2023 Volume 118 Issue ANTARCTICA Article ID: 221124
We determined the metamorphic age and pressure-temperature conditions recorded in a sillimanite-garnet-bearing pelitic gneiss from Niban-nishi Rock, which is part of Niban Rock, on the Prince Olav Coast, eastern Dronning Maud Land, East Antarctica. Niban-nishi Rock is recognized as a component of the Lützow-Holm Complex (LHC), which is characterized by metamorphic age of 600-520 Ma and metamorphic grade of amphibolite to granulite facies. Electron microprobe U-Th-Pb monazite dating of the examined gneiss revealed that, unlike the typical exposures of the LHC, Niban-nishi Rock experienced Tonian metamorphism at 940.1 ± 9.8 Ma (2σ level), and is thus more similar to the neighboring exposures of Cape Hinode and Akebono Rock. Small numbers of younger monazite ages of 827-531 Ma were also detected, and some of which might relate to the metamorphism of the LHC. Phase equilibrium modeling and geothermobarometry indicate metamorphic conditions of 690-730 °C/0.38-0.68 GPa and 620-670 °C/0.42-0.60 GPa for peak and retrograde stages, respectively. The obtained peak temperature is lower than that of typical exposures in transitional- and granulite-facies zones of the LHC, such as Akarui Point and Skallen. Metamorphic features of Niban-nishi Rock, such as the upper amphibolite-facies condition and occurrences of garnet with retrograde zoning and sillimanite in the matrix, differ from those of Cape Hinode and Akebono Rock, with the former belonging to granulite facies and the latter showing kyanite in the matrix and garnet with growth zoning. Investigating Niban-nishi Rock is key in revealing the tectonic relation between the LHC and Cape Hinode (the Hinode Block). Further field surveys are required to reveal the metamorphic variations and the relations among exposures on the Prince Olav Coast.
An area that extends over approximately 400 km from the Prince Olav Coast to Prince Harald Coast in eastern Dronning Maud Land, East Antarctica was previously considered to be occupied by the Lützow-Holm Complex (LHC), which is characterized by a metamorphic age of 600-520 Ma (Dunkley et al., 2020 and reference therein) with the metamorphic grade of amphibolite to granulite facies (e.g., Hiroi et al., 1991) (Fig. 1a). Hiroi et al. (1991), who found a continuous increase in the metamorphic grade from the Prince Olav Coast to Prince Harald Coast, divided the area into three metamorphic zones referred to as the amphibolite-facies zone, transitional zone, and granulite-facies zone from northeast to southwest (Fig. 1a). Suzuki and Kawakami (2019) recently revealed that the metamorphic grade of the transitional zone is similar to that of the granulite-facies zone.
In contrast, the distinct discrepancies in metamorphic age and lithology observed in other exposures of the LHC has been confirmed in Cape Hinode, an exposure located in the amphibolite-facies zone (Fig. 1a), which shows Tonian metamorphism of 1017-960 Ma and metatonalite-dominant lithology (Yanai and Ishikawa, 1978; Hiroi et al., 2006, 2008; Dunkley et al., 2014). In addition, the metamorphic grade of Cape Hinode is granulite facies evidenced by the presence of orthopyroxene in the metamorphosed mafic enclaves within the metatonalite, and it is apparently higher grade than other exposures in the amphibolite-facies zone (Hiroi et al., 2006, 2008). Thus, Cape Hinode is presently denoted as Hinode Block and is now regarded independent of the LHC (Dunkley et al., 2020). Recently, Baba et al. (2022) revealed that Akebono Rock, an exposure located approximately 12 km northeast of Cape Hinode in the amphibolite-facies zone (Fig. 1a), records Tonian metamorphism of 977-917 Ma. Although both Cape Hinode and Akebono Rock also record Tonian metamorphism (Hiroi et al., 2006, 2008; Dunkley et al., 2014; Baba et al., 2022), the metamorphic age of Sinnan Rocks, which is located in the amphibolite-facies zone and at northeastern end of the LHC, reported at 533 Ma (Shiraishi et al., 1994), and is thus consistent with the metamorphic age of the LHC. Due to the complexity of the metamorphic conditions and ages in the amphibolite-facies zone, exposures in the zone are likely to be a key in revealing the tectonic evolution of the eastern Dronning Maud Land in East Antarctica from the Mesoproterozoic to the Neoproterozoic.
Despite its importance, the metamorphic pressure (P)-temperature (T) conditions and the age of the exposures in the amphibolite-facies zone are poorly understood as compared with those in transitional and granulite-facies zones. Niban-nishi Rock, the studied exposure in this study, is located approximately 15 km southwest of Cape Hinode (Fig. 1a), has also been considered a component of the LHC. However, the metamorphic age and P-T conditions recorded by the major lithology in Niban-nishi Rock have not as yet been determined.
In this study, we deal with a sillimanite-garnet-bearing pelitic gneiss, one of the most dominant lithologies in Niban Rock, which was collected from Niban-nishi Rock on the Prince Olav Coast. The timing of the metamorphism was determined by electron microprobe (EMP) U-Th-Pb monazite dating and the P-T conditions of the peak and retrograde stages were estimated with geothermobarometry and phase equilibrium modeling. The relation among Niban-nishi Rock, the LHC, and the neighboring exposures of Cape Hinode and Akebono Rock is also discussed. Mineral abbreviations follow those in Whitney and Evans (2010).
The LHC is a Neoproterozoic-Cambrian high-grade metamorphic complex that covers the area from the Prince Olav Coast to Prince Harald Coast in eastern Dronning Maud Land, East Antarctica (Fig. 1a). The main lithology of the LHC is well-layered pelitic to psammatic and intermediate gneisses that include minor amounts of metamorphosed mafic to ultramafic gneisses occurring as lenticular masses or thin layers within the metasedimentary gneisses (e.g., Hiroi et al., 1991). Typical features among exposures of the LHC include (1) metamorphism of ∼ 600-520 Ma (Dunkley et al., 2020 and references therein), (2) a clockwise P-T path and isothermal decompression (Fig. 2) characterized by the presence of symplectite around garnet as a breakdown product (e.g., Hiroi et al., 1986a; Iwamura et al., 2013; Takahashi and Tsunogae, 2017; Takamura et al., 2020), and (3) a change in the stable phase aluminosilicate phase from kyanite to sillimanite (e.g., Motoyoshi et al., 1985; Suzuki and Kawakami, 2019). Chemical zoning of garnet in the LHC is commonly affected by homogenization under high-T conditions and re-equilibration during the retrograde stage (e.g., Ikeda, 2004; Hiroi et al., 2019). The LHC is considered to show progressive metamorphism from northeast to southwest and has been divided into three zones of the amphibolite-facies, transitional, and granulite-facies zones (Hiroi et al., 1991) (Fig. 1a). However, recent studies have revealed that Akarui Point in the transitional zone has attained high-temperature conditions similar to those in the granulite-facies zone (Kawakami et al., 2008; Iwamura et al., 2013; Nakamura et al., 2014; Suzuki and Kawakami, 2019) (Fig. 2). The ultra-high temperature metamorphism was confirmed in Rundvågshetta and Skallevikshalsen, which are located in the granulite-facies zone (e.g., Motoyoshi et al., 1985; Kawakami and Motoyoshi, 2004; Kawasaki et al., 2011; Hiroi et al., 2019) (Figs. 1a and 2).
