Zircon U-Pb dating and petrogenesis of the São José do Campestre Granite Complex, NE Brazil: an example of neoarchean mantle-derived post-collisional magmatism

Abstract

This article discusses the geochemical and petrological evolution of the São José do Campestre granite complex (SJCgr), the last Neoarchean plutonic event so far described in the São José do Campestre massif in NE Brazil. We report field, petrographic, zircon U-Pb dating, and whole rock and mineral chemistry for representative SJCgr samples. Laser ablation zircon U-Pb data indicate that the granite emplacement took place at 2664 ± 13 Ma. The rocks comprising the SJCgr have relatively well-preserved primary textures and fabrics and compositions varying from gabbro to syenogranite. Major and trace element contents reveal a metaluminous, calc-alkaline through transitional to alkaline signature, and LILE- and LREE-enriched series analogous to late Archean sanukitoid and modern arc granitoid. The evolution of the SJCgr is envisaged as follows: (1^st) partial melting of a metasomatized mantle (2.5–3.0 GPa, ~85–102 km, 1,000–1,200°C), generating a basaltic to basaltic andesitic magma; (2^nd) fractional crystallization (FC) of olivine at mantle or lower crustal depth, leading to the parental magmas of the magmatic series; and (3^rd) 40–15% FC of olivine gabbro-norite and olivine monzonite cumulates (400–600 MPa, 15–23 km). The SJCgr shares similarities with post-collisional granitoids and, thus, would represent the last Neoarchean episode of mantle-derived magma in Northeastern Brazil.

KEYWORDS:
sanukitoid; geochronology; geochemistry; petrogenesis; tectonics

INTRODUCTION

The Archean Eon has long been considered the most important period of continental growth and is characterized by much higher geothermal gradients than today (Richter 1988, Abbott et al. 1994, Vanderhaeghe et al. 2019). These features resulted in the genesis of komatiites and an enormous volume of tonalite-trondhjemite-tonalite-granodiorite (TTG) magmatism (Barker and Arth 1976, Condie 1981, Martin 1986, 1994, Herzberg 2016, Laurent et al. 2014, 2020). Potassium-rich alkaline granites (Achaean crustal progeny granites; cf. Nebel et al. 2018) appear in the late Archean, and their generation has been associated with partial melting of older crustal sources, mainly TTG (Jahn et al. 1981, Martin 1986, Sylvester 1994, Laurent et al. 2014). On the contrary, studies in well-investigated areas suggest that these granites are mantle-derived, formed through variable degrees of interaction between mantle peridotite and TTG melts. In this case, they are generally referred to as sanukitoid or Closepet-type plutons (Stern and Hanson 1991, Jayananda et al. 1995, 2000, 2018, Smithies and Champion 2000, Moyen et al. 2001, 2003, Martin et al. 2005, 2009, Moyen et al. 2001, 2003, Laurent et al. 2014, Guo et al. 2018, Nebel et al. 2018).

Recent analogs of Archean TTG, the so-called adakites, are formed in modern subduction zones and are distinguished by higher Mg, Ni, and Cr than Archean TGG (Martin 1999, Smithies 2000, Martin et al. 2005, Castillo 2006). Their generation involves a complex interaction system, in which the adakitic magma, once formed by partial melting of subducted oceanic crust, interacts with the overlying mantle and/or the lower arc crust (Defant and Drummond 1990, Drummond and Defant 1990, Sen and Dunn 1994, Maury et al. 1996, Stern and Kilian 1996, Rapp et al. 1999, Castillo 2012).

The gradual cooling of the Earth led to the progressive generation of the continental crust, of which two-thirds to three-quarters may have formed in the time range 2.5–2.0 Ga (Taylor and McLennan 1985, Santosh et al. 2015, Sun et al. 2019, Ganade et al. 2021). Hence, the Archean–Proterozoic transition has long been considered a marker of relevant changes in the geothermal gradient that led to major formation of continental crust in the Archean and evolved to a regime of progressively lower geothermal gradient and dominance of crustal recycling in the Middle and Late Proterozoic (Taylor and McLennan 1985, Martin 1994, Laurent et al. 2014, Nebel et al. 2018).

