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YC is Zincblende, Sphalerite structured and crystallizes in the cubic F-43m space group. The structure is three-dimensional. Y3+ is bonded to four equivalent C3- atoms to form corner-sharing YC4 tetrahedra. All Y–C bond lengths are 2.37 Å. C3- is bonded to four equivalent Y3+ atoms to form corner-sharing CY4 tetrahedra.
YC is Tetraauricupride structured and crystallizes in the cubic Pm-3m space group. The structure is three-dimensional. Y3+ is bonded in a body-centered cubic geometry to eight equivalent C3- atoms. All Y–C bond lengths are 2.69 Å. C3- is bonded in a body-centered cubic geometry to eight equivalent Y3+ atoms.
Y3C4 crystallizes in the tetragonal P4/mnc space group. The structure is three-dimensional. there are four inequivalent Y3+ sites. In the first Y3+ site, Y3+ is bonded to six C+2.25- atoms to form YC6 octahedra that share corners with two equivalent YC6 octahedra, corners with four equivalent YC7 pentagonal bipyramids, and edges with four equivalent YC7 pentagonal bipyramids. The corner-sharing octahedral tilt angles are 0°. There are two shorter (2.51 Å) and four longer (2.62 Å) Y–C bond lengths. In the second Y3+ site, Y3+ is bonded to six C+2.25- atoms to form YC6 octahedra that share a cornercorner with one YC6 octahedra and edges with four equivalent YC7 pentagonal bipyramids. The corner-sharing octahedral tilt angles are 0°. There are a spread of Y–C bond distances ranging from 2.36–2.64 Å. In the third Y3+ site, Y3+ is bonded to seven C+2.25- atoms to form distorted YC7 pentagonal bipyramids that share a cornercorner with one YC6 octahedra, corners with two equivalent YC7 pentagonal bipyramids, edges with three YC6 octahedra, edges with two equivalent YC7 pentagonal bipyramids, and faces with two equivalent YC7 pentagonal bipyramids. The corner-sharing octahedral tilt angles are 1°. There are a spread of Y–C bond distances ranging from 2.50–2.78 Å. In the fourth Y3+ site, Y3+ is bonded in a 7-coordinate geometry to seven C+2.25- atoms. There are a spread of Y–C bond distances ranging from 2.36–2.86 Å. There are seven inequivalent C+2.25- sites. In the first C+2.25- site, C+2.25- is bonded to six Y3+ atoms to form CY6 octahedra that share corners with nine CY6 octahedra and edges with eight CY5C octahedra. The corner-sharing octahedra tilt angles range from 0–85°. In the second C+2.25- site, C+2.25- is bonded to six Y3+ atoms to form a mixture of corner and edge-sharing CY6 octahedra. The corner-sharing octahedra tilt angles range from 0–1°. In the third C+2.25- site, C+2.25- is bonded to six Y3+ atoms to form a mixture of corner and edge-sharing CY6 octahedra. The corner-sharing octahedra tilt angles range from 0–1°. Both C–Y bond lengths are 2.36 Å. In the fourth C+2.25- site, C+2.25- is bonded to six Y3+ atoms to form a mixture of corner and edge-sharing CY6 octahedra. The corner-sharing octahedra tilt angles range from 0–1°. There are a spread of C–Y bond distances ranging from 2.36–2.64 Å. In the fifth C+2.25- site, C+2.25- is bonded to five Y3+ and one C+2.25- atom to form a mixture of corner and edge-sharing CY5C octahedra. The corner-sharing octahedra tilt angles range from 3–85°. The C–C bond length is 1.35 Å. In the sixth C+2.25- site, C+2.25- is bonded in a 2-coordinate geometry to four equivalent Y3+ and two equivalent C+2.25- atoms. In the seventh C+2.25- site, C+2.25- is bonded in a 6-coordinate geometry to five Y3+ and one C+2.25- atom. The C–C bond length is 1.30 Å.