In studying the LHC, exposures showing different features from the LHC have been found in the area where previously considered to belong to the LHC. Shiraishi et al. (1994) reported 1017 ± 13 Ma as the igneous age of the metatrondjemite in Cape Hinode and mentioned that there is no evidence for zircon crystallization during the 550-500 Ma event. Dunkley et al. (2014) conducted U-Pb zircon dating of samples from Cape Hinode and reported metamorphic ages of 970 Ma for the garnet metapelite and 960 Ma for the metatrondjemite, both of which are much older than the metamorphic age of 600-520 Ma obtained for the LHC. Hiroi et al. (2006) described orthopyroxene-bearing mafic enclaves with a metamorphic grade corresponding to the granulite facies in the metatonalite, which is higher metamorphic grade than that of exposures in the amphibolite-facies zone. Thus, Hiroi et al. (2006, 2008) considered Cape Hinode to be an exotic block. Hiroi et al. (2008) applied a hornblende-plagioclase thermometer and an Al-in-hornblende geobarometer to metatonalite from Cape Hinode and determined 640-800 °C and 0.4-0.7 GPa as the conditions during the last stage of tonalite magmatic crystallization and its overprint with subsequent metamorphism (Fig. 2). In Cape Hinode, kyanite that reported to replace sillimanite in the matrix of the pelitic gneiss is considered to have formed during a deformation event following peak metamorphism (Hiroi et al., 2008) (Fig. 2). In Akebono Rock that is located approximately 12 km northeast of Cape Hinode (Fig. 1a), Baba et al. (2022) reported Tonian metamorphism of 937 ± 6 Ma based on U-Pb zircon dating and 977-917 Ma based on EMP U-Th-Pb monazite dating. Baba et al. (2022) found that garnet preserves prograde zoning with an increase in almandine and pyrope and a decrease in grossular and spessartine contents from core to rim and kyanite is stable phase in the matrix of pelitic gneisses from Akebono Rock. Peak metamorphic temperature and timing of 642 °C at 937 Ma were also obtained for Akebono Rock using a Ti-in-zircon geothermometer and U-Pb zircon dating (Baba et al., 2022). The P-T path of the regional metamorphism in Akebono Rock was illustrated by correlating the results of the garnet-biotite geothermometer and that of the garnet-biotite-plagioclase-quartz geobarometer to various combinations of compositions recorded in different parts of the grain (core or rim) and the mode of occurrence (inclusion or matrix) for garnet, biotite, and plagioclase (Baba et al., 2022) (Fig. 2). Investigation on the shear zone in Akebono Rock (Baba et al., 2020) indicated that the mineral chemical compositions of the deformed metamorphic rock are re-equilibrated during deformation and the P-T conditions during the final deformation event at the shear zone are 610-660 °C and 0.4-0.5 GPa (Fig. 2). The shear zone contains amphibolite with large sigmoidal garnets, the core of which is considered to have formed at a deeper crustal depth (∼ 700 °C and ∼ 0.8 GPa) prior to deformation and transferred to shallower levels during formation of the shear zone (Baba et al., 2020) (Fig. 2). A metamorphic age of 533 ± 6 Ma has been reported for Sinnan Rocks at the northeastern end of the LHC (Fig. 1a), based on U-Pb zircon dating (Shiraishi et al., 1994).
Therefore, reconstruction of the framework of the LHC remains necessary, and several studies have proposed new frameworks based on protolith age, metamorphic age, lithology, and metamorphic conditions (e.g., Takahashi et al., 2018; Dunkley et al., 2020). Dunkley et al. (2020) subdivided the LHC into six suites and one block based on protolith age and lithology: the Akarui Suite, the East Ongul Suite, the Langhovde Suite, the Skallevikshalsen Suite, the Rundvågshetta Suite, the Inhovde Suite, and the Hinode block (Fig. 1a). The Hinode block is composed of Cape Hinode and is located between Niban Rock and Akebono Rock in the Akarui Suite (Fig. 1a); thus, Dunkley et al. (2020) excluded Cape Hinode from the LHC.
Niban Rock is a 2.5 km × 3.5 km exposure located in the amphibolite-facies zone defined by Hiroi et al. (1991) and is composed of Niban-higashi Rock and Niban-nishi Rock (Fig. 1b). Located approximately 15 km southwest of Cape Hinode, Niban Rock belongs geographically to the Akarui Suite (Dunkley et al., 2020) and is underlain mainly by sillimanite-garnet-biotite gneiss, biotite gneiss, and biotite-hornblende gneiss along with minor metabasite, calc-silicate gneiss, granite, and aplite (Kizaki et al., 1983) (Fig. 1b). Dunkley et al. (2014) conducted U-Pb zircon dating and U-Th-Pb monazite dating on samples from Niban Rock and reported magmatic and metamorphic ages of 551 ± 11 Ma and 532 ± 7 Ma, respectively, from the metagranitic dyke and the protolith age of 940 ± 6 Ma from the granitic augengneiss. Recently, preliminary U-Pb zircon age of garnet-sillimanite-biotite gneiss from Niban-nishi Rock was reported as 998 ± 9.7 Ma as a metamorphic age and 1940-1760, 1300, and 1160-1040 Ma as detrital age (Kitano et al., 2021).
The chemical compositions of the constituent minerals in the studied gneiss were determined using a JEOL JXA-8530F electron probe microanalyzer equipped with wavelength dispersive X-ray spectrometers (WDS) at Kyushu University, Fukuoka, Japan. Both of natural and synthetic materials were used for standards. An accelerating voltage of 15 kV, beam current of 10 nA, and probe diameter of 2-6 µm were used for analysis. The ZAF method was used for data correction. Counting time were 60 and 30 s for F, 30 and 15 s for Cl, and 10 and 5 s for other elements for the peak and backgrounds, respectively.
