38, Nr 8, 1217 1241 (2003) Abstract available online at http://meteoritics.org Minor element zoning and trace element geochemistry of pallasites Weibiao HSU1, 2* 1 Purple Mountain Observatory, Chinese Academy of Sciences, 2 West Beijing Road, Nanjing 210008, China 2 Division of Geological & Planetary Sciences, Mail Code 170 25, California Institute of Technology, Pasadena, California 91125, USA * Corresponding author. E-mail: wbxu@pmo.ac.cn (Received 1 April 2003; revision accepted 9 July 2003) Abstract I report here on an ion probe study of minor element spatial distributions and trace element concentrations in six pallasites. Pallasite olivines exhibit ubiquitous minor element zoning that is independent of grain size, morphology, and adjacent phases. Ca, Cr, Ti, V, and Ni concentrations decrease from center to rim by factors of up to 10, while Mn is generally unzoned or increases slightly at the very edge of some olivine grains. The maximum concentrations of these elements at the center of olivine vary from grain to grain within the same meteorite and among the pallasites studied. These zoning profiles are consistent with thermal diffusion during rapid cooling. The inferred cooling rates at high temperature regimes are orders of magnitude faster than the low-temperature metallographic cooling rates (~0.5 to 2 C/Ma). This suggests that pallasites, like mesosiderites, have experienced rather complicated thermal histories, i.e., cooling rapidly at high temperatures and slowly at low temperatures. Pallasite olivines are essentially free of REEs. However, the phosphates display a wide range of REE abundances (0.001 to 100 CI) with distinct patterns. REEs are generally homogeneous within a given grain but vary significantly from grain to grain by a factor of up to 100. Albin and Imilac whitlockite are highly enriched in HREEs (~50 CI) but are relatively depleted in LREEs (~0.1 to 1 CI). Eagle Station whitlockite has a very unusual REE pattern: flat LREEs at a 0.1 CI level, a large positive Eu anomaly, and a sharp increase from Gd (0.1 CI) to Lu (70 CI). Eagle Station stanfieldite has a similar REE pattern to that of whitlockite but with much lower REEs by a factor of 10 to 100. Springwater farringtonite has relatively low REE concentrations (0.001 to 1 CI) with a highly fractionated HREE-enriched pattern (CI-normalized Lu/La ~100). Postulating any igneous processes that could have fractionated REEs in these phosphates is difficult. Possibly, phosphates were incorporated into pallasites during mixing of olivine and IIIAB-like molten Fe. These phosphates preserve characteristics of a previous history. Pallasites have not necessarily formed at the mantle-core boundary of their parent bodies. The pallasite thermal histories suggest that pallasites may have formed at a shallow depth and were subsequently buried deep under a regolith blanket. INTRODUCTION Pallasites are highly differentiated meteorites with a very simple mineralogy. They consist of roughly equal amounts of olivine and Fe-Ni metal with minor phosphates, pyroxenes, chromite, schreibersite, and troilite (Buseck 1977). Despite their simple mineralogy, various models have been proposed regarding the origin of pallasites and their physical position in the parent bodies. A popular idea is that pallasites are samples of the mantle-core boundary of the differentiated parent bodies (Mason 1962, 1963; Anders 1964; Buseck and Goldstein 1969). This is mainly because olivine and Fe-Ni metal are physically and chemically dissimilar and immiscible due to their large contrast in density if they both crystallized from a melt. Thus, one can legitimately expect that olivine at the bottom of the mantle mixes with Fe-Ni metal from the core to form a layer of pallasite in a differentiated parent body. Urey (1956), however, posed a different view of the origin for pallasites. He noted that pallasites were too abundant to represent a single boundary layer and suggested that they came from discrete metal pools throughout the parent bodies, the so-called raisin-bread model. A third theory is that pallasites originated near the surface of their parent bodies (Mittlefehldt 1980). 1217 Meteoritical Society, 2003. Printed in USA.
1218 W. Hsu Mittlefehldt (1980) found that mesosiderite and pallasite olivines have similar major element compositions and suggested that pallasite olivine, like mesosiderite olivine, was a product of external heating of a chondritic parent body. The olivine layer produced by extensive partial melting and fractional crystallization near the surface was invaded by molten metal from the upper portion of the parent body. Thermal histories of pallasites reveal information essential to a deeper understanding of their physical position in the parent bodies as well as their origin. Cooling rates estimated from kamacite-taenite diffusion profiles in pallasite Fe-Ni metal are in the range of 0.5 2 C/Ma from 700 to 300 C (Buseck and Goldstein 1969), which indicates a burial depth of ~200 km (Fricker et al. 1970). Such slow cooling rates are generally compatible with pallasite formation at the mantle-core boundary in their parent bodies. However, studies of major and minor element concentrations in pallasite olivines suggested much high cooling rates at high temperatures (Scott 1977a; Reed et al. 1979; Miyamoto 1997). Pallasite olivines are well-known to exhibit chemical zoning of minor elements (e.g., Ca, Cr, Ti, and Ni), which were thought to represent diffusion during rapid cooling processes under subsolidus temperatures (Leitch et al. 1979; Reed et al. 1979; Zhou and Steele 1993; Steele 1994). Diffusion induced chemical zoning in silicates potentially provides constraints on their thermal histories. Most analyses of minor element zoning in pallasite olivines were obtained with an electron microprobe (Zhou and Steele 1993; Steele 1994; Miyamoto 1997). As these minor element concentrations are usually very low and close to the detection limit of an electron microprobe, high sensitivity ion microprobe analyses are desired. The ion microprobe data collected to date include analyses of 7 minor element concentrations in a Springwater olivine (Leitch et al. 1979), the Ni concentration in 7 pallasite olivines (Reed et al. 1979), and 11 minor element concentrations in olivines of Springwater and Mount Vernon (Floss 2002). In this study, I extended the ion microprobe analyses to 6 minor element spatial distributions in olivine crystals from 6 pallasites. I will compare this data with previously reported results and discuss their implications to pallasite thermal histories. Trace elements, particularly the rare earth elements (REEs), provide much information about the crystallization, fractionation, and differentiation processes that led to the formation of igneous rocks. Pallasites are highly differentiated igneous meteorites. The REE distributions in pallasites provide additional constraints on their origin. The major REE sources in these meteorites are olivine (dominant phase) and phosphates (REE carrier). Early studies showed that pallasite olivine has a V-shaped CI-normalized REE pattern with abundances at 10 2 to 1 CI (Schmitt et al. 1964; Masuda 1968). Recent analyses by Minowa and Ebihara (2002) indicated that much of the light REE (LREE) in pallasite olivines are due to terrestrial contamination and that this mineral has extremely low REE abundances (10 5 to 10 2 CI) with a highly fractionated HREE-enriched pattern. While the REEs are virtually excluded from olivine, they can be highly enriched in phosphates. Ion microprobe analyses revealed that pallasite phosphates exhibit diverse REE patterns with highly variable abundances (Davis and Olsen 1991, 1996). Most phosphates have relatively low REE abundances (0.01 to 10 CI) with an HREE-enriched pattern (CI-normalized Lu/La ~100). A few of the phosphate grains in Springwater and Santa Rosalia are highly enriched in REEs (up to 300 CI) with an LREE-enriched pattern (CI-normalized La/Lu 3 to 10). A merrillite grain from Giroux has an unusual REE pattern