The Borborema Province in northeastern Brazil (Almeida et al. 1981) presents an outstanding opportunity to study the Neoarchean–Paleoproterozoic transition. This territory is located at the margin of a large continental landmass formed during the amalgamation of the Congo and São Francisco cratons (inset in Fig. 1). This vast region comprises Paleo-to Neoproterozoic units and some Archean fragments, the last ones represented by the São José do Campestre, Campo Grande, Granjeiro, Tróia, Mombaça, and Cruzeta remnants (Neves 2003, 2011, Van Schmus et al. 2003, 2008, 2011, Dantas et al. 2004, 2013, Souza et al. 2007, Ferreira et al. 2020, 2021, Ganade et al. 2021).

Figure 1
Simplified geology of the northern portion of the Borborema Province (integrated after Angelim et al. 2004, Arthaud et al. 2008, Souza et al. 2016, Oliveira and Medeiros 2018, Ganade et al. 2021). Inset shows the study region in a pre-drift reconstruction (after Sá 1994).

In this article, we present new zircon U-Pb geochronological, whole rock, and microprobe data for the São José do Campestre granite complex (Fig. 1) and discuss the following topics:

1. the evolution mechanism of a Late Neoarchean granite;

2. the magma genesis and associated tectonics;

3. the relative role of upper mantle and lower crustal materials as sources of magmas at that time.

GEOLOGICAL SETTING

The Borborema Province comprises several metasupracrustal sequences overlying Archean to Paleoproterozoic gneissic basement that have been subsequently intruded by several plutons during the Late Neoproterozoic (Figs. 1 and 2). This province resulted from the convergence of the West Africa — Amazonian and São Francisco — Congo cratons during the assembly of West Gondwana and extends from central and northeastern Brazil to the Trans-Saharan Orogen in northwestern Africa (Tuareg and Benino-Nigerian shields) and the Central African Orogen of Cameroon, Chad, and the Central African Republic (Caxito et al. 2020). All units forming this province are controlled and/or reworked by a complex system of continental-scale high-temperature shear zones (Caby et al. 1991, Vauchez et al. 1995, Neves et al. 2021).

Figure 2
Simplified geologic map of the study area.

The São José do Campestre Massif (SJCM; Fig. 3) comprises:

1. an Archean complex in its central part, which is composed of mafic to felsic metaplutonics and metasupracrustals;

2. Paleoproterozoic orthogneisses (the Caicó Complex; 2.25–2.11 Ga) surrounding the Archean core;

3. late Neoproterozoic metasupracrustal sequences (the Seridó Group with detrital zircons with age ≥ 640 Ma);

4. voluminous Pan-African granitoids and pink granite dykes (zircon U-Pb ages ranging from 600 to 543 Ma) (Dantas 1996, Van Schmus et al. 2003, Dantas et al. 2004, 2013, Hollanda et al. 2011, Souza et al. 2016, Ferreira et al. 2021).

Figure 3
Field aspects of the São José do Campestre granite complex, NE Brazil. (A) SJCgr crosscut by late hololeucocratic magnetite-rich pegmatite. (B) Gray SJCgr with late biotite-rich pegmatite injected along the linear fabric defined by the orientation of hornblende + recrystallized K-feldspar. (C) Quartz diorite interleaved with granite. (D) SJCgr with quartz diorite autolith. (E) Incomplete mixing of granite and quartz diorite. (F) SJCgr crosscut by late interlayered pegmatite and leucotonalite (below the scale). (A, B, and C) Site ES1-2, São José do Campestre city (6.308^oS/35.715^oW). (D, E, and F) Site ES580, Santa Maria city (5.854^oS/35.701^oW). Abbreviations for minerals (Whitney and Evans 2010): Kfs K-feldspar; Bt: Biotite; Mag: Magnetite. All pictures display horizontal surface.

The SJCM Archean complex is composed of the following units (Dantas 1996, Dantas et al. 2004, 2013, Souza et al. 2016, Ferreira et al. 2020): Senador Elói de Souza complex — diopside-grossular-bearing anorthosite, garnet-bearing metagabbro, and garnet-bearing paragneiss (3.03 Ga); São Pedro do Potengi biotite-bearing granodiorite (3.12 Ga); Brejinho complex — gabbro to tonalite metaplutonics (3.18 Ga); Serra Caiada biotite ± hedenbergite ± hornblende-bearing granitic orthogneiss (3.36 Ga); marbles and paragneisses of unknown age (inherited zircons of 3.4 and 3.2 Ga, and metamorphic zircon of ~3.0, 1.9, and 0.58 Ga); and Bom Jesus metagabbro to quartz diorite to tonalite gneisses (c. 3.41 Ga). Ediacaran plutons and N-S trending granite dykes are the last Precambrian magmatic events in the region.

