Journal of Petrology | Volume 38 | Number 7 | Pages 843-876 | 1997
© Oxford University Press 1997
Eclogites and Blueschists of the Pam Peninsula, NE New Caledonia: a Reappraisal
1 Department of Geology and Geophysics, University of SydneySydney, NSW 2006, Australia
2 Department of Earth Sciences, University of Hong KongPokfulam Road, Hong Kong
3 Laboratoire de Géologie, Université FranÇaise du PacifiqueBP 4477, Noumea, New Caledonia
Received May 8, 1996; Revised typescript accepted February 19, 1997
 |
ABSTRACT |
|---|
 TOP ABSTRACT INTRODUCTIONIntroductionRegional GeologyThe Pouebo terranePetrologyMineral ChemistryThermobarometryDiscussionReferences |
|---|
High-P rocks of the Pam Peninsula, NE New Caledonia, are divided into three zones: (1) an uppermost ferroglaucophane–lawsonite zone of Cretaceous to Eocene metasediments and metavolcanics of the Diahot terrane that experienced peak conditions involving P=7–9 kbar and T=400±58°C; (2) albite–epidote–omphacite zone Diahot terrane rocks that experienced blueschist facies conditions of P=12.6±1.2 kbar and T=570±36°C; (3) lowermost metabasic eclogites of uncertain age that form the Pouebo terrane which experienced high-P conditions of P-23.9±3.0 kbar and T 600°C. Eclogite occurs as metre- to kilometre-scale pods in coarse-grained hydrous mineral-rich ‘glaucophanite’ formed during hydration and decompression of the Pouébo terrane. Metamorphism and deformation were consequent to 44–51 Ma Eocene convergence, when sedimentary and ophiolitic nappes were thrust over the eclogites in a SW direction; white mica ages constrain metamorphism to have ended by 37±1 Ma. Large steps in metamorphic grade are coincident with SW-dipping and NE-dipping faults that separate the three zones and were formed during two stages: (1) comparatively slow uplift and hydration of the Pouébo terrane before it was juxtaposed with the albite–epidote–omphacite zone at P 14 kbar; (2) comparatively rapid uplift of both the Pouebo terrane and the albite–epidote–omphacite zone to form a domal core of eclogite flanked by significantly lower-grade rocks to the SW and NE.
KEY WORDS: blueschist; eclogite; high-P metamorphism; New Caledonia; thermobarometry
|
INTRODUCTION |
|---|
|
TOPABSTRACTINTRODUCTIONIntroductionRegional GeologyThe Pouebo terranePetrologyMineral ChemistryThermobarometryDiscussionReferences |
|---|
MINERAL ABBREVIATIONS: ab, albite; act, actinolite; aeg, aegirine; alm, almandine; amph, amphibole; ap, apatite; barr, barroisite; cats, Ca-tschermakite molecule; cel, celadonite; chl, chlorite; clin, clinochlore; cpx, clinopyroxene; cr, crossite; ct, chloritoid; cz, clinozoisite; daph, daphnite; di, diopside; ep, epidote; fgl, ferroglaucophane; gl, glaucophane; gr, grossular; grt, garnet; jd, jadeite; ky, kyanite; law, lawsonite; mu, muscovite; omph, omphacite; pa, paragonite; parg, pargasite; phen, phengite; py, pyrope; q, quartz; ru, rutile; sp, spessartine; tit, titanite; win, winchite; zo, zoisite.
|
Introduction |
|---|
|
TOPABSTRACTINTRODUCTIONIntroductionRegional GeologyThe Pouebo terranePetrologyMineral ChemistryThermobarometryDiscussionReferences |
|---|
One of the world's largest and most continuous exposures of eclogite and blueschist facies rocks occurs in an elongate anticlinal range (Lillie, 1975
) in the northeastern corner of New Caledonia (Fig. 1). This belt comprises metamorphosed Cretaceous to Eocene sediments and volcanics and has been the subject of numerous petrological studies (e.g. Brothers & Blake, 1972; Black & Brothers, 1977; Brothers & Yokoyama, 1982; Brothers, 1985; Bell & Brothers, 1985; Yokoyama et al., 1986; Ghent et al., 1987a, b; Black et al., 1988) that have documented a sequence of mineral zones ranging from lawsonite+albite-bearing schists in the southwest to omphacite+garnet-bearing eclogites in the northeast. The sequence has been interpreted as representing an intact metamorphic profile generated during southwest-dipping Eocene subduction, yet the increase in grade occurs over such a comparatively short distance that it can only be explained by anomalous pressure and/or temperature gradients (Brothers, 1985; Yokoyama et al., 1986). On the basis of new mapping and quantitative P–T estimates we divide high-grade rocks exposed in part of the anticline into three zones: an uppermost ferroglaucophane–lawsonite zone, an intermediate albite–epidote–omphacite zone and a lowermost eclogite massif. It is presumed that the ferroglaucophane–lawsonite and albite–epidote–omphacite zones contain metamorphosed Cretaceous to Eocene sedimentary and volcanic rocks (Maurizot et al., 1989), whereas the eclogite zone is formed mostly from metabasic rocks of uncertain age (see Black et al., 1988). The eclogites experienced high-P metamorphism followed by hydration and decompression that resulted in coarse-grained secondary crossite–rich assemblages. The eclogites are now exposed along the axis of a regional anticline, interpreted to have been formed as a metamorphic core that was shedding lower-grade rocks including the ophiolite nappe and overlying blueschists (Aitchison et al., 1995a). This interpretation satisfactorily explains the juxtaposition of anomalously low-grade rocks and eclogites (e.g. Carroué, 1971), which presented problems in previous interpretations of the area. Moreover, the new interpretation is similar to the tectonic history recently inferred for other, significantly younger, eclogite facies rocks in the region (Davies & Warren, 1988; Hill et al., 1992).
|
Below, we outline the petrology of the three zones, building on the substantial database provided by the previous work. As our work differs mostly in the interpretation of eclogite facies exposures, we concentrate on these rocks.
|
Regional Geology |
|---|
|
TOPABSTRACTINTRODUCTIONIntroductionRegional GeologyThe Pouebo terranePetrologyMineral ChemistryThermobarometryDiscussionReferences |
|---|
Eocene collisional orogenesis grossly affected the continental basement terranes of New Caledonia, when this dispersed Gondwana fragment collided with an intraoceanic island-arc system (Cluzel et al., 1994
; Aitchison et al., 1995a; see Brothers & Blake, 1972). This convergence choked an established subduction zone and resulted in overthrusting of the New Caledonian ophiolite coeval with the generation of high-grade eclogite and blueschist facies rocks currently exposed in the northeast of the island (Black & Brothers, 1977; Yokoyama et al., 1986; Ghent et al., 1994). The Palaeozoic to Mesozoic Téremba, Boghen and Koh terranes (Fig. 1) are formed from rocks of varying metamorphic grade, transgressively overlain by upper Cretaceous to Eocene sedimentary rocks deposited during Gondwana dispersal (Paris, 1981; Cluzel et al., 1994; Aitchison et al., 1995a). Eocene convergence resulted in the southwest-directed overthrusting of two nappes and the development of a collisional foredeep in which turbidite, olistostromal and coarse clastic sedimentation occurred (Cluzel et al., 1994). Dolerites, pillow basalts and Late Cretaceous (Campanian) siliceous shales and cherts (Aitchison et al., 1995b) of the ‘formation des basaltes’ (Paris, 1981; Fig. 1) form the lowermost nappe. A subsequent ophiolitic nappe dominates outcrop in the south of the island and structurally overlies all other pre-Neogene elements of New Caledonia, including the eclogite facies rocks in the north of the island (Carroué, 1971). Upper sections of the ophiolite sequence are absent, with outcrop being restricted to ultramafic and rare gabbroic rocks. The ophiolitic rocks give a range of ages, with clusters around 100–77 and 48 Ma (Prinzhofer, 1981), indicating a complex history. We interpret the Cretaceous ages as representing the timing of formation of the rocks constituting the ophiolite, whereas Eocene ages reflect thermal overprinting associated with ophiolite emplacement (Ghent et al., 1994).
Blueschist to eclogite facies rocks in northeastern New Caledonia are exposed in an elongate anticlinal range (Fig. 1) that is in fault contact with Cretaceous to Eocene sedimentary and volcanic rocks (Maurizot et al., 1989; Fig. 2a). In this paper, the blueschist to eclogite facies rocks in northeastern New Caledonia are divided into two terranes (after Cluzel et al., 1994): (1) blueschist facies metasedimentary and metavolcanic rocks of the Diahot terrane, which are correlated with tectonically disrupted Cretaceous to Eocene sedimentary and volcanic rocks further west that were not significantly affected by the Eocene metamorphism (Maurizot et al., 1989; Figs 1 and 2a); and (2) metabasic eclogite, glaucophanite and rare quartzite of the Pouébo terrane. A 40Ar/39Ar study of white micas from the two terranes yielded consistent cooling ages of 37±1 Ma (Ghent et al., 1994). Previous work has documented a progressive northeastward increase in the metamorphic grade of rocks forming the Diahot terrane through ordered carbon, lawsonite, graphite and spessartine isograds (Black, 1977; Brothers & Yokoyama, 1982; Fig. 2a). There is a dramatic series of mineralogical changes over a comparatively small distance where the Diahot terrane is in fault contact with the Pouébo terrane. A confusing pattern involving discontinuous and truncated almandine and omphacite isograds has been defined in these high-grade rocks (Black, 1977; Brothers & Yokoyama, 1982; Maurizot et al., 1989). In the Pam Peninsula, the confusing pattern can be understood in terms of what were once shallowly dipping faults being coincident with most isograds in the high-grade rocks, with the faults separating distinct tectonic elements. The shallowly dipping faults have themselves been extensively disrupted by steeply dipping faults with throws of the order of 100 m. However, the shallowly dipping contacts can still be observed in some places, and are inferred to be the significant contacts on the basis of grade variations established below. In this paper, we are only concerned with rocks that vary from the spessartine zone (Fig. 2a) through what has been mapped as the lawsonite–epidote transition, almandine and omphacite zones (after Yokoyama et al., 1986).
|
The extent of post-metamorphic tectonic disruption in the Pam Peninsula means that many structural fabrics do not correlate between the three tectonic units defined below. Structural chronologies are established for each unit and shared deformation events that time the amalgamation of the various units summarized in Table 1. SL1 refers to the earliest foliation in the ferroglaucophane–lawsonite zone, SA2 to the foliation produced by the second deformation event in the albite–omphacite–epidote zone, and SE3 to the foliation produced by the third deformation event in eclogites of the Pouébo terrane.
|
The Diahot terrane
Here we concentrate on two metamorphic zones in exposures of these rocks in the Pam Peninsula: ferroglaucophane–lawsonite schist broadly equivalent to rocks of the spessartine zone of Yokoyama et al., (1986
, and references therein; Fig. 2a), and schist of the albite–epidote–omphacite zone. Rocks in the albite–epidote–omphacite zone lie above the epidote, almandine and omphacite isograds of Yokoyama et al., (1986) and include rocks described as ‘felsic eclogites’ by Black et al., (1988). However, these rocks are not eclogites in the currently accepted nomenclature (Carswell, 1990), and we find that the metamorphic zones are less continuous than was inferred in previous work, owing to them being comparatively thin and having experienced significant post-metamorphic tectonic disruption.
