Petrologic and geochronological constraints on the polymetamorphic evolution of the collisional granites, Araçuaí Orogen (SE Brazil)

Abstract

Collisional granites of the Araçuaí Orogen, southeastern Brazil, record petrological and geochronological evidence for multiple crustal melting during the orogeny evolution. U-Pb zircon data indicate that these granites crystallized at 586 ± 2 M.y. High-grade metamorphism (M₁) involved partial melting by fluid-absent reactions that produced the first generation of garnet in temperatures of approx. 750°C. Preservation of the mineral assemblage A₁ (garnet-biotite-plagioclase-K-feldspar-quartz-ilmenite-melt) indicates that most of the generated melt was lost from these rocks at or near peak metamorphic conditions. A second metamorphic event (M₂) is characterized by growth of a second generation of garnet in preserved A₂ assemblage (garnet-sillimanite-biotite-plagioclase-K-feldspar-quartz-ilmenite-melt). Mineral equilibria modeling constrains conditions of M₂ metamorphism to 713-729 °C and 6.2-7.3 kbar. Retrograde assemblage (A₃) records equilibrium conditions at 610-660 °C. The Hf isotope composition indicates significant crustal contribution to the genesis of the collisional granites. The elevated geotherms in thickened crust provide enough heat for the M₁ event at 562 ± 2 M.y. Subsequent heating probably associated to the transfer of mantle heat to the crust during the extensional thinning and gravitational collapse of the orogen lead to the M₂ event at 526 ± 4 M.y. This event is concomitant to the emplacement of the post-collisional magmas in the orogen.

Key words
Araçuaí Orogen; partial melting; pseudosection; U-Pb zircon; collisional granite; recycling

INTRODUCTION

Orogenic belts are elongated crustal portions characterized by intense deformation, metamorphism and granitic magmatism, resulting from the collision of continental lithospheric plates with intra-plate shortening, crustal thickening and elevation of the relief (Kearey et al. 2009). In active orogens, such as the Himalayas, the study of their internal structure is limited by topography and depth. In contrast, deeply eroded ancient orogenic belts expose their roots, and their investigation provides important information about the development of ancient orogenic structures and are of great value for continental reconstructions.

Geothermobarometry of high-grade rocks is an important tool for the understanding of tectonic and rock forming processes and large-scale transport of matter at great depths within the continental crust. The record of changes in metamorphic conditions could be preserved as microstructures, mineral relicts, and chemical zoning. Several conventional geothermometers and geobarometers have been used to estimate the equilibrium conditions of determined minerals (e.g., Thompson 1976, Ferry & Spear 1978, Newton & Haselton 1981, Hodges & Spear 1982, Pigage & Greenwood 1982, Koziol 1989, Williams & Grambling 1990). Theriak-Domino (De Capitani & Petrakakis 2010) computer program package is commonly applied in petrology to constrain P-T conditions of metamorphic rocks. Combining internally consistent thermodynamic data set with composition-activity models for critical minerals have taken the more precise estimates of metamorphic conditions. U-Pb zircon geochronology is a fundamental tool for obtaining crystallization and/or metamorphism ages for a wide range of igneous and/or metamorphic rocks. The combination of these approaches provide the reconstruction and analysis of P-T-t paths, contributing to interpreting the development of orogens with a polycyclic history.

The Araçuaí-West Congo Orogen is one of the various components of the Brasiliano/Pan-African orogenic system generated during the amalgamation of West Gondwana (Trompette 1994, Alkmim et al. 2006). The Araçuaí Orogen (southeastern Brazil) records a long period of granitic magmatism (ca. 150 M.y.) associated to the different stages of the orogeny, with a large production of collisional granitoids (Pedrosa-Soares et al. 2001, 2011, Gradim et al. 2014). In this study, we present new petrographic evidence, geochronological data and thermobarometric estimates for samples from the Carlos Chagas Suite collisional granites, which preserve mineral assemblages of two high-grade metamorphic events (M₁ and M₂). Calculated phase diagrams in combination with conventional geothermobarometry and geochronological methods are used to investigate the ages and metamorphism conditions. These data provide important constraints on the P-T-t paths of these rocks and contribute to the understanding of the crustal recycling in Neoproterozoic orogens.

Geological setting

The Araçuaí Orogen, located in southeastern Brazil, represents the largest portion of the Brasiliano/Pan-African Araçuaí-West Congo Orogenic System, developed between the São Francisco and Congo cratons during the amalgamation of West Gondwana in late Proterozoic time (Alkmim et al. 2006, Pedrosa-Soares et al. 2011) (Figure 1a). Rocks exposed in Espírito Santo, Bahia and eastern of Minas Gerais states contain a record of the evolution of an active continental margin, Rio Doce Magmatic Arc-related rocks and syn- to post-collisional magmatism from the Araçuaí Orogen (De Campos et al. 2004, 2016, Alkmim et al. 2006, Pedrosa-Soares et al. 2011, Peixoto et al. 2015, Richter et al. 2016, Tedeschi et al. 2016, Melo et al. 2017 a-b, 2020, Serrano et al. 2018, Araújo et al. 2020, Soares et al. 2020) (Figure 1b). However, this subduction-collision model has been subject of important discussions (e.g., Cavalcante et al. 2019, Fossen et al. 2020). The main problems presented by Fossen et al. (2020) for this model are: (i) lack evidence of high-P metamorphism; (ii) and the abrupt termination of an ocean with no realistic way to transfer the large amount of oceanic opening displacement and subsequent convergence required by the model. The intracontinental hot orogen model has been proposed by several authors as an alternative model for the Araçuaí Orogen (Vauchez et al. 2007, 2019, Cavalcante et al. 2013, 2014, 2019, Fossen et al. 2020).

Figure 1
Gondwana scenario during amalgamation of the Araçuaí-West Congo belt. (a) Araçuaí-West Congo orogenic system and adjacent cratonic region in the context of the West Gondwana (after Alkmim et al. 2006). (b) Simplified geological map of the Araçuaí orogen (modified from Pedrosa-Soares et al. 2011), showing location of the study area (dashed rectangle).

Despite the differences, both models show vast amounts of granites that range in crystallization age from 630 to 480 M.y. The first granitoids were emplaced around 630 Ma and they were formed by crustal anatexis with minor involvement of mantle magmas (Pedrosa-Soares et al. 2011, Tedeschi et al. 2016). Large amounts of peraluminous melts were produced by partial melting of metasedimentary rocks. Previous geochronological data yield of crystallization age between 600 and 540 M.y. (Pedrosa-Soares et al. 2011, Gradim et al. 2014, Melo et al. 2017 a-b, Cavalcante et al. 2019, Araújo et al. 2020). Paragneisses (Nova Venécia and Jequitinhonha complexes) show widespread evidence of partial melting and they are considered as main source of peraluminous melts (Gradim et al. 2014, Pedrosa-Soares et al. 2011, Araújo et al. 2020). Geothermobarometric estimates from different techniques indicate peak metamorphic conditions of 712-930 °C and 5.0-7.5 kbar (Munhá et al. 2005, Moraes et al. 2015, Richter et al. 2016). Emplacement of widespread, post-collisional intrusions occurred from 530 to 480 M.y. (De Campos et al. 2004, 2016, Gradim et al. 2014, Melo et al. 2020) after the regional metamorphism that took place at ca. 575-550 M.y. (Gradim et al. 2014, Richter et al. 2016, Melo et al. 2017 a-b).

Carlos Chagas Suite

The Carlos Chagas Suite occurs as a large batholith (~ 14,000 km²), consisting of S-type, peraluminous, leucocratic, whitish gray to yellowish granites. Most of the suite consists of progressively deformed facies, since the weakly foliated granites to strongly stretched ultramylonites, passing through the prevalent foliated granite with well-developed augen structure. Subordinately non-deformed facies with well-preserved igneous fabrics and free from the regional ductile S_n foliation occur (Pedrosa-Soares et al. 2007, 2011).

On the map, the Carlos Chagas Suite is surrounded by the foliated peraluminous granites of the Ataléia suite and Al-rich paragneisses (Jequitinhonha and Nova Venécia complexes) (Figure 1). The contacts of Carlos Chagas Suite with their host rocks vary from intrusive, with the Nova Venécia and Jequitinhonha paragneisses, to gradational or sharp, with the Ataléia foliated granites, and can also be marked by ductile shear zones elsewhere (Pedrosa-Soares et al. 2007, 2011). Xenoliths of the host rocks are common along the borders of the batholith. The Carlos Chagas Suite was intruded by several post-collisional bodies (ca. 530-480 M.y.) during the gravitational collapse of the orogen (De Campos et al. 2004, 2016, Pedrosa-Soares et al. 2011, Gradim et al. 2014, Melo et al. 2020).

