九州大学学術情報リポジトリ
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Petrochronology of the high temperature metamorphic rocks in the Higo metamorphic complex and Kontum Massif
ブイ, ティ, シン, ブオン
https://doi.org/10.15017/4060251
出版情報：九州大学, 2019, 博士（理学）, 課程博士 バージョン：
権利関係：
Petrochronology of the high temperature metamorphic rocks in the Higo metamorphic complex and Kontum Massif
Graduate School of Integrated Sciences for Global Society Kyushu University
Bui Thi Sinh Vuong
January 2020 March, 2020
Contents
Abstract
1. Introduction………...1–11 2. Geological and geochronological outline………...12–25 3. Sample descriptions………....………26–62 3.1. Modes of occurrence………....………26–34 3.2. General petrography………....……… 34–62 4. Analytical methods………....……….………63–65
4.1. LA–ICP–MS U–Pb zircon dating …………....………63–64 4.2. FE–EPMA U–Th–Pb monazite dating………....……….……….... 64 4.3. REE geochemistry of zircon and garnet……....………...……….. 64–65 5. Geochronology………..………....……….………..66–157 5.1. LA–ICP–MS separated zircon geochronology………….……….66–137 5.2. LA–ICP–MS in-situ zircon geochronology..……….……..…………133–148 5.3. FE–EPMA monazite geochronology……....……….………..………148–157 6. X-ray elemental mapping for major and trace elements of garnet….………158–166 7. REE geochemistry……..………...………....……….………167–182 7.1. Zircon……..………...……….………167–176 7.2. Garnet……..………...……….………176–182 8. Discussion……..………….………...……….………183–220
8.1. Constraints on the timing of high-grade metamorphism ….………183–208 8.2. Zircon–monazite–garnet behavior during high temperature metamorphism……….209–220 9. Concluding remarks...…….………...……….………221–222
Acknowledgements..…….………...……….….………223–224 References……..………...….………...……….………225–237
Appendix 1.1. Collected samples and their mineral assemblages in the Higo metamorphic complex……..………….………...……...……….………238–239 Appendix 1.2. Collected samples and their mineral assemblages in the Kontum Massif..………...240–241
Appendix 2.1. LA–ICP–MS zircon U–Pb age data for analyzed samples in the Higo metamorphic complex.………...……...……….………242–264 Appendix 2.2. LA–ICP–MS zircon U–Pb age data for analyzed samples in the Kontum Massif………...……...……….………..265–293 Appendix 3. Chemical compositions and U–Th–Pb age of monazite for analyzed samples in the Kontum Massif………...……...……….………..294–303 Appendix 4.1. Major chemical composition of garnet in the Higo metamorphic complex..……… 304–305 Appendix 4.2. Major chemical composition of garnet in the Kontum massif..…………..………..306 Appendix 4.3. Major chemical composition of biotite in the Higo metamorphic complex..…………..……….307–308 Appendix 5.1. REE composition of zircon in the Higo metamorphic complex ……... 309 Appendix 5.2. REE composition of zircon in the Kontum Massif ………310–311 Appendix 5.3. REE composition of garnet in the Higo metamorphic complex ………….312 Appendix 5.4. REE composition of garnet in the Kontum Massif ……….…313
Abstract
The Higo metamorphic complex in Japan and the Kontum Massif in Vietnam are well- known high temperature (HT)–ultrahigh temperature (UHT) metamorphic areas in the world (e.g., Kelsey and Hand, 2015; Osanai et al., 2017). They play an important role in revealing tectonic evolution related with the large-scale growth of the Asian continent throughout a geological time (Osanai et al., 2008). Though the pressure–temperature (P–T) evolution and timing of the HT–UHT metamorphism in these areas were well reported (e.g., Hamamoto et al., 1999; Osanai et al., 2001a, 2001b, 2004, 2006; Maki et al., 2004; Nakano et al., 2007b, 2013), the timeline of the metamorphic evolution such as the timing of prograde, peak and retrograde stages is still obscured due to the lack of detail petrochronological studies. This study carried out U–Pb dating for both separated and in-situ zircon using laser ablation inductively coupled plasma mass spectrometry (LA–ICP–MS) and U–Th–Pb monazite dating by field-emission electron probe microanalyzer (FE–EPMA) for some representative high- grade metamorphic samples from these two mentioned study areas. Additionally, the X-ray chemical mapping and major element composition of garnet grains and rare earth elements (REE) geochemical analyses of zircon, monazite and garnet grains were also conducted. The combination the petrological context provided by garnet with geochronological ones from monazite and zircon allows to precisely constrain on the timings of the specific metamorphic stages for the HT to UHT metamorphism in the Higo metamorphic complex and Kontum Massif.
As the results, the timings of the HT–UHT metamorphism in the Higo metamorphic complex (ca. 105–120 Ma) and Kontum Massif (ca. 230–265 and ca. 430–450 Ma) are consistent with those of previous studies. However, compiling with previous P–T evolution, this study provided critical chronological information about the metamorphic evolution in each study area as follows. Firstly, in the Higo metamorphic complex the prograde metamorphism occurred at ca. 120 Ma immediately after precursor deposition at ca. 130 Ma and peak metamorphic conditions were attained during the period between 110 and 120 Ma. The ca. 110 Ma age might be indicative of the retrograde stage which promoted new zircon crystallization from melts and garnet breakdown. Finally, the hydration stage involved granitic intrusion into pelitic gneisses at ca. 105 Ma, which was characterized by the abundance of secondary chlorite, prehnite or even xenotime.
In case of the Kontum Massif, the results investigated that the pelitic granulites in the western Kannak Complex might have undergone the prograde M0 metamorphic stage at ca.
