Ediacaran high-pressure collision metamorphism and tectonics of the southern Ribeira Belt (SE Brazil): Evidence for terrane accretion and dispersion during Gondwana assembly

Precambrian Research 189 (2011) 263–291

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Ediacaran high-pressure collision metamorphism and tectonics of the southern Ribeira Belt (SE Brazil): Evidence for terrane accretion and dispersion during Gondwana assembly Frederico Meira Faleiros a,b,∗ , Ginaldo Ademar da Cruz Campanha b , Lucelene Martins b , Silvio Roberto Farias Vlach b , Paulo M. Vasconcelos c a

CPRM – Geological Survey of Brazil, Rua Costa 55, Consolac¸ão, CEP 01304-010, São Paulo, SP, Brazil Departamento de Mineralogia e Geotectônica, Instituto de Geociências, Universidade de São Paulo, Rua do Lago 562, CEP 05508-900, São Paulo, SP, Brazil c University of Queensland, Department of Earth Sciences, Brisbane, Qld., 4072 Australia b

a r t i c l e

i n f o

Article history: Received 1 June 2010 Received in revised form 15 June 2011 Accepted 20 July 2011 Available online 30 July 2011 Keywords: Ribeira Belt Collision metamorphism P–T–t path Nappe stacking Strike-slip dispersion Kyanite migmatites

a b s t r a c t The Curitiba Terrane represents a major segment of the southern Ribeira Belt (SE Brazil), which was derived from the collision between the São Francisco, Congo, Paranapanema and Luís Alves Cratons during the Neoproterozoic (Brasiliano/Pan-African Orogeny). The tectonic setting and the metamorphic records of two major juxtaposed units from the Curitiba Terrane, a Neoproterozoic shallow continental-shelf metasedimentary assemblage (Turvo-Cajati Formation) and an Archaean to Paleoproterozoic TTG-type orthogneiss assemblage (Atuba Complex), were investigated. Migmatitic paragneisses from the TurvoCajati Formation underwent a deep collision metamorphism. Conventional geothermobarometry and petrological modelling in the system NCKFMASHTi indicate peak metamorphic conditions between 670 and 810 ◦ C at 9.5–12 kbar. Metamorphic paths calculated from zoned garnet and plagioclase using the Gibbs method of differential thermodynamics indicate distinct evolution for two major groups of migmatites from the Turvo-Cajati Formation: (i) kyanite migmatites evolved from low-temperature eclogite to high-pressure granulite facies conditions following near isobaric heating; (ii) sillimanite migmatites underwent near isothermal decompression and apparently evolved from high-temperature eclogite facies conditions. Chemical dating of monazite indicates that the peak metamorphism of the Turvo-Cajati Formation occurred at 589 ± 12 Ma, followed by a greenschist facies metamorphic overprint at 579 ± 8 Ma related with late transcurrent shear zones. 40 Ar–39 Ar biotite ages indicate that the TurvoCajati Formation cooled below 250–300 ◦ C at 555 ± 4 Ma. P–T data and petrological evidence of rocks from the Atuba Complex suggest a retrograde metamorphic path with cooling from 730 to 630–650 ◦ C at 6–7 kbar. Available K–Ar and 40 Ar–39 Ar data indicate that the Atuba Complex had cooled to below 300–500 ◦ C between ca. 590 and 580 Ma. Geochronological data indicate that the main metamorphism of the Turvo-Cajati Formation and the Atuba Complex are coeval, but very contrasting metamorphic signatures reflect formation in different parts of a collisional suture. The integration of structural and petrological data indicates that the structural pattern of the Curitiba Terrane is related to Ediacaran westward directioned nappes during the late- to postmetamorphic period. This is concomitant with a main, crustal-scale, strike-slip regime, dominant throughout the Ribeira Belt. The nappe stack was later deformed by cylindrical folds with E–W trending sub-horizontal axes parallel to the synthrusting stretching lineation and was dismembered and dispersed by late sinistral strike-slip shear zones. The late tectonic assembly of the Ribeira Belt was controlled by significant postcollision terrane dispersion along major strike-slip shear zones. © 2011 Elsevier B.V. All rights reserved.

1. Introduction

∗ Corresponding author at: CPRM – Geological Survey of Brazil, Rua Costa 55, Consolac¸ão, CEP 01304-010 São Paulo, São Paulo, Brazil. Tel.: +55 11 37755119; fax: +55 11 32568430. E-mail addresses: [email protected], [email protected] (F.M. Faleiros). 0301-9268/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.precamres.2011.07.013

The Ribeira Belt in southeastern South America is a NE-trending collisional orogen related to the assembly of West Gondwana during the Neoproterozoic (Brito Neves et al., 1999). The orogen strikes subparallel to the coastline of southeastern Brazil and has counterparts in the Kaoko and West Congo Belts in western Africa (Almeida et al., 1973) (Fig. 1). The structure of the Ribeira Belt

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Fig. 1. Regional geological context of southeastern Brazil and the position of West Africa at Gondwana times, showing the relationships between the main cratonic units and fold belts.

was controlled by a late, orogen-parallel, crustal-scale transcurrent shear system related to an oblique collision between the Paranapanema, São Francisco, Luís Alves and Congo Cratons (e.g., Dürr and Dingeldey, 1996; Campanha and Brito Neves, 2004; Faleiros et al., 2010). As a result, the Ribeira Belt is dominated by faultbounded terranes of diverse origins, including reworked blocks of allochthonous Archaean to Paleoproterozoic basement rocks, Mesoproterozoic and Neoproterozoic suites of volcanosedimentary rocks, intrusive granitoids that are interpreted as arc assemblages, passive margin sedimentary sequences and late kinematic granites (e.g., Campos Neto and Figueiredo, 1995; Campanha and Sadowski, 1999; Campos Neto, 2000; Heilbron et al., 2004). Since the late 1980s, the southern Ribeira Belt has been interpreted in the light of plate tectonics concepts (e.g., Soares, 1987; Campanha et al., 1987; Campanha, 1991; Reis Neto, 1994; Siga Junior, 1995; Fassbinder, 1996; Campanha and Sadowski, 1999), based on the assumption that the observed spatial relationships between the juxtaposed terranes represent genetic relationships (i.e., partially preserved paleogeography). However, this interpretation is contradicted by recent geochronological and isotopic data (Hackspacher et al., 2000; Harara, 2001; Basei et al., 2008; Siga Junior et al., 2009, 2011a,b; Weber et al., 2004; Campanha et al., 2008a,b) which show a tectonic configuration that cannot be explained by reconstructions based on simple events of opening and inversion of oceanic basins, or by a single Wilson Cycle. The discrepant geological and geochronological histories for the juxtaposed terranes suggest that the Ribeira Belt consists of a complex tectonic collage of several suspect and exotic terranes (cf., Coney et al., 1980; Schermer et al., 1984; Howell, 1995). Detailed studies of the metamorphic history (P–T–t evolution) of individual terranes can provide valuable information about the orogenic cycle independent of any stratigraphic or structural models (Spear et al., 2008); essential when studying intensely deformed/metamorphic terranes where the primary rock characteristics were strongly overprinted. The metamorphic history of

the terranes present in the southern Ribeira Belt has been understudied, yet it is a critical issue in evaluating the different tectonic models. This contribution uses detailed structural and petrological analyses combined with geochronological data to constrain the tectonic setting and evolution of the Curitiba Domain (Terrane), a major segment of the southern Ribeira Belt in SE Brazil. We present the first detailed description and P–T–t quantification of the main collisional metamorphic event of the southern Ribeira Belt, including the first discovery of high-pressure granulites in the Ribeira Belt. These data contribute to our understanding of the tectonic evolution of the Ribeira Belt, with implications for the understanding of the development of thickened crust in ancient collisional orogens. 2. Geological setting The southern Ribeira Belt can be divided into two main tectonic segments bounded by major strike-slip shear zones: the Apiaí and the Curitiba Domains (or composite terranes; cf. Howell, 1995; referred to below simply as terranes) (Fig. 2). The Curitiba and the Apiaí Terranes are separated by the Lancinha-Cubatão Fault Zone, which is the surface trace of a boundary between tectonic plates (e.g., Basei et al., 2008). The Paranapanema Craton, one of the main cratonic references to the southern Ribeira Belt, is completely covered by Phanerozoic sedimentary rocks from the Paraná Basin (Fig. 1) and was recognised by data from deep boreholes and geophysical studies (Mantovani and Brito Neves, 2005). The Luís Alves Terrane (Fig. 1) is an exposed small cratonic remnant that occurs south of the Curitiba Terrane and is composed of Archaean–Paleoproterozoic (ca. 2.7–2.0 Ga.; Basei et al., 1998; Hartmann et al., 2000) mafic and leucocratic granulitic gneisses. K–Ar hornblende and biotite ages between 2100 and 1700 Ma indicate that the Luís Alves Terrane was stable during the Brasiliano/Pan-African Orogeny (Siga Junior, 1995).

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Fig. 2. Simplified geological map of the southern Ribeira Belt and localisation of areas of detailed studies (modified after Faleiros, 2008) (see localisation on Fig. 1). The ages in the legend refer to crystallisation ages for the basement units (Tigre Gneiss, Atuba Complex and Serra Negra Complex) and granitic rocks, and depositional ages for metasedimentary units.

The Apiaí Terrane (Figs. 1 and 2) was formed by the amalgamation of Calymmian (ca. 1450–1500 Ma, U–Pb zircon ages; Weber et al., 2004; Campanha et al., 2008b; Siga Junior et al., 2011a,b), Tonian (ca. 900–1000 Ma, U–Pb zircon ages; Siga Junior et al., 2009) and Ediacaran (ca. 630–580 Ma, U–Pb zircon ages; Hackspacher et al., 2000; Siga Junior et al., 2009; Campanha et al., 2008a) terranes. These terranes include remnants of back-arc basin assemblages, carbonate shelf deposits, turbidite with flysch-like slump deposits, and deeper water assemblages associated with

mafic magmatism (Campanha and Sadowski, 1999). Rhyacian and Statherian basement rocks occur in the cores of broad antiforms (Cury et al., 2002). The metasedimentary units of the Apiaí Terrane (Fig. 2), display a Barrovian metamorphism varying from lowergreenschist to middle amphibolites facies, with clockwise P–T paths and peak pressures of approximately 8 kbar (Faleiros et al., 2010). In its eastern portion the Apiaí Terrane is in tectonic contat with the Embu Terrane (Fig. 1), a major tectonic unit dominated by metasedimentary rock assemblages (micaschists and migmatitic

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paragneisses) of unknown ages of sedimentation. The Embu Terrane presents a Cryogenian granitic magmatism (811 ± 13 Ma, U–Pb zircon age; Cordani et al., 2002) and a main metamorphism dated at 789 ± 10 Ma (U–Th–PbT chemical monazite age; Vlach, 2001). The Curitiba Terrane is primarily composed of a TTG-type migmatitic orthogneiss suite (Atuba Complex) and Neoproterozoic shallow continental-shelf metasedimentary assemblages (TurvoCajati and Capiru Formations) (Fig. 2). The Atuba Complex constitutes a microcontinent fragment that was formed during the Archaean (3.1–2.7 Ga) and was reworked primarily during the following three time periods: (i) a main Rhyacian (ca. 2.1 Ga, U–Pb ages) high-grade metamorphic event related to the amalgamation of the Atlantica Supercontinent; (ii) a Statherian (ca. 1.7 Ga, U–Pb ages) event related to the break-up of Atlantica; and (iii) an Ediacaran (ca. 630–590 Ma, U–Pb, Rb–Sr, Sm–Nd and Ar–Ar dates) high-grade metamorphic event associated with the assembly of West Gondwana (Sato et al., 2003, 2009; Machado et al., 2007). U–Pb ages of detrital zircons from the Turvo-Cajati Formation indicate a maximum age of sedimentation of ca. 900 Ma (Campanha et al., 2009).

3. Structure The Curitiba Terrane is limited by two orogen-parallel transcurrent shear zones of opposite kinematics (Figs. 1 and 2): the Lancinha-Cubatão Fault Zone (dextral) in the north and the Serra do Azeite Shear Zone (sinistral) in the south. Both shear zones truncate all the other structures present in the rocks of the Curitiba Terrane (Sections 3.1 and 3.2) and they were responsible for greenschist facies mylonitic overprint in previous higher-temperature fabrics, corroborating that these shear zones represent a late stage deformation in the southern Ribeira Belt. The Lancinha-Cubatão Fault Zone is the main segment of a crustal-scale dextral shear system, which extends for about 2100 km (800 km exposed and 1300 km covered by the Paraná Basin) in southeastern Brazil (Sadowski, 1991). In the study area the Lancinha-Cubatão Fault Zone comprises a NE-trending vertical zone with a width of approximately 200 m with mylonites and cataclasites. Recent zircon U–Pb geochronology (Weber et al., 2004; Basei et al., 2008; Campanha et al., 2009; Siga Junior et al., 2011a,b) indicates that the Lancinha-Cubatão Fault Zone represents an important boundary between regional geochronological domains, separating terranes dominated by Mesoproterozoic metasedimentary units in the north and Neoproterozoic metasedimentary units in the south. The Serra do Azeite Shear Zone comprises a domain with 1–2 km of width, with a subvertical ENE-trending mylonitic foliation and an associated ENE-plunging sub-horizontal stretching lineation. A greenschist mylonitic fabric obliterates previous higher-temperature fabrics of the Curitiba and Luís Alves Terranes. The Curitiba Terrane was internally deformed by an anastomosing network of greenschist facies sinistral strike-slip shear zones that are parallel to the Serra do Azeite Shear Zone and also obliterated previous fabrics. The internal units of the Curitiba Terrane in the study area are composed of strongly deformed and sheared metasedimentary rocks of the Turvo-Cajati Formation (TCF) and mylonitic orthogneisses of the Atuba Complex (AC). The TCF presents two main metamorphic units: a higher-grade unit (Higher-TCF) composed of mylonitic paragneiss and a lower-grade unit (Lower-TCF) composed of mylonitic micaschist, phyllite and slate (Figs. 3 and 4). A regional ENE-trending mylonitic foliation (Sm ) and an associated ESE-plunging sub-horizontal stretching lineation (Lm ) are present.