In the ferroglaucophane–lawsonite zone, west- to northwest-trending FL3 folds preserve a steeply dipping SL3 axial planar foliation that deforms a layer-parallel SL2 (Maurizot et al., 1989) and a southwest-trending LL2 mineral and stretching lineation (Cluzel et al., 1995). Evidence for SL1 can only be recognized in thin section, where SL1 glaucophane and white mica are oblique to, and enveloped by, SL2. Southwest of the Pam Peninsula, SL3 is cut by shallowly NE-dipping DL4 faults that thrust Cretaceous metasediments of the Diahot terrane over lower-grade Eocene metasediments (Fig. 2a) of the same terrane. Contacts between the ferroglaucophane–lawsonite zone and either the albite–epidote–omphacite zone or Pouébo terrane (Fig. 2a) are mostly steeply dipping shear zones that are parallel to SL3. For example, tectonic windows of Pouébo terrane have been upfaulted along SL3 through ferroglaucophane–lawsonite zone schist at and northwest of Ouégoa; throws of the order of 100 m are inferred. Between Ouégoa and the Col d'Amoss (Fig. 2a), a thin layer of ferroglaucophane–lawsonite schist is inferred to dip parallel to a pervasive, shallowly southwest-dipping SL2. In this area, asymmetric shear sense indicators are consistent with southwest-directed shear during the development of SL2.
Rocks in the albite–epidote–omphacite zone are significantly coarser grained than rocks in the ferroglaucophane–lawsonite zone, but bedding is mostly well preserved. The effects of at least four periods of deformation (DA1-DA4) can be recognized (see also Maurizot et al., 1989). Bedding and SA1 are isoclinally folded by north-trending, reclined or recumbent FA2 folds and transposed into parallelism with a shallowly dipping SA2 that is broadly coplanar with the upper and lower boundaries of the zone. Open to tight north-trending FA3 folds with a steeply dipping axial plane and amplitudes of tens of metres and wavelengths of the order of 100 m deform SA2. Steeply dipping shear zones oriented parallel to SA3 define some contacts between albite–epidote–omphacite zone schist and the Pouébo terrane (Fig. 2a). West-trending FA4 folds, also with amplitudes of tens of metres and wavelengths of the order of 100 m, extensively crenulate both SA2 and SA3, and mesoscopic FA2-FA4 fold interference patterns are common (Fig. 3a). The effects of DA3 are not observed in the ferroglaucophane–lawsonite zone schists. On the basis of style and orientation criteria, SA4 is equivalent to SL3, consistent with the two zones of the Diahot terrane having been together when this deformation event occurred. The central part of the Pam Peninsula has extensive exposures of fine-grained, metasedimentary glaucophane–albite–garnet schist that contain SA3 (Fig. 2a). These schists are carbonaceous and more fine grained than typical albite–epidote–omphacite zone schists further west. In the absence of diagnostic lawsonite or epidote, they have been placed in the albite–epidote–omphacite zone on the basis of glaucophane and garnet compositions (Table 2) and their preservation of the SA3 foliation.
|
|
On the southwestern flank of the Pam Peninsula, the lower contact of the albite–epidote–omphacite zone is inferred to dip shallowly to the southwest closely matching the slope of the ground (Section B–B', Fig. 2b). On the northeastern tip of the Pam Peninsula, the same contact is inferred to dip shallowly northeastwards but is folded and faulted by post-DA2 deformation (section A–A', Fig. 2). Projection of the structural section from 50 km along strike towards Hienghène (Fig. 1) has this zone overlain by northeast-dipping klippen of lawsonite zone rocks, which preserve a northeast-plunging LL2 mineral and stretching lineation (Cluzel et al., 1995
). Thus, an antiform cored by eclogites of the Pouébo terrane is draped by folded klippen of albite–epidote–omphacite and ferroglaucophane–lawsonite zones on a macroscopic scale (Cluzel et al., 1995; Fig. 2b). This simple regional setting is locally complicated by late faulting and limited post-metamorphic folding. |
The Pouébo terrane |
|---|
|
TOPABSTRACTINTRODUCTIONIntroductionRegional GeologyThe Pouebo terranePetrologyMineral ChemistryThermobarometryDiscussionReferences |
|---|
The Pouébo terrane (Cluzel et al., 1994
) mostly contains metabasic eclogite as metre- to kilometre-scale pods in coarse-grained hydrous-mineral-rich ‘glaucophanite’ (Maurizot et al., 1989), formed by hydration of eclogite. Boulders of quartzite interlayered with metabasic eclogite are observed at a few locations. The terrane is intensely dissected by structurally complex crossite–phengite-rich hydrous shear zones that controlled the fluid influx and, most probably, uplift of the Pouébo terrane. Near the hydrous shear zones, glaucophanite is commonly intensely foliated, but random late hydrous minerals cut eclogite facies foliations distal to the shear zones. Mineral assemblages of the Pouébo terrane lie above the almandine and omphacite zones of Yokoyama et al., (1986), but can be discriminated from the albite–epidote–omphacite zone defined above on the basis of protolith, grain size and the lack of S2E plagioclase. The terrane includes part of what has been previously mapped as the epidote zone (e.g. Yokoyama et al., 1986), but the role of tectonic disruption and timing of mineral growth was not fully considered in previous zone definitions. There are abrupt variations in metamorphic grade at the tectonic boundaries of the Pouébo terrane, as noted in earlier works. For example, Briggs et al., (1978) recognized ‘fault emplaced’ blocks of eclogite within low-grade schists at Ouégoa and Col d'Amoss (Fig. 2a) and low-grade schists within eclogite near Pouébo. Carroué, (1971) reported very low-grade blocks of ophiolite and Inoceramid-bearing sediments apparently surrounded by eclogite southeast of Pouébo (Fig. 1). The presence of these blocks cannot be reconciled with the previous interpretation that metamorphic grade progressively increased with distance northeast, but they may be explained as tectonic windows of high-grade Pouébo terrane in klippen of the lower-grade Diahot terrane.
Three types of eclogite are recognized. Type I eclogite is dark green and barroisite rich, intensely foliated and has a bulk rock composition consistent with a basaltic protolith. Type II eclogite is light coloured, weakly foliated and has a bulk rock geochemistry consistent with it being a metamorphosed gabbro cumulate (Reid, 1994). Type III eclogite is rare. It is found in boulders, tectonically interlayered with Type I eclogite. Both Type I and II eclogite contain garnet+omphacite-bearing assemblages that, where minimally affected by the secondary hydration, lack plagioclase and are true eclogites (after Carswell, 1990). Type I eclogite contains a shallowly dipping SE3 foliation and rare, tight to intrafolial FE3 folds that transpose a pervasive SE2 foliation (Fig. 3b). Evidence for SE1 is preserved by curved to spiral inclusion trails in garnet (also Bell & Brothers, 1985; Fig. 4a). Type II eclogites outcrop in the northwestern Pam Peninsula (Fig. 2a) as large discontinuous pods tectonically interlayered with Type I eclogites. Type II eclogite is commonly less intensely foliated than Type I eclogite, with a weak SE3 foliation defined by glaucophane, paragonite and phengite cut by post-SE3 hornblende with or without omphacite (Fig. 3c). Near boundaries with Type I eclogite, Type II eclogite contains a pervasive SE3 and is tectonically interlayered with Type I eclogite; thus the difference in foliation intensity would seem to mostly reflect DE3 strain partitioning.
|
The recrystallization of metabasic eclogite was related to fluid influx along large, shallow to steeply dipping DE4 and DE5 hydrous shear zones that cut all eclogites and preserve diverse mineral assemblages. These zones commonly contain large crystals of hydrous minerals such as crossite, epidote, actinolite–winchite and phengite, and they are mostly distinct to the shear zones that cut the Diahot terrane. Individual shear zones are commonly between 50 and 100 m across and have complex internal structural and metamorphic relationships; pods of partially to pervasively recrystallized eclogite are enveloped by crossite-chlorite-phengite-rich blastomylonites that commonly preserve several in-foliation mineral and stretching lineations oblique to each other. Large DE4 shear zones are commonly separated by low strain domains of several hundred metres width, in which the secondary recrystallization is commonly randomly oriented and patchily developed. Eclogite facies assemblages persist in these domains, presumably owing to irregularities in strain and/or restricted fluid penetration.
