Metamorphic pressure–temperature conditionsof the Lützow–Holm Complex of East Antarctica deduced

from Zr–in–rutile geothermometerand Al2SiO5 minerals enclosed in garnet

Kota SUZUKI and Tetsuo KAWAKAMI

Department of Geology and Mineralogy, Graduate School of Science, Kyoto University, Kyoto 606–8502, Japan

The metamorphic pressure–temperature (P–T ) conditions of high–grade pelitic gneisses from Akarui Point,Skarvsnes, Skallen, and Rundvågshetta in the Lützow–Holm Complex (LHC), East Antarctica are re–examinedby applying the Zr–in–rutile geothermometer to rutile inclusions in garnet enclosing Al2SiO5 minerals and tomatrix rutile grains. By utilizing the P zoning of garnet to indicate isochronous surface, samples from AkaruiPoint, Skarvsnes, and Skallen were shown to have experienced almost the same P–T conditions around thekyanite/sillimanite transition boundary (~ 830–850 °C/~ 11 kbar). From Rundvågshetta, higher–T condition(850 ± 15 °C/0.1 kbar to 927 ± 16 °C/12.5 kbar) was confirmed from rutile inclusions in garnet enclosingsillimanite. Matrix rutile yielded similar temperature as inclusion rutile for Akarui Point, Skarvsnes, and Skallensamples. Therefore, the traditional metamorphic zone mapping based on matrix mineral assemblages, whichclassified Akarui Point as belonging to the transitional zone between the upper–amphibolite and the granulitefacies zones, does not reflect the highest metamorphic conditions attained. The P–T–t evolution of the LHCneeds to be re–evaluated utilizing detailed petrochronological approaches.

Keywords: Chemical zoning, Phosphorus, Al2SiO5 minerals, Zr–in–rutile geothermometer

INTRODUCTION

The Lützow–Holm Complex (LHC) of East Antarctica iscomposed of high–grade metamorphic rocks of variouscomposition and formed during the late Neoproterozoicto early Cambrian continental collision events of theGondwana amalgamation (e.g., Shiraishi et al., 2008).Based on matrix mineral assemblages of mafic– to inter-mediate–gneisses, Hiroi et al. (1991) divided the complexinto three metamorphic zones; the upper–amphibolite fa-cies zone, the transitional zone, and the granulite facieszone (Fig. 1). Because of the continuous change in meta-morphic grade from the upper–amphibolite to the granu-lite facies, the LHC has been considered as a cross sectionof the lower continental crust at the active margin (e.g.,Hiroi et al., 1991). However, recent studies based on the

ternary–feldspar thermometry of a pelitic gneiss (Naka-mura et al., 2014) and a pseudosection analysis of a meta-

Figure 1. Simplified metamorphic zone map of the Lützow–HolmComplex (LHC; Hiroi et al., 1991) showing the localities ofAkarui Point, Skarvsnes, Skallen and Rundvågshetta. A rectan-gle in an inset represents the location of the LHC in Antarctica.

doi:10.2465/jmps.190801K. Suzuki, suzuki.kouta.22u@st.kyoto–u.ac.jp Corresponding au-

thorT. Kawakami, kawakami.tetsuo.3e@kyoto-u.ac.jp

Journal of Mineralogical and Petrological Sciences, Volume 114, page 267–279, 2019

mafic rock (Iwamura et al., 2013) from Akarui Point (Fig.1) have revealed that the high–T condition above 850 °Cwas likely attained even in the ‘transitional zone’. Further-more, it was suggested that such a temperature was pos-sibly attained under relatively high–P condition aroundthe kyanite/sillimanite transition boundary (Iwamura etal., 2013) (Fig. 2a). Such a high–T under the kyanite sta-bility field was recently reported from elsewhere in theLHC (Skallevikshalsen, Kawakami et al., 2016; Run-dvågshetta, Hiroi et al., 2019) (Figs. 2c and 2d). In orderto comprehensively understand the P–T–t evolution of theentire LHC, precise estimates focused on the inclusionminerals in garnet from the entire region are required.

Previous studies estimating the metamorphic P–Tconditions of the LHC have commonly applied the Fe–Mg exchange type geothermometers. These studies utiliz-ed pairs of matrix minerals which are not in direct contactwith each other in order to minimize the effect of retro-grade re–equilibrium (e.g., Yoshimura et al., 2004). Aspointed out by Ikeda (2004), however, the assumption thatthe cores of matrix minerals preserve chemical composi-tions at the peak metamorphism is suspicious. Chemicalzoning of garnet in high–T metamorphic rocks in terms ofmajor elements (Fe, Mg, Ca, and Mn) is often obscured bythe volume diffusion under prolonged high–T conditions(e.g., Hiroi et al., 1994). On the other hand, chemical zon-ing of garnet in terms of P is less affected by the laterdiffusion and commonly preserves sharp zoning (e.g.,Hiroi et al., 1997; Kawakami and Motoyoshi, 2004). Uti-lizing the zoning in P as a contemporaneous surface dur-ing the garnet growth, relic minerals included in garnetsuch as Al2SiO5 minerals and rutile can be utilized to con-strain the P–T conditions of each metamorphic stage evenif the chemical zoning of garnet in divalent major cationsis modified by the later diffusion (e.g., Kawakami andHokada, 2010). It is known that Zr content in rutile ismainly a function of temperature when it is in equilibriumwith zircon and quartz (e.g., Zack et al., 2004; Tomkins etal., 2007). Since the diffusion of Zr in garnet is too slow tochange Zr content in rutile by diffusion (Zack et al., 2004),once included in garnet and chemically isolated from zir-con and quartz, rutile can preserve the Zr contents at thetime of entrapment, that is, the temperature condition ofthe garnet growth.