However, the foliation trajectories vary significantly in restricted areas. Detailed structural studies were performed on two areas: Cajati and Barra do Turvo (Figs. 3 and 4). 3.1. The Cajati area The AC and the TCF in the Cajati area show parallel structures (Sm , Lm , fold elements) (Fig. 3). Petrographic evidence indicates that the Sm and the Lm were developed concomitantly with the peak metamorphic conditions in the TCF units. In the AC the Sm and the Lm represent a high-temperature fabric, but frequently it was replaced by a greenschist facies mylonitic fabric. The macrostructure is composed of overturned folds with an ESE-plunging sub-horizontal fold axis (Fig. 3a, b). These form a north-verging asymmetrical anticlinorium where the AC overlies the Higher-TCF (a stratigraphic inversion since the AC is older) and the Lower-TCF overlies the AC (Fig. 3b). Along the contacts between these units a greenschist facies mylonitic foliation replaces a higher-grade mylonitic foliation of the same orientation. A metamorphic inversion occurs in the Higher-TCF, where sillimanite paragneisses overlies kyanite–staurolite-bearing micaschists (Fig. 3b). In stereographic projection, the folded surface (Sm ) of the AC distributes along a great circle girdle, suggesting cylindrical folding along an estimated axis S20◦ E/20◦ (Fig. 3c). Outcrop and hand-sample scale parasitic folds with the same pattern of orientation are present (Fig. 3g). The orientation of the foliation of the TCF is scattered (Fig. 3e). However, the data are coherent with the fold pattern observed in the AC, primarily for outcrop-scale parasitic folds (Fig. 3h). The Lm of both units shows a very consistent orientation parallel to the fold hinges (Fig. 3d, f). The mylonitic rocks from the TCF and AC show several types of shear sense indicators (Fig. 5): domino-type and shear bandtype asymmetric boudins (Passchier and Trouw, 2005), sigmoid objects, SC and SC’ shear band cleavage (e.g., Passchier and Trouw, 2005), SC1 fabrics (Ramsay and Lisle, 2000), ␴-type and ␦-type porphyroclasts with asymmetric strain shadows (e.g., Passchier and Trouw, 2005), and asymmetric quartz c-axis fabrics. On the map, the shear sense indicators show a domainal distribution that is demonstrated by the following (Fig. 3a): (a) top-to-the-WNW thrusting (Fig. 5a–d) in normal limbs (southern portion of the anticlinorium) and (b) an apparent sinistral strike-slip shear with a top-to-the-SE extensional component (Fig. 5e–g) in inverted limbs (northern portion of the anticlinorium). The distribution associated with the fold geometry and the parallelism between the fold axis and stretching lineation indicates that the shear sense indicators were inverted in the inverted limbs of the anticlinorium. 3.2. The Barra do Turvo area In this area the Putunã Shear Zone (Fig. 4a), a greenschist facies sinistral transcurrent fault, separates two domains, both characterized by the occurrence of the Higher-TCF in the south and the Lower-TCF in the north of each segment. The AC outcrops in the northern segment in the core of a regional open antiform (Fig. 4a), forming a structure like a gneiss dome. The AC and the TCF rocks exhibit contrasting structural orientations (Fig. 4a–e). The TCF displays a main ENE-trending subvertical Sm and an EW-trending sub-horizontal Lm (Fig. 4a, c, d), concordant with the regional structures. Broad isoclinal overturned folds and restricted normal open folds with hinge lines parallel to the Lm are present (Fig. 4f). In general, these folds exhibit northward vergence. Frequently the Sm of the TCF units was heterogeneously replaced by a greenschist facies mylonitic foliation with the same orientation. In stereographic projection, the Sm data suggest cylindrical folding along a constructed axis N80◦ E/20◦ (Fig. 4c). Outcrop and hand-sample scale parasitic folds display the same pattern of orientation (Fig. 4e). The AC rocks

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Fig. 3. Simplified geological-structural map (a) and structural profile (b) of the Cajati area (see localisation on Fig. 2). Stereograms are projection of the lower hemisphere onto equal area nets with structural data of the Atuba Complex (c–d) and of the Turvo-Cajati Formation (e–f). The stereoplots in (g) and (h) refer to structural data of outcrop-scale parasitic folds of the sites 140, 245 and 246. Also shown are the localisations of sites with kinematic indicators displayed in Fig. 5.

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Fig. 4. Simplified geological-structural map (a) and structural profile (f) of the Barra do Turvo area (see localisation on Fig. 2), and stereoplots with structural data of the Atuba Complex (b) and of the Turvo-Cajati Formation (c–d). The stereoplot in (e) refer to structural data of outcrop-scale parasitic folds of the site 201. Also shown are the localisations of sites with kinematic indicators displayed in Fig. 5.

show a NNE-trending steeply dipping Sm and a shallowly SSWplunging Lm (Fig. 4b). Nevertheless, structures (Sm , Lm and fold elements) that are parallel to the structures of the TCF occur in restricted outcrops in the vicinity of the contact between both units. Unambiguous kinematic indicators in the Barra do Turvo area occur along the late transcurrent shear zones, such as the Putunã Shear Zone (Fig. 5j, k). However, inversions of kinematic indicators in inverted limbs of overturned folds within the TCF also occurred (SE portion of the Fig. 4a). The kinematic data of the TCF allow for interpretation of a dominant top-to-the-WSW shearing, if we consider the inversion of the kinematic indicators. Due to the many contrasting sense of shear indicators associated with folds that have hinge lines parallel to the stretching lineation in individual outcrops (Fig. 5h, i), a consistent kinematic interpretation was not possible for the AC in the Barra do Turvo area.

3.3. Structural interpretation The cylindrical folding of Sm along the sub-horizontal axes, as indicated by the stereoplots, indicates that Sm was originally sub-horizontal, which could be associated with early low-angle thrusting or extensional faulting. The stratigraphic inversion with the AC (Archaean to Paleoproterozoic; Campagnoli, 1996; Sato et al., 2003) overlying the Higher-TCF (maximum age of sedimentation of ca. 900 Ma; Campanha et al., 2009) in the Cajati area is unequivocal evidence for thrusting during the Neoproterozoic. Another evidence is a metamorphic inversion present in the Higher-TCF in the Cajati area, where staurolite–kyanite micaschists were overlaid by sillimanite paragneiss (Fig. 3b). The occurrence of inverted kinematic indicators on inverted limbs of overturned folds that have axes parallel to the stretching lineation in the Cajati area indicates that the folding occurred after the development of the Sm and Lm . Thus, the geometric and kinematic data suggest

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Fig. 5. Macro and microscopic shear sense indicators observed in sections parallel to the stretching lineation. Site localisations are shown in Figs. 3 and 4. (a–b) Sigmoid amphibole aggregates with ␴-type strain shadow, top-to-the-WSW movement (site 298, Turvo-Cajati Formation, Cajati area). (c) Garnet porphyroblast with ␴-type strain shadow, top-to-the-WSW shear sense (site 243, Turvo-Cajati Formation, Cajati area). (d–e) Asymmetric c-axis fabrics indicating top-to-the-WNW (d; site 127) and top-tothe-ESE (e; site 257) shear sense (Atuba Complex, Cajati area). (f) Asymmetric boudin of amphibolite, top-to-the-ESE shear sense (site 137, Atuba Complex, Cajati area). (g) Porphyroclast of plagioclase with ␴-type strain shadow and SC fabric, sinistral shear sense with top-to-the-ESE component (site 156, Atuba Complex, Cajati area). (h) Asymmetric boudin of quartz vein, dextral shear sense with top-to-the-ENE component (site 530, Atuba Complex, Barra do Turvo area). (i) ␦-type mantled porphyroclast of feldspar, sinistral shear sense with top-to-the-WSW component (site 530, Atuba Complex, Barra do Turvo area). (j) Porphyroblast of garnet with ␴-type strain shadow, sinistral shear sense with top-to-the WSW component (site 592, Turvo-Cajati Formation, Putunã Shear Zone, Barra do Turvo area). (k) C-type shear bands with mica fish arranged between the bands, sinistral shear sense with top-to-the WSW component (site 592, Turvo-Cajati Formation, Putunã Shear Zone, Barra do Turvo area). (l) C-type shear bands with mica fish arranged between the bands, top-to-the WNW shear sense (Turvo-Cajati Formation, east of the Cajati area).

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Fig. 6. (a) Geometric and kinematic block diagrams for the Cajati area. (b) Nappe stacking model for the study area.

westward thrusting overprinted by a folding phase (Fig. 6a). The overall structure was later segmented by sinistral strike-slip faults including the Putunã Shear Zone that separates two main domains with different structural relationships between the AC and the TCF units. Combined with geological mapping the structural data of the domain to the south of the Putunã Shear Zone suggest stacking of three principal nappes (Fig. 6b): (i) the Higher-TCF in the base; (ii) an allochthonous AC nappe overlaying the Higher-TCF; and (iii) the Lower-TCF overlaying the AC and Higher-TCF nappes. The domain to the north of the Putunã Shear Zone displays a different stacking (Fig. 6b): (1) the AC in the base; (2) the Higher-TCF overlying the AC; and (3) the Lower-TCF overlaying the AC and Higher-TCF nappes. In both domains the Lower-TCF override the AC and the Higher-TCF and the tectonic contacts between them, suggesting an out-of-sequence thrusting (Fig. 6b). The contrast of structural orientation of the AC and the TCF units in the Barra do Turvo area (north of the Putunã Shear Zone) indicates that the fabrics of both units were not developed contemporaneously.

4. Metamorphism Approximately 150 thin sections were examined and eight representative samples (six from the TCF and two from the AC; Fig. 7, Table 1) were evaluated for mineral chemistry and thermobarometry. Chemical analyses were performed with a JEOL JXA-8600S EPMA (electron probe microanalyzer) at the University of São Paulo (Brazil). Operation conditions were: 15 kV, 20 nA and 5–10 ␮m as the column accelerating voltage, beam current and beam diameter, respectively. Representative analyses are shown in Tables 2 and 3. 4.1. Petrography, mineral chemistry and reaction history 4.1.1. The Turvo-Cajati Formation The six samples of the TCF are related with three metamorphic zones of the Higher-TCF (staurolite–kyanite, kyanite–K-feldspar and sillimanite–K-feldspar zones; Fig. 7) and are representative of proposed P–T–t evolution for this unit.

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Fig. 7. Maps of metamorphic zoning for the Turvo-Cajati Formation in the Cajati (a) and Barra do Turvo (b) areas and location of studied samples (circles: samples with mineral chemistry; squares: samples with reference to photomicrographs in Figs. 8 and 10). See localisation on Fig. 2.

4.1.1.1. Staurolite–kyanite zone (sample 242B). This comprises micaschist with the assemblage zone Qtz + Ms + Bt + Ky + Grt + Pl + St ± Ilm ± Rt (abbreviations after Kretz, 1983) and a schistosity defined by the orientation of micas and porphyroblastic kyanite. Granitic veins (centimetre-thick) consisting of feldspar, muscovite, quartz, tourmaline and minor garnet are parallel to, or truncate, the schistosity. Garnet porphyroblasts (3–4 mm) are euhedral or corroded, the latter replaced by Bt + Sil aggregates or Bt + Ms ± Qtz coronas. The presence of poikiloblastic cores with wavy inclusion trails and inclusion-poor rims (Fig. 8a) suggest two main growth pases. Garnet is rich in almandine (Alm79–83 ) and displays prograde chemical zoning, with bell-shaped profiles for spessartine, grossular and XFe (Fe/Fe + Mg) contents (Fig. 9a). A reversed trend in the outer rim indicates (Fig. 9a) late diffusional reequilibration. Plagioclase (An7–11 ) and staurolite occur in the matrix.

The occurrence of muscovite-rich leucosomes and melt inclusions in garnet are evidence for partial melting, which may be explained by the reaction (1) (Spear et al., 1999). Ms + Bt + fluid = St + Grt + Ky + melt

(1)

Stable staurolite suggests that the fluid-absent staurolite melting reaction (2) has not been crossed. Reactions (1) and (2) limit a field of 660–690 ◦ C and 7–13 kbar (Spear et al., 1999). St + Ms = Grt + Bt + Ky + melt

(2)

Corroded grains of garnet, kyanite and staurolite enveloped by Bt + Ms ± Qtz coronas suggest that the reaction (1) has been crossed in the reverse sense, requiring consumption of melt. This is corroborated by the presence of thin films of quartz, pseudomorphing the former melt, surrounding the same corroded minerals. However, garnet rims replaced by Sil + Bt ± Qtz aggregates indicate a late decompression.

Table 1 Modes of metamorphic mineral assemblages of the Turvo-Cajati Formation and the Atuba Complex. Prograde metamorphic assemblage Sample

Area

Unit

Metamorphic zone

Rock

Ms

Qtz

Bt

056 129A 158B 242B 297D 298E 133A 164B

Barra do Turvo Cajati Cajati Cajati Cajati Cajati Cajati Cajati

TCF TCF TCF TCF TCF TCF AC AC

Sil–Kfs Ky–Kfs Ky–Kfs St–Ky Sil–Kfs Sil–Kfs

Paragneiss Paragneiss Paragneiss Micaschist Paragneiss Paragneiss Amphibolite Ortogneiss

– tr 15 25 – – – –

20 25 37 29 26 30 tr 20

tr 10 30 13 28 10 23 5 21 1 35 2 5 – 10 –

tr – trace.

Grt

Pl

Kfs

3 10 20 8 5 – 5 – 18 24 tr 17 40 – 49 1

Retrograde Ky

Sil

St

Hbl

Ttn

Ilm

Rt

Ser/Ms

Chl

Ep

– tr 5 10 tr tr – –

tr – tr tr 10 15 – –

– – – 1 – – – –

– – – – – – 49 10

– – – – – – 5 5

tr – – – tr – tr tr

– tr tr tr tr tr – –

30 3 – – tr tr 1 1

25 tr tr tr – tr tr 1

– – – – – – tr 3

272

Table 2 Representative microprobe analyses of garnet and plagioclase of rocks from the Turvo-Cajati Formation and Atuba Complex. Garnet

Sample

056

129A

129A

158B

158B

242B

242B

297D

297D

298E

129A

129A

158B

158B

242B

242B

297D

297D

133A

133A

164B

164B

Core

Core

Rim

Core

Rim

Core

Rim

Core

Rim

Core

Core

Rim

Core

Rim

Core

Rim

Core

Rim

Core

Rim

Core

Rim

37.43 0.04 21.81

36.78 0.03 21.16

36.82 0.08 21.51

37.37 0.02 21.23

37.57 0.03 21.82

36.58 0.07 21.17

36.92 b.d.l. 21.65

36.90 0.06 21.52

36.53 b.d.l. 21.55

37.28 b.d.l. 21.64

63.00 b.d.l. 23.78 0.31

57.28 b.d.l. 27.02 0.07

61.55 b.d.l. 24.21 b.d.l.

58.88 b.d.l. 26.47 b.d.l.

65.86 0.04 21.23 b.d.l.

65.35 0.09 21.69 b.d.l.

61.49 b.d.l. 23.73 b.d.l.

58.72 0.04 26.52 b.d.l.

