Late amalgamation in the central part of West Gondwana: new geochronological data and the characterization of a Cambrian collisional orogeny in the Ribeira Belt (SE Brazil)

Precambrian Research 133 (2004) 29–61

Late amalgamation in the central part of West Gondwana: new geochronological data and the characterization of a Cambrian collisional orogeny in the Ribeira Belt (SE Brazil) Renata da Silva Schmitt a,∗ , Rudolph A.J. Trouw b , William Randall Van Schmus c , Márcio Martins Pimentel d a b

DGRG, Faculdade de Geologia, Universidade do Estado do Rio de Janeiro—UERJ, Rio de Janeiro, RJ, Brazil Departamento de Geologia, IGEO, Universidade Federal do Rio de Janeiro—UFRJ, Rio de Janeiro, RJ, Brazil c Departament of Geology, University of Kansas, 120 Lindley Hall, Lawrence, KS 66045, USA d Instituto de Geociˆ encias, Universidade de Bras´ılia—UnB, Bras´ılia, DF, Brazil Received 19 August 2003; accepted 29 March 2004

Abstract New U–Pb data reveal that during the mid-Cambrian the central part of West Gondwana was still undergoing a high-grade tectonometamorphic event corresponding to collision. The studied area is located in the southeastern part of the Pan-African– Brasiliano Ribeira Belt, in SE Brazil. The area is part of the Cabo Frio Tectonic Domain (CFTD) which is limited to the NW by a major NE–SW striking thrust zone which separates it from the Neoproterozoic “Oriental terrane”, whereas to the SE it is covered by the Atlantic Ocean. The domain comprises a Paleoproterozoic orthogneissic basement tectonically interleaved with younger supracrustal rocks, folded and metamorphosed at upper amphibolite to granulite facies during the mid-Cambrian. The supracrustal rocks are subdivided in two successions: Búzios (Al-metapelites, calcsilicates and amphibolites) and Palmital (quartz–feldspathic metasediments with minor metapelites). These successions were deposited in a deep oceanic environment between ca. 620 and 525 Ma as indicated by SHRIMP U–Pb data for detrital zircons and by TDM model ages. The metamorphic peak, defined by the mineral associations Ky + Kfs in metapelites and Cpx + Grt + Qz in amphibolites, occurred at minimum pressure of 9 kbar and temperature in excess of 780 ◦ C. At this stage migmatites were generated by partial melting in all lithostratigraphic units, including the amphibolites. The metamorphic peak was also contemporaneous with top to the NW thrusting, testified by mineral and stretching lineations related to progressive deformation phases D1 and D2. The metamorphic peak was dated between 525 and 520 Ma, as determined by U–Pb analyses of zircons of leucosomes. During deformation phase D3, large recumbent folds developed with NW–SE axes, parallel to the main direction of movement. The CFTD was juxtaposed at this stage to the “Oriental terrane” by a major NE–SW striking thrust fault. U–Pb dating of monazites from metapelites and of sphenes from amphibolites revealed ages of about 510 Ma for the mineral growth. The sillimanite, aligned as L3 , partially replaced kyanite, indicating a clockwise P–T–t path for the central and eastern areas of the CFTD. After docking into the Ribeira belt, during the late Cambrian, the western limit of CFTD was affected by a transcurrent dextral shear zone that developed a NE–SW stretching lineation related to D4, under amphibolite facies conditions. This is recorded in monazites and zircons within this shear zone with U–Pb ages ranging from 505 to 490 Ma. At this stage, the central and eastern parts of CFTD were already cooling at a rate of 10 ◦ C/Ma. After 480 Ma, the cooling rate diminished to 5 ◦ C/Ma. A 207 Pb/206 Pb

∗ Corresponding author. Tel.: +55-21-9638-8859/2587-7102; fax: +55-21-2587-7704. E-mail addresses: [email protected], [email protected] (R. S. Schmitt).

0301-9268/$ – see front matter © 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.precamres.2004.03.010

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age in rutile (480 ± 5 Ma) and a U–Pb zircon age in a post-tectonic pegmatite (440 ± 11 Ma) mark the stabilization of the area during the Ordovician–Silurian boundary. This well-constrained Cambrian collisional orogeny is here denominated the Búzios Orogeny and is the youngest tectonometamorphic event recorded in the Brasiliano belts of Brazil. This orogeny is contemporaneous with some other marginal orogenies (e.g. Pampeana; Ross) that probably accompanied the final adjustments of the Precambrian continents to complete the formation of Gondwana in the Ordovician. © 2004 Elsevier B.V. All rights reserved. Keywords: High-grade metamorphism; U–Pb; Orogeny; Cambrian; Brazil; Gondwana

1. Introduction The formation of Gondwana in the Neoproterozoic– Early Paleozoic resulted from a series of thermotectonic events related to the subduction of oceanic lithosphere, accretion of exotic terranes, collision of continental blocks, mantle plume activities, and other events registered in the Pan-African–Brasiliano belts (e.g. Almeida et al., 2000; Cordani et al., 2000; Unrug, 1997; Hoffman, 1991; Doblas et al., 2002). Various of these belts result from the amalgamation of tectonic domains with diachronous events generated by several orogenies. This paper presents new structural and geochronological data from the eastern domain of the Ribeira Belt (southeast Brazil) revealing the record of a Cambrian tectonometamorphic event (Schmitt, 2001). The characterization of this young tectonic event has implications with respect to the history of Gondwana, mainly because of its location. It is generally considered (e.g. Unrug, 1997; Condie, 1997) that, by this time, the central part of West Gondwana was already stabilized and that orogenic activity only continued along the margins (e.g. Pampeana and Ross orogens). The first stages of sedimentary deposition in the Paleozoic–Mesozoic intracratonic South American basins (e.g. Paraná basin) occurred during the Ordovician (Milani and Zálan, 1999). Recent geochronological studies both in the Pan-African and Brasiliano belts (e.g. Jung et al., 2000; Seth et al., 1998; Schmitt et al., 1999; this paper) indicate that the assembly of Gondwana was only completed by the end of the Cambrian and Early Ordovician.

2. Regional setting The Ribeira belt represents a complex orogenic domain which extends more than 1400 km along the

South Atlantic margin and is approximately 300 km wide in its central segment (Heilbron et al., 2000; Trouw et al., 2000a). It forms one of the margins of the Precambrian São Francisco Craton (SFC) in Brazil (Fig. 1) (Cordani et al., 2000). Its continuation to the north is known as the Araçua´ı belt (Pedrosa-Soares et al., 1998; Pedrosa-Soares and Wiedmann-Leonardos, 2000) with similar structural trend, sedimentary successions, and timing of collisional events. To the west, the Ribeira belt deformation partially overprints the slightly older Bras´ılia belt, where the formation of orthogneisses and metavolcanic juvenile arcs between 900 and 600 Ma, resulted from convergence and oceanic lithosphere consumption between various cratons (Pimentel and Fuck, 1992). The main collisional event of the Bras´ılia belt occurred around 625 Ma (Campos Neto, 2000; Pimentel et al., 1999), whereas in the Ribeira belt collisions took place between 590 and 520 Ma (Machado et al., 1996; Heilbron et al., 2000; this paper). This diachronous evolution is registered by an interference zone between these belts, to the south of the SFC (Trouw et al., 1994), where low- to medium-pressure metamorphic mineral assemblages, related to the Ribeira belt, overprint higher pressure metamorphic mineral assemblages of the Bras´ılia belt (Ribeiro et al., 1995; Trouw and Oliveira Castro, 1996; Campos Neto and Caby, 1999). Therefore, the Ribeira Belt is among the younger products of the Pan-African–Brasiliano collage with collisional events taking place from the late Neoproterozoic through the Paleozoic, ending with Ordovician–Silurian late-tectonic granites (Heilbron et al., 2000). The Ribeira belt in southeastern Brazil is divided into four main tectonic domains (Heilbron et al., 2000; Trouw et al., 2000b) (Fig. 2):

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Fig. 1. Gondwana paleocontinent with its cratonic blocks and Pan-African–Brasiliano mobile belts. Cratons: RDP, Rio de La Plata; AMZ, Amazonas; ARQ, Arequipa; WA, West African; CH, Chad; SF, São Francisco; CG, Congo; KAL, Kalahari; EAN, East Antarctica; IND, Indian; WAS, West Australian; NAS, North Australian: GAW, Gawler. Mobile belts: Moç, Moçambique; Zb, Zambezi; Lf, Lufilian; ROS, Ross; Kan, Kanmatoo; CF, Cape Fold; Sal, Saldania: Gar, Gariep; Dm, Damara; Kk, Kaoko; SP, Sierra Pampeanas; SA, Sierra Australes (modified from Powell, 1993).

(A) Adjacent to the SFC, in the NW of the area considered here, is the “Occidental terrane”, constituted of Meso- to Neoproterozoic metasedimentary rocks of a marine passive margin. According to Heilbron et al. (1995, 1998), this terrane is internally organized into two crustal scale thrust sheets, Andrelˆandia (to the west) and Juiz de Fora (to the southeast), that override the SFC foreland. The Andrelˆandia thrust sheet (partially equivalent to the Liberdade nappe system; Fig. 2) is alternatively interpreted by Trouw et al. (2000b) as belonging to the Bras´ılia belt because of its stretching lineations and shear-sense indicators that show movement with top to the NE. (B) Overlying the Juiz de Fora thrust sheet, to the southeast, is the Para´ıba do Sul klippe, a terrane constituted of post-1.8 Ga granulitic metasediments and their Paleoproterozoic basement, thrust over the “Occidental terrane”. A late NE–SW axis megasynform structure, with its core in the center of the Para´ıba do Sul klippe, folds the Juiz de Fora domain that crops out also to the SE, with the foliation dipping NW (Fig. 2).

(C) A major mylonitic shear zone, more than 200 km long and dipping 35◦ NW, named the Central Tectonic Boundary, separates the “Oriental terrane” from the “Occidental terrane” (Almeida et al., 1998) (Fig. 2). The “Oriental terrane” contains two main lithotectonic units: orthogneisses, interpreted as a magmatic arc (Rio Negro arc; Tupinambá et al., 1998), and metasedimentary rocks intruded by synto late-collisional granitoids (named Costeiro Domain). (D) To the SE, the Cabo Frio Tectonic Domain overrides the “Oriental terrane” through a SE dipping thrust fault, denominated here the Cabo Frio Tectonic Boundary (Fig. 3).

3. Geology of the Cabo Frio Tectonic Domain The Cabo Frio Tectonic Domain is composed of Neoproterozoic–Cambrian supracrustal rocks tectonically interleaved at low angle with a reworked Paleoproterozoic basement (Fig. 3). The CFTD does not exhibit Brasiliano plutons.

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Fig. 2. Tectonic setting of São Francisco Craton, Bras´ılia and Ribeira belts with their tectonic domains, including the Cabo Frio Tectonic Domain. CTB means Central Tectonic Boundary (modified from Heilbron et al., 2000; Trouw et al., 2000b). Marked box shows the area of this study (Fig. 3).

The orthogneissic basement is subdivided into a predominant felsic unit (Região dos Lagos Unit) and a subordinate mafic unit (Forte de São Mateus Unit). The Região dos Lagos Unit comprises mainly metagranitoids with subordinate metaquartz-diorite and metatonalite bodies. In less deformed domains, the igneous protoliths can easily be identified, whereas in strongly deformed domains the metagranitoids become banded gneisses with migmatitic structures (Fig. 4a and b). The interleaved metaquartz-diorites and metatonalites show medium to coarse-grained texture with amphibole and biotite as varietal and accessory minerals. The metagranitoid group shows pre-

dominant monzogranitic composition, with subordinate syenogranite, quartz-monzonite and granodiorite varieties. They present two main textures: porphyritic and equigranular. The porphyritic metamonzogranites have amphibole and biotite, with microcline phenocrysts up to 7 cm long. The equigranular metagranitoids have a medium to coarse-grained texture. All these lithotypes are cross-cut by quartz–feldspar hololeucocratic aplitic veins, with localized pegmatitic texture. These veins are granodioritic in composition within the metatonalites and metaquartzdiorites, and monzogranitic to alkali-feldspar granitic in the metagranitoids.

