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Abstract

  1. Top of page
  2. Abstract
  3. Introduction
  4. Analytical Methods
  5. Results
  6. Discussion
  7. Conclusions
  8. References

Abstract– Dhofar 458 is a lunar meteorite consisting mainly of olivine-plagioclase intergrowths, pyroxene-plagioclase intergrowths, and plagioclase fragments. Pyroxene-plagioclase globules are also common. In this study, we report the discovery of a polycrystalline zircon in this lunar meteorite. The polycrystalline zircon contains small vesicles and rounded baddeleyite grains at its margin. The polycrystalline and porous texture of the zircon indicates high-pressure shock-induced melting and degassing. Baddeleyite grains are derived from decomposition of zircon under high postshock temperature. The shock features in zircon indicates that the shock pressure in Dhofar 458 was greater than approximately 60 GPa and the postshock temperature greater than approximately 1700 °C. The polycrystalline and degassing texture and decomposition zircon also strongly indicates that Dhofar 458 is a clast-rich impact melt rock. During this shock event, most components were melted and grains of mafic minerals are interstitial to lath-like plagioclase grains. Large fragments of olivine and chromite also formed polycrystalline texture at margins and chemically reequilibrated with surrounding melts. We suggest that pyroxene-plagioclase globules could be remains of melted target clasts, whereas vesicles may form during shock-induced degassing of the rock. The U-Pb isotopic data plot on a well-defined discordant line, yielding the age of the zircon of 3434 ± 15 Ma (2σ). This age is interpreted as the time of the impact event that melted Dhofar 458 and caused decomposition and recrystallization of this zircon in Dhofar 458, which reset this zircon’s U-Pb age.


Introduction

  1. Top of page
  2. Abstract
  3. Introduction
  4. Analytical Methods
  5. Results
  6. Discussion
  7. Conclusions
  8. References

Dhofar 458 is a feldspathic lunar meteorite found in 2001. It is paired with Dhofar 026 and Dhofar 457–468 based on their petrologic and mineralogic similarities (Warren et al. 2005). Oxygen isotopic compositions indicate Dhofar 026 has a lunar origin (Taylor et al. 2001). These lunar meteorites have a distinctly different petrographic texture from most other feldspathic lunar meteorites. Up to date, only a few investigations were performed on Dhofar 026 and most of them are conference abstracts (Cohen et al. 2001, 2002, 2004; James et al. 2003, 2007; Taylor et al. 2001; Warren et al. 2001, 2005; Fernandes et al. 2004; Demidova et al. 2007). These investigations show that Dhofar 026 was heavily shocked; however, its complex texture obscures the nature of the shock metamorphic events that affected it. Based on comparison with Apollo sample 15418, Cohen et al. (2004) and James et al. (2007) suggested that Dhofar 026 is a strongly shocked granulitic breccia or a strongly shocked fragmental breccia consisting almost entirely of granulitic-breccia clasts. Cohen et al. (2004) interpreted that plagioclase in Dhofar 026 is devitrified maskelynite. Based on this interpretation, they infered that the shock pressure Dhofar 026 experienced was between 30 and 45 GPa and that the postshock temperature was about 1200 °C. However, Warren et al. (2005) suggested that Dhofar 026 was an unusual variety of impact melt breccia (now called “impact melt rock,”Stöffler and Grieve 2007) whose precursor could be a regolith breccia. A key difference between Cohen et al. (2004) and Warren et al.’s (2005) petrographic interpretations of Dhofar 026 is the heat source, which melted this rock (James et al. 2007). Cohen et al. (2004) and James et al. (2007) suggested that the major evidence of a heavy shock event in Dhofar 026 is the fact that the plagioclase is devitrified maskelynite and the melting in Dhofar 026 was caused by thermal conduction from an external, hotter source. However, Warren’s petrologic interpretation of Dhofar 026 is based on the evidence for an important melt component and the strong likelihood that an impact was the proximal cause of melting although no regions were explicitly pointed out as the impact melt (James et al. 2007).