Previous geochronological studies (U-Pb zircon) have constrained the age of crystallization of the Carlos Chagas Suite to between ca. 597 and 565 M.y. (Vauchez et al. 2007, Gradim et al. 2014, Melo et al. 2017 a-b, Cavalcante et al. 2019). Recent geochronological investigations of the zircon and monazite texture and U-Pb dating have shown that the Carlos Chagas Suite was recorded by two metamorphic events: ca. 570-550 M.y. and ca. 535-515 M.y. (Melo et al. 2017 a-b). The oldest metamorphic event (M₁) records peak P-T conditions at 790-820 °C and 9.5-10.5 kbar, whilst the youngest event (M₂) was attained at 770 °C and 6.6 kbar (Melo et al. 2017a).

MATERIALS AND METHODS

Mineral compositions were determined using the Cameca SX-50 microprobe located at the Universidade de Brasília with an acceleration voltage of 20kV, beam current of 40nA, and counting times of 15s. Whole-rock chemistry analysis was performed at the Department of Earth Sciences, Stellenbosch University (South Africa), using XRF spectrometry on a Philips 1404 Wavelength Dispersive spectrometer.

Zircons were separated from one rock sample at the Department of Geology (DEGEO), Universidade Federal de Ouro Preto, Brazil. The zircon extraction technique makes use of jaw crusher, milling, manual panning, heavy liquids separation, hand-picking under a binocular microscope and mounting on 25 mm epoxy (SpeciFix) mounts. Mounts were polished to expose the grain centers and imaged under SEM-cathodoluminescence (CL) in a Scanning Electron Microscope hosted at Universidade de São Paulo. U-Pb analyses were carried out using a Thermo-Finnigan Element 2 sector field ICP-MS coupled to a CETAC ultraviolet laser system (LA-SF-ICP-MC) housed at DEGEO. Detailed analytical procedures were those described by Melo et al. (2017b).

Lu-Hf analyses were carried out using a Thermo-Scientific Finnigan Neptune SF-ICP-MS Thermo-Scientific Neptune Plus multi-collector ICP-MS coupled to a Photon Machines 193 Excimer laser ablation system at DEGEO, Universidade Federal de Ouro Preto (Brazil), following the instrumentation and analytical methods described by Gerdes & Zeh (2006) and Moreira et al. (2016).

RESULTS

Field relationships and petrography

The studied area (Figure 2) is included in an important production pole of ornamental rocks, with over hundred active quarries. The Carlos Chagas granites have been exploited for ornamental rocks and these quarries contains excellent exposures for petrologic studies. The rock is holo to leucocratic, coarse-to medium grained and whitish gray in hand specimen, although yellowish varieties occur locally. The predominant texture is granoblastic, although lepidogranoblastic and nematogranoblastic textures also occur (Figure 3a). Sometimes mylonitic features are observed, such as K-feldspar porphyroclasts form eye-shaped augen (Figure 3b) due to the development of pressure shadows during rotation of the crystal. In some outcrops, granites have leucosomes crystallized from anatectic melt, consisting of peritectic garnet-rich quartz-feldspathic patches (up to 30 cm long) that are free of the regional solid-state foliation (Figure 3c). Schlieren, nebulitic and stromatic structures are also observed (Figure 3d).

Figure 2
Simplified geologic map for study area, showing the location of samples (modified from Pedrosa-Soares et al. 2007).

Figure 3
Macroscopic features of studied rocks. (a) Leucocratic granite showing foliation (S_n) defined by orientation of biotite. (b) Sigma shape of K-feldspar grains in granite. (c) Neoformed leucocratic patches (dashed line) with subhedral to euhedral garnet crystals. (d) Stromatic structure that shows leucocratic segregations. (e) The contact of the xenolith with host rock is sharp. Note the presence of garnet-bearing leucosome. (f) Mafic dikes (dashed line) cutting the Carlos Chagas granite.

Migmatitic paragneisses (Nova Venécia Complex) and foliated syn-collisional granites of the Ataléia suite surround and underlie the Carlos Chagas Suite (Figure 2). The contact of the studied granite with its host rocks varies from intrusive to gradational. Xenoliths of the host rocks are locally found (Figure 3e). Post-collisional bodies are intrusive in collisional granites (Figure 3f).

Eleven representative thin sections were used for a detailed petrographic study, including thin sections previously studied by Melo et al. (2017a). The mineralogy consists of K-feldspar (48-67 vol.%), quartz (16-28 vol.%), plagioclase (10-21 vol.%) and minor amount of garnet (3-8 vol.%) and biotite (1-6 vol.%). Accessory phases include sillimanite, hercynite, apatite and ilmenite (~ 2 vol.%). Common alteration products are carbonate and muscovite (1-4 vol.%). On the QAP ternary classification diagram (Streckeisen 1974), the Carlos Chagas Suite is classified mainly as syenogranite (Figure 4).

Figure 4
Modal QAP diagram of Streckeisen (1974), showing the composition of the studied granites.

The studied samples contain three preserved mineral assemblages (A₁ to A₃). The A₁ assemblage consists of course-to medium grained plagioclase, K-feldspar, quartz, garnet, biotite and ilmenite. The A₂ assemblage differs from A₁ assemblage by the presence of sillimanite and by generation of a second type of garnet. The A₃ assemblage is characterized by retrograde phases (biotite and hercynite), which occasionally replace garnet boundaries. The samples show granolepidoblastic texture and occasionally nematoblastic texture (Figure 5a).

Figure 5
Photomicrographs of microstructures observed in the studied granites. (a) Granolepidoblastic and nematoblastic textures showing the orientation of biotite (Bt₂) and sillimanite. (b) Intergrowth myrmekite. (c) Films of quartz in perthitic K-feldspar boundaries in the matrix. (d) Polygonal contacts between plagioclase (Pl) grains. (e) Symplectite intergrowth between biotite and quartz in garnet boundaries. Note quartz-filled pools adjacent to garnet. (f) Quartz films along the grain boundaries of garnet in the matrix. (g) Quartz films along the grain boundaries of K-feldspar in the matrix. (h) Quartz ribbon shows intense grain boundary migration recrystallization. Note ameboid to lobate contacts between quartz and feldspar grains. Mineral abbreviations are after Whitney & Evans (2010).

Anhedral K-feldspar grains are medium-to coarse-grained and often show vermicular quartz hosted in plagioclase denotes myrmekite texture (Figure 5b). Some of them are perthitic with exsolved plagioclase lamellae (Figure 5c). Evidence of ductile deformation includes undulose extinction and recrystallized aggregates. of plagioclase with polygonal contacts (Figure 5d).

Quartz grains are anhedral with variable grain sizes, and sometimes concentrate as polygonal aggregates. Symplectic intergrowth involving biotite and quartz are common (Figure 5e). Thin films of quartz are observed along the boundaries of garnet crystal (Figure 5f) and between plagioclase and K-feldspar grains in the matrix (Figure 5g). Ribbons formed by quartz grains in mylonitic samples are common (Figure 5h). The quartz grains commonly exhibit undulose extinction (Figure 6a), subgrains and new grains. Some plagioclase grains exhibit undulose extinction (Figure 6b) and subgrain boundaries. Microfractures filled with quartz are observed cutting K-feldspar grains (Figure 6c).

Figure 6
Photomicrographs of microstructures observed in the studied granites. (a) Evidence of deformation in quartz grains include undulose extinction, subgrain formation and grain boundary migration. (b) Recrystallized plagioclase by subgrain rotation. (c) Microfracture filled by quartz. (d) Garnet poikiloblastic (Grt₁) host quartz, biotite (Bt₁) and ilmenite inclusions. (e) Sillimanite needles wrap around Grt₁, forming a second stage of garnet growth (Grt₂). (f) Garnet poikiloblastic (Grt₂) showing inclusions of sillimanite needles. (g) Hercynite (Spl) inclusions in prismatic sillimanite associated with biotite flakes (Bt₂). (h) Hercynite partially replaces the garnet boundaries.