The DE4 zones are truncated by the contact between the Pouébo terrane and the albite–epidote–omphacite zone, and rocks of the Diahot terrane show limited post-peak recrystallization (e.g. Bell & Brothers, 1985). The upper boundary of the Pouébo terrane commonly involves a discontinuous zone of intensely foliated glaucophanite 100–300 m thick. A pervasive SE5 foliation in these rocks is sub-parallel to the inferred shallowly dipping contact with the albite–epidote–omphacite zone and SA2. In the northwestern Pam Peninsula, uppermost portions of the Pouébo terrane may be weakly recrystallized Type II eclogite (Fig. 2a), and strain related to their juxtaposition with the Diahot terrane has apparently been accommodated in a narrow (<20 m thick) zone of intensely deformed albite–epidote–omphacite zone schist. Thus SE5 may be defined by blueschist, eclogite intensely recrystallized to glaucophanite, or greenschist. However, the shallowly dipping SE5-SA2 foliation is extensively folded about the west- to northwest-trending SA4-SL3 foliation (Table 1), movement along which controls the shape of the eclogite windows in the southern Pam Peninsula (Fig. 2a). Juxtaposition of the Pouébo and Daihot terranes is thus constrained to post-date DE4 and the main phase of hydrous recrystallization of the Pouébo terrane, but predate deformation that resulted in the SE5-SA2 foliation. Mineral assemblages that define SE4 and SE5 record P–T conditions similar to those of the albite–epidote–omphacite zone and it seems probable this polyphase deformation occurred through a large range in pressure during exhumation of the Pouébo terrane.
|
Petrology |
|---|
|
TOPABSTRACTINTRODUCTIONIntroductionRegional GeologyThe Pouebo terranePetrologyMineral ChemistryThermobarometryDiscussionReferences |
|---|
Type I eclogite
Dark green Type I eclogite and recrystallized varieties form most of the Pouébo terrane (Fig. 2a). Eclogite with minimal recrystallization contains large grains of idioblastic to xenoblastic garnet (
5–10 mm diameter) enveloped by a granoblastic SE2 foliation defined by barroisitic hornblende, quartz, epidote, phengite and rutile, with or without omphacite, apatite and titanite (Fig. 4b). Omphacite, barroisite, epidote and mica grains are commonly 2–5 mm in diameter, rutile and titanite grains commonly being <1 mm in diameter. Titanite may enclose rutile, and post-SE3 chlorite commonly partially pseudomorphs garnet. Omphacite is typically xenoblastic and nearly colourless. Where abundant it may occur as large porphyroblasts that contain inclusions of glaucophane and epidote. Garnet preserves complex inclusion patterns. Commonly, a fractured inclusion-rich core contains random inclusions involving all or some of glaucophane, epidote, omphacite, titanite and quartz. This core is then enclosed by what is a comparatively inclusion-poor domain that commonly forms the outer half or one-third of the garnet (Fig. 4c). Partial overgrowths on the inclusion-poor domain may be continuous with the enveloping SE2 or SE3, consistent with syn-tectonic growth. Rare large garnet grains (1–2 cm diameter) may have optically distinguishable grossular- and spessartine-rich cores not present in smaller grains in the same rock. Abundant oriented titanite inclusions in narrow zones at the inner part of the inclusion-poor domain (e.g. Fig. 4a) preserve evidence for SE1. Rutile appears as inclusions in garnet between the inner part of the inclusion-poor domain and the grain boundary.
Type II eclogite
Large random laths of omphacite and/or light green hornblende commonly cut a weak SE2 or SE3 foliation defined by phengite, clinozoisite and/or epidote, garnet, quartz, rutile and glaucophane with or without paragonite and/or titanite. SE2 and SE3 cannot be discriminated in Type II eclogite. Garnet-rich seams and laminae define SE2-SE3 in field observation. Quartz is modally less significant, and titanite and rutile modally more significant, than in Type I eclogite. Idioblastic omphacite laths are optically and chemically zoned (Fig. 4d) and may contain oriented clinozoisite and glaucophane inclusions. Large, subhedral to xenoblastic hornblende laths are also commonly zoned, and have a compositional range that overlaps the boundaries of the actinolite–winchite–barroisite fields (see below). Hornblende may enclose and partially pseudomorph omphacite. SE2-SE3 glaucophane is pale violet or nearly colourless. Where present in the matrix of these rocks, titanite commonly encloses rutile or occurs in titanite-rich laminae that define an SE2 crenulation cleavage (Fig. 4e). Garnet is commonly idioblastic and preserves complex inclusion relationships. Large grains (>2 cm diameter) may have a core zone dominated by intergrown garnet and titanite that commonly preserve spiral SE1 inclusion trails and may or may not be enclosed by an inclusion-poor zone. Small garnet grains adjacent to the large grains commonly lack the titanite-rich cores. Within, or immediately adjacent to the titanite-rich domain, are inclusions of glaucophane, epidote, quartz and lawsonite (Fig. 4f). If present as inclusions in garnet, rutile occurs in the inclusion-poor zone. Rare inclusions of omphacite are only found near the rim of garnet grains.
Intensely foliated Type II eclogite separates Type I eclogite from discontinuous pods of weakly foliated Type II eclogite. The intense SE2-SE3 foliation, which envelopes garnet, is defined by blue–green barroisitic hornblende, clinozoisite, titanite, rutile, omphacite, paragonite and quartz. Garnet grains are mostly smaller than in weakly foliated Type II eclogite and contain inclusions of quartz and clinozoisite.
Type III eclogite
Quartz, andraditic garnet, epidote, glaucophane, titanite and pyrite define an intense SE3 in rare quartzite, which occurs interlayered with metabasic eclogite. Andraditic garnet commonly forms poikiloblasts that have abundant quartz inclusions. Quartz commonly forms >70% of these samples, which are inferred to be metamorphosed cherts.
Glaucophanite
Previous reports (e.g. Yokoyama et al., 1986; Bell & Brothers, 1985) have emphasized the secondary growth of albite at the expense of omphacite in New Caledonian ‘eclogites’, but many of the rocks so described come from the albite–epidote–omphacite zone. Moreover, recrystallized Type I eclogite, here called glaucophanite (after Maurizot et al., 1989), may contain post-SE3 omphacite and rutile that are interpreted to be part of a secondary assemblage. Coarse-grained, random blue crossite, garnet, idioblastic omphacite and epidote, phengite, rutile and quartz are intergrown and interpreted as being in textural equilibrium in such samples. Garnet commonly preserves growth zoning identical to that observed in the minimally recrystallized eclogites (see below). However, the idioblastic matrix omphacite is aegirine rich and green, chemically zoned (see below), and optically distinct from nearly colourless aegirine–poor omphacite present in the same sample as inclusions in garnet. The mineral chemistry of the omphacite inclusions in garnet (see below) is consistent with them having formed at the same time as SE3, and with the post-SE3 matrix omphacite being part of a secondary assemblage. Such assemblages were probably developed early in the hydration event, and most probably at high pressures.
In domains of low DE4 strain, most glaucophanites contain random epidote, crossite and phengite, with or without titanite and/or albite. SE2-SE3 omphacite is commonly partially to completely pseudomorphed by post-SE3 crossite and epidote. Albite is intergrown with the pseudomorphous crossite and epidote in some samples. Chlorite and phengite, with or without albite, commonly partially to completely pseudomorph garnet. Garnet may persist in some assemblages as large (>5 mm diameter) grains that lack the growth zonation (Table 2) characteristic of eclogitic garnet (see below). These garnets may have inclusion-rich cores and inclusion-poor rims, though the inclusions are generally coarser than inclusions in eclogitic garnet. Amphiboles in the glaucophanites are complex. Individual grains may be zoned from green barroisitic cores to blue crossite rims. Blue or green barroisite may be intergrown with blue crossite, without any apparent reactive relationship. Large quartz- or phengite-rich veins containing random, coarse-grained epidote and crossite, with or without chlorite are common.
Intensely foliated secondary assemblages and monomineralic layers envelop pods of partially to pervasively recrystallized eclogite in the hydrous shear zones. Amphiboles in these zones are diverse, though crossite is usually present. Porphyroclastic green barroisite may have crossite rims, or be intergrown with crossite. Shear zones that cut Type II eclogite contain clots of coarse-grained actinolite and may lack glaucophane. Garnet may be pseudomorphed by chlorite with or without albite (Fig. 3d). En-echelon quartz or carbonate veins may cut less recrystallized blocks, indicating that at least some of the deformation was brittle and was accompanied by high fluid activities. The mineral assemblages developed in these hydrous shear zones may be similar to those developed in the albite–epidote–omphacite zone and field relationships can be ambiguous near the contact between the two units. However, minerals in the hydrous shear zones are larger than equivalents in the albite–epidote–omphacite zone, and mineral assemblages in the hydrous shear zones commonly lack quartz.
Albite–epidote–omphacite zone
As described in previous studies, diverse mineral assemblages occur in this zone, reflecting a variety of metasedimentary protoliths. The common rock type in the zone is a glaucophane–albite–phengite–garnet schist that contains an intense SA2 foliation, but pods of patchily recrystallized metabasaltic schist and siliceous chloritoid schist are observed. The most useful rock type for constraining the P–T conditions is a siliceous glaucophane–albite–omphacite schist that has previously been used to constrain the ‘eclogite’ facies conditions (Yokoyama et al., 1986; Black et al., 1988; Ghent et al., 1994). It contains albite and garnet porphyroblasts (2–3 mm diameter) enveloped by fine-grained (<0.5 mm) phengite, epidote, quartz, glaucophane and titanite, with or without paragonite, omphacite and rutile, which all define SA2. SA1 is defined by curved inclusion trails of titanite and epidote in garnet, and curved inclusion trails of glaucophane, omphacite, quartz and epidote in albite (Fig. 5a and b). Many rocks contain omphacite both as comparatively large (1 mm diameter) idioblastic matrix grains that define SA2 (Fig. 5c) and are inferred to be in textural equilibrium with all other minerals, and as smaller (0.1 mm diameter) idioblastic inclusions in albite (Fig. 5d; Bell & Brothers, 1985) that either define SA1 or cut SA1 (Fig. 5a and b). Rutile is commonly mantled by titanite. Garnet does not display systematic inclusion patterns, but may have titanite-rich cores overgrown by comparatively inclusion-poor domains where quartz is the commonly included mineral. Inclusion-poor, syn-DA2 overgrowths, which may include omphacite, occur on many garnets. Previous interpretations involve omphacite in these rocks having crystallized early (pre-SA2) and being partially pseudomorphed by albite (e.g. Bell & Brothers, 1985). However, as omphacite also occurs as large grains that define SA2, we infer that it was stable through the change in conditions that accompanied the shift from SA1 to SA2.
|
What have previously been referred to as ‘felsic eclogites’ (Black et al., 1988
) are exposed in the vicinity of the peak Bouehndep, and were mapped by Maurizot et al., (1989) as a ‘meta-rhyolite’ (silicious glaucophane-epidote schist in Fig. 2a). These rocks are feldspar-rich variations of the omphacite schists described above, but they may contain jadeite instead of, or in addition to, omphacite (Black et al., 1988). Chloritoid occurs in siliceous schists that additionally contain chlorite, albite and phengite (sample RS14, Table 2), also as part of the ‘meta-rhyolite’ and elsewhere throughout the albite–epidote–omphacite zone (Black et al., 1993).