265 Ma, reached up to peak UHT condition of M1 stage between ca. 265 and 250 Ma by the
continental collision in the Trans Vietnam Orogenic Belt (e.g., Osanai et al., 2004). The subsequent decompression relating to the M2 stage during retrograde stage and granite emplacement occurred at ca. 250 Ma, then followed by the later intense hydration involved the appearance of andalusite, muscovite and chlorite at ca. 230 Ma during its exhumation at the upper crust level. On the other hand, two high-grade metamorphic events were recorded at Permian–Triassic and Ordovician–Silurian in the western Ngoc Linh Complex. Evaluating the equilibrium between zircon, monazite and garnet proposed the first prograde HT metamorphism at ca. 440–450 Ma, overprinted by the second one at ca. 240–250 Ma under the granulite-facies condition with following decompression on retrograde stage after ca. 240 Ma.
The decompression texture was indicated by the breakdown of garnet and sillimanite to form cordierite and spinel in the pelitic gneiss or garnet into orthopyroxene and plagioclase in the mafic granulite. Pelitic gneisses in the Kham Duc Complex recorded medium pressure metamorphism at ca. 230–240 Ma without any traces of early metamorphic event. There is no growth of zircon during this metamorphic condition. The timing of metamorphism at ca. 450 Ma was detected from two pelitic gneisses in the Dai Loc Complex, which is ca. 20 million years older than the magmatism event at ca. 420–430 Ma (Hieu et al., 2016).
Revealing the behavior of zircon–monazite–garnet during the HT to UHT metamorphism in these areas allows to verify the driving force or factor which control their behavior, and provides the critical information to understand the tectonic evolution. The interaction of internal and external fluid as well as melt plays an important role as the driving force for the formation or deduction of these three minerals. The most likely reasons to decide whether they can be formed or not during each metamorphic stage are the REE competition by the evolution (growth, breakdown or dissolution) and amount of other REE and Y-bearing minerals (xenotime and apatite), the local bulk rock composition as well as the formation of Zr-bearing phases such as ilmenite for the case of zircon.
1. Introduction
1.1. Significance of the high to ultrahigh temperature metamorphism
The metamorphic rock, accounting for seventeen percent of the rocks exposed on the earth surface (Blatt and Jones, 1975), is the witness for the tectonic processes in the crustal Earth. It records the pressure–temperature (P–T) affinity of the metamorphism that reflects the internal thermal structure in the crustal Earth in dependence upon the tectonic events (Brown, 2014). Metamorphic rocks have a range of P–T conditions. The P/T ratios define the geothermal gradient which contribute to classify three types of metamorphism correlating with the tectonic settings (Fig. 1.1). The low geothermal gradient of 5–10 °C km−1 caused by the subduction of oceanic lithosphere features the high P/T metamorphism of the blueschist- to eclogite-facies grade. The moderate (10–20 °C km−1) and high (20–50 °C km−1) gradients are led by the high radioactive heat production in thicken crust orogenesis or extra heating of magma upwelling, and ascertain the medium to low P/T metamorphism of amphibolite- to granulite-facies grade (e.g., Brown, 2007) (Fig. 1.1). Especially, producing the elevated gradient of 20–50 °C km−1 is required some hot heat source within the crust because the steady– state geotherm for continental crust is curved with average gradients of approximately 15 °C km−1 (Fig. 1.1). The occurrence of the ultrahigh temperature (UHT) metamorphic rocks, which experienced at >
900 °C and 7–13 kbar (e.g., Harley, 1998), indicates the specific tectonics to supply the extreme thermal condition such as the interaction between the crust and asthenospheric mantle (Harley, 2008 and reference therein). In the correlation between the peak metamorphic P–T conditions for all metamorphic belts with the well-constrained ages, the high temperature (HT)–UHT metamorphism is remarkably concomitant with the formation of four main supercraton (Superia) and supercontinent cycles (Nuna, Rodinia, and Gondwana) on the earth history (Brown, 2007) (Fig. 1.2). Thus, it is supposed to occur in the collisional plate boundaries and accretionary orogens (Kelsey and Hand, 2015). The HT to UHT metamorphic rocks, therefore, play the important role in revealing the principal understanding of the development and stabilization of continents through the collisional process and the historical variant of the earth crustal behavior (Harley et al., 2008; Kelsey, 2008; Osanai et al., 2008; Kelsey and Hand, 2015;
Harley, 2016). It has demanded the accurate and precise ascertainment of their metamorphic evolution via assessing the pressure–temperature–time (P–T–t) path. Thus, the main task of metamorphic petrology requires the comprehensive integration of geochronological information and qualitative or quantitative metamorphic conditions (timings of onset and peak
o
5 C km
-1
10
o
C km
-1
20
oCkm
-1
50
oCkm
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St ea dy -st at e co nt in en ta l geot
her m
Fig. 1.1. P–Tplot of metamorphic facies and typical steady–state continental geotherm (Fitzsimons,
0 2 4
200 6
8 10 12 14 16 18 20
400 600 800 1000
10 20 30 40 50 60 70
D ep th ( km )
Temperature (
oC)
Pr es sur e (kb ar )
.2. Plot of thermal gradient against age for high–grade metamorphic belts with low (squares, <10 °C km−1), moderate (diamonds, 10–20 °C km−1) and high (circles, °C km−1) gradients (Brown, 2007; Fitzsimons, 2016). The moderate and hot gradients belts match periods of supercontinent assembly (shaded age bands).
Supe ria
Nuna
Rodin a
Gond wan a
Pang ea
Ag e (M a)
05001000150020002500300010
20
30
40
1750 1500 1250 1000 750 500 250
50
T condition of the metamorphism or cooling and exhumation stage) for the modelling of the tectonic process and crustal evolution. Linking geochronological data to the P–T path information provides important constraint on timescales of the HT–UHT metamorphic processes and plausible tectonic settings (e.g., Engi, 2017).