Two textural varieties of garnet porphyroblasts are observed (Grt₁ e Grt₂). The first variety occurs as subhedral to anhedral poikiloblastic crystals that contain lobate to rounded biotite, quartz and ilmenite (Figure 6d). A second type of garnet commonly occurs as overgrowths in Grt₁ porphyroblasts (Figure 6e) and as euhedral to anhedral crystals, both with inclusions of sillimanite needles (Figure 6f).

Biotite occurs both in the matrix and as inclusions in garnet crystals. Those in the matrix (Bt₂) are brown, aligned and fine-grained, defining the foliation (Figure 5a). Rounded biotite inclusions (Bt₁) are reddish brown in colour (Figures 6 d-e), implying high Ti contents. In local domains of retrograde re-equilibrium, garnet is partially replaced by biotite (Bt₃) (Figures 5e and 6d).

Oriented aggregates of prismatic sillimanite are associated with biotite, whilst sillimanite needles (fibrolite) occur as inclusions in plagioclase and Grt₂. Hercynite occurs with biotite and sillimanite in the matrix or partially replacing garnet boundaries. Rare hercynite occurs as inclusions within garnet porphyroblasts and prismatic sillimanite (Figures 6 g-h).

Mineral chemistry

Representative mineral analyses are presented in Tables I to V. Both generations of garnet in studied rocks are almandine-rich (Figure 7a) and individual crystals do not display any significant zoning (Figure 8). The Grt₁ crystals composition is similar in all samples, with X_Alm values of 0.74-0.80 (CCS-21) and 0.77-0.82 (CCS-25), X_Prp values of 0.15-0.21 (CCS-21) and 0.13-0.19 (CCS-25). Both samples have low X_sps (0.02-0.03) and X_Grs (0.02-0.03) contents and Mg# of 0.16-0.22 (CCS-21) and 0.13-0.20 (CCS-25). Typical compositional ranges for Grt₂ crystals in the samples are X_Alm 0.73-0.82, X_Prp 0.13-0.21, X_Grs 0.02-0.03, X_Sps 0.03-0.04 (CCS-21) and X_Alm 0.76-0.82, X_Prp 0.13-0.18, X_Grs 0.02-0.04, X_Sps 0.02-0.03 (CCS-25). Mg# values range between 0.13 and 0.22 (CCS-21) and 0.13-0.19 (CCS-25).

Table I
Average (Av) and standard deviation (SD) of the mineral chemistry of garnets. Calculations based on 12 oxygens, following the methodology of Droop (1987).

	CCS-21				CCS-25
	Grt1 (n=18)		Grt2 (n=13)		Grt1 (n=19)		Grt2 (n=20)
	Av	SD	Av	SD	Av	SD	Av	SD
SiO2	36.08	0.30	36.46	0.17	36.42	0.22	36.32	0.25
TiO2	0.01	0.02	0.01	0.02	0.01	0.02	0.01	0.02
Al2O3	21.82	0.18	22.0	0.24	22.13	0.19	22.06	0.20
Cr2O3	0.01	0.02	0.01	0.01	0.01	0.01	0.01	0.02
FeO	35.25	0.86	34.68	1.03	34.65	0.67	35.08	0.81
MnO	1.23	0.11	1.41	0.10	1.12	0.11	1.08	0.06
MgO	4.31	0.53	4.52	0.57	4.19	0.34	3.84	0.40
CaO	0.80	0.11	0.87	0.12	0.92	0.08	1.01	0.21
Total	99.50	0.57	99.95	0.49	99.45	0.37	99.40	0.44
Si	2.90	0.02	2.91	0.02	2.92	0.02	2.92	0.01
Ti	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
Al	2.06	0.01	2.07	0.02	2.09	0.01	2.09	0.01
Cr	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
Fe3+	0.14	0.04	0.11	0.03	0.06	0.04	0.06	0.03
Fe2+	2.23	0.06	2.20	0.09	2.27	0.04	2.30	0.06
Mn	0.08	0.01	0.09	0.01	0.08	0.01	0.07	0.00
Mg	0.51	0.06	0.54	0.07	0.50	0.04	0.46	0.05
Ca	0.07	0.01	0.07	0.01	0.08	0.01	0.09	0.02
XAlm	0.76	0.02	0.75	0.02	0.77	0.01	0.78	0.02
XSps	0.03	0.00	0.03	0.00	0.02	0.00	0.02	0.00
XPrp	0.18	0.02	0.18	0.02	0.17	0.01	0.16	0.01
XGrs	0.02	0.00	0.02	0.00	0.03	0.00	0.03	0.00
XAdr	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
XUv	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00

Table II
Average (Av) and standard deviation (SD) of the mineral chemistry of feldspars in matrix. Calculations based on 8 oxygens, according to parameters of Deer et al. (2020).

	CCS-21				CCS-25
	Pl (n=13)		Kfs (n=9)		Pl (n=32)		Kfs (n=8)
	Av	SD	Av	SD	Av	Sd	Av	Sd
SiO2	63.69	10.14	64.04	0.77	60.01	1.26	63.91	0.38
TiO2	0.00	0.01	0.01	0.01	0.00	0.01	0.01	0.02
Al2O3	23.06	6.92	19.56	0.18	25.76	0.71	19.76	0.18
FeO	0.04	0.02	0.04	0.09	0.03	0.03	0.04	0.07
MnO	0.02	0.02	0.01	0.02	0.02	0.02	0.04	0.03
MgO	0.00	0.00	0.00	0.01	0.00	0.01	0.01	0.01
CaO	4.71	1.47	0.14	0.09	5.67	0.36	0.12	0.05
Na2O	8.21	2.46	2.65	1.72	8.54	0.24	1.61	0.32
K2O	0.23	0.10	13.18	2.33	0.23	0.06	14.51	0.60
BaO	0.03	0.05	0.06	0.06	0.03	0.05	0.17	0.07
ZnO	0.02	0.04	0.04	0.06	0.03	0.05	0.01	0.03
Total	100.00	0.75	99.74	0.79	100.34	0.95	100.20	0.38
Si	2.80	0.36	2.95	0.01	2.66	0.03	2.94	0.01
Al	1.21	0.36	1.06	0.01	1.35	0.05	1.07	0.01
Ti	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
Fe	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
Mn	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
Mg	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
Zn	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
Ca	0.22	0.07	0.01	0.00	0.27	0.02	0.01	0.00
Na	0.71	0.21	0.23	0.15	0.73	0.02	0.14	0.03
K	0.01	0.01	0.77	0.14	0.01	0.00	0.85	0.04
Ba	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
XOr	0.04	0.06	0.76	0.15	0.01	0.00	0.85	0.03
XAb	0.74	0.02	0.23	0.15	0.73	0.01	0.14	0.03
XAn	0.22	0.05	0.01	0.40	0.26	0.01	0.01	0.27

Table III
Average (Av) and standard deviation (SD) of the mineral chemistry of biotite. Calculations based on 20 oxygens, following the methodology of Tindle & Webb (1990).