Zircon has been used as a shock indicator for rocks from meteorite impact craters on Earth (e.g., El Goresy 1965; Bohor et al. 1993; Wittmann et al. 2006, 2009). With increasing degree of shock metamorphism, zircon may show a variety of different shock-induced features, such as reidite and planar microstructures, polycrystalline textures, and decomposition to PbO2 and silica (El Goresy 1965, 1968; Chao 1968; Bohor et al. 1993; Corfu et al. 2003; Wittmann et al. 2006, 2009). Besides as a shock indicator, zircon is particularly suitable for U-Pb geochronology because it can incorporate trace amounts of U and Th, but little or no Pb (Ireland and Williams 2003). A few authors have studied the effect of shock metamorphism of zircon on the disturbance of U-Pb isotopic system. Deutsch and Schärer (1990) stated that the U-Pb isotopic system of zircon shocked in laboratory experiments up to 59 GPa were not appreciately fractionated. In contrast, the U-Pb isotopic systems of shocked zircon from terrestrial impactites were found to be partly or totally reset (e.g., Kamo and Krogh 1995; Kamo et al. 1996; Gibson et al. 1997; Deloule et al. 2001; Mänttäri and Koivisto 2001; Kalleson et al. 2009). Zircon grains had also observed in lunar rocks and lunar meteorites and were dated for U-Pb isotopic ages (e.g., Meyer et al. 1996; Gnos et al. 2004; Pidgeon et al. 2007; Nemchin et al. 2008, 2009a, 2009b; Grange et al. 2009). Although a few U-Pb isotopic ages of zircons were considered as the time of impact events (e.g., Gnos et al. 2004; Liu et al. 2010a), however, only a few studies reported shock-induced features of zircon (Pidgeon et al. 2007). Pidgeon et al. (2007) described a zircon aggregate from lunar breccia 73235 that contains shock-induced parallel deformation features. They further attributed the younger one of two ages (4.31 Ga and 4.18 Ga) of the zircon aggregate to a severe shock event. In addition, almost all U-Pb ages of zircons from lunar rocks and lunar meteorites are older than approximately 3.9 Ga, which was considered as the time of bombardment in the Earth–Moon system (Cohen et al. 2000). Although seven to nine different impact events were suggested by Cohen et al. (2000) based on their Ar-Ar ages, no younger ages of zircons than approximately 3.9 Ga were reported. It leads to a question whether impact events after the bombardment did not have enough energy to disturb the U-Pb isotopic system of zircon.

In this study, we report a polycrystalline zircon in Dhofar 458 which not only shows a degassing texture but also contains baddeleyite grains at its margin. The polycrystalline and degassing texture and decomposition of the zircon indicates that Dhofar 458 is a clast-rich impact melt rock. The U-Pb isotopic compositions of the zircon were also determined by using ion microprobe to check whether the U-Pb isotopic system was partly or totally reset by the impact event.

Analytical Methods

  1. Top of page
  2. Abstract
  3. Introduction
  4. Analytical Methods
  5. Results
  6. Discussion
  7. Conclusions
  8. References

Petrography of Dhofar 458 was studied using the Hitachi S-3400N II scanning electron microscope (SEM) at Purple Mountain Observatory. Backscattered electron (BSE) images were taken using the same SEM. Mineral compositions of Dhofar 458 were measured with the JEOL 8100 electron microprobes at Nanjing University and the China University of Geosciences, Wuhan. The operating conditions are 15 kV accelerating voltage and 20 nA beam current. Peak and background counting times for most elements are 10 s and 5 s, respectively. Natural and synthetic minerals were used as standards. Typical detection limits for oxides of most elements are 0.03 wt%. All data were reduced by ZAF procedures.

Raman spectra of zircon and baddeleyite were acquired at Nanjing University with a Renishaw RM2000 micro-Raman spectrometer. The operating conditions were: 514 nm Ar+ laser, 5 mW laser energy, and 25 μm spectral slit. The 50× objective was used on a Leica DM/LM microscope. With this objective, the lateral spot-size of the laser beam was about 1 μm. The acquisition time for each spectrum was 100 s. Silicon (520 cm−1 Raman shift) was used as a standard.

The U-Pb dating of the zircon in Dhofar 458 was performed by using the Cameca IMS-1280 secondary ion mass spectrometer (SIMS) at the Institute of Geology and Geophysics of the Chinese Academy of Sciences following the procedure of Liu et al. (2010b). The O2primary ion beam was accelerated at −13 kV, with an intensity of approximately 220 pA. The ellipsoidal spot is about 5 × 8 μm in size and a Gaussian illumination mode was used in order to obtain this small beam size. Positive secondary ions were extracted with a 10 kV potential. Oxygen flooding was used to increase the O2 pressure to approximately 5 × 10−6 Torr in the sample chamber to enhance Pb+ sensitivity (Li et al. 2010). Single electron multiplier was used to measure secondary ion beam intensities by peak-jumping in sequence: 195.94 (180Hf 16O, high intensity peak for centering the secondary ion beam as well as energy and mass), 199.80 (92Zr216O, reference mass), 200.50 (background of detector), 203.97 (204Pb), 205.97 (206Pb), 206.98 (207Pb), 207.98 (208Pb), 238.05 (238U), 248.03 (232Th16O), 270.04 (238U16O2). Each measurement contains 12 cycles, with a total analytical time of approximately 15 min. A nuclear magnetic resonance magnet controller was used to stabilize the magnetic field. The mass resolution power is fixed to 7000. The Pb/U ratios were calibrated using an empirical correlation between Pb+/U+ and UO2+/U+ ratios, normalized to the 561.3 Ma zircon standard M257 (Nasdala et al. 2008). Before measuring secondary ions, an area of 20 × 20 μm was sputtered by a primary ion beam of 3 nA to remove possible contaminants. Correction of common Pb was made by measuring 204Pb and using a common lead composition of 206Pb/204Pb = 18 ± 2 and 207Pb/206Pb = 0.858 ± 0.2 (cf. Nemchin et al. 2009a, 2009b).