1.2. Integration of metamorphic condition and timing of metamorphism
To obtain the P–T metamorphic conditions, several methods have been used such as conventional geothermobarometer, petrogenetic grid and pseudosection on the basis of the coexisting mineral assemblage and its mineral chemistry and/or bulk chemistry on each metamorphic stage. In this case, garnet, which is a common porphyroblast in the metamorphic rocks, makes a great contribution to extract the information of petrological context due to several following reasons: 1) its occurrence in diverse lithologies and P–T conditions, 2) its durability and capability to retain other index metamorphic minerals as inclusions during the garnet growth as well as the chemical zoning (Fe, Mg, Ca, Mn, P and Y) responding to the metamorphic condition and 3) its sensitivity to the metamorphic reaction involving the reaction texture at its rim as the indicator for the retrograde path (e.g., Baxter et al., 2017). Therefore, the detail P–T information is often obtained from garnet-bearing metamorphic assemblages.
Additionally, garnet is also regarded as the most important major mineral sink for the heavy rare earth elements (HREE) and yttrium in the metamorphic rocks (e.g., Bea et al., 1994). The growth and breakdown of garnet play an important role in controlling the HREE and Y budgets for the geneses of accessory mineral chronometers of zircon, monazite and xenotime (Rubatto and Hermann, 2007; Taylor et al., 2015).
Geochronologically, several isotopic dating methods have been applied to the metamorphic rocks to determine the timing of metamorphism depending on the isotopic characteristics and metamorphic conditions. For examples, the K–Ar, Rb–Sr methods on muscovite, biotite or amphibole are often useful for the low-grade metamorphism due to their low closure temperature and ease for the isotopic resetting under the low-grade metamorphic condition (e.g., Dodson, 1973; Villa, 2015). On the other hand, the U–Th–Pb chronology for zircon and monazite has been concerned as the most convinced age determination method for the HT–UHT metamorphic rocks due to their chemical and physical resistance and high closure temperature equivalent to the granulite-facies condition (e.g., Harley et al., 2007; Engi, 2017;
Rubatto, 2017). Moreover, in the case of metamorphic rocks, their stubbornness enables to hold
the identifiable growth zoning that consists of the relic protolithic domain (detrital or igneous grain) and (re)crystallized metamorphic domain. Through the spot isotopic analyses, the ages related with protolith formation and metamorphism can be provided, which is the advantage of this method to obtain precisely the timing of the metamorphism. Especially, the clustered metamorphic ages obtained from the zircon and/or monazite in the HT to UHT metamorphic rocks have often been regarded as the timing of the peak metamorphism because of their high closure temperature and robustness (e.g., Harley et al., 2007, 2016). However, it is not always satisfied to indicate the age of the peak metamorphism since these two minerals can be formed or grown at the different portion of P–T evolution during the metamorphism, such as prograde, peak or retrograde stages (e.g., Hoskin and Black, 2000; Rubatto, 2002, 2017; Rubatto and Hermann, 2003). Some mechanisms for the geneses of zircon and monazite under the various conditions have been suggested: the dissolution-recrystallization from the pre-existing zircon and monazite grains (Nemchin and Bodorkos, 2000; Nemchin et al., 2001; Kawakami et al., 2013), solid-state reaction by the breakdown of garnet, ilmenite, rutile, apatite or allanite (Degeling et al., 2001; Rubatto et al., 2001; Wing et al., 2003), crystallization from anatectic melt (Hokada and Harley, 2004; Harley et al., 2007; Rubatto, 2017) or precipitation from fluids (e.g., Schaltegger, 2007). In particular for the metamorphic zircon, Harley et al. (2007) proposed its genesis during the HT metamorphism using “bracketing” approach (Fig. 1.3), as follows: the dissolution and later precipitation of zircon during the partial melting on the prograde stage; the chemical reactions involving zircon and other accessory minerals (e.g., monazite, xenotime, apatite, rutile, ilmenite) during the growth of major minerals such as garnet; the thermal peak metamorphism; the crystallization of zircon from the anatectic melt or by the solid-state mineral reactions at the post-peak stage (Fig. 1.3). Therefore, understanding the behavior or response of zircon during the metamorphism gives an important contribution to decipher the metamorphic history of the rocks. However, the responses of monazite and zircon against the (re)crystallization/modification during the HT to UHT metamorphism are slightly different. Zircon is an extremely refractory mineral that can retain the older protolith isotopic composition, but it is not always a reliable recorder for the timing of lower T metamorphism due to the less reactive ability (e.g., Rubatto, 2017). Though monazite has less ability to retain the older protolith, it may have greater potential to be recrystallized even under the lower T condition via the interaction with fluid or remnant melt due to its sensitive reactivity (e.g., Rubatto et al., 2013; Yakymchuk et al., 2017; Skrzypek et al., 2018). Therefore, the combination of both zircon and monazite geochronology can provide the insight into the
“Bracketing”approachappliedformetamorphicepisodeevolutionduringhightemperaturemetamorphism(Harleyetal.,2007).Duringheatingrocksmayundergo lting,causingdissolutionoffinerpre-existingzircon(A)followedbylaterprecipitationofnewzircononsurvivorgrains(B).Crystallizationofhigh-Tmeltsand entgrowthofzirconwillgenerallyoccuronthepost-peakcoolingpathatdifferentintervalsdependingonthewatercontentofthemelt(reflectingaH2O)or withwallrocks(C,D).NewzirconmayalsogrowasaresultofZr-liberatingreactions,forexamplegarnetbreakdownorrutiledecomposition(E),dependingon path.The‘residual’meltscancontributetocrystallizesomenewzircon(F),andliberatedfluidsmaycauseselectiverecrystallizationofexistingzircon(G).
temporal evolution of the HT to UHT metamorphism rather than only zircon or monazite one (e.g., Kelsey et al., 2008).
Consequently, the accurate integration of metamorphic condition and its timing requires the reasonable evaluation of the co-existence between the petrological and geochronological key minerals of garnet–zircon–monazite.