similar to the volatilitycontrolled group II REE pattern commonly observed in CAIs of CV chondrites (Davis and Olsen 1996). Interpretations of this diversity include phosphate formation by sub-solidus redox reaction between olivine and metal, crystallization from a trapped liquid, and the late-stage crystallization on pallasite parent bodies (Davis and Olsen 1991). In this study, I carried out an investigation of REE concentrations in phosphates and olivines from 6 pallasites. With the extensive data set on minor element zoning of olivine and REE abundances of phosphates, I will evaluate previous models and seek additional constraints on the origin of pallasites. ANALYTICAL METHODS Six pallasites, Albin, Brenham, Eagle Station, Imilac, Glorieta Mountain, and Springwater, were studied in this work. Samples used are 1 inch round polished thick sections. All samples were studied with a JEOL JSM 35-CF scanning electron microscope (SEM) equipped with a Tractor Northern energy dispersive (EDS) X-ray analysis system. The spatial distributions of minor elements in pallasite olivines were obtained with a Cameca 3f ion microprobe at the California Institute of Technology. A mass resolving power was set at ~3000. A liquid nitrogen trap was used near the secondary ion source to further depress the hydride signal. No energy filtering was used in the measurements. A primary beam current of 1 na was used, which resulted in a beam spot of ~5 µm. Standardization was done on San Carlos olivine, NBS-610 glass, and a synthetic Ti-pyroxene. The singly charged positive secondary ions were collected at masses 30 Si, 40 Ca, 47 Ti, 48 Ti, 51 V, 52 Cr, 55 Mn, and 60 Ni. The minor element concentrations were obtained by normalizing ion signal intensities to the silicon content of olivine. The analysis precisions (1σ) for Ca, Ti, V, Cr, Mn, and Ni are 3%, 5%, 5%, 1%, 1%, and 20%, respectively. The precision error for Ni is relatively large. This is a combination of counting statistics and the matrix effect on the ionization efficiency of Ni. Reed et al. (1979) reported a strong and systematic dependence of the secondary ion yield for Ni on the fayalite (Fa) content of
Minor element zoning and trace element geochemistry of pallasites olivine. The Fa values of the olivines in this study range from Fa10 in San Carlos olivine to Fa20 in Eagle Station olivine. Minor element distribution profiles were obtained from linear traverses across olivine grains. REE measurements were also made with the Caltech ion microprobe using the techniques of Fahey et al. (1987). Primary beam currents of 10 na (~20 µm in beam size) for phosphates and 20 na (~40 µm in beam size) for olivine were used to sputter positive secondary ions from the sample surfaces. The secondary ions were collected at low mass resolution using 80 V of energy filtering, which efficiently suppresses complex molecular interferences (Zinner and Crozaz 1986). A deconvolution calculation was carried out to eliminate the simple oxide interferences in the mass range of REEs. Synthetic Ti-pyroxene glass and Durango apatite standards were measured periodically to account for any variation of ionization efficiencies caused by minor changes of operating conditions. Reference element concentrations of pallasite olivine (SiO2) and phosphates (CaO and MgO) were taken from the literature (Buseck and Goldstein 1969; Buseck and Holdsworth 1977). 1219 RESULTS Minor Element Spatial Distributions in Pallasite Olivines I studied minor element spatial distributions in olivine grains from 6 pallasites (Table 1 and Appendix). The size of pallasite olivine ranges from 250 to 4300 µm. Both round and angular olivine grains were included in the analyses. These olivine crystals are adjacent to Fe-Ni metal, troilite, chromite, or phosphate (Fig. 1). I found that minor element zoning is ubiquitous among pallasite olivines (Fig. 2). It is independent of olivine grain sizes, morphologies, and adjacent phases. However, each element behaves distinctly. Ca: Ca is zoned within 300 500 µm from the edge in all pallasite olivines. Its concentration decreases from the core to the rim by a factor of up to 3. Within a given meteorite, the maximum Ca concentration at the center of olivine varies by a factor of up to 3 (e.g., from 25 to 70 ppm in Albin) and the minimum concentration at the rim by a factor of up to 4 (e.g., from 12 to 47 ppm in Albin). Among the pallasites studied, the Fig. 1. Reflected light microscope photographs of ion probe profile measurements in Albin olivine (Ol). The visible dots on the olivine crystals are the ion probe pits: a) olivine adjacent to Fe metal (Fe). The field of view is 2.25 mm; b) olivine adjacent to chromite (Chrm) and Fe metal (Fe). The field of view is 2.25 mm; c) olivine adjacent to troilite (FeS). The field of view is 2.25 mm; d) olivine adjacent to troilite (FeS) and chromite (Chrm). The field of view is 1.13 mm.
1220 W. Hsu Table 1. Minor element spatial distributions (ppm) in pallasite olivines. Ca Ti V Cr Mn Ni Size (µm) Adjacent phases Core Rim Core Rim Core Rim Core Rim Core Rim Core Rim Albin 920 Fe, Ni Fe, Ni 32 26 4 2 5 4 74 33 2250 2310 16 15 1220 Fe, Ni Fe, Ni 36 26 5 2 6 5 100 40 2240 2365 15 10 625 FeS FeS 51 24 18 2 8 5 158 35 2300 2400 15 10 310 Chromite FeS 25 12 2 1 5 3 33 16 2345 1140 12 6 480 FeS FeS 70 47 14 2 5 3 124 36 2360 2170 17 15 500 Chromite Fe, Ni 60 45 2 3 4 2 45 30 2300 2465 20 13 Brenham 4300 Fe, Ni Fe, Ni 150 100 3 1 4 3 90 25 1535 1500 15 12 Eagle Station 950 Fe, Ni Fe, Ni 280 92 8 3 4 3 70 20 2360 2370 55 45 530 Fe, Ni Fe, Ni 270 96 12 4 5 3 94 27 2360 2290 55 36 1100 Fe, Ni Fe, Ni 380 120 15 5 6 3 150 38 2245 2270 74 33 330 Chromite Fe, Ni 186 87 3 1 3 1 27 13 2330 1880 65 40 Glorieta Mountain 970 Fe, Ni Olivine 71 48 15 5 8 6 118 54 2540 2800 28 25 250 Fe, Ni Fe, Ni 108 86 7 4 6 5 72 48 2290 2677 40 32 Imilac 275 Fe, Ni Fe, Ni 68 43 16 8 5 3 80 42 2290 1585 12 7 480 Fe, Ni Fe, Ni 90 70 25 7 8 5 173 62 2230 2445 15 12 300 Fe, Ni Fe, Ni 76 61 20 10 7 6 135 68 2400 2500 20 10 Springwater 1300 Fe, Ni Olivine 193 65 5 1 6 4 144 36 3400 3580 50 27 2250 FeS Farringtonite 115 67 6 1 6 4 100 30 3450 3750 46 30
Minor element zoning and trace element geochemistry of pallasites 1221 Fig. 2. Ca, Ti, V, Cr, Mn, and Ni zoning profiles in Albin (open circle) and Eagle Station (filled circle) olivines. Note that the maximum concentrations at the center of the grains are significantly different among pallasite olivines. maximum Ca concentration at the center of olivine also varies significantly, e.g., 25 ppm in Albin and 380 ppm in Eagle Station. Such large variations suggest that pallasite olivines were heterogeneous with respect to the initial Ca concentration. Cr: Cr zoning in olivine is very similar to that of Ca but with a more pronounced zoning profile. Its concentration typically decreases by a factor of up to 4 from core to rim. Eagle Station olivine exhibits the largest variations of Cr concentration. In this pallasite, the maximum Cr concentration at the center of olivine varies from 27 to 150 ppm, and the minimum Cr concentration at the rim of olivine ranges from 13 to 38 ppm from grain to grain. Imilac olivine (core) has Cr concentrations up to 170 ppm, and Eagle Station olivine (core) has Cr concentrations down to 27 ppm. Ti: Ti zoning in pallasite olivines can be very extensive. From the core to the rim, Ti content decreases by a factor of up to 9. In Albin, the maximum Ti concentration at the center of olivine ranges from 2 to 18 ppm and the minimum at the rim from 1 to 3 ppm. Imilac olivine (core) has much higher Ti concentration (16 to 25 ppm) than that of Breham (~3 ppm). V: Albin, Eagle Station, and Imilac olivines show clear V zoning, but olivines from Brenham, Glorieta Mountain, and Springwater do not. Some pallasite olivines have essentially homogeneous V within the grains. The variation of V concentration (core) is generally insignificant within a given meteorite and among the pallasites studied. The concentration (~5 ppm) is roughly the same among pallasite olivines (core).