Results

  1. Top of page
  2. Abstract
  3. Introduction
  4. Analytical Methods
  5. Results
  6. Discussion
  7. Conclusions
  8. References

Petrography and Mineralogy

Dhofar 458 consists mainly of olivine-plagioclase intergrowths, pyroxene-plagioclase intergrowths, and plagioclase fragments (Fig. 1). Olivine-plagioclase intergrowth and pyroxene-plagioclase intergrowth usually have irregular outlines and the boundaries between them could not be precisely determined. Pyroxene-plagioclase globules, large grains of olivine, and ulvöspinel-chromite-spinel solid solution are scattered throughout the whole section studied. A few metal-sulfide eutectics (5–30 μm in size) were also observed in this rock. The rock is tough; however, abundant vesicles occur in the rock. The size of vesicle varies from submicron to approximately 200 μm.

image

Figure 1.  Backscattered electron image of a typical area in Dhofar 458 conaining both olivine-plagioclase intergrowth (ol-pl intergrowth) and pyroxene-plagioclase intergrowth (px-pl intergrowth). A pyroxene-plagioclase globule (px-pl globule) with a subophitic texture contains a large, rounded vesicle.

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Olivine-plagioclase intergrowth consists dominantly of olivine and plagioclase. At first glance, olivine-plagioclase intergrowth shows a granulitic texture. However, detailed observations show that olivine grains in this kind of intergrowth usually have irregular rims and small acicular olivine grains (<5 μm in length) are interstitial to the plagioclase (Fig. 2a). This texture could be due to melting and recrystallization of plagioclase and olivine. Plagioclase grains grow fast, and olivine grains grow later with shapes conforming to the lath-like plagioclase grains. Olivine grains in olivine-plagioclase intergrowth have a limited variation of Mg# [≡100 × Mg/(Mg+Fe)] value from 63.2 to 67.5 (Table 1). Plagioclase in this intergrowth is calcic and ranges from An96.0 to An98.1 (Table 2).

image

Figure 2.  Backscattered electron images of lithologies in Dhofar 458. a) Plagioclase (Pl) surrounding olivine (Ol) in olivine-plagioclase intergrowth usually contains small acicular olivine grains (indicated by arrows). b) Pyroxene (Px) and plagioclase grains form a sieve-like texture in the domain between the pyroxene-plagioclase (px-pl) intergrowth and globule. Plagioclase laths penetrate into pyroxene grains and the rims of globules. c) At the margin of a pyroxene-plagioclase globule, plagioclase laths are perpendicular to and penetrating into the boundary of pyroxene-plagioclase globules. d) A close-up view of a pyroxene-plagioclase globule in which olivine and pyroxene grains are interstitial to the plagioclase. A few Si-rich blebs occur at contacts between minerals.

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Table 1.   Representative compositions of olivine in Dhofar 458.
 Pyroxene-plagioclase intergrowthPyroxene-plagioclase globuleOlivine-plagioclase intergrowthOlivine fragment
CoreRimCoreRimCoreRim
Major elements, wt% (electron microprobe)
SiO239.438.338.339.036.636.536.238.1
TiO20.03<0.030.030.050.070.04<0.03<0.03
Al2O30.200.030.050.080.080.11<0.030.12
Cr2O30.240.230.210.350.070.100.100.23
MgO44.338.741.633.932.131.731.640.5
FeO14.221.718.526.329.730.229.818.7
MnO0.170.290.210.410.330.310.340.19
CaO0.410.400.350.440.480.410.240.49
Total98.999.699.3100.599.599.398.398.4
Mg#84.976.380.269.966.165.465.679.5
Table 2.   Representative compositions of plagioclase in Dhofar 458.
 FragmentPyroxene-plagioclase globuleOlivine-plagioclase intergrowthPyroxene-plagioclase intergrowth
Major elements, wt% (electron microprobe)
SiO243.343.945.647.442.843.644.744.6
Al2O335.835.833.831.836.635.833.934.5
MgO0.250.200.550.900.060.150.280.17
FeO0.370.330.660.870.220.510.570.47
CaO20.319.919.618.820.620.019.419.5
Na2O0.300.440.480.600.220.350.350.19
K2O<0.030.04<0.03<0.03<0.030.040.05<0.03
Total100.3100.6100.7100.4100.5100.599.399.4
An97.395.995.894.498.196.796.698.1
Ab2.63.84.25.51.93.03.11.7
Or0.10.20.00.10.10.20.30.1