The method used for electron microprobe analysis of monazite followed Hokada and Motoyoshi (2006). Analysis was performed using a JEOL JXA-8200 electron probe microanalyzer equipped with WDS at the National Institute of Polar Research, Tokyo, Japan. Concentrations of U, Th, Pb and other elements that monazite possibly contains (P, Si, Ca, Y, La, Ce, Pr, Nd, Sm, Gd, Dy, Er, and Yb) were determined with a polished thin section that included monazite grains. Elements that generally of extremely low concentrations (S, Zr, Hf, Al, Mg, and Na) in monazite were measured simultaneously during analysis to check for contamination with other minerals. An accelerating voltage of 15 kV, beam current of 200 nA, and probe diameter of 10 µm were used. We confirmed that counting rate of characteristic X-ray did not decrease during the analysis and could minimize analytical errors when the analytical condition was applied. Monazite from Namaqualand in South Africa (1040-1033 Ma; Knoper et al., 2001) was also analyzed at the start, middle, and end of the series of analyses as an internal age-reference.
Major and minor element compositions of whole rock samples were determined using an X-ray fluorescence (XRF) spectrometer (Rigaku-ZSX Primus IV) at Kyushu University, Fukuoka, Japan. The rock sample utilized for XRF analysis was powdered with cylinder mortar, agate mortar, and agate pestle. Loss on ignition was measured after the sample was heated at 1000 °C for 6 h in an electric furnace. Powdered rock sample and anhydrous dilithium tetraborate flux was fused at a weight ratio of 1:5 to prepare a glass bead. Glass discs were utilized for calibration using a technique described in Nakada (1985) and Nakada et al. (1985).
The studied pelitic gneiss (sample no. TM11020804A) was collected from Niban-nishi Rock during the 52nd Japanese Antarctic Research Expedition (JARE 52) (Fig. 1b). The pelitic gneiss is intruded by a pegmatite dyke (Fig. 3a). On an outcrop scale, the pelitic gneiss are largely divided into melanosomes and lenticular leucosomes, with the long axis of the leucosome reaching to several tens of centimeters in size (Fig. 3b). The gneiss examined in this study corresponds to the melanosome and contains leucocratic patches (up to 3 cm long) (Fig. 4). We refer to melanosome excluding the leucocratic patch as ‘melanocratic portion’. The melanocratic portion is composed mainly of quartz, plagioclase, biotite, sillimanite, K-feldspar, muscovite (primary), and garnet with minor amounts of zircon, monazite, apatite, and ilmenite (Figs. 4, 5a-5c, and 6). An areal proportion of garnet in the melanocratic portion is ∼ 0.6% measured under two thin sections and a slab photograph (Fig. 4). The melanocratic portion shows foliation that is defined by the shape-preferred orientations of sillimanite, biotite, muscovite, and quartz (Figs. 4 and 5a-5c). The leucocratic patch is composed mainly of coarse grains of plagioclase and quartz (up to 2.5 mm in diameter) with minor amounts of biotite (up to 0.5 mm long) and K-feldspar (up to 0.5 mm in diameter) (Figs. 4 and 5d). Rare monazite occurs in the leucocratic patch as well (Fig. 5d). Foliation of the leucocratic patch defined by shape-preferred orientation of biotite and the long axis of the leucosome and the leucocratic patch are parallel to foliation of the melanocratic portion (Figs. 3b, 4, and 5). The whole-rock chemistry of the melanocratic portion is shown in Table 1. Petrography and mineral chemistry described below is about the melanocratic portion. Iron in minerals reported below was treated as ferrous and denoted as Fetotal. The representative analyses of garnet, biotite, muscovite, K-feldspar, and plagioclase are shown in Table 2.
| SiO2 (wt%) | 69.74 |
| TiO2 | 0.82 |
| Al2O3 | 14.84 |
| Fe2O3 | 6.85 |
| MnO | 0.10 |
| MgO | 2.14 |
| CaO | 0.78 |
| Na2O | 1.68 |
| K2O | 3.23 |
| P2O5 | 0.05 |
| Total | 100.23 |
| LOI (wt%) | 0.69 |
| Sc (ppm) | 19.4 |
| V | 113.7 |
| Cr | 208.3 |
| Ni | 82.5 |
| Cu | 2.8 |
| Zn | 85.1 |
| Rb | 130.9 |
| Sr | 75.1 |
| Y | 26.5 |
| Zr | 176.5 |
| Nb | 12.9 |
| Ba | 355.7 |
| Pb | 11.1 |
LOI, loss on ignition.
| Mineral | Grt | Grt | Grt | Grt | Grt | Grt |
| Occurrence | Mtx | Mtx | Mtx | Mtx | Mtx | Inc in Pl |
| Part | Core | Rim | Rim | Rim | Rim | Core |
| In contact with | - | Bt | Felsic | Felsic | Felsic | - |
| Analysis no. | p063-L16 | p051 | p063-L01 | p198*1 | p060*1 | p214 |
| SiO2 | 37.56 | 36.57 | 36.91 | 37.06 | 36.94 | 37.31 |
| TiO2 | 0.01 | 0.00 | 0.05 | 0.01 | 0.03 | 0.01 |
| Al2O3 | 20.58 | 20.31 | 20.55 | 20.38 | 20.64 | 20.99 |
| Cr2O3 | 0.00 | 0.08 | 0.05 | 0.03 | 0.00 | 0.02 |
| FeO | 32.54 | 32.58 | 32.16 | 32.23 | 32.69 | 32.58 |
| MnO | 5.15 | 7.15 | 7.26 | 5.60 | 6.66 | 5.65 |