1222 W. Hsu Mn: Mn behaves differently from other minor elements in pallasite olivines. Most olivines have a constant Mn concentration across the crystal. In some cases (e.g., Albin), it increases slightly at the very edge (within 50 µm) of the grain. Springwater olivine has a higher (~3500 ppm) and Brenham olivine has a lower Mn concentration (~1500 ppm) than the rest (~2300 ppm). Ni: Ni is clearly zoned in pallasite olivines with high Ni concentration (~50 ppm, Eagle Station and Springwater). The concentration decreases from the core to the rim by a factor of 2. However, Ni zoning is not pronounced in pallasite olivines with low Ni concentration (~20 ppm). Minor element zoning usually occurs within 300 500 µm from the edge of the grain. Large olivine grains (>2 mm) seem to have a flat top of minor element concentrations in the middle of the zoning profiles. Note that not all zoning profiles are as symmetric as in Fig. 2. Asymmetric zoning profiles are also commonly observed. A few grains exhibit anomalous profiles that do not show a simple increase or decrease in concentration relative to grain edges. Sometimes, the zoning profile encounters exceptionally high values (Fig. 2). This is due to the fact that pallasite olivines contain tubular inclusions (<1 µm) (Buseck 1977). However, it is not clear what phases these inclusions are. When I traced back to check the spots with a SEM, the inclusions were not detected. I presume that they were very small and were eroded by the ion beam. The observed minor element zoning profiles in pallasite olivines are consistent with previously reported data. Ca and Cr zoning profiles have been found in all pallasite olivines studied, such as Admire, Eagle Station, Esquel, Imilac, Pavlodar, and Springwater (Leitch et al. 1979; Zhou and Steele 1993; Steele 1994; Miyamoto 1997). Ion probe analyses also revealed clear Ni zoning in olivines of Eagle Station (Reed et al. 1979) and Springwater (Leitch et al. 1979). Reed et al. (1979) found that Ni in an Eagle Station olivine shows a decrease from 40 ppm at the core to 25 ppm at the rim. Mn has been found either unzoned in Springwater olivine (3175 ppm; Leitch et al. 1979) and Pavlodar olivine (2700 ppm; Steele 1994) or increasing at the edge (50 µm) of Springwater olivine (Zhou and Steele 1993). Ion probe analyses show that Ti (~6 ppm) is zoned but V (~10 ppm) is unzoned in Springwater olivine (Leitch et al. 1979). REE Concentrations in Pallasite Phosphates and Olivine Phosphates were found in 4 pallasites: Albin, Eagle Station, Imilac, and Springwater. They have highly variable chemical compositions, from Ca-rich whitlockite (Ca3[PO4]2) to Ca, Mg-bearing stanfieldite (Mg3Ca3[PO4]4) to Mg-rich farringtonite (Mg3[PO4]2). In Albin, I found 4 whitlockite grains ranging in size from 100 to 500 µm. They tend to occur interstitially between olivine and metal (Fig. 3a). In Eagle Station, 1 whitlockite grain and 8 stanfieldite grains were found; they are generally small (30 to 100 µm) and in irregular Fig. 3. Back scattered electron images of phosphate occurrences in pallasites: a) whitlockite (whit) occurs interstitially between olivine (ol) and Fe metal (Fe) in Albin; b) stanfieldite (stan) is adjacent to the Eagle Station olivine with co-existing low-ca pyroxene (pyx); c) whitlockite is highly fractured in Imilac; d) a round farringtonite (farr) grain occurs interstitially between olivine grains in Springwater.
Minor element zoning and trace element geochemistry of pallasites 1223 shapes. These grains are commonly plastered onto the surfaces of olivines, and sometimes, low-ca pyroxene is co-existing (Fig. 3b). Three whitlockite grains (200 to 500 µm) found in Imilac are extensively fractured with numerous iron oxide veins (Fig. 3c). They also occur interstitially between olivine and metal. No phosphate was found in Brenham and Glorieta Mountain. Although widespread, phosphates are usually present in minor amounts in pallasites. However, phosphates are relatively abundant in Springwater. Buseck (1977) reported 4 vol% phosphates in this meteorite. These phosphates are large in size (up to 1.5 mm) and occur interstitially between olivine grains (Fig. 3d). The round boundary of farringtonite grains indicates that this mineral was present initially as molten droplets. One farringtonite grain was found in a crack of an olivine grain. No whitlockite and Mn-rich silico-phosphates (Davis and Olsen 1991; Hutcheon and Olsen 1991) were found in the Springwater samples. Pallasite phosphates display distinct REE abundances and patterns (Table 2 and Fig. 4). In Albin, 8 REE measurements were made in 4 whitlockite grains. No significant variation of REEs was found within a given grain. Inter-grain variation is generally small, up to a factor of 5. Four whitlockite grains have essentially the same REE pattern that is highly enriched in Sm, Eu, and HREEs (~ 50 CI) but relatively depleted in La to Nd (~1 CI) (Fig. 4a). In general, CI-normalized REE abundances increase gradually from La to Nd and from Sm to Lu. A clear cut in abundances exists between Nd (~1 CI) and Sm (~30 CI). The Eu anomaly is negligible. In Eagle Station, the whitlockite grain has a very unusual REE pattern. It has essentially flat LREEs with depleted abundances (~0.1 CI), a large positive Eu anomaly (Eu/Eu* ~10, where Eu* is the interpolated value between CInormalized Sm and Gd abundances), and a sharp increase from Gd (0.1 CI) to Lu (70 CI) (Fig. 4b). The other 5 stanfieldite grains analyzed have a similar REE pattern to that of whitlockite but with much lower REE abundances, by a factor of 10 100 (Fig. 4b). Three whitlockite grains were found in Imilac. Four analyses were made in 1 grain and 1 Fig. 4. REE abundances and patterns of pallasite phosphates: a) Albin whitlockite; b) Eagle Station whitlockite and stanfieldite; c) Imilac whitlockite; d) Springwater farringtonite. Pallasite phosphates have highly variable REE abundances and patterns. In general, they all have an HREE-enriched pattern.