Pyroxene-plagioclase intergrowth has a subophitic texture and consists mainly of anorthite and pyroxene with minor olivine (Fig. 2b). The outline of pyroxene-plagioclase intergrowth is irregular and pyroxene and plagioclase grains around the intergrowth sometimes form a sieve-like texture (Fig. 2b). A few submicron Si-rich blebs were observed in this kind of intergrowth, but are rare. Olivine and pyroxene grains in pyroxene-plagioclase intergrowth are usually zoned and relatively Fe-rich near grain margins. The Mg# value of olivine ranges from 66.7 to 84.9. Pyroxene in this kind of intergrowth has a large compositional varation from pigeonite through subcalcic augite to augite. No orthopyroxene was detected. Pigeonite (En57.4–60.5Fs26.3–28.2Wo12.8–14.6) has a limited Mg# range (67.1–69.3). Both subcalcic augite (En47.1–63.0Fs18.5–35.2Wo15.3–24.4) and augite (En28.1–55.2Fs18.9–37.2Wo25.7–41.0) have a large range of Mg# value (57.8–77.2 and 43.0–74.5, respectively). Pyroxene in pyroxene-plagioclase intergrowths contains 2.3–6.8 wt% Al2O3, 0.4–5.9 wt% TiO2, and 0.1–1.5 wt% Cr2O3 (Table 3). The compositional range of anorthite grains in this intergrowth is An93.7 to An98.1.

Table 3.   Representative compositions of pyroxene in Dhofar 458.
 Pyroxene-plagioclase intergrowthPigeonitePyroxene-plagioclase globule
AugiteSubcalcic augiteSubcalcic augiteAugite
Major elements, wt% (electron microprobe)
SiO244.751.048.950.348.251.447.952.149.750.445.849.949.5
TiO24.850.591.410.722.290.650.600.771.040.893.341.401.31
Al2O35.313.823.524.415.593.162.962.813.164.134.578.685.30
Cr2O30.931.040.101.451.401.060.400.640.450.810.410.771.57
MgO12.118.78.8619.422.121.722.819.814.519.68.0011.015.9
FeO12.711.521.111.511.713.518.115.919.6813.121.011.49.28
MnO0.280.290.330.340.320.360.430.370.450.330.390.250.22
CaO18.612.315.311.29.137.366.736.8410.810.615.817.016.2
Na2O<0.03<0.030.070.05<0.03<0.03<0.030.04<0.03<0.030.060.120.04
Total99.599.299.699.4100.799.299.999.399.899.999.4100.599.3
Mg#63.174.543.075.277.274.369.369.257.072.840.663.575.5
Wo41.025.934.723.618.615.312.814.623.222.136.541.335.5
En37.355.228.157.462.963.060.559.143.856.825.837.248.7
Fs21.818.937.218.918.521.826.826.333.021.237.721.415.8

Pyroxene-plagioclase globules are composed of olivine, pyroxene, and plagioclase. They show a subophitic texture. Most pyroxene-plagioclase globules have rounded outline and their diameter varies from approximately 100 μm to approximately 600 μm. A few compound globules were also observed. Although at first glance, the boundary between globule and surrounding lithology is smooth, detailed observations show that plagioclase laths penetrate into the margin of pyroxene-plagioclase globules in most cases (Figs. 2b and 2c). Pyroxene and plagioclase around pyroxene-plagioclase globules usually form a sieve-like texture (Fig. 2b). Within pyroxene-plagioclase globules, abundant Si-rich blebs occur along boundaries between olivine, pyroxene, and plagioclase (Fig. 2d). However, most of these blebs are too small (submicron in size) for quantitative analyses. Olivine grains in pyroxene-plagioclase globules are usually zoned with Mg# value ranging from 66.4 to 82.6. Pyroxene in pyroxene-plagioclase globules is augite-dominant and subcalcic augite is minor. The Mg# value of augite (En16.2–52.1Fs13.2–58.6Wo25.1–41.3) varies from 21.7 to 78.2. The two subcalcic augite data points (En43.8–56.8Fs21.2–33.0Wo22.1–23.2) show a large compositional variation (Mg# = 57.0–72.8). Pyroxene in pyroxene-plagioclase globules contains 2.9–9.6 wt% Al2O3 and 1.0–3.8 wt% TiO2 (Table 3). Plagioclase in pyroxene-plagioclase globules is also calcic with An value varying from 93.6 to 95.8.

Plagioclase fragments usually have a smooth outline. No chemical zoning was detected and the variation of An value of different grains are from An95.8 to An98.4. Large fragments of olivine contain mosaics of 5–20 μm euhedral subgrains at margins (Fig. 3a). The rims of olivine subgrains (Mg# = 72.7–79.5) are usually more Mg-rich than the cores of olivine subgrains and relict olivine (Mg# = 65.6–67.5) (Table 1). Plagioclase grains surrounding olivine fragments are cloudy with submicron olivine grains (Fig. 3a). Fragments of ulvöspinel-chromite-spinel solid solution occur mainly as fine-grained aggregates (30–120 μm) and in a few cases relict core could be observed (Fig. 3b). This texture could be interpreted as a result of shock-induced recrystallization (Walton and Herd 2007; Walton and Shaw 2009). The relict cores (Usp48.2–53.4Chr32.9–38.0Spl13.1–14.6) contain more ulvöspinel and less spinel components than the fine subgrains (Usp28.3–43.2Chr32.7–43.3Spl24.1–31.6) (Table 4).

image

Figure 3.  Backscattered electron images of fragments of olivine and ulvöspinel-chromite-spinel solid solution in Dhofar 458. a) A large olivine fragment has been transformed into polycrystalline olivine subgrains along its rim. The rims of olivine subgrains are more magnesian than the relict cores. b) An irregularly shaped, large grain of ulvöspinel-chromite-spinel solid solution (Usp-Chr-Spl SS) contains a relict core and a lot of subgrains at the margin.