59.34 b.d.l. 25.50 0.09

61.00 b.d.l. 24.89 0.21

63.73 b.d.l. 22.87 0.21

62.22 b.d.l. 23.85 0.11

28.04 6.21 4.79 1.87

29.83 6.25 0.87 4.84

35.60 0.30 2.87 3.79

29.29 4.07 1.37 7.13

33.29 0.58 3.21 5.22

35.69 2.48 1.98 2.29

36.95 0.51 3.64 0.67

35.50 1.13 2.94 2.01

35.73 1.63 2.36 2.10

34.58 1.04 4.25 1.78

0.03 b.d.l. 5.17 8.95 0.12

0.05 b.d.l. 8.69 6.58 0.13

b.d.l. b.d.l. 5.78 8.24 0.27

0.03 b.d.l. 8.04 7.04 0.12

b.d.l. b.d.l. 1.65 10.72 0.24

b.d.l. 0.02 2.44 10.48 0.08

b.d.l. b.d.l. 5.22 8.39 0.34

b.d.l. b.d.l. 7.97 7.04 0.06

0.03 b.d.l. 7.44 7.03 0.34

0.02 b.d.l. 6.33 8.00 0.16

0.02 b.d.l. 4.10 9.33 0.20

0.02 b.d.l. 4.93 8.81 0.28

b.d.l.

b.d.l.

b.d.l.

b.d.l.

b.d.l.

101.73

100.26

100.34 100.10

99.91

100.58

101.37

100.33

100.04

100.58

99.74

100.15

99.19

100.35

99.78

100.63

100.46

100.22

12 2.958 0.001 2.056

12 2.965 0.001 2.029

8 2.758

8 2.579

8 2.731

8 2.614

8 2.697

8 2.806

8 2.753

0.266 0.001

1.385 0.002

8 2.611 0.001 1.391 0.003

8 2.654

1.421 0.002

8 2.871 0.003 1.123 0.002

8 2.749

1.227 0.010

8 2.900 0.001 1.102 0.001

1.344 0.003

1.297 0.007

1.187 0.007

1.243 0.004

2.420 0.112 0.285 0.182

2.300 0.070 0.504 0.151

0.001 0.001 0.242 0.759 0.006

0.002

0.001

0.001

0.001

0.001

0.416 0.570 0.007

0.275 0.709 0.015

0.382 0.606 0.007

0.078 0.915 0.013

0.001 0.115 0.893 0.004

0.250 0.728 0.020

0.380 0.607 0.003

0.356 0.610 0.020

0.300 0.685 0.009

0.193 0.796 0.011

0.234 0.756 0.016

8.013

8.020

5.005

4.997

4.997

4.998

5.011

5.012

4.999

4.996

4.987

4.996

5.002

5.006

80.70 3.74 9.49 6.07

76.03 2.31 16.67 5.00

24.04 75.31 0.64

41.87 57.39 0.73

27.53 70.97 1.50

38.39 60.90 0.70

7.75 90.95 1.29

11.36 88.24 0.40

25.05 72.95 2.00

38.38 61.31 0.30

36.15 61.86 1.99

30.16 68.94 0.90

19.31 79.55 1.14

23.23 75.17 1.59

SiO2 TiO2 Al2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O Cr2 O3 Total Oxygen Si Ti Al Fe3+ Fe2+ Mn Mg Ca Na K Cr Total alm sps prp grs

Plagioclase

0.03

0.03

0.05

0.03

100.24

99.77

101.02

100.50

0.03

12 2.970 0.002 2.037

12 2.986 0.002 2.025

12 2.942 0.005 2.025

12 2.991 0.001 2.002

12 2.957 0.002 2.024

12 2.965 0.004 2.022

12 2.962 2.047

12 2.969 0.004 2.040

1.858 0.417 0.566 0.158

2.025 0.429 0.105 0.421

2.378 0.020 0.342 0.325

1.960 0.276 0.164 0.611

2.191 0.038 0.377 0.440

2.419 0.170 0.239 0.199

2.479 0.035 0.435 0.058

2.388 0.077 0.353 0.173

0.002

0.002

0.003

0.002

8.009 61.96 13.90 18.86 5.28

7.994 67.95 14.41 3.52 14.12

8.039 77.59 0.67 11.15 10.59

8.006 65.10 9.16 5.45 20.30

1.251 0.002

0.001

0.002 8.029

8.019

8.015

71.92 1.26 12.37 14.45

79.90 5.62 7.90 6.58

82.44 1.16 14.48 1.92

Mineral abbreviations from Kretz (1983). b.d.l.–below detection limit.

8.006 79.84 2.57 11.80 5.79

an ab or

F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

Mineral

Table 3 Representative microprobe analyses of K-feldspar, biotite, muscovite, staurolite and amphibole of rocks from the Turvo-Cajati Formation and Atuba Complex. K-feldspar

Sample

056

129A

297D

298E

056

129A

158B

242B

297D

298E

129A

158B

242B

242B

242B

133A

133A

164B

164B

SiO2 TiO2 Al2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O Cr2 O3 ZnO

64.40 0.04 18.61 b.d.l.

64.65 b.d.l. 18.04 0.02

64.20 b.d.l. 18.72 b.d.l.

64.40 b.d.l. 18.66 0.02

35.73 3.59 18.27

34.03 4.26 17.80

35.85 2.17 18.82

35.25 2.75 19.04

34.63 4.39 18.43

34.81 2.30 19.36

46.01 1.20 34.66

46.06 0.75 35.25

46.30 0.99 35.95

27.56 0.60 55.87

27.57 0.76 55.27

b.d.l. b.d.l. b.d.l. 0.81 15.46

b.d.l. b.d.l. b.d.l. 0.71 15.64

b.d.l. b.d.l. 0.02 1.79 14.13

b.d.l. 0.02 0.10 1.19 15.01

18.21 0.36 10.04 b.d.l. 0.04 9.64

24.07 0.10 5.95 0.06 0.08 9.47

20.15 0.09 8.87 b.d.l. 0.13 9.67

22.08 0.07 7.47 b.d.l. 0.19 9.43

21.84 0.10 6.56 b.d.l. 0.12 9.58

22.04 0.07 7.32 b.d.l. 0.11 9.74

1.55 0.02 0.64 b.d.l. 0.40 10.54

1.42 b.d.l. 0.85 b.d.l. 0.65 10.04

1.16 b.d.l. 0.61 b.d.l. 1.01 9.65

10.66 0.14 0.68 b.d.l. 0.11 b.d.l. 0.03 2.49

10.46 0.17 0.61 b.d.l. 0.11 b.d.l. 0.07 2.93

43.64 0.69 12.59 5.11** 11.29 0.26 10.95 11.63 1.28 1.26

42.80 0.85 12.79 6.39** 10.69 0.38 10.94 11.75 1.33 1.38

44.71 0.50 10.26 4.98** 10.87 0.40 11.77 11.72 1.27 1.05

45.68 0.43 8.71 6.61** 9.90 0.33 12.42 11.66 1.16 0.70

Total

99.34

99.07

98.88

99.41

95.90

95.79

95.77

96.28

95.66

95.77

95.02

95.05

95.69

98.16

97.95

98.70

99.30

97.53

97.60

8 3.009

8 2.981

8 2.984

0.990 0.001

4.025

1.019 0.001

11 2.695 0.204 1.624

11 2.651 0.249 1.632

11 2.726 0.124 1.687

11 2.690 0.158 1.713

11 2.667 0.254 1.674

11 2.677 0.133 1.756

11 3.076 0.060 2.728

11 3.069 0.037 2.769

11 3.056 0.049 2.797

46 7.599 0.125 18.154

46 7.634 0.159 18.034

1.149 0.023 1.129

1.568 0.006 0.690 0.005 0.012 0.940

1.281 0.006 1.005

1.409 0.004 0.850

1.407 0.006 0.753

1.418 0.004 0.839

0.087 0.001 0.063

0.079

0.064

0.085

0.060

2.458 0.032 0.278

2.421 0.040 0.251

0.020 0.938

0.027 0.918

0.018 0.942

0.017 0.956

0.052 0.899

0.084 0.854

0.130 0.813

0.057

0.060

23 6.400 0.076 2.176 0.564* 1.385 0.033 2.394 1.827 0.363 0.236

23 6.269 0.094 2.208 0.704* 1.309 0.047 2.388 1.844 0.378 0.258

23 6.619 0.055 1.791 0.555* 1.346 0.050 2.596 1.859 0.363 0.198

23 6.735 0.048 1.514 0.733* 1.221 0.041 2.729 1.842 0.332 0.132

0.006 0.508

0.015 0.598 15.454

15.499

15.432

15.327

0.650

0.658

0.666

0.707

Oxygen Si Ti Al Fe3+ Fe2+ Mn Mg Ca Na K Cr Zn

2.988 0.001 1.018

0.073 0.915

0.064 0.929

Biotite

0.001 0.162 0.837

0.001 0.005 0.106 0.887

Total

4.996

4.992

5.006

5.004

an ab or

0.00 7.39 92.61

0.00 6.41 93.59

0.10 16.20 83.70

0.50 10.62 88.88

0.005 0.928

XMg

Muscovite

7.757

7.753

7.786

7.769

7.722

7.800

0.496

0.306

0.440

0.376

0.349

0.372

XK

Staurolite

6.966

6.978

6.969

0.945

0.910

0.862

Amphibole

XMg

F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

Mineral

Mineral abbreviations from Kretz (1983). * Fe3+ estimated according to Holland and Blundy (1994). ** Fe2 O3 has been recalculated from Fe3+ estimates.

273

274

F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

Fig. 8. Optical photomicrographs showing microstructures of rocks from the higher-grade unit of the Turvo-Cajati Formation. (a) Garnet porphyroblast with a poikiloblastic core and an inclusion-free rim, sample 242B, staurolite–kyanite zone. (b) Submillimetric layering with alternating biotite-rich mesocratic layer (residual neosome), biotite-free granitic leucosome, and muscovite-rich restricted layers. Note in the centre a corroded garnet porphyroblast with biotite-rich reaction-rim corona, sample 129A, kyanite–Kfeldspar zone. (c) Corroded garnet porphyroblasts surrounded by double corona with an internal simplectitic Bt + Pl + Kfs domain, and an external coarse muscovite selvage, sample 129A, kyanite–K-feldspar zone. (d) ‘Melt pocket’ of microcline surrounding rounded crystals of quartz, plagioclase and biotite and corroded garnet, sample 129A, kyanite–K-feldspar zone. (e) Quartz–feldspar aggregates included in garnet porphyroblasts, where films of microcline surrounds residual grains of quartz, plagioclase and biotite, sample 129A, kyanite–K-feldspar zone. Note crystal faces locally developed on quartz. (f) Corroded grains of kyanite surrounded by moat of plagioclase or corona with plagioclase + muscovite. The coronas are surrounded by films of K-feldspar mimicking former melt; sample 015, kyanite–K-feldspar zone. (g) Megacrystal of plagioclase with vermicular inclusions of quartz; sample 355, kyanite–K-feldspar zone. (h) Euhedral plagioclase and intersertal quartz inferred to record crystallisation from a melt in a trondhjemitic leucosome; sample 325, kyanite–K-feldspar zone.

F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

Fe/(Fe+Mg)

0.4

(b)

129A Fe/(Fe+Mg)

158B Fe/(Fe+Mg)

0.4

0.8 0.4

Alm

XAlm, XFe

0.2

0.8

0.8

Alm

Alm

0.6 0.2

Grs

0.2

Grs

Prp Alm

0.6

0.6

Grs

Sps

Prp Sps

0

(e)

3

(d)

056

Fe/(Fe+Mg)

0.4

Prp

2

0

4

6

297D Fe/(Fe+Mg)

0.8 0.4

0

0

4.5

(f)

9

XAlm, XFe

0

Sps

XAlm, XFe

0.4

XAlm, XFe

0

(c)

XAlm, XFe

242B

XAlm, XFe

XAlm, XFe

(a)

275

298E Fe/(Fe+Mg)

0.8 0.4

Alm

0.8 Alm

Alm

0.6 0.2

0.2

0.6 0.2

0.6

Prp

Prp Sps

Prp

Grs

Grs Grs

0

0

0.4 1.0

0.5

(g)

129A

0.8

Sps

Sps

0.4

0

0.75

0.25

(h)

1.25

0.5

(i)

158B

0.8

0,4

0

1.5

2.5

297D

0.8

Ab

An

0.2

0.6 0.4

0.6

XAn, Xab

0.2

0 0.3

0.5

0.7

164B

0.8

(k)

0 0

0.5

XAn, Xab

XAn, Xab

0.4

0.2

An

0.2

0

0 0

0.15

0.3

0.45

Distance (mm)

1

Tr Act

Mg-Hbl

Fe-Act

Fe-Hbl

Ab

An

0

0.1

0.2

Distance (mm)

0.5

0.9

Distance (mm)

(l)

Ab

0.4

0.1

1.0

133A

0.8 0.6

An

0.2

0 0.1

(j)

0.6

0.4

An

Mg/(Mg+Fe)

0.4

Ab

XAn, Xab

0.6

XAn, Xab

Ab

0.3

133A 164B

Ts

0.5

0

8

7.5

7

Fe-Ts

6.5

6

5.5

Si

Fig. 9. (a–f) Chemical profiles across garnet porphyroblasts from samples of the Turvo-Cajati Formation: (a) sample 242B, St–Ky zone, (b–c) samples 129A and 158B, Ky–Kfs zone, (d–f) samples 056, 297D and 298E, Sil–Kfs zone. Mole fractions of spessartine, grossular, almandine and pyrope end-members are plotted against distance across the grain. (g–k) Chemical profiles across plagioclase grains from samples of the Turvo-Cajati Formation (g–i) and of the Atuba Complex (j–k). (l) Classification of Ca-amphiboles of samples 133A and 164B, Atuba Complex, according to Leake et al. (1997).