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Fig. 3. Geological map of the Cabo Frio Tectonic Domain (CFTD) and the eastern part of the “Oriental terrane”. The inserts present geological cross sections A–A ; B–B and C–C .

The Forte de São Mateus Unit is constituted by banded amphibolitic gneisses up to 50 m thick. It is interleaved tectonically with the Região dos Lagos Unit, and contains amphibolitic dikes that cross-cut the felsic orthogneisses. The banded am-

phibolitic gneisses include two lithotypes: (1) massive garnet–orthoamphibolite; (2) amphibole–garnet– diopside banded gneiss. The massive amphibolite is medium to coarse-grained with andesine and garnet, plus sphene, diopside, and zircon as accessories. The

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Fig. 4. Photos and microphotos from the Cabo Frio Tectonic Domain lithostratigraphic units. (a) Metagranitoids from Região dos Lagos Unit with penetrative quartz–feldspar stretching lineation parallel to the hammer; (b) boudins of orthoamphibolites (O) from the Forte de São Mateus Unit within the Região dos Lagos orthogneisses. Note the felsic pegmatitic material in the necks; (c) metasedimentary rocks from the B´uzios succession. A garnet gneiss more competent layer (in grey) show the pattern of the D3 folds in a hinge area. Note that some leucosome veins are also folded and some are parallel to the axial plane. (d) Paragneiss from Palmital succession (in Ponta Negra region). Note the leucosome concentration at the bottom. The tight fold is probably related to D3 and the open fold (to the right), to D4. (e) Kyanite crystal (K) and perthite (P) in equilibrium in a metapelite from B´uzios succession (width of view 2 mm: plane polarized light); (f) Pseudomorphic structure shown by a kyanite crystal (K) partially replaced by sillimanite (S) in the metapelites from the B´uzios succession (width of view 2 mm; plane polarized light).

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Fig. 4. (Continued )

amphibole–garnet–diopside banded gneiss is a laminated rock with calcsilicate (diopside, andesine and almandine) and amphibolitic layers of 0.5–2 cm thick. The amphibolitic dikes, with thickness ranging from

5 to 2 m, were originally tabular bodies, but due to deformation they usually appear as boudins in limbs of isoclinal folds (Fig. 4b). The dikes are medium to coarse-grained and are composed of amphibole

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Fig. 4. (Continued ).

and andesine-labradorite plagioclase, with garnet, diopside, sphene, and quartz as accessories. The Supracrustal Unit is at least 500 m thick and is divided into three compositional groups

(quartz–feldspathic, aluminous with kyanite/sillimanite and calcsilcate rocks) that correspond to the original sedimentary facies, interleaved with amphibolitic bodies. The contact with the basement units is always

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strongly tectonized. The compositional layering of these gneisses is interpreted to reflect the original bedding; therefore this unit was subdivided into two main sedimentary successions that grade laterally one into the other based on the frequency of the interbedded lithotypes. The Búzios succession is a thick, aluminous metasediment package (sillimanite–kyanite–garnet–biotite gneiss) (Fig. 4c) with numerous calcsilicate and amphibolitic intercalation. The sillimanite–kyanite– garnet–biotite gneiss also contains orthoclase, minor microcline, oligoclase-andesine, and quartz. Trace minerals are sphene, monazite, rutile, zircon, tourmaline, and apatite. The kyanite varies from microscopic scale to 7 cm long crystals (Figs. 4e and 4f). The sillimanite occurs as prismatic crystals and also as fibrolite partially substituting for kyanite (Fig. 4f). The calcsilicate layers show two gneiss types: one composed of diopside (40–70%), garnet and minor biotite; and a second composed of diopside (15–25%), amphibole, quartz, biotite, plagioclase, orthoclase, and garnet. The amphibolite layers are subdivided into garnet amphibolites, diopside amphibolites, and metahornblendites. The amphibolites occur in layers of a few centimeters to meters or grouped in packages up to 15 m thick interbedded with the aluminous metasediments and the calcsilicate rocks. Actinolite, scapolite, sphene, epidote, rutile, chlorite, and zircon are trace minerals. The Palmital succession is constituted mainly of quartz–feldspathic metasedimentary (paragneiss) (Fig. 4d) thick packages (>300 m), with some aluminous intercalation, calcsilicate rocks and feldspathic quartzite layers. The paragneiss shows a regular compositional layering varying from 5 to 30 cm, determined by felsic layers (rich in quartz and feldspar) and more biotite-rich layers with kyanite and sillimanite, which are interpreted as representative of the bedding of a probably turbiditic succession (Fig. 4d). The main mineralogical constituents are orthoclase, quartz, biotite, plagioclase, and sillimanite. Accessories are garnet, apatite, sphene, zircon, and kyanite (only in the metamorphic KY zone). Within the calcsilicate lenses, there is a 1 m thick layer of sphene-bearing amphibolite interpreted to be of sedimentary origin. It is constituted mainly of plagioclase (bytownite–anortite) and amphibole, with garnet, sphene, and ilmenite as accessories.

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These two successions, Búzios and Palmital, are interpreted as deposited as lateral varieties in the same basin (Búzios–Palmital basin). Minor layers of garnet–quartz gneiss and feldspathic quartzite within both successions could represent interdigitations. Based mainly on facies descriptions, the lithotypes are thought to be derived from turbiditic hemipelagic to pelagic sediments of a submarine fan. The quartzo-feldspathic dominated Palmital succession may represent deposits of the medium portion of the fan, and the pelitic Búzios succession may correspond to the pelagic distal facies. The calcsilicate layers might represent chemical deposits possibly related to volcanic events. The amphibolites, interpreted as mafic subvolcanic intrusions and/or lava flows, are restricted to the pelagic environment of the Búzios succession. 3.1. Structural geology All lithostratigraphic units of CFTD were affected by a low angle top to the NW thrusting (D1–D3), which generated a NW–SE stretching/mineral lineation with low plunge to both directions, a subhorizontal tectonic foliation, subhorizontal shear zones, and recumbent folds of variable size. The penetrative character of this NW–SE lineation, associated with several kinematic indicators, is taken as evidence that the main direction of thrust movement was top to NW (Fig. 4a). During D1 and D2 phases, ductile thrust sheets of orthogneisses were superimposed over the supracrustal rocks in a ductile flow regime. The detachment surfaces and shear zones are located preferentially along the compositional contacts (basement/ supracrustal rocks) and preexistent structures, such as sedimentary bedding. As a result, the basement rocks are strongly deformed in 50–150 m thick zones along the contact with the supracrustal rocks (Fig. 4a). Within the basement sheets, isolated mylonitic shear zones with some L-tectonites are preserved. The intrafolial folds in the D1/D2 foliation in the basement and supracrustal rocks could be related to the passive folding and be responsible for the thrust transport as described by Hatcher and Hooper (1992) for F-type thrust sheets developing under metamorphic conditions of medium to high-grade, in the lower crust.

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The supracrustal rocks also develop a foliation (S1/2 ) subparallel to the original compositional layering (S0 ). Although the stretching lineation is mainly defined by quartz and feldspar in the metasedimentary rocks, it is also reinforced by mineral alignment of kyanite, large sillimanite crystals, and amphibole. During the progressive deformation D3 phase, the foliation S1/2 , the thrust zones, and the boudinage structures were folded into meso to macro-scale recumbent folds with axes parallel to the main mineral lineation (L1/2 ) (Fig. 4c). The vergence of these D3 folds varies from ENE, in the eastern portion of the area, to WSW in the western portion. This opposite vergence probably means that the main transport was still top to NW (parallel to the D3 fold axes), possibly related to the geometry of a major basement sheet in a sheath or oblique fold model (Passchier and Trouw, 1996). There is also an E–W mineral lineation (L3 ) formed by fibrolitic sillimanite, plunging 5–25◦ , almost orthogonal to L1/2 , and restricted to the F3 axial plane foliation. Kinematic indicators such as S/C foliations, rotated porphyroblasts, and mica fish indicate transport of both top to west and top to east in alternating fold limbs. The thrust contact between the CFTD and the “Oriental terrane” to the west is interpreted to be developed during D3 (Fig. 3, geological section C–C ). In this shear zone, kinematic indicators, such as S/C foliations, disrupted fold limbs and C surfaces show transport of top to NW. These structures were developed at lower temperatures than the D1/D2 thrust zones, in a transition between F and C thrust types (Hatcher and Hooper, 1992). In this case, the crystalline thrust sheet probably had its roots below the ductile–brittle transition and was transported through the transition during the progressive movement. The presence of sillimanite oriented down dip in the plane of the shear zone indicates that the metamorphic grade was still high (amphibolite facies). Shear-sense indicators are frequent and well-preserved in this D3 thrust, whereas the D1/D2 shear zones show very few kinematic indicators, probably because they developed during the metamorphic peak (granulite facies). After the “docking” of the CFTD, a transpressional, dextral, subvertical NE–SW shear zone developed along the southwestern contact with the “Oriental terrane”. Mineral and stretching lineations plunging 10–30◦ to SW developed parallel to tight to open

F4 fold axes, with axial planes dipping steeply to NW and SE. The main S1/2 foliation was folded and reoriented to attitudes with dips 50–90◦ to SE and NW (Fig. 3). In strongly deformed domains there is a transposition foliation parallel to the axial plane of the F4 folds. Sillimanite and amphibole were realigned and quartz and feldspar were stretched (L4 ) in D4 shear zones. In some folds the L3 sillimanite lineation can still be recognized in addition to new growth of sillimanite along L4 (Rocha, 2002). D4 folded the original contact between CFTD and the “Oriental terrane” in the Ponta Negra region, to a position with 75◦ dip to NW (Fig. 3). D4 is restricted to the western part of the area, but the remaining part of the CFTD presents gentle to open folds with NE–SW axes that could be also related to this phase. 3.2. Metamorphism The CFTD lithostratigraphic units show mineral assemblages indicative of a metamorphic peak of at least 9 kbar and 780 ◦ C, in the transitional field between amphibolite to granulite facies in the eastern area, and in the upper amphibolite facies field in the western area (Fig. 5). The metamorphic peak was contemporaneous with D1 and D2 deformation phases because the leucosomes in all units are parallel to the S1/2 foliation and present L1/2 stretching lineation. During phase D3, the metamorphic conditions were at lower P and T on a clockwise return path, mainly related to decompression, as shown by the growth of sillimanite parallel to L3 and folding of the leucosomes by F3 (Fig. 4c). In the western part of CFTD, the growth of sillimanite and stretching of quartz–feldspar indicate that during D4 transpressional dextral shear, metamorphic conditions were still in the amphibolite facies. In the eastern part of the CFTD, aluminous metasedimentary rocks and aluminous layers of the paragneiss have a mineral assemblage of kyanite + sillimanite + garnet + biotite + orthoclase + plagioclase (oligoclase/andesine) + rutile, with quartz feldspathic leucosome forming a typical stromatic migmatite structure. The leucosomatic material, mainly composed of quartz, K-feldspar, and garnet, is usually surrounded by restite of garnet–sillimanite–biotite. The percentage of partial melting is estimated to be between 10 and 30%. The absence of muscovite, the

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Fig. 5. P–T–t paths for metapelites of the CFTD in P–T diagram for KASH in metapelites (Spear et al., 1999), with metamorphic facies fields (Bucher and Frey, 1994). The trajectories indicate minimum P–T conditions. Each path is correlated with the deformation phases (D1–D4). Biot, biotite; Ms, muscovite; Kfs, K-feldspar; Als, aluminosilicate; Qz, quartz; L, liquid phase; V, vapor phase; Ky, kyanite; Sil, sillimanite; And, andaluzite.