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Table 4.   Representative compositions of ulvöspinel-chromite-spinel solid solution in Dhofar 458.
 Relict coreSubgrains
Major elements, wt% (electron microprobe)
SiO21.600.230.030.060.100.840.411.04
TiO221.121.120.121.513.217.715.712.0
Al2O36.867.557.297.0313.912.613.0617.1
Cr2O324.625.730.128.431.725.426.732.3
MgO6.977.508.778.487.728.038.859.23
FeO37.837.733.834.532.734.733.828.6
MnO0.310.290.370.340.300.310.300.26
CaO0.700.070.100.070.110.680.690.35
Total99.9100.1100.6100.499.7100.399.5100.9
Chromite32.933.338.035.640.932.735.240.0
Spinel13.714.613.713.126.824.125.631.6
Ulvöspinel53.452.048.251.232.343.239.228.3

Occurrence and Dating Results of Polycrystalline Zircon

A polycrstalline zircon aggregate, about 60 μm in size, was observed in recrystallized plagioclase (Fig. 4a). The polycrystalline zircon is porous (Figs. 4a and 4b), which could be a result of incipient melting and degassing (Bohor et al. 1993). It is noteworthy that a few ovoid baddeleyite grains (approximately 0.1 μm to approximately 1 μm in size) occur at the margin of the polycrystalline zircon aggregate (Fig. 4c). The micro-Raman spectra of the zircon and baddeleyite (Fig. 5) are similar to those of unshocked zircons and baddeleyite, respectively (e.g., Gucsik et al. 2004; Wittmann et al. 2006). No Raman peaks corresponding to reidite (e.g., 839 cm−1 and 878 cm−1, Wittmann et al. 2006), cubic high temperature polymorph of ZrO2 (Raman band at 616 cm−1, López et al. 2001), or high-pressure orthohombic II polymorph of ZrO2 (e.g., 661 cm−1, López et al. 2001; Wittmann et al. 2006) were observed in this study.

image

Figure 4.  Backscattered electron image of polycrystalline zircon. a) An irregular zircon grain entrained in plagioclase shows a polycrystalline texture. b) The zircon subgrains are porous, indicating a melting and degassing origin. c) A few fine-grained (less than 1 μm) baddeleyite grains occur at the margin of the zircon.

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image

Figure 5.  Raman spectra of zircon (zrn) and baddeleyite in Dhofar 458.

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Nine analyses were performed on the polycrystalline zircon and gave 207Pb/206Pb ages ranging from 3417 to 3483 Ma (average at 3434 ± 15 Ma). Although all analyses are discordant, these data plot on a well-defined discordia yielding an upper intercept of 3425 ± 29 Ma (Fig. 6) and a lower intercept almost through the coordinate origin. We prefer the 207Pb/206Pb age (3434 ± 15 Ma) as the age of the polycrystalline zircon because this age is not affected by the possible fractionation of U and Pb during the measurements. The nine data points show a restricted range of Th (232–406 ppm) and U (249–414 ppm) concentrations, and Th/U varying from 0.67 to 1.06 (Table 5).

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Figure 6.  Concordia plot of the ion microprobe U-Pb age data from the polycrystalline zircon in Dhofar 458 showing the time of the impact event and recent lead loss. All errors are at 2σ level.

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Table 5.   Ion probe U-Pb results for the polycrystalline zircon in Dhofar 458.
SpotU (ppm)Th (ppm)Th/UPb (ppm)204Pb/206Pb207Pb/235U±1σ (%)206Pb/238U±1σ (%)207Pb/206Pb±1σ (%)206Pb/238U±1σ207Pb/206Pb±1σ
DHO458@12852540.891400.00020113.12.20.3251.8330.29361.1650181529342619
DHO458@24012840.71325021.52.00.5311.8360.29400.8522274541343913
DHO458@34143260.79361023.34.00.5753.5560.29371.9267292684343730
DHO458@43952660.672740.00008318.32.20.4571.8790.29081.0379242638341716
DHO458@54083420.84250016.22.20.4011.7910.29261.2802217433343120
DHO458@63854061.06215014.82.00.3701.7640.28960.8393202931341513
DHO458@72922700.92144013.22.60.3232.4540.29540.9473180439344615
DHO458@82492320.93131014.13.00.3392.6230.30251.4098188043348322
DHO458@92011510.75101013.32.70.3261.9460.29621.8822182031345129