The aim of this study is to deduce the P–T condi-tions recorded as inclusion minerals in garnet and as ma-trix minerals in high–grade pelitic gneisses from selectedareas of the LHC (Akarui Point, Skarvsnes, Skallen, andRundvågshetta) (Fig. 1) and re–evaluate the P–T evolu-tion of the LHC. The temperature conditions were con-strained by applying the Zr–in–rutile geothermometer(Zack et al., 2004; Tomkins et al., 2007), which is robust

against retrograde re–equilibrium (e.g., Kelsey and Hand,2015), to rutile grains enclosed in garnet and present inthe matrix. In this study, experimentally calibrated Zr–in–rutile geothermometer that takes the pressure effect intoaccount (Tomkins et al., 2007) is preferred. Mineral ab-breviations are after Whitney and Evans (2010).

GEOLOGICAL SETTING

The LHC is located in eastern Dronning Maud Land andis composed mainly of high–grade metamorphic rockssuch as pelitic–psammitic gneisses and mafic–intermedi-ate gneisses, along with subordinate amount of ultramaficlenses, marbles, and calcsilicate rocks. Based on the met-amorphic field mapping utilizing matrix mineral assem-blages in rocks with various bulk compositions, the met-amorphic grade of the LHC is considered to progres-sively increase southwestwards from the upper–amphib-olite facies in the Prince Olav Coast to the granulite faciesin Lützow–Holm Bay (Hiroi et al., 1991) (Fig. 1). Theoccurrence of kyanite and staurolite as relict inclusionsenclosed in garnet or plagioclase, and the development ofsymplectites replacing garnet rims indicating isothermaldecompression have been interpreted to be the evidencefor a clockwise P–T path (e.g., Hiroi et al., 1991). A ther-mal axis of the peak metamorphism is estimated to lie atthe southern end of Lützow–Holm Bay around Rundvåg-shetta (Motoyoshi, 1986) (Fig. 1).

Akarui Point in the Prince Olav Coast is about 180km northeast of Rundvågshetta and located in the transi-tional zone (Hiroi et al., 1991; Fig. 1). Peak metamorphicP–T condition was previously estimated at 770–790 °C/7.7–9.8 kbar by applying Grt–Bt geothermometers andGrt–Al2SiO5–Qtz–Pl (GASP) geobarometers to a Sil–Bt–Grt leucogneiss (Kawakami et al., 2008) (Fig. 2a). Re-cently, Nakamura et al. (2014) estimated slightly highertemperature condition of 825–900 °C by utilizing a ter-nary–feldspar thermometry (Fig. 2a), and Iwamura et al.(2013) reported the prograde metamorphic condition of~ 900 °C/11–12 kbar, followed by isothermal decompres-sion to the peak metamorphic condition of 900–920 °C/5–6 kbar by applying the pseudosection approach to a met-amafic rock (Fig. 2a).

Skarvsnes is about 50 km northeast of Rundvågshet-ta and located in the granulite facies zone (Hiroi et al.,1991; Fig. 1). The metamorphic P–T condition ofSkarvsnes has not been quantitatively estimated in detail.

Skallen is about 30 km northeast of Rundvågshettaand located in the granulite facies zone (Fig. 1). Skallenand neighboring Skallevikshalsen have some lithologiesand structures in common (Yoshida, 1977). MetamorphicP–T conditions and P–T paths of Skallen and Skallevik-

K. Suzuki and T. Kawakami268

shalsen have been estimated by several previous studies.Satish–Kumar and Wada (2000) estimated the peak tem-perature condition of 848 ± 55 °C by applying carbonisotope thermometry between calcite and graphite to mar-bles from Skallen (Fig. 2c). Yoshimura et al. (2004) esti-mated the peak P–T condition of 770–940 °C/6.5–12.0kbar by applying Grt–Bt geothermometers and GASP ge-obarometers to a Grt–Bt gneiss and 780–960 °C/6.0–11.0

kbar by applying Grt–Opx/Cpx geothermometers andGrt–Opx/Cpx–Pl–Qtz geobarometers to a mafic gneissfrom Skallevikshalsen (Fig. 2c). On the other hand, Ikeda(2004) estimated the retrograde condition of 710–800 °C/4.8–6.4 kbar by applying Grt–Bt geothermometers andGASP geobarometers to the rims of matrix garnet andbiotite which are not in direct contact with each otherin a garnet–rich pelitic gneiss from Skallen (Fig. 2c). Sat-