Y4C7 crystallizes in the monoclinic P2_1/c space group. The structure is three-dimensional. there are four inequivalent Y3+ sites. In the first Y3+ site, Y3+ is bonded to six C+1.71- atoms to form distorted YC6 octahedra that share corners with two equivalent YC7 octahedra and edges with seven YC6 octahedra. The corner-sharing octahedral tilt angles are 2°. There are a spread of Y–C bond distances ranging from 2.52–2.65 Å. In the second Y3+ site, Y3+ is bonded in a 8-coordinate geometry to eight C+1.71- atoms. There are a spread of Y–C bond distances ranging from 2.43–2.70 Å. In the third Y3+ site, Y3+ is bonded in a 8-coordinate geometry to nine C+1.71- atoms. There are a spread of Y–C bond distances ranging from 2.41–2.94 Å. In the fourth Y3+ site, Y3+ is bonded to seven C+1.71- atoms to form distorted YC7 octahedra that share corners with two equivalent YC6 octahedra, a cornercorner with one CY4C trigonal bipyramid, and edges with five YC6 octahedra. The corner-sharing octahedral tilt angles are 2°. There are a spread of Y–C bond distances ranging from 2.48–3.06 Å. There are seven inequivalent C+1.71- sites. In the first C+1.71- site, C+1.71- is bonded in a 6-coordinate geometry to four Y3+ and two C+1.71- atoms. Both C–C bond lengths are 1.33 Å. In the second C+1.71- site, C+1.71- is bonded in a 4-coordinate geometry to three Y3+ and two C+1.71- atoms. There is one shorter (1.32 Å) and one longer (1.35 Å) C–C bond length. In the third C+1.71- site, C+1.71- is bonded to six Y3+ atoms to form CY6 octahedra that share a cornercorner with one CY4C trigonal bipyramid, edges with six CY6 octahedra, and edges with two equivalent CY4C trigonal bipyramids. In the fourth C+1.71- site, C+1.71- is bonded in a 4-coordinate geometry to four Y3+ and one C+1.71- atom. In the fifth C+1.71- site, C+1.71- is bonded in a 5-coordinate geometry to four Y3+ and one C+1.71- atom. In the sixth C+1.71- site, C+1.71- is bonded to four Y3+ and one C+1.71- atom to form distorted CY4C trigonal bipyramids that share a cornercorner with one YC7 octahedra, corners with five CY6 octahedra, corners with two equivalent CY4C trigonal bipyramids, and edges with three CY6 octahedra. The corner-sharing octahedra tilt angles range from 2–96°. In the seventh C+1.71- site, C+1.71- is bonded to five Y3+ and one C+1.71- atom to form distorted CY5C octahedra that share corners with four equivalent CY4C trigonal bipyramids, edges with four CY6 octahedra, and an edgeedge with one CY4C trigonal bipyramid.
CI1 is Cyanogen Chloride-derived structured and crystallizes in the tetragonal P4_2/n space group. The structure is zero-dimensional and consists of eight diiodoacetylene molecules. there are two inequivalent C sites. In the first C site, C is bonded in a linear geometry to one C and one I atom. The C–C bond length is 1.23 Å. The C–I bond length is 1.99 Å. In the second C site, C is bonded in a linear geometry to one C and one I atom. The C–I bond length is 1.99 Å. There are two inequivalent I sites. In the first I site, I is bonded in a single-bond geometry to one C atom. In the second I site, I is bonded in a single-bond geometry to one C atom.
Y15C19 crystallizes in the tetragonal P-42_1c space group. The structure is three-dimensional. there are five inequivalent Y3+ sites. In the first Y3+ site, Y3+ is bonded to six C+2.37- atoms to form a mixture of edge and corner-sharing YC6 octahedra. The corner-sharing octahedra tilt angles range from 0–1°. There are a spread of Y–C bond distances ranging from 2.51–2.71 Å. In the second Y3+ site, Y3+ is bonded in a 7-coordinate geometry to seven C+2.37- atoms. There are a spread of Y–C bond distances ranging from 2.34–2.78 Å. In the third Y3+ site, Y3+ is bonded in a 7-coordinate geometry to seven C+2.37- atoms. There are a spread of Y–C bond distances ranging from 2.34–2.82 Å. In the fourth Y3+ site, Y3+ is bonded to six C+2.37- atoms to form a mixture of edge and corner-sharing YC6 octahedra. The corner-sharing octahedral tilt angles are 0°. There are a spread of Y–C bond distances ranging from 2.59–2.63 Å. In the fifth Y3+ site, Y3+ is bonded to six C+2.37- atoms to form a mixture of edge and corner-sharing YC6 octahedra. The corner-sharing octahedral tilt angles are 0°. There are two shorter (2.53 Å) and four longer (2.63 Å) Y–C bond lengths. There are