	CCS-21						CCS-25
	Bt1 (n=3)		Bt2 (n=9)		Bt3 (n=17)		Bt1 (n=4)		Bt2 (n=15)		Bt3 (n=10)
	Av	SD	Av	SD	Av	SD	Av	SD	Av	SD	Av	SD
SiO2	36.16	0.67	35.06	0.20	35.63	0.30	34.83	0.60	34.59	0.26	34.67	0.24
TiO2	4.13	0.89	4.59	0.24	3.46	1.03	3.69	0.68	4.00	0.30	4.00	0.39
Al2O3	16.60	0.10	16.49	0.23	17.21	1.21	16.37	0.13	16.85	0.25	17.26	0.20
FeO	13.64	1.94	18.36	0.31	17.42	5.02	16.73	2.85	19.04	0.35	18.29	1.09
MnO	0.01	0.02	0.03	0.03	0.10	0.30	0.00	0.01	0.04	0.03	0.03	0.03
MgO	14.28	1.32	10.67	0.30	11.83	2.22	11.70	2.34	9.77	0.33	10.20	0.70
CaO	0.02	0.01	0.01	0.01	0.05	0.14	0.01	0.01	0.01	0.01	0.02	0.02
Na2O	0.40	0.20	0.15	0.04	0.21	0.11	0.15	0.06	0.17	0.04	0.22	0.08
K2O	9.34	0.15	9.47	0.13	8.95	2.19	9.44	0.06	9.37	0.11	9.46	0.14
SrO	0.00	0.00	0.02	0.03	0.03	0.04	0.00	0.00	0.04	0.03	0.03	0.02
BaO	0.07	0.06	0.05	0.04	0.05	0.06	0.11	0.10	0.07	0.05	0.07	0.09
ZnO	0.09	0.04	0.05	0.07	0.12	0.11	0.05	0.06	0.09	0.08	0.15	0.12
F	2.50	0.11	1.88	0.18	2.20	0.59	2.79	0.56	2.27	0.12	2.30	0.07
Cl	0.24	0.04	2.06	5.23	0.22	0.07	0.31	0.07	0.37	0.02	0.35	0.01
Cr2O3	0.03	0.01	0.03	0.03	0.01	0.02	0.01	0.01	0.03	0.03	0.02	0.02
NiO	0.01	0.02	0.02	0.03	0.03	0.03	0.04	0.04	0.02	0.04	0.03	0.04
Total	100.04	1.32	100.84	3.40	100.12	1.35	97.91	0.52	98.78	0.46	99.20	0.32
Si	5.35	0.02	5.33	0.02	5.34	0.03	5.37	0.04	5.34	0.02	5.31	0.03
Aliv	2.65	0.02	2.67	0.02	2.66	0.03	2.63	0.04	2.66	0.02	2.69	0.03
Alvi	0.24	0.09	0.28	0.03	0.39	0.18	0.34	0.01	0.40	0.04	0.42	0.05
Ti	0.46	0.09	0.52	0.03	0.39	0.12	0.43	0.08	0.46	0.04	0.46	0.04
Cr	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
Fe	1.69	0.28	2.33	0.04	2.18	0.61	2.16	0.39	2.46	0.05	2.34	0.15
Mn	0.00	0.00	0.00	0.00	0.01	0.04	0.00	0.00	0.01	0.00	0.00	0.00
Mg	3.15	0.22	2.42	0.07	2.65	0.50	2.68	0.51	2.25	0.07	2.33	0.14
Zn	0.01	0.00	0.01	0.01	0.01	0.01	0.01	0.01	0.01	0.01	0.02	0.01
Ni	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
Li	0.49	0.10	0.31	0.03	0.41	0.05	0.27	0.10	0.23	0.05	0.25	0.04
Ca	0.00	0.00	0.00	0.00	0.01	0.02	0.00	0.00	0.00	0.00	0.00	0.00
Na	0.12	0.05	0.05	0.01	0.06	0.03	0.04	0.02	0.05	0.01	0.07	0.02
K	1.76	0.06	1.83	0.02	1.71	0.42	1.85	0.01	1.84	0.02	1.85	0.04
Sr	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.00
Ba	0.00	0.00	0.00	0.00	0.00	0.00	0.01	0.01	0.00	0.00	0.00	0.01
OH	2.77	0.03	2.68	1.01	2.90	0.29	2.56	0.24	2.80	0.06	2.80	0.03
F	1.17	0.03	0.91	0.08	1.05	0.28	1.36	0.26	1.11	0.06	1.11	0.03
Cl	0.06	0.01	0.52	1.33	0.06	0.02	0.08	0.02	0.10	0.01	0.09	0.00
Total	19.93	0.05	19.87	0.33	19.83	0.19	19.80	0.13	19.73	0.05	19.75	0.03

Table IV
Average (Av) and standard deviation (SD) of the mineral chemistry of ilmenite. Calculations based on 3 oxygens, following the methodology of Droop (1987).

	CCS-21 (n=4)		CCS-25 (n=8)
	Av	SD	Av	SD
SiO2	0.01	0.01	0.01	0.01
TiO2	51.26	0.17	52.11	0.53
Al2O3	0.00	0.00	0.02	0.01
Cr2O3	0.01	0.01	0.02	0.02
FeO	47.18	0.56	45.88	0.80
MnO	0.31	0.03	0.22	0.08
MgO	0.61	0.19	0.42	0.22
CaO	0.01	0.01	0.01	0.01
Na2O	0.01	0.01	0.03	0.04
K2O	0.01	0.01	0.01	0.02
F	0.13	0.05	0.09	0.08
Cl	0.00	0.00	0.01	0.01
V2O3	0.35	0.05	0.40	0.04
Totals	99.88	0.31	99.22	0.60
Si	0.000	0.000	0.000	0.000
Ti	0.97	0.00	0.99	0.01
Al	0.00	0.00	0.00	0.00
Cr	0.00	0.00	0.00	0.00
Fe2+	0.92	0.01	0.96	0.02
Fe3+	0.08	0.01	0.02	0.02
Mn	0.01	0.00	0.01	0.00
Mg	0.02	0.01	0.02	0.01
Ca	0.00	0.00	0.00	0.00
Na	0.00	0.00	0.00	0.00
K	0.00	0.00	0.00	0.00
F	0.01	0.00	0.01	0.01
Cl	0.00	0.00	0.00	0.00
V	0.00	0.00	0.00	0.00

Table V
Average (Av) and standard deviation (SD) of the mineral chemistry of spinel. Calculations based on 4 oxygens, following the methodology of Droop (1987).

	CCS-21 (n=8)
	Av	SD
SiO2	0.01	0.01
TiO2	0.01	0.01
Al2O3	59.14	0.31
Cr2O3	0.10	0.08
FeO	28.92	1.93
MnO	0.10	0.03
MgO	3.76	0.32
CaO	0.02	0.02
Na2O	0.32	0.08
K2O	0.00	0.01
F	0.05	0.07
Cl	0.00	0.00
ZnO	7.56	2.33
Totals	99.99	0.46
Si	0.00	0.00
Ti	0.00	0.00
Al	1.97	0.01
Cr	0.00	0.00
Fe2+	0.64	0.04
Fe3+	0.05	0.01
Mn	0.00	0.00
Mg	0.16	0.01
Ca	0.00	0.00
Na	0.02	0.00
K	0.00	0.00
F	0.00	0.01
Cl	0.00	0.00
Zn	0.16	0.05

Figure 7
Mineral chemistry of the studied rocks. (a) Alm-Prp-Grs garnet ternary diagram. (b) Mica classification according to Tieschendorf et al. (1997). (c) Biotite compositions plotted on a Ti (a.p.f.u.) vs. Mg#. (d) An-Ab-Or feldspar classification diagram (after Deer et al. 2010). (e) TiO₂-FeO-Fe₂O₃ ternary diagram (after Deer et al. 2010).

Figure 8
Chemical profile in garnets showing content of X_Alm, X_Prp and X_Sps in section along the grains.

According to the binary diagram of Tischendorf et al. (1997) (Figure 7b), most of the biotite analyses plot into the domain of the Mg-biotite. Biotite inclusions (Bt₁) in garnet show Mg# values of 0.59-0.67 (CCS-21) and 0.49-0.66 (CCS-25) and Ti values of 0.37-0.55 (CCS-21) and 0.34-0.50 (CCS-25). Two generations of biotite exist in the matrix of the granites: a predominant peak assemblage biotite (Bt₂) characterized by Mg# values of 0.50-0.53 (CCS-21) and 0.46–0.49 (CCS-25) and Ti values of 0.46-0.56 (CCS-21) and 0.41-0.42 (CCS-25); and a retrograde biotite (Bt₃) that partially replaced garnet boundaries, with Mg# of 0.54-0.62 (CCS-21) and 0.48-0.50 (CCS-25) and Ti values of 0.37-0.57 (CCS-21) and 0.43-0.57 (Figure 7c). Ti in the biotite inclusions is slightly higher than Ti in the biotite of the matrix but varies from sample to sample (Figure 7c).

Plagioclase of the matrix is unzoned and displays X_an values of 0.19–0.25 (CCS-21) and 0.24-0.28 (CCS-25) (Figure 7d). Matrix K-feldspar crystals have low anorthite content (~ 0.01) and variable X_Or values of 0.56-0.93 (CCS-21) and 0.72-0.91 (CCS-25) (Figure 7d). Ilmenite grains are classified as Fe-Ti ilmenite (Figure 7e), with Fe²⁺ contents of 0.91-0.92 apfu (CCS-21) and 0.92-0.9 apfu (CCS-25) and Fe³⁺ values of 0.06-0.08 apfu (CCS-21) and 0.01-0.04 apfu (CCS-25) Spinel crystals have relatively high Zn-content with X_Ghn of 0.10-0.22 (CCS-21).