1224 W. Hsu Table 2. REE concentrations (ppm) in pallasite phosphates. (The errors quoted are 1σ standard deviation from counting statistics only.) Albin Whitlockite Whitlockite Whitlockite Whitlockite Whitlockite Whitlockite Whitlockite Whitlockite La 0.123 ± 0.010 0.11 ± 5 0.01 0.109 ± 0.012 0.102 ± 0.01 0.218 ± 0.014 0.183 ± 0.013 0.163 ± 0.014 0.174 ± 0.014 Ce 0.438 ± 0.023 0.462 ± 0.023 0.357 ± 0.025 0.234 ± 0.018 0.683 ± 0.028 0.687 ± 0.028 0.58 ± 0.028 0.607 ± 0.03 Pr 0.082 ± 0.009 0.078 ± 0.009 0.06 ± 0.008 0.05 ± 0.006 0.101 ± 0.01 0.102 ± 0.009 0.075 ± 0.008 0.101 ± 0.012 Nd 0.61 ± 0.033 0.623 ± 0.035 0.535 ± 0.045 0.391 ± 0.027 0.966 ± 0.045 0.953 ± 0.044 0.863 ± 0.046 0.83 ± 0.048 Sm 3.318 ± 0.108 3.436 ± 0.11 3.301 ± 0.122 2.728 ± 0.085 6.312 ± 0.162 6.415 ± 0.158 5.607 ± 0.164 5.62 ± 0.172 Eu 1.952 ± 0.075 1.739 ± 0.072 1.506 ± 0.045 2.27 ± 0.08 1.504 ± 0.039 1.631 ± 0.039 1.785 ± 0.043 1.856 ± 0.046 Gd 9.22 ± 0.262 9.201 ± 0.268 6.519 ± 0.196 5.534 ± 0.141 14.075 ± 0.369 14.767 ± 0.349 12.7 ± 0.37 12.73 ± 0.397 Tb 1.662 ± 0.074 1.571 ± 0.072 1.461 ± 0.073 0.659 ± 0.032 2.527 ± 0.091 2.659 ± 0.088 2.33 ± 0.094 2.238 ± 0.095 Dy 9.089 ± 0.294 9.157 ± 0.299 8.172 ± 0.246 5.909 ± 0.189 14.75 ± 0.399 14.964 ± 0.379 12.204 ± 0.395 12.964 ± 0.418 Ho 1.794 ± 0.083 2.121 ± 0.089 1.141 ± 0.066 1.089 ± 0.048 2.931 ± 0.123 3.223 ± 0.122 2.585 ± 0.127 2.505 ± 0.131 Er 6.094 ± 0.303 6.46 ± 0.302 4.751 ± 0.264 3.297 ± 0.18 9.689 ± 0.382 9.557 ± 0.37 8.713 ± 0.419 8.197 ± 0.427 Tm 1.059 ± 0.100 1.118 ± 0.094 0.943 ± 0.074 0.42 ± 0.069 1.85 ± 0.082 1.725 ± 0.077 1.58 ± 0.099 1.55 ± 0.103 Yb 10.969 ± 0.259 12.399 ± 0.259 9.611 ± 0.254 7.495 ± 0.195 20.39 ± 0.402 21.116 ± 0.393 16.996 ± 0.401 15.244 ± 0.409 Lu 1.865 ± 0.113 1.955 ± 0.11 0.83 ± 0.072 1.294 ± 0.086 3.438 ± 0.107 3.449 ± 0.102 3.028 ± 0.109 2.612 ± 0.109 Eagle Station Stanfieldite Stanfieldite Whitlockite Stanfieldite Stanfieldite Stanfieldite La 0.008 ± 0.004 0.026 ± 0.007 0.027 ± 0.008 0.007 ± 0.004 0.008 ± 0.004 0.012 ± 0.004 Ce 0.009 ± 0.004 0.042 ± 0.007 0.083 ± 0.012 0.016 ± 0.004 0.022 ± 0.004 0.013 ± 0.004 Pr <0.007 0.012 ± 0.004 0.004 ± 0.004 <0.004 <0.004 Nd 0.007 ± 0.004 0.04 ± 0.012 0.071 ± 0.014 0.01 ± 0.003 0.035 ± 0.007 0.01 ± 0.003 Sm 0.036 ± 0.033 0.005 ± 0.01 0.017 ± 0.008 0.011 ± 0.005 0.009 ± 0.006 0.011 ± 0.006 Eu 0.059 ± 0.007 0.002 ± 0.005 0.005 ± 0.003 0.015 ± 0.005 Gd 0.014 ± 0.007 0.022 ± 0.011 0.02 ± 0.013 0.032 ± 0.013 0.003 ± 0.003 Tb <0.007 0.008 ± 0.004 0.008 ± 0.004 0.009 ± 0.004 Dy 0.008 ± 0.004 0.019 ± 0.006 0.16 ± 0.018 0.018 ± 0.004 0.045 ± 0.008 0.007 ± 0.004 Ho <0.006 0.051 ± 0.007 0.007 ± 0.004 0.019 ± 0.004 0.004 ± 0.004 Er 0.009 ± 0.004 0.041 ± 0.007 0.435 ± 0.028 0.036 ± 0.008 0.076 ± 0.008 0.012 ± 0.004 Tm 0.026 ± 0.007 0.287 ± 0.019 0.022 ± 0.004 0.028 ± 0.004 0.012 ± 0.004 Yb 0.071 ± 0.01 0.366 ± 0.03 5.522 ± 0.148 0.472 ± 0.026 0.368 ± 0.023 0.201 ± 0.018 Lu 0.014 ± 0.007 0.144 ± 0.022 1.762 ± 0.065 0.14 ± 0.019 0.088 ± 0.014 0.066 ± 0.013 Imilac Whitlockite Whitlockite Whitlockite Whitlockite Whitlockite Whitlockite La 0.102 ± 0.012 0.099 ± 0.012 0.123 ± 0.015 0.094 ± 0.012 0.063 ± 0.014 0.099 ± 0.017 Ce 0.448 ± 0.028 0.325 ± 0.027 0.45 ± 0.029 0.322 ± 0.025 0.045 ± 0.015 0.054 ± 0.009 Pr 0.054 ± 0.009 0.068 ± 0.008 0.071 ± 0.01 0.061 ± 0.008 0.022 ± 0.007 0.009 ± 0.009 Nd 0.876 ± 0.052 0.895 ± 0.047 0.897 ± 0.052 0.83 ± 0.041 0.105 ± 0.025 0.036 ± 0.015 Sm 2.716 ± 0.128 3.149 ± 0.102 2.843 ± 0.138 2.543 ± 0.089 0.327 ± 0.051 0.097 ± 0.024 Eu 0.202 ± 0.021 0.245 ± 0.018 0.178 ± 0.019 0.178 ± 0.015 0.063 ± 0.018 0.037 ± 0.016 Gd 3.381 ± 0.159 3.938 ± 0.165 2.593 ± 0.18 3.311 ± 0.144 0.925 ± 0.107 0.176 ± 0.028 Tb 1.501 ± 0.087 1.559 ± 0.05 1.405 ± 0.087 1.408 ± 0.043 0.302 ± 0.031 0.091 ± 0.018 Dy 11.631 ± 0.453 12.77 ± 0.359 10.441 ± 0.459 11.78 ± 0.326 2.907 ± 0.113 0.735 ± 0.047 Ho 3.179 ± 0.203 3.267 ± 0.113 2.746 ± 0.207 2.917 ± 0.103 0.665 ± 0.053 0.194 ± 0.023 Er 9.935 ± 0.331 11.299 ± 0.275 8.658 ± 0.345 9.857 ± 0.241 3.323 ± 0.126 1.165 ± 0.061 Tm 1.536 ± 0.085 1.449 ± 0.041 1.285 ± 0.083 1.396 ± 0.039 0.524 ± 0.042 0.264 ± 0.025 Yb 10.835 ± 0.358 13.52 ± 0.25 9.151 ± 0.36 12.546 ± 0.233 3.049 ± 0.114 2.548 ± 0.097 Lu 0.648 ± 0.089 1.405 ± 0.077 1.235 ± 0.073 0.451 ± 0.047 0.462 ± 0.042 Springwater Farrringtonite Farringtonte Farringtonite Farringtonite Farringtonite La 0.005 ± 0.002 0.001 ± 0.001 0.002 ± 0.001 Ce 0.015 ± 0.002 0.001 ± 0.001 0.002 ± 0.001 Pr <0.001 0.007 ± 0.002 <0.001 <0.001 Nd 0.005 ± 0.002 0.028 ± 0.006 0.003 ± 0.001 0.004 ± 0.002 0.002 ± 0.001 Sm 0.003 ± 0.002 0.012 ± 0.005 0.002 ± 0.002 0.004 ± 0.002 0.003 ± 0.002 Eu 0.001 ± 0.001 0.001 ± 0.001 <0.001 0.002 ± 0.001 0.001 ± 0.001
Minor element zoning and trace element geochemistry of pallasites 1225 Table 2. REE concentrations (ppm) in pallasite phosphates. (The errors quoted are 1σ standard deviation from counting statistics only.) Continued. Springwater Farrringtonite Farringtonte Farringtonite Farringtonite Farringtonite Gd 0.005 ± 0.003 0.006 ± 0.01 0.009 ± 0.005 0.006 ± 0.004 0.005 ± 0.002 Tb 0.003 ± 0.001 0.007 ± 0.002 0.001 ± 0.001 0.004 ± 0.001 0.001 ± 0.001 Dy 0.014 ± 0.003 0.06 ± 0.009 0.022 ± 0.003 0.012 ± 0.002 0.012 ± 0.002 Ho 0.003 ± 0.001 0.017 ± 0.004 0.005 ± 0.001 0.002 ± 0.001 0.002 ± 0.001 Er 0.026 ± 0.004 0.091 ± 0.01 0.017 ± 0.002 0.014 ± 0.002 0.011 ± 0.002 Tm 0.002 ± 0.001 0.016 ± 0.002 0.006 ± 0.001 0.003 ± 0.001 0.003 ± 0.001 Yb 0.043 ± 0.004 0.18 ± 0.013 0.064 ± 0.005 0.04 ± 0.004 0.038 ± 0.004 Lu 0.01 ± 0.002 0.057 ± 0.008 0.017 ± 0.004 0.013 ± 0.002 0.012 ± 0.002 each in the other 2 grains. Within the grain, REE abundances are essentially homogeneous. Large inter-grain variations (up to a factor of 30) of REEs exist among 3 Imilac whitlockite grains. The observed REE pattern is similar in shape to that of the Albin whitlockite but with a negative Eu anomaly (Fig. 4c). In general, Imilac whitlockite is enriched in HREEs (10 to 80 CI) and relatively depleted in LREEs (0.1 to 1 CI). Four farringtonite grains in Springwater were analyzed. Two measurements were made in 1 farringtonite grain and 1 each in the other 3 grains. Springwater farringtonite grains have relatively low REEs (0.001 