1.3. Petrochronology for the high to ultrahigh temperature metamorphic rocks
The petrochronology, which is simple combination of petrology and geochronology, has developed, of which main topic is to put the chronometer dates provided by monazite and zircon into the petrological context extracted from garnet (Fraser et al., 1997). For many UHT terranes/localities, the petrological framework and P–T evolution have been adequately revealed as well as the vast increase in geochronological data sets over the past decade, which provide the great basement to unravel their petrochronological characteristics. Indeed, the combination of zircon and monazite geochronology coupled with their trace element (i.e. REE) partitioning/distribution data to determine whether the equilibrium existed between zircon (or monazite) and major minerals (e.g., Rubatto, 2002; Whitehouse and Platt, 2003; Harley et al., 2007; Rubatto et al., 2013), and phase diagrams (Yakymchuk and Brown, 2014), currently contributes the thorough approach to decipher the petrochronology. Therefore, understanding the behavior between zircon, monazite and garnet is the reliable and effective way to link geochronological data to metamorphic condition. Regarding to zircon and garnet, the REE partitioning between these two minerals (Dzircon/garnet) has been applied to connect zircon age with metamorphic condition through the assessment of their equilibrium (e.g., Harley, 2002;
Rubatto, 2002, 2017; Whitehouse and Platt, 2003; Hokada and Harley, 2004; Taylor et al., 2015). As both minerals prefer HREE relative over light to middle REE (L–MREE), the
HREEDzircon/garnet values are close to unity for Gd to Lu at the high temperature and Ca-poor garnet (Harley, 2002; Whitehouse and Platt, 2003; Hokada and Harley, 2004; Taylor et al., 2015). The yttrium concentration in monazite is about 1.5 orders of magnitude higher than that in garnet (Bea et al., 1994). However, in most cases, garnet is substantially more modally abundant than monazite and hence, the breakdown or growth of garnet exerts first-order controls on the Y and HREE budget of the rocks. Monazite and zircon crystallization during garnet growth result in relatively low contents of Y and HREE in these minerals, while the garnet breakdown during accessory mineral formation can lead to the Y- and HREE-enriched concentrations of the new growth for these minerals, respectively (e.g., Engi, 2017). Though the above statement for the
trace elements equilibria of zircon–monazite–garnet is a fundamental tool in connecting the P–
T and time information, the occurrences, external and internal textures are also necessary to precisely interpret the significance of zircon and monazite U–Th–Pb ages. The zircon grains separated from the crushed samples are conventionally dated to obtain the abundant dates easily and quickly, but it lacks of the relation between textural occurrence and its age. Therefore, the in-situ zircon and monazite geochemical and chronological analyses from the thin section are more powerful strategy to reveal the timeline of the HT–UHT metamorphic process through the direct correlation of their occurrence, morphology, internal texture, and U–Pb age as well as intact textural relationship, trace element analyses (e.g., Y, REE).
Recently, by approaching the behavior of zircon–monazite–garnet during the HT–UHT metamorphism, the precise metamorphic evolution has been constrained for several HT–UHT metamorphic areas, such as Anápolis-Itauçu Complex, Brazil (Moraes et al., 2002; Baldwin et al., 2007; Baldwin and Brown, 2008); western Betic Cordillera, Southern Spain (Whitehouse and Platt, 2003); Napier Complex, East Antarctica (Hokada and Harley, 2004); Rauer Island, East Antarctica (Hokada et al., 2016). However, the Higo metamorphic complex in Japan and the Kontum Massif in Vietnam, are well-known HT–UHT metamorphic areas in the world (Kelsey and Hand, 2015; Fig. 1.4), have not been applied for the detail petrochronological analyses yet. In these two areas, the metamorphic evolution in terms of petrography has been well unraveled, of which the prograde and retrograde metamorphic processes for each metamorphic zone were reported in detail on the basis of reaction textures, chemical zoning of minerals, and mineral inclusions in porphyroblasts. The extensive geochronological studies provides the adequately comprehended information about timing of the HT–UHT metamorphism in these areas. The Higo metamorphic complex consists of the high-grade metamorphic rocks up to the UHT granulite-facies that are characterized by a clockwise P–T evolution and peak metamorphic conditions of 7.8–9.0 kbar and 900–960 oC (Osanai et al., 1998). Previously, this complex was proposed to have metamorphosed at ca. 230–260 Ma (Osanai et al., 1998, 2006; Hamamoto et al., 1999), however, the recent zircon geochronology has suggested that the timing of the HT–UHT metamorphism is approximately 110–120 Ma due to the clustered age data (Sakashima et al., 2003; Takagi and Arai, 2003; Dunkley et al., 2008; Maki et al., 2014; Osanai et al., 2017; Suga et al., 2017). On the other hand, the Kontum Massif, which is a member of the Trans Vietnam Orogenic Belt (TVOB) (Osanai et al., 2008), has been known to have a complicated metamorphic history which composed of various metamorphic rocks from greenschist- to granulite-facies (e.g., Osanai et al., 2004; Nakano et
1.4.DistributionoftheUHTmetamorphicterranesfromKelseyandHand(2015)andOsanaieral.(2017).TerranesthathavebeeninvestigatedbyJapanese archersareshownannotatedwiththeterranenames.Inthisstudy,thestudyareasoftheHigometamorphiccomplexandKontumMassifarehighlightedbythe wboxeswithredletters.
KontumMassifHigo metamorphic complex
Massif basically experienced the clockwise P–T–t paths and can be recognized as three metamorphic events of M0, M1, and M2. The M0 metamorphism is led by the continent–
continent collision and is characterized by high- to ultrahigh-pressure conditions. The metamorphic rocks relating to M1 and M2 events experienced the UHT granulite and the granulite-facies metamorphism, respectively (Nakano et al., 2004; Osanai et al., 2004, 2008).