Discussion

  1. Top of page
  2. Abstract
  3. Introduction
  4. Analytical Methods
  5. Results
  6. Discussion
  7. Conclusions
  8. References

Warren et al. (2005) reported major and trace element compositions of Dhofar 026 and Dhofar 457–468 including Dhofar 458. Their results indicate that Dhofar 458 is paired with Dhofar 026. In this study, Dhofar 458 consists mainly of olivine-plagioclase intergrowths, pyroxene-intergrowths, and plagioclase fragments. Pyroxene-plagioclase globules and vesicles are also very common. These petrographic and mineralogical features are very similar to those of Dhofar 026. Based on these petrologic and mineralogical similarities (Cohen et al. 2004; Warren et al. 2005, this study), we consider that Dhofar 458 is paired with Dhofar 026 although the cosmic ray exposure age of Dhofar 458 has not been determined so far and could not be compared with that of Dhofar 026 (Nishiizumi and Caffee 2001).

Shock Metamorphic Origin of the Polycrystalline Zircon

Corfu et al. (2003) reviewed textures of zircon formed in different geological environments. Distinctly different from most terrestrial igneous and metamorphic zircons, zircon that experienced shock events usually contains multiple sets of planar deformation features (PDF). With increasing shock pressure, reidite, the impact-induced high-pressure polymorph of zircon could be observed (Glass et al. 2002; Wittmann et al. 2006); further, zircon can develop polycrystalline (granular) textures (Bohor et al. 1993; Corfu et al. 2003; Wittmann et al. 2006). Under the most extreme conditions that zircon starts to melt (Corfu et al. 2003), zircon would develop degassing and melting textures (Bohor et al. 1993). Zircon would decompose to ZrO2 and silica (El Goresy 1965) at temperatures around 1775–1900 °C (El Goresy 1968). The decomposition of zircon has been regarded as a diagnostic feature of impact melt (Chao 1968; El Goresy 1968; Marvin and Kring 1992; Koeberl and Reimold 2003; Wittmann et al. 2006, 2009).

In Dhofar 458, zircon is a polycrystalline aggregate and porous. Furthermore, baddeleyite grains also occur at the margin of the aggregate. It is very difficult to interpret these features as formed by igneous crystallization and thermal metamorphism. Instead, they can be interpreted as a product of shock-induced melting, as suggested for similar features found in terrestrial impactites (e.g., Bohor et al. 1993; Wittmann et al. 2006, 2009). Although the polycrystalline texture of zircon could be diaplectic (formed by shock without melting, Bohor et al. 1993), the porous texture of zircon cannot be diaplectic and must involve melting and degassing processes (Bohor et al. 1993; Corfu et al. 2003). Occurrence of small baddeleyite grains at the margin of the aggregate also supports the melting origin. In terrestrial impactites, baddeleyite grains were also observed occurring at margins of zircon grains (Wittmann et al. 2006) and were interpreted as decomposition products of zircon at extremely high postshock temperature. This interpretation could also be suitable to the case in this study. Similar to other decomposed zircons reported in terrestrial impactites, no silica was observed in Dhofar 458. The lack of silica could be due to vaporation of silica or that silica was lost from the decomposed zircon domains (Wittmann et al. 2006). The zircon aggregate observed in this study could be the first report of shock-induced zircon with polycrystalline and degassing texture and decomposition in extraterrestrial materials.

Based on comparison with results of shock recovery experiments, Wittmann et al. (2006) constrained possible ranges of shock pressures and postshock temperatures for various shock-metamorphic features of zircon from terrestrial impact craters. The granular (polycrystalline) textures of zircon could occur at shock pressure higher than 50 GPa and postshock temperature higher than approximately 1200 °C. Decomposition of zircon to ZrO2 with granular textures in nonporous quartzo-feldspathic rocks starts at approximately 60 GPa shock pressure and approximately 1700 °C (Wittmann et al. 2006). Applying these suggestions to the case in this study, because the whole rock of Dhofar 458 is feldspathic (Warren et al. 2005), the presence of decomposition of zircon to ZrO2 with granular textures might indicate shock pressure is higher than approximately 60 GPa and postshock temperature is above approximately 1700 °C. Based on the classification of shocked quartzo-feldspathic rocks (Stöffler and Grieve 2007), the inferred shock pressure and postshock temperature are high up to shock stage IV and enough to melt whole rock. At the same time, the inferred postshock temperature is much above the stable temperature of reidite, the high-pressure polymorph of ZrSiO4 which is stable at temperature below approximately 1200 °C (Kusaba et al. 1985). This could be the reason why reidite is absent. In addition, as pointed out by Wittmann et al. (2006), because baddeleyite is a monoclinic low-temperature and low-pressure polymorph, it cannot be the primarily resultant polymorph of temperature and/or pressure-induced decomposition of ZrSiO4. Besides baddeleyite, ZrO2 has a high-pressure polymorph (orthorhombic II polymorph) that is stable at ambient conditions (Arashi et al. 1990; Wittmann et al. 2006). However, although the decomposition of zircon in Dhofar 458 was attributed to postshock temperature, no high-temperature and high-pressure polymorphs of ZrO2 were detected by Raman spectroscopy. This could be due to relatively slow quenching that caused transformation to baddeleyite but no or not complete reversion to zircon or due to low spectral resolution (Wittmann et al. 2006, 2009).