six inequivalent C+2.37- sites. In the first C+2.37- site, C+2.37- is bonded to six Y3+ atoms to form CY6 octahedra that share corners with six CY6 octahedra and edges with twelve CY5C octahedra. The corner-sharing octahedra tilt angles range from 0–1°. In the second C+2.37- site, C+2.37- is bonded to five Y3+ and one C+2.37- atom to form distorted CY5C octahedra that share corners with seven CY5C octahedra and edges with eight CY6 octahedra. The corner-sharing octahedra tilt angles range from 0–90°. The C–C bond length is 1.35 Å. In the third C+2.37- site, C+2.37- is bonded in a 6-coordinate geometry to four Y3+ and two C+2.37- atoms. The C–C bond length is 1.34 Å. In the fourth C+2.37- site, C+2.37- is bonded to five Y3+ and one C+2.37- atom to form distorted CY5C octahedra that share corners with seven CY5C octahedra and edges with eight CY6 octahedra. The corner-sharing octahedra tilt angles range from 2–92°. In the fifth C+2.37- site, C+2.37- is bonded to six Y3+ atoms to form a mixture of edge and corner-sharing CY6 octahedra. The corner-sharing octahedra tilt angles range from 0–92°. In the sixth C+2.37- site, C+2.37- is bonded to six Y3+ atoms to form a mixture of edge and corner-sharing CY6 octahedra. The corner-sharing octahedra tilt angles range from 0–1°.
We have determined the Rb-Sr age of basaltic shergottite NWA 1460 to be 312 +/- 3 Ma, and the Sm-Nd age to be 352 +/- 30 Ma. The initial Sr and Nd isotopic compositions of NWA 1460 suggest it is an earlier melting product of a Martian mantle source region similar to those of the Iherzolitic shergottites and basaltic shergottite EETA79001, lithology B. The new ages of NWA 1460 and other recently analyzed Martian meteorites leads us to reexamine the paradox that most of the Martian meteorites appear to be younger from the majority of the Martian surface. This paradox continues to pose a challenge to determining a reliable Martian chronology.
The Sm-Nd and Rb-Sr ages and also the initial Nd and Sr isotopic compositions of MIL 05035 are the same as those of A-881757. Comparing the radiometric ages of these meteorites to lunar surface ages as modeled from crater size-frequency distributions as well as the TiO2 abundances and initial Sr-isotopic compositions of other basalts places their likely place of origin as within the Australe or Humboldtianum basins. If so, a fundamental west-east lunar asymmetry in compositional and isotopic parameters that likely is due to the PKT is implied.
Chronological information about Martian processes comes from two sources: Crater-frequency studies and laboratory studies of Martian meteorites. Each has limitations that could be overcome by studies of returned Martian rocks and soils. Chronology of Martian volcanism: The currently accepted chronology of Martian volcanic surfaces relies on crater counts for different Martian stratigraphic units [1]. However, there is a large inherent uncertainty for intermediate ages near ~2 Ga ago. The effect of differing preferences for Martian cratering chronologies [1] is shown in Fig. 1. Stoeffler and Ryder [2] summarized lunar chronology, upon which Martian cratering chronology is based. Fig. 2 shows a curve fit to their data, and compares to it a corresponding lunar curve from [3]. The radiometric ages of some lunar and Martian meteorites as well as the crater-count delimiters for Martian epochs [4] also are shown for comparison to the craterfrequency curves. Scaling the Stoeffler-Ryder curve by a Mars/Moon factor of 1.55 [5] places Martian shergottite ages into the Early Amazonian to late Hesperian epochs, whereas using the lunar curve of [3] and a Mars/Moon factor ~1 consigns the shergottites to the Middle-to-Late Amazonian, a less probable result. The problem is worsened if a continually decreasing cratering rate since 3 Ga ago is accepted [6]. We prefer the adjusted St ffler-Ryder curve because it gives better agreement with the meteorite ages (Fig.
Here, we compare Sm-147-Nd-143 and Sm-146-Nd-142 data for lunar norite 78236 to those for approximately 4.54-4.56 Ga old cumulate eucrite Yamato 980318/433 and show that the norite data are compatible with its derivation from an isotopic reservoir similar to that from whence the eucrite pair came. Thus, lunar-like Sm-Nd isotopic systematics are not unique to the Earth-Moon system.