Late Neoproterozoic high-temperature shear zones (Picuí-João Câmara Shear Zone) juxtapose the SJCM and the Rio Piranhas-Seridó Domain, which is composed of Neoproterozoic metasedimentary rocks (Seridó Group) overlying a Paleoproterozoic basement (Caicó Complex). Late Ediacaran structures are westward plunging synforms and antiforms in the western side of the Archean block and a southern plunging inverted antiform in the southern portion (Viegas 2007, Viegas and Souza 2007).

The São José do Campestre granite complex (SJCgr) is a hornblende-biotite-bearing granite with subalkaline and metaluminous affinity, previously analyzed by TIMS (zircon U-Pb), which yielded ages of 2685 ± 9 Ma (n = 6, MSWD of 8.8) and 2655 ± 4 Ma (n = 6, MSWD of 0.28), and whole rock Nd model ages of 3.5–3.3 Ga with ε_Nd(t) between –4.2 and –6.2 (Dantas 1996, Dantas et al. 2004). The SJCgr has been interpreted as derived from the melting of either earlier gneisses (Dantas 1996, Dantas et al. 2004) or an enriched mantle in a subduction tectonic setting (Dantas et al. 2013, Souza et al. 2016).

FIELD GEOLOGY

The São José do Campestre granite (SJCgr) crops out in the southern limit of the SJCM near the city of São José do Campestre, with the northern and eastern limits nearby Santa Maria and Boa Saúde cities, respectively (Fig. 2), in the Rio Grande do Norte state. It is a holocrystalline, medium- to coarse-grained, equigranular to slightly inequigranular, gray to pink deformed granite (Figs. 3A and 3B). A prominent subhorizontal lineation (Lγ) with a WSW-ENE trend is marked by the orientation of hornblende + biotite + K-feldspar + magnetite. Late millimetric to centimetric size magnetite + K-feldspar ± biotite-rich veins crosscut (Figs. 3A and 3C) or follow the linear fabric (Fig. 3B).

Hornblende-rich quartz diorite occurs as autoliths or in interdigitated contacts with the host granite (Figs. 3C and 3D). Incomplete mixing of granite and quartz diorite occurs in some places (Fig. 3E). Igneous layering (Sγ) is characterized by alternate layers of biotite-bearing pegmatite and hornblende-biotite-bearing tonalite crosscutting an earlier magmatic linear fabric (Lγ) of the hosting SJCgr, and both units are overprinted by a NW-SE tectonic planar fabric S₁ (Fig. 3F).

METHODS AND ANALYTICAL TECHNIQUES

Analytical methods included classical field and petrographic descriptions, as well as additional whole rock and mineral chemistry for selected SJCgr samples. These data were integrated with former results from Viegas (2007), Dantas et al. (2013), Souza et al. (2016), and Ribeiro (2019).

Microprobe analyses of selected amphibole, biotite, plagioclase, microcline, and Fe-Ti oxide crystals were done on a Cameca SX50 electron microprobe hosted at the Geosciences Institute, Universidade de Brasília (UnB). Analytical conditions were 15 kV of accelerating voltage, a beam current of 25 nA, a 10 s counting time, and a spot size of 1 μm. In-house standards included synthetic and natural minerals. Amphibole structural formula, Fe²⁺/Fe³⁺ratio, and (OH) calculations were done with the spreadsheet of Locock (2014), according to the IMA nomenclature scheme (Hawthorne et al. 2012). Cation proportions for biotite and feldspars were computed based on 24 and 8 oxygen molecules, respectively (Deer et al. 2013). Fe-Ti oxides were computed for 4 oxygen molecules/3 cations and Fe₂O₃ and FeO from charge balance after Carmichael (1967). For the feldspars and biotite, all iron was assumed as FeO. The analytical errors are ± 0.5–2% for SiO₂, Al₂O₃, Fe₂O₃, MgO, MnO, CaO, and TiO₂ and 4.5–5.6% for Na₂O and K₂O. All microprobe results are reported in Suppl. Data Table 1.

Table 1
Major and trace element fractional crystallization modeling of the Neorchean São José do Campestre granite complex, NE Brazil.