Shock-Induced Melting versus Thermally Induced Melting of Dhofar 458

To understand the formation process of Dhofar 458, two key questions that must be answered are whether the whole rock was melted and what process causes the melting of the rock. Granulitic breccias are interpreted to be formed by thermally induced subsolidus recrystallization of breccia precursors (Stöffler et al. 1980; Stöffler and Grieve 2007). However, melting in impact melt rocks is attributed to high postshock temperatures (James et al. 2007 and references therein). For Dhofar 026, Cohen et al. (2004) and James et al. (2007) concluded that it should be a strongly shocked granulitic breccia or a strongly shocked fragmental breccia consisting almost entirely of granulitic-breccia clasts. Their conclusion was based on their interpretation that olivine-plagioclase intergrowths which are abundant in volume (40–45%) were not melted during the impact. They suggested that the impact event only caused melting of plagioclase, whereas melting of pyroxene and subsequent recrystallization of plagioclase and pyroxene were due to thermal diffusion from an external, hotter material (Cohen et al. 2004). However, in this study, the high shock pressure (greater than approximately 60 GPa) and postshock temperature (greater than approximately 1700 °C) deduced from the polycrystalline and degassing texture and decomposition of zircon are enough to melt the whole rock of the feldspathic Dhofar 458 (Stöffler and Grieve 2007). It is impossible that thermal metamorphism by conduction causes such a special texture. The decomposition of zircon to ZrO2 has also been considered as a diagnostic feature of impact melt (Chao 1968; El Goresy 1968; Marvin and Kring 1992; Koeberl and Reimold 2003; Wittmann et al. 2006).

Our petrographic observations (e.g., Fig. 2a) show that olivine grains at the margins of olivine-plagioclase intergrowths are interstitial to the plagioclase, indicating melting and recrystallization of both plagioclase and olivine. Similar textures have been reported in the shocked lherzolitic Martian meteorite Grove Mountains 99027 (Wang and Chen 2006) and were attributed to shock-induced melting and recrystallization of plagioclase. Similar textures along margins of pyroxene-plagioclase globules and in pyroxene-plagioclase intergrowths (Figs. 2b and 2c) could also due to melting and recrystallization. Why Cohen et al. (2004) and James et al. (2007) did not observe similar textures in olivine-plagioclase intergrowths in Dhofar 026 could be due to the fact that they used lower magnifications than we do in this study when they took BSE images. In addition, large polycrystalline rims of olivine and ulvöspinel-chromite-spinel solid solution in Dhofar 458 could be also closely related to impact melting. Similar texture of chromite and spinel grains has been described in shock melt pockets in Martian meteorites (e.g., Walton and Herd 2007; Walton and Shaw 2009). The compositional difference of the polycrystalline rims of olivine and ulvöspinel-chromite-spinel solid solution compared to their respective relict cores could be a result of chemical re-equilibrium with melt (Cohen et al. 2004; Walton and Shaw 2009). To summarize, based on the above interpretations, melting of most components in Dhofar 458 were shock induced and Dhofar 458 should be classified as a clast-rich impact melt rock according to the classification by Stöffler and Grieve (2007).

Other issues which should be addressed about the formation of special texture of Dhofar 458 are the origins of globules and vesicles. Cohen et al. (2004) and James et al. (2007) proposed that pyroxene-plagioclase globules were formed in situ after the impact event; however, Warren et al. (2005) suggested that they could be relict clasts before the impact event. Based on our observations that almost all margins of pyroxene-plagioclase globules exhibit a penetrating texture and a few compound pyroxene-plagioclase globules also occur, these pyroxene-plalgioclase globules could be relict clasts as Warren et al. (2005) suggested. Vesicles are very common in Dhofar 458 and 026. They could be analogs of vesicles in polycrystalline zircon and are products of impact melting and degassing of plagioclase, pyroxene, and olivine.

Implication of the U-Pb Age of the Polycrystalline Zircon

A few authors have studied the U-Pb isotopic system of shocked zircon. The U-Pb system of most shocked zircon has been highly reset (e.g., Krogh et al. 1993; Kamo and Krogh 1995; Kamo et al. 1996; Gibson et al. 1997; Deloule et al. 2001; Mänttäri and Koivisto 2001; Gnos et al. 2004; Pidgeon et al. 2007; Liu et al. 2010a), although shocked zircon without any impact effect on the U-Pb system was also reported in literature (e.g., Åberg and Bollmark 1985). Recently, Kalleson et al. (2009) studied the effects of shock events on the U-Pb system in zircon from the Gardnos impact structure, Norway. They found that unshocked zircon does not have any appreciable Pb loss effect and fractured zircon usually has substantial Pb loss. However, granular zircon has the highest degree of resetting of the U-Pb system; a few were totally reset.