coexistence in equilibrium of kyanite and orthoclase (Fig. 4e), and the partial melting structures define the minimum P–T conditions. The secondary growth of sillimanite over relict crystals of kyanite (Fig. 4f) in pseudomorphic structures is interpreted as related to a clockwise metamorphic path due to a diminishing pressure; this is also indicated by sphene growing over rutile and ilmenite. To the west of a N–S oriented “kyanite isograd” (Fig. 3), the metasedimentary rocks present sillimanite in equilibrium with K-feldspar, absence of muscovite, and presence of leucosomes. The garnet has sillimanite inclusions in this area, while in the “kyanite area” the garnet contains kyanite as inclusions. The presence of retrograde muscovite partially substituting for sillimanite, K-feldspar and plagioclase indicates that the return P–T–t path occurred above the invariant point (Fig. 5). The calcsilicate rocks present mineral parageneses comprising diopside + quartz + garnet + plagioclase (labradorite–bytownite) + scapolite in the garnet metaclinopyroxenite lithotype and quartz + diopside + biotite + plagioclase + garnet + hornblende in the diopside gneiss. The amphibolites interbedded with the supracrustal rocks present the following parageneses: garnet + plagioclase

+ hornblende + quartz (garnet amphibolite), diopside + garnet + plagioclase + hornblende + quartz (diopside amphibolite) and hornblende + diopside + plagioclase (metahornblendite). The orthoamphibolite dikes and the garnet amphibolites (banded gneiss) of the basement have: hornblende + plagioclase (andesine/labradorite) + quartz ± diopside ± garnet ± biotite. The garnet amphibolites show also trondhjemitic veins (andesine + quartz) surrounded by concentrations of pure hornblende, interpreted as “in situ” partial melting of the amphibolite. The mineral paragenesis garnet + diopside + plagioclase + amphibole + quartz in metamafic rock is stable in upper amphibolite facies and also in high-pressure granulite facies, probably indicating the transition zone (Yardley, 1989) coherent with the generation of partial melting (Bucher and Frey, 1994). The absence of orthopyroxene does not indicate that the granulite facies was not reached, because orthopyroxene is usually absent in high-pressure metamorphic rocks. These high-pressure granulites are commonly referred to as garnet granulites (Bucher and Frey, 1994). Therefore, the orthoamphibolites present a metamorphic assemblage consistent with the metamorphic conditions of the supracrustal rocks.

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The basement felsic orthogneiss (Região dos Lagos Unit) does not show a diagnostic paragenesis for these conditions, although there are leucosomes in some domains, usually at the contacts with the supracrustal rocks. The main mineral assemblage is K-feldspar + plagioclase (oligoclase/andesine) + quartz + hornblende + biotite. The strongest evidence that the felsic basement was also affected by the same metamorphic conditions as the other units is the deformation structures such as stretching and mineral lineations, as well as the folding of leucosomes.

4. Geochronology 4.1. Laboratory procedures The minerals were analyzed at the Isotope Geochemistry Laboratory (University of Kansas, USA), using a VG-Sector multicollector thermal ionization mass spectrometer, following standard U–Pb techniques (Krogh, 1973, 1982; Parrish, 1987). The detailed description of the technique is in Appendix A. The data are presented in Table 1. Sm–Nd data were obtained for seven whole-rock samples of the supracrustal unit. The detailed description of the technique is in Appendix B. The data are presented in Table 2. Ion microprobe analyses were carried out using SHRIMP RG at the Research School of Earth Sciences, Australian National University, Canberra, Australia. Zircon grains were mounted in epoxy resin and polished. Transmitted and reflected light microscopy, as well as scanning electron microscope cathodoluminescence imagery was used to investigate the internal structures of the zircon crystals prior to spot analysis. Data were collected and reduced as described by Williams and Claesson (1987) and Compston et al. (1992). Uncertainties reported in Table 3 are given at 1σ level, and final age is quoted at 95% confidence level. Reduction of raw data was carried out using Squid 1.02 (Ludwig, 2001a), and final ages and concordia plots were calculated using Isoplot-Ex (Ludwig, 2001b). U–Pb ratios were referenced to the RSES standard zircon FC1 (1099 Ma, ∗206 Pb/238 U = 0.1859). U and Th concentrations were determined relative to those measured in the RSES standard SL13.

4.2. Basement data 4.2.1. Felsic Orthogneiss Região dos Lagos Unit Zircon morphology is extremely varied in this unit. One rock sample may contain four to six distinct populations varying from needle-shaped crystals, through oval grains, to round, equant crystals. The zircon crystals are pink, transparent, and have very few inclusions. Although heterogeneous, these populations all gave similar U–Pb data, proving that they grew during the same geological event. The zircons are discordant, but in sample from less deformed domains (BUZ-48) the fractions plot close to the upper intercept. The sample from deformed domains displays stronger Pb loss and the lower intercept ages have smaller uncertainties. The upper intercepts indicate Paleoproterozoic ages, interpreted as crystallization ages for the protoliths (Fig. 6a and b): 1960 ± 6 Ma and 1971±5 Ma. The lower intercepts show Paleozoic age, 465±52 Ma and 525±37 Ma, respectively. However, this is not the best evidence to prove that the orthogneisses were deformed and metamorphosed during the Cambrian. One sample from a leucocratic, deformed vein (BUZ-62-4) within this unit presents two distinct zircon populations: (POP.1), acicular pink transparent fractured zircons (similar to some populations described in the orthogneiss); (POP.2), anhedral, transparent, colorless to slightly pink crystals, with less fractures (not observed in the orthogneiss samples). Both populations were abraded and analyzed, giving different ages. Population 1 yielded an upper intercept of 1977 ± 9 Ma, and a lower intercept of 519 ± 11 Ma (Fig. 6c); both intercepts are coherent with the orthogneiss data. Population 2 presented an upper intercept of 518 ± 5 Ma (Fig. 6d). Therefore, these Cambrian zircons grew during the crystallization of those leucocratic veins, and the Paleoproterozoic zircons (POP.1) were probably inherited from the orthogneiss during the intrusion of the vein (thickness of 2–60 cm). In addition, the leucocratic veins are folded by the F3 generation, so this deformation phase (D3) occurred after 518 Ma. This Cambrian age also coincides with the lower intercept age of the inherited zircon population. 4.2.2. Amphibolites—Forte de São Mateus Unit The Forte de São Mateus amphibolites produced small amounts of zircon. Crystals from sample

Table 1 U–Pb data Sample fractiona

Size (mg)

U (ppm)b

Observed Pb (ppm)b

Radiogenic ratiosc 206 Pb/204 Pb

207 Pb∗ /235 U

206 Pb∗ /238 U

207 Pb∗ /206 Pb∗

206 Pb∗ /238 U

207 Pb∗ /235 U

207 Pb∗ /206 Pb∗

0.317375 0.295957 0.29935 0.33131 0.310858 0.340575 0.305996

0.118622 0.116948 0.118133 0.119002 0.117673 0.119719 0.117588

1776.9 1671.2 1688.1 1844.7 1744.9 1889.4 1721

1851.1 1780 1798.1 1890.6 1826.7 1919.5 1812.7

1936+−19 1910+−3 1928+−4 1941+−3 1921+−2 1952+−5 1920+−2

0.117124 0.119161 0.118539 0.116516 0.119552 0.113103

1700.5 1822.9 1670.1 1644.9 1814.3 1417.2

1797.9 1880 1790.7 1761.8 1878.2 1600

1913+/−4 1944+/−1 1934+/−2 1903+/−1 1949+/−3 1850+/−2

476.1 508.39 491.21 513.74 1250.4 812.89

483.65 510.2 495.03 514.4 1456 989.07

520+/−10 518+/− 5 513 +/− 4 517 +/− 5 1770 +/− 2 1404 +/− 3

1819.3 1656.3 1643.7 801.72

1877.8 1768.2 1759.5 975.55

1943 1903 1900 1391

UI = 1960 +−6 Ma, LI = 465 +− 52 Ma (2σ); MSWD = 1.2; P = 0.3 (model 1; all fractions except∗ ) BUZ-62—metaquartz-monzonite (zircon) Loc: Rio das Ostras city—BUZ-62 (Fig. 3)—Lat: 22◦ 31 56,1 S Long: 41◦ 57 23,4 W nm(−1)[1]x 0.004 146.23 60.03 185.662 4.87477 M(−1)x[1] 0.006 255.23 90.02 1541.34 5.36944 M(−1)[1] 0.005 292.81 93.205 1837.24 4.83334 M(−1)x[3] 0.019 185.73 57.585 1852.33 4.66964 M(−1)[1] 0.007 139.4 50.327 545.174 5.35796 M(0)[1] 0.005 255.19 76.906 362.051 3.83444

0.301861 0.326807 0.295722 0.290668 0.325043 0.245883

UI = 1971+/− 5 Ma, LI = 525 +/− 37 Ma (2σ); MSWD =1.8; P = 0.18 (model 1; n(x) = 3) BUZ-62-4—Leucocratic vein pre-D3 (monzogranite) (zircon) Loc: Rio das Ostras city—BUZ-62 (Fig. 3)—Lat: 22◦ 31 56.1 S Long: 41◦ 57 23.4 W nm(−1)[1] 0.011 208.17 16.697 235.389 0.610149 0.076651 0.0577321 M(−1)[1] 0.01 251.63 19.461 556.877 0.652792 0.082057 0.0576978 M(−1)[1] 0.016 276.88 20.113 1197.92 0.628295 0.079177 0.0575526 M(0)[1] 0.009 178.4 14.162 557.702 0.659647 0.082956 0.0576715 M(1)x[1] 0.005 387.51 96.766 1519.76 3.1952 0.214064 0.108256 M(1)x[1] 0.006 479.37 66.572 510.549 1.64875 0.134394 0.0889762 UI = 518 +/− 5 Ma, LI = 67 +/− 110 Ma (2σ); MSWD =1.6; P = 0.2 (model 1; four bypiramidal crystals)

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

BUZ-48—metasienogranite (zircon) Loc: Maca´e city—BUZ-48 (Fig. 3)—Lat: 22◦ 23 56 S Long: 41◦ 50 4,4 W nm(−1)[2] 0.009 61.085 25.521 629.841 5.19085 nm(−1)[2] 0.008 29.053 11.128 905.727 4.77226 nm(−1)[1]∗ 0.004 132.3 49.899 669.851 4.87589 M(−1)[2] 0.006 19.894 8.2987 789.436 5.43613 M(−1)[1] 0.008 40.239 16.645 1112.85 5.04359 M(−1)[2] 0.014 33.609 17.168 205.916 5.62183 M(1)[1] 0.009 28.108 10.611 1055.98 4.96113

Calculated ages (Ma)d

UI = 1977 +/− 9 Ma; LI = 519 +/− 11 Ma (2σ); MSWD =0; P= 1 (model 1, two acicular inherited crystals, x) BUZ-44—garnet–amphibole–diopside gneiss (banded orthoamphibolite) (ZIRCON) Loc: Cabo Frio city—BUZ-44 (Fig. 3)—Lat: 22◦ 53 13 S Long: 42◦ 0 19.4 W nm(3)[1] 0.004 235.55 91.466 426.303 5.35558 M(3)[1] 0.004 156.57 52.216 842.159 4.70541 M(3)[3] 0.007 184.63 62.174 649.061 4.65675 M(3)[1] 0.005 120.18 20.874 173.867 1.61372

0.326076 0.292968 0.290439 0.132431

0.119121 0.116487 0.116286 0.0883766

+/− +/− +/− +/−

2 3 2 8 41

42

Table 1 (Continued ) Sample fractiona

Size (mg)

U (ppm)b

Observed Pb (ppm)b

Radiogenic ratiosc 206 Pb/204 Pb

207 Pb∗ /235 U

Calculated ages (Ma)d 206 Pb∗ /238 U

207 Pb∗ /206 Pb∗

206 Pb∗ /238 U

207 Pb∗ /235 U

207 Pb∗ /206 Pb∗

UI = 1969 +/− 4 Ma; LI = 519 +/− 19 Ma (2σ); MSWD = 1.14; P = 0.32 (model 1, all fractions)