Ferroglaucophane–lawsonite zone
The comparatively low-grade rocks exposed to the southwest of the Pam Peninsula (Fig. 2a) will not be described in detail here as they have been extensively studied and well described in previous reports [Yokoyama et al., (1986) and references therein]. Below, we summarize features that distinguish them from the other zones. Rocks above the spessartine isograd shown in Fig. 2a comprise fine-grained ferroglaucophane–albite–phengite metasedimentary schist and patchily recrystallized metabasalt. Lawsonite mostly occurs in the partially recrystallized metabasalt in association with ferroglaucophane. However, it also occurs in metasedimentary schist that contains SL2 lawsonite, albite, phengite, chlorite and titanite, with or without rutile, garnet or paragonite. The average grain size in these rocks is 0.1 mm (diameter), making the rocks appreciably finer than rocks in the albite–epidote–omphacite zone. Individual mica grains cannot be observed in hand specimen, and ferroglaucophane is commonly <0.3 mm in length.
Near the contact with the albite–epidote–omphacite zone, the ferroglaucophane–albite–phengite schists contain comparatively coarse-grained (1–2 mm in length) SL1 ferroglaucophane and albite porphyroblasts. These are enveloped by an intense southwest-dipping SL2 foliation defined by fine-grained ferroglaucophane, albite, titanite, phengite and quartz, with or without garnet or lawsonite. Albite may contain curved SL1 inclusion trails. Distal to the albite–epidote–omphacite zone, SL2 is folded by upright northwest-trending FL3 folds, and random glaucophane cuts SL2.
|
Mineral Chemistry |
|---|
|
TOPABSTRACTINTRODUCTIONIntroductionRegional GeologyThe Pouebo terranePetrologyMineral ChemistryThermobarometryDiscussionReferences |
|---|
Analyses were performed on the Cameca SX-50 Camebax microprobe housed at the University of New South Wales, operating with an accelerating voltage of 15 kV, a beam width of 1–5 µm and PAP data reduction software supplied by the manufacturer. The major features of mineral chemistry for samples used for thermobarometry are presented below (see also Black, 1973a
, b, 1975), with representative compositions summarized in Table 2 and representative microprobe analyses presented in Table 3. Many of the minerals are zoned, and there may be restricted ranges in mineral chemistry of individual minerals in any one sample. Generalizations are made below for the case of simplicity, but reference is given to type examples on which the generalization is based to allow cross referencing.
|
Garnet
The major features of garnet have been reviewed by Black (1973a)
, but are discussed here in terms of the newly defined zones and chemical zoning patterns. Garnet is almost invariably zoned with cores enriched in spessartine relative to almandine, but the nature of zoning varies with grain size, protolith and metamorphic grade. Garnet in the ferroglaucophane–lawsonite zone metapelitic rocks is spessartine-rich, pyrope-poor relative to the higher-grade rocks and preserves pronounced zoning from cores with Xspess=Mn/(Fe+Mn+Mg+Ca)0.4–0.6 and Xalm=Fe/(Fe+Mn+Mg+Ca)0.2–0.4 to rims with Xspess0.2–0.4 and Xalm0.3–0.5. Pyrope and grossular contents vary only subtly through this commonly bell-shaped profile with Xpy=Mg/(Fe+Mn+Mg+Ca)0.01 and Xgross=Ca/(Fe+Mn+Mg+Ca)0.2–0.3 (both core to rim). In the lawsonite zone, garnet XFe=Fe/(Fe+Mg)>0.95. Garnet in metapelitic schists of the albite–epidote–omphacite zone can also have pronounced bell-shaped zoning profiles (Fig. 6a) with antithetic zoning of spessartine relative to almandine, grossular and pyrope contents. The amplitude of the zoning profile varies between samples and between grains within individual samples. Whereas the rim compositions of garnet tend to be broadly similar, larger grains commonly preserve a fuller profile involving cores of (sample 9403) Xalm0.39, Xspess0.35, Xpy0.02, Xgross0.24 changing through smooth profiles to rims of Xalm0.63, Xspess0.02, Xpy0.07, Xgross0.28. There is a subtle variation in XFe0.95–0.90 from core to rim. Xpy of garnet in all eclogites is <0.3, making them Type C (crustal) eclogites (Coleman et al., 1965; Mottana et al., 1990). Garnet in Type I eclogite commonly preserves inclusion-rich cores of (sample 9341) Xalm0.50, Xspess0.10, Xpy0.05, Xgross0.30, XFe0.90 changing to inclusion-poor shoulders of Xalm0.60, Xspess0.05, Xpy0.10, Xgross0.25, XFe0.85, and then rims of Xalm0.58, Xspess0.02, Xpy0.12, Xgross0.29, XFe0.83 (Fig. 6a). Garnet may have variable core compositions both between and within samples, but the rim compositions are uniform within samples and between samples on a local scale. The optically distinguishable grossular-rich cores (sample WC68) of garnet in Type I eclogite coincide with steps in the zoning profile, which changes from the common core composition given above to Xalm0.41, Xspess0.18, Xpy0.03, Xgross0.37, XFe0.94. There is no apparent systematic variation in Type I eclogite garnet compositions throughout the Pam Peninsula, consistent with there being little variation in grade or bulk rock composition. Type II eclogite may be MnO poor in comparison with Type I eclogite, meaning that garnet commonly has less well-developed bell-shaped spessartine zoning profiles. Instead, garnets with comparatively grossular- and spessartine-rich cores are zoned continuously to pyrope-rich, grossular-poor rims. For example, sample WC110 has cores of Xalm0.57, Xspess0.06, Xpy0.07, Xgross0.30 and rims of Xalm0.54, Xspess0.01, Xpy0.21, Xgross0.24. Sample 9409b has spessartine-rich cores (Xspess0.16), and sample 9408c has higher almandine and lower grossular contents in all garnets than other samples. Garnet in the glaucophanites is more variable: it can preserve growth zoning similar to that described above for the Type I eclogite, show no zoning, or preserve rim compositions enriched in pyrope and diminished in grossular relative to garnet in the common Type I eclogite. For example, sample RS31c preserves garnet with core compositions similar to that of garnet in Type I eclogite but has rim compositions of Xpy0.23 and Xgross0.15.
|
Representative zoning profiles of individual garnet grains are shown in Fig. 6a. XFe is the factor that changes least in these profiles, so the zonation patterns can be summarized on a ternary diagram with apices Xgross, Xspess and (Xalm+Xpy)/2 (Fig. 6b). The core to rim decrease in spessartine content in Type I eclogite is compensated for mostly by increasing almandine+pyrope content, although some profiles then show a subtle decrease in grossular content. In comparison, the decreasing spessartine content in Type II eclogite garnet is initially compensated by increasing grossular and almandine+pyrope content before a marked increase in almandine+pyrope content (Fig. 6b). Garnet in the albite–epidote–omphacite zone has cores with significantly higher spessartine and lower grossular contents than the eclogites, with the rimward decrease in spessartine content compensated by significantly increasing grossular and increasing almandine+pyrope contents (Fig. 6b).
Clinopyroxenes
Clinopyroxenes in the Pam Peninsula are mostly omphacite, though aegirine and jadeite have been reported by Black et al., (1986). For the purpose of comparison, clinopyroxene compositions have been calculated following Morimoto, (1988) and plotted on a ternary diagram with apices jadeite, aegirine and diopside+hedenbergite (Fig. 7). Omphacite in Type II eclogite contains minor aegirine but may be zoned in jadeite, diopside and hedenbergite compositions. Large omphacite grains in sample 9408c are atypical in that they are strongly zoned from comparatively diopside+hedenbergite-poor cores with Jd75 to diopside+hedenbergite-rich rims with Jd51. Omphacite in other Type II eclogites shows a restricted range of compositions between Jd43 and Jd51 (Fig. 7). Clinopyroxene in Type I eclogite is commonly omphacite (Jd40–45), but clinopyroxene in some samples straddles the omphacite–aegirine–augite boundary (Fig. 7). There is a general trend between samples of increasing aegirine content at nearly constant diopside+hedenbergite content, meaning that the controlling substitution is simply Fe3+Al-1. There is limited within-sample variation in clinopyroxene composition. Large omphacites in WC68 are zoned from cores involving Jd47 to rims involving Jd40. Aegirine–augite inclusions in garnet in sample WC11 preserve lower jadeite and higher aegirine (Jd18Aeg23) contents than matrix omphacite grains (Jd34Aeg15). Idioblastic omphacite grains in glaucophanite have significant aegirine content and may be zoned. For example, matrix omphacite in sample 9342 is zoned from cores with Jd42Aeg21 to rims with Jd32Aeg24, both being distinct from omphacite inclusions in garnet with Jd34Aeg31. Omphacite in samples of epidote–albite–omphacite zone schist used in this study has negligible aegirine content and Jd39–43. Where found as inclusions in garnet, omphacite had an identical composition to matrix S2A omphacite.
|
Epidote group minerals
These include clinozoisite in Type II eclogite, and epidote in Type I and II eclogite and albite–epidote–omphacite schist. Ferric iron in epidote analyses was recalculated assuming two-site ordering with a total of six silica, aluminium and ferric cations per 12.5 oxygens using the computer program AX (shareware written by T. J. B. Holland, http://www.esc.cam.ac.uk/software.html). As a generalization, clinozoisite or epidote in Type II eclogite contains less Fe3+ than epidote in Type I eclogite, which contains less Fe3+ than epidote in glaucophanite: Cz=(Al - 2)/(Al - 2+Fe3+) content commonly ranges from 1 to 0.85 for clinozoisite in Type II eclogite; Type I eclogites contain epidote commonly in the range Cz54–63; and glaucophanites contain epidote commonly in the range Cz25–50. Exceptions to these generalizations occur (Table 2), but the trend of increasing pistacite end-member in moving from Type II eclogite to Type I eclogite to glaucophanite matches a similar pattern established above for aegirine content in clinopyroxene. Epidote in the albite–epidote–omphacite zone schist lies in the range Cz49–60.