| MgO | 2.42 | 2.01 | 2.21 | 2.71 | 2.25 | 2.97 |
| CaO | 1.11 | 1.18 | 1.20 | 1.16 | 1.10 | 1.10 |
| ZnO | 0.00 | 0.03 | 0.11 | 0.00 | 0.10 | 0.00 |
| BaO | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Na2O | 0.02 | 0.05 | 0.04 | 0.03 | 0.03 | 0.04 |
| K2O | 0.00 | 0.00 | 0.00 | 0.00 | 0.02 | 0.00 |
| F | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Cl | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| -O=F | - | - | - | - | - | - |
| -O=Cl | - | - | - | - | - | - |
| Total | 99.39 | 99.94 | 100.54 | 99.19 | 100.47 | 100.67 |
| Number of O | 12 | 12 | 12 | 12 | 12 | 12 |
| Si | 3.05 | 2.99 | 2.99 | 3.02 | 2.99 | 3.00 |
| Ti | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 |
| Al | 1.97 | 1.96 | 1.96 | 1.96 | 1.97 | 1.99 |
| Cr | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 |
| Fetotal | 2.21 | 2.23 | 2.18 | 2.20 | 2.22 | 2.19 |
| Mn | 0.35 | 0.50 | 0.50 | 0.39 | 0.46 | 0.38 |
| Mg | 0.29 | 0.24 | 0.27 | 0.33 | 0.27 | 0.36 |
| Ca | 0.10 | 0.10 | 0.10 | 0.10 | 0.10 | 0.09 |
| Zn | 0.00 | 0.00 | 0.01 | 0.00 | 0.01 | 0.00 |
| Ba | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Na | 0.00 | 0.01 | 0.01 | 0.00 | 0.00 | 0.01 |
| K | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 |
| F | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Cl | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Total cation | 7.97 | 8.03 | 8.02 | 8.00 | 8.02 | 8.01 |
| Xalm | 0.75 | 0.73 | 0.71 | 0.73 | 0.73 | 0.72 |
| Xprp | 0.10 | 0.08 | 0.09 | 0.11 | 0.09 | 0.12 |
| Xgrs | 0.03 | 0.03 | 0.03 | 0.03 | 0.03 | 0.03 |
| Xsps | 0.12 | 0.16 | 0.16 | 0.13 | 0.15 | 0.13 |
| XMg | 0.12 | 0.10 | 0.11 | 0.13 | 0.11 | 0.14 |
| Xan | - | - | - | - | - | - |
| Xab | - | - | - | - | - | - |
| Xor | - | - | - | - | - | - |
| Mineral | Bt | Bt | Bt | Bt | Bt | Bt | Bt | Ms | Ms | Ms |
| Occurrence | Mtx | Mtx | Mtx | Mtx | Mtx | Mtx | Inc in Grt | Mtx | Mtx | Inc in Pl |
| Part | Core | Rim | Rim | Core | Core | Rim | Core | Core | Core | Core |
| In contact with | Grt | Ms | Ms | Felsic | Felsic | Felsic | - | Bt | Bt | - |
| Analysis no. | p147 | p046*2 | p077 | p102*1 | p208*1 | p118 | p199 | p068 | p072 | p104 |
| SiO2 | 34.51 | 34.76 | 34.56 | 34.88 | 34.89 | 34.78 | 35.39 | 45.14 | 45.70 | 45.24 |
| TiO2 | 2.20 | 3.76 | 2.88 | 2.94 | 2.64 | 3.02 | 2.00 | 0.79 | 0.69 | 0.75 |
| Al2O3 | 19.60 | 17.42 | 19.32 | 19.26 | 19.74 | 19.09 | 19.91 | 35.18 | 34.91 | 34.59 |
| Cr2O3 | 0.11 | 0.12 | 0.05 | 0.12 | 0.19 | 0.06 | 0.03 | 0.08 | 0.04 | 0.28 |
| FeO | 19.32 | 21.25 | 19.92 | 20.42 | 20.02 | 20.89 | 16.82 | 0.94 | 1.35 | 1.37 |
| MnO | 0.09 | 0.17 | 0.18 | 0.21 | 0.11 | 0.22 | 0.14 | 0.02 | 0.01 | 0.00 |
| MgO | 8.06 | 7.57 | 7.63 | 7.36 | 7.71 | 7.54 | 10.27 | 0.48 | 0.68 | 0.66 |
| CaO | 0.01 | 0.04 | 0.02 | 0.02 | 0.05 | 0.00 | 0.00 | 0.01 | 0.00 | 0.00 |
| ZnO | 0.00 | 0.00 | 0.15 | 0.00 | 0.00 | 0.00 | 0.09 | 0.01 | 0.00 | 0.00 |
| BaO | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Na2O | 0.08 | 0.05 | 0.14 | 0.07 | 0.12 | 0.15 | 0.17 | 0.39 | 0.46 | 0.36 |
| K2O | 9.87 | 9.91 | 9.65 | 9.82 | 10.05 | 9.49 | 9.82 | 10.79 | 10.58 | 11.17 |
| F | 0.22 | 0.22 | 0.18 | 0.16 | 0.16 | 0.19 | 0.17 | 0.04 | 0.06 | 0.06 |
| Cl | 0.07 | 0.06 | 0.05 | 0.07 | 0.07 | 0.07 | 0.06 | 0.00 | 0.01 | 0.00 |
| -O=F | 0.09 | 0.09 | 0.08 | 0.07 | 0.07 | 0.08 | 0.07 | 0.02 | 0.02 | 0.03 |
| -O=Cl | 0.01 | 0.01 | 0.01 | 0.02 | 0.02 | 0.01 | 0.01 | 0.00 | 0.00 | 0.00 |
| Total | 94.03 | 95.21 | 94.62 | 95.24 | 95.65 | 95.40 | 94.79 | 93.85 | 94.46 | 94.45 |
| Number of O | 22 | 22 | 22 | 22 | 22 | 22 | 22 | 22 | 22 | 22 |
| Si | 5.36 | 5.40 | 5.34 | 5.37 | 5.34 | 5.35 | 5.37 | 6.11 | 6.15 | 6.12 |
| Ti | 0.26 | 0.44 | 0.33 | 0.34 | 0.30 | 0.35 | 0.23 | 0.08 | 0.07 | 0.08 |
| Al | 3.59 | 3.19 | 3.52 | 3.49 | 3.56 | 3.46 | 3.56 | 5.61 | 5.53 | 5.51 |
| Cr | 0.01 | 0.01 | 0.01 | 0.01 | 0.02 | 0.01 | 0.00 | 0.01 | 0.00 | 0.03 |
| Fetotal | 2.51 | 2.76 | 2.58 | 2.63 | 2.56 | 2.69 | 2.13 | 0.11 | 0.15 | 0.16 |
| Mn | 0.01 | 0.02 | 0.02 | 0.03 | 0.01 | 0.03 | 0.02 | 0.00 | 0.00 | 0.00 |
| Mg | 1.87 | 1.75 | 1.76 | 1.69 | 1.76 | 1.73 | 2.32 | 0.10 | 0.14 | 0.13 |
| Ca | 0.00 | 0.01 | 0.00 | 0.00 | 0.01 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 |
| Zn | 0.00 | 0.00 | 0.02 | 0.00 | 0.00 | 0.00 | 0.01 | 0.00 | 0.00 | 0.00 |
| Ba | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Na | 0.02 | 0.01 | 0.04 | 0.02 | 0.03 | 0.04 | 0.05 | 0.10 | 0.12 | 0.09 |
| K | 1.96 | 1.96 | 1.90 | 1.93 | 1.96 | 1.86 | 1.90 | 1.86 | 1.82 | 1.93 |
| F | 0.11 | 0.11 | 0.09 | 0.08 | 0.08 | 0.09 | 0.08 | 0.02 | 0.03 | 0.03 |
| Cl | 0.02 | 0.02 | 0.01 | 0.02 | 0.02 | 0.02 | 0.01 | 0.00 | 0.00 | 0.00 |
| Total cation | 15.58 | 15.55 | 15.53 | 15.51 | 15.56 | 15.52 | 15.60 | 13.98 | 13.98 | 14.05 |