Preliminary Rb-Sr and Sm-Nd ages reported by [1] for the NWA 1460 basaltic shergottite are refined to 336+/-14 Ma and 345+/-21 Ma, respectively. These concordant ages are interpreted as dating a lava flow on the Martian surface. The initial Sr and Nd isotopic compositions of NWA 1460 suggest it is an earlier melting product of a Martian mantle source region similar to those of the lherzolitic shergottites and basaltic shergottite EETA79001, lithology B. We also examine the suggestion that generally "young" ages for other Martian meteorites should be reinterpreted in light of Pb-207/Pb-206 - Pb-204/Pb-206 isotopic systematics [2]. Published U-Pb isotopic data for nakhlites are consistent with ages of approx.1.36 Ga. The UPb isotopic systematics of some Martian shergottites and lherzolites that have been suggested to be approx.4 Ga old [2] are complex. We nevertheless suggest the data are consistent with crystallization ages of approx.173 Ma when variations in the composition of in situ initial Pb as well as extraneous Pb components are considered.
Chondrules and chondrites are interpreted as objects formed in the early solar system, and it is important to study them in order to elucidate its evolution. Here, we report the study of the Mn-Cr systematics of the R-Chondrite NWA753 and compare the results to other chondrite data. The goal was to determine Cr isotopic and age variations among chondrite groups with different O-isotope signatures. The Mn-53-Cr-53 method as applied to individual chondrules [1] or bulk chondrites [2] is based on the assumption that 53Mn was initially homogeneously distributed in that portion the solar nebula where the chondrules and/or chondrites formed. However, different groups of chondrites formed from regions of different O-isotope compositions. So, different types of chondrites also may have had different initial Mn-53 abundances and/or Cr isotopic compositions. Thus, it is important to determine the Cr isotopic systematics among chondrites from various chondrite groups. We are studying CO-chondrite ALH83108 and Tagish Lake in addition to R-Chondrite NWA753. These meteorites have very distinct O-isotope compositions (Figure 1).
Significantly large mass fractionations between chlorine isotopes (Cl-35, Cl-37) have been reported for terrestrial materials including both geological samples and laboratory materials. Also, the chlorine isotopic composition can be used as a tracer for early solar system processes. Moreover, chlorine is ubiquitous on the Martian surface. Typical chlorine abundances in Gusev soils are approx.0.5 %. The global surface average chlorine abundance also is approx.0.5 %. Striking variations among outcrop rocks at Meridiani were reported with some chlorine abundances as high as approx.2%. Characterizing conditions under which chlorine isotopic fractionation may occur is clearly of interest to planetary science. Thus, we have initiated development of a chlorine isotopic analysis technique using TIMS at NASA-JSC. We present here a progress report on the current status of development at JSC and discuss the possible application of chlorine isotopic analysis to Martian meteorites in a search for fluid- and possibly biological activity on Mars.
A significantly large mass fractionation between two stable chlorine isotopes is expected during planetary processes In addition, in view of the isotopic heterogeneity of other light elements, the chlorine isotopes can potentially be used as a tracer for the origins and evolutionary processes of early solar system materials. Due to analytical difficulties, however, current chlorine isotope studies on planetary materials are quite controversial among IRMS (gas source mass spectrometry) and/or TIMS (Thermal Ionization Mass Spectrometry) groups [i.e. 1-3]. Although a cross-calibration of IRMS and TIMS indicates that both techniques are sufficiently consistent with each other [4], some authors have claimed that the Cl-37/Cl-35 ratio of geological samples obtained by TIMS technique are, in general, misleadingly too high and variable compared to those of IRMS [3]. For example, almost no differences of Cl isotope composition were observed among mantle materials and carbonaceous meteorites by [3]. On the other hand, according to more recent IRMS work [2], significant Cl isotope variations are confirmed for mantle materials. Therefore, additional careful investigation of Cl isotope analyses are now required to confirm real chlorine isotope variations for planetary materials including carbonaceous chondrites [5]. A significantly large mass fractionation between two stable chlorine isotopes is expected during planetary processes In addition, in view of the isotopic heterogeneity of other light elements, the chlorine isotopes can potentially be used as a tracer for the origins and evolutionary processes of early solar system materials. Due to analytical difficulties, however, current chlorine isotope studies on planetary materials are quite controversial among IRMS (gas source mass spectrometry) and/or TIMS (Thermal Ionization Mass Spectrometry) groups [i.e. 1-3]. Although a cross-calibration of IRMS and TIMS indicates that both techniques are sufficiently consistent with each other [4], some authors have claimed that the 37Cl/35Cl ratio of geological samples obtained by TIMS technique are, in general, misleadingly too high and variable compared to those of IRMS [3]. For eample, almost no differences of Cl isotope composition were observed among mantle materials and carbonaceous meteorites by [3]. On the other hand, according to more recent IRMS work [2], significant Cl isotope variations are confirmed for mantle materials. Therefore, additional careful investigation of Cl isotope analyses are now required to confirm real chlorine isotope variations for planetary materials including carbonaceous chondrites [5]. In order to clarify the stable chlorine isotope features of early solar system materials, we have initiated development of the TIMS technique at NASA JSC applicable to analysis of small amounts of meteoritic and planetary materials. We report here the current status of chlorine isotope analysis at NASA JSC.