Amphiboles
Amphiboles show a wide range in composition (Black, 1973b), though they are mostly hornblende or glaucophane–crossite (after Leake, 1978). Type I eclogite with minimal effects of the blueschist hydration contain barroisitic hornblende, with between 7.5 and 6.8 Si cations per 23 oxygens at XNa=Na/(Na+Ca)0.5 (Fig. 8a; after Yokoyama et al., 1986). When Fe3+ content and site distribution are calculated on the basis of charge balance following Robinson et al., (1982), Fe2+:Fe3+ ratios are of the order of 3:1 and the varying Si content mostly reflects antithetic edenitic and pargasitic substitution at constant tschermakite content. Type II eclogite with minimal effects of the blueschist hydration contains glaucophane with a restricted compositional range (Fig. 8) and minimal ferric iron content. Hornblende in Type II eclogite shows a trend of increasing Si (7.4–7.8 cations) and decreasing Na (XNa0.4–0.2) content from the silica-rich end of the barroisite field, through winchite to actinolite. There can be considerable within-sample variation in Si content, which mostly reflects tschermakitic, edenitic and glaucophane exchanges. Ferrous to ferric iron ratios range from approximately 2:1 to 3:1 in these hornblendes. Amphiboles in glaucophanite and the hydrous shear zones may be considerably more complex, with individual grains commonly involving green barroisitic cores and blue barroisite, winchite or crossite rims. Alternatively, crossite may be intergrown with a blue or green hornblende, apparently in textural equilibrium. Crossite in glaucophanite is appreciably more Fe rich than glaucophane in Type II eclogite (Table 2) and shows a significant range in Si content, with the charge imbalance created by substitution of Si compensated for mostly by ferritschermakitic and glaucophane exchanges, with limited antithetic edenitic exchange. Whereas the Si-XNa trends of the two amphibole series converge in Fig. 8, the series remain distinct when tetrahedral alumina is plotted against either octahedral alumina or A-site sodium (after Robinson et al., 1982). Hornblende in glaucophanite has a similar, but less tightly constrained XNa-Si trend to that established above in the minimally recrystallized eclogites, with the scatter mostly reflecting varying ferritschermakitic exchange. There is an appreciable range in antithetic glaucophane and edenite substitutions. Blue amphibole in the epidote–albite–omphacite zone is mostly glaucophane. Ferroglaucophane in the ferroglaucophane–lawsonite zone has minor ferric iron content.
|
Micas
Silica content in phengite ranges from 6.7 to 7.0 cations per formula unit (p.f.u.) (22 oxygens), with moderate content of octahedrally coordinated cations (Table 2). A general antithetic trend can be established for silicon and soda contents, ranging from 7.0 Si cations and XNa=Na/(Na+K)
0.05 for phengite in ferroglaucophane–lawsonite schist to 6.6 Si cations and XNa0.2 for phengite in the eclogites (Fig. 9). Phengite in epidote–albite–omphacite schist is intermediate in composition between that in lawsonite schist and eclogite, with phengite in the eclogites generally having <6.8 Si cations and XNa>0.1. Isopleths of sodium content in phengite mapped by Yokoyama et al., (1986) in the northeastern Pam Peninsula define the lower boundary of the albite–epidote–omphacite zone (Fig. 2a). Similarly, paragonite in the albite–epidote–omphacite schist has 6.0–6.2 Si cations per formula unit, in comparison with 5.8–6.0 Si cations for paragonite in Type II eclogite (Fig. 9).
|
|
Thermobarometry |
|---|
|
TOPABSTRACTINTRODUCTIONIntroductionRegional GeologyThe Pouebo terranePetrologyMineral ChemistryThermobarometryDiscussionReferences |
|---|
On the basis of 16O/18O fractionation between metamorphic minerals, Black, (1974)
inferred metamorphic temperatures for rocks of the Diahot terrane to range from 250°C in the ferroglaucophane–lawsonite zone to 400°C for the albite–epidote–omphacite zone. The observation of aragonite within lawsonite-bearing rocks from the Diahot terrane (Brothers, 1974) is consistent with pressures for the higher-grade rocks in this zone having been above 6 kbar and most probably having been 9.5 kbar (Yokoyama et al., 1986). The conditions of formation of the SL2 mineral assemblages in the high-grade rocks of the ferroglaucophane–lawsonite zone may also be determined using the average pressure–temperature (P–T) method, using microprobe analyses of the minerals and the computer software Thermocalc (version 2.5; Powell & Holland, 1988), with the internally consistent thermodynamic dataset of Holland & Powell, (1990; datafile created April 1996). In finding an average P–T of a mineral assemblage, an independent set of reactions between the end-members of the minerals in the assemblage is used. Different sets of independent reactions that define a series of P–T lines may be found under different conditions, and the software selects the appropriate set to minimize errors on the calculated P and T conditions. In calculating an average P–T, the lines are displaced so as to coincide at this P–T. These displacements are mainly made by varying the activities of the end-members of the minerals, in proportion to their uncertainties. All mineral end-member activities were calculated using the computer program AX and the defaults suggested by Powell & Holland, (1988). Sample 9338 is a metasediment taken from near the northeastern boundary of the ferroglaucophane–lawsonite zone (Fig. 2a) and contains the SA2 assemblage garnet–glaucophane–phengite–chlorite–albite–quartz–titanite–rutile. Average P–T calculations return P=6.9±4.0 kbar and T=415±152°C (all 2
errors), though more precise estimates for the conditions of SA2 may be made by running average T and average P separately: for a pressure of 6.5 kbar the average T method returns T= 400±58°C, and for T=400°C the average P returns P=6.5±1.6 kbar (Table 4). When these error ranges are overlain with the lawsonite breakdown reaction defined by Chatterjee et al., (1984), peak conditions in the high-grade end of the ferroglaucophane–lawsonite zone involved P=7–9 kbar for T400°C (Fig. 10b).
|
|
The simplest P–T estimates for conditions appropriate to mineral assemblages present in the albite–epidote–omphacite zone can be made by combining P estimates made on the basis of coexisting albite and omphacite (after Holland, 1980
) with T estimates made on the basis of coexisting garnet and clinopyroxene [rim compositions; after Krogh, (1988)]. Figure 10a shows the position in P–T of the reaction jadeite+quartz=albite for omphacite of composition Jd55 and appropriate Kd isopleths for FeMg-1 exchange between coexisting garnet and clinopyroxene for garnet XCa=Ca/(Ca+Fe+Mg)=0.3. Peak P–T conditions for the albite–epidote–omphacite can be inferred to have involved T=610°C at P=14±1 kbar during D2A. Garnet (rim) and omphacite (matrix) compositions in sample WC21 give consistent results. Somewhat higher temperature estimates would be obtained using the Powell, (1985) or Ellis & Green, (1979) calibration of the garnet–clinopyroxene exchange equilibria. Temperature estimates made on the basis of garnet–clinopyroxene thermometry in the text below cite results after Krogh, (1988), but results for equivalent XCa in garnet and Kd after Ellis & Green, (1979) and Powell, (1985) are shown in Table 5. The composition of minerals forming the SA2 assemblage garnet (rim)–omphacite–phengite–paragonite–albite–quartz–epidote–glaucophane –rutile–titanite in sample WC21 were also used to obtain P–T estimates with Thermocalc. Average P–T calculations return a result with a satisfactory fit of P=14.3±3.4 and T=588±115 (both 2 error). Average P calculations estimate the SA2 conditions to be P=12.6±1.2 for T=550°C, whereas average T calculations estimate T=565±36°C for P=12.5 kbar. Calculations made using microprobe results of a similar mineral assemblage in sample 9403 give statistically identical results (Table 5), and the results of the Thermocalc calculations made on both rocks agree well with the P–T estimates made above on the basis of garnet–clinopyroxene thermometry and albite-quartz-clinopyroxene barometry (Fig. 10b). Similar P–T conditions were inferred for samples taken from this zone by Ghent & Stout, (1994).