Figure 2. Summary of the P–T conditions and P–T paths of the studied areas of the LHC determined by previous studies and by this study.Phase diagram of Al2SiO5 minerals is after Pattison (1992). (a) Akarui Point. k2008, Kawakami et al. (2008); I2013, Iwamura et al. (2013);N2014, Nakamura et al. (2014). (b) Skarvsnes. Dotted arrow shows P–T path determined by this study. (c) Skallen and Skallevikshalsen.SW2000, Satish–Kumar and Wada (2000); I2004, Ikeda (2004); Y2004, Yoshimura et al. (2004); S2006, Satish–Kumar et al. (2006);Ks2013, Kawasaki et al. (2013); Kk2016, Kawakami et al. (2016). (d) Rundvågshetta. The vapor saturated solidus is after Johannes andHoltz (1996). MI1997, Motoyoshi and Ishikawa (1997); KsM2006, Kawasaki and Motoyoshi (2006); Y2008, Yoshimura et al. (2008);H2019, Hiroi et al. (2019). Color version is available online from https://doi.org/10.2465/jmps.190801.

Pressure–temperature estimates of the Lützow–Holm Complex, Antarctica 269

ish–Kumar et al. (2006) revealed the metamorphic fluidhistory along a clockwise P–T path (Fig. 2c) by focusingon scapolite boudins within marbles from Skallen andestimated the minimum P–T condition of CO2–rich fluidinfiltration as 600 °C/3.8 kbar during the retrograde de-compression stage. Kawasaki et al. (2013) proposed theultrahigh–T (UHT ) condition exceeding 1000 °C basedon the occurrence of possible armalcolite pseudomorphin a Grt–Sil gneiss from Skallevikshalsen (Fig. 2c). Re-cently, Kawakami et al. (2016) constrained metamorphicP–T condition of 820–850 °C at >11 kbar by applying Zr–in–rutile geothermometer to rutile coexisting with kyaniteboth enclosed in the P–poor core of garnet porphyroblastsin a Grt–Sil gneiss from Skallevikshalsen (Fig. 2c). Thepossibility of polymetamorphism was proposed in Skall-en and Skallevikshalsen (Kawakami et al., 2016) basedon two populations of metamorphic ages detected by U–Th–Pb monazite and U–Pb zircon datings of metapeliticgneisses from Skallen and Skallevikshalsen (Hokada andMotoyoshi, 2006; Kawakami et al., 2016).

Rundvågshetta records the highest metamorphicgrade in the LHC. Some characteristic mineral assem-blages for the UHT metamorphism such as Opx + Sil +Qtz (Motoyoshi and Ishikawa, 1997) and Spr + Qtz(Yoshimura et al., 2008) have been reported from Crd–bearing Grt–Opx and Grt–Opx–Sil granulites occurringexclusively in the northern part of Rundvågshetta. Moto-yoshi and Ishikawa (1997) estimated the peak metamor-phic condition of 1000 °C/11 kbar followed by subse-quent isothermal decompression by utilizing a petrogen-etic grid for the SiO2–Al2O3–MgO–FeO (FMAS) systemand a Grt–Opx geothermometer to pelitic granulites (Fig.2d). Kawasaki and Motoyoshi (2006) experimentallyconstrained the peak and retrograde metamorphic condi-tions of 925–1039 °C/11.5–15.0 kbar and 824–1010 °C/6.5–10.8 kbar, respectively, by utilizing the piston cylin-der apparatus for Sil–Crd–Spr granulites (Fig. 2d). Yoshi-mura et al. (2008) estimated the peak metamorphic con-dition of 1000–1100 °C/10–12 kbar based on the findingof Spr + Qtz and Opx + Sil + Qtz assemblages, and onapplication of ternary–feldspar geothermometry and Al–in–orthopyroxene geothermometry to a Spr–bearing Grt–Opx–Sil granulite (Fig. 2d). Recently, Hiroi et al. (2019)estimated the peak metamorphic condition of >940 °C byapplying the published univariant reaction lines for therutile–bearing K2O–FeO–MgO–Al2O3–SiO2–H2O–TiO2–

Fe2O3 (KFMASHTO) model system to the mineral as-semblages of UHT granulites. Hiroi et al. (2019) alsosuggested that such high–T condition was already attainedaround the kyanite/sillimanite transition boundary (Fig.2d) because high–XCa domain of garnet includes multi-phase inclusions of Ky + Spr and Sil + Spr, which are

interpreted as the breakdown products of Mg–rich stau-rolite under high–T and high–P condition. On the otherhand, Tsunogae et al. (2014) estimated the peak meta-morphic condition of <900 °C by applying the pseudo-section approach to a charnockite, which is a predomi-nant lithology there. They proposed that the UHT con-dition exceeding 1000 °C might be a local event limitedto the northern part of Rundvågshetta.