	Group 2 – Enriched LREE and HREE				Group 3 – Less enriched LREE and flat HREE
	C0	CL2	CL2’	Cumulate C2	C0	CL3	CL3’	Cumulate C3
	ES467	ES572	FC = 40%	Σr2 = 0.31	ES471	CE116B	FC = 15%	Σr2 = 0.02
SiO2 (wt. %)	59.70	66.19	66.23	49.82	69.89	74.10	74.11	48.15
TiO2	0.90	0.54	0.57	1.40	0.53	0.28	0.29	1.75
Al2O3	15.87	17.70	17.60	13.26	13.40	12.56	12.57	17.73
Fe2O3	6.96	3.09	3.12	12.77	5.38	3.16	3.16	16.83
MnO	0.10	0.04	0.05	0.18	0.09	0.04	0.08	0.14
MgO	4.30	1.39	1.39	8.68	0.31	0.19	0.19	0.93
CaO	6.62	4.44	4.45	9.91	1.70	0.99	0.99	5.33
Na2O	4.12	4.28	4.71	3.23	3.01	3.28	3.26	1.75
K2O	1.13	2.16	1.88	0.00	5.61	5.36	5.35	6.90
P2O5	0.30	0.17	0.00	0.74	0.08	0.05	0.00	0.50
Cs (ppm)	0.37	0.50	0.80		0.52	2.6	2.8
Hf	4.99	5.3	11.0		12.5	11.5	12.2
Nb	9.42	9.5	6.1		34.0	34.4	54.2
Ni	96.30	17.4	6.0		2.79	2.4	0.01
Rb	20.80	55.4	37.5		138.0	182.0	152.9
Sr	628.0	540.9	610.2		104.0	68.1	41.2
Ta	0.45	0.7	0.1		1.65	2.4	0.3
Th	5.14	12.2	8.9		11.4	23.7	45.8
U	0.23	0.6	0.5		1.36	4.3	8.1
Y	22.7	10.5	23.0		87.9	75.0	182.0
Zr	218.0	228.4	277.7		535.0	403.2	433.3
Cumulate (wt. %)				Ol Gabbro-norite				Ol monzonite
Plagioclase (An38)				49.38				30.55
K-feldspar (Or92)								42.06
Hortonolite (Fo30)				19.11				6.25
Orthopyroxene (En90)				7.59
Dio-Hed (Wo50-43)				19.43
Hornblende								7.75
Magnetite								9.09
Ilmenite				2.73				3.10
Apatite				1.76				1.18
Zircon								0.02

• LREE: light rare earth element; HREE: heavy rare earth elements; C₀ and C_L: less evolved and most evolved samples, respectively; C_L’: calculated liquid after extraction of the cumulate; Σr²: statistical error adjustment of the theoretical model through matrix inversion; FC: % of fractional crystallization. K-feldspar, amphibole, magnetite, and ilmenite compositions used in modeling fractional crystallization are from this article. The other phases were compiled from et al. The partition coefficients used for modeling are mostly those compiled by Nielsen (2022). We also used partition coefficients for zircon after Martin (1987), et al, Rubatto and Hermann (2007), Burnham and Berry (2012), and et al, and apatite after Prowatke and Klemme (2006). Symbols for minerals (according to Whitney and Evans 2010): Ol: olivine; Di: diopside; Hd: hendebergite; An: anorthite; Wo: wollastonite; Fo: forsterite; En: enstatite; Or: orthoclase.

A selected sample (CLZ40) was analyzed for major and trace elements at ALS Minerals Laboratories (www.alsglobal.com). Major elements were measured by inductively coupled plasma atomic emission spectroscopy (ICP-AES), whereas trace element abundances were computed by inductively coupled plasma mass spectroscopy (ICP-MS) after fusion with lithium tetraborate. Loss on ignition (LOI) was determined by the gravimetric method after 1 h of heating at 1,000°C. The analytical errors are < 7% for oxides and < 6% for trace elements. Other 14 whole rock analyses were compiled from Dantas et al. (2013). Supplementary Data Table 2 displays all whole-rock chemical data.