0.078046 0.079196 0.082063 0.082808

0.0550678 0.0585503 0.058551 0.0577197

484.45 491.32 508.43 512.86

472.52 501.9 516.38 514.01

415 550 551 519

+/− +/− +/− +/−

57 20 11 9

0.079809 0.081567 0.080218 0.080887 0.079016 0.082119 0.08138 0.078354 0.082881 0.081518

0.0567781 0.0566256 0.0566023 0.056907 0.0575965 0.0577731 0.0577599 0.0581959 0.0578726 0.0576788

494.99 505.47 497.42 501.41 490.24 508.76 504.35 486.29 513.3 505.18

492.85 500.35 493.63 499 494.53 511.03 507.31 495.29 515.44 507.43

483 477 476 488 514 521 521 537 525 517

+/− +/− +/− +/− +/− +/− +/− +/− +/− +/−

6 2 3 5 3 4 6 9 5 3

0.082514 0.08216 0.0833 0.082185 0.081343

0.0575015 0.0575833 0.0578163 0.0576827 0.0581114

511.11 509 515.79 509.15 504.14

511.06 509.9 517.08 510.72 509.56

511 514 523 518 534

+/− +/− +/− +/− +/−

4 3 3 2 11

BUZ-46—kyanite–sillimanite–garnet–biotite Gneiss (aluminous metassediment) (monazite) Loc: Tartaruga beach, B´uzios city—BUZ-46 (Fig. 3)—Lat: 22◦ 45 0 S Long: 41◦ 53 53 W 0.5 A 0.0044 2687 705.24 1151.75 0.655773 0.08263 0.55 A 0.006 3391.7 882.97 3379.64 0.651234 0.081958 0.55 A 0.002 4276.4 1061.8 2055.24 0.654919 0.082509 0.6 A 0.0044 2930.2 843.73 498.909 0.644452 0.081061 0.6 A 0.004 2021 451.97 1590.17 0.651943 0.082102

0.0575595 0.0576297 0.0575688 0.0576606 0.0575912

511.8 507.8 511.08 502.45 508.66

512.03 509.24 511.5 505.06 509.67

513 516 513 517 514

+/− +/− +/− +/− +/−

3 2 3 6 3

UI = 551 +/− 10 Ma forced through zero (2σ); MSWD = 0; P = 1 (model 1, n = 2∗ ) BUZ-8-8—leucosome sin-D1/D2 (monazite–Mn and zircon) Loc: Foca beach, B´uzios city—BUZ-08 (Fig. 3)—Lat: 22◦ 45 54 S Long: MN-Allen 0.0084 3072.4 1220.1 1205.83 MN-0.65A 0.006 1.35% 1581 7355.18 MN-0.70A 0.0025 1.04% 1451.6 3210.86 MN-0.70 A 0.0049 8490.7 2064.3 5705.02 nm(0)[1]∗ 0.012 1187.6 87.404 2372.17 0.005 1246.4 95.144 1135.77 nm(0)[1]∗ nm(0)[1]∗ 0.007 1150.7 87.212 1354.55 M(0)[1] 0.008 953.62 69.007 740.187 M(0)[1] 0.006 657.29 62.929 243.183 M(0)[1]∗ 0.005 1100.7 85.456 931.068

41◦ 52 42,3 W 0.624793 0.63684 0.626045 0.634666 0.627495 0.654143 0.648101 0.628714 0.661346 0.648293

UI = 525 +/− 9 Ma; LI = 167 +/−110 Ma (2σ); MSWD = 0.62; P= 0.54 (model 1; n = 4 only∗ ) BUZ-8-5—kyanite–sillimanite–garnet–biotite gneiss (aluminous metassediment) (monazite) Loc: Foca beach, B´uzios city—BUZ-08 (Fig. 3)—Lat: 22◦ 45 54 S Long: 41◦ 52 42,3 W 0.003 6720.6 2395.4 1591.94 0.654194 0.65 A∗ 0.65 A 0.004 4381.7 1329.4 761.651 0.652316 0.65A 0.0042 3117 1095.1 1349.07 0.664044 0.65 A 0.005 1.14% 3687.4 2329.19 0.653639 0.65 A 0.003 2921.2 1080.7 182.495 0.651755

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

BUZ-01-04—garnet–amphibole–diopside gneiss (banded orthoamphibolite) (sphene) Loc: Cabo Frio city—BUZ-01 (Fig. 3)—Lat: 22◦ 53 5 S Long: 42◦ 0 29 W M(4)[s] 0.308 11.211 2.6609 51.1925 0.592585 M(4)[s]∗ 0.604 9.1032 1.4862 90.469 0.639344 0.78 8.9731 1.4527 104.305 0.662885 M(5)[s]∗ M(6)[s] 0.58 9.945 1.5226 113.374 0.659017

UI = 513 +/− 2 Ma; LI = −328 +/440 Ma (2σ); MSWD = 0.116; P = 0.95 (model 1; all fractions) 07-LAG—garnet gneiss (quartz–feldspathic metassediments) (Zircon; Monazite–Mn–Rutile–Rt) Loc: Lagoinha Point, B´uzios city—BUZ-07 (Fig. 3)—Lat: 22◦ 46’3,8 S Long: 41◦ 52 39”W nm(1)[1] 0.001 453.28 33.482 167.688 0.625018 nm(1)[4] 0.003 573.88 42.982 360.793 0.648386 nm(1)[1] 0.003 226.37 20.249 75.5143 0.652151 nm(1)[1] 0.003 644.16 54.005 272.19 0.628703 MN-0.6 A 0.003 443.71 289.81 164.776 0.627286 MN-0.8 A 0.004 1224.6 824.96 659.6 0.630537 RT-1.5A 0.71 334.51 22.087 21270.9 0.566841

0.080169 0.081462 0.079565 0.078484 0.080086 0.080058 0.07253

497.13 504.84 493.53 487.07 496.63 496.47 451.38

492.99 507.49 509.8 495.29 494.4 496.43 455.97

474 519 583 533 484 496 479

+/− +/− +/− +/− +/− +/− +/−

33 12 65 10 14 6 5

0.0582557 0.570497 0.0570852 0.0581093 0.0572986 0.0563227

493.79 487.94 491.93 495.01 499.97 396.1

501.96 488.91 492.44 501.98 500.52 406.36

539 493 495 534 503 465

+/− +/− +/− +/− +/− +/−

21 41 22 45 19 22

UI = 504 +/− 21 Ma; LI= 160 +/− 100 Ma (2σ); MSWD =0.78; P = 0.5 (model 1, with n(x)) BUZ-06-ANF—diopside amphibolite (zircon; sphene—Sp) Loc: Gerib´a Point, B´uzios city—BUZ-06 (Fig. 3)—Lat: 22◦ 46 51,2 S Long: 41◦ 54 23 W nm(0)[1] 0.007 13.725 5.7335 28.8475 0.657982 nm(0)[1] 0.009 24.759 4.4503 50.5744 0.577458 nm(0)[7] 0.008 14.219 5.9346 30.2522 0.561413 nm(0)#[1] 0.004 1223.6 89.992 1077.77 0.619116 nm(0)[3] 0.024 20.133 1.7714 180.893 0.630486

0.083371 0.083738 0.083269 0.077843 0.079098

0.0572397 0.0500145 0.0488988 0.0576834 0.0578105

516.21 518.4 515.61 483.24 490.74

513.38 462.83 452.45 489.29 496.4

501+/−240 196 +/− 150 143 +/− 130 518 +/− 3 523 +/− 17

UI = ∗∗∗ Ma; LI = 487 +/− 6 Ma (2σ); MSWD = 0.648; P = 0.58 (model 1; all fractions) SP-M(1) 0.445 25.489 3.9591 165.433 0.627596 SP-M(3) 0.528 15.059 2.3469 185.315 0.636044 SP-M(4) 0.614 15.421 2.4457 179.741 0.64287 SP-M(5)∗ 0.5 14.056 2.3407 162.985 0.658499

0.079367 0.080125 0.081126 0.082243

0.0573508 0.057573 0.0574731 0.0580702

492.34 496.87 502.84 509.5

494.6 499.85 504.08 513.7

505 513 510 532

+/− +/− +/− +/−

13 9 6 8

BUZ-11—sillimanite–biotite gneiss (quartz–feldspathic metassediments) (monazite) Loc: Ponta Negra beach, Maric´a city—BUZ-11 (Fig. 3)—Lat: 22◦ 57 40.5 S Long: 42◦ 41 38 W 0.002 2862 659.51 1294.12 0.643688 0.081291 0.55A∗ 0.55A∗ 0.003 1319.8 310.79 576.642 0.640984 0.08065 0.55A∗ 0.001 6976.9 1742.4 849.88 0.642065 0.081268 0.55A∗ 0.002 3765.5 929.96 787.607 0.65052 0.082574 0.65A 0.0017 5046.5 757.94 202.806 0.426247 0.053978 0.65A 0.002 6027.1 1616.6 197.895 0.659948 0.082901

0.0574288 0.0576423 0.0573008 0.0571367 0.0572718 0.0577366

503.83 500 503.68 511.47 338.9 513.41

504.59 502.92 503.58 508.8 360.51 514.58

508 516 503 497 502 520

+/− +/− +/− +/− +/− +/−

3 6 4 5 24 5

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

18-PARA—titanite amphibolite (calcsilicate rock) (Zircon) Loc: Jacon´e beach, Ponta Negra district—BUZ-18 (Fig. 3)—Lat: 22◦ 57 14.6 S Long: 42◦ 41 2.7 W nm(0)[1] 0.009 40.752 3.4364 146.992 0.63944 0.079609 0.006 52.988 4.2934 138.818 0.618502 0.07863 nm(0)∗ x[1] nm(0)∗ x[1] 0.009 52.671 4.338 263.156 0.624142 0.079297 nm(0)x[2] 0.005 141.01 24.363 49.968 0.639481 0.079814 0.009 45.259 3.4458 173.895 0.63712 0.080645 M(0)∗ x[3] M(0)x[1] 0.01 90.849 5.3689 575.926 0.492128 0.063371 With n = 3 (∗ )—concentrated age at 494 +/− 8 Ma (2σ); MSWD = 0.27; P =0.93

0.0565437 0.0577267 0.0594462 0.0580982 0.0568081 0.057122 0.0566819

UI = 510 +/− 5 Ma forced through zero; MSWD =1.2 (n = 3, without∗ )

43

44

Table 1 (Continued ) Sample fractiona

Size (mg)

U (ppm)b

Observed Pb (ppm)b

Radiogenic ratiosc 207 Pb∗ /206 Pb∗

206 Pb∗ /238 U

207 Pb∗ /235 U

207 Pb∗ /206 Pb∗

BUZ-20-01—sillimanite–biotite gneiss (quartz–feldspathic metassediments) (monazite–Mn; zircon) Loc: Ponta Negra coast, Maric´a city—BUZ-20 (Fig. 3)—Lat: 22◦ 57 35.7 S Long: 42◦ 41 16.8 W MN-0.55A∗ 0.003 3278 762.22 274.661 0.622286 0.079228 0.003 2940.5 799.05 689.813 0.62027 0.078729 MN-0.55A∗ MN-0.55A 0.0017 2662.8 691.12 334.297 0.591807 0.075075 MN-0.55A 0.0046 3467.8 854.3 1742.64 0.634436 0.080271 MN-0.55A∗ 0.0056 2694.3 613.81 1421.67 0.624319 0.079244 0.0028 3795.5 947.28 1122.06 0.623137 0.079092 MN-0.55A∗ MN-0.65A 0.001 94.177 26.944 55.3226 2.62813 0.2068