ANALYTICAL METHODS

Quantitative analysis of rock–forming minerals and X–rayelemental mapping were performed by the JEOL superp-robe JXA8105 at Department of Geology and Mineralo-gy, Kyoto University. Analytical conditions for quantita-tive analysis were 15.0 kV acceleration voltage, 10 nAprobe current with probe diameter of 3 µm. Natural andsynthetic minerals (Astimex MINM25–53) were utilizedfor standards, and ZAF correction was applied. Analyticalconditions for X–ray elemental mapping were 15.0 kVacceleration voltage with beam diameter up to 10 µm.Probe currents were 60 and 800 nA for major elementsand trace elements, respectively. Analytical conditions ofrutile for the Zr–in–rutile geothermometer followed thatrecommended by Zack et al. (2004), which were 20.0kV acceleration voltage, 120 nA probe current, and probediameter of 5 µm in Supplementary Table S1 (availableonline from https://doi.org/10.2465/jmps.190801). Count-ing error of X–ray intensities for Zr by the superprobe wasabout ±1%, which yielded negligible small temperaturevariation (±1 °C) in the result of the Zr–in–rutile geother-mometer (Tomkins et al., 2007). All analytical resultsof rutile utilized for calculation of temperature are givenin Supplementary Tables S2–S5 (available online fromhttps://doi.org/10.2465/jmps.190801). The polymorphsof Al2SiO5 minerals were identified by Raman spectrosco-py at Kyoto University (JASCO NRS 3100).

PETROGRAPHY

All samples utilized in this study were collected duringthe 44th Japanese Antarctic Research Expedition (JARE44) during the summer season of 2002–2003. Samplesselected for this study are representative of typical lithol-ogies from each outcrop.

Sillimanite–biotite–garnet gneiss from Akarui Point(sample TK2002122304)

This is the same sample studied by Kawakami et al.(2008) and Nakamura et al. (2014). The matrix of thissample consists mainly of garnet (~ 10 mm), biotite, sil-

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limanite, plagioclase, K–feldspar, and quartz with minorilmenite, rutile (~ 100–500 µm), apatite, zircon, monazite,and spinel. As noted by Nakamura et al. (2014), garnetporphyroblasts {Alm58.8–69.4Prp13.6–27.1Grs2.1–22.1Sps1.4–5.6,XMg[=Mg/(Mg + Fetotal)] = 0.19–0.28} are commonly as-sociated with pressure shadows consisting of leucosome,and the intergrowth of Bt (3.30–4.15 wt% TiO2) + Pl(An15–18) partly replaces the margin of garnet (Fig. 3a).From the center to margin of garnet, Fe, Mg, and XMg

increase, while Ca and Mn decrease. X–ray elementalmaps of garnet shows obscured chemical zoning in Fe,Mg, Ca, and Mn in the vicinity of the aggregated poly-phase inclusion of Pl (An20) + Kfs + Qtz (polyphase in-clusions are represented by using ‘+’, hereafter), whereassharp chemical zoning in P is not affected by the presenceof the polyphase inclusion (Figs. 3b and 3c). Sharp anddiscontinuous zoning in terms of P content in garnet de-fines the boundary between the P–poor cores and P–richrims (Figs. 3c and 3e). The boundary is locally irregu-larly–shaped (lower middle of Fig. 3c). Gradual increaseof P content towards the core/rim boundary is observed inthe core side, and this is discontinuously cut by the poly-phase inclusion (lower right of Fig. 3c). The P–poor coresof garnet include quartz, rutile (~ 50 µm), ilmenite, apa-tite, and zircon, whereas the P–rich rims include quartz,rutile (~ 50 µm), ilmenite, apatite, kyanite, sillimanite,biotite (3.03–5.26 wt% TiO2), and zircon (Figs. 3a, 3d,and 3e). Rutile inclusions separately enclosed in the P–poor cores and P–rich rims of garnet show Zr content of~ 1860 ± 320 ppm (average ± 2σ ppm hereafter; 14 pointson 14 grains) (Fig. 3c, Tables 1 and S1), and ~ 2040 ± 60ppm Zr (5 points on 1 grain) (Figs. 3c and 3d, Tables 1and S1), respectively. Rutile grains present in the matrixshow ~ 1940 ± 360 ppm Zr (16 points on 8 grains) (Ta-bles 1 and S1).

Sillimanite–garnet gneiss from Skarvsnes (sampleTK2003012803)