|
The prograde path of the Pouébo terrane passed through the blueschist facies, on the basis of lawsonite and chlorite inclusions in garnet of Type II eclogite, and glaucophane inclusions in garnet of Type I eclogite that lacks SE2 glaucophane. Peak conditions that accompanied the development of SE2 in Type II eclogite involved temperatures of 550–600°C on the basis of garnet–clinopyroxene and garnet-amphibole (after Graham & Powell, 1984
) thermometry (Table 5). Owing to the textural complexity of Type II eclogite, analyses from adjacent grains were used to obtain temperature estimates. Part of the prograde path that led to Type I eclogite can be inferred by pairing the compositions of inclusion omphacite with adjacent garnet, and the compositions of matrix omphacite with adjacent garnet rims. Caution is required owing to the high spessartine content in most garnet cores, the significant aegirine content in some clinopyroxene grains, and the clinopyroxene being omphacite. Where garnet core, shoulder or rim compositions can be paired with appropriate omphacite grains, there is a systematic decrease in Kd value in moving from inferred core equilibria formed before or during SE1 to matrix equilibria formed before or during SE3. Sample WC68, which has aegirine–poor omphacite present as inclusions in garnet and as large zoned matrix grains, is a good example to illustrate these variations in temperature. A clinopyroxene inclusion in the inclusion-poor domain of a garnet grain is itself zoned from Jd43 to Jd40 with only a subtle change in XFe. The surrounding inclusion-poor zone of garnet is characterized by a 'shoulder’ in the zoning profile of the grain (Fig. 8a). If these features are assumed to record growth zoning, the paired garnet shoulder–omphacite core compositions give Kd=15.1, compared with Kd=13.4 for the garnet shoulder–omphacite inclusion rim compositions. An identical result of Kd=13.3 is given by pairing the garnet shoulder and matrix core omphacite compositions, and compositions of garnet rim–matrix omphacite rim yield Kd=9.0. These variations are consistent with increasing temperature conditions through a spectrum of 100°C, ranging from 580°C for the post-SE1 garnet shoulder–clinopyroxene inclusion to 694°C for the garnet rim–SE3 omphacite. This final estimate is high in comparison with results from other Type I eclogites, and reflects the comparatively pyrope-rich rim of garnet in sample WC68. Obtaining consistent temperature estimates on the basis of garnet–clinopyroxene thermometry in Type I eclogite with aegirine-rich omphacite is more problematic. The pattern of decreasing Kd established for aegirine-poor–omphacite-bearing assemblages can only be consistently reproduced if the ferric iron component of the clinopyroxene, calculated on the basis of four cations per six oxygens, is removed and the ferrous component rather than total Fe used in the calculation. This can drop temperature estimates made on the basis of garnet–clinopyroxene equilibria by more than 250°C. For example, an exceptionally high temperature estimate of 780°C on the basis of garnet rim and matrix SE2 clinopyroxene compositions in sample WC11 is reduced to 480°C when only ferrous iron is used in the calculation. In summary, the systematic decrease in Kd values for the garnet–clinopyroxene exchange equilibria most probably reflects the prograde growth of garnet, and the core to rim T estimates overlap the lawsonite breakdown as defined by Chatterjee et al., (1984) in a manner that is consistent with the textural information provided by the Type II eclogites. However, owing to the significant variation in spessartine content of the garnets involved, and the clinopyroxene being omphacite, caution is required in using the T estimates in more than a general manner.
Pressure estimates for conditions that accompanied the development of SE3 in Type I eclogite can be made using the end-member activities for the inferred assemblage garnet(rim)–omphacite(matrix rim)–epidote–phengite–quartz–water and Thermocalc. End-member activities were calculated using the computer program AX assuming a temperature of 550°C. Average T calculations made using mineral compositions in sample 9341a and 9341b and assuming P=25 kbar give results with a good fit of T=478±90°C and T=517±96°C, respectively. Average P calculations for 9341a and 9341b give results P=25.6±3.2 kbar for T=480°C, and P=23.9±3.0 kbar for T=517°C (Table 4), respectively. Similar results can be obtained using the compositions of the inferred SE3 mineral assemblages in samples WC11 and WC68 (Table 5). The temperature estimates seem too low; though within 2 error of the results obtained on the basis of garnet–amphibole and garnet–clinopyroxene thermometry after Krogh, (1988), they lie in the lawsonite stability field when the lawsonite breakdown reaction is calculated by Thermocalc (Fig. 10b). Whereas running average P calculations for higher temperatures results in no change in the estimated P conditions, there is a rapid expansion of the 2 error envelope (Fig. 10b). On the basis of textural information provided by the Type II eclogites, P–T conditions that accompanied the development of SE3 lay at higher T than lawsonite stability, and at lower P than paragonite breakdown. Thus the most realistic quantitative estimates using sample 9341b as an example (Fig. 10b) are P=23.8±4.0 kbar for T=600°C. Average P–T calculations have not been made on the Type II eclogites; owing to the lack of a penetrative foliation and presence of reaction textures, it is difficult to justify the interpretation of a suitable 'stable’ equilibrium mineral assemblage. None the less, there is good agreement between the temperature estimates obtained for the two types of eclogite, with conditions that accompanied the impersistent development of SE2-SE3 in Type II eclogite having been 50°C cooler than contemporary conditions in Type I eclogite (Table 5).
The use of quantitative methods in constraining P–T conditions witnessed by the glaucophanites is hampered by difficulties in inferring a 'stable’ mineral assemblage, mostly owing to the glaucophanites involving incompletely recrystallized Type I eclogite. However, samples 9342 and RS103 are inferred to have been pervasively recrystallized during the post-SE3 blueschist hydration, only preserving eclogite remnants as inclusions in garnet. Sample 9342 can be used to constrain the change in conditions from eclogite to glaucophanite, because omphacite occurs as inclusions in garnet and as idioblastic grains intergrown with post-SE3 glaucophane and epidote. Omphacite splays in the crossite–rich matrix are aegirine rich and systematic variations in Kd values for garnet–clinopyroxene thermometry need to be established using only ferrous iron. Following this procedure, there is again a systematic decrease in Kd values from paired garnet core–omphacite inclusion that give Kd=11.5, to garnet shoulder–matrix omphacite core that give Kd=9.74, to garnet rim–matrix omphacite rim that give Kd=7.21. This would broadly correspond to a temperature increase over the range of 150°C, with the highest temperature inferred for the post-SE3 glaucophanite assemblage. Average P–T calculations made using the inferred assemblage garnet–omphacite–crossite–epidote–phengite–quartz give a result with a good fit and conditions involving P=23.3±5.0 kbar and T= 619±137°C. For reasons outlined above, a ‘best fit’ result would involve the lower-P, higher-T section of the error ellipse (Fig. 10b). A similar glaucophanite assemblage involving zoned idioblastic omphacite occurs in sample RS103, and a similar pattern involving decreasing Kd values can be established using what are interpreted to be progressively younger pairs of garnet and clinopyroxene analyses (Table 5). Results from both the garnet–clinopyroxene thermometry and Thermocalc are consistent with the change from eclogite to glaucophanite assemblages having involved similar, or increasing temperature conditions.
Qualitative indications that the recrystallization of Type I eclogite to glaucophanite involved temperatures somewhat higher than the eclogite event involve garnet in the glaucophanites being commonly more pyrope rich than SE3 garnet (Table 2) and having lost growth zoning. Most glaucophanite retains some record of having witnessed the eclogite event in the form of omphacite inclusions in garnet or garnet with growth zoning identical to that in eclogite. However, large grains of garnet (>2 cm in diameter) in glaucophanite, which contain coarse-grained inclusions of epidote and glaucophane, commonly have chemical zoning significantly reduced in amplitude in comparison with similar garnet in eclogite or may have no zoning at all (samples 9337a, 9337b, RS31, RS26ab, Table 2). The simplest explanation to account for these large, unzoned garnet grains involves temperatures having been high enough for comparatively rapid intragranular diffusion to have erased the growth zoning (e.g. Ghent, 1988). Whereas SE2-SE3 glaucophane in Type II eclogite has a very restricted compositional range (Fig. 8b), crossite in glaucophanite shows a trend involving decreasing XNa and increasing tetrahedral Al content such that it approaches the Si-XNa compositions of the hornblende series. It is tempting to argue that this trend reflects conditions for the glaucophanite recrystallization having been higher than where the two amphibole series are clearly separated and having almost exceeded those appropriate for the two-amphibole solvus inferred by Reynard & Ballèvre, (1988). However, the two series remain distinct when tetrahedral alumina is plotted against either octahedral alumina or A-site sodium (after Robinson et al., 1982). It is difficult to obtain quantitative pressure estimates on most glaucophanite assemblages, owing to problems outlined above. However, some samples contain coexisting omphacite (Jd42–45) and albite (e.g. 9334) that represent conditions of P=14±1 kbar (after Holland, 1980) for T=650°C.
|
Discussion |
|---|
|
TOPABSTRACTINTRODUCTIONIntroductionRegional GeologyThe Pouebo terranePetrologyMineral ChemistryThermobarometryDiscussionReferences |
|---|
High-grade rocks of the Pam Peninsula can be divided into three tectonic units, on the basis of metamorphic grade and protolith. Metasedimentary and metavolcanic rocks of the Cretaceous to Eocene Diahot terrane form two of these units: (1) ferroglaucophane–lawsonite zone schists that show a gradational northeastwards increase in metamorphic grade [Yokoyama et al., (1986)
and references therein] to rocks that reflect conditions of P=7–9 kbar and T=400±58°C [see also Black (1973a) and Yokoyama et al., (1986)], and which are in fault contact with a narrow, discontinuous zone of (2) albite–epidote–omphacite zone schists that reflect conditions of P=12.6±1.2 kbar and T=570±36°C [see also Ghent & Stout, (1994)]. The Diahot terrane is tectonically underlain by eclogite facies metabasalts and metagabbros of the Pouébo terrane. On the basis of pre-SE1 lawsonite and glaucophane inclusions in eclogite facies garnets, the prograde path of the Pouébo terrane passed through the blueschist facies and involved rapid early changes in pressure. Pairing compositions of omphacite grains with that of adjacent garnet grains from Type I eclogite (after Ellis & Green, 1979) is consistent with the change in conditions that accompanied the shift from SE1 to SE3 having involved 100°C (Fig. 10a). Eclogite facies conditions that accompanied the development of SE3 involved P=23.8±4.0 kbar for T=600°C. Extensive post-SE3 hydration, oxidation and recrystallization occurred during prograde decompression of the terrane to T650°C and P=14±1 kbar. Eclogites exposed between Pouébo and Cape Colnett (Fig. 1) have similar mineral chemistry to the eclogites of the Pam Peninsula (Table 2), so there does not appear to be a large change in grade or protolith along the length of the Pouébo terrane. Glaucophanite recrystallization of the Pouébo terrane preceded its juxtaposition with the albite–epidote–omphacite zone, before or during the DA2-DE5 event (Table 1) at T650°C and P14 kbar. Both the Pouébo terrane and the albite–epidote–omphacite zone were then affected by deformation that produced the north-trending, steeply dipping SA3-SE6 foliation and foliation-parallel faults before being joined to the ferroglaucophane–lawsonite zone at P<8 kbar. All rocks are affected by the east- to south-east-trending SL3-SA4-SE7 foliation and foliation-parallel faults, movement along which resulted in tectonic windows of the Pouébo terrane near Ouégoa (Fig. 2a).