276

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4.1.1.2. Kyanite–K-feldspar zone (samples 129A and 158B). This zone comprises migmatitic Bt + Pl + Qtz + Mc + Grt ± Ms ± Ky ± Ilm ± Rt paragneiss. The description of the constituent parts of the migmatites follows the nomenclature of Sawyer (2008). The paragneiss shows a millimetre-thick tabular to lenticular layering (Fig. 8b) composed of mesocratic neosome with Pl + Qtz + Bt + Grt ± Mc ± Ky; an in-source leucosome composed of Qtz + Ab ± Mc; a biotite-rich selvedge surrounding leucosomes; and coarse-grained muscovite-rich lenses in the neosome. Porphyroblastic garnet (5–30 mm) from samples 129A and 158B exhibit corroded boundaries (Fig. 8b), it is rich in almandine (Alm65–79 ) and features prograde chemical zoning in the major elements (Tracy, 1982) (Fig. 9b, c). Porphyroblasts from sample 158B shows an increase in spessartine and XFe contents in the outer rim (Fig. 9c), indicating late diffusional reequilibration. Matrix plagioclase of sample 129A shows more anorthite content (An37–50 ) than grains included in garnet (An24–39 ), whereas porphyroblasts show flat core compositions (∼An50 ) and an abrupt decrease in anorthite rimward (An42–44 ) (Fig. 9g). Porphyroblastic plagioclase from sample 158B displays anorthite increase from the core (An27 ) to the rim (An34 ) (Fig. 9h). The grains included in garnet show a larger compositional variation (An26–39 ), in which the inclusions near the rims exhibit higher anorthite contents than inclusions in the cores of garnets. Garnet porphyroblasts from sample 129A show reactionrim coronas composed of biotite or simplectites with Bt + Pl + Mc ± Qtz ± Ms (Fig. 8b, c). Microcline in these coronas commonly occurs as thin films surrounding round-shaped minerals (Qtz, Pl, Bt) and corroded garnet boundaries (Fig. 8d), forming concave–convex relations. In many examples this forms larger pockets with a continuous crystallographic orientation. Such a microstructural relationship suggests that the microcline mimics the former melt (e.g., Sawyer, 1999, 2001; Brown, 2002; Marchildon and Brown, 2002; Holness and Sawyer, 2008). Therefore, the loss of garnet should have occurred due to a melt-consuming reaction reversal (Kriegsman, 2001; Brown, 2002). Microcline and plagioclase with the same microstructural relationship as in the quartz–feldspar aggregates also occur in the matrix of leucosomes and in mesocratic neosome and within inclusions located in the rim of garnet (Fig. 8e). The latter is evidence that the garnet rim grew in the presence of a melt. Additional evidence supporting the former presence of a melt includes leucosomes with plagioclase megacrystals with vermicular quartz inclusions (Fig. 8g) and euhedral plagioclase against intersertal quartz (Fig. 8h); and garnet rich in melt inclusions. Late fibrolitic sillimanite is intergrown with biotite and form restricted aggregates, being possibly the result of reaction between melt and residual phases during cooling (e.g., Waters, 2001). The scarcity of prograde muscovite, evidence for partial melting and the matrix assemblage Kfs–Qtz–(±Ky) require that the reaction (3) has been reached, which should occur at ∼710–730 ◦ C and 8–13 kbar (Spear et al., 1999). Samples that remain rich in muscovite (∼15% in sample 158B) may indicate that reaction (3) has not been completed, probably as a result of the bulk rock composition and the relatively high-pressure conditions (see Section 4.3). Ms + Pl + Qtz = Kfs + Ky + melt

(3)

Additional evidence for melting involving plagioclase is the general trend of rimwards anorthite increase in matrix plagioclase, which occurs once the melt is enriched in Na/Ca relative to the coexisting plagioclase (Spear and Kohn, 1996). Corroded grains of kyanite enveloped by albite moats or coronas with Ab + Ms + Qtz, both surrounded by Kfs pseudomorphing melt (Fig. 8f), indicate that reaction (3) has been crossed on cooling. Sample 129A shows garnet porphyroblasts with melt inclusions and inclusions of microcline–quartz aggregates where the feldspar

mimics former melt (Fig. 8e), suggesting that the garnet might have grown due to the fluid-absent biotite melting reaction (4) (e.g., Vielzeuf and Holloway, 1988). Bt + Ky + Pl + Qtz = Grt + Kfs + melt

(4)

Sample 129A also shows corroded garnet porphyroblasts surrounded by pockets where microcline mimics melt, which are enclosed by simplectites with Bt + Pl + Qtz. This situation suggests that reaction (4) was crossed on cooling. As an alternative explanation, Le Breton and Thompson (1988) observe that, due to different dP/dT slopes, reactions (3) and (4) will intersect at a given pressure condition (kyanite stability field), and above that, reactions (3) and (4) will compete, creating reaction (5). Ms + Bt + Pl + Qtz = Grt + Kfs + Ky + melt

(5)

Evidence for the back reaction (5) is represented in sample 129A in the form of corroded garnet rims surrounded by a double corona, which are formed by Bt + Pl + Qtz + Kfs ± Ms (inner corona) and muscovite selvedge (outer corona) (Fig. 8c). Alternatively, this feature may be the result of an increase in the activity of H2 O outwards in the corona and a decrease in the activity of K+ . 4.1.1.3. Sillimanite–K-feldspar zone (samples 056, 297D and 298E). This zone is composed of migmatitic paragneiss with sillimanite as the dominant aluminosilicate. Prograde muscovite occurs only as inclusions in the garnet porphyroblasts. The peak metamorphic assemblage is Qtz + Bt + Sil + Kfs + Grt ± Pl ± Rt. The structure of the paragneiss consists of an anastomosing schistosity/banding in which biotite-rich and sillimanite-rich layers wrap sigmoid lenses of quartz–feldspar and porphyroblastic garnet. Plagioclase-poor (samples 056 and 298E) and plagioclase-rich (sample 297D) paragneisses were distinguished. In sample 298E porphyroblastic garnet (1–7 mm) is rounded or partially corroded and displays a poikiloblastic core (with inclusions of kyanite) and an inclusion-free rim. It is enveloped by coronas composed of Bt + Sil ± Qtz ± feldspar. In sample 297D garnet occurs as dismembered corroded fragments surrounded by thin films or moats of plagioclase (Fig. 10a), or one surrounded by double coronas with an inner domain of Pl + Qtz and an outer one with Bt + Qtz ± Pl ± Sil. Garnet from samples 056, 297D and 298E features diffusional zoning with relatively flat chemical traverses in core and an increase in spessartine and XFe contents rimward (Fig. 9d–f). A relict growth zoning occur in a garnet grain of sample 298E (Fig. 9f). Perthitic feldspar (Or84–95 ) from samples 056 and 298E forms highly elongated and irregularly shaped megacrystals (up to 1 cm in diameter), oriented along the schistosity. Commonly, K-feldspar forms pseudomorphs after a former melt (Fig. 10b). In sample 297D microcline (Or83–96 ) and plagioclase (An25–30 , Fig. 9i) are exclusive to the matrix. Kyanite occurs as inclusions in K-feldspar (sample 298E) and in plagioclase (sample 297D). Sample 056 was intensely retrometamorphosed along the Serra do Azeite Shear Zone (Barra do Turvo area). In this sample, almost all biotite and sillimanite have been replaced by chlorite and muscovite aggregates, respectively. The absence of matrix muscovite and the assemblage Grt + Sil + Kfs + Bt + Qtz ± Pl indicate that temperatures in excess of the fluid-absent muscovite melting reaction (3) have been reached. Inclusions of muscovite and kyanite in garnet and of kyanite in K-feldspar are evidence for this reaction. The fact that matrix muscovite was completely eliminated indicates that the fluid-absent biotite melting reaction (4) has been crossed, which should have been responsible by the prograde development of Grt + Kfs. The fact that garnet and K-feldspar enclose kyanite and never sillimanite suggests that reactions (3) and (4) have been crossed within the stability field of kyanite. The development of sillimanite probably

F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

277

Fig. 10. Optical photomicrographs and photographs showing microstructures of rocks from the Turvo-Cajati Formation (a–b) and field structures (c–d) and microstructures (e–f) of rocks from the Atuba Complex. (a) Rounded garnet fragment surrounded by moat of plagioclase, sample 297D, Sil–Kfs zone, Cajati area. (b) Arc-shaped megacrystal of K-feldspar surrounding a rounded grain of quartz. The former is inferred to record crystallisation from a melt; sample 296, Sil–Kfs zone, Cajati area. (c) Strongly banded tabular centimetre-thick layered fabric with alternation of dark tonalitic mesocratic layer rich in biotite and hornblende and pale granitic leucosome; site 530, Barra do Turvo area. (d) Granitic veins truncating the mylonitic foliation of gneiss, Cajati area. (e) Lenticular microstructure in a mesocratic layer, sample 156, Cajati area. Shear bands and sericite-rich domains are present. (f) Plagioclase grain with oscillatory zoning from an amphibolitic boudin, sample 133A, Cajati area.

occurred by the back reaction (4), which is corroborated by (i) corroded garnet enveloped by coronas with Bt + Sil ± Qtz ± feldspar; (ii) dismembered garnet fragments enveloped by moats of plagioclase; and (iii) the absence of sillimanite within leucosomes or as inclusions in K-feldspar. 4.1.2. The Atuba Complex The AC is composed of mylonitic orthogneiss with alternating grey tonalitic, trondhjemitic or plagioclase-rich mesocratic neo-

somes with biotite and hornblende (residual neosome) (Fig. 10c), white granitic leucosome (Fig. 10c), and amphibolitic lenses (Fig. 4f). Reddish leucogranitic and pegmatitic veins that range from undeformed to slightly deformed commonly truncate the gneissic fabric (Fig. 10d). Amphibole and plagioclase in representative samples of a mesocratic neosome (164B) and an amphibolitic layer (133A) were analysed. The mesocratic neosome shows lenticular microstructure (Fig. 10e) with fish-shaped porphyroclasts of plagioclase moulded

6–6.5

6–7

600–620

610–640

670–730 690–730 (core) 630–680 (rim)

T (◦ C)

11 11.5

8–10.3 8–10 760–820 790–820

3

2

Calibrations of Newton and Haselton (1981), Hodges and Crowley (1985) and Koziol and Newton (1988). Calibration of Holland and Blundy (1994). According to Zenk and Schulz (2004). 1

TCF TCF TCF TCF TCF TCF AC AC AC 056 129A 158B 242B 297D 298E 133A 164B 164B

Sil–Kfs Ky–Kfs Ky–Kfs St–Ky Sil–Kfs Sil–Kfs

792 740 761 637 805 786

42 41 32 24 56 38

7.8 10.4 10.5 7.9 9.5 8.4

1.7 1.2 1.1 1.8 1.1 1.6

0.934 0.537 0.730 −0.397 0.238 0.927

10.5–13 10–12 8.5–10.5 7.5–11.5

800 770

T (◦ C) P (kbar P (kbar) Metamorphic zone Unit Sample

Table 4 Summary of the geothermobarometry data.

4.2. Conventional thermobarometry

sd (T) (◦ C)

avP (kbar)

sd (P) (kbar)

Corr avT (◦ C)

T (◦ C)

Hbl-Pl thermometer2 Phase diagrams and geothermobarometry

(6)

Peak metamorphic conditions for the High-TCF samples 242B (St–Ky zone), 129A and 158B (Ky–Kfs zone) and 056, 297D and 298E (Sil–Kfs zone) were calculated using the software THERMOCALC (Powell and Holland, 1988) version 3.26 in the average PT mode (avPT), and the garnet–plagioclase–kyanite–quartz (GASP) barometer in conjunction with the grid of Spear et al. (1999). The six samples show considerably smaller garnet volumes (1–13%) compared to biotite (21–35%), suggesting that the composition of the matrix biotite has most likely not been significantly modified by diffusional effects after peak conditions (e.g., Spear, 1989). For calculations, we selected garnet compositions that were most likely equilibrated under the highest temperature conditions for each sample (i.e., average lowest values of XFe for garnet coupled to the average lowest contents of grossular and spessartine), and analyses of the coexisting matrix minerals. Near-rim and core garnet com-

GASP barometer1

Qtz + Pl + Kfs + H2 O = melt

Amphibole thermobarometer3

by the matrix composed of dynamically recrystallised quartz and disseminated biotite and hornblende. Plagioclase (An19–24 ) from sample 164B shows a slight increase in anothite rimward (Fig. 9j) and has various degrees of replacement by sericite or sericite + epidote. Sample 164B presents porphyroblastic Mghornblende (cf. Leake et al., 1997; Fig. 9l) oriented parallel to the gneissic layers, which shows a weak compositional zoning with decreasing Na, Ti and Al coupled with increasing Si from the core to the rim, indicating that the rims record lower temperatures than the cores (e.g., Spear, 1980). Fe-tschermakite occurs in restricted mafic layers, with abundant retrograde chlorite, biotite, epidote and titanite, and could represent an unusual bulk composition or a relict of higher P–T conditions. The leucosome is composed of Mc + Qtz ± Pl ± Bt. Microcline is perthitic, coarse-grained and slightly elongated which defines a shape-preferred orientation parallel to the gneissic layering. Inclusions of rounded, or corroded plagioclase partially replaced by sericite are common in the microcline. The microstructure of the plagioclase present in leucosomes suggests that it is a residual phase or may be attributable to either an early crystallised phase or a phase inherited from another magma. Quartz aggregates display grain boundary migration microstructures replaced by bulging recrystallisation, inferred to be evidence of mylonitisation during decreasing temperature. In the vicinity of the contact zones with the TCF in the Cajati area, the banded orthogneiss was commonly mylonitised under greenschist facies conditions. Amphibolitic lenses contain Hbl + Pl + Bt + Ttn oriented along a schistosity that is parallel to the foliation in the host gneiss. Sample 133A is dominated by tschermakite amphibole with little compositional variation (Fig. 9l). Plagioclase (An30–36 ) from this sample exhibits oscillatory optical zoning (Fig. 10f), and a chemical zoning with decrease in anorthite content from the core to the rim (Fig. 9k). The large volume of granitic leucosomes in the AC suggests that partial melting occurred here. This is confirmed by the presence of restricted microcline and quartz pseudomorphing melt pockets surrounding corroded grains of plagioclase within mesocratic neosome. The absence of residual K-feldspar (all of the K-feldspar present mimics melt pockets) may be explained by the H2 O-saturated melting reaction (6); K-feldspar is the first phase to disappear during experimental wet melting of tonalite and granodiorite (Wyllie et al., 1976). Therefore, microcline has difficulty in nucleating relative to quartz or plagioclase, and so crystallises last and thus takes the shape of remaining melt pore. Reaction (6) is a low-temperature melting reaction (680–700 ◦ C; Watkins et al., 2007).