| Xalm | - | - | - | - | - | - | - | - | - | - |
| Xprp | - | - | - | - | - | - | - | - | - | - |
| Xgrs | - | - | - | - | - | - | - | - | - | - |
| Xsps | - | - | - | - | - | - | - | - | - | - |
| XMg | 0.43 | 0.39 | 0.41 | 0.39 | 0.41 | 0.39 | 0.52 | 0.48 | 0.47 | 0.46 |
| Xan | - | - | - | - | - | - | - | - | - | - |
| Xab | - | - | - | - | - | - | - | - | - | - |
| Xor | - | - | - | - | - | - | - | - | - | - |
| Mineral | Kfs | Kfs | Kfs | Pl | Pl | Pl | Pl | Pl | Pl | Pl |
| Occurrence | Mtx | Mtx | Mtx | Mtx | Mtx | Mtx | Mtx | Inc in Grt | Inc in Kfs | Inc in Qz |
| Part | Rim | Rim | Core | Core | Core | Rim | Rim | Core | Core | Core |
| In contact with | - | - | - | - | - | - | - | - | - | - |
| Analysis no. | p008 | p010 | p015 | p096 | p113 | p033*1 | p106*1 | p189 | p009 | p090 |
| SiO2 | 64.62 | 64.03 | 64.48 | 63.89 | 62.27 | 62.55 | 62.59 | 62.76 | 62.81 | 62.75 |
| TiO2 | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Al2O3 | 18.68 | 18.16 | 18.36 | 22.13 | 22.99 | 23.22 | 22.65 | 23.23 | 23.09 | 22.90 |
| Cr2O3 | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| FeO | 0.03 | 0.00 | 0.06 | 0.05 | 0.00 | 0.34 | 0.07 | 0.34 | 0.00 | 0.01 |
| MnO | 0.00 | 0.01 | 0.02 | 0.02 | 0.05 | 0.05 | 0.00 | 0.02 | 0.00 | 0.00 |
| MgO | 0.00 | 0.00 | 0.02 | 0.00 | 0.01 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 |
| CaO | 0.00 | 0.07 | 0.05 | 3.34 | 4.66 | 4.56 | 4.06 | 4.11 | 4.29 | 4.38 |
| ZnO | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| BaO | 0.63 | 0.39 | 0.57 | 0.00 | 0.00 | 0.00 | 0.11 | 0.00 | 0.06 | 0.10 |
| Na2O | 1.21 | 1.32 | 1.40 | 9.76 | 8.90 | 8.94 | 9.23 | 9.16 | 9.21 | 9.21 |
| K2O | 15.32 | 15.11 | 15.25 | 0.24 | 0.26 | 0.16 | 0.24 | 0.19 | 0.20 | 0.14 |
| F | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Cl | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| -O=F | - | - | - | - | - | - | - | - | - | - |
| -O=Cl | - | - | - | - | - | - | - | - | - | - |
| Total | 100.49 | 99.09 | 100.20 | 99.43 | 99.13 | 99.81 | 98.95 | 99.81 | 99.67 | 99.49 |
| Number of O | 8 | 8 | 8 | 8 | 8 | 8 | 8 | 8 | 8 | 8 |
| Si | 2.98 | 2.99 | 2.98 | 2.84 | 2.78 | 2.78 | 2.80 | 2.78 | 2.79 | 2.79 |
| Ti | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Al | 1.02 | 1.00 | 1.00 | 1.16 | 1.21 | 1.22 | 1.19 | 1.21 | 1.21 | 1.20 |
| Cr | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Fetotal | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.01 | 0.00 | 0.01 | 0.00 | 0.00 |
| Mn | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 |
| Mg | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 |
| Ca | 0.00 | 0.00 | 0.00 | 0.16 | 0.22 | 0.22 | 0.19 | 0.20 | 0.20 | 0.21 |
| Zn | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Ba | 0.01 | 0.01 | 0.01 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 | 0.00 |
| Na | 0.11 | 0.12 | 0.13 | 0.84 | 0.77 | 0.77 | 0.80 | 0.79 | 0.79 | 0.79 |
| K | 0.90 | 0.90 | 0.90 | 0.01 | 0.01 | 0.01 | 0.01 | 0.01 | 0.01 | 0.01 |
| F | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Cl | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. | n.d. |
| Total cation | 5.02 | 5.02 | 5.03 | 5.01 | 5.00 | 5.00 | 5.01 | 5.01 | 5.01 | 5.01 |
| Xalm | - | - | - | - | - | - | - | - | - | - |
| Xprp | - | - | - | - | - | - | - | - | - | - |
| Xgrs | - | - | - | - | - | - | - | - | - | - |
| Xsps | - | - | - | - | - | - | - | - | - | - |
| XMg | - | - | - | - | - | - | - | - | - | - |
| Xan | 0.00 | 0.00 | 0.00 | 0.16 | 0.22 | 0.22 | 0.19 | 0.20 | 0.20 | 0.21 |
| Xab | 0.11 | 0.12 | 0.12 | 0.83 | 0.76 | 0.77 | 0.79 | 0.79 | 0.79 | 0.79 |
| Xor | 0.89 | 0.88 | 0.88 | 0.01 | 0.01 | 0.01 | 0.01 | 0.01 | 0.01 | 0.01 |
Notations of ‘In contact with’ are described as follows. For garnet rim analyses: mineral in contact with the analysis point. For biotite: Grt, grain is in contact with garnet; Ms, grain is in contact with muscovite without garnet; Felsic, the grain is in contact with felsic minerals such as plagioclase, K-feldspar, or quartz without garnet or muscovite. For muscovite: Bt, grain is in contact with biotite without garnet. All the analysis points are far from ilmenite.
Abbreviations: Bt, biotite; Grt, garnet; Inc, inclusion; Kfs, K-feldspar; Ms, muscovite; Mtx, matrix; Pl, plagioclase.
Xalm = Fetotal/(Fetotal + Mg + Ca + Mn) in atom, Xprp = Mg/(Fetotal + Mg + Ca + Mn) in atom, Xgrs = Ca/(Fetotal + Mg + Ca + Mn) in atom, Xsps = Mn/(Fetotal + Mg + Ca + Mn) in atom, XMg = Mg/(Fetotal + Mg) in atom, Xan = Ca/(Ca + Na + K) in atom, Xab = Na/(Ca + Na + K) in atom, and Xor = K/(Ca + Na + K) in atom.
n.d., not determined.
*1 Compositions used for garnet-biotite geothermometry and garnet-sillimanite-plagioclase-quartz geobarometry.
*2 Composition used for Ti-in-biotite geothermometry.