For in-situ zircon U-Pb dating, one sample (ES471) was fragmented with a jaw crusher, followed by 250 mm grinding, manual panning, and mineral separation by methylene iodide and a Frantz magnetic separator. Zircon grains were then manually selected with a binocular loupe, mounted on epoxy resin, and imaged by cathodoluminescence. All procedures were conducted at the Geoscience Institute of the Universidade Estadual de Campinas (IG/UNICAMP). The isotopic data were also obtained at IG/UNICAMP on an ICP-MS Element XR (Thermo Scientific) coupled to an Excite193 laser ablation system (Photon Machines) with a HelEx ablation cell (25 μm laser beam). Data were reduced with the Iolite software and compared with the reference zircon 91500 (1065.4 ± 0.3 Ma; Wiedenbeck et al. 1995) and the Peixe zircon (571 ± 10 Ma; Navarro et al. 2017) for data quality control. Supplementary Data Table 3 reports all zircon U-Th-Pb isotopic results.

RESULTS

Petrography

The SJCgr rocks vary from quartz diorites to syenogranites (Streckeisen 1976), the latter being the main litotype, with variable modal contents of felsic minerals. According to Dantas et al. (2013), the modal composition of the SJCgr follows the calc-alkaline intermediate K-enrichment trend of Lameyre and Bowden (1982).

On textural grounds, these rocks present equigranular, medium-grained textures, and a discrete alignment of hornblende (Fig. 4A), which is usually anhedral and has a skeletal to poikilitic texture involving feldspar and quartz and late transformation to biotite (Figs. 4A–4D). Sodic, myrmekitic plagioclase (An_10–25 by the Michel-Lévy method), perthitic to mesoperthitic microcline (Fig. 4C), and quartz often show minor evidence of late recrystallization. This includes earlier-formed crystals showing irregular shapes and subhedral polygonal mosaics. Quartz appears interstitial to feldspar grains or forms discontinuous millimetric-size venulations concordant with the magmatic fabric.

Figure 4
Petrographic and textural features (sample ES471; 6.236°S/35.710°W). (A) Green hornblende, apatite, zircon, and a colorless medium-grained quartz-feldspathic groundmass. (B) Mesoperthitic microcline, green hornblende, and zircon. (C) Detail of cumulate texture of hornblende and rounded apatite. (D) Detail of cumulate texture of hornblende and magnetite and late biotite lamella. Abbreviations for minerals (Whitney and Evans 2010): Mag: magnetite; Zrn: zircon; Hbl: hornblende; Mc: microcline. (A, B, and C) uncrossed polarizers; (D) crossed polarizers.

Mafic phases (up to 65% in volume) are composed mainly of deep green to brown hornblende (poikilitic grains enclosing microcline, plagioclase, and zircon; up to 2–3 mm in size; Figs. 4C and 4D), plus brown biotite (< 1–5%; usually related to alteration of hornblende; locally, it may be altered to green chlorite + titanite + magnetite) and magnetite (< 1–3%; irregular, skeletal to interstitial grains up to 4 mm in length). Hornblende predominates in the less evolved quartz dioritic facies, whereas biotite surpasses hornblende in the most evolved ones.

Other accessory phases include apatite (< 1% in the syenogranites and up to 3% in quartz diorites; euhedral to rounded crystals; Fig. 4C), titanite (< 1–2%; anhedral grains replace biotite), allanite (< 1%; anhedral crystals with epidote rim and crosscutting earlier feldspar and biotite; it may reach up to 2 mm in diameter), garnet (< 1%; irregular rounded grains associated with plagioclase), zircon (euhedral to subhedral, quadratic to rectangular-shaped gray grains crosscutting or hosted in hornblende or magnetite; it may reach up to 0.5 mm in length; Fig. 4B), and rare relicts of uralitized pale green diopside.

All analyzed samples show a marked fabric defined by continuous millimetric thick bands of feldspar plus quartz alternated with mafic-rich bands (cf. Figs. 3A to 3F). A new generation of hornblende postdates earlier poikilitic hornblende, and features of recrystallization of plagioclase and microcline point to the overprinting of magmatic minerals and fabrics.

Mineral chemistry

Electron microprobe analyses of feldspar (plagioclase, n = 5; K-feldspar, n = 3), amphibole (n = 8), biotite (n = 4), and oxides (n = 2) of sample ES36 (see location in Fig. 3) are given as a supplementary data file (Suppl. Data Table 1). The analyzed spots were located on the core and crystalline rims of plagioclase and amphibole, and on K-feldspar, biotite, and Fe-Ti oxide crystal cores.

Feldspar

Plagioclase is sodic and shows no significant chemical zoning between the core and rims, and the anorthite content of five single crystals ranges from 12.68 to 15.33% (average: 13.85 ± 0.92%). Microcline has the composition Or_92-71Ab_8-29; perthitic exsolutions for one grain are still more sodic (Ab₉₉An₁).