0.0569651 0.0571404 0.0571724 0.0573227 0.0571398 0.0571414 0.0921709

491.51 488.53 466.66 497.74 491.61 490.7 1211.7

491.28 490.02 472.02 498.86 492.55 491.81 1308.5

490 +/− 5 497 +/− 5 498 +/− 10 504 +/− 2 497 +/− 2 497 +/− 5 1471+/− 50

UI = 497 +/− 3 Ma; LI = −27 +/−210 Ma (2σ); MSWD = 0.00117; P = 0.99 (model 1, n = 4) Concentrated age with n(∗ ) at 491 +/− 1 Ma (2σ); MSWD = 1.9; P = 0.07 nm(−1)[1] 0.008 516.12 62.459 893.096 0.959115 0.110078 nm(−1)[1] 0.004 150.51 16.917 574.968 1.01134 0.10959 M(−1)[1] 0.006 70.523 11.07 109.943 1.41812 0.14416 M(4)[1] 0.004 124.92 18.911 438.172 1.11044 0.117165

0.0631931 0.0669306 0.0713458 0.0687376

673.2 670.37 868.14 714.23

682.84 709.55 896.57 758.38

715 836 967 891

+/− +/− +/− +/−

3 9 44 8

0.0556575 0.0568189 0.0560764 0.0554448 0.0555013 0.0556345 0.0558083 0.0661364

309.97 275.12 267.35 227.94 279.59 286.11 297.9 437.23

325.58 298.38 287.58 246.8 296.57 303.31 315.15 502.27

439 484 455 430 432 438 445 811

+/− +/− +/− +/− +/− +/− +/− +/−

6 18 7 19 8 4 4 7

0.0585053 0.0585836 0.0584786 0.05865 0.0587316 0.0616241

541.92 501.28 539.82 550.26 534.89 561.77

543.23 510.43 541.34 551.02 539.14 581.86

549 552 548 554 557 661

+/− +/− +/− +/− +/− +/−

2 6 8 5 10 120

206 Pb/204 Pb

207 Pb∗ /235 U

Calculated ages (Ma)d 206 Pb∗ /238 U

UI = 505 +/− 2 Ma, no LI (2σ); MSWD = 0.82; P = 0.44 (model 1; n = 4 ∗ )

∗∗∗ Ma;

LI = 515 +/−

∗∗∗ Ma

(2σ); MSWD = 18; P = 0 (model 2)

BUZ-19-11A—post-tectonic pegmatite (zircon) Loc: Jacon´e beach, Ponta Negra district—BUZ-19 (Fig. 3)—Lat: 22◦ 56 55 S Long: 42◦ 41 0 W M(1)x[1] 0.008 4012 192.88 1078.95 0.378015 0.049259 M(2)[1] 0.005 4035.1 192.57 468.372 0.341589 0.043602 M(3)[1] 0.008 4410.4 198.54 454.193 0.327397 0.042344 M(3)x[1] 0.003 4629.3 163.26 687.827 0.275143 0.035991 M(3)x[1] 0.003 4627.7 253.07 205.447 0.339207 0.044326 M(3)x[1] 0.007 5071.3 248.24 433.745 0.348128 0.045383 M(3)[1] 0.007 3114.5 151.56 510.901 0.363936 0.047296 M(3)[1] 0.008 3343.3 278.15 261.039 0.639944 0.070178 UI = 447 +/− 20 Ma; LI = 21 +/− 37 Ma (2σ); MSWD = 0.56; P = 0.57 (model 1, n – x) BUZ-79—sillimanite–garnet–biotite gneiss (“Oriental terrane”) (monazite–MN; zircon) Loc: C´orrego do Ouro district, Maca´e city—BUZ-79 (Fig. 3)—Lat: 22◦ 14 8.11 S Long: 5888.2 1416.1 1256.56 0.707446 MN-0.55 A∗ 0.0022 MN-0.55 A∗ 0.0037 5399.7 1128.9 2888.51 0.65318 2443.1 1602.4 396.609 0.704273 MN-0.60 A∗ 0.003 MN-0.60 A 0.0026 2995.3 976.13 1656.33 0.720593 MN-0.65 A 0.0025 988.93 632.15 369.274 0.700583 MN-0.65 A 0.0015 3280.2 2224.6 308.35 0.773664

41◦ 58 27.6 W 0.087699 0.080864 0.087346 0.089109 0.086514 0.091054

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

UI = 1090 +/−

0.106043 0.090422 0.087609 0.098614 0.087944 0.094064 0.203288 0.152823 0.090764 0.088266

0.063824 0.0622797 0.0588604 0.0615697 0.0588564 0.0620858 0.121169 0.114448 0.0590352 0.0592802

649.73 558.03 541.38 606.28 543.37 579.52 1193 916.77 560.05 545.27

669.31 583.46 545.34 617.57 546.93 599.77 1503.5 1246 561.7 551.52

736 684 562 660 562 677 1973 1871 568 577

+/− +/− +/− +/− +/− +/− +/− +/− +/− +/−

6 44 14 5 11 26 2 8 25 11

Asterisk (∗ ) denotes radiogenic Pb. Loc.= location; UI = upper intercept; LI = lower intercept. a nm= nonmagnetic; M= magnetic, numbers in parentheses indicate side tilt used on Franz separator at 1.5 A power; [1] = numbers of grains analyzed. b Total U and Pb concentrations, corrected for analytical blank. c Pb corrected for blank and non-radiogenic Pb (see text). d Ages using decay constants recommended by Steiger and Jäger (1977); uncertainties at 2σ.

Table 2 Sm–Nd data Sample

Location

Lithotype

Age (Ma)

Nd (ppm)

Sm (ppm)

147 Sm/144 Nd

143 Nd/144 Nd

εNd (0)

εNd (t)

TDM (Ga)

06-ANF 18-PARA LAG-10 07-LAG BUZ-46 BUZ-8-5 BUZ-79

B´uzios P. Negra B´uzios B´uzios B´uzios B´uzios Oriental terrane

Diopside amphibolite Calcsilicate Granite–diopside gneiss Garnet gneiss Kyanite–sillimanite gneiss Kyanite–sillimanite gneiss Sillimanite–garnet–biotite gneiss

900 900 530 520 520 530 550

17.56 21.79 30.16 24.55 39.11 27.13 53.18

3.894 4.964 6.027 4.992 7.327 5.189 9.813

0.1341 0.1378 0.1208 0.1229 0.1133 0.1156 0.1116

0.512475 0.512522 0.511968 0.512031 0.51191 0.512019 0.51195

–3.19 –2.26 –13.08 –11.84 –14.2 –12.08 –13.43

4.03 4.54 –7.95 –6.95 –8.67 –6.6 –7.45

1.09 1.05 1.75 1.69 1.7 1.58 1.6

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

UI = 548 +/− 3 Ma; LI = −44 +/− 96 Ma (2σ); MSWD = 0.073; P = 0.79 (model 1; n∗ ) nm(−2)[1] 0.007 292.1 30.361 1133.77 0.933186 M(−2)[1] 0.013 166.01 22.205 97.8162 0.776462 M(−2)∗ [1] 0.005 207.92 21.328 172.861 0.711004 M(−2)[1] 0.012 214.04 21.141 573.866 0.837156 M(−2)∗ [1] 0.008 375.38 32.017 1147.04 0.713675 M(−2)[1] 0.006 251.82 33.698 95.5157 0.805221 M(−2)[1] 0.005 245.77 55.22 914.094 3.39627 M(−2)[1] 0.003 591.46 104.85 259.141 2.41155 M(-2)∗ [1] 0.003 119.27 14.009 95.8401 0.738797 M(0)[1] 0.008 206.87 23.512 161.116 0.721445 UI = 1905 +/− 810 Ma; LI= 539 +/−12 Ma (2σ); MSWD = 0.72; P = 0.58 (model 1) Concentrated age from n = 3(∗ ) = 553 +/− 7 Ma (2σ); MSWD = 1.9; P = 0.093

Note. 143 Nd/144 Nd normalized to 146 Nd/144 Nd = 0.72190. εNd (0) calculated relative to CHUR(0) = 0.512638. Model ages (TDM ) were calculated according to the single-stage depleted mantle model of DePaolo (1981). Primary ages used for εNd (t) are based on U–Pb ages known or estimated (italics) based on regional geology.

45

46

Table 3 Summary of SHRIMP U–Pb zircon data for sample BUZ-20 Grain spot

1.21 0.11 0.10 0.29 0.09 0.01 0.09 0.11 1.23 0.10 0.29 −0.03 0.67 0.35 0.46 0.30 0.07 0.42 11.42 0.15 0.06 0.12 −0.03 0.17 −0.01 0.04 0.09 0.16 0.07

ppm (U)

ppm (Th)

232 Th/238 U ppm 206 Pb/238 U 206 Pb∗ agea

1.00 0.27 0.42 0.66 0.15 0.37 0.32 0.72 1.02 0.34 0.19 0.06 0.60 0.42 0.04 0.29 0.00 0.67 0.00 0.09 0.03 0.13 0.06 0.68 0.06 0.50 10.30 0.13 0.00

2497 187 146 193 178 149 157 1281 300 98 219 428 164 196 416 166 105 621 265 256 246 212 537 116 142 174 69 234 94

185 94 22 133 106 135 76 337 142 116 226 126 10 6 138 77 85 47 132 136 278 168 11 2 73 53 24 146 33

0.08 0.52 0.15 0.71 0.61 0.94 0.50 0.27 0.49 1.23 1.07 0.30 0.06 0.03 0.34 0.48 0.83 0.08 0.51 0.55 1.16 0.82 0.02 0.02 0.53 0.31 0.37 0.64 0.37

280.5 910.7 735.7 851.7 1021.8 1047.6 745.8 607.1 823.1 1984.9 997.6 2389.6 471.7 523.2 2013.7 678.4 1816.6 815.1 801.3 721.0 1018.5 671.8 483.9 516.4 1978.1 661.9 352.3 974.1 572.8

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

6.2 20.6 17.4 18.9 23.4 23.1 16.9 13.5 18.2 60.4 21.8 45.6 11.0 12.2 39.6 15.4 38.6 17.6 18.8 15.9 23.9 16.0 11.8 12.4 40.9 14.9 12.9 21.2 15.6

206 Pb/238 U ageb 276.1 909.7 734.5 853.5 1021.0 1053.2 745.5 603.0 814.5 1994.7 998.0 2303.3 468.9 523.8 2016.5 679.3 1801.0 810.7 800.8 718.7 1019.1 671.4 483.3 518.9 1949.2 663.0 355.3 974.7 570.7

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

6.2 21.4 17.9 19.6 24.5 24.4 17.4 13.7 18.6 72.8 22.8 60.6 11.1 12.3 47.8 15.8 44.3 18.1 19.4 16.3 25.0 16.4 11.9 12.6 48.2 15.3 12.2 22.1 15.8

207 Pb/206 Pb Percent agea Disc. 752 935 778 798 1039 929 758 793 1062 1934 988 2615 648 489 2000 638 1915 946 817 802 1005 686 519 363 2120 617 1118 961 677

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

51 40 60 72 27 34 43 38 62 29 33 5 74 82 8 45 18 35 22 32 25 30 25 133 14 55 35 24 47

168 3 6 −6 2 −11 2 31 29 −3 − 9 37 −6 −1 −6 5 16 2 11 −1 2 7 −30 7 −7 −90 −1 18

238 U/ 206 Pb

±% Total 207 Pb/ 206 Pb

±% 207 Pb∗ / 206 Pb∗

±%

22.26 6.57 8.24 7.03 5.81 5.65 8.13 10.05 7.27 2.76 5.96 2.23 13.09 11.78 2.73 8.99 3.07 7.37 7.56 8.44 5.84 9.09 12.82 11.91 2.78 9.20 15.97 6.12 10.76

2.3 2.4 2.5 2.4 2.5 2.4 2.4 2.3 2.3 3.5 2.4 2.3 2.4 2.4 2.3 2.4 2.4 2.3 2.5 2.3 2.5 2.5 2.5 2.5 2.4 2.4 3.1 2.3 2.8

0.6 1.0 1.7 2.3 1.2 1.3 1.6 0.6 1.0 1.1 1.1 0.3 1.7 1.7 0.4 1.4 1.0 0.7 1.1 1.1 1.1 1.3 1.0 2.1 0.8 1.5 5.2 1.0 2.2

2.4 0.39 2.0 1.47 2.8 1.09 3.4 1.28 1.3 1.75 1.6 1.70 2.0 1.09 1.8 0.89 3.1 1.40 1.6 5.89 1.6 1.66 0.3 10.88 3.4 0.64 3.7 0.66 0.5 6.22 2.1 0.93 1.0 5.26 1.7 1.31 1.1 1.21 1.5 1.07 1.2 1.72 1.4 0.94 1.1 0.62 5.9 0.62 0.8 6.52 2.6 0.90 46.7 0.36 1.2 1.60 2.2 0.80

0.0720 0.0723 0.0683 0.0708 0.0751 0.0729 0.0669 0.0711 0.0825 0.1210 0.0735 0.1763 0.0659 0.0602 0.1233 0.0633 0.1173 0.0757 0.0663 0.0666 0.0729 0.0633 0.0582 0.0591 0.1320 0.0642 0.1280 0.0721 0.0621

Error in standard calibration was 0.70% (not included in above errors but required when comparing data from different mounts). Disc = discordant. a Common Pb corrected using measured 204Pb. b Common Pb corrected by assuming 206 Pb/238 U–207 Pb/235 U age-concordance.