to 1 CI) with a highly fractionated HREE-enriched pattern (CI-normalized Lu/La ~100). Within the grain, REEs are essentially homogeneous. One grain has relatively higher REE abundances (0.02 to 2 CI) than the other 3, by a factor of 10 (Fig. 4d). All phosphates analyzed in this study are enriched in HREEs relative to LREEs. Phosphates with LREE enrichments, as reported by Davis and Olsen (1991, 1996), were not observed. Although phosphates exhibit distinct REE abundances and patterns among the pallasites studied, they tend to have essentially the same REE pattern within a given meteorite. REE concentrations in pallasite olivines are extremely low and, thus, required a large primary beam current (20 na) and long measurement time (4 hours). A total of 50 measurements were made in olivines of the 6 pallasites. Most olivines have REE concentrations below the ion microprobe detection limit. Fig. 5 shows some representative REE abundances and patterns of pallasite olivines. They are well below the 0.1 CI level with a gradual increase toward HREEs (Fig. 5). DISCUSSION Comparison with Previously Reported Minor Element Concentrations in Pallasite Olivines Minor element concentrations in pallasite olivines were previously analyzed with electron probe, instrumental neutron activation analysis (INAA), spark-source mass spectrometer, X-ray fluorescence, and ion probe techniques (Lovering 1957; Nichiporuk et al. 1967; Buseck and Goldstein 1969; Mason and Graham 1970; Goles 1971; Fig. 5. Representative REE abundances and patterns of pallasite olivines. The REEs plotted here can only be taken as upper limits. They are close or below the ion probe detection limits in most cases. Leitch et al. 1979; Reed et al. 1979; Mittlefehldt 1980; Zhou and Steele 1993; Floss 2002). Table 3 lists some of the previously published results. These data generally fall into 2 categories: bulk analyses with spark-source mass spectrometer (Mason and Graham 1970), INAA (Goles 1971; Mittlefehldt 1980) and X-ray fluorescence (Lovering 1957; Nichiporuk et al. 1967), and in situ measurements with electron probe (Buseck and Goldstein 1969; Zhou and Steele 1993) and ion probe (Leitch et al. 1979; Reed et al. 1979; Floss 2002). Relatively abundant data are available for Springwater olivine. Large discrepancies are apparent among the reported results. For example, the reported Ca concentration in Springwater olivine varies from 20 to 980 ppm and Ni from 26 to 350 ppm. Several possible reasons exists to explain such large discrepancies. Pallasite olivines are well-known to contain tubular inclusions (Buseck 1977; Steele 1994). Bulk analyses, therefore, have a tendency to yield high minor element concentrations. In Springwater olivine, the reported Ca concentration with X-ray
1226 W. Hsu Table 3. Comparison of ion probe minor element analyses (ppm) with previously reported data of pallasite olivines. Ca Ti V Cr Mn Ni Albin This study 12 70 1 18 2 8 16 158 1140 2465 6 20 Literature 60 a 15 a 150 a 2000 a 150 a Brenham This study 100 150 1 3 3 4 25 90 1500 1530 12 15 Literature 10 a 17 a ; 9.2 b 100 a ; 196 b 1500 a 350 a ; <40 b ; 22 c Eagle Station This study 92 380 1 15 1 6 13 150 1880 2370 33 74 Literature 35 d 211 ± 18 e 1250 d 33 43 c ; 55 d Glorieta Mountain This study 48 108 4 15 5 8 54 118 2290 2800 25 46 Literature 430 ± 10 e 75 500 d Imilac This study 43 90 7 25 3 8 42 173 1585 2445 7 20 Literature 28 35 c Springwater This study 65 193 1 6 4 6 30 144 3400 3750 27 50 Literature 20 d ; 980 ± 10 f ; 32 g ; 64 h ; 22 195 j 10a ; 6g ; 40i ; 2.2 5.8 j 10 a ; 10 g ; 23 i ; 100 a ; 184 ± 8 e ; 193 ± 16 f ; 185 g ; a Lovering 1957. b Mittlefehldt 1980. c Reed et al. 1979. d Buseck and Goldstein 1969. e Goles 1971. f Nichiporuk et al. 1967. g Leitch et al. 1979. h Zhou and Steele 1993. i Mason and Graham 1970. j Floss 2002. 140 h ; 270 i ; 340 490 j 2000 a ; 2300 d ; 3000 ± 30 f ; 3175 g ; 2750 h ; 2600 i ; 2630 2970 j 350 a ; 28 34 c ; 50 d ; 97 ± 10 f ; 26 g ; 110 i fluorescence is 980 ppm (Nichiporuk et al. 1967), much higher than in situ measurements (20, 32, 64 ppm) by Buseck and Goldstein (1969), Leitch et al. (1979), and Zhou and Steele (1993). The same is true for Ni concentration, for which the bulk analyses (97, 110, and 350 ppm) are higher than the probe analyses (26, 28 34, 50 ppm). On the other hand, in situ analysis, such as electron microprobe, suffers low detection limits and X-ray fluorescence effect from adjacent phases. An electron probe has the detectability limits of 15, 10, and 20 ppm for Ca, Ti, and Ni, respectively, with large precision errors from 20% to 100% (Buseck and Goldstein 1969). With its high sensitivity, ion microprobe overcomes these difficulties and yielded consistent results for pallasite olivines (Leitch et al. 1979; Reed et al. 1979). In Springwater olivine, the ion probe analyses of Ni are 26 and 28 34 ppm, respectively, by Leitch et al. (1979) and Reed et al. (1979). My results range from 27 to 50 ppm. They are in excellent agreement. Small variations could be due to either the inter-grain concentration variations or the crystal orientation and cross section position, as pallasite olivines commonly display minor element zoning. My ion probe analyses of minor element contents in pallasite olivines are broadly in line with the previously reported results. Chemical Zoning in Olivine: Diffusion and Thermal Histories of Pallasites In this study, I observed widespread minor element zoning in pallasite olivine (Fig. 2). With the high sensitivity of the ion microprobe, I am able to reveal, on a microscale, the chemical zoning of minor elements (e.g., Ti, V) with very low concentrations (~10 ppm). My results are basically consistent with the previous reports (Leitch et al. 1979; Reed et al. 1979; Zhou and Steele 1993; Miyamoto 1997). The origin of the zoning is intimately tied with the origin of the olivine itself and the process or processes by which it was incorporated into the metal. Olivine is completely solid below 1600 1700 C, well above the temperature of metal solidification, which is ~1500 to ~1000 C depending on the sulfur content (Hansen and Anderko 1958). Thus, most workers believe that pallasites formed when molten metal intruded a crystalline olivine layer, although their models