P (kbar)

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THERMOCALC results

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F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

positions were used for crystals with growth (samples 129A, 158B and 242B) and diffusional (samples 056, 297D and 298E) zoning, respectively. The results are shown in Table 4. Sample 242 (St–Ky zone) yields avPT of 637 ± 24 ◦ C and 7.9 ± 1.8 kbar (Fig. 11a), whereas the combination of the field delimitated by the reactions (1) and (2) using the grid of Spear et al. (1999) and the GASP barometer (calibrations of Newton and Haselton, 1981; Hodges and Crowley, 1985; and Koziol and Newton, 1988) yields conditions of 660–690 ◦ C and 8.5–10.5 kbar (Fig. 11a). The avPT results for samples 158B and 129A (Ky–Kfs zone) (740–760 ◦ C and 10–11 kbar; Table 4) exceed the fluid-absent muscovite melting reaction (3) within the kyanite stability field. The GASP barometer provides pressures between 10 and 13 kbar at 740–760 ◦ C (Table 4). The avPT results from samples 056, 297D and 298E (Sil–Kfs zone) (790–800 ◦ C and 8–9.5 kbar; Table 4) exceed the fluid-absent muscovite melting reaction (3) within the sillimanite stability field. The GASP barometer yields a pressure range of 7.5–11.5 kbar for sample 297D (Table 4). Peak metamorphic conditions for samples 133A and 164B from the AC were estimated using the hornblende–plagioclase thermometer of Holland and Blundy (1994) and the amphibole thermobarometer of Zenk and Schulz (2004). The former yields results (at 7 kbar) in the range of 670–730 ◦ C (133A; Fig. 11b) and 630–730 ◦ C (164B; Fig. 11c). For sample 164B, rim compositions yield temperatures that are 50–100 ◦ C lower than core compositions, indicating a cooling path. The amphibole thermobarometer yields conditions of 600–660 ◦ C and 6–7 kbar for both samples (Table 4; Fig. 11b, c). 4.3. Phase equilibria modelling Isochemical phase diagrams in the NCKFMASHTi system were constructed for four bulk compositions (samples 129A, 158B, 297D and 298E) using the Perple X software (Connolly, 2005) and the updated version (hp04ver.dat) of the internally consistent thermodynamic database of Holland and Powell (1998). The following solution models were used: Grt, Bt and melt (White et al., 2007); Opx, Crd, St and Chl (Holland and Powell, 1998); feldspar (Fuhrman and Lindsley, 1988); Ms (Coggon and Holland, 2002). The H2 O content for each sample was estimated taking into account the content of hydrous minerals and so that the rocks were saturated in H2 O immediately below the solidus (cf., White et al., 2001). The phase diagrams calculated for sample 158B (Fig. 12a, b) indicate that melting during prograde metamorphism would first occur by the fluid-present muscovite melting reaction between 670 and 690 ◦ C (above 6 kbar) and K-feldspar-bearing residual assemblages occur at temperatures above 800 ◦ C (above 8 kbar) for the considered bulk composition. The peak metamorphic assemblage is represented by the tri-variant field Bt + melt + Pl + Ms + Grt + Ky + Qtz + Rt and the avPT result for this sample (761 ± 32 ◦ C and 10.5 ± 1.1 kbar) is almost completely within this stability field (Fig. 12a). These data, combined with isopleths of Xgrs (Fig. 12a, b), suggest peak conditions of approximately 770 ◦ C and 11.5 kbar (Table 4). These data corroborate the field and petrographic evidence of partial melting for this rock, despite the absence of matrix K-feldspar. In the phase diagram for sample 129A (Fig. 12c) the first melting would occur by the fluid-present muscovite melting reaction in temperatures between 670 and 710 ◦ C (above 6 kbar). The peak metamorphic assemblage (Bt + melt + Pl + Kfs + Grt + Ky + Qtz + Rt) is stable between 775 and 825 ◦ C and above 9.5 kbar (Fig. 12c). This field combined with the avPT result and isopleths of Xalm and Xgrs provide peak conditions of approximately 800 ◦ C and 11 kbar (Fig. 12c–d). The P–T phase diagram for sample 297D (Fig. 13a) indicates that muscovite-free assemblages occur above 670–800 ◦ C

279

for this bulk composition. The field of peak mineral assemblage (Bt + melt + Pl + Kfs + Grt + Sil + Qtz + Rt) combined with the avPT result yield conditions between 760 and 820 ◦ C and 8–10.3 kbar. Inclusions of kyanite in plagioclase suggest that the rock underwent a path of isothermal decompression (Fig. 13a). The retrograde P–T path was inferred from the transition between Rt- and Ilm-bearing assemblages (Fig. 13a). The equilibrium assemblage for sample 298E (Fig. 13b) is represented by the quadri-variant field Bt + melt + Kfs + Grt + Sil + Qtz + Ilm which coincides with the avPT result. A P–T range of approximately 8–10 kbar and 790–820 ◦ C is estimated considering both data. Fig. 13b shows that biotite-free assemblages occur at temperatures above 860–870 ◦ C. The phase diagram also gives information regarding the prograde mineral assemblage enclosed in porphyroblasts. Inclusions of muscovite, kyanite and rutile in garnet and of kyanite in K-feldspar, suggest that the rock was previously equilibrated in the tri-variant field Bt + melt + Kfs + Ms + Grt + Ky + Qtz + Rt, which occurs between 750 and 810 ◦ C and above 10 kbar. The two stability fields attained for the rock suggest a near isothermal decompression (Fig. 13b). 4.4. P–T paths from garnet zoning P–T paths for samples 129A, 158B, 242B and 298E were calculated using the Gibbs method of differential thermodynamics (Spear and Selverstone, 1983) with the GIBBS software (Spear and Menard, 1989; Spear et al., 2001). Calculations were made considering the chemical system MnNCKFMASH and utilising activity and mixing models provided by the program. The input data consist of the composition and modes of minerals considered in equilibrium at the calculated peak P–T conditions. Xalm , Xgrs , Xsps and Xan were considered as monitor variables for samples 129A, 158B and 242B, and Xalm , Xgrs , Xsps and Xannite were monitor variables for sample 298E. Garnet and plagioclase end-members were calculated according to Berman (1990) and Fuhrman and Lindsley (1988), respectively. Water, quartz, kyanite, sillimanite, K-feldspar and muscovite were considered to be pure substances. As correlations of the incremental variations between garnet and plagioclase compositions cannot be precisely established, we considered regular intervals of anorthite increments between the extreme compositions (Spear, 1989). Two types of clockwise P–T paths were calculated (Fig. 14a): (i) near isobaric heating (242B, 158B and 129A); and (ii) near isothermal decompression (298E). Sample 242B follows a near isobaric heating path from 510 ◦ C at 10.2 kbar to 585 ◦ C. It then follows an episode of heating-decompression until 670 ◦ C at 9 kbar. The path calculated for sample 158B indicates a near isobaric heating from 575 to 760 ◦ C at approximately 11.5 kbar (Fig. 14a), despite the strong compositional variation shown by garnet and plagioclase (Fig. 9c). The calculated metamorphic path for sample 129A initiates at around 620 ◦ C and 11.6 kbar and is followed by an episode of heating coupled to slight compression until the baric peak is reached at around 760 ◦ C and 11.9 kbar (Fig. 14a). Subsequently, heating with decompression occurs until 800 ◦ C and 11 kbar (Fig. 14a). The path for sample 298E starts at 760 ◦ C and 12.5 kbar and follows a strong decompression (up to 3 kbar) with subordinate heating (40 ◦ C), reaching the conditions of 800 ◦ C and 9 kbar (Fig. 14a). 5. Geochronology 5.1. EPMA Th–U–PbT monazite dating Samples 056 and 297D of the Higher-TCF, both from the sillimanite–K-feldspar zone, were selected for EPMA monazite dat-

F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

St Ms Grt Bt As L

(a)

10 2 3

Ky l Si

10

GA

Ky And

500

133A

(c)

164B

8

core 6

rim

1

5

(b)

SP 4 St As Chl M Bt V s

Pressure (kbar)

Ms Bt V St Grt As L

15

Ms A s K Ab fs L

280

s s M tA S t rt B G

600

2

V

242B 700

Temperature(ºC)

800

400

600

800

1000

Temperature(ºC)

600

800

1000

Temperature(ºC)

Fig. 11. (a) P–T diagram with selected NKFMASH reaction (after Spear et al., 1999), avPT THERMOCALC result (grey ellipse) and GASP barometer results (calibrations: (1) Koziol and Newton, 1988; (2) Newton and Haselton, 1981; (3) Hodges and Crowley, 1985) for the sample 242B, Higher-TCF, St–Ky zone. (b–c) Results of the hornblende–plagioclase thermometer of Holland and Blundy (1994) (lines) and the amphibole thermobarometer of Zenk and Schulz (2004) (squares) for samples 164B and 133A.

ing. Sample 297D is a paragneiss that lacks evidence for retrograde metamorphic overprint and was selected to tentatively date the peak metamorphic conditions. Sample 056 comprises a paragneiss that has been intensely retrograded along the Serra do Azeite Shear Zone (Barra do Turvo Area) and was selected to tentatively date this retrograde event. In both samples monazite and apatite are abundant in the matrix and allanite and xenotime absent in the matrix and as inclusions. Wavelength-dispersive spectroscopy (WDS) spot analyses for representative monazite grains were performed with a JEOL JXA8600S EPMA under analytical conditions of 15 kV, 300 nA and 5 ␮m for the column accelerating voltage, beam current and beam diameter, respectively. Compositional zoning was monitored with BSE imaging, performed at 15 kV and 20 nA. A review of the chemical dating method, analytical protocol details (e.g., element set-ups, standards, counting times) and data treatment are presented in Vlach (2010). Attained Pb, Th, and U detection limits are close to 90, 100, and 100 ppm, respectively. An in-house 587 Ma natural monazite standard (E 0013A; Janasi et al., 2003) was used for EPMA U–Th–PbT dating calibration and the standard data are reported in Table S1 (Supplementary material). Th, U, and total Pb measurements and spot ages are presented in Table 5. Weighted average ages for monazite populations were calculated using the Isoplot 3 software (Ludwig, 2003). Sample 297D contains monazite grains exclusive to the matrix. The grains are subhedral, elongated (25 × 45 ␮m to 40 × 95 ␮m) and generally aligned with the mylonitic foliation (Fig. 15a–d), suggesting synkinematic growth. Four monazite grains were selected for BSE imaging and WDS spot analyses (Fig. 15a–d). The grains display a weak chemical variation, with Y2 O3 and ThO2 contents primarily in the range of 2–2.8 and 2–3.3 wt.%, respectively (Table 5; Figs. 15c–d and 16a–c). Two grains show rimward decrease of Y2 O3 content and rimward increase of ThO2 content (Fig. 16a–c). Sample 056 contains abundant monazite grains in the matrix and subordinate grains included or partially included in garnet, feldspar or quartz. The monazite grains are subhedral, subequant (∼50 × 55–60 ␮m) to elongated (∼35–40 × 60–80 ␮m), generally aligned with the mylonitic foliation (Fig. 15e, g, h), indicating synkinematic growth. Six monazite grains of sample 056 were selected to analyse, including grains partially included in chlorite (after biotite) and quartz (Fig. 15e), grains surrounded by muscovite ± quartz (Fig. 15g–j) and a grain included in garnet (Fig. 15f). BSE images and WDS analyses reveal up to three distinct compositional zones (Z1–Z3; Fig. 15e–j) with characteristic variation in Y2 O3 and ThO2 contents (Fig. 16d, e). One (Fig. 15i), two (Fig. 15e,

h) or three zones (Fig. 15f, j) can occur in a single grain and only Z3 is present in all analysed grains as outer rims or homogeneous grains. Z2 (core) is enriched in ThO2 and depleted in Y2 O3 in relation to Z3 (5.5–7.2 against 4.2–5.0 wt.% and 0.2–0.7 against 0.5–1.5 wt.%, respectively) (Fig. 16d, e). Z1 (inner core) is depleted in ThO2 in relation to Z3 (2.7–3.6 against 4.2–5.0 wt.%) (Fig. 15e, j). Since monazite in sample 297D is a stable phase restricted to the matrix, elongated grains are aligned with the hightemperature foliation and weakly zoned monazite grains show rimward decrease in Y, we infer that monazite grew during the prograde metamorphism. Taking into account the strong affinity of garnet for Y (e.g., Pyle and Spear, 2003) the Y zonation combined with the absence of monazite inclusions in garnet suggest that monazite grew during the late stage of garnet growth and before the stage of garnet consumption (indicated by microstructural evidence), and thus during near-peak P–T conditions. Monazite from sample 056 shows a more complex history, with at least three distinct growth phases (Z1, Z2 and Z3). As matrix monazite is abundant and only one grain included in garnet was found, monazite should have grew during near-peak metamorphism and during the retrograde path. A general decrease in Y content from Z1 to Z2 is consistent with monazite crystallisation during the prograde metamorphism. However, a concomitant strong increase in Th content suggests the breakdown of a Th-bearing phase (allanite?) or influx of Th-rich fluids. An increase in Y content in Z3 monazite (Fig. 16d, e) suggests that it grew during a retrograde event with garnet consumption. This is reinforced by the fact that the Z3 of the monazite grain included in garnet, which is interconnected with the matrix through a microfracture (Fig. 15f), shows the highest Y2 O3 content of sample 056 (Fig. 16d). Twenty-four analyses of monazite from sample 297D define a weighted average date of 589 ± 12 Ma (±2 error; MSWD = 0.91; Fig. 16f), which is inferred as a near-peak metamorphic age for this sample. For Z2 and Z3 monazites (sample 056), nine and 21 analyses define weighted mean dates of 597 ± 12 Ma (MSWD = 1.4) and 579 ± 8 Ma (MSWD = 1.15), respectively (Fig. 16g, h), inferred as a prograde near-peak metamorphic age and a retrograde monazite growth age, respectively. 5.2.

40 Ar–39 Ar

biotite dating

Biotite grains from one hornblende–biotite paragneiss (298A) from the TCF were selected for 40 Ar–39 Ar analysis. Ten biotite grains of 0.5–2 mm in size were placed into aluminium irradiation

F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

281

Fig. 12. NCKFMASHTi phase diagrams calculated with the Perple X software for samples 158B (a) and 129A (c). The water content was chosen so that rocks were H2 Osaturated immediately below the solidus (cf. White et al., 2001). The fields of equilibrium peak assemblages are superimposed hachured domains. Also shown are avPT THERMOCALC results (superimposed ellipses) and the results of GASP barometer (superimposed lines) (calibration: (1) Koziol and Newton, 1988; (2) and (4) Hodges and Crowley, 1985; (3) Newton and Haselton, 1981). Subplots (b) and (d) show calculated isopleths of grossular (sample 158B) and almandine and grossular (sample 129A), respectively. The superimposed black stars represent the best fit peak P–T conditions of each sample estimated taking into account the combination of all results.

282

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Fig. 13. NCKFMASHTi phase diagrams calculated with the Perple X software for samples 297D (a) and 298E (c). The water content was chosen so that rocks were H2 Osaturated immediately below the solidus (cf. White et al., 2001). The fields of equilibrium peak assemblages are superimposed hachured domains. Also shown are avPT THERMOCALC results (superimposed ellipses) (a–b), the results of the GASP barometer (superimposed lines) (calibration: (1) Koziol and Newton, 1988; (2) and (4) Hodges and Crowley, 1985; (3) Newton and Haselton, 1981) for sample 297D (b), and the best fit peak P–T conditions (black stars). The superimposed grey (a) and white (b) arrows are the inferred metamorphic paths.

disks along with Fish Canyon sanidine standards (28.02 ± 0.28 Ma; Renne et al., 1998). Two aliquots were analysed by laser incremental 40 Ar–39 Ar step heating following procedures detailed in Vasconcelos et al. (2002). Mass spectrometric analyses (Table 6) were performed at the University of Queensland Argon Geochronology in Earth Sciences laboratory. From the two analysed grains of biotite, one produced a plateau age (cf., Fleck et al., 1977) of 553 ± 4 Ma (Fig. 17a). The other grain

yielded a six steps pseudo-plateau age of 558 ± 5 Ma (Fig. 17a) with 48.3% of the total 39 Ar released. An age probability diagram plotted for the plateau and pseudo-plateau steps of both grains yields a maximum probability peak at 554.9 Ma (Fig. 17b), and it defines a mean-weighted age of 555 ± 4 Ma, which is compatible at the 2 level with the ages given by the plateaus. The latter value is inferred as the age of the biotite isotopic closure at temperatures below 300–250 ◦ C.