Amphibole

The two types of amphibole crystals are hastingsite and ferro-pargasite (Fig. 5A; cf. Hawthorne et al. 2012). They have Si contents between 6.206 and 6.325 cations per formula unit (cpfu), Al between 1.937 and 1.997 cpfu, and Na+K between 0.839 and 1.003 cpfu. They are relatively homogeneous from core to rim, have low Ti contents (0.163–0.248 cpfu), and are Fe-enriched, with the fe# number [Fe/(Fe + Mg)] ranging from 0.90 to 0.91.

Figure 5
Mineral chemistry of amphibole (A) and biotite (B–D) according to different classifications. (A) C* (cpfu) vs. A* after Hawthorne et al. (2012). (B) Al_t (cpfu) vs. Fe²⁺/(Mg + Fe²⁺) after Rieder et al. (1998) and Deer et al. (2013). (C) Ternary plot TiO₂-(FeO + MnO)-MgO (all in wt. %) after Nachit et al. (2005). (D) Mg vs. Al_total (both in cpfu) of biotite from magmas of peralkaline to aluminum-potassic affinity after Stussi and Cuney (1996).

Biotite

Biotite crystals are relatively enriched in Ti (0.484–0.548 cpfu), Fe (4.88–4.958 cpfu), and Mg-depleted (0.573–0.646 cpfu), leading to fe# ratios between 0.89 and 0.9, which put them between annite and syderophyllite (Rieder et al. 1998, Deer et al. 2013; Fig. 5B). They plot within the ‘primary biotite’ field in the diagram after Nachit et al. (2005; Fig. 5C) and show chemical signatures like those found in micas of alkaline to slightly subalkaline magmas (Stussi and Cuney 1996; Fig. 5D).

Iron oxides

Representative electron microprobe data of two oxide crystals indicate magnetite (TiO₂ 0.94 wt. %, Al₂O₃ 0.32 wt. %) and ilmenite (TiO₂ 48.37 wt. %, FeO_t 47.51 wt. %, MnO 1.27 wt. %) with contents of the ulvöspinel molecule of 2.33 and 1.55, respectively.

Whole rock chemistry

Geochemical characterization

The analyzed samples were divided into three groups as follows (see Suppl. Data Table 2):

1. group 1 (n = 1) is the least evolved, with SiO₂ = 52.83 wt. %, and MgO = 5.35 wt. %, has the highest anorthite, diopside, and hypersthene and the lowest quartz and orthoclase normative contents;

2. group 2 (n = 5) with SiO₂ and MgO in the respective ranges of 58.49–67.8 wt. % and 3.64–0.57 wt. %;

3. group 3 (n = 10) with SiO₂ and MgO of 69.19–75.12 wt. % and 0.31–0.27 wt. %, respectively, and the highest quartz and orthoclase and the lowest anorthite and diopside normative contents.

Oxide composition (in wt. %) show that the SJCgr rocks have a wide SiO₂ variation (52.8–75.1); low TiO₂ (0.24–1.84) and P₂O₅ (0.05–0.29); and low-to-moderate MgO (0.10–5.35), CaO (0.84–9.02), Fe₂O₃t (2.45–8.96), and Al₂O₃ (11.7–15.55). They are moderately enriched in alkalis (Na₂O+K₂O of 4.5–8.6); less evolved samples (SiO₂ = 52.8–65.3) are Na-enriched (Na₂O/K₂O ratios of 2.0–3.6), whereas more evolved Si-rich facies (SiO₂ = 66.6–75.1) is K-enriched (Na₂O/K₂O ratios of 0.5–0.8). Fe₂O₃t contents are higher than MgO for individual samples, leading to Mg# [ = 100MgO/(0.8998*Fe₂O₃t+MgO) molar] in the range of 3–55. All samples are silica-saturated, with normative quartz ranging from 2.8 to 35%.

Figure 6 exhibits major element plots (Barker and Arth 1976, Miyashiro 1978, Cox et al. 1979, Maniar and Piccoli 1989, Frost et al. 2001). In the TAS diagram (Fig. 6A), the samples are subalkaline (groups 1 and 2) or transitional to alkaline (group 3). This is corroborated by the SiO₂ vs. MALI plot (Fig. 6B), in which the samples scatter from calcic (the least evolved) to transitional to alkali-calcic (the most evolved). In the molar A/CNK vs. A/NK plot (Fig. 6C), which measures aluminum saturation, all samples are metaluminous. Coherently, in the ternary cationic diagram K-Na-Ca, the least evolved (group 1 and part of group 2) plot close to the trondhjemitic field, while the most evolved ones (part of group 2 and the entire group 3) tend to follow the calc-alkaline (K-enriched) trend or toward the Na-K edge (Fig. 6D).