Garnet occurs as porphyroblast (up to 2 mm in diameter) or rare inclusion in plagioclase (up to 0.1 mm in diameter) (Figs. 5a and 5b). Porphyroblastic garnet is slightly to highly elongated along the foliation and commonly contains abundant inclusions of biotite, plagioclase, or quartz (Figs. 5a and 5b). Porphyroblastic garnet represents a monotonic increase in spessartine content [Xsps = Mn/(Fetotal + Mg + Ca + Mn)] and a decrease in almandine content [Xalm = Fetotal/(Fetotal + Mg + Ca + Mn)] from core to rim irrespective of grain size, except for around cracks that are connected to the matrix (Fig. 7). Around the cracks in garnet, Xsps shows a similar value to the rim of garnet (Fig. 7). Local depletion in XMg [= Mg/(Mg + Fetotal)] is recognized at the rims of garnet, which is in contact with biotite in some cases. Garnet is poor in grossular content [Xgrs = Ca/(Fetotal + Mg + Ca + Mn)] as 0.03-0.04 (Fig. 7). The XMg of inclusion garnet in plagioclase is slightly higher than that of porphyroblastic garnet (Table 2).
Biotite occurs in the matrix (up to 2 mm long) or as inclusions in garnet (up to 0.3 mm long) (Figs. 5a-5c) and is chemically homogeneous within grains. Biotite included in garnet shows higher XMg of 0.46-0.52 and lower Ti of 0.23-0.32 apfu (O = 22) than that of matrix biotite (Fig. 8). Matrix biotite that is in contact with the garnet rim shows slightly higher XMg of 0.39-0.43 than that of matrix biotite that contacts muscovite (0.39-0.40) or felsic minerals such as plagioclase, K-feldspar, and quartz (0.39-0.41) (Fig. 8). Ti contents of matrix biotite are 0.26-0.36 apfu for that in contact with garnet, 0.33-0.44 apfu for that in contact with muscovite, and 0.28-0.35 apfu for that in contact with felsic minerals (Fig. 8).
Plagioclase occurs in the matrix (up to 1 mm in diameter) or as inclusions in garnet (up to 0.3 mm in diameter) (Figs. 5a-5c). Matrix plagioclase cores have low anorthite content [Xan = Ca/(Ca + Na + K)] (0.16-0.22, average = 0.19), whereas that of rims is higher (0.19-0.22, average = 0.21) (Fig. 9). The Xan of inclusion plagioclase in quartz, K-feldspar, or garnet is 0.19-0.21 (Fig. 9).
Muscovite occurs in the matrix (up to 1.8 mm long) and as inclusions in plagioclase (up to 0.1 mm long) (Figs. 5a and 5c). Matrix muscovite is often associated with biotite (Figs. 5a and 5c). Muscovite is not chemically distinguished in terms of occurrence and exhibits compositions of Si = 6.09-6.20 apfu (O = 22), FeO + MgO = 1.43-2.38 wt%, XMg = 0.43-0.49, and Na = 0.09-0.14 apfu (Table 2).
K-feldspar is also found in the matrix (up to 1 mm in diameter) (Figs. 5a and 5c). K-feldspar shows narrow perthitic lamellae; however, the composition of each lamella could not be determined because of their narrow width as compared with the beam diameter. Therefore, the composition reported here is that of mixed analysis. The chemical composition of K-feldspar was Xor [= K/(Ca + Na + K)] of 0.84-0.91, Xab [= Na/(Ca + Na + K)] of 0.09-0.16, and Xan of 0.00 (Table 2). K-feldspar contains BaO of 0.39-0.70 wt% also (Table 2).
Sillimanite occurs as fibrolite in the matrix (up to 3 mm long) and is often associated with biotite (Figs. 5a and 5c).
Monazite occurs both in the matrix and as inclusions in biotite and garnet (up to 0.1 mm long) (Fig. 6). Some of the grains in the matrix are highly irregular in shape and incompletely include sillimanite ± biotite (Figs. 6b and 6c). Monazite is generally composed of three domains of dark, medium, and bright parts in the back-scattered electron image (Fig. 6). Monazite shows complex (Figs. 6a and 6b) or concentric zonation (Figs. 6c-6f) defined by the brightness in the back-scattered electron image and the rims are generally dark.
A total of 72 analyses were conducted on 29 grains, yielding ages ranging from 1015 to 531 Ma (Fig. 10). The results of the monazite analysis are summarized in Supplementary Table S1 (Table S1 is available online from https://doi.org/10.2465/jmps.221124). All analytical spots are shown in Supplementary Figure S1 (Fig. S1 is available online from https://doi.org/10.2465/jmps.221124), and some in Figure 6 also. According to the mode of occurrence of monazite, 974-531 Ma was obtained from matrix monazite (number of analysis point, n = 45), 1015-651 Ma from inclusions in biotite (n = 26), and 926 Ma from an inclusion in garnet (n = 1). All 72 ages were input into IsoplotR (Vermeesch, 2018) to calculate a weighted mean age of 940.1 ± 9.8 Ma (MSWD = 0.31, 2σ level) using 62 of the 72 ages of 1015-889 Ma (Fig. 10b). The youngest 10 of the 72 ages obtained (827-531 Ma) was rejected from the calculation by outlier detection in IsoplotR (Vermeesch, 2018) (Fig. 10b). The adopted 62 ages overlap each other within their 2σ errors and MSWD is smaller than unity (Fig. 10b), indicating the timing of a specific geological event. Focusing on shape of monazite, some grains show highly irregular shape and incompletely include sillimanite ± biotite in the matrix, as described previously (Figs. 6b and 6c). Sillimanite and biotite are major phases within the gneiss and are likely of metamorphic origin, and thus we consider that the monazite showing such texture was metamorphic rather than detrital in origin. In addition, analyses on both sides of the incompletely included sillimanite show ages of 945 ± 36 Ma and 906 ± 36 Ma for points m16.5 and m16.6, respectively (Fig. 6b). Therefore, we regard that the weighted mean age of 940.1 ± 9.8 Ma is the timing of metamorphism of Niban-nishi Rock.
The 10 ages that were rejected are 827, 742, 659, 651, 627, 592, 567, 547, 532, and 531 Ma (Fig. 10b and Table S1). These younger ages were mostly obtained from the darker parts in the back-scattered electron images (Figs. 6a-6c, S1, and Table S1), which are most often found at the rim of monazite (Figs. 6a and 6c) and locally dig into the interior of the grain (Fig. 6b). The youngest 10 data show a weighted mean age with a MSWD of >3, indicating that they do not represent the timing of a specific single event. The implication of these younger ages is discussed in a later section.