The three zones, defined largely on the basis of the field relationships involving grain size, protolith and inferred mineral assemblage, may additionally be characterized on the basis of mineral chemistry (Table 2). Amphibole is probably the mineral that shows the greatest variation; with increasing grade it changes from ferroglaucophane to glaucophane in the metasedimentary blueschists of the Diahot terrane (Black, 1973b), then to SE3 barroisite in Type I eclogite and comparatively Mg-rich, Fe3+-poor SE3 glaucophane with or without SE3 hornblende in Type II eclogite (Fig. 8). Though almost all garnet has bell-shaped zoning profiles involving a rimward decrease in spessartine content, the ferroglaucophane–lawsonite zone contains comparatively pyrope-poor, spessartine-rich garnet (Table 2; Black, 1973a). Whereas garnet in the albite–epidote–omphacite zone has cores with appreciably more spessartine and less grossular than eclogitic garnet cores (Fig. 6b), rim compositions of garnet in the albite–epidote–omphacite zone have similar grossular but significantly lower pyrope content than rim compositions of SE3 garnet (Table 2). Post-SE3 glaucophanite garnet approaches, but is commonly more pyrope rich than, the composition of garnet rims in the albite–epidote–omphacite zone. Ignoring aegirine content, omphacite has a restricted compositional range. Clinozoisite and epidote in Type II eclogite have significantly lower pistacite contents than the overlapping compositional range shown by epidote in the Type I eclogite and albite–epidote–omphacite zone (Table 2). This lower pistacite content, together with aegirine–poor omphacite and Fe3+-poor glaucophane, reflects the Fe2O3-poor gabbroic protolith of the Type II eclogite. Phengite shows a trend of increased tetrahedral substitution of Si by Al with the inferred increase in pressure. Though not clearly separated, phengite in the ferroglaucophane–lawsonite zone commonly has 6.9–7.0 Si p.f.u., compared with 6.7–6.9 Si p.f.u. in the albite–epidote–omphacite zone and 6.5–6.8 in eclogites and glaucophanites (Fig. 9).
Whereas P–T conditions inferred for the ferroglaucophane–lawsonite and albite–epidote–omphacite zones can be placed on a single geotherm involving temperature increasing at the rate of 15°C/km, the conditions inferred to have accompanied the development of SE3 in the Pouébo terrane involved a geotherm of the order of 7°C/km. The boundary between the Diahot and Pouébo terranes joins rocks that were in profoundly different sections of a subduction zone: a continental leading edge for Diahot terrane rocks, versus subducted oceanic material for the Pouébo terrane (Aitchison et al., 1995a). The comparatively cool geotherm recorded by SE3 assemblages in the Pouébo terrane could be explained by T conditions having been reduced by continued subduction. Water for the glaucophanite recrystallization could have been provided by dehydration of the underlying slab. Deep geophysical anomalies that currently overlie subducted portions of the Juan de Fuca plate in western Canada (Hyndmann et al., 1990; Kurtz et al., 1990; Calvert & Clowes, 1991) are thought to represent fluids migrating under conditions similar to those inferred for the recrystallization of the Pouébo terrane.
A clockwise P–T path can be inferred for the Pouébo terrane on the basis of early blueschist facies conditions (SE1 lawsonite), high-P eclogite facies conditions that accompanied the development of SE3, and prograde, post-SE3 decompression that accompanied the formation of the glaucophanites. This is consistent with previous interpretations (Bell & Brothers, 1985; Yokoyama et al., 1986) but developed from different criteria and assumes that the rocks were affected by only one metamorphic event. The only age constraint on the deformation sequence is a minimum age: on the basis of a 40Ar/39Ar study of white micas, Ghent et al., (1994) inferred that rocks from the Pam Peninsula above the epidote isograd cooled below 350°C by 37±1 Ma. Faults parallel to the SA3-SE6 and SL3-SA4-SE7 foliations, which involve displacements of the order of 100 m, form most contacts between the three metamorphic zones in the Pam Peninsula. However, these steeply dipping faults are comparatively small-scale structures on a regional anticline involving a core of Pouébo terrane overlain by tectonically disrupted, lower-pressure metamorphic zones of the Diahot terrane (Fig. 2b). The high-P conditions experienced by the Pouébo terrane cannot be accounted for by the currently thin veneer of Diahot terrane, with omissions of crustal section consistent with significant tectonic thinning during exhumation of the Pouébo terrane (e.g. Platt, 1986, 1993). Though the form of the Pouébo terrane may be consistent with it being a metamorphic core complex (Cluzel et al., 1995), the positive density contrast between it and Diahot terrane rocks makes it difficult for buoyancy forces to have emplaced it in the upper crust. Prograde decompression of the Pouébo terrane to mid-crustal levels would be consistent with it having been initially exhumed slowly, similar to high-P metamorphism consequent to continental collision (Selverstone et al., 1984; Dal Piaz, 1993; Platt, 1993; Searle et al., 1994) and possibly during continued convergence that was slowed as a consequence of arc–continent collision (Aitchison et al., 1995a). As an alternative explanation to the magmatically driven crustal extension inferred for many core complexes (Lister & Baldwin, 1993), late doming and uplift of the Pouébo and Diahot terranes could have been driven by processes similar to that suggested for similar rocks in NW America where high-density lower-crustal or upper-mantle material has been uplifted beneath southern Vancouver Island and the Juan de Fuca Strait (Dehler & Clowes, 1992). This is inferred to have happened in response to subduction stepping at 40 Ma and continued convergence that resulted in anticlinal buckling of the Crescent terrane, which is inferred to represent fossil subducted slab material.
|
Acknowledgements |
|---|
Many samples used in this work were collected by C. Gerakiteys, R. Davies, S. Gotley and W. Reid during the course of field work undertaken for their Honours year. R. Powell, C. Carson, M. C. Blake, Jr, S. Meffre and D. Nobes are thanked for discussion. R. W. White and C. Noccolds are thanked for their considerable help with the microprobe work. Thorough reviews by P. J. O'Brien and A. Barnicoat, and comments by D. Shelley and R. White improved the text. Field work was funded through Australian Research Council grants to G.L.C. and R. Powell (A39600827) and G.L.C. and J.C.A. (University of Sydney Institutional grant).
* Corresponding author Email: geoffc{at}ucc.s.oz.au
|
References |
|---|
|
TOPABSTRACTINTRODUCTIONIntroductionRegional GeologyThe Pouebo terranePetrologyMineral ChemistryThermobarometryDiscussionReferences |
|---|
Aitchison J. C., Clarke G. L., Cluzel D., Meffre S. Eocene arc–continent collision in New Caledonia and implications for regional SW Pacific tectonic evolution. Geology (1995a) 23:161–164.
Aitchison J. C., Meffre S., Cluzel D. Cretaceous/Tertiary radiolarians from New Caledonia. Geological Society of New Zealand, Miscellaneous Publication (1995b) 81A:70.
Bell T. H., Brothers R. N. Development of P–T prograde and P-retrograde, T-prograde isogradic surfaces during blueschist to eclogite regional deformation/metamorphism in New Caledonia, as indicated by progressively developed porphyroblast microstructures. Journal of Metamorphic Geology (1985) 3:59–78.[Web of Science]
Black P. M. Mineralogy of New Caledonian metamorphic rocks. I. Garnets from the Ouégoa district. Contributions to Mineralogy and Petrology (1973a) 38:221–235.[Web of Science]
Black P. M. Mineralogy of New Caledonian metamorphic rocks. II. Amphiboles from the Ouégoa district. Contributions to Mineralogy and Petrology (1973b) 39:55–64.[Web of Science]
Black P. M. Oxygen isotope study of metamorphic rocks from the Ouégoa district, New Caledonia. Contributions to Mineralogy and Petrology (1974) 47:197–206.[Web of Science]
Black P. M. Mineralogy of New Caledonia metamorphic rocks. IV. Sheet silicates from the Ouégoa district. Contributions to Mineralogy and Petrology (1975) 49:269–284.[Web of Science]
Black P. M. Regional high-pressure metamorphism in New Caledonia: phase equilibria in the Ouégoa district. Tectonophysics (1977) 43:89–107.[Web of Science]
Black P. M., Brothers R. N. Blueschist ophiolites in the melange zone, northern New Caledonia. Contributions to Mineralogy and Petrology (1977) 65:69–78.[Web of Science]
Black P. M., Brothers R. N., Yokoyama K. Mineral parageneses in eclogite-facies meta-acidites in northern New Caledonia. In: Eclogites and Eclogite-Facies Rocks—Smith D. C., ed. (1988) Amsterdam: Elsevier. 271–289.
Black P. M., Maurizot P., Ghent E. D., Stout M. Z. Mg–Fe carpholites from aluminous schists in the Diahot region and implications for preservation of high pressure–low temperature schists, northern New Caledonia. Journal of Metamorphic Geology (1993) 11:455–460.[Web of Science]
Briggs P. M., Lillie A. R., Brothers R. N. Structure and high-pressure metamorphism in the Diahot region, northern New Caledonia. Bulletin du BRGM (1978) 2:171–189.
Brothers R. N. High-pressure schists in northern New Caledonia. Contributions to Mineralogy and Petrology (1974) 46:109–127.[Web of Science]
Brothers R. N. Regional mid-Tertiary blueschist–eclogite metamorphism in northern New Caledonia. Géologie de la France (1985) 1985:37–44.