The age of the metamorphism, however, is still controversy because the metamorphic ages of the Kontum Massif indicate two peaks, the Ordovician to Silurian and the Late Permian to Early Triassic (Carter et al., 2001; Nagy et al., 2001; Nam et al., 2001; Lan et al., 2003; Owada et al., 2006; Roger et al., 2007; Osanai et al., 2008; Usuki et al., 2009; Nakano et al., 2013; Faure et al., 2018). This inconsistency is mainly due to the dependence on which metamorphic ages were detected by the single dating method such as zircon U–Pb or monazite U–Th–Pb dating, although metamorphic rocks with various metamorphic conditions occur in the Kontum Massif.
Therefore, to reveal precisely the timing of the amphibolite- to granulite-facies metamorphism in the Kontum Massif, combining reliable zircon and monazite U–Th–Pb dating methods is required. In addition, the detail study concerning the zircon–monazite–garnet behaviors during the HT–UHT metamorphism in two study areas allows to precisely constrain on the timings of the specific metamorphic stages and reveal the comprehensive history of metamorphic evolution. Consequently, the results of this study can contribute to the development of the petrochronology by the feedback of the behavior of zircon–monazite–garnet during the HT–
UHT metamorphism in the Higo metamorphic complex and Kontum Massif as the cases.
Therefore, in this study, the following purposes will be considered:
1) Confirming the timing of the HT–UHT metamorphism in the Higo metamorphic complex and Kontum Massif by combining both monazite and zircon age dating.
2) Correlating the zircon and monazite ages to the petrological context provided by garnet to constrain the metamorphic evolution and re-establish the P–T–t paths with integrating previous reports for each study area.
3) Revealing the general behavior of zircon–monazite–garnet during the HT–UHT metamorphism in both study areas.
For obtaining these expected results above, this study carried out the U–Pb separated zircon dating for some 14 representative samples from the Higo metamorphic complex and 15 ones from the Kontum Massif using laser ablation inductively coupled plasma mass spectrometry (LA–ICP–MS). The in-situ method of the zircon U–Pb dating by LA–ICP–MS is also applied for one pelitic gneiss from the Higo metamorphic complex and three pelitic
by field-emission electron probe microanalyzer (FE–EPMA) is also conducted for five metamorphic rocks in the Kontum Massif. The X-ray chemical mapping and major element composition of garnet grains and REE geochemical analyses of zircon, monazite and garnet grains from some representative samples in these two study areas are also carried out.
Considering the REE geochemistry behaviors between garnet and the accessory chronometer minerals of zircon and monazite, it allows to link the individual ages to the metamorphic stage of these analyzed metamorphic rocks, in order to constrain on the metamorphic evolution for the Higo metamorphic complex and the Kontum Massif. Minerals abbreviation in this study follows the works of Whitney and Evans (2010).
2. Geological and geochronological outline
2.1. Higo metamorphic complex
The Higo terrane located near Kumamoto city in the central-western part of Kyushu island, southwest Japan (Fig. 2.1). It extends for about 40 km in length with a maximum width of 20 km (Fig. 2.1) (e.g., Osanai et al., 1998). From the north to south, this terrane is divided into four major units (Yamamoto, 1962): the Manotani metamorphic complex, Higo metamorphic complex, Higo plutonic complex, and Ryuhozan metamorphic complex (Fig.
2.1). The boundary between the Manotani and Higo metamorphic complexes is still under controversial, which is supposed to be separated by ENE-trending high-angle faults or a Ndipping low-angle fault (Miyazaki, 2004), or a gradual change with no faults between them (Okamoto et al., 1989), or a thrust fault accompanied by serpentinite intrusion (Osanai et al., 2001a). On the other hand, the Higo metamorphic complex borders with the Ryuhozan metamorphic complex via the intrusions of the Miyanohara tonalite and the Shiraishino granodiorite (Higo plutonic complex) (Osanai et al., 2001a) (Fig. 2.1).
The mafic greenschists and lawsonite–pumpellyite pelitic schists with subordinate siliceous and ultramafic schists mainly compose the Manotani metamorphic complex (Karakida et al., 1989) (Fig. 2.1). The presence of some diagnostic minerals such as lawsonite in pelitic schists and sodic amphibole (riebeckite–ferro-glaucophane) and the pumpellyite + actinolite assemblage in mafic schists indicates the high pressure, low temperature metamorphism in this complex (Karakida et al., 1989). Karakida et al. (1989) suggested the prograde history with increasing temperature and decreasing pressure for the mafic and pelitic metamorphic rocks in this complex based on relict minerals of sodic amphibole (crossite and riebeckite), pumpellyite and lawsonite remaining in some porphyroblastic minerals. The southern margin of the Manotani metamorphic complex is overprinted by the Early Cretaceous low P/T metamorphism (Higo metamorphism) which contributed to form biotite in pelitic schists (Karakida et al., 1989; Okamoto et al., 1989; Nagakawa et al., 1997).
The Higo plutonic complex, which intruded into the Higo and Ryuhozan metamorphic complexes (Miyazaki, 2004), consists of granodiorite (the Shiraishino granodiorite) and tonalite (the Miyanohara, Manzaka, and Joyama tonalites) (Fig. 2.1). The boundary of these two types of granitic rocks (the Shiraishino granodiorite and Miyanohara tonalite) seems to be gradual change but in some places, the granodiorite crosscuts the tonalite (Yamamoto, 1962).
The southern margin of the Miyanohara tonalite has gneissic foliation (Miyazaki, 2004).