Conventional geothermobarometer

In order to establish the path of metamorphism, temperatures were calculated by using the garnet-biotite geothermometer. Peak metamorphic conditions (M₁) were estimated using garnet core compositions (Grt₁) together with inclusions of biotite (Bt₁). For peak metamorphic temperature estimates of the M₂ event, garnet core compositions (Grt₂) and matrix biotite (Bt₂) were used. Rim compositions of garnet and biotite (Bt₃) give rough estimates of temperatures of the retrograde reactions associate to the M₂ event. Considering pressures of 7-10 kbar, estimates geothermometric indicates equilibrium conditions at 633-643 °C (CCS-21) and 736-746 °C for the M₁ event (Table VI). Temperatures around 757-775 °C (CCS-21) and 718-735 °C (CCS-25) were attained during the M₂ event (Table VI). Estimate pressures of the peak metamorphic of the M₂ event have been determined using the garnet-sillimanite-quartz-plagioclase (GASP) geobarometer. Calculated peak pressures are around 5.4-8.5 kbar (CCS-21) and 5.2-8.2 kbar (CCS-25) at temperatures between 700 and 900 °C (Table VI). Retrograde re-equilibrium indicates lower temperature of about 600-609 °C (CCS-21) and 665-664 °C (CCS-25) at pressures of 4-6 kbar (Table VI).

Table VI
Temperature and pressure estimates obtained by conventional geothermobarometry.

Temperature (°C) for A1
Thermometer	Calibration	CCS-21			CCS-25
		7 Kbar	9 Kbar	11 Kbar	7 Kbar	9 Kbar	11 Kbar
Grt-Bt	Thompson (1976)	629	643	658	692	709	723
	Ferry & Spear (1978)	-	-	-	694	703	711
	Hodges & Spear (1982)	622	629	637	732	741	749
	Pigage & Greenwood (1982)	643	651	658	777	786	794
	Williams & Grambling (1990)	640	648	656	785	794	803
Average	-	633	643	653	736	746	756
Temperature (°C) for A 2
Thermometer	Calibration	CCS-21			CCS-25
		5 Kbar	7 Kbar	9 Kbar	5 Kbar	7 Kbar	9 Kbar
Grt-Bt	Thompson (1976)	696	712	728	667	683	698
	Ferry & Spear (1978)	712	721	730	673	681	689
	Hodges & Spear (1982)	723	732	741	685	694	702
	Pigage & Greenwood (1982)	753	762	770	708	717	725
	Williams & Grambling (1990)	757	766	775	718	726	735
Average	-	738	739	749	690	700	710
Pressures (Kbar) for the A 2
Barometer	Calibration	CCS-21			CCS-25
		500°C	700°C	900°C	500°C	700°C	900°C
Grt-Sil-Qz-Pl	Newton & Haselton (1981)	1.7	5.2	8.7	1.5	4.9	8.4
	Hodges & Spear (1982)	1.9	4.1	6.1	1.8	3.9	6.0
	Hodges & Crowley (1985)	2.3	5.1	7.9	2.2	5.0	7.8
	Koziol & Newton (1988)	2.7	6.0	9.2	2.6	5.8	9.0
	Koziol (1989)	3.3	6.8	10.4	2.9	6.5	10.0
Average	-	2.4	5.4	8.5	2.2	5.2	8.2
Temperatures (°C) for the A3
Thermometer	Calibration	CCS-21			CCS-25
		4 Kbar	6 Kbar	8 Kbar	4 Kbar	6 Kbar	8 Kbar
Grt-Bt	Thompson (1976)	594	608	623	637	652	667
	Ferry & Spear (1978)	582	589	597	638	647	655
	Hodges & Spear (1982)	593	601	608	650	658	666
	Pigage & Greenwood (1982)	618	525	633	670	678	687
	Williams & Grambling (1990)	614	622	630	678	687	695
Average		600	609	618	655	664	674

Phase equilibrium modeling

Phase equilibria modelling was undertaken in the chemical system MnO-Na₂O-CaO-K₂OFeO-MgO-Al₂O₃-SiO₂-H₂O-TiO₂-O₂ (MnNCKFMASHTO) using Theriak-Domino software (De Capitani & Petrakakis 2010), in combination with the updated Holland & Powell (1998) data set that includes silicate melt. The following a-x models were used: feldspar (Holland & Powell 2003); garnet and biotite (White et al. 2005); orthopyroxene (White et al. 2002); ilmenite (White et al. 2000, 2005); cordierite (Holland & Powell 1998); muscovite (Coggon & Holland 2002); and melt (White et al. 2007).

The investigated samples (CCS-21 and CCS-25) preserve their A₁ assemblage only as isolated porphyroblasts in a foliated matrix that consists of the A₂ assemblage (Kfs+Pl+Qz+Grt₂+Bt₂+Sil+Ilm+melt). This mineral assemblage was used to estimate P-T conditions during the peak metamorphic of the M₂ event. Table VII presents the bulk composition used for metamorphic modeling. H₂O in the bulk composition were set by using T-X pseudosections and the bulk rock H₂O content was chosen based on the methodology of White & Powell (2002), such that the solidus was located just below the lowest temperature of the interpreted peak assemblage. The H₂O values were of 0.14 wt.% (CCS-21) and 0.18 wt.% (CCS-25). The Fe₂O₃ content were determined using P-X pseudosection and the values chosen for Fe³⁺ concentration in each bulk composition (CCS-21 = 0.05 wt.% and CCS-25 = 0.09 wt.%) results in isopleths of Fe³⁺ in garnet, within the field of the preserved assemblage.

Table VII
Bulk composition used for metamorphic modelling.

Sample	Al2O3	CaO	FeO	Fe2O3	K2O	MgO	MnO	Na2O	SiO2	TiO2	H2O	Total
CC21	13.42	0.92	1.83	0.05	5.89	0.31	0.05	2.48	73.89	0.20	0.14	99.18
CC25	13.65	1.02	2.43	0.09	5.86	0.43	0.07	2.37	72.84	0.26	0.18	99.20

The peak metamorphic assemblage of the M₂ event is stable over a wide range of pressures (4.2 to >8 kbar) and relatively restricted temperatures (690 to 760 ˚C) (Figure 9a). The stability of biotite places an upper temperature constraint of < 760 °C (Figure 9a). The lower pressure limit of the A₂ assemblage is bounded by stability of cordierite. The absence of A₂ cordierite constrains pressures to < 4.2 kbar at 700 °C and < 5.3 kbar at > 700 °C (Figure 9a). X_An values for modeled plagioclase range from 0.24 to 0.28 in in the preserved assemblage field (Figure 9b). Modeled garnet composition shows Mg# of 0.11–0.25, X_Prp of 0.10–0.23, X_Alm of 0.70-0.81 and X_Sps of 0.02–0.05 in the preserved assemblage field (Figures 9 c-d-e). The mode of melt at preserved metamorphic assemblage is less than 3 vol% (Figure 9f). These values encompass those of measured in plagioclase and Grt₂, showing X_An, Grt Mg#, X_Prp, X_Alm and X_Sps values of 0.19-0.25, 0.13-0.22, 0.13-0.21, 0.73-0.82 and 0.029-0.04, respectively. The use of X_An, Grt Mg#, X_Prp, X_Alm and X_Sps values help to more tightly constrain the P–T conditions of the preserved (A₂) assemblage to 713–719 °C and 6.4-7.1 kbar (Figure 9f).

Figure 9
P–T pseudosection calculated using the bulk composition from sample CCS-21 (A₂ assemblage). (a) P–T fields showing the preserved assemblage Kfs+Pl+Qz+Grt₂+Bt₂+Sil+Ilm+melt (bold line). (b) Plots of X_An. (c) Plots of Mg# in garnet. (d) Plots of X_Prp. (e) Plots of X_Alm and X_Sps. (f) P–T conditions of equilibrium based on the use of overlapping ranges of phase composition.

In sample CCS-25, pseudosection modelling constrains the preserved (A₂) metamorphic assemblage of Kfs+Pl+Qz+Grt₂+Bt₂+Sil+Ilm+melt to a field of stability between 700-760 ˚C and pressures ranging from 4.3 to >8 kbar (Figure 10a). These P-T conditions are like those of the sample CCS-21 (Figure 10a). Modeled plagioclase, biotite, garnet compositions show X_An of 0.28-0.31, Grt Mg# of 0.12-0.26, X_Prp of 0.10–0.23, X_Alm of 0.69-0.80 and X_Sps of 0.02–0.04 in the preserved assemblage field (Figures 10 b-c-d-e). Within this field, the mode of melt is less than 3 vol% (Figure 10f). Plagioclase, biotite and garnet compositional isopleths show that the measured compositions these mineral phases intersect the preserved A₂ assemblage field between 719 ˚C and 729 °C and 6.1- 7.3 kbar (Figure 10f).