Brothers R. N., Blake M. C. Tertiary plate tectonics and high-pressure metamorphism in New Caledonia. Tectonophysics (1972) 17:359–391.[Web of Science]
Brothers R. N., Yokoyama K. Comparison of high-pressure schist belts of New Caledonia and Sanbagawa. Contributions to Mineralogy and Petrology (1982) 79:219–229.[Web of Science]
Calvert A. J., Clowes R. M. Seismic evidence for the migration of fluids within the accretionary complex of western Canada. Canadian Journal of Earth Sciences (1991) 28:542–556.
Carroué J. P. Carte géologique à l'échelle du 1/50000, feuille Pouébo. 12. Paris: Bureau des Recherches Géologiques et Minières, map sheet and explanatory notes.
Carswell D. A. Eclogites and the eclogite facies: definitions and classification. In: Eclogite Facies Rocks—Carswell D. A., ed. (1990) London: Blackie. 1–13.
Chatterjee N. D., Johannes W., Leistner H. The system CaO-Al2O3-SiO2-H2O: new phase equilibria data, some calculated phase relations, and their petrological applications. Contributions to Mineralogy and Petrology (1984) 88:1–13.[Web of Science]
Cluzel D., Aitchison J. C., Clarke G. L, Meffre S., Picard C. Point de vue sur l'évolution tectonique et géodynamique de la Nouvelle-Calédonie. Comptes Rendus Hebdomadaires des Séances de l'Académie des Sciences, Série II (1994) 319:683–690.
Cluzel D., Aitchison J. C., Clarke G. L., Meffre S., Picard C. Dénudation tectonique du complexe à noyau métamorphique de haute pression Tertiaire du Nord de la Nouvelle-Calédonie (Pacifique, France) données cinématique. Comptes Rendus Hebdomadaires des Séances de l'Académie des Sciences, Série II (1995) 321:57–64.
Coleman R. G., Lee D. E., Beatty L. B., Brannock W. W. Eclogites and eclogites: their differences and similarities. Geological Society of America Bulletin (1965) 76:483–507.
Dal Piaz D. V. Evolution of Austro-Alpine and Upper Penninic basement in the northwestern Alps from Variscan convergence to post-Variscan extension. In: Pre-Mesozoic Geology of the Alps—von Raumer J. F., Neubauer F., eds. (1993) Heidelberg: Springer-Verlag. 327–344.
Davies H. L., Warren R. G. Eclogites of the D'Entrecasteaux Islands. Contributions to Mineralogy and Petrology (1988) 112:463–474.
Dehler S. A., Clowes R. M. Integrated geophysical modelling of terranes and other structural features along the western Canadian margin. Canadian Journal of Earth Sciences (1992) 29:1492–1508.
Ellis D. J., Green D. H. An experimental study of the effect of Ca upon garnet–clinopyroxene Fe–Mg exchange equilibria. Contributions to Mineralogy and Petrology (1979) 71:13–22.[Web of Science]
Ghent E. D. A review of chemical zoning in eclogite garnets. In: Eclogites and Eclogite-Facies Rocks. Development in Petrology 12—Smith D. C., ed. (1988) Amsterdam: Elsevier. 207–236.
Ghent E. D., Stout M. Z. Geobarometry of low-temperature eclogites: applications of isothermal pressure–activity calculations. Contributions to Mineralogy and Petrology (1994) 116:500–507.[Web of Science]
Ghent E. D., Black P. M., Brothers R. N., Stout M. Z. Eclogite and associated albite–epidote–garnet paragneisses between Yambe and Cape Colnett, New Caledonia. Journal of Petrology (1987a) 28:627–643.
Ghent E. D., Stout M. Z., Black P. M., Brothers R. N. Chloritoid-bearing rocks associated with blueshists and eclogites, northern New Caledonia. Journal of Metamorphic Geology (1987b) 5:239–254.[Web of Science]
Ghent E. D., Roddick J. C., Black P. M. 40Ar/39Ar dating of white micas from the epidote to omphacite zones, northern New Caledonia: tectonic implications. Canadian Journal of Earth Science (1994) 31:995–1001.
Graham C. M., Powell R. A garnet–hornblende geothermometer: calibration, testing and application to the Pelona Schist, Southern California. Journal of Metamorphic Geology (1984) 2:13–31.[Web of Science]
Hill E. J., Baldwin S. L., Lister G. S. Unroofing of active metamorphic core complexes in the D'Entrecasteaux Islands, Papua New Guinea. Geology (1992) 20:907–910.
Holland T. J. B. The reaction albite=jadeite+quartz determined experimentally in the range 600–1200°C. American Mineralogist (1980) 65:129–134.[Abstract]
Holland T. J. B., Powell R. An enlarged and updated internally consistent dataset with uncertainties and correlations: the system K2O-Na2O-CaO-MgO-MnO-FeO-Fe2O3-Al2O3-TiO2-SiO2-C-H2-O2. Journal of Metamorphic Geology (1990) 8:89–124.[Web of Science]
Hyndmann R. D., Yorath C. J., Clowes R. M., Davies E. E. The structure and tectonic history of the northern Cascadia subduction zone at Vancouver Island. Canadian Journal of Earth Sciences (1990) 27:313–329.
Krogh E. J. The garnet–clinopyroxene geothermometer-a reinterpretation of existing experimental data. Contributions to Mineralogy and Petrology (1988) 99:44–48.[Web of Science]
Kurtz R. D., Delaurier J. M., Gupta J. C. The electrical conductivity distribution beneath Vancouver Island: a region of active plate subduction. Journal of Geophysical Research (1990) 95:10929–10946.
Leake B. E. Nomenclature of amphiboles. Mineralogical Magazine (1978) 42:533–563.[Web of Science]
Lillie A. R. Structures in the lawsonite-glaucophane schists of New Caledonia. Geological Magazine (1975) 112:225–340.[Abstract]
Lister G. S., Baldwin S. L. Plutonism and the origin of metamorphic core complexes. Geology (1993) 21:607–610.
Maurizot P., Eberlé J-M., Habault C., Tessarollo C. Carte géologique à l'échelle du 1/50000, feuille Pam-Ouégoa. (1989) 81. Paris: Bureau des Recherches Géologiques et Minières, map sheet and explanatory notes.
Morimoto N. Nomenclature of pyroxenes. Mineralogical Magazine (1988) 52:535–550.[Web of Science]
Mottana A., Carswell D. A., Chopin C., Oberhansli R. Eclogite facies mineral parageneses. In: Eclogite Facies Rocks—Carswell D. A., ed. (1990) London: Blackie. 14–52.
Paris J. P. Géologie de la Nouvelle-Calédonie. Mémoires du Bureau des Recherches Géologiques et Minières (1981) 113:279.
Platt J. P. Dynamics of orogenic wedges and the uplift of high-pressure metamorphic rocks. Geological Society of America Bulletin (1986) 97:1037–1053.
Platt J. P. Exhumation of high-pressure rocks: a review of concepts and processes. Terra Nova (1993) 5:119–133.[Web of Science]
Powell R. Regression diagnostics and robust regression in geothermometer/geobarometer calibration: the garnet–clinopyroxene thermometer revisited. Journal of Metamorphic Geology (1985) 3:327–342.[Web of Science]
Powell R., Holland T. J. B. An internally consistent dataset with uncertainties and correlations: 3. Applications to geobarometry, worked examples and a computer program. Journal of Metamorphic Geology (1988) 6:173–204.[Web of Science]
Prinzhofer A. Structure et pétrologie d'un cortège ophiolitique: le massif du sud (nouvelle Calédonie); la transition manteau–croûte en milieu océanique. (1981) 185. Ph.D. thesis, L'École Nationale Supérieure des Mines de Paris.
Reid W. Characterisation of the protolith and peak metamorphic conditions of eclogites in the Pam Peninsula, NE New Caledonia. (1994) 73. Unpublished B.Sc.(Hons) Thesis, University of Sydney, N.S.W.
Reynard B., Ballèvre M. Coexisting amphiboles in an eclogite from the western Alps: new constraints on the miscibility gap between sodic and calcic amphiboles. Journal of Metamorphic Geology (1988) 6:333–350.[Web of Science]
Robinson P., Spear F. S., Schumacher J. C., Laird J., Klein C., Evans B. W., Doolan B. L. Phase relations of metamorphic amphiboles: natural occurrence and theory. In: Mineralogical Society of America, Reviews in Mineralogy—Veblen D. R., Ribbe P. H., eds. (1982) 9B:1–227.[Medline]
Searle M. P., Waters D. J., Martin H. N., Rex D. C. Structure and metamorphism of blueschist–eclogite facies rocks from the northeastern Oman Mountains. Journal of the Geological Society, London (1994) 151:555–576.
Selverstone J., Spear F. S., Franz G., Morteani G. High-pressure metamorphism in the SW Tauern Window, Austria: P–T paths from hornblende–kyanite–staurolite schists. Journal of Petrology (1984) 25:501–531.
Yokoyama K., Brothers R. N., Black P. M. Regional facies in the high-pressure metamorphic belt of New Caledonia. Geological Society of America, Memoir (1986) 184:407–423.
CiteULike 
Connotea 
Del.icio.us What's this?
This article has been cited by other articles:
|
|
 |
|
  M. Brown Metamorphic patterns in orogenic systems and the geological record Geological Society, London, Special Publications, January 1, 2009; 318(1): 37 - 74. [Abstract] [Full Text] [PDF]
|
|
|
|
 |
|
 S. L. Baldwin, T. Rawling, and P. G. Fitzgerald Thermochronology of the New Caledonian high-pressure terrane: Implications for middle Tertiary plate boundary processes in the southwest Pacific Geological Society of America Special Papers, January 1, 2007; 419(0): 117 - 134. [Abstract] [Full Text] [PDF]
|
|
|
|
 |
|
 J. F. Molina, J. F. Molina, and S. Poli Singular Equilibria in Paragonite Blueschists, Amphibolites and Eclogites J. Petrology, July 1, 1998; 39(7): 1325 - 1346. [Abstract] [Full Text] [PDF]
|
|
| ||||||||||||||||||||||||||||||||||||||||||||||||||