We have established a precise analytical technique for stable chlorine isotope measurements of tiny planetary materials by TIMS (Thermal Ionization Mass Spectrometry) [1], for which the results are basically consistent with the IRMS tech-nique (gas source mass spectrometry) [2,3,4]. We present here results for Martian shergottites and nakhlites; whole rocks, HNO3-leachates and residues, and discuss the chlorine isotope evolution of planetary Mars.
The Asuka 881394 achondrite contains fossil Al-26 and Mn-53 and has a Pb-206/Pb-207 age of 4566.5 +/- 0.2 Ma, the oldest for an achondrite. Recent re-investigation of A881394 yielded revised initial Sm-146/Sm-144 = (9.1 +/- 1.4) x 10(exp -3), a Sm-147-Nd-143 age of 4525 +/- 58 Ma, a Rb-87-Sr-87 age of 4490 +/- 130 Ma, and initial Sr-87/Sr-86 = 0.698991 +/- 19, respectively. The relatively large uncertainties in the Sm-Nd and Rb-Sr ages are due to disturbances of the isotopic systematics of tridymite and other minor phases. A preliminary value for the Sm-147-Nd-143 age of the Yamato 980318 cumulate eucrite of 4560 +/- 150 Ma was refined in later work to 4567 +/- 24 Ma as reported orally at LPSC 35. Similarly, a preliminary value for Sm-146/Sm-144 = (7.7 +/- 1.2) x 10 (exp -3) was refined to (6.0 +/- 0.3) x 10(exp -3). For Yamato 980433, a Sm-147-Nd-143 age of 4542 +/-42 Ma and Sm-146/Sm-144 = (5.7 +/- 0.5) x 10(exp -3) has been reported. Because these two cumulate eucrites are paired, we consider them to represent one igneous rock and present their combined isotopic data here.
Textural variations in the shergottite Zagami were initially interpreted as evidence that it formed in a heterogeneous lava flow. Variations in initial Sr-87/Sr-86 ratios between a Coarse Grained (CG) and a Fine Grained (FG) lithology and evidence for more extensive fractionation of the Rb/Sr ratio in a Dark Mottled Lithology (DML) are consistent with such an interpretation. More recently, Niihara et al. and Misawa et al. have reported the mineralogy and Sr-isotopic systematics of an Olivine Rich Lithology (ORL) found in association with the coarse-grained DML lithology in the Kanagawa Zagami specimen [6,7]. Here we call this lithology DML(Ka) to maintain a distinction with DML(USNM) as studied. An Ar-Ar study by Park et al. of a late stage K-rich melt enriched in K2O to approx 7% and intruded into ORL yielded an Ar-Ar age of 202+/0 7 Ma. The present work extends the study of Kanagawa Zagami to Nd-isotopes.
Certain anorthositic rocks that are rare in the returned lunar samples have been identified among lunar meteorites. The variety of anorthosites in the Apollo collection also is more varied than is widely recognized. James eta. identified three lithologies in a composite clast o ferroan anorthosite (FAN)-suite rocks in lunar breccia 64435. They further divided all FANs into four subgroups: anorthositic ferroan (AF), mafic magnesian (MM), mafic ferroan (MF), and anorthositic sodic (AS, absent in the 64435 clast). Here we report Sm-Nd isotopic studies of the lithologies present in the 64435 composite clast and compare the new data to our previous data for lunar anorthosites incuding lunar anorthositic meteorites. Mineralogy-petrography, in situ trace element studies, Sr-isotope studies, and Ar-Ar chronology are included, but only the Nd-isotopic studies are currently complete.