Figure 10
P–T pseudosection calculated using the bulk composition from sample CCS-25 (A₂ assemblage). (a) P–T fields showing the preserved assemblage Kfs+Pl+Qz+Grt₂+Bt₂+Sil+Ilm+melt (bold line). (b) Plots of X_An and Mg# in biotite. (c) Plots of Mg# in garnet. (d) Plots of X_Prp. (e) Plots of X_Alm and X_Sps. (f) P–T conditions of equilibrium based on the use of overlapping ranges of phase composition.

Zircon U-Pb and Lu-Hf data

In order to evaluate the metamorphic history from Carlos Chagas Suite, we investigated the internal structure of the zircon grains (sample CCS-35a) through cathodoluminescence (CL) imaging (Figure 11). Spot analyses were carried on zircon core and/or rim domains via LA-SF-ICP-MS. The analytical results are presented in Table VIII and ages presented in figure 12.

Figure 11
Representative cathodoluminescence (CL) images showing LA-ICP-MS spots for Hf isotopes. U-Pb error reported as 2σ.

Figure 12
U-Pb ages and Hf isotopic composition for the Carlos Chagas Suite. (a and b) Concordia diagrams and weighted mean ²⁰⁶Pb/²³⁸U age plots containing the results of zircon U-Pb dating of the sample CCS-35b. (c) Initial ¹⁷⁶Hf/¹⁷⁷Hf ratios of zircon populations from the studied sample plotted against their age (Ma). The error bars are at 2σ level. (d) εHf(_t) vs. age diagrams for zircon grains this study. Data compiled from Gonçalves et al. (2016), Tedeschi et al. (2016), Melo et al. (2017b) and Araújo et al. (2020).

Table VIII
U-Pb data determined by LA-ICP-MS.

	ppm			Ratios						Ages (M.y.)						Rho	Conc
Spot	206Pb	238U	Th/ U	207Pb/ 206Pb	± 2σ	206Pb/ 238U	± 2σ	207Pb/ 235U	± 2σ	207Pb/ 206Pb	± 2σ	206Pb/ 238U	± 2σ	207Pb/ 235U	± 2σ
Group I
34	22	139	0.16	0.0600	1.7367	0.0981	2.6423	0.8115	3.1619	567	47	604	8	597	8	0.84	101
59	9	76	0.17	0.0598	1.3021	0.0983	2.6040	0.8100	2.9113	573	75	604	9	598	11	0.89	101
73	47	63	0.30	0.0601	1.1328	0.1001	2.5732	0.8301	2.8115	579	56	615	9	608	9	0.92	101
74	13	70	0.98	0.0602	1.6922	0.1000	2.6983	0.8298	3.1850	579	55	614	8	607	9	0.85	101
76	9	45	0.27	0.0601	1.8307	0.0996	2.7231	0.8252	3.2813	586	71	612	9	607	11	0.83	101
87	10	51	0.28	0.0598	1.6314	0.0984	2.9976	0.8109	3.4128	576	69	605	9	599	11	0.88	101
88	58	94	0.25	0.0604	1.2322	0.1027	1.8151	0.8555	2.1939	762	117	631	10	660	19	0.83	95
Group II
9	57	153	0.09	0.0595	1.5846	0.0942	1.9470	0.7752	2.5038	550	49	582	7	575	8	0.78	101
10	23	100	0.77	0.0595	1.6679	0.0956	1.9384	0.7856	2.5570	624	79	589	8	597	12	0.76	99
11	37	241	0.15	0.0596	1.9130	0.0948	2.2431	0.7772	2.9533	551	44	583	7	576	7	0.76	101
24	29	143	0.14	0.0594	1.4315	0.0952	2.6597	0.7794	3.0216	560	61	585	7	580	9	0.88	101
26	32	195	1.11	0.0597	1.5948	0.0967	2.6319	0.7964	3.0772	567	46	593	7	587	7	0.86	101
27	51	92	0.30	0.0591	2.1254	0.0941	2.6417	0.7737	3.3908	546	58	581	8	574	9	0.78	101
29	35	125	0.41	0.0594	1.5846	0.0953	2.0374	0.7790	2.5805	553	52	585	7	579	8	0.79	101
41	28	117	0.20	0.0594	1.5706	0.0950	3.1107	0.7786	3.4835	555	51	585	8	579	8	0.89	101
49	34	411	0.32	0.0597	1.5643	0.0953	1.8930	0.7830	2.4557	552	31	585	6	579	5	0.77	101
52	18	319	0.65	0.0593	1.5837	0.0949	1.9412	0.7756	2.5040	551	35	582	6	576	6	0.78	101
61	15	60	1.06	0.0582	1.7091	0.0956	1.5104	0.7764	2.1300	560	61	589	9	583	9	0.71	101
64	24	54	0.43	0.0598	1.6077	0.0964	1.7352	0.7945	2.3654	560	67	592	8	585	10	0.73	101
83	36	56	0.34	0.0596	1.7193	0.0952	2.1891	0.7823	2.7838	555	58	587	9	581	9	0.79	101
Group III
3	25	118	0.28	0.0582	1.6583	0.0916	2.7752	0.7444	3.2226	537	57	564	7	558	8	0.86	101
5	58	739	0.25	0.0590	1.6525	0.0916	2.7702	0.7443	3.2283	551	37	564	6	561	6	0.86	100
8	37	142	0.13	0.0589	1.3692	0.0905	1.9914	0.7315	2.4298	532	55	557	7	552	8	0.82	101
43	29	119	0.19	0.0589	1.4753	0.0907	2.1785	0.7306	2.6336	529	56	558	7	552	8	0.83	101
44	18	166	0.29	0.0512	1.9455	0.0907	2.8621	0.7355	3.4607	528	43	560	7	554	7	0.83	101
45	16	206	0.31	0.0576	1.6546	0.0918	2.7731	0.7442	3.2285	538	40	564	6	559	6	0.86	101
46	12	185	0.12	0.0598	1.8609	0.0914	2.1659	0.7431	2.8553	533	44	563	7	557	7	0.76	101
82	27	63	0.37	0.0581	1.0621	0.0910	2.0970	0.7333	2.3460	523	58	558	10	551	9	0.89	101
84	21	95	0.14	0.0587	2.1770	0.0911	2.9625	0.7459	3.6821	538	51	566	11	560	9	0.80	101
Group IV
2	24	281	0.08	0.0578	1.3320	0.0841	3.4408	0.6699	3.6896	501	54	521	7	518	8	0.93	101
6	14	377	0.65	0.0581	1.2789	0.0865	2.7388	0.6930	3.0222	611	80	531	7	536	11	0.91	99
13	28	589	0.04	0.0579	1.2908	0.0852	3.3066	0.6801	3.5496	499	29	527	6	522	5	0.93	101
17	16	273	0.42	0.0590	1.2779	0.0876	2.7388	0.6930	3.0222	508	100	531	8	528	13	0.91	101
36	25	215	0.22	0.0581	1.5561	0.0854	2.1445	0.6842	3.6496	501	44	529	7	523	7	0.59	101
38	17	157	0.14	0.0577	1.0799	0.0844	2.8735	0.6716	3.1625	497	53	523	7	518	8	0.91	101
68	32	145	0.50	0.0577	1.1159	0.0837	3.2513	0.6658	3.4375	493	75	520	9	518	10	0.95	100

Zircon grains are subhedral to euhedral prismatic, light pink, translucent, ranging from 80 to 500 µm, with length/width ratios in average of 4:1. CL images of the internal structure of the zircons show that they have a core with oscillatory-zoned patterns or irregular zoning, which are generally surrounded by low luminescence and structureless rim (Figure 11). Some rounded and unzoned cores appear luminescent in CL images. Thirty-seven analytical spots yield a range of concordant ages that spread along the Concordia diagram from ca. 631 to 520 M.y. (Figure 12a). Four populations have been distinguished for zircon ²⁰⁶Pb/²³⁸U ages (groups I to IV). The oldest population (Group I; n = 8) is represented by rounded cores with bright luminescence with ages ranging from ca. 631 to 604 M.y. and Th/U ratios of 0.16-0.98 (Table VIII). Other thirteen spot analyses in core domains (Group II) give weighted mean ²⁰⁶Pb/²³⁸U age of 586 ± 2 M.y. (MSWD = 1.02; Figures 12 a-b), with Th/U ratio ranging from 0.09 to 1.11 (Figure 12a). Nine spot analyses in rims (overgrowths) and unzoned domains (Group III) yield a weighted mean of 562 ± 2 M.y. (MSWD = 0.72; Figures 12 a-b) and Th/U ratio between 0.12 and 0.37 (Table VIII). Other seven spot analyses in zircon rims (Group IV) give a weighted mean of 526 ± 4 M.y. (MSWD = 1.5; Figures 12 a-b) and Th/U ratio between 0.04 and 0.65 (Table VIII).