15

(a)

(b) 50

40

ies fac

30

Depth (km)

Pressure (kbar)

10

20

5

10

0 300

500 700 Temperature (ºC)

900 200

600 Temperature (ºC)

0 1000

Fig. 14. (a) P–T paths calculated from zoned garnet and plagioclase of samples 242B (St–Ky zone), 158B (Ky–Kfs zone), 129A (Ky–Kfs zone) and 298E (Sil–Kfs zone) using the Gibbs method. Metamorphic facies fields from Brown (2001). (b) Generalised P–T paths for the Turvo-Cajati Formation (TC) and the Atuba Complex (AC), and for the Apiaí Terrane (ApT, Faleiros et al., 2010). Geochronologic data from: (1) Th–U–PbT chemical dating of monazite, this work, (2) U–Pb zircon ages, Sato et al. (2003, 2009), (3) Ar–Ar hornblende and biotite ages, Machado et al. (2007).

F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

6. Discussion

283

Wakabayashi, 2004). Similar metamorphic paths have also been described for many collisional orogens, such as the Limpopo Belt of South Africa (Zeh et al., 2004) and the NE Greenland Caledonites (Jones and Strachan, 2000). The preserved portions of the prograde metamorphic paths from the kyanite migmatites (samples 242B, 158B and 129A) of the Higher-TCF start at approximately 500–600 ◦ C and 10–12 kbar, within the field of eclogite facies metamorphism (Fig. 14a). Considering a lithostatic gradient of 3.75 km kbar−1 the maximum pressure attained by each sample indicates maximum burial depth between 38 and 45 km (Fig. 14b). No record of the tectonic burial stage was preserved in the kyanite migmatites, but the prograde paths suggest that these rocks evolved from blueschist or low-temperature eclogite facies condi-

6.1. Metamorphic signatures and tectonic implications The two types of P–T paths calculated for the Higher-TCF (isobaric heating and isothermal decompression) can be inferred as parts of a general clockwise path. The generalised P–T paths, including the retrograde portions inferred from petrographic evidence, are shown in Fig. 14b. The near isobaric heating at relatively high-pressure conditions (10–12 kbar), reconstructed for the kyanite-bearing rocks (Fig. 14b), is similar to the predictions of collision-related metamorphic paths related with crustal thickening (e.g., England and Thompson, 1984; Beaumont et al., 2001; Table 5 Th, U and Pb chemical data and spot ages of monazite from samples 056 and 297D. Sample-grain point analysis

ThO2 (wt%)

Y2 O3 (wt%)

056-Mnz1-P1 056-Mnz1-P2 056-Mnz1-P3 056-Mnz3-P1 056-Mnz3-P10 056-Mnz3-P2 056-Mnz3-P3 056-Mnz3-P4 056-Mnz3-P5 056-Mnz3-P6 056-Mnz3-P7 056-Mnz3-P8 056-Mnz3-P9 056-Mnz4-P1 056-Mnz4-P2 056-Mnz4-P3 056-Mnz4-P4 056-Mnz5-P1 056-Mnz5-P10 056-Mnz5-P11 056-Mnz5-P2 056-Mnz5-P3 056-Mnz5-P4 056-Mnz5-P5 056-Mnz5-P6 056-Mnz5-P7 056-Mnz5-P8 056-Mnz5-P9 056-Mnz6-P1 056-Mnz6-P2 056-Mnz7-P1 056-Mnz7-P2 056-Mnz7-P3 056-Mnz7-P4 297D-Mnz1-P6 297D-Mnz1-P7 297D-Mnz1-P8 297D-Mnz2-P7 297D-Mnz2-P8 297D-Mnz4-P1 297D-Mnz4-P2 297D-Mnz4-P3 297D-Mnz1-P1 297D-Mnz1-P2 297D-Mnz1-P3 297D-Mnz1-P4 297D-Mnz1-P5 297D-Mnz12-P1 297D-Mnz12-P2 297D-Mnz12-P3 297D-Mnz12-P4 297D-Mnz12-P5 297D-Mnz2-P1 297D-Mnz2-P2 297D-Mnz2-P3 297D-Mnz2-P4 297D-Mnz2-P5 297D-Mnz2-P6

3.42 3.63 4.73 4.42 4.38 6.85 6.11 5.53 7.19 4.44 4.24 5.54 3.24 4.97 4.49 4.69 4.16 4.40 5.03 4.59 4.54 6.93 4.32 4.52 4.43 4.54 4.89 6.83 4.54 4.38 4.54 6.53 2.74 6.43 2.28 2.81 1.76 2.67 2.99 2.85 3.00 2.91 1.94 2.23 2.03 1.05 1.66 3.26 3.73 3.09 3.10 2.91 2.70 2.72 2.52 2.61 3.03 2.76

0.69 0.69 0.57 1.33 1.48 0.71 0.42 0.09 0.32 1.23 0.24 0.40 0.54 0.52 1.04 0.13 0.10 0.51 0.55 0.48 0.64 0.18 0.35 0.55 0.54 0.74 0.49 0.21 0.74 0.76 0.83 0.70 0.46 0.72 2.63 2.58 2.23 2.09 1.73 1.83 0.84 1.16 2.14 2.77 2.33 2.17 2.04 2.29 2.39 2.19 2.19 2.23 2.49 2.44 2.34 2.31 1.64 2.40

Th (ppm) 30,030 31,920 41,580 38,850 38,520 60,200 53,690 48,610 63,180 38,980 37,280 48,670 28,480 43,640 39,490 41,170 36,540 38,620 44,210 40,310 39,910 60,910 37,960 39,750 38,960 39,900 42,960 60,040 39,850 38,480 39,880 57,340 24,080 56,460 20,040 24,720 15,470 23,490 26,270 25,050 26,330 25,560 17,060 19,620 17,810 9,270 14,580 28,630 32,780 27,110 27,250 25,570 23,750 23,860 22,170 22,920 26,590 24,210

2Sig 681 719 919 861 855 1308 1172 1063 1370 863 828 1064 648 960 876 911 815 856 975 891 884 1321 845 881 863 884 948 1305 883 854 883 1247 557 1227 478 573 386 547 605 579 606 588 417 471 430 265 365 651 739 620 622 588 551 553 519 532 611 559

U (ppm)

2Sig

2251 2502 7002 3010 3084 2320 4247 1860 1870 3549 3006 2639 2797 3721 3784 2736 3324 5613 5325 5625 6189 2793 5459 7051 6919 6359 5258 3472 4640 3274 3330 1560 1412 1819 1584 1356 1051 1728 1710 1482 1219 1137 1485 1798 1317 1135 890 3456 2193 1251 1590 1167 1886 2115 1832 1784 2117 1941

121 124 197 138 139 125 154 120 121 144 138 135 135 146 146 136 141 170 167 170 178 138 168 198 196 188 166 145 157 141 141 118 115 120 116 115 113 117 117 115 114 114 115 117 114 113 112 142 121 114 116 114 118 120 117 117 120 118

Pb (ppm) 947 985 1692 1221 1333 1773 1827 1482 1843 1306 1166 1604 988 1433 1317 1258 1245 1438 1650 1588 1639 1923 1418 1705 1606 1543 1615 1972 1444 1254 1309 1667 717 1597 645 733 554 803 841 787 783 785 547 728 595 354 446 975 1055 794 823 753 829 842 750 851 916 834

2Sig

Th–U–PbT age (Ma)

83 83 87 84 85 88 89 86 89 85 84 87 83 86 85 84 84 86 87 87 87 89 85 88 87 86 87 90 86 84 85 87 72 87 72 82 71 82 82 82 82 82 71 72 72 71 71 83 84 82 82 72 82 83 72 83 83 82

563 546 582 557 608 581 600 601 591 573 550 621 583 570 564 558 583 561 594 600 604 609 564 602 579 565 595 613 582 567 572 593 555 569 568 559 648 611 586 584 574 595 555 632 597 605 567 543 586 565 564 569 615 607 591 655 606 605

2Sig 47 44 27 36 36 26 26 32 25 35 37 30 46 31 34 35 37 31 29 30 29 25 31 28 29 29 29 25 32 36 35 28 53 28 60 60 80 60 55 58 57 59 69 59 69 115 87 44 44 56 54 52 58 56 54 60 52 57

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Fig. 15. Backscatter electron images of the analysed monazite grains of samples 297D (a–d) and 056 (e–j) (Sil–Kfs zone, Turvo-Cajati Formation). The top of each figure shows the relationship between the monazite grain and other minerals in paragenesis, and the bottom shows enlarged high-contrast BSE images (the lighter grey scale, the higher Th content). The top of (c) is an optical photomicrograph. Numbers of samples and analysed grains are indicated on the right side of the high-contrast BSE images. White circles with white numbers mark the position of EMPA analyses shown in Table 5, which is labelled in black with wt% Y2 O3 and ThO2 in brackets for some images (c–e, g, i–j). Also shown are the positions of chemical profiles across the monazite grains Mnz01 and Mnz02 (sample 297D) and Mnz03 and Mnz05 (sample 056) shown in Fig. 16a–e. Z1, Z2 and Z3 (e–j) are chemical zones of monazite grains of sample 056 (see text for explanation).

tions. Near isobaric heating suggests that the thermal relaxation of the thickened crust occurred in the absence of uplift/erosion until reaching conditions close to the steady-state geotherm supported by the thickened crust. Rocks from the sillimanite–K-feldspar zone (samples 056, 297D and 298E), which represent the dominant metamorphic zone of the Higher-TCF, show a contrasting metamorphic history compared with the kyanite migmatites (samples 242B, 158B and 129A), indicating that the latter represents a different tectonic sliver. In the rocks of the sillimanite–K-feldspar zone only the metamorphic record of exhumation was preserved. The preserved portion of the prograde metamorphic path from the sample 298E starts at approximately 760 ◦ C at 12.5 kbar, corresponding to the transitional conditions between high-pressure amphibolite and eclogite facies metamorphism. The maximum pressure recorded in sample 298E suggests that this rock was buried to a minimal depth of approximately 47 km, while the absence of records of the burial and thermal relaxation parts of the metamorphic path suggests that this rock was buried to significantly deeper levels (e.g., England and Thompson, 1984). Metamorphic assemblages and near-peak metamorphic conditions for three Higher-TCF samples (129A, 297D and 298E) (∼800 ◦ C and 11–12.5 kbar) are indicators of high-pressure granulite facies

rocks (cf., O’Brien and Rötzler, 2003). According to O’Brien and Rötzler (2003), high-pressure granulites mostly represent rocks formed as a result of short-lived tectonic events that led to crustal thickening or subduction of the crust into the mantle. For the Higher-TCF rocks, a short-lived tectonic event of crustal thickening is supported by (i) the very well preserved prograde zoning patterns in garnet and plagioclase from the kyanite migmatites (Fig. 9a–c, g, h) (e.g., O’Brien and Rötzler, 2003) and (ii) the generalised preservation of microstructures indicative of the former presence of melt (at the grain-scale) in regionally metamorphosed rocks (Fig. 8d–h) (e.g., Sawyer, 2001). The baric regime of the metamorphic event that affected the Atuba Complex (peak pressure of 6–7 kbar, maximum depth of burial of 22–26 km) contrasts with the baric regime of the underlying Higher-TCF (Fig. 14b). Thermobarometric results indicate cooling from ∼730 to ∼630 ◦ C with slightly decreasing pressure (Fig. 14b), which contrasts with a path of heating-decompression predicted for the hanging wall rocks of a nappe system in which colder rocks overlies hotter rocks (Spear et al., 1984). On the other hand, commonly the contacts between the TCF and the AC are marked by metamorphic overprint of a greenschist facies mylonitic foliation in a high-temperature mylonitic foliation, suggesting that the high-grade metamorphic events of both units could not have

Table 6 40 Ar/39 Ar biotite analytical results for sample 298A, J factors, system blanks, and 40 Ar/36 Ar discrimination for the UQ-AGES. The uncertainties are given at the 1 level. Run

Ar/39 Ar

37

Ar/39 Ar

38

Ar/39 Ar

40

Ar/39 Ar

40

Ar* /39 Ar

%Ar40*

Age Ma

EMV

Ar39

Ar40

J

Ar40 Disc

ID

Error 1

Error 1

Error 1

Error 1

Error 1

Error 1

Error 1

Axial

Moles

Moles

Error 1

Error 1

01A 01B 01C 01D 01E 01F 01G 01H 01I 01J 01K 01L 01M 01N 02A 02B 02C 02D 02E 02F 02G 02H 02I 02J 02K 02L

0.0019 ± 0.00041 0.0004 ± 0.00056 0.0002 ± 0.00069 0.0013 ± 0.00075 0.0003 ± 0.00072 0.0005 ± 0.00075 0.0002 ± 0.00080 0.001 ± 0.00090 0.0001 ± 0.00096 0.0004 ± 0.00089 0.0015 ± 0.00085 0.00 ± 0.00440 0.0011 ± 0.00580 0.0013 ± 0.00340 0.0018 ± 0.00034 0.0006 ± 0.00033 0.0004 ± 0.00042 0.0013 ± 0.00060 0.001 ± 0.00032 0.0007 ± 0.00035 0.0007 ± 0.00050 0 ± 0.00048 0.0004 ± 0.00039 3E-05 ± 0.00050 0.004 ± 0.01000 0.0044 ± 0.00390

0.149 ± 0.011 0.165 ± 0.017 0.233 ± 0.031 0.339 ± 0.040 0.316 ± 0.049 0.278 ± 0.035 0.207 ± 0.032 0.198 ± 0.039 0.232 ± 0.049 0.165 ± 0.063 0.131 ± 0.040 0.77 ± 0.310 0.49 ± 0.310 0.23 ± 0.240 0.001 ± 0.013 0.03 ± 0.016 0.077 ± 0.021 0.052 ± 0.031 0.043 ± 0.021 0.041 ± 0.018 0.031 ± 0.020 0.048 ± 0.022 0.048 ± 0.017 0.086 ± 0.019 0.34 ± 0.580 0.37 ± 0.220