Minor element zoning and trace element geochemistry of pallasites 1227 differ in detail (e.g., Scott 1977a, b; Buseck 1977; Mittlefehldt 1980; Scott and Taylor 1990). At the time of metal intrusion, the chemical composition of the olivine would have reflected its previous igneous history. The observed chemical zoning is almost certainly not original igneous zoning. For example, the distribution coefficient for Ca in olivine relative to silicate melts is considerably less than unity (e.g., Jurewicz and Watson 1988a), which implies that the melt should increase in Ca content as it crystallizes olivine. Thus, one would expect the Ca concentration in olivine to increase as the melt crystallized, producing a profile of increasing concentration toward the crystal edge. I observe a decrease in Ca concentration from center to edge in all pallasite olivines (Fig. 2). Also, Hirschmann and Ghiorso (1994) have shown that the distribution coefficients of Ni, Mn, and Co in olivine relative to silicate melts increase with decreasing temperature at constant composition and, independently, with decreasing MgO content in the liquid. Thus, concentrations of Ni, Mn, and Co would be expected to increase toward the edges of the crystals. In contrast, I observe the concentration of Ni to decrease toward the edges of the crystals, while the Mn concentration remains essentially constant (e.g., Fig. 2). In addition, diffusion rates for the elements in this study at 1000 1500 C, which is the temperature range of molten iron, are too high to preserve unmodified igneous profiles in the olivine grains for more than a few days to a few weeks (e.g., Jurewicz and Watson 1988b). Thus, the zoning profiles probably developed by solid-state diffusion during cooling, as assumed by several previous workers (e.g., Leitch et al. 1979; Reed et al. 1979; Zhou and Steele 1993). A first-order test of the diffusion hypothesis is to invert the concentration profiles through the error function. When inverted through the error function, profiles that were produced by diffusion should plot as linear functions of distance from the crystal boundary. The current data do not permit quantitative evaluation of diffusion hypothesis because measurements were made on random slices through the crystals, not through the centers, and because the initial concentrations of the elements in the crystals are unknown at present. I now consider whether the observed zoning profiles are consistent with diffusion during cooling after the intrusion of molten iron into a previously crystallized olivine layer. Most of the elements, except for Mn, have lower concentrations at the surface of the olivine grains, consistent with loss of the elements to other phases. Ni is highly siderophile and is expected to concentrate in metal over silicates. However, the other elements do not partition preferentially into the metal. Pallasites contain several minor phases that could serve as sinks for the elements lost from the olivine. Chromite, troilite, and phosphates are common, and traces of rutile and low-ca pyroxene are also observed (e.g., Buseck 1977). These phases are often not in direct contact with the measured olivines, but this may not be necessary because grain boundary diffusion is known to be orders of magnitude faster than diffusion within a crystal (e.g., Freer 1981). Sinks would, thus, appear to be available for all of the elements depending on where they appeared in the crystallization sequence, including in the subsolidus. The outward decrease of Cr may reflect diffusion from olivine to chromite and troilite. The depletions of Ca and Ti at olivine grain boundaries indicate migration to phosphates and chromite/rutile, respectively. The zoning profile of V may suggest its siderophile nature under reducing conditions (Drake et al. 1989) or diffusion to chromite. Mn concentration sometimes increases at the very edge of the crystals. This probably reflects the strong decrease in compatibility of Mn in the metal phase with decreasing temperature coupled with the freezing in of the Mn concentration profile at low temperature. Cation diffusion in olivine was studied experimentally under various physical conditions (e.g., temperature and oxygen fugacity) (Morioka 1981; Jurewicz and Watson 1988b; Hain et al. 1996). Olivine was found to be anisotropic with respect to the diffusion of Ca, Fe, Mg, and Mn, i.e., fastest along the c-axis and slowest along the a-axis. Therefore, the diffusion-induced chemical zoning in olivine in a random section should be smaller than that along the c- axis but larger than that along the a-axis. As expected, I observed in this study that the degrees of minor element zoning (e.g., Ca and Cr) vary from grain to grain, from extensive to moderate zoning. It was found that, at a given temperature in olivine, the diffusion rates decrease significantly in the order of Fe, Mn, Co, Ni, Mg, and Ca (Morioka 1981). Cations (e.g., Mn) with higher diffusion rates more easily establish equilibrium between olivine and the receptors than those with lower rates (e.g., Ca). My results are basically consistent with this conclusion. Mn is essentially unzoned, while Ca and Cr exhibit ubiquitous zoning in pallasite olivines. Other elements such as Ni and V behave intermediately, sometimes zoned and sometimes unzoned. Further, note that in a given meteorite (e.g., Albin), the maximum concentrations of Ca and Cr at the center of olivine vary widely but Mn concentration is essentially the same. This indicates that Mn was fully equilibrated between olivine and the receptors and that Ca and Cr were not, during subsolidus cooling periods. Previous studies (Davis 1977; Scott 1977a) of olivine chemistry and Ni content in metal showed that pallasites can be divided into 2 sub-groups: main group (e.g., Albin, Brenham, Glorieta Mountain, Imilac) and Eagle Station trio (e.g., Eagle Station). In this study, I found that Springwater olivine has higher (~3500 ppm) and Brenham olivine has lower Mn concentrations (~1500 ppm) than those of Albin, Glorieta Mountain, and Imilac (~2300 ppm). This further suggests that Springwater and Brenham may be anomalous and do not fit in the main group pallasites (Davis 1977; Scott 1977a).