Compositional zoning in garnet shows an outward decrease in Xalm and an increase in Xsps (Fig. 7). The rim of garnet is partially replaced by biotite (Figs. 5a and 5b), and biotite contacting garnet shows relatively low Ti content (Fig. 8). Ti content in biotite generally decreases with temperature (e.g., Guidotti, 1984). Plagioclase in the examined gneiss shows slightly higher anorthite content at the rim as compared with that in the core (Fig. 9). Anorthite content in plagioclase of pelitic rocks generally increases during cooling or decompression (e.g., Spear et al., 1990). These features can be interpreted as garnet being affected by resorption and modification of the garnet and plagioclase compositions during the retrograde stage. The compositions of rims comprising other matrix minerals can also be considered to have re-equilibrated, with the retrograde conditions of regional metamorphism preserved as their compositions. Therefore, we used the compositions of the rim of porphyroblastic garnet, matrix biotite, and the rim of matrix plagioclase to estimate the P-T conditions of the retrograde stage. The garnet-biotite geothermometer of Holdaway (2000) and the garnet-sillimanite-plagioclase-quartz geobarometer of Holdaway (2001) were applied for estimation. For garnet and biotite, compositions that were not in contact with each other or other Fe-Mg minerals (e.g., muscovite or ilmenite) were selected to eliminate the effect of local re-equilibration at the closure temperature. The minimum and maximum values of XMg in garnet and biotite and those of anorthite content in plagioclase were taken into account for the calculation. The analyses used for the P-T calculation are listed in Table 2. PTQuick software (Simakov and Dolivo-Dobrovolsky, 2009) was used for the calculations and results of 620-670 °C and 0.42-0.60 GPa were thus obtained as the retrograde conditions of regional metamorphism (Figs. 2 and 11).
We do not calculate the P-T condition of the prograde to peak stages of metamorphism based on conventional geothermobarometers that consider cation-exchange and net-transfer reactions by using mineral chemical compositions of the core or inclusions because there is no evidence that they equilibrated with each other (e.g., Ikeda, 2004).
Phase equilibrium modeling was performed to obtain stable mineral assemblages under each P-T condition. P-T pseudosections were computed using the free energy minimization software Perple_X version 6.9.1. (Connolly, 2005, 2009, updated September 24, 2022) and the thermodynamic dataset (hp62ver.dat) of Holland and Powell (2011) in the simplified K2O-Na2O-CaO-FeO-MgO-MnO-Al2O3-TiO2-SiO2-H2O system. The solid-solution models for garnet, biotite, white mica, cordierite, chloritoid, chlorite, staurolite, ilmenite, and melt of White et al. (2014) and that for plagioclase and K-feldspar of Fuhrman and Lindsley (1988) were used in the calculation. Other phases were treated as pure. All iron was treated as ferrous. We used whole-rock composition of the melanocratic portion shown in Table 1 for the modeling. We consider that it preserves the composition around and after the peak metamorphic stages because of the following reasons. The examined gneiss shows migmatitic appearance (Figs. 3b and 4); however, it is difficult to identify the occurrence of partial melting because the pelitic gneiss has been intruded by a pegmatite dyke (Fig. 3a). Two possible origins for the leucosome and leucocratic patch in the melanosome can be considered. One is that melt has been derived from partial melting of the pelitic gneiss, and the other is that melt infiltrated from outside the gneiss. We cannot evaluate which theory is appropriate interpretation for the origin in the present study. Nevertheless, the foliation in the melanocratic portion and the long axis of the leucosome, leucocratic patch, and biotite in the leucocratic patch are parallel (Figs. 3b, 4, and 5). These results indicate that the leucosome and leucocratic patch formed prior to foliation. Sillimanite, biotite, muscovite, and quartz that show shape-preferred orientation and define the foliation in the melanocratic portion are the major constituting phases of the gneiss and should have formed around the peak stage of the metamorphism. Thus, we consider that the whole-rock composition of the melanocratic portion would not change near or after the peak metamorphic stage, and that constructing the pseudosection from the composition is valid for estimating the stable mineral assemblages at these stages. Water content cannot be directly determined from the value of loss on ignition, which is affected by loss of other volatile elements such as sulfur and carbon and oxidation during ignition. Attempts at modelling using whole-rock chemistry with a water content of 0.69, which is the value of loss on ignition implied that the present matrix mineral assemblage (garnet + biotite + muscovite + plagioclase + K-feldspar + quartz + sillimanite + ilmenite) did not appear in the P-T pseudosection. The assemblage also did not appear when the water content was set to >1.3 wt%. Thus, we assumed it to be 1 wt%, allowing the composition of SiO2, 69.395; TiO2, 0.815; Al2O3, 14.763; FeO, 6.131; MnO, 0.095; MgO, 2.129; CaO, 0.780; Na2O, 1.675; K2O, 3.218; H2O, 1.000 in wt% (normalized to 100 wt%) to be used for the calculation. The obtained P-T pseudosection is shown in Figure 11. The volume fraction of garnet was also calculated using Perple_X and the contours given in Figure 11. The present mineral assemblage appears to be in the range 538-737 °C and 0.37-0.82 GPa (Fig. 11). This result is consistent with the P-T estimation of the retrograde condition using the garnet-biotite geothermometer of Holdaway (2000) and the garnet-sillimanite-plagioclase-quartz geobarometer of Holdaway (2001). Therefore, we use the whole-rock composition containing 1.000 wt% H2O to construct P-T pseudosection.
Ti content in the matrix biotite varied from 0.26-0.44 apfu (O = 22) despite a similar XMg (Fig. 8). The variation in Ti content can be interpreted as variation in the crystallization temperature (e.g., Guidotti, 1984). The application of Ti-in-biotite geothermometers occasionally results in revealing higher metamorphic temperatures than those obtained using garnet-biotite geothermometers based on the Fe-Mg exchange reaction, which may be due to sluggish diffusion of Ti in biotite after crystallization (e.g., Wu et al., 2017; Chen et al., 2021). Thus, Ti content in biotite likely preserves the temperature of formation of grain rather than that of later re-equilibration. The presence of ilmenite in the examined sample allows application of the geothermometer of Wu and Chen (2015), which was therefore applied to matrix biotite to estimate the highest temperature recorded in the sample. Using the composition with the highest Ti content (Fig. 8 and Table 2), the highest temperature of 690-722 °C was obtained, assuming the pressure range of 0.50-0.80 GPa, which corresponds to stability fields of garnet and sillimanite in the P-T pseudosection (Fig. 11). The volume fraction of garnet was further applied in the P-T space to constrain the peak P-T condition. The areal proportion of ∼ 0.6% for garnet in the examined gneiss has been described previously. Garnet is partially replaced by biotite at the periphery, but the amount of the biotite observed is minimal; hence, the volume fraction of garnet remains largely unchanged in the examined gneiss, as determined from textural observation (Figs. 5a and 5b). These results indicate that the present areal proportion of ∼ 0.6 vol% is almost the same as that of peak conditions. In addition, we believe that metamorphism occurred outside of the cordierite stable conditions because cordierite was absent in the examined gneiss. Cordierite is stable at the high-T/low-P conditions observed in the pseudosection shown in Figure 11. Considering the volume fraction of garnet in the P-T pseudosection, temperature estimation using Ti-in-biotite geothermometer, and the cordierite stability fields in the P-T pseudosection, the peak metamorphic condition can be constrained to 690-730 °C at 0.38-0.68 GPa (Fig. 11). Overall, the metamorphic grade of Niban-nishi Rock would be lower than the typical exposures of the LHC that are located in the transitional and granulite-facies zones (e.g., Akarui Point and Skallen) (Fig. 2), which is consistent with the metamorphic zonation in Hiroi et al. (1991).