0.0022 ± 0.00010 0.0017 ± 0.00018 0.0019 ± 0.00023 0.0017 ± 0.00036 0.0015 ± 0.00033 0.0014 ± 0.00031 0.0012 ± 0.00031 0.0012 ± 0.00030 0.001 ± 0.00036 0.0019 ± 0.00038 0.0015 ± 0.00035 0.0031 ± 0.00160 0.0032 ± 0.00180 0.0014 ± 0.00140 0.0014 ± 0.00014 0.0012 ± 0.00011 0.0009 ± 0.00018 0.0018 ± 0.00023 0.0011 ± 0.00016 0.001 ± 0.00016 0.0011 ± 0.00020 0.0014 ± 0.00019 0.0013 ± 0.00020 0.0023 ± 0.00022 0.0043 ± 0.00260 0.0021 ± 0.00120

101.42 ± 0.55 96.8 ± 0.54 97.28 ± 0.78 96.91 ± 0.89 98.09 ± 0.92 96.35 ± 0.71 96.98 ± 0.73 94.26 ± 0.79 97.59 ± 0.98 97.9 ± 1.10 95.61 ± 0.80 99.4 ± 3.50 73.7 ± 2.60 85 ± 2.20 98.17 ± 0.48 96.14 ± 0.53 96.14 ± 0.53 96.33 ± 0.66 96.46 ± 0.59 95.36 ± 0.51 96.13 ± 0.60 94.68 ± 0.52 97.78 ± 0.54 98.01 ± 0.54 115.6 ± 6.90 109.1 ± 2.50

100.89 ± 0.57 96.72 ± 0.57 97.25 ± 0.81 96.58 ± 0.92 98.04 ± 0.95 96.25 ± 0.74 96.96 ± 0.77 93.98 ± 0.84 97.6 ± 1.00 97.8 ± 1.10 95.19 ± 0.84 99.5 ± 3.8 73.5 ± 3.1 84.7 ± 2.4 97.64 ± 0.5 95.96 ± 0.54 96.03 ± 0.54 95.96 ± 0.68 96.16 ± 0.59 95.15 ± 0.52 95.93 ± 0.62 94.68 ± 0.54 97.68 ± 0.55 98.01 ± 0.56 114.6 ± 7.5 107.8 ± 2.7

99.47 ± 0.21 99.9 ± 0.22 99.96 ± 0.27 99.63 ± 0.32 99.93 ± 0.29 99.87 ± 0.29 99.96 ± 0.30 99.69 ± 0.35 99.98 ± 0.36 99.88 ± 0.37 99.55 ± 0.34 100.1 ± 1.4 99.6 ± 2.4 99.6 ± 1.3 99.46 ± 0.16 99.81 ± 0.18 99.88 ± 0.2 99.61 ± 0.27 99.69 ± 0.2 99.78 ± 0.18 99.78 ± 0.2 100 ± 0.22 99.89 ± 0.17 100 ± 0.22 99 ± 2.6 98.8 ± 1.1

578.1 ± 2.8 557.5 ± 2.8 560.2 ± 4.0 556.8 ± 4.6 564.1 ± 4.7 555.2 ± 3.7 558.7 ± 3.8 543.9 ± 4.2 561.8 ± 5.1 562.7 ± 5.4 549.9 ± 4.2 571 ± 19. 438 ± 17. 497 ± 12. 562.1 ± 2.5 553.8 ± 2.7 554.1 ± 2.7 553.8 ± 3.4 554.8 ± 3.0 549.7 ± 2.6 553.6 ± 3.1 547.4 ± 2.7 562.3 ± 2.7 563.9 ± 2.8 644 ± 35. 612 ± 13.

1.533687 1.533692 1.533685 1.533691 1.533692 1.533689 1.533684 1.533681 1.533679 1.533666 1.533663 1.533664 1.533670 1.533668 1.533674 1.533676 1.533676 1.533674 1.533682 1.533682 1.533685 1.533691 1.533688 1.533694 1.533694 1.533681

6.72E-16 3.43E-16 2.05E-16 1.42E-16 1.33E-16 2.09E-16 1.90E-16 1.71E-16 1.33E-16 1.11E-16 1.50E-16 2.15E-17 2.03E-17 2.87E-17 5.66E-16 4.66E-16 3.22E-16 2.40E-16 3.52E-16 4.07E-16 3.45E-16 3.20E-16 3.92E-16 3.93E-16 1.25E-17 3.53E-17

6.81E-14 3.32E-14 2.00E-14 1.37E-14 1.30E-14 2.01E-14 1.85E-14 1.61E-14 1.30E-14 1.09E-14 1.44E-14 2.13E-15 1.49E-15 2.44E-15 5.56E-14 4.48E-14 3.09E-14 2.31E-14 3.40E-14 3.88E-14 3.32E-14 3.03E-14 3.83E-14 3.86E-14 1.44E-15 3.85E-15

0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014 0.003744 ± 0.000014

1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 1.0016 ± 0.0034 0.9991 ± 0.0034 0.9991 ± 0.0034 0.9991 ± 0.0034 0.9991 ± 0.0034 0.9991 ± 0.0034 0.9991 ± 0.0034 0.9991 ± 0.0034 0.9991 ± 0.0034 0.9991 ± 0.0034 0.9991 ± 0.0034 0.9991 ± 0.0034 0.9991 ± 0.0034

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40

36

Ar* is the total amount of radiogenic 40 Ar obtained by substracting the atmospheric 40 Ar and 40 ArK from the total 40 Ar measured from the sample.

285

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F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

4 297D-Mnz01 3

6 4

5

4 297D-Mnz02 4 297D-Mnz01 (b)

(a) core

rim

rim

rim

8 3

3

2

1

3 8 4

1

2

2

2

1

1

1

0

20 40 60 distance (µm)

80 0

20 40 0 distance (µm)

740

297D - near-peak metamorphic age

(f) 660

8 056-Mnz03

(c) core

7

3

2

core

3

rim 4

Z1 8

4

YO ThO UO PbO

10

10

Age (Ma)

6

1

2

4 Z3

20 30 40 distance (µm)

50

10

8

4

Z2

Z3

2

2 0

20 40 distance (µm)

(g) 650

056 - prograde metamorphic age

60 0

20 40 60 distance (µm)

80

(h)

056 - retrograde metamorphic age

630

640 660

5

1

7

9

(e)

9

3

6

3

8

8 056-Mnz05

Z3 2

6

5

(d)

Z2

610

620

590

600 580

570 580 N=24

500

Mean = 589±12 [2.0%] 95% conf. MSWD = 0.91 Probability = 0.59

420

550 N=9

560

Mean = 597±12 [2.0%] 95% conf. MSWD = 1.4 Probability = 0.20

540

530 510

N=21 Mean = 579±8 [1.3%] 95% conf. MSWD = 1.15 Probability = 0.29

490

520

Fig. 16. (a–e) Chemical profiles across monazite grains of samples 297D and 056 (Sil–Kfs zone, Turvo-Cajati Formation). (f–h) Monazite chemical ages of samples 297D and 056 (Turvo-Cajati Formation). The mean ages (at 95% confidence level) were calculated using the weighted average tool of Isoplot 3.0, considering only the internal errors. Grey rectangles are ages from individual spot analyses with 2 errors bars determined from counting statistics.

(a) 100

Sample 298A

Apparent Age (Ma)

600

K/Ca

100 10 1 0.1

0

558 ± 5 Ma

553 ± 4 Ma 500 0

(b)

50 Cumulative % Ar released

0.6 0.4 0.2 0

100

Sample 298A

5

Probability

1.2

Analysis #

15

554.9

0.8

555 ± 4 Ma MSWD=1.10 Prob.=0.36

0.4

0 500

550

600

Age (Ma)

Fig. 17. (a) Incremental-heating spectra for two biotite grains of the sample 298A (Sil–Kfs zone, Turvo-Cajati Formation), and corresponding ideogram and cumulative probability plot (b).

been related in space and may be not related in time (see discussion in Sections 9.2 and 9.3). The metamorphic signatures from the TCF are also strongly contrasting in relation to the patterns recognised for the Apiaí Terrane that is juxtaposed with the Curitiba Terrane along the LancinhaCubatão Fault Zone (Fig. 1), which is characterised by Barrovian metamorphism ranging from lower-greenschist to amphibolite facies conditions (Faleiros et al., 2010). The metamorphic paths calculated with the Gibbs method indicate a contrasting tectonic regime (Fig. 14b). The P–T path calculated by Faleiros et al. (2010) for sample 198A (Fig. 14b) provides evidence for the burial stage (heating-compression stage) until reaching peak pressure conditions of approximately 8.4 kbar at 530 ◦ C (maximum depth of burial of 31 km). This is followed by an exhumation-heating stage until the peak thermal conditions are reached: 550 ◦ C at 6.9 kbar (Fig. 14b). The fact that the Lancinha-Cubatão Fault Zone was responsible for a greenschist facies metamorphic overprint on the units of the Curitiba and the Apiaí Terranes suggests that the juxtaposition occurred after the main metamorphic events of each terrane. 6.2. Nappe stacking, strike-slip faulting and exhumation history Metamorphic stack of units such as observed in the Curitiba Terrane (Cajati area), where lower-grade rocks override higher-grade rocks (Fig. 6b), is normally used as evidence for normal faulting (Ring et al., 1999). However, the presence of older rocks of the AC (Archaean to Paleoproterozoic) overlying younger rocks of the Higher-TCF (Neoproterozoic) in the Cajati area (Fig. 6b) is unequivocal evidence for an episode of thrusting during the Neoproterozoic. On the other hand, the metamorphic paths of the Higher-TCF and the AC (Fig. 14b) do not fit a synmetamorphic nappe stacking model where cold rocks (AC) override hot rocks (Higher-TCF), in which the hanging wall rocks would undergo heating with decompression (versus observations of simultaneous decrease in temperature and pressure in the AC) and the footwall rocks would undergo simultaneous cooling and compression (versus observations of isothermal decompression in the sillimanite–K-feldspar zone, the dominant

F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

rock assemblage of the Higher-TCF) (Spear et al., 1984). These data indicate that the metamorphism related with crustal thickening recorded in the High-TCF rocks was not contemporaneous with the nappe stacking responsible for the final juxtaposition of the units in the study area. Considering all these features combined with the strong metamorphic overprint of a greenschist facies mylonitic foliation in a previous high-temperature mylonitic foliation along the contacts between the AC and the TCF we interpret that the nappe emplacement in the Curitiba Terrane occurred during postmetamorphic exhumation tectonics, as demonstrated for others tectonic settings such as the Menderes nappes of western Turkey (Ring et al., 2001) and the Bronson Hill anticlinorium in central New England (Spear et al., 2008). This interpretation is reinforced by (i) the fact that regionally the TCF and the AC underwent moderate to strong greenschist facies metamorphic overprint also outside the contact zones and outside the late strike-slip shear zones; and (ii) the well preserved high-grade rocks of the Higher-TCF present in the Cajati area represent an exception. Thus, the episode of postmetamorphic nappe stacking occurred after a first episode of rapid exhumation, as indicated by the isothermal decompression recorded in the sillimanite–K-feldspar zone rocks of the Higher-TCF. Possibly this rapid exhumation was associated with the installation of the transcurrent shear system in the Ribeira Belt, as described for similar settings (e.g., Ailao Shan-Red River shear zone in SE Asia, Leloup et al., 2001). Corroborate this interpretation the fact that mylonitic rocks of some transcurrent shear zones in the southern Ribeira Belt preserve evidence of an early phase of mylonitisation under moderate-temperature (500–600 ◦ C) replaced by a late lowtemperature fabric, as the Ribeira shear zone (Fig. 2) in the Apiaí Terrane (Faleiros et al., 2007, 2010) and the Putunã Shear Zone (Figs. 2, 4a, 6b) in the Curitiba Terrane (Faleiros and Faleiros, 2008). The sinistral transcurrent shear zones in the interior of the Curitiba Terrane also should have contributed to the late exhumation, once these faults dip southward and show an ESE-plunging stretching lineation indicating top-to-the-ESE extensional component. In this scenario the opposite transcurrent shearing along the LancinhaCubatão Fault Zone (dextral) to the north of the Curitiba Terrane and the Serra do Azeite Shear Zone (sinistral) to the south suggest that these faults could have been acted as lateral ramps for the west-directed thrusting in the Curitiba Terrane. An alternative model of restricted transtensional deformation associated with a regional transpression was presented by Dehler et al. (2007) for the Cajati area, but this model did not take into account the metamorphic evidence and failed to explain the stratigraphic inversion with the AC overlying the Higher-TCF and the pattern of regional fold.

6.3. Timing of accretionary, collisional and exhumation stages in the southern Ribeira Belt The Apiaí and the Embu Terranes (Fig. 1) were intruded by voluminous arc-related granitic batholiths (Cunhaporanga, Três Córregos and Agudos Grandes Batholiths) between 630 and 600 Ma (U–Pb zircon data of Gimenez Filho et al., 2000; Janasi et al., 2001; Prazeres Filho, 2005; Silva et al., 2005; Leite et al., 2007), setting the subduction-related magmatic stage in the southern Ribeira Belt. The Agudos Grandes Batholith intruded the contacts between the Apiaí and the Embu Terranes, indicating that these terranes were already juxtaposed during 630–600 Ma. Field relationships between the metamorphic rocks and the granitic batholiths (e.g., crosscutting relationships, contact aureoles) suggest that there are metamorphic events older than 630–600 Ma in the units of the Apiaí Terrane (Faleiros et al., 2010). The Rio Piên arc-related suite (Fig. 2), previously considered as a part the Curitiba Terrane (Siga Junior, 1995), was formed at the same time interval (630–600 Ma; U–Pb zircon and titanite data from Harara, 2001). However, the