As discussed above, in the case from Dhofar 458, the polycrystalline and degassing texture indicates that zircon must have been melted and recrystallized. The shock pressure and the high postshock temperature are also very high (greater than approximately 60 GPa and greater than approximately 1700 °C, respectively). During this process, the U-Pb system must have been extensively reset, if not totally reset. The U-Pb isotopic data of the polycrystalline zircon plot on a well-defined discordia between 3425 Ma and 0 Ma (Fig. 6). This result supports that the U-Pb system of the zircon in Dhofar 458 with polycrystalline and degassing texture was totally reset. The zircon totally lost its radiogenic Pb. Although the duration of the impact event was short (Kalleson et al. 2009), the rapid Pb loss could be related to the high postshock temperature, recrystallization into polycrystalline aggregate (Deutsch and Schärer 1994; Kalleson et al. 2009). At the same time, we note that the data did not plot on the concordia; this result may suggest that recent Pb-loss took place in this polycrystalline zircon.

Dhofar 026 was previously dated by the Ar-Ar method and feldspathic clasts therein show a large variation in age from 0.569 ± 0.011 Ga (Cohen et al. 2002) to 6.81 ± 0.32 Ga (Fernandes et al. 2004). Fernandes et al. (2004) obtained age ranges for two feldspathic clasts (called clasts A and E by Fernandes et al. 2004), 2.16 ± 0.21 to 5.07 ± 1.71 Ga and 1.56 ± 0.16 to 3.94 ± 0.60 Ga, respectively. They suggested that 2.16 Ga and 1.56 Ga could be the maximum ages for the impact event that formed these two clasts, respectively. These Ar-Ar ages of feldspathic clasts are much younger than the zircon U-Pb age reported here. It is possible that these Ar-Ar ages reflect disturbance of K-Ar systems to various degrees in different clasts after the heavy impact that formed the polycrystalline zircon in Dhofar 458.

Compared to zircons from lunar rocks and other lunar meteorites (e.g., Meyer et al. 1996; Gnos et al. 2004; Pidgeon et al. 2007; Nemchin et al. 2008, 2009a, 2009b; Grange et al. 2009; Liu et al. 2010a), the polycrystalline zircon in Dhofar 458 is the youngest so far. This age is consistent with one of the age peaks (approximately 3.4 Ga) of impact melt clasts determined by Cohen et al. (2000) and the age of Luna 16 basalt surface (Stöffler et al. 2006). Considering the difference of the lithologies between the Luna 16 basalt surface and Dhofar 458 and other feldspathic lunar meteorites reported in Cohen et al. (2000), this age consistency may indicate that approximately 3.4 Ga recorded one of wide and intense impact events on the lunar surface.

Conclusions

  1. Top of page
  2. Abstract
  3. Introduction
  4. Analytical Methods
  5. Results
  6. Discussion
  7. Conclusions
  8. References
  • 1
    The petrographic and mineralogical results suggest that Dhofar 458 is a stone paired with Dhofar 026, consistent with the conclusion in previous investigations.
  • 2
    The zircon in Dhofar 458 has a polycrystalline and degassing texture. Fine baddeleyite grains are present at the margin of the polycrystalline zircon. We interpret that the polycrystalline and degassing texture of the zircon is a result of shock-induced melting and recrystallization. At the same time, the presence of baddeleyite indicates the high-temperature decomposition of zircon. Based on comparison with results from shock-recovery experiments and textures of zircon in terrestrial impactites, the polycrystalline and degassing texture and decomposition of zircon indicate that shock pressure could be greater than approximately 60 GPa and the postshock temperature could be greater than approximately 1700 °C. The polycrystalline and degassing texture and decomposition of zircon also indicate that Dhofar 458 is a clast-rich impact melt rock.
  • 3
    The SIMS U-Pb isotopic results show a well-defined discordia between 3425 Ma and 0 Ma. The 3425 Ma age recorded the time of the impact event that caused melting, decomposition, and recrystallization of zircon.

Acknowledgments— The senior author thanks Mr. Tiangang Wang for his assistance during Raman spectroscopic analyses. The authors would like to express their gratitude to Dr. Axel Wittmann and an anonymous reviewer for their constructive and helpful comments and to the associate editor Dr. Christian Koeberl for editorial efforts. This work was supported by the Natural Science Foundation of China (grant nos. 41703052, 40703015, 40773046) and the State Key Laboratory of Lithospheric Evolution at the Institute of Geology and Geophysics, Chinese Academy of Sciences.

Editorial Handling— Dr. Christian Koeberl

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  5. Results
  6. Discussion
  7. Conclusions
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