1228 W. Hsu In principle, diffusion-induced chemical zoning in silicates potentially provides constraints on their thermal histories. Because of the difficulties mentioned above, my data do not permit a quantitative estimation of cooling rates for pallasites. In addition, the cooling rate calculation is very sensitive to the initial temperature (T o ) and the parameters for the diffusion coefficient (D o ) and activation energy of diffusion (E) in the Arrhenius relation: D = D o exp( E/RT). Inspection of the literature shows that there are serious discrepancies between experimental studies (Morioka 1981; Jurewicz and Watson 1988b; Hain et al. 1996). In particular, the extrapolation to low temperatures is very unreliable. However, the determination of (D) at individual elevated temperatures appears to be more reliable. If one takes the result for Ni (4.7 10 13 cm 2 sec 1 ) at 1200 C from Morioka (1981), then one estimates that the maximum time that an olivine of 1 mm could remain at that temperature and not completely equilibrate Ni with the metal is ~170 yr. I noted in this study that the maximum concentrations of some minor elements at the center of olivine vary significantly from grain to grain within a given meteorite and among the pallasites studied. For example, the maximum concentration of Ca at the core of Albin olivine ranges from 25 to 70 ppm, much lower than that (180 to 380 ppm) in Eagle Station olivine. Such large variations require that the initial Ca concentration in pallasite olivines be highly heterogeneous and that the olivines have experienced rapid cooling after mixing with molten metal. For an olivine grain of 300 to 500 µm, comparable to those of Albin, the homogenization time for Ca is less than 4 yr at 1300 C and less than 8 yr at 1100 C (Jurewicz and Watson 1988b). To preserve the Ca zoning and the large variation of maximum concentrations among the different grains, Albin olivine had cooled from 1300 to 1100 C within less than 4 yr, which is equivalent to a cooling rate of >50 C/year. In evaluating the equilibration temperature for Ni between olivine and metal, Scott (1977a) and Reed et al. (1979) noted a wide range of apparent metalolivine equilibrium temperatures (800 1200 C) among pallasites. If pallasites cooled very slowly at the rates equivalent to the metallographic cooling rates (0.5 2.0 C/ Ma) obtained in pallasite Fe-Ni alloys, Ni would be fully equilibrated between olivine and metal, and a unique equilibration temperature should be expected for all pallasites. Scott (1977a) then suggested that Ni concentrations in the olivines may have been established before metal-olivine mixing and preserved by rapid cooling down to ~800 C. Evidence for fast cooling at high temperatures was also provided by Miyamoto (1997). He noted that the major and minor element zoning profiles in pallasite olivines could be understood if they experienced fast cooling (20 to 100 C/yr) through 1100 to 600 C. Evidence for the presence of live 53 Mn (Birck and Allègre 1988; Hsu et al. 1997; Lugmair and Shukolyukov 1998) and 107 Pd (Chen and Wasserburg 1996) in pallasites suggests that these meteorites formed very early, probably within the first 10 Ma of the solar system history. The Re-Os chronology of pallasites (4.60 ± 0.05 AE) further supports this suggestion (Shen et al. 1998). The mean life of 53 Mn is 5.3 Ma. Thus, if 53 Mn were alive when the molten iron was intruded, the cooling would have been fast enough to preserve the evidence in the olivines. This also implies that pallasites cooled rapidly down to the Mn-Cr closure temperature shortly after the pallasite formation. The inferred cooling rate at high temperatures is much faster than the metallographic cooling rates (0.5 2.0 C/Ma) obtained in pallasite Fe-Ni alloys. Such a fast cooling rate is required to prevent immiscible separation of the metal phase from olivines due to their large contrast in density. Despite the uncertainties in calculated cooling rates, a general consensus exists that pallasites cooled much faster at high temperatures than at low temperatures, by several orders of magnitude. However, the nature of transition from a rapid cooling at high temperatures to a very slow cooling at low temperatures is not clear. It could be smooth or, more likely, mark a catastrophic event. Thus, pallasites, like mesosiderites (Stewart et al. 1994), seem to have experienced rather complicated thermal histories involving very rapid cooling at high temperatures and slow cooling at low temperatures. REE Distributions in Pallasites Pallasite olivines probably have the lowest REE abundances among meteoritic and terrestrial olivines. A recent study indicated that pallasite olivines have extremely low REEs (10 5 to 10 2 CI; Minowa and Ebihara 2002). Ion probe analyses showed that pallasite olivines have very low REEs that are below the detection limit (Davis and Olsen 1991; this study). REE concentrations of pallasite phosphates were previously determined by Davis and Olsen (1991, 1996) with an ion probe. They found that phosphates have large variations of REE abundances (0.01 to 300 CI) and display distinct REE patterns, ranging from highly HREE-enriched to LREE-enriched. In this study, I also observed that pallasite phosphates have large variations of REEs (0.001 to 100 CI) with different REE patterns (Fig. 4). Within a given phosphate grain, REEs are generally homogeneous. REEs of phosphate may vary from grain to grain within the same meteorite by a factor of 10 to 100. Basically, the phosphates analyzed in this study have an HREE-enriched pattern with either a negative or a positive Eu anomaly (Fig. 4). But, the phosphate with an LREE-enriched pattern, as reported by Davis and Olsen (1991, 1996), was not observed. The REE abundances and patterns of pallasite phosphates potentially shed light on the process or processes that led to the formation of pallasite. Davis and Olsen (1991) suggested that the HREE-enriched pattern in phosphates is compatible with formation of this mineral by subsolidus