Five analyses on the oldest cores (Group I) yield initial ¹⁷⁶Hf/¹⁷⁷Hf ratios ranging from 0.28227 to 0.28229, with ɛHf(t) values between -4.3 and -5.1. The distribution of model age (T_DM) spans from 1490 to 1530 M.y. (Figures 12 c-d and Table IX). Six analyses performed on zircon cores (Group II) gives initial ¹⁷⁶Hf/¹⁷⁷Hf ratios ranging from 0.28226 to 0.28230, ɛHf(t) between -5.9 and -4.5 and T_DM age varying from 1490 to 1540 M.y. (Figure 12c and Table IX). Three analyses performed on unzoned rims (Group III) yield initial ¹⁷⁶Hf/¹⁷⁷Hf ratios between 0.28227 and 0.28233, with ɛ_Hf(_t) values ranging from -6.0 to -4.1 and Hf T_DM model ages from 1440 to 1540 M.y. (Figures 12 c-d and Table IX). Seven analyses performed on youngest zircon rims (Group IV) gives initial ¹⁷⁶Hf/¹⁷⁷Hf ratios ranging from 0.28225 to 0.28232 and ɛ_Hf(_t) values between -7.1 and -4.8 and Hf T_DM model ages from 1460 to 1560 M.y. (Figures 12 c-d and Table IX).

Table IX
Hf isotopes results determined by LA-ICP-MS.

Spot	176Yb/ 177Hf	±2σ	176Lu/ 177Hf	±2σ	178Hf/ 177Hf	180Hf/ 177Hf	Sig Hf	176Hf/ 177Hf	±2σ	176Hf/ 177Hf(t)	ƐHf (t)	±2σ	T DM	age	±2σ
Group I
59	0.01959	15.78209	0.00069	4.18489	1.46726	1.88671	10.90207	0.28227	10.60610	0.28226	-5.1	0.6	1.53	604	2.6040
73	0.04165	33.45867	0.00134	8.04034	1.46730	1.88568	14.64688	0.28228	13.53183	0.28226	-4.8	0.5	1.52	615	2.5732
74	0.03176	25.63953	0.00112	6.70399	1.46718	1.88687	14.71458	0.28227	14.07013	0.28226	-5.0	0.7	1.53	614	2.6983
76	0.04003	32.41101	0.00141	8.45557	1.46729	1.88635	13.95330	0.28229	13.80431	0.28228	-4.3	0.7	1.49	612	2.7231
87	0.03424	27.70324	0.00120	7.17500	1.46731	1.88583	10.49419	0.28229	12.70900	0.28228	-4.4	0.9	1.49	605	2.9976
Group II
9	0.01585	12.70560	0.00053	3.16601	1.46721	1.88653	17.36774	0.28226	16.05120	0.28226	-5.7	0.6	1.54	582	1.9470
11	0.03270	26.18538	0.00115	6.92113	1.46724	1.88708	13.65326	0.28230	17.63084	0.28229	-4.5	0.6	1.48	583	2.2431
24	0.01904	15.29406	0.00066	3.99424	1.46735	1.88531	9.34459	0.28229	11.24185	0.28228	-4.7	0.7	1.49	585	2.6597
29	0.04384	35.13823	0.00152	9.19881	1.46726	1.88677	15.45630	0.28227	14.87631	0.28226	-5.9	0.6	1.54	585	2.0374
41	0.04346	35.06024	0.00164	9.85693	1.46723	1.88676	12.03026	0.28228	14.46894	0.28226	-5.8	0.7	1.53	585	3.1107
49	0.04167	39.86563	0.00151	11.92437	1.46725	1.88704	16.83198	0.28229	14.19479	0.28228	-5.3	0.6	1.51	585	1.8930
Group III
5	0.04680	37.53494	0.00175	10.60401	1.46727	1.88562	9.62315	0.28229	13.28910	0.28228	-5.4	0.6	1.51	564	2.7702
44	0.02751	22.03994	0.00098	5.90390	1.46729	1.88474	6.94840	0.28233	12.67070	0.28232	-4.1	1.1	1.44	560	2.8621
84	0.02862	23.06660	0.00101	6.07993	1.46727	1.88665	12.93633	0.28227	12.89225	0.28226	-6.0	0.7	1.54	566	2.9625
Group IV
2	0.04714	38.31740	0.00166	10.04826	1.46729	1.88599	10.14293	0.28230	15.73147	0.28228	-6.0	0.8	1.51	521	3.4408
6	0.02637	21.15595	0.00098	5.88142	1.46730	1.88679	12.89078	0.28228	16.60829	0.28227	-6.9	0.7	1.54	531	2.7388
13	0.03416	29.28333	0.00124	7.52412	1.46729	1.88656	11.28913	0.28228	15.08119	0.28227	-6.6	0.6	1.54	527	3.3066
17	0.02186	17.52403	0.00078	4.66181	1.46733	1.88673	10.62554	0.28225	14.30773	0.28224	-7.1	1.0	1.58	531	2.7388
36	0.04714	37.89595	0.00170	10.31284	1.46730	1.88669	13.90255	0.28232	15.87397	0.28230	-4.8	0.5	1.46	529	1.1445
38	0.03624	29.05170	0.00132	7.89625	1.46728	1.88662	11.34966	0.28230	15.40726	0.28229	-6.1	0.7	1.50	523	2.8375
68	0.03706	29.77998	0.00132	7.95821	1.46731	1.88645	10.62597	0.28227	13.35866	0.28225	-6.7	0.8	1.56	520	3.2513

DISCUSSION

P-T conditions of metamorphism

The P-T history of the Carlos Chagas granites can be summarized in a number of stages (M₁ to M₂) based on textural evidence, mineral chemistry (Figures 5 to 8) and geothermobarometric data presented above (Figures 9, 10 and Table VI). Two main preserved metamorphic assemblages are recognized: A₁, given by Grt₁ + Bt₁ + Pl + Kfs + melt; and A₂, given by Grt₂ + Bt₂ + Pl + Kfs ± Ilm ± Sil ± melt.

Estimates geothermometric indicates that the garnet and biotite (mineral assemblage A₁) reached equilibrium conditions at 740-750 °C during the peak metamorphic M₁ event. The presence of biotite, plagioclase and quartz as inclusions in garnet (Figures 6 d-e) indicates that garnet growth in granite at the expense of these minerals according to the dehydration reaction: Bt₁ + Pl + Qz = Grt₁ ± Kfs + melt. However, experimental studies have shown that this reaction only occurs at temperatures between 780 and 830 °C (Patiño Douce & Johnston 1991, Vielzeuf & Montel 1994, Stevens et al. 1997), implying that the values found here are significantly lower than expected. Taken into consideration the uncertainties of the Grt-Bt thermometer (~ 60 °C), our estimates of temperature are like the recent results of Melo et al. (2017a). The peak pressure was attained at 9.5-10.5 kbar (Melo et al. 2017a). The lack of extensive post-peak retrogression is consistent with progressive melt loss and dehydration during metamorphism (White & Powell 2002), producing an almost completely anhydrous and refractory rock.

For the subsequent anatectic process, these rocks probably underwent rehydration after M₁ melt loss and prior to M₂ event. The development of muscovite-biotite fabric is interpreted as a strong evidence of this rehydration process (Melo et al. 2017a). Shear zones are considered tectonic corridors for fluid percolation. In the field, such structures are characterized by mylonitic rocks that show evidence of ductile deformation. The rehydration of the sheared rocks must be occurred through hydration of K-feldspar and some minor fraction of garnet (Grt₁), producing a muscovite-biotite fabric (Melo et al. 2017a). Hydrothermal gem deposits from districts of the Eastern Pegmatite Province in Araçuaí Orogen attest fluid movement through the crust (Pedrosa-Soares et al. 2011).