287

relationship between the Rio Piên suite and the adjacent units (Curitiba and Luís Alves Terranes) is still debated. The monazite chemical age obtained for the Higher-TCF records a deep collisional metamorphism (40–47 km of depth) with climax at 589 ± 12 Ma, which represents the main collisional stage of the southern Ribeira Belt that lead to crustal thickening. The age of 597 ± 12 Ma obtained in monazite cores of sample 056 of the Higher-TCF (sillimanite–K-feldspar zone) can be inferred as related with a period of the prograde metamorphism before the peak conditions. The geochronological data from the AC (U–Pb zircon, Rb–Sr, Sm–Nd dates; Sato et al., 2003, 2009) provided inaccurate Brasiliano ages between 646 and 525 Ma related with a high-temperature metamorphic (partial melting) event. A U–Pb zircon age (lower intercept) of 584 ± 28 Ma (Sato et al., 2009) is inferred as the best estimate for this metamorphic event. However, a minimum metamorphic age of 594 ± 1 Ma is indicated by 40 Ar–39 Ar hornblende data obtained in the Cajati area (Machado et al., 2007), limiting to ca. 595–610 Ma the age of the high-grade event. These data suggest that the high-temperature metamorphic events from the Higher-TCF and the AC were coeval, although the evidence for postmetamorphic juxtaposition of these units (Section 9.2) may indicate that these metamorphic events were not spatially related. We interpret the monazite chemical age of 579 ± 8 Ma (monazite rims) obtained in sample 056, a paragneiss with strong greenschist facies metamorphic overprint associated with the Serra do Azeite Shear Zone, as the minimum age for the postmetamorphic juxtaposition of units in the Curitiba Terrane. The fact that monazite grains from this sample show a rimward increase in Y content, suggesting that the rims grew during a retrograde phase of garnet consumption, corroborates this interpretation. A minimum age for the juxtaposition in the Curitiba Terrane can also be inferred from the age of the Graciosa Province, an extensional A-type granitic suite (Gualda and Vlach, 2007) in which some plutons truncate the tectonic contacts between the AC and the TCF within the Curitiba Terrane and also the contacts between the Luís Alves and the Curitiba Terranes (e.g., Guaraú pluton, Fig. 2). Recent U–Pb zircon data (ID-TIMS) obtained for diorites coeval to the main granites in three plutons of the Graciosa Province, which were emplaced at pressures no greater than ∼2 kbar (maximum depth of ca. 7.5 km) (Gualda and Vlach, 2007), yield concordant ages of 580 ± 2 and 583 ± 3 Ma, interpreted as representative of magmatism for the Graciosa Province (Vlach et al., 2011). These data limit the final juxtaposition between the AC and TCF and between the Curitiba and the Luís Alves Terranes at ca. 580 Ma, indicating a short-lived evolution from crustal thickening to exhumation. Considering an average age of 590 Ma for the peak of collisional metamorphism recorded in sample 297D (at depth of ∼40–47 km) and an average age of 580 Ma for the intrusion of the granites from the Graciosa Province (at depth of ∼7.5 km) we calculate an exhumation rate of 3.9–3.2 mm year−1 for the Higher-TCF. Minimum exhumation rates of 1.7–1.4 mm year−1 were calculated considering the error bars of the geochronological data. For comparison, modern exhumation rates for the Himalaya vary between 0.8 ± 0.3 mm year−1 (southern edge of the Tibetan Plateau) and 2.7 ± 0.3 mm year−1 (High Himalaya) (Vance et al., 2003). The 40 Ar–39 Ar biotite age (555 ± 4 Ma) from sample 298A (Higher-TCF) combined with geochronological data from the Graciosa Province and petrological information indicate that the Higher-TCF rocks remained at temperatures between 400 and 300–250 ◦ C for a period of approximately 25 Ma. The thermochronological evolution of the Higher-TCF and the AC is shown in Fig. 18. Average cooling rates of 40 and 4 ◦ C/Ma for the periods between 590 and 580 Ma and between 580 and 555 Ma, respectively, are determined from the monazite and biotite ages of the Higher-TCF. Available petrological and geochronological data indicate strongly constrasting tectonic evolutions for the southern, cen-

288

F.M. Faleiros et al. / Precambrian Research 189 (2011) 263–291

1000

Temperature (ºC)

800

Data from: 1 - This work 2 - Campagnoli (1996) 3 - Machado et al. (2007)

U-Th-Pb (Mnz) - TCF

Turvo-Cajati Formation - TCF Atuba Complex - AC TCF AC

600 U-Th-Pb (Mnz) - TCF

Ar-Ar (Hbl) - AC

400 K-Ar (Bt) - CA

Ar-Ar (Ms) - AC

200 Ar-Ar (Bt) - TCF

500

550

600

650

Age (Ma) Fig. 18. Plot of temperature versus time for thermal history of the Curitiba Terrane units (Turvo-Cajati Formation and Atuba Complex) in the study area.

tral and central-northern portions of the Ribeira Belt. Studies on the Oriental Terrane of the central Ribeira Belt (approximately 600 km northeast from the present area), which was accreted to the reworked continental margin of the São Francisco Craton, reveal a complex history for the main collisional event in this sector, with three high-grade metamorphic episodes with ages of ca. 605, 580 and 550 Ma (U–Pb ages, Heilbron and Machado, 2003). Petrological studies in the Oriental Terrane show a 590–550 Ma high T/P metamorphism that reached peak conditions of 750–800 ◦ C at 7 kbar (Kühn et al., 2004). Geochronological and petrological studies on the São Fidelis-Santo Antônio de Pádua sector of the centralnorthern Ribeira Belt reveal an early collisonal stage between 623 ± 14 and 591 ± 17 Ma (U–Pb SHRIMP zircon ages), followed by a main metamorphic event with peak conditions of 850 ± 50 ◦ C at 8 ± 1 kbar in the period between 572 ± 13 and 562 ± 11 Ma (U–Pb SHRIMP zircon ages) (Bento dos Santos et al., 2010, 2011). Late granite with U–Pb SHRIMP zircon age of 493 ± 14 Ma was interpreted as an age constraint to the onset of the thermal collapse in this sector of the Ribeira Belt (Bento dos Santos et al., 2010). Schmitt et al. (2004) describe for the Cabo Frio domain of the central Ribeira Belt a 525–520 Ma collisional metamorphic event (U–Pb zircon ages) that reached peak condition in excess of 780 ◦ C and 9 kbar, followed by a late deformation episode related with high-temperature dextral transcurrent shear zones between 505 and 490 Ma (U–Pb zircon and monazite ages). Summarising, these data indicate that the central and central-northern portions of the Ribeira Belt underwent a longer tectonic history with coeval and younger metamorphic and magmatic events in relation to the southern Ribeira Belt, indicating a complex and diachronic collisional evolution for the Ribeira Belt. In this tectonic scenario it is most probable that the short-lived collisional metamorphic event recorded in the Curitiba Terrane was related with the collision of a microcontinent, in which possibly the Luís Alves Terrane is a remnant. 7. Tectonic evolution Based on new data presented in this work combined with previous works we summarise the main tectonic events recorded in the southern Ribeira Belt. During the period between 630 and 600 Ma a continental subduction-related magmatic arc was installed in the Apiaí and the Embu Terranes (the Paranapanema accrescionary orogen of Campos Neto, 2000), associated with a Barrovian metamorphism (Faleiros et al., 2010). Considering the localisation of the Paranapanema Craton (covered by the Paraná Basin) to the

north of the Apiaí and Embu Terranes (Fig. 1), we infer a northwestward subduction. The end of this period (∼600 Ma) marks the end of the subduction and the initiation of the collisional process. In the period between 600 and 590 Ma the southern Ribeira Belt underwent an oblique collisional regime. It is intuitive to relate the collision metamorphism (crustal thickening) recorded in the Turvo-Cajati Formation (from 597 ± 12 to 589 ± 12 Ma) and in the Atuba Complex (584 ± 28 Ma; Sato et al., 2009) with a collisional suture zone between the Luís Alves Terrane, which constitutes a cratonic fragment, and the Apiaí Terrane. Corroborate this interpretation the existence of possible ophiolite remnants with U–Pb zircon age of 618 ± 17 Ma described for the Piên mafic–ultramafic suite in the border between the Luís Alves and the Curitiba Terranes ∼100 km to the southwest of the present study area (Harara, 2001; Harara et al., 2004). However, it is clear from the available data that the current space relationships between the forming units of the Curitiba Terrane do not represent genetic relationships, and the Turvo-Cajati Formation, as well as the Atuba Complex, can be considered as exotic terranes accreted to the active margin bordering the Paranapanema Craton. The generalised greenschist facies metamorphic overprint present in the units of the Curitiba Terrane is related with the late strike-slip regime and it is recorded by a 579 ± 8 Ma monazite chemical age from an overprinted paragneiss from the Turvo-Cajati Formation. The A-type granitic rocks from the Graciosa Province (U–Pb zircon ages of 580 ± 2 and 583 ± 3 Ma, Vlach et al., 2011) is most probably related with the extensional branches of the strike-slip system, as it is the case of small pullapart basins of this period, as the Iporanga Formation (U–Pb zircon age of 579 ± 34 Ma, Campanha et al., 2008a) in the Apiaí Terrane. The latest ductile movements of the Lancinha-Cubatão Fault Zone are constrained at 534 ± 16 Ma by K–Ar dating of micas (M.A.S. Basei, personal communication), although ruptil-ductile deformation may have continued until about 500 Ma. The current data suggest that significant terrane dispersion occurred in the period after ca. 590 Ma through the Lancinha-Cubatão fault system, which extends for about 2100 km, juxtaposing units of strongly contrasting ages and tectonic settings of formation and deformation. This dispersion tectonics controlled the final tectonic configuration of the southern Ribeira Belt. The collisional suture zone was dismembered during this continentalscale strike-slip system and the Lancinha-Cubatão Fault Zone, which separates the Curitiba Terrane of the Apiaí and Embu Terranes, probably does not represent exactly the suture zone as suggested in previous works (e.g., Basei et al., 2008), but it is mostly likely a transcurrent reactivation of the suture zone during the late oblique collisional period. During the Late Ediacaran the southern Ribeira Belt undewent an evolution similar to that first described for the North America Cordillera (Coney et al., 1980). In this context, we outline the following model for the evolution of the southern Ribeira Belt, with five main stages: 1) Magmatic continental arc stage (630–600 Ma): During this period the Apiaí and Embu Terranes were the basements to arc-related granitic magmatism that resulted from a northwestward subduction. This suggests that these terranes were already juxtaposed and accreted to the active margin bordering the Paranapanema Craton. 2) Collisional stage (600–590 Ma): The Turvo-Cajati Formation (a Neoproterozoic shallow continental-shelf metasedimentary assemblage) and Atuba Complex (an Archaean to Paleoproterozoic TTG-type orthogneiss assemblage) were accreted to the Apiaí-Embu Terrane as distinct exotic terranes. A short-lived event from crustal thickening to exhumation in the southern Ribeira Belt was related with the collision of a microcontinent, in which the Luís Alves Terrane probably is a remnant. The Turvo-Cajati Formation and the Atuba Complex displays con-

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Fig. 19. Block diagrams for tectonic models of evolution of the Curitiba Terrane (a) and (b) and tectonic assemblage diagram illustrating the history of terrane juxtaposition in the southern Ribeira Belt (c).

trasting metamorphic records of crustal thickening related with this collision, reflecting deformation in different parts of the collisional suture. Different slivers of the higher-grade unit of the Turvo-Cajati Formation were buried to depths ranging from 38 to >47 km, producing low-temperature eclogite facies rocks that evolved mainly to high-pressure granulites following clockwise paths. Rocks from the Atuba Complex were buried to maximum depths of 26 km. 3) Early exhumation stage (after 590 Ma): An early stage of rapid exhumation was recorded in the metamorphic paths from rocks of the higher-grade unit from the Turvo-Cajati Formation. Sillimanite migmatites record isothermal decompression from ca. 12.5 to 9.5 kbar, whereas scarce evidence for decompression after the peak metamorphism was recorded in kyanite migmatites, suggesting also a very rapid exhumation. Rocks from the Atuba Complex record a path of simultaneous cooling and decompression. We infer that the rapid exhumation was associated with the installation of the crustal-scale strike-slip regime. 4) Strike-slip tectonics, postcollision nappe stacking and folding (590–580 Ma): After the early period of exhumation the rocks from the collisional suture zone were dismembered and dispersed along the crustal-scale strike-slip system (lateral extrusion) that caused a generalised greenschist facies metamorphic overprint in the rocks from the Turvo-Cajati Formation and Atuba Complex. In the Curitiba domain the deformation associated with opposite movements of the transcurrent shear zones in its limits (dextral in the Lancinha-Cubatão Fault Zone to the north and sinistral in the Serra do Azeite Shear Zone to the south) was accommodated by westward thrusting with the transcurrent shear zones having acted as lateral ramps (Fig. 19a). This was followed by a northward verging folding phase with the squeezing of the Curitiba Terrane as a consequence of a NS compression between the Luís Alves and Apiaí Terranes (Fig. 19b). Late sinistral strike-slip faults would have dismembered and dispersed the folded nappe stack structure. 5) Extensional environment (ca. 580 Ma): The A-type granitic rocks from the Graciosa Province are inferred as related to the extensional braches of the strike-slip system and mark the final juxtaposition between de Curitiba and Luís Alves Terranes. The dispersion along the Lancinha-Cubatão Fault Zone continued up to ca. 530 Ma.

A tectonic assemblage diagram illustrating the history of terrane juxtaposition in the southern Ribeira Belt is shown in Fig. 19c. Acknowledgements Financial support was provided by FAPESP (Fundac¸ão de Amparo à Pesquisa do Estado de São Paulo – Brazil) grants 02/13654-4 to FMF and 01/13457-1 and 06/01327-0 to GACC. Field work collaboration by the geologists Rodrigo Meira Faleiros, Rodrigo Prudente de Melo and Diego Antonio Rodrigues Tamborim was of great importance. Thanks are also given to Marcos Mansueto (Electron Microprobe Laboratory, Geosciences Institute, University of São Paulo) for his assistance with the microprobe analysis. The paper was substantially improved following detailed comments and suggestions by Jennifer Chambers and an anonymous reviewer and editorial comments of Randall Parrish. Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at doi:10.1016/j.precamres.2011.07.013. References Almeida, F.F.M.de, Amaral, G., Cordani, U.G., Kawashita, K., 1973. The Precambrian evolution of the South American Cratonic margin south of Amazon River. In: Nairn, E.M., Stehli, F.G. (Eds.), The Ocean Basins and Margins, vol. 1. Plenum Press, New York, pp. 411–446. Basei, M.A.S., Frimmel, H.E., Nutman, A.P., Preciozzi, F., 2008. West Gondwana Amalgamation Based on Detrital Zircon Ages from Neoproterozoic Ribeira and Dom Feliciano Belts of South America and Comparison with Coeval Sequences from SW Africa. Special Publications 294. Geological Society, London, pp. 239–256. Basei, M.A.S., McReath, I., Siga Junior, O., 1998. The Santa Catarina granulite complex of Southern Brazil. Gondwana Research 1, 383–391. Beaumont, C., Jamieson, R.A., Nguyen, M.H., Lee, B., 2001. Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to focused surface denudation. Nature 414, 738–742. Bento dos Santos, T.M., Munhá, J.M., Tassinari, C.C.G., Fonseca, P.E., Dias Neto, C., 2010. Thermochronology of central Ribeira Fold Belt, SE Brazil: petrological and geochronological evidence for long-term high temperature maintenance during Western Gondwana amalgamation. Precambrian Research 180, 285–298. Bento dos Santos, T.M., Munhá, J.M., Tassinari, C.C.G., Fonseca, P.E., Dias Neto, C., 2011. Metamorphic P-T evolution of granulites in the central Ribeira Fold Belt, SE Brazil. Geosciences Journal 15, 27–51. Berman, R.G., 1990. Mixing properties of Ca-Mg-Fe-Mn garnets. American Mineralogist 75, 328–344.

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