Sampling points
Impure Marble Impure marble
430T01A,B, C 43003A 430T03 43007B
430T04C 50201A, C
42901B–E, G Btschist Btgneiss Grt–Btgneiss Grt–Sil–Btgneiss Grt–Crd–Btgneiss Hbl–Btschist Hbl–Btgneiss Opx–Grt–Btgneiss Grt–Crd–Bttonalite Grt–Bttonalite Amphibolite Serpentinite Metamorphosed peridotitePelitic, psammitic and tuffaceousschist
DoleriteChl–Msschist
QtzschistHbl–Act–Chlschist
Ep–Act–ChlschistEp–ChlschistEp–Act–Pmpschist
Chert Manzakatonalite Shiraishinogranodiorite Hblgabbro
Mitanoharatonalite
Alluvium Asowelded–tuff Diluvium MifuneGroup Cretaceous Sediments MizukoshiFormation Higo plutonic complex
Higo metamorphic complex
Manotani metamorphic complex Ryuhozan metamorphic complex
Shiraishino
Tomochi Mt. Manzaka
Mt. Kamakura
MiyanooTabata
Mt. KosaMt. Manotani Mt. Suishou Uchida Ogawa Miyahara Mt. Ryuho
Mt. Tori
After Osanai et al. (2004) Fault
Hi g o t e rr a n e
5 km Fig. 2.1. Geological map of the Higo terrane (modified after Osanaiet al., 2006). The location of analyzed samples are indicated by the blue points with the sample numbers presented in the yellow box.
After Osanaiet al. (2006)
intrusion and solidification of the Higo plutonic complex was Early Cretaceous (100–121 Ma) (Nakajima et al., 1995). In addition, the zircon U–Pb ages of the Miyanohara and Manzaka tonalities have been defined by SHRIMP analyses at 110–111 and 113 Ma, respectively (Sakashima et al., 2003).
The Ryuhozan metamorphic complex consists of mafic–felsic schists, pelitic and psammitic schists, impure marble (Sakashima et al., 2003) (Fig. 2.1). This metamorphic complex is characterized by the low pressure and temperature metamorphism and some rocks are strongly mylonitized (Yamamoto, 1962). The intrusion of the Higo plutonic complex produced cordierite- and andalusite-bearing hornfelses in this complex (Miyazaki, 2004).
Dating from the metamorphic rocks gave the hornblende K–Ar ages of 80–113 Ma (Sakashima et al., 1999).
The Higo metamorphic complex, of which high-grade metamorphic rocks are the main objects in this study, is an imbricated crustal section increasing metamorphic grade from north (greenschist-facies) to south (granulite-facies) (Osanai et al., 1998). Pelitic to psammitic metamorphic rocks are predominate throughout the sequence with lesser amounts of calcareous metasediments, impure marble, and mafic to intermediate metaigneous rocks and orthogneisses (Osanai et al., 1998) (Fig. 2.1). This metamorphic complex also contains abundant migmatites which associated with leucogranitic dykes and pods throughout the complex and the Miyanohara tonalite at the south (Obata et al., 1994). The petrological characteristics of the complex have been well studied by many authors (Yamamoto, 1962; Karakida et al., 1989;
Okamoto et al., 1989; Obata et al., 1994; Osanai et al., 1998, 2006, 2017; Hamamoto et al., 1999; Maki et al., 2004; Miyazaki, 2004; Dunkley et al., 2008). The metamorphic zonation was defined based on the mineral assemblages of metapelites in the Higo and Manotani metamorphic complexes, which divides the complexes into five zones: zone A by chlorite + muscovite, zone B by biotite + muscovite + andalusite, zone C by K-feldspar + sillimanite + biotite, zone D by garnet + cordierite + biotite, zone E by the presence of orthopyroxene and clinopyroxene in mafic gneisses (Obata et al., 1994) (Fig. 2.2). In a few areas, it contains serpentinite with blocks of peridotite and sapphirine-bearing granulites (Osanai et al., 1998).
Osanai et al. (2001a, 2006) proposed the new zone as zone F on the basis of the presence of sapphirine and corundum-bearing highly aluminous granulite. The granulites only occur as blocks in meta-peridotites (serpentinite) and these blocks have been regarded as the deeper portion of the crustal level than zone E (Osanai et al., 1998) (Fig. 2.2). The zone A corresponds to the Manotani metamorphic complex and zones B to F accord to the Higo metamorphic
Fig.2.2.MetamorphiczonationoftheManotaniandHigometamorphiccomplexesdefinedbythemineralassemblagesofthemetapelites(Obataetal.,1994;Osanai etal.,2006):ZoneAbychlorite+muscovite,zoneBbybiotite+muscovite+andalusite,zoneCbyK–feldspar+sillimanite+biotite,zoneDbygarnet+cordierite+ biotite,zoneEbyorthopyroxene+garnet,andzoneFbysapphirine+corundum+spinel+cordierite.Thebluepointsindicatethesamplelocalitiesinthisstudy.
Zo ne F
Zone A–I Zone A–II, III Zone B Zone C
Zone D Zone E Impure marble Granite intrusive
2 km Alluvium MifuneGroup MizukoshiFm.
Matsubase Asowelded tuff
Shiraishino
Kosa
Mt. Kosa
Mt. Manotani 430T01A, B, C43003A 430T03 43007B
430T04C 50201A, C
Samplingpoints
42901B-E, G
Text
23–24 km and consists mainly of garnet–cordierite–biotite gneiss, garnet–orthopyroxene gneiss, orthopyroxene-bearing amphibolite, and orthopyroxene-bearing S-type tonalite (Osanai et al., 2006). Osanai et al. (2006) investigated the progressive changes in metamorphic grade from zone A to F via the following reactions:
Zone A to B: phengite + chlorite = muscovite + biotite + quartz + vapor, Zone B to C: muscovite + quartz = sillimanite + K-feldspar + vapor, Zone C to D: biotite + sillimanite + quartz = garnet + cordierite + K-feldspar + vapor, Zone D to E: biotite + quartz = orthopyroxene + garnet + melt (or K-feldspar + vapor), Zone E to F: fluorine-rich phlogopite + sillimanite + quartz + garnet + plagioclase = sapphirine + corundum + cordierite + melt.