This sample was collected from an interboudin portion of aboudinaged Grt–Bt gneiss (Fig. 4a). The matrix of the sam-ple consists mainly of garnet (~ 3 mm), sillimaite, quartz,plagioclase (An25–26), and K–feldspar with minor rutile(~ 100–1000 µm), biotite (4.93–6.37 wt% TiO2), zircon,apatite, ilmenite, spinel, and pyrite. Garnet grains are com-positionally mostly homogeneous in terms of major ele-ments (Alm54.1–57.2Prp39.0–42.0Grs2.4–3.6Sps0.4–1.3), whereasthey are clearly zoned in terms of P. Therefore, garnetcan be divided into P–poor cores and P–rich rims (Figs.4c and 4f). The P–poor cores of garnet include Ky + Spl,rutile (~ 10–50 µm), biotite (4.74–7.70 wt% TiO2), apatite,zircon, and quartz (Figs. 4b and 4c). The P–rich rims of

garnet include Ky + Spl, sillimanite, rutile (~ 30 µm),quartz, zircon, and pyrite (Figs. 4e and 4f). Rutile enclosedin the P–poor cores of garnet, enclosed in the P–rich rims,and present in the matrix shows Zr content of ~ 2250 ± 500ppm (29 points on 29 rutile grains), 2320 ppm (1 point on 1rutile grain) (Figs. 4c and 4d), and ~ 2150 ± 660 ppm (10points on 5 rutile grains), respectively (Tables 1 and S2).

Spinel–sillimanite–garnet gneiss from Skallen (sampleTK2003011203)

In the outcrop observation, this sample contains abundantcm–sized garnet porphyroblasts up to ~ 5 cm in diameterand Spl + Sil aggregates (Fig. 5a). The matrix of thissample consists mainly of garnet, sillimanite, plagioclase(An33–39), and K–feldspar with minor spinel, ilmenite,apatite, zircon, monazite, and rutile (~ 500 µm). TheSpl + Sil aggregates are surrounded by plagioclase whoseanorthite content increases up to An42 at the contact withthe aggregates. The aggregates are never in direct contactwith quartz and K–feldspar (Fig. 5b). The cm–sized gar-net porphyroblast in the outcrop appears to be an aggre-gate of smaller–sized garnet of ~ 5 mm in diameter underthin section (Figs. 5a and 5b). The cm–sized garnet por-phyroblasts are mostly homogeneous in terms of majorelements (Alm67.8–71.9Prp22.4–25.7Grs4.0–5.8Sps0.6–1.9). Onthe other hand, chemical zoning in terms of P is compli-cated. P–rich domains of garnet are patchily distributed inthe inner part of the cm–sized garnet porphyroblast (Figs.5f and 5g). The P–rich patchy domains of garnet are sur-rounded by P–poor rims of garnet (Figs. 5b–5e). Bounda-ries between the P–rich patches and the P–poor rims aresharp and irregularly–shaped (Figs. 5c, 5e, and 5f). TheP–rich patches include kyanite, whereas the P–poor rimsinclude both kyanite and sillimanite (Figs. 5e and 5g).The grain size of rutile enclosed in garnet is almost thesame (~ 100–200 µm) regardless of the P zoning of gar-net. The intervening matrix within the cm–sized garnetporphyroblasts consists of plagioclase (An33–39), K–feld-spar, and quartz (Fig. 5b).

Rutile enclosed in the P–rich patches, enclosed inthe P–poor rims, and present in the matrix shows averageZr content of ~ 2950 ± 760 ppm (14 points on 12 rutilegrains) (Fig. 5e), ~ 2080 ± 140 ppm (7 points on 1 rutilegrain) (Fig. 5f), and ~ 2780 ± 60 ppm (5 points on 2rutile grains), respectively (Tables 1 and S1).

Garnet–sillimanite gneiss from Rundvågshetta (sam-ple TK2003010307)

This sample was collected from the northern part of Run-dvågshetta. The matrix consists mainly of sillimanite, gar-

Pressure–temperature estimates of the Lützow–Holm Complex, Antarctica 271

Figure 3. A BSE image, X–ray elemental maps and a photomicrograph of garnet–bearing domain in a Sil–Bt–Grt gneiss from Akarui Point(sample TK2002122304). Four digit numbers represent Zr concentration in rutile (in ppm). (a) BSE image of garnet–bearing domain. (b)and (c) X–ray elemental maps of garnet–bearing domain in terms of Ca and P. (d) Photomicrograph of the boxed part of (a)–(c). Note that theP–rich rim of the garnet includes both kyanite and sillimanite. Plane polarized light (PPL). (e) Enlargement of the boxed part of (c), showingthe gradual outward increase of P content in the core, that is discontinuously bounded by an irregularly–shaped sharp boundary with theP–rich rim where kyanite and sillimanite are enclosed.

K. Suzuki and T. Kawakami272

Tab

le1.

Averagedcompo

sitio

nsof

rutilein

metapeliticsamples

from

theselected

areasof

theLHC

Pressure–temperature estimates of the Lützow–Holm Complex, Antarctica 273

net (~ 5 mm), quartz, K–feldspar, and plagioclase (An38–41)with minor rutile (~ 100–500 µm), zircon, ilmenite, andbiotite (4.69–5.35 wt% TiO2). Garnet grains are mostlyhomogeneous in terms of major elements (Alm56.2–58.4