The metamorphic age of 940.1 ± 9.8 Ma obtained in this study is a significant weighted mean age. Preliminary reports of LA-ICP-MS U-Pb zircon dating for garnet-biotite-sillimanite gneiss from Niban-nishi Rock by Kitano et al. (2021) showed a metamorphic age of 998 Ma, which corresponds to the Tonian age. Thus, Niban-nishi Rock has experienced the Tonian metamorphic event. The metamorphic age of Niban-nishi Rock is consistent with that of neighboring exposures of Akebono Rock of 937 ± 6 Ma deduced using pelitic gneiss reported by Baba et al. (2022) and Cape Hinode of 970 Ma deduced using garnet metapelite and 960 Ma deduced using metatrondjemite reported by Dunkley et al. (2014). The metamorphic age of 940 Ma determined in this study is significantly older than that of the LHC (600-520 Ma). Therefore, we consider that Niban-nishi Rock should be treated as an exposure that would have tectonic relation to Cape Hinode and Akebono Rock rather than typical exposures of the LHC.
In addition to the Tonian metamorphic age of 940 Ma, a few younger monazite ages of 827-531 Ma were also detected, which partially overlap the metamorphic age of the LHC. Dunkley et al. (2014) reported 532 ± 7 Ma as the metamorphic age of a metagranitic dyke in Niban Rock, although no information is given describing the exact locality in Niban Rock or the petrography of the metagranitic dyke. Baba et al. (2022) also reported a few younger U-Th-Pb monazite ages of 769-502 Ma for Akebono Rock using garnet-biotite gneisses. Based on these lines of evidence, Niban-nishi Rock and Akebono Rock might have experienced later metamorphism that is related to the LHC. However, it is difficult to determine whether the younger ages obtained in this study represent the timing of a later metamorphic event because an insufficient number of data points from younger ages are available. A cluster comprising a significant number of ages is required for the results of EMP monazite dating to be recognized as a specific event because the method used does not provide information that can be used to identify whether each obtained age is concordant or discordant. Further, the younger ages vary from 531 to 827 Ma. If the variation can be explained by mixed analysis of domains of monazite with older and younger ages due to the large beam diameter applied in this study, there is a possibility that more datapoints indicating younger domains are present within the monazite. Therefore, we consider that Niban-nishi Rock might have experienced the later metamorphism stage that is related to the LHC, and we suggest that further study on the metamorphic ages of Niban-nishi Rock is required.
Although the metamorphic ages of Niban-nishi Rock, Cape Hinode, and Akebono Rock overlap each other in the Tonian period, the stable phase of the aluminosilicate minerals in the matrix and the compositional zoning in garnet vary. The pelitic gneiss examined in this study represents retrograde zoning in garnet (Fig. 7) and the presence of sillimanite in the matrix as a stable aluminosilicate phase (Figs. 5a-5c). These results differ from the pelitic gneisses in Akebono Rock, which contain garnet representing prograde zoning and kyanite as a stable matrix phase (Baba et al., 2022). Cape Hinode is known to have experienced granulite facies metamorphism, as evidenced by the presence of orthopyroxene in the metamorphosed mafic enclaves in the metatonalite (Hiroi et al., 2006, 2008) and the high metamorphic grade as compared with that in Niban-nishi Rock. Thus, Niban-nishi Rock is an exposure showing a new character in terms of metamorphism on the Prince Olav Coast. Sillimanite in the matrix was also reported in sillimanite-garnet-biotite gneiss from Akebono Rock by Hiroi et al. (1986b) and in pelitic gneiss from Cape Hinode by Yanai and Ishikawa (1978) and Hiroi et al. (2006). These gneisses may be related to our studied gneiss from Niban-nishi Rock; however, the relationship cannot be clarified because no information is available on the gneisses from Akebono Rock and Cape Hinode, except for that inferred by textural observation.
This study revealed that Niban-nishi Rock records Tonian metamorphism at 940 Ma; thus, we concluded that Niban-nishi Rock is related to the Tonian metamorphic terrane proposed for Cape Hinode (the Hinode Block; Dunkley et al., 2020) and Akebono Rock (Baba et al., 2022). We clarified that the examined pelitic gneiss experienced peak conditions of 690-730 °C/0.38-0.68 GPa, which correspond to upper amphibolite facies, and retrograde condition of 620-670 °C/0.42-0.60 GPa, and that it contains garnet that represents retrograde zoning and stable sillimanite in the matrix. The metamorphic features recorded in the examined gneiss differ from those of Cape Hinode, which show a higher metamorphic grade that corresponds to granulite facies (Hiroi et al., 2006, 2008), and that of Akebono Rock, whose pelitic gneiss includes stable kyanite in the matrix and garnet with prograde zoning (Baba et al., 2022), although they show similar metamorphic ages. Further field survey and detailed mapping of the variations in metamorphic conditions and ages in Niban Rock are required to clarify the boundary between the LHC and areas showing Tonian metamorphism, the relations among exposures on the Prince Olav Coast, and the tectonic evolution of the eastern Dronning Maud Land of East Antarctica from the Mesoproterozoic to the Neoproterozoic.
This study was part of the Science Program of the Japanese Antarctic Research Expedition (JARE) and was supported by the National Institute of Polar Research (NIPR) under MEXT. We acknowledge Daniel J. Dunkley (JARE 52, Polish Academy of Sciences) for his support in the geological fieldwork at Niban Rock during JARE 52, all the members of JARE 52, and the crew of the icebreaker Shirase. We are grateful to Kazuhiko Shimada (Kyushu University) for their technical support with the mineral chemical analyses. We are grateful to the two anonymous reviewers for constructive reviews and Tetsuo Kawakami (Kyoto University) for the editorial efforts. This study was financially supported by JSPS KAKENHI Grant Numbers JP21K14015 to YM, JP21H01182 to TH, and JPC25400518 to TI and NIPR through General Collaboration Project nos. 28-23 and 4-17 to TI and no. 31-24 to YM and TI.
Supplementary Table S1 and Figure S1 are available online from https://doi.org/10.2465/jmps.221124.