0.0643 0.0702 0.0651 0.0657 0.0739 0.0700 0.0645 0.0656 0.0747 0.1186 0.0721 0.1759 0.0612 0.0569 0.1230 0.0610 0.1173 0.0706 0.0663 0.0659 0.0727 0.0623 0.0577 0.0538 0.1317 0.0604 0.0467 0.0711 0.0621

207 Pb∗ / 235 Ua

±%

206 Pb∗ / 238 Ua

±% Error corrected

3.3 3.1 3.8 4.2 2.8 2.9 3.2 3.0 3.9 3.9 2.9 2.3 4.2 4.4 2.3 3.2 2.6 2.9 2.7 2.8 2.8 2.9 2.8 6.4 2.5 3.5 46.8 2.6 3.6

0.0445 0.1517 0.1209 0.1412 0.1718 0.1765 0.1227 0.0988 0.1362 0.3606 0.1674 0.4487 0.0759 0.0845 0.3667 0.1110 0.3255 0.1348 0.1324 0.1183 0.1712 0.1098 0.0779 0.0834 0.3591 0.1081 0.0562 0.1631 0.0929

2.3 2.4 2.5 2.4 2.5 2.4 2.4 2.3 2.4 3.5 2.4 2.3 2.4 2.4 2.3 2.4 2.4 2.3 2.5 2.3 2.5 2.5 2.5 2.5 2.4 2.4 3.8 2.3 2.8

0.681 0.776 0.661 0.568 0.882 0.824 0.763 0.787 0.606 0.911 0.827 0.991 0.575 0.545 0.979 0.749 0.928 0.800 0.921 0.837 0.902 0.872 0.914 0.391 0.949 0.679 0.081 0.892 0.788

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

BUZ-20-1.1 BUZ-20-1.2 BUZ-20-2.1 BUZ-20-3.1 BUZ-20-4.1 BUZ-20-5.1 BUZ-20-6.1 BUZ-20-7.1 BUZ-20-7.2 BUZ-20-8.1 BUZ-20-9.1 BUZ-20-10.1 BUZ-20-11.1 BUZ-20-12.1 BUZ-20-13.1 BUZ-20-14.1 BUZ-20-15.1 BUZ-20-16.1 BUZ-20-17.1 BUZ-20-17.2 BUZ-20-18.1 BUZ-20-19.1 BUZ-20-20.1 BUZ-20-21.1 BUZ-20-21.2 BUZ-20-22.1 BUZ-20-23.1 BUZ-20-24.1 BUZ-20-25.1

Percent 206 Pb ∞a c

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

47

Fig. 6. Plot of U–Pb data from samples of Região dos Lagos and Forte de São Mateus Unit. Uncertainties at 2σ (n, number of fractions analysed).

48

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

BUZ-44 banded amphibolite are variably discordant and gave an upper intercept of 1969 ± 4 Ma and a lower intercept of 519 ± 19 Ma (Fig. 6e). The data indicate that these mafic rocks crystallized contemporaneously with the Paleoproterozoic felsic igneous rocks and were probably metamorphosed and deformed together during this possible Cambrian event (corresponding to the lower intercept). An additional banded amphibolite sample provided further evidence for the Cambrian metamorphism. One concordant sphene fraction gave a 207 Pb/206 Pb age of 519 ± 9 Ma (Fig. 6f).

4.3. Supracrustal data 4.3.1. Provenance ages Sm–Nd data from whole-rock samples of the supracrustal unit indicate TDM model ages ranging from 1.75 to 1.0 Ga (Table 2). These could be interpreted as maximum ages for the basin formation. More accurate data came from SHRIMP U–Pb analyses on a paragneiss sample (BUZ-20-01) from the Palmital succession at Ponta Negra (Table 3), thought to be deposited by turbiditic processes. The

Fig. 7. SHRIMP and ID–TIMS U–Pb data for detrital zircons from a paragneiss of Palmital succession—sample BUZ-20-01. (a) U–Pb analysis from detrital zircon cores—SHRIMP; (b) plot of number of analysis against 207 Pb/206 Pb ages indicating at least four sources for the zircons; (c) SHRIMP analysis in metamorphic overgrowth; (d) U–Pb conventional ID–TIMS analysis (n, number of fractions analyzed).

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

49

Fig. 8. Cathodoluminescence images of zircons from a paragneiss sample of Palmital succession (BUZ-20-01). Ion microprobe sample spots and derived 207 Pb/206 Pb ages are indicated.

cores of the detrital zircons gave at least four age groups related to the source rocks: 2.5, ∼2.0, ∼1.0 Ga and a younger and most abundant population with ages ranging from 800 to 620 Ma (Figs. 7a,b and 8). The metamorphic overgrowth is normally very thin and allowed only four spot analyses giving an upper intercept age of 511 ± 31 Ma (Fig. 7c). Conventional ID–TIMS U–Pb analyses of four detrital zircons from the paragneiss are all discordant and reveal 207 Pb/206 Pb ages between 950 and 700 Ma (Fig. 7d). Only one monazite from the same paragneiss sample (BUZ-20-01) is clearly inherited, with a 207 Pb/206 Pb age of 1471 ± 50 Ma (Table 1). Morphologically this monazite is similar to other metamorphic monazites.

4.3.2. Metamorphic ages The oldest metamorphic age was observed for sample BUZ-8-8, a syntectonic leucosome in a migmatitic outcrop of the Búzios succession (Fig. 4c). A very homogeneous and consistent population of zircons of anhedral/subhedral, very transparent, pink crystals was identified. The subhedral grains show some facets indicating a bypiramidal shape. Six grains were abraded and analyzed yielding an upper intercept of 525 ± 9 Ma (Fig. 9c). Since the leucosomes were generated during the metamorphic peak, coeval with deformation phases D1/D2, this age is interpreted as the crystallization age of those veins, before phase D3, when they were folded (Fig. 4c). Monazites from the same sample are reversely discordant. In this case, the 207 Pb/235 U ages, constrained between 500 and

50

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

(b)

(a) 0.085

524 520

0.084

BUZ - 8-5 Kyanite Gneiss Búzios Succession (monazite; n = 4)

0.084

BUZ-46 Kyanite Gneiss Búzios Succession (monazites; n=5)

520

516

0.083

516

512

0.083 206

Pb/

238

206

512

U

U 508

0.082

511 +/- 2 Ma (n =2) MSWD = 1.19

508

0.082

Pb/

238

Intercepts at 513 +/- 2 Ma & -328 +/- 440 Ma MSWD= 0.116

504

504 0.081

0.081

500

500

0.080 0.63

0.64

0.66

0.65 207

Pb/

235

0.67

0.080

0.68

0.63

0.64

(c)

0.66

0.65 207

U

(d) 540

0.088

BUZ-8-8 Leucosome Búzios Succession (monazites-gray - n =4) (zircons - black- n=6)

0.085

0.083

206

Pb/

238

520 0.084

510 206

520

500

Pb/

238

U

480

0.076

500

0.081

07-LAG Garnet Gneiss Búzios Succession (zircon - black; monazite-gray; rutile-filled ellipse)

0.080

U Upper Intercept at 525 +/- 9 Ma MSWD= 0.62 (zircons; n = 4)

490

0.079

0.67

Pb/ 235 U

460 0.072

0.068

0.077

Rutile 479 +/- 5 Ma 207 Pb/ 206 Pb

440

420

0.064

0.60

0.62

0.66

0.64 207

Pb/

235

0.68

0.48

0.52

0.56

0.60 207

U

Pb/

0.64 235

0.68

0.72

U

Fig. 9. Plot of U–Pb data from samples of B´uzios succession. Uncertainties at 2σ (n, number of fractions analyzed). In part (d) of the figure, two ellipses in gray represent two monazite crystals indicating a concordant age at 496 ± 4 Ma (MSWD, 0.97).

493 Ma (Table 1), are considered to be more reliable, based on the arguments of Parrish (1990). The aluminous metasedimentary rocks from the Búzios succession yielded upper intercepts U–Pb monazite ages of 511 ± 2 Ma (sample BUZ-8-5) and 513 ± 2 Ma (sample BUZ-46) (Fig. 9a and b). The 10 million years gap between ages of monazites from the leucosome and those from the metasedimentary rocks could be related to the origin of those monazites: the first crystallized from a partial melt, and the second recrystallized in the solid state. However, a quartz

feldspathic metasedimentary rock (07-LAG) from the Búzios succession (garnet gneiss) has two concordant monazite analyses at 496 ± 4 Ma (Fig. 9d). This similarity with the leucosome age could be related to the fact that this sample was collected in the migmatitic domain. On the other hand, there is always the possibility of more than one generation of monazites, which should be detailed, in future studies. The garnet gneiss has a population of round pristine very small zircons, interpreted as metamorphic crystals that are very distinct from the detrital zircons grains of the

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

other metasedimentary rocks, which are larger, pink and acicular. The ages from these round zircons have large uncertainties, but one crystal is almost concordant and gave a 207 Pb/206 Pb age of 519 ± 12 Ma (Fig. 9d). This age is compatible with data from the leucosome zircons, meaning that during the metamorphic peak growth of new zircon crystals took place within the metasedimentary rocks. This same sample also yields a 207 Pb/206 Pb age of 479 ± 5 Ma for a multigrain fraction of discordant rutile (Fig. 9d). A diopside amphibolite within the Búzios succession (BUZ-06-ANF) presents a population of round

51

metamorphic zircons with very low Pb contents (1–89 ppm) (Fig. 10c). One zircon crystal yields a 207 Pb/206 Pb age of 518 ± 3 Ma (Fig. 10d). The same sample has an upper intercept age of 510 ± 5 Ma for three fractions of sphene (Fig. 10b), one of which is almost concordant at 510 ± 6 Ma (Fig. 10d). This sphene age is coherent with the sphenes from amphibolites of the basement (Fig. 6f). The supracrustal rocks from the western portion of CFTD present younger metamorphic ages. A sphene-amphibolite sample (BUZ-18-PARA) within a calc-silicate lens from the Palmital succession has

Fig. 10. Plot of U–Pb data from amphibolites samples within the Supracrustal sequence. Uncertainties at 2σ (n, number of fractions analyzed). Zircon fractions with thicker lines were used in regression in (a).