Prp36.2–39.0Grs3.5–6.5Sps0.4–1.2), whereas they preserve sharpzoning in terms of P. The garnet grains can be divided intoP–poor cores and P–rich rims (Fig. 6b). The P–poor coresof garnet enclose quartz, plagioclase (An38–41), nanogran-ite inclusions (NIs; Cesare et al., 2009), rutile (~ 10–100µm), apatite, zircon, biotite (4.65–5.60 wt% TiO2), andsillimanite (Figs. 6a–6c). The NIs are not arranged andexclusively included in the garnet cores (Fig. 6a). TheNIs mainly consist of Qtz + Pl (An20–26) + Kfs + Bt(3.37–4.18 wt% TiO2). Rare sillimanite in contact withquartz is enclosed in the garnet cores (Fig. 6c). The P–richrims of garnet are almost free of inclusion mineral exceptfor coarse–grained sillimanite up to ~ 500 µm (Figs. 6a and6b). Kyanite was absent as an inclusion mineral in garnetin this sample. Zirconium content of rutile enclosed in theP–poor cores of garnet was ~ 3990 ± 460 ppm on average(14 points on 10 rutile grains) (Tables 1 and S4), while thatof rutile present in the matrix was ~ 2630 ± 420 ppm (14

points on 5 rutile grains) (Figs. 6b–6d, Tables 1 and S4).

APPLICATION OF ZR–IN–RUTILEGEOTHRMOMETER

The Zr–in–rutile geothermometer (Tomkins et al., 2007)was applied to rutile grains in all samples describedabove in order to estimate the temperature conditions.

In the pelitic gneiss sample from Akarui Point, theP–rich rims of garnet include both kyanite and sillimanite(Figs. 3d and 3e). Therefore, the P–T condition can beestimated as the intersection of the kyanite/sillimanitetransition line (Pattison, 1992) with the result of Zr–in–rutile geothermometry (Tomkins et al., 2007) applied torutile inclusions separately enclosed in the P–rich rims ofgarnet. This gave 834 ± 4 °C at 10.5 kbar, representingthe P–T condition for the growth of P–rich rims (Fig. 2a).Selection of the Al2SiO5 diagrams (e.g., Holdaway, 1971;Bohlen, 1991; Pattison, 1992) did not give any significantdifference in P–T estimates. Since metastable kyanite cansurvive to the sillimanite stability field around the transi-tion line, the pressure condition for the growth of P–rich

Figure 4. A field photo of the sample locality and BSE images, X–ray elemental maps, and a photomicrograph of garnet–bearing domain fromSkarvsnes. The core/rim boundary of garnet is indicated by white broken lines. Four digit numbers represent Zr concentration in rutile (inppm). (a) Field occurrence of a Sil–Grt gneiss (sample TK2003012803) that fills interboudin partitions of a Grt–Bt gneiss. Hammer as ascale. (b) and (e) BSE images of garnet–bearing domain. (c) and (f ) X–ray elemental maps of garnet–bearing domain in terms of P. Note thatthe P–rich rims of garnet include both kyanite and sillimanite, whereas the P–poor cores of garnet include kyanite alone. (d) Photomicro-graph of the boxed part of (b) and (c). Rutile is included in the P–poor core of the garnet with kyanite. PPL.

K. Suzuki and T. Kawakami274

rims of garnet may extend to the lower pressure side to acertain extent (Fig. 2a). On the other hand, the pressurecondition could not be constrained in the case of P–poorcores of garnet since no Al2SiO5 minerals was enclosedthere. Applying the Zr–in–rutile geothermometer (Tom-kins et al., 2007) for the assumed pressure range of 0–14.0 kbar, rutile inclusions separately enclosed in the P–poor cores of garnet yielded 764 ± 19 °C/0 kbar to 844 ±21 °C/14.0 kbar (Fig. 2a).

Coarser–grained nature of matrix rutile compared tothe inclusion rutile indicates that matrix rutile rims atleast grew under the sillimanite stability field when the

matrix minerals recrystallized. Therefore, the temperatureestimate by the Zr–in–rutile geothermometry utilizing ma-trix rutile rims (773 ± 21 °C/1.0 kbar to 828 ± 22 °C/10.5kbar; Fig. 2a) is interpreted to represent the temperaturecondition of final equilibrium with zircon and quartz inthe matrix.

The P–T conditions for other localities can be esti-mated in the same way. In the pelitic gneiss sample fromSkarvsnes, the P–poor cores of garnet include kyanitealone (Figs. 4b and 4c). Therefore, by applying Zr–in–rutile geothermometry (Tomkins et al., 2007) to rutile in-clusions separately enclosed in the P–poor cores of gar-

Figure 5. A field photo of the studied sample, BSE images and X–ray elemental maps of garnet–bearing domain from Skallen. The outline ofgarnet is drawn in white solid line in (e) and (g). Four digit numbers in (d) and (g) represent Zr concentration in rutile (in ppm). (a) Fieldoccurrence of a Spl–Sil–Grt gneiss (sample TK2003011203) containing the garnet porphyroblast up to ~ 5 cm in diameter. (b) BSE image ofa cm–sized garnet porphyroblast. Kyanite and sillimanite crystals included in the garnet are shown as white triangle and circle, respectively.(c) X–ray elemental map of the cm–sized garnet porphyroblast in terms of P. (d) Enlargement of the BSE image of boxed part of (b). Rutile isincluded in the P–poor rim of the garnet as shown in (e). (e) Enlargement of the boxed part in (c). Note that the P–poor rim includes bothkyanite and sillimanite, whereas the P–rich patch includes kyanite alone. (f ) BSE image of one of the apparent ~ 5 mm–sized garnet grainsconstituting a cm–sized garnet porphyroblast. (g) X–ray elemental map of the same garnet grain shown in (f ) in terms of P. Note that all therutile grains are enclosed in the P–rich patch of the garnet. The P–rich patch also encloses kyanite, while the P–poor rim includes sillimanite.Color version is available online from https://doi.org/10.2465/jmps.190801.