Minor element zoning and trace element geochemistry of pallasites 1229 redox reactions between Fe-metal and olivine. Phosphorus from schreibersite and iron metal could have been oxidized by oxygen originally dissolved in molten metal and incorporated with Ca, Mg, and Fe from olivine to form whitlockite, stanfieldite, and farringtonite (Olsen and Fredriksson 1966; Davis and Olsen 1991). In this case, phosphates would concentrate REEs from olivine and inherit its REE pattern. Davis and Olsen (1991, 1996) also found that stanfieldite and whitlockite in Springwater and Santa Rosalia are highly enriched in REEs (10 to 300 CI) with an LREE-enriched pattern. This type of pattern is expected from REE equilibrium partitioning between phosphates and silicate melts. However, the pallasite parent body mantle probably did not have enough phosphorus to crystallize phosphates when cumulate olivine formed (Davis and Olsen 1991). Davis and Olsen (1991) argued that the phosphates could have crystallized from a trapped liquid interstitial between cumulate olivines. The second alternative is that these phosphates represent the late-stage products of crystallization from a silicate melt of high density and low viscosity on pallasite parent bodies (Davis and Olsen 1991). This liquid could percolate downward but not as deep as the core-mantle boundary. This suggestion requires that pallasites formed much nearer the surface than the core-mantle boundary on their parent bodies. In pallasites, the main REE source is olivine. If all phosphates were produced in situ by redox reactions between iron metal and olivine, they are expected to have relatively low REE concentrations with an HREE-enriched pattern. The calculation by Davis and Olsen (1991) seemed to match the REE abundances and pattern observed in Eagle Station whitlockite. Davis and Olsen (1991) noted that REE partitioning between 1.2 vol% whitlockite and 98.8 vol% olivine would result in relatively low REEs (0.1 to 10 CI) in whitlockite with an HREE-enriched pattern. In their calculation, Davis and Olsen (1991) assumed that REE partitioning between phosphate and olivine was fully equilibrated under subsolidus conditions. I consider this assumption to be fundamentally flawed, and the mechanism proposed by Davis and Olsen (1991) to explain the REE patterns observed in pallasite phosphates is implausible. To see if REEs were completely equilibrated between phosphate and olivine during redox reactions, one has to consider the diffusion rates of REEs in olivine under subsolidus conditions. However, up to this date, these diffusion rates are largely unknown. A related study of selfdiffusion in haplobasaltic melts revealed that REEs have much lower diffusion rates than Mg and Ca over a temperature range of 1350 to 1500 C (LaTourrette et al. 1996). A general observation was also made that, at a given temperature, a systematic, order of magnitude decrease occurs in diffusion rates with both increasing ionic radius and increasing cation charge. If this were true for olivine, one would expect that REEs would have much lower diffusion rates than Ca at a given temperature. Yet, this mineral displays extensive and ubiquitous Ca zoning among all the pallasites studied. This indicates that Ca did not diffuse fast enough in pallasite olivine to fully equilibrate with other phases such as phosphate. It follows that REEs would not have reached equilibrium partitioning between olivine and phosphates under subsolidus conditions. The diffusion of REEs from olivine to phosphates would be negligible. Phosphates produced in situ by redox reactions between metal and olivine would be virtually free of REEs. However, it is possible that pre-existing phosphates could have participated in the redox reaction. In this case, REE abundances of the phosphate would be diluted, but the pattern would remain the same. Even if one assumes that REEs have much higher diffusion rates than Ca, so that REEs were well-equilibrated between olivine and phosphates under these conditions, it is still hard to explain how pallasite phosphates can have large inter-grain variations of REE abundances with significantly different patterns. One would expect to see that all pallasite phosphates have the same REE concentrations with an identical pattern. The same arguments can be made for the liquidus conditions. I now consider whether pallasite phosphates could have crystallized from a trapped liquid remaining from olivine formation. If olivine crystallized from a melt with chondritic REE abundances, the residual liquid would be rich in REEs with a relatively flat but slightly LREE-enriched pattern. Phosphate grows from the liquid would be rich in REEs with a relatively LREE-enriched pattern. This kind of phosphate was not observed in this study. Most pallasite phosphates have an HREE-enriched pattern. Davis and Olsen (1991) reported that some Springwater and Santa Rosalia phosphates have high REE concentrations with an LREE-enriched pattern. Thus, a small amount of pallasite phosphates likely could have crystallized from a trapped liquid interstitial between olivine grains. But, such a mechanism is not responsible for the formation of the majority of the pallasite phosphates. The metallic phase of pallasites is chemically related to IIIAB irons. It was even suggested that pallasites might have formed through the mixing of IIIAB-like molten metal with an olivine layer (Scott 1977a, b). IIIAB irons commonly contain small amounts of phosphates, such as sarcopside and graftonite ([Fe,Mn] 3 [PO 4 ] 2 ), johnsomervilleite (Na 2 Ca[Fe,Mn] 7 [PO 4 ] 6 ), and galileiite (Na 2 [Fe,Mn] 8 [PO 4 ] 6 ) (Olsen and Fredriksson 1966; Davis and Olsen 1990; Olsen et al. 1999). It is possible that phosphates were introduced into pallasites during mixing and were subsequently involved in liquidus or subsolidus redox reactions between olivine and phosphorus from metal to form pallasitic phosphates. In this case, pallasite phosphates would inherit the REE patterns of IIIAB-like phosphates. Davis and Olsen (1990) found that 4 phosphate grains in El Sampal (IIIA) have low but variable REE abundances (0.1 to 2 CI) with an LREE-enriched pattern and a large positive Eu