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Fig. 11. Plot of U–Pb data from samples of Palmital succession (a–c) and a late-tectonic pegmatite (d). Uncertainties at 2σ (n, number of fractions analyzed).

a very homogeneous population of round colourless zircons with very low Pb contents (1–24 ppm). Despite the large uncertainties, a regression with three fractions yielded an upper intercept age of 494 ± 8 Ma (Fig. 10a). Two samples from a typical Palmital paragneiss (BUZ-11 and BUZ-20-01) produced concordant monazites at 503 ± 4 Ma and 490 ± 5 Ma (Fig. 11a and b), with a weighted 207 Pb/206 Pb average at 491 ± 1.5 Ma in the last sample (Fig. 11c). Regressions with other monazite fractions in both samples yielded ages of 505±2 Ma and 497±3.5 Ma, respectively (Fig. 11a and b). The zircon U–Pb age from the amphibolite and the monazite ages from the paragneiss indicate that the metamorphic peak in the western area of the CFTD

is at least 10 million years younger than that in the eastern area (considering the uncertainties). This is coherent with the structural evidence suggesting that the western area was affected by a late transpressional shear zone, attributed to D4. 4.4. Late-tectonic pegmatite The CFTD units are cross-cut by several latetectonic pegmatite dikes, with almost no evidence of deformation. Zircons from one sample of these dikes (BUZ-19-11) indicated an upper intercept age of 440 ± 11 Ma (Fig. 11d) with five fractions and forcing the lower intercept through zero. The zircons from this dike have very high U concentrations

R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

53

Fig. 12. Plot of U–Pb data from a metapelite sample of the “Oriental terrane”. Uncertainties at 2σ (n, number of fractions analyzed). Fractions filled with gray are detrital zircon grains.

(3115–5071 ppm), which is probably responsible for their strongly discordant behavior. Taking only four fractions and not forcing the lower intercept leaves an age of 447 ± 20 Ma. 4.5. “Oriental terrane” sample One sample from an aluminous metasedimentary rock from the “Oriental terrane”, adjacent to the CFTD to the NW, was also investigated. The sillimanite–garnet–biotite gneiss sample (BUZ-79) belongs to the São Fidélis Unit (Costeiro Domain; Heilbron et al., 2000; Fig. 2) and was collected from an outcrop 5 km NW of the contact between the CFTD and the “Oriental terrane” in the vicinities of Macaé (Fig. 3). Three monazite fractions yield an upper intercept age of 548 ± 2 Ma (Fig. 12a). A concordant monazite crystal indicates a 207 Pb/206 Pb age of 554 ± 5 Ma (Fig. 12a). Ten zircon crystals were analyzed in the same sample, and two populations were recognized: discordant zircons with 207 Pb/206 Pb ages ranging from 1973 to 659 Ma (Table 1) and almost concordant zircons with 207 Pb/206 Pb ages between 577 and 561 Ma (Fig. 12b). An age of 553 ± 7 Ma was calculated using three fractions (Fig. 12b). One concordant zircon yielded the 207 Pb/206 Pb age of 562 ± 11 Ma (Fig. 12b). The first group of zircons is interpreted to be detrital; whereas the second group is considered to be metamorphic.

5. Discussion The Neoproterozoic–Early Paleozoic evolution of the CFTD started with the deposition (sediments) and crystallization (volcanics) of the supracrustal rocks in a marine sedimentary basin after 620 Ma (Pb–Pb age in the core of detrital zircon) and before 525 Ma (metamorphic event). At 525 Ma, the rocks from the Búzios–Palmital basin were metamorphosed at high-grade and deformed together with several Paleoproterozoic basement thrust sheets that had been tectonically interleaved with them. This low angle thrust regime of deformation, subdivided into three deformational phases (D1–D3), developed isoclinal and disrupted folds, mylonitic zones, and NW–SE stretching lineations with shear-sense to NW, in both lithostratigraphic sequences. During D1/D2 these rock units reached their metamorphic peak with the generation of partial melt veins oriented parallel to the main foliation (S1/2 ). A new zircon population grew in leucosomes derived from the basement and from the metapelites, yielding upper intercepts ages of 518 ± 5 Ma and 525 ± 9 Ma, respectively. Within the supracrustal rocks, there are zircon populations in the amphibolites with an age of 518 ± 3 Ma and in the metasedimentary rocks with an age of 519 ± 12 Ma. These Cambrian ages are interpreted as indicative of the crystallization of the leucosome and re-crystallization of zircons in the supracrustal rocks.

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R. S. Schmitt et al. / Precambrian Research 133 (2004) 29–61

The estimated temperature for the metamorphic peak in the CFTD is similar to zircon closure temperature, at medium to high-pressure, suggesting that these ages indicate the peak of metamorphism and do not represent exhumation, as interpreted in other granulitic regions (Roberts and Finger, 1997). In addition, these Cambrian zircons have at least three origins: some derive from partial melting and subsequent magmatic crystallization (leucosomes), some from breakdown of other minerals, such as garnet and hornblende (metamorphic reactions that liberate Zr enough for zircon growth during the metamorphic peak) (Fraser et al., 1997), and some from re-crystallization of detrital zircons. Both zircons gave ages compatible with the metamorphic peak temperatures. Therefore, the period of time between 525 and 518 Ma corresponds to a tectonometamorphic event that overprinted all preexistent deformation structures and lower grade metamorphic assemblages in these rocks. After the metamorphic peak, the P–T–t path shows a clockwise pattern coeval with the D3 phase that generated recumbent folds (folding also the leucosomes), shear zones, and alignment of sillimanite crystals. The minerals with lower closure temperatures (monazite, sphene, and rutile) show an array of metamorphic ages related to the decompression and cooling. Monazites yielded ages a few million years younger than the zircons, reflecting their closure temperature at approximately 700 ◦ C (Copeland et al., 1988). Monazites of the aluminous metasedimentary rocks of the eastern part of the CFTD yielded older ages (513–511 Ma) than those in the western portion (505–497 Ma). One hypothesis to explain this is the development of a high-temperature D4 shear zone along the contact with the “Oriental terrane”. In this model, the western area would have taken more time to cool, which is reflected in younger monazites and zircon ages (sphene-amphibolite). Another hypothesis is that more than one generation of monazites grew and they may not be identified by their morphology. An additional possibility is that the stability of monazites in metamorphic rocks could be altered by retrograde metamorphism associated with deformation (Lanzirotti and Hanson, 1996). Sphenes from amphibolites of the basement and the supracrustal rocks yielded U–Pb ages at around 510 Ma. The closure temperature for this mineral depends on several factors such as grain size and cooling

rate (Mezger et al., 1991; Scott and St-Onge, 1995). In the CFTD, sphenes occur mainly as aureoles around ilmenite and rutile, probably related to decompression. The ages are similar to the monazite ages, indicating a high closure temperature for sphenes (around 700–650 ◦ C), consistent with fast cooling. The subsequent recumbent folding of D3 phase occurred probably between 518 and 505 Ma. The upper limit (518 Ma) is estimated from the leucosome solidification, folded by D3. The lower limit (505 Ma) is inferred from the relation with D4 that probably took place between 505 and 495 Ma (monazite and zircon ages in the western area). Therefore, after crustal thickening at depths in excess of 30 km, the CFTD rock units started to cool while a major thrust zone juxtaposed this domain with the “Oriental terrane” during D3. The adjustment of stresses within both tectonic domains after collision was the possible cause for the development of the D4 dextral shear zone along the western contact. Ordovician ages were obtained from a discordant rutile fraction (207 Pb/206 Pb age: 479 ± 5 Ma) and a late-tectonic pegmatite dike (447 ± 20 Ma). They represent the low closure temperature of rutile for the U–Pb system (420–380 ◦ C, Mezger et al., 1989). The U–Pb data in this study allow the reconstruction of a preliminary cooling curve for the CFTD, combining the age of the minerals and their estimated closure temperatures for geochronologic systems (Mezger et al., 1991; Fig. 13). The calculation of the cooling rate considers the errors of U–Pb ages and estimated closure temperatures. The eastern area shows rocks that underwent high temperatures (850–600 ◦ C) during 20 million years at most. For the first 10 million years, it is possible to construct two average cooling rate curves: (1) one representing a 17 ◦ C/Ma rate, using zircon and sphene ages, and (2) one representing a 30 ◦ C/Ma rate, using only monazite and sphene ages. Considering this data, the best rate would be a combination between monazite/sphene data against zircon data, admitting a drop of 200 ◦ C in 10 million years, giving a cooling rate of 20 ◦ C/Ma. After 510 Ma, the cooling curve went down to 10 ◦ C/Ma until 480 Ma and then to 5 ◦ C/Ma. These last rates were based on two K/Ar analyses in biotite (Delhal et al., 1969; Zimbres et al., 1990) together with the rutile Pb–Pb age. The decrease in cooling rates is consistent with a considerably fast decompression af-

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Fig. 13. Cooling path for the central-eastern part of the Cabo Frio Tectonic Domain (CFTD) in a temperature–time diagram. Each point represents the age of a sample; circles are zircons, boxes are monazites, triangles are sphenes, empty circle is rutile. Cross represents K–Ar in biotite from Zimbres et al. (1990) and Delhal et al. (1969). The empty boxes encircling the points represent error boxes. Samples BUZ-11 and BUZ-20-01 are from the western area of CFTD. Dashed lines represent cooling paths with the calculated rate above. Closure temperatures for all minerals were estimated using data from Mezger et al. (1989, 1991), Copeland et al. (1988), Gebauer and Grünenfelder (1979), Lanzirotti and Hanson (1995), Hawkins and Bowring (1999), Parrish (1990), Scott and St-Onge (1995).

ter the metamorphic peak in the eastern area, which is also demonstrated by the well preserved, medium to high pressure parageneses and the structural development of the D3 thrust zones that placed CFTD units in higher crustal levels. In this kind of diagram (Fig. 13) it is possible to observe more clearly the slightly distinct evolution for the eastern and western portions of CFTD; with the latter being at higher temperatures around 500 million years ago. It is clear that the CFTD has not only different petrogenetic assemblages from the “Oriental terrane” but presents also distinct ages. The metapelite sample (BUZ-79) from this terrane indicates a metamorphic peak at 555 Ma with growth of zircon and monazite (Fig. 12a and b). This is consistent with U–Pb ages obtained in syntectonic granites of this terrane in Rio de Janeiro city (Silva et al., 2000; 559 ± 4 Ma and 560 ± 7 Ma), and with U–Pb ages from granulitic zircons in Esp´ırito Santo State (to northeast) (Söllner et al., 1989; 558 ± 2 Ma). These ages also coincide

with the collisional period of the Rio Doce orogeny defined by Campos Neto and Figueiredo (1995) between 560 and 530 Ma. This age (∼550 Ma) is intermediate between M1 (590–570 Ma) and M2 (535–520 Ma) metamorphic phases defined by Machado et al. (1996) for the central segment of the Ribeira Belt. The M2 period should be, therefore, revised. M1 is characterized by medium-pressure metamorphism with anatexis increasing from west to east and affecting both Occidental and Oriental terranes; M2 affects only the “Oriental terrane”, generating parageneses of high temperature and low pressure (Heilbron et al., 2000). These authors identify five stages in the tectonic evolution of the central segment of the Ribeira belt: (1) pre-collisional stage—SE dip subduction zone developed close to the SFC passive margin generating the Rio Negro magmatic arc at 630 Ma (Tupinambá et al., 1998); (2) syn-collisional stage—frontal collision between the “Oriental terrane” (Rio Negro arc and