Pressure–temperature estimates of the Lützow–Holm Complex, Antarctica 275

net, the P–T condition of >845 ± 28 °C at >10.6 kbar wasestimated (Fig. 2b). As for the P–rich rims of garnet, al-though kyanite and sillimanite do not coexist as inclusionminerals in a single garnet grain, adjacent garnet grainsseparately include kyanite and sillimanite in the P–richrims (Figs. 4e and 4f). Therefore, the P–rich rims of gar-net were interpreted to have formed at the P–T conditionunder which kyanite and sillimanite can coexist. The P–Tcondition of ~ 849 °C and 10.7 kbar was estimated as theintersection of the kyanite/sillimanite transition line withthe result of Zr–in–rutile geothermometry (Tomkins et al.,2007) applied to rutile enclosed in the P–rich rims of gar-

net. This P–T condition represents that for the growth ofP–rich rims (Fig. 2b). Applying the Zr–in–rutile geother-mometer (Tomkins et al., 2007), rutile rims in the matrixyielded 783 ± 38 °C/0.8 kbar to 841 ± 40 °C/10.7 kbarunder the sillimanite stability field (Fig. 2b).

In the pelitic gneiss sample from Skallen, the P–poorrims of garnet include both kyanite and sillimanite (Fig.5e). Therefore, the Zr–in–rutile geothermometry (Tom-kins et al., 2007) was applied to rutile inclusions sepa-rately enclosed in P–poor rims of garnet, and the P–Tcondition of 836 ± 8 °C/10.5 kbar was estimated as theintersection of the kyanite/sillimanite transition line with

Figure 6. BSE images and an X–ray elemental map of garnet–bearing domain in a Grt–Sil gneiss (sample TK2003010307) from Rundvåg-shetta. The outline of garnet and the core/rim boundary are drawn in white solid and broken lines, respectively. Four digit numbers representZr concentration in rutile (in ppm). (a) BSE image of garnet–bearing domain. (b) X–ray elemental map of the area shown in (a) in terms of P.(c) Enlargement of the boxed part in (a) and (b). Note that rutile is included in the P–poor core of garnet with sillimanite and NIs. (d) BSEimage of rutile in the matrix.

K. Suzuki and T. Kawakami276

the result of the Zr–in–rutile geothermometry. The resultrepresents the P–T condition for the growth of P–poorrims (Fig. 2c). On the other hand, P–T condition of thegrowth of P–rich patches of garnet was estimated by ap-plying the same geothermometer to rutile inclusions sep-arately enclosed in P–rich patches. The presence of kyan-ite inclusions in the P–rich patches (Fig. 5g) was utilizedfor the pressure constraint, yielding P–T estimate of >883± 36 °C and >11.5 kbar (Fig. 2c). Rutile grains in thematrix yielded 810 ± 3 °C/0.5 kbar to 875 ± 3 °C/11.4kbar under the sillimanite stability field (Fig. 2c) by ap-plying the Zr–in–rutile geothermometer (Tomkins et al.,2007). However, as occurrence of quartz is limited to theintervening matrix within the cm–sized garnet grain, it issuspicious whether rutile grains in the matrix were inequilibrium with quartz.

In the pelitic gneiss sample from Rundvågshetta,based on the occurrence of NIs and sillimanite inclusionsin the P–poor cores of garnet (Figs. 6a–6c), the P–T con-dition of the growth of P–poor cores can be constrained atleast to be higher than the vapor saturated solidus underthe sillimanite stability field. Applying the Zr–in–rutilegeothermometer (Tomkins et al., 2007) to rutile inclu-sions enclosed in P–poor cores of garnet yielded 850 ±15 °C/0.1 kbar to 927 ± 16 °C/12.5 kbar (Fig. 2d). On theother hand, applying the Zr–in–rutile geothermometer(Tomkins et al., 2007) to rutile in the matrix yielded804 ± 19 °C/0.5 kbar to 868 ± 20 °C/11.3 kbar underthe sillimanite stability field (Fig. 2d).