For the peak metamorphic of the M₂ event, the thermodynamic modeling (Figures 9 and 10) of the preserved mineral assemblage A₂ suggests that all samples (CCS-21 and CCS-25) depict a similar P-T conditions, recording temperatures between 713 and 729 °C and pressures from 6.3 to 7.2 kbar. According to conventional geothermobarometry data, both samples were formed within temperature and pressure ranges corresponding to medium pressure (5.2-5.5 kbar) and relatively high temperature (710°C). Grt₂ crystals have sillimanite inclusions (Figs. 6e and 6f), indicating that the garnet was produced by dehydration reaction: Pl + Qz + Sil + Bt₂ = Grt₂ ± Kfs + melt. According to the experimental studies of Vielzeuf & Holloway (1988), this reaction occurs at temperatures of approximately 750-800 °C. Pseudomorphed melt films and Bt-Qz symplectites are observed surrounding Grt₂ crystals (Figure 5e), corroborating that garnet grew in the presence of melt in studied samples. This remain residuum is confirmed by thermodynamic modelling, which shows < 2 vol.% of melt preserved in mineral assemblage A₂ (Figures 9 and 10). The studied granites contain a few leucocratic patches with diffuse margins which may be considered as leucosomes produced by a second metamorphic event (Figure 3c). In fact, they are free of the regional solid-state foliation and occur as autochthonous patches, indicating that the granites have undergone renewed melting. Considering the average density of the crust to be 2.75 g/cm³ and the barometric gradient to be 270 bar/km, we estimated depths of 23-27 km with a geothermal gradient up to 27-31 °C/km for the M₂ event that generated the mineral assemblage A₂ (Figure 13).

Figure 13
P–T path for collisional rocks from the Carlos Chagas Suite. Compilation of the metamorphic conditions from paragneisses (Nova Venécia Complex) and pre-collisional and collisional granitoids (Richter et al. 2016, Melo et al. 2017a, Gouvêa et al. 2020). Reaction curves after partial melting experiments (Patiño Douce & Johnston 1991, Stevens et al. 1997, Vielzeuf & Montel 1994).

If the melt is not removed via an open system, H₂O amounts of the system will favor significant retrogression of the mineral assemblage. The occurrence of retrograde assemblages led us to the infer that the studied rocks must have been rehydrated after M₂ metamorphic peak. Some garnet crystals (Grt₂) are partially replaced by biotite (Bt₃), suggesting retaining enough H₂O from crystallizing melt to develop the observed texture (Stevens et al. 1997). The application of Grt₂-Bt₃ geothermometer indicates low temperature (600-660 °C – Table IX, Figure 13), resulting from retrograde Fe-Mg exchange between garnet and biotite (Ferry & Spear 1978). This temperature is consistent with the hercynite formation in garnet rims at temperatures below 700 °C (Spear 1993).

The occurrence of two generations of garnet in combination with microstructures evidencing the former presence melt in studied samples may suggest temperatures at > 750 °C (Vielzeuf & Holloway 1988, Patiño Douce & Johnston 1991, Vielzeuf & Montel 1994, Stevens et al. 1997). In fact, the lack of major element zoning in garnet (Figure 8) indicates that temperatures of metamorphism exceeded 650 °C, the temperature at which growth zoning is typically homogenized by volume diffusion (Tracy 1982). Therefore, we interpreted that the metamorphic peak in all samples was attained at granulite-facies conditions and the mineral assemblages equilibrated close to the transition between upper amphibolite to granulite conditions (Figure 13).

Interpretation of scattered ages and source rocks

The investigation of zircon textures via CL images associated with U-Pb data suggest a complex magmatic/metamorphic history. The U-Pb data from this study reveal a wide range of zircon ages (ca. 631-520 M.y.) (Figures 12 a-b). The oldest rounded zircon grains (Group I: ca. 631-604 M.y.) show resorbed domains and are interpreted as inherited. Oscillatory to weakly zoned zircon cores (Group II) indicate the magmatic crystallization age of the granite at 586 ± 2 M.y. (Figures 12 a-b). Zircon grains commonly show structureless rims and low luminescence in CL images, which may be produced from dissolution-reprecipitation and overgrowth processes in response to interactions with silicate melts. The youngest zircon populations with ages peaks at 562 ± 2 M.y. (Group III) and 526 ± 4 M.y. (Group IV) represent overgrowth rims, formed during the regional metamorphism and later thermal episodes, respectively. Previous geochronological studies (U-Pb zircon and monazite) have constrained the age of the regional metamorphism at ca. 570-550 M.y. (Gradim et al. 2014, Richter et al. 2016, Melo et al. 2017 a-b, Araújo et al. 2020). Several studies have reported the record of a later thermal event (ca. 530- 515 M.y.), coeval with the emplacement of post-collisional melts (Richter et al. 2016, Melo et al. 2017a, b, Peixoto et al. 2018, Araújo et al. 2020, Melo et al. 2020). The Lu-Hf data from zircon grains of the studied rock show significant crustal contribution to the genesis of the collisional granites. As can see in Figures 12b and 12c, the protholiths involved during crustal melting for Carlos Chagas granites generation could be the metasedimentary rocks like the Nova Venécia complex.

Solutions of thermobarometric and geochronological data

The anatectic episodes recorded in Carlos Chagas Suite require a sustained heat source to maintain high temperatures. High crustal heat production can be associated with elevated amounts of heat-producing elements in a thickened crust during the collisional stage. Fertile paragneisses of the Nova Venécia complex have been interpreted as the main source of S-type granites of the Araçuaí Orogen (Gradim et al. 2014, Pedrosa-Soares et al. 2011, Araújo et al. 2020). However, the oldest Carlos Chagas granites crystallized before (ca. 600-590 M.y.: Melo et al. 2017b, Cavalcante et al. 2019) the regional metamorphic event (ca. 570-550 M.y.: Gradim et al. 2014, Richter et al. 2016, Melo et al. 2017 a-b, Peixoto et al. 2018, Araújo et al. 2020. Gouvêa et al. 2020), implying that collision must have started before 600 M.y. (Cavalcante et al. 2019; Fossen et al. 2020).

As a consequence of crustal thickening and gravitational overloading, the Araçuaí Orogen underwent extensional collapse. A huge number of post-collisional plutons (felsic and mafic) were emplaced between 530 and 480 M.y. (Gradim et al. 2014, De Campos et al. 2016, Aranda et al. 2020, Araújo et al. 2020, Melo et al. 2020). The heat source to generate the post-collisional rocks can be associated with asthenospheric mantle ascent probably related to slab break-off followed by delamination of lithosphere mantle during the extensional thinning and gravitational collapse of the orogen (Gradim et al. 2014) or the onset of a hotspot due to the destabilization of the asthenospheric mantle (De Campos et al. 2016). Another hypothesis is the tectonic switching that is induced by changes in the angle of the subducting slab (Moraes et al. 2015). This must have driven the high heat flux into the crust, raising the geotherms, and contributed to the generation of post-collisional rocks, as well as a second metamorphism in the studied area. In fact, the CL-images in combination with U-Pb data clearly show two different ages of overgrowth formation zircon (Figs. 11 and 12), being the younger age (M₂ event) concomitant to the emplacement of the post-collisional magmas in the orogen. Petrological and geochronological evidence of this late metamorphism event have been recorded in several units from the Araçuaí orogen (e.g., Melo et al. 2017 a-b, Peixoto et al. 2018, Serrano et al. 2018).

ACKNOWLEDGMENTS

This work is part of the Project nº 9486/2019 registered in PRPPG (UFES). Sandro L.M. Ferreira acknowledges a scholarship from PIIC/UFES (Edital 2019/2020). We appreciate financial support provided by Conselho Nacional de Desenvolvimento Científico (CNPq 402852/2012-5 and 401334/2012-0). We would like to thank associate editor and anonymous reviewer for the detailed and constructive reviews and Editor-in-Chief Claudio Riccomini for the editorial handling.

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Publication Dates

• Publication in this collection
  29 Oct 2021
• Date of issue
  2021

History

• Received
  27 Apr 2020
• Accepted
  18 Feb 2021