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associated sediments) and the “Occidental terrane” (passive margin sediments of the São Francisco Craton; SFC; and their basement), developing ductile thrust shear zones with top to NW and metamorphic assemblages M1; (3) late-collisional stage—between 565 and 540 Ma, with multiple granitoid intrusions during a oblique compressive regime; (4) post-collisional stage—intrusion of small calcalkaline plutons within vertical dextral shear zones between 540 and 520 Ma associated with M2; (5) transitional stage—orogen thermal relaxation between 520 and 480 Ma with calkalkaline plutons associated with tholeiitic rocks (Junho et al., 1993; Wiedmann, 1993). The thermotectonic event characterized in the CFTD is coeval with Heilbron’s stages 4 and 5, but with distinct tectonic meaning: the post-collisional stage is a pre-collisional stage in the CFTD and the transitional stage corresponds to the syn-collisional stage of the CFTD. In fact, one of the sources for the CFTD sediments could well be the Rio Negro magmatic arc. Therefore, during the syn-collisional stage in the internal parts of the Ribeira belt, the Búzios–Palmital basin was active. The 520-490 Ma event characterized in this paper defines a new collisional period. The CFTD contains assemblages of high temperature and medium to high pressures that do not appear in the adjacent “Oriental terrane”. Several tectonic hypothesis could be modeled for the “docking” of the CFTD, but the geological data from adjacent domains, “Oriental terrane” to the west, and in Angola in the African continent to the east, are at present not well-known. It is clear, however, that a collisional event was responsible for the parageneses developed in the Búzios–Palmital sediments and that reworked basement (margin of a large continent or a microplate) participated in this crustal thickening event. The high-pressure mineralogy also suggests that the supracrustal rocks were transported to low crustal depths possibly along an active subduction zone. The question whether the subduction was to NW or SE remains unsolved. 5.1. The Búzios Orogeny and the Gondwana assembly The metamorphic and structural data, together with the ages, provide strong arguments for the characterization of this Cambrian event in this segment of the

Ribeira belt. This tectonometamorphic event is the youngest identified in this belt and probably represents one of the final orogenic events of the Gondwana amalgamation. We propose here to designate this event as a separate orogeny because we consider that it satisfies the definition of an orogeny (Brown, 1993; Burg and Ford, 1997). Crustal thickening in the CFTD is represented by thrust sheets and nappes as a consequence of collision. The sediments related to this orogeny is represented by the Búzios–Palmital basin, probably a submarine turbidite fan deposited in a relatively short period at the end of the Neoproterozoic and the beginning of the Cambrian. The deformation and regional metamorphism are well constrained by phases D1–D4 and the P–T–t paths. The magmatism is restricted to partial melts, at least in the CFTD, but more of this event could be characterized in the adjacent tectonic domains. Late Precambrian–early Paleozoic events are based on geological evidence from ca. 930 Ma (Arenópolis belt, central Brazil; Pimentel and Fuck, 1992) up to ca. 450 Ma (this paper). Therefore, the Brasiliano Orogenies lasted more than 480 million years. The suggestion to use the term orogeny for the Búzios Orogeny comes as an additional and relevant piece of Pan-African–Brasiliano puzzle, attempting to elucidate one final stage of the amalgamation of the Gondwana. The location of this young orogeny, in the central part of West Gondwana (Fig. 1), indicates that, during the Cambrian, the future Gondwana supercontinent was not only surrounded by marginal orogenies but was also still amalgamating in its interior part. Recently other data supporting Cambrian thermotectonic events in Pan-African–Brasiliano belts have been presented (e.g. Dürr and Dingeldey, 1996; Jung et al., 2000; Ring et al., 2002), together with paleomagnetic data (Li and Powell, 1993; Powell, 1993). In the African counterparts, two tectonometamorphic events are described in the Kaoko belt at 550–530 Ma and 460 Ma (Dürr and Dingeldey, 1996). The high-grade metamorphic peak was dated in syn-tectonic granites with ages between 567 and 552 Ma (Seth et al., 1998). The Damara belt has some syn- to late-tectonic S- and A-type granites with ages between 526 and 516 Ma (Jung et al., 2000) attributed to the metamorphic peak and post-tectonic alkaline granites of 494 and 488 Ma. Kock (1992) also de-

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scribes continental collision in the Damara belt in the age interval 550–485 Ma. The Cambrian–Ordovician tectonic activity described in this study is coeval with several well-known orogenies along the margins of Gondwana. The Pampean Orogeny (532–518 Ma; Rapela et al., 1998a,b) and the subsequent Famatinian Orogeny (ca. 490 Ma) are responses to subduction in the ancient Andean region in Argentina (Rapela et al., 1998a,b). The first is the result of the subduction of an oceanic plate to the west, generating 530 Ma granitoids in a passive margin sequence. The collision between the margin with the Pampean terrane occurred at 525 Ma and was accompanied by crustal thickening, ophiolite obduction and metamorphic conditions of 8.6 kbar and 810 ◦ C (Rapela et al., 1998b). The Famatinian Orogeny contains a magmatic arc formed by the subduction of an oceanic plate to the west resulting in the accretion of the Precordillera terrane (Pankhurst et al., 1998). Both orogenies are possibly related to a connection between Laurentia and Gondwana by mid- to late-Cambrian times. These orogenies could have induced the Mozambique orogeny (Grunow et al., 1996). In Antarctica and Australia, the Ross and the Delamerian orogenies are coeval and related to the subduction of Pacific Ocean floor after rifting of Laurentia (Dalziel, 1991). The Neoproterozoic continents that had not yet collided were forced to amalgamate by the marginal orogenies “trap”, leading to the closure of young seas. This could well be the cause of the Búzios Orogeny, since its time span coincides with most of the “marginal” orogenies. It is important to emphasize that the eastern limit of the CFTD is unknown. Rough estimates indicate that at least 250 km of continental crust is covered in Brazilian and African coastal platforms (Chang et al., 1992; Austin and Uchupi, 1982). This would correspond to almost the total width onshore of the Ribeira Belt (Fig. 2). The occurrence of Cambrian high-grade belts along present African (Kaoko and West Congo) and Brazilian coast lines (Ribeira, Araçua´ı and Dom Feliciano) shows that the Mesozoic South Atlantic rifting closely follows Paleozoic sutures of West Gondwana. Therefore, the tectonic and sedimentary processes that started to take place in the Paraná Basin, in SE Brazil, during the Early Paleozoic, should not only be related to the marginal orogenies mentioned above (Milani

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and Zálan, 1999), but also to the contemporaneous orogenies in central Gondwana.

6. Conclusions The Búzios Orogeny started with the subduction followed by collision and crustal thickening reaching conditions of high pressure and temperature at a depth of at least 27 km. During the metamorphic peak, the supracrustal rocks of the Búzios–Palmital basin were covered by thrust sheets of reworked Paleoproterozoic basement during D1 and D2 deformational phases at about 525 Ma. The juxtaposition of continental basement over ocean floor rocks and the clockwise P–T–t path involving relatively high pressures indicate a collisional setting. During the subsequent D3 phase, large recumbent folds were developed with subordinated thrust zones, one of which caused the CFTD to override the “Oriental terrane”. Pseudomorphic sillimanite over kyanite indicates the return P–T–t path. This occurred between 518 and 505 Ma. The cooling rate until 505 Ma was as high as 20 ◦ C/Ma. After the collision, a NE–SW trending dextral subvertical shear zone deformed the western portion of the CFTD between 505 and 495 Ma still at amphibolite facies conditions. Therefore, the collisional stage of the Búzios Orogeny occurred between 525 and 495 Ma and the orogenic collapse stage from 495 to 440 Ma, as indicated by late-tectonic pegmatite age.

Acknowledgements This work was sponsored by CNPq Grant n.200978/97.7 to Dr. Schmitt. Operation of the Isotope Geochemistry Laboratory at the University of Kansas was supported in part by funds from the Department of Geology and by various NSF grants to Van Schmus. We thank Dr. B.B. Brito Neves and Dr. E.P. de Oliveira for their thoughtful reviews.

Appendix A. Analytical methods: geochronologic U–Pb data, ID–TIMS U–Pb data was obtained at the Isotope Geochemistry Laboratory (IGL), Department of Geology and

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Kansas University Center for Research, University of Kansas, Lawrence, KS. Zircon fractions used for isotope dilution analyses were carefully selected by hand prior to dissolution. Some fractions were air-abraded according to Krogh (1982) methodology. Zircons were dissolved with a HF–HNO3 mixture and Pb and U were separated using procedures modified after Krogh (1973) and Parrish (1987). All samples were total-spiked with a mixed 205 Pb/235 U tracer solution. The U–Pb isotopic analyses were done in a single-collector ion-counting mode using a newly fitted ion-counting Daly system in a VG-Sector mass spectrometer. Most analyses included ion exchange purification of Pb and U, with both loaded on the same Re filament using phosphoric acid–silica gel and measured as Pb+ and UO2 + . Pb compositions were corrected for mass discrimination as determined by analysis of NBS-SRM-982 (equal-atom) Pb and monitored by analysis of NBS SRM-983 (radiogenic) Pb. Uranium fractionation was monitored by analyses of NBS SRM U-500. Uncertainties in Pb/U ratios due to uncertainties in fractionation and mass spectrometry for typical analyses are ±0.5%; in some instances weak signals caused uncertainties to range up to ±2%. Radiogenic 208 Pb, 207 Pb and 206 Pb were corrected for modern blank Pb and for non-radiogenic original Pb corresponding to Stacey and Kramers (1975) model Pb for the approximate age of the sample. Uncertainties in radiogenic Pb ratios are typically less than ±0.1% at the 2σ level unless the samples had a low 206 Pb/204 Pb ratio, in which case uncertainties in the common Pb correction could cause greater uncertainties. Decay constants used here were 0.155125 × 10−9 per year for 238 U and 0.98485 = 10-9 per year for 235 U. Blanks ranged from 5 to 30 pg total Pb; in most cases they do not contribute significantly to uncertainties in the ages of the samples, although some of the single-crystal analyses may show effects of blank Pb as larger uncertainties in the calculated ages. Zircon data were regressed using the Microsoft Excel version of Isoplot (Ludwig, 1999). Model 1 regressions were accepted if probabilities of fit were better than 20%; model 2 regressions were used if probabilities of fit were less than 20%. Uncertainties in concordia intercept ages are given at the 2σ level.

The procedures differ a little in the analyses of the other minerals. Monazite analyses were single crystals or even aliquots. They were dissolved with nitric acid in beakers and not microcapsules. The column procedures were the same used for zircons. Sphene fractions were multigrain (around 150 crystals), carefully selected by hand. The process differed from the zircon process in the column stage. In sphene, the U-collection is done with HBr. The first analyses had very weak signal at the spectrometer. After deciding to pass the fractions twice through the column procedure, the analyses came out better with very strong signals. The rutile analysis was taken in a multigrain fraction, carefully selected by hand. The U-collection in columns is done purified twice with HBr.

Appendix B. Analytical methods: geochronologic Sm–Nd data Rock powders for Sm/Nd analyses were dissolved and REE were extracted using general methods of Patchett and Ruiz (1987). Isotopic compositions for Nd were measured with a VG Sector multicollector mass spectrometer. Sm was loaded with H3 PO4 on a single Ta filament and typically analyzed as Sm+ in static-multicollector mode or single-collector mode. Nd was loaded with phosphoric acid on a single Re filament having a thin layer of AGW-50 resin beads and analyzed as Nd+ using dynamic-multicollector mode. External precision based on repeated analyses of the internal standard is ±40 ppm (2σ) or better; all analyses are adjusted for instrumental bias determined by measurements of the internal standard for periodic adjustment of collector positions; on this basis the analyses of La Jolla Nd average 0.511860 ± 0.000010. Eight recent analyses of BCR-1 yielded Nd = 2944 ± 0.70 ppm, Sm = 6.77 ± 0.21 ppm, 147 Sm/144 Nd = 0.1393 ± 0.00071, and 143 Nd/144 Nd = 0.512641 ± 0.000007, yielding εNd (0) = 0.07 ± 0.12 (all at 1σ). Sm/Nd ratios are correct to within ±0.5%, based on analytical uncertainties; εNd (t) values were calculated using the U–Pb ages defined from zircons and other minerals where available or estimated ages based on the regional geology and current results from nearby samples. Crustal residence ages (TDM ) were calculated following the model of De Paolo (1981).

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