DISCUSSION

Metamorphic P–T conditions recorded as inclusionminerals in garnet

In the pelitic gneiss sample from Skallen, the cm–sizedgarnet porphyroblasts show characteristic P zoning, i.e.,the P–rich patches and the P–poor rims (Figs. 5c, 5e, and5g). The formation process of such a complicated P zon-ing is still unclear. However, this complicated texture issimilar to ‘P–rich patch’ reported by Kawakami et al.(2016) in that kyanite alone is included in the P–rich do-mains in garnet. Based on the difference of trace elementcomposition from surrounding garnet, and also based onthe presence of two metamorphic age populations in zir-con and monazite from the same sample, the ‘P–richpatch’ was interpreted as a relict garnet domain of thepre–existed kyanite–grade garnet possibly formed duringthe former kyanite–grade metamorphism (Kawakami etal., 2016). Taking into account the textural similarity ofthe P–rich patch with Kawakami et al. (2016) in that itincludes kyanite, and the fact that two populations of

metamorphic ages were detected by U–Th–Pb monazitedating of metapelitic gneisses from Skallen (Hokada andMotoyoshi, 2006), polymetamorphic origin of the studiedsample is possible. However, petrochronological study isfurther required to constrain the formation process of thecomplexly–zoned garnet of this study. Nevertheless, themetamorphic P–T conditions of the P–rich patches andthe P–poor rims of garnet are well constrained, whichare >883 ± 36 °C/>11.5 kbar and 836 ± 8 °C/10.5 kbar,respectively (Fig. 2c).

As reported above, the P–T conditions of the forma-tion of garnet rims in pelitic gneiss samples from AkaruiPoint, Skarvsnes, and Skallen are estimated as 834 ± 4°C/10.5 kbar, ~ 849 °C/10.7 kbar and 836 ± 8 °C/10.5kbar, respectively and are almost the same (Fig. 7). TheseP–T conditions are similar to that estimated for P–poorcores of garnet in a Grt–Sil gneiss from Skallevikshalsen(Kawakami et al., 2016). Because homogenization oforiginally zoned rutile in terms of Zr contents during pro-longed high–T metamorphism through diffusion processand re–equilibrium between the host garnet cannot beruled out, the metamorphic temperature conditions esti-mated from inclusion rutile should be considered as rep-resenting the minimum temperature attained. However, ifalmost the same temperature condition estimated fromAkarui Point, Skarvsnes, and Skallen were the result ofcomplete diffusional re–equilibrium between rutile andgarnet after entrapment in the host garnet, it is not con-sistent with the preservation of outstanding high–Zr con-centrations (i.e., high–T ) recorded in inclusion rutile en-closed in garnet from Rundvågshetta and in P–richpatches of garnet from Skallen (Fig. 7). Therefore, it ishighly probable that the estimated metamorphic P–T con-dition for garnet rim formation (~ 830–850 °C/~ 11 kbar)from Akarui Point, Skarvsnes, and Skallen represent theactual P–T conditions attained in each area.

Matrix rutile also yielded high–T conditions, whichare only slightly lower than that of inclusion rutile inAkarui Point, Skarvsnes, and Skallen samples. In the caseof Akarui Point, the temperature condition estimated byapplying the Zr–in–rutile geothermometer to the matrixrutile is higher than that estimated by the Fe–Mg ex-change type geothermometers for the same sample(Kawakami et al., 2008) (Fig. 2a). Combining the P–Tconditions estimated by inclusion minerals and matrixminerals, isothermal decompression paths starting fromthe kyanite stability field into the sillimanite stability fieldwere confirmed in these areas (Figs. 2a–2c).

Although these areas attained almost the same P–Tconditions corresponding to the granulite facies, AkaruiPoint has been traditionally classified as the transitionalzone based on matrix mineral assemblages of mafic– to

Pressure–temperature estimates of the Lützow–Holm Complex, Antarctica 277

intermediate–gneisses (Hiroi et al. 1991). In this study,the information of the metamorphic conditions attainedaround the kyanite/sillimanite transition boundary is onlypreserved in minerals enclosed in garnet, whereas the ma-trix minerals record the P–T conditions of post–peak de-compression stage in the case of pelitic gneisses. There-fore, the traditional metamorphic zone mapping bymatrix mineral assemblages of mafic– to intermediate–gneisses may not represent the highest metamorphic con-ditions attained in each area. In order to fully understandthe P–T–t evolution of the LHC, petrochronological stud-ies focusing on the inclusion minerals are further requiredthroughout the Complex.

ACKNOWLEDGMENTS

This study is a part of the science program of Japanese

Antarctic Research Expedition. We would like to sincere-ly thank members of JARE44 for the supports during thefieldwork. We are also grateful to T. Ikeda and an anon-ymous reviewer for constructive reviews, and M. Satish–Kumar for editorial efforts. This study was financiallysupported by Kyoto University research developmentprogram ISHIZUE, General Collaboration Project of theNational Institute of Polar Research (No. 28–25), andJSPS KAKENHI Grant Number JP19H01991 to T.K.

SUPPLEMENTARY MATERIALS

Color versions of Figures 2, 5, and 7 and SupplementaryTables S1–S5 are available online from https://doi.org/10.2465/jmps.190801.

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Manuscript received August 1, 2019Manuscript accepted November 7, 2019Published online December 28, 2019Manuscript handled by M. Satish–Kumar

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