The metamorphic grade for each zone corresponds to greenschist-facies in zone A, amphibolite-facies in zones B and C, and granulite-facies in zones D, E and F (Obata et al., 1994; Osanai et al., 1998, 2006). The P–T conditions for pelitic metamorphic rocks from zones C, D, E and F were estimated by using geothermobarometers and the petrogenetic grid of the KFMASH system (e.g., Spear and Cheney, 1989) (Fig. 2.3). The results pointed out peak P–
T conditions for each zone, which are 3.5–4.2 kbar, 640–720 ºC for zone C, 4.8–6.0 kbar, 740–
820 ºC for zone D, 6.0–7.2 kbar, 800–870 ºC for zone E and 7.8–9.0 kbar, 900–960 ºC for zone F (Osanai et al., 1998, 2001a, 2006) (Fig. 2.3). The progressive change in metamorphic grade from upper to lower crustal levels during the specific period is indicated by metamorphic field gradient from the amphibolite-facies zone C to UHT granulite-facies zone F (Osanai et al., 2006) (Fig. 2.3). In addition, Osanai et al. (2001a) identified the prograde and retrograde metamorphic processes in zone D by the detailed study of the reaction textures, chemical zoning and mineral inclusions in garnet porphyroblasts in garnet–cordierite–sillimanite–biotite gneiss. Based on the low-XMg (0.12), high-XCa (0.10) garnet core that includes some inclusions of staurolite, tourmaline, ilmenite, rutile, quartz, apatite and sillimanite; and high-XMg (0.24) and low-XCa (0.02) garnet rim which includes only fibrolite inclusions, they stated that the garnet porphyroblast has grown during temperature increasing and pressure decreasing prograde processes by the staurolite consuming reactions: staurolite + quartz = garnet + sillimanite (or kyanite) + vapor or staurolite = garnet + biotite + sillimanite (or kyanite). At the decompression and cooling process, cordierite–spinel symplectite surrounding the garnet porphyroblasts was formed as a retrograde product through the reaction of garnet + sillimanite (fibrolite) = cordierite + spinel (Osanai et al., 2001a). The partial melting is recognized by the cordierite melt pod in the hand specimen, and also by the occurrence of inclusion-free cordierite
Retrograde path
Progradepath Prograde path (inferred) Progressivemetamorphic fieldgradient Fig.2.3.Pressure–temperaturepathoftheHigometamorphiccomplex(Osanaietal.,2006andreferencestherein).Peakmetamorphic conditionsofeachzoneareestimatedbygeothermobarometersandapetrogeneticgridforpeliticmetamorphicrocks.
1200
by sub-solidus metamorphic reactions (Osanai et al., 2006). For the timing of the metamorphism, many radiometric ages have been reported from the Higo metamorphic complex (e.g., Nagakawa et al., 1997; Osanai et al., 1998, 2017; Suzuki et al., 1998; Hamamoto et al., 1999; Dunkley and Suzuki, 2001; Sakashima et al., 2003; Suga et al., 2017). The timing of the metamorphism has been considered to occur during Permian–Triassic by K–Ar, Rb–Sr, Sm–Nd (mineral isochron), CHIME and zircon U–Pb methods (e.g., Osanai et al., 1998, 2006;
Suzuki et al., 1998; Hamamoto et al., 1999), or Cretaceous (100–117 Ma) through Rb–Sr and Sm–Nd mineral-whole rock isochron, U–Th–Pb monazite and U–Pb zircon dating (e.g., Sakashima et al., 2003; Takagi and Arai, 2003; Dunkley et al., 2008; Maki et al., 2014).
Recently, Osanai et al. (2017) and Suga et al. (2017) have acceded the Cretaceous age (ca.
110–120 Ma) as timing of HT–UHT metamorphism and migmatization with the youngest detrital zircon age of ca. 170 Ma.
2.2. Kontum Massif
In the northcentral Vietnam, the developed northwest to southeast trending shear zones divide the geological framework of Vietnam into several units (Lepvrier et al., 1997), in which metamorphic rocks distribute mainly in the Red River shear zone, Song Ma suture zone, Truong Son fold belt and Kontum Massif (Fig. 2.4). The Trans Vietnam Orogenic Belt (TVOB) defined by Osanai et al. (2008) is situated between the South China and Indochina blocks as the result of their collision. The belt includes several high-grade and highly deformed metamorphic provinces along the Red River shear zone, Song Ma suture zone, Tam Ky-Phuoc Son shear zone, and Dak To Kan shear zone (Figs. 2.4 and 2.5a). The Kontum Massif is located at the southern part of the TVOB, central Vietnam, and extends partly into western Vietnam and southeast Laos (Hutchison, 1989) (Figs. 2.4 and 2.5a). It has been divided into three complexes on the basis of metamorphic grade: the Kannak Complex in the southeast (granulite- facies); the Ngoc Linh Complex in the center (amphibolite- to granulite-facies) and the Kham Duc Complex in the north and southwest (greenschist- to amphibolite-facies) (United Nations, 1990) (Fig. 2.5b). The Dai Loc Complex is also recognized in the northern Kontum Massif, further to the north of Kham Duc Complex (Nakano et al., 2013; Hieu et al., 2016) (Fig. 2.5b).
Numerous scientists have investigated the metamorphic perspective of the Kontum Massif (Carter et al., 2001; Nagy et al., 2001; Lan et al., 2003; Lepvrier et al., 2004; Osanai et al., 2004, 2008; Maluski et al., 2005; Owada et al., 2006; Nakano et al., 2007a, 2007b, 2008, 2009, 2010a, 2010b, 2013; Roger et al., 2007; Usuki et al., 2009; Sanematsu et al., 2011), of which
Fig. 2.4. Location of the Kontum Massif within Southeastern Asia. After Osanai et al. (2008).
Sibum asu Block
1: Red River shear zone 2: Song Ma suture zone 3: Danang-Dai Loc shear zone 4: Tam Ky-Phuoc Son suture zone 5: Kham Duc shear zone 6: Dak To Kan shear zone 7: Po Ko suture zone
200 km 73004A3, D, E
72902E
72903G
72801B, C, D T01A, T01B 73004A3, D, E
72903G 72902E
72901E
72802C-E ca. 230-260 Ma
ca. 430-460 Ma Both ages Metamorphic ages
Sampling points Osanai et al. (2008)
Sibum asu Block
Kham Duc Osanai et al. (2008)
(a)
Sibumasu Block
(b)
73102E, I
Migmatite Complex Dai Loc Complex
Fig. 2.5. Location of the Kontum Massif in Vietnam (a) after Osanai et al. (2008) and distribution of metamorphic rocks corresponds to the Kannak, Ngoc Linh, Kham Duc and Dai Loc Complexes as well as intrusive rocks in the Kontum massif (b). Modified after United Nations (1990). Metamorphic ages reported in previous studies are shown in (b), which is compiled from Nagy et al. (2001); Nam et al., (2001); Leprvier et al. (2004); Osanai et al. (2004);
Nakano et al. (2007a, 2013); Roger et al. (2007); Usuki et al. (2009); Tran et al. (2014). Locations of samples analyzed in this study are also shown as yellow spots.
2013; Faure et al., 2018). Osanai et al. (2008) presented the clockwise P–T–t paths for the metamorphic rocks from the Kon Tum Massif, which can be recognized as three metamorphic events; M0, M1, and M2 (Fig. 2.6). The M0 metamorphism is induced by the continent–
continent collision, and is featured high- to ultrahigh-pressure conditions. The metamorphic rocks from M1 and M2 underwent the ultrahigh-temperature granulite and the granulite facies metamorphism, respectively (Nakano et al., 2004; Osanai et al., 2004, 2008). Particularly, Nakano et al. (2013) summarized the metamorphic evolution of the large-scale metamorphic rocks on the basis of their occurrences, microstructures, metamorphic pressure–temperature conditions, and monazite chemical dating. Considering metamorphic textures and ages, they concluded that the Late Permian to Early Triassic eclogite- to granulite-facies event was triggered by the continental collision, whereas the Ordovician to Silurian thermal event occurred in an active continental margin. The metamorphic ages reported from the Kontum Massif are compiled in figure 2.5b.
2.2.1. Kannak Complex
The Kannak Complex is mainly composed by HT to UHT granulite-facies and minor amphibolite-facies metamorphic rocks which show the clockwise P–T path (e.g., Osanai et al., 2001b, 2004, 2008) (Fig. 2.6). The pelitic rocks are predominant in the Kannak Complex, such as garnet–orthopyroxene–sillimanite–cordierite granulite, garnet–cordierite–sillimanite–
biotite gneiss, garnet–orthopyroxene ± biotite gneiss (charnockitic gneiss), and graphite- bearing garnet–sillimanite gneiss (khondalitic gneiss) (Osanai et al., 2004). The minor amounts of mafic (clinopyroxene–orthopyroxene-bearing mafic granulite, orthopyroxene–biotite enderbitic gneiss) and calc-silicate rocks (olivine–spinel–humite or wollastonite-bearing calc- silicate rocks) intercalating within the pelitic ones are also present (Osanai et al., 2001b, 2004;
Nakano et al., 2007b, 2013). Intrusive rocks including granitoid and gabbro exposed in this complex occur as stocks and lenses within granulite-facies metamorphic rocks (Owada et al., 2016). The granitoids are divided into two types: garnet-bearing granite and orthopyroxene- bearing granite or charnockite (Owada et al., 2006, 2016). In term of the metamorphic aspect, the highest-grade pelitic granulite (garnet–orthopyroxene–sillimanite granulite; 1050 °C at 12 kbar) in the Kontum Massif was reported from the western Kannak Complex (Osanai et al., 2004). Whereas, the lower pressure condition was found in the eastern portion of the complex (6.3–8.4 kbar; Nakano et al., 2007b) with increasing temperature condition from the north (ca.
800 °C) to south (ca. 900 °C). Osanai et al. (2001b) investigated the UHT evolution, isothermal decompression and exhumation history for this complex through mineral assemblage
Fig. 2.6. CompiledP–Tpath for the Kannak, Ngoc Linh and Kham Duc Complexes in the Kontum Massif. The various metamorphic P–Tgradients are obtained from the Permian–Triassic rocks but all of them follow the typical clockwise P–T trajectory. The near isobaric heating under high T/P condition is seen in the Ordovician–Silurian rocks. The source of data are from Osanai et al., (2004, 2008); Nakano et al., (2007a, 2013, 2016) and Owada et al. (2020)
500 600 700 800 900 1000 1100
20 40 60 80 100
Depth (Km)
0.0 0.5 1.0 1.5 2.0 2.5 3.0
Pressure (GPa)
Temperature (ToC)
KD KD
NL
KN KN
NL And
246 Ma Grt-Ky-Bt gneiss
248 Ma Grt-Opx-Crd- Sil granulite (W. KN)
247 Ma Ky eclogite- granulite (W. NL)
442 Ma Grt-Crd-Sil-Bt gneiss (E. KN)
0.0
NL KN
KD
Kannak Complex Ngoc Linh Complex Kham Duc Complex
P–Tpath at Permian–Triassic P–Tpath at Ordovician–Silurian
M0 ca. 270 Ma
M2 stage 240–255 Ma
Grt granite emplacement ca.240 Ma
429 Ma Grt-Crd-Sil-Bt gneiss (W. NL)
M1 stage 250–260 Ma M0 stage
.