Figure 6
Major element geochemical characteristics of the studied rocks. (A) TAS (total alkalis vs. SiO₂ in anhydrous basis) diagram after Cox et al. (1979); alkaline-subalkaline boundary from Miyashiro (1978). (B) SiO₂ vs. MALI as proposed by Frost et al. (2001). (C) Saturation alumina index (Shand 1943) in the diagram of Maniar and Piccoli (1989). (D) Ternary cationic K-Na-Ca (Barker and Arth 1976) showing the classical calc-alkaline differentiation (labeled CA after Nockolds and Allen 1953) and trondhjemitic (named Trj from Barker and Arth 1976) trends; the trondhjemitic field is from Martin et al. (2005).

In the fe* vs. SiO₂ plot (Fig. 7A), less evolved samples show magnesian affinity, whereas the more evolved are ferroan (SiO₂ > 65 wt. %). They scatter within the I- to A-type granites (the most evolved ones). In the fe* vs. Al₂O₃ diagram (Fig. 7B), most samples have been classified from slightly oxidized to reduced A-type granite, while some less evolved samples are calc-alkaline. In the ternary plot (NaO₂ + K₂O)-5Fe₂O₃t-5(CaO + MgO) of Fig. 7C, these less evolved samples are akin to subducted and active continental margin related granites, whereas the most evolved correspond to A₂-type granites that are commonly associated with post-collisional geodynamic settings. In the bilogarithimic (Y + Nb) vs. Rb diagram (Fig. 7D), the studied rocks scatter within the post-collisional field of Pearce (1996).

Figure 7
Geochemical affinity and tectonic setting of the studied rocks. (A) SiO₂ vs. fe* and the fields of Ferroan, Magnesian, and I-, A- and S-type granitoids (Frost et al. 2001). (B) Al₂O₃ vs. fe* and fields of calc-alkaline, oxidized A-type and reduced A-type granitoids (Dall’Agnol & Oliveira 2007). (C) (Na₂O + K₂O)-5Fe₂O₃t-5(CaO + MgO) ternary plot showing the fields of A₁- and A₂-type granites (Grebennikov 2014). (D) Bilogarithmic Y + Nb vs. Rb plot with the fields of volcanic arc, syn- and post-collisional, within plate, and ocean ridge granitoids (Pearce 1996).

Binary Harker-Type diagrams

Groups 1–3 as defined above show similar behavior for TiO₂, CaO, Fe₂O₃t, and P₂O₅ (Fig. 8) that decrease regularly with differentiation (i.e., increasing SiO₂). Al₂O₃ shares similar behavior, decreasing toward high silica, although the curve is less pronounced. Alkalis (Na₂O + K₂O) have opposite patterns as their abundances increase with differentiation. Of note, MgO content decreases strongly for group 2 but is very low (< 0.4 wt. %) and constant for group 3. Considering all samples, they define concave upward (for MgO) and concave downward (for K₂O) curves peaking at ~65 wt. % SiO₂. This may imply that magmatic evolution has been achieved by at least two main stages of crystallization (Wilson 1989).

Figure 8
Major element Harker-type plots for the studied rocks (in wt. %).

Trace element data (Fig. 9) reveal lower Rb (< 33 ppm), Sc (25 ppm), Zr (72 ppm), and Ba (197 ppm) and high V (212 ppm) for group 1 (n = 1); group 2 (n = 5) has higher Sr (up to 685 ppm) and Ba (up to 4,588 ppm), with one sample having the highest Zr (910 ppm) and Nb (45.4 ppm); and group 3 (n = 10) has the highest Rb (182 ppm); Ni is usually low (< 40 ppm), except for one sample of group 1 and another one of group 3 (81.0–96.3 ppm). Figure 9 demonstrates the compatible behavior of V, Ni, and Sc, and the incompatible behavior of Rb and Ba in group 2. In group 3, Ba, Nb, Zr, and Sc behave as compatible elements, whereas Rb and Ni are incompatible.

Figure 9
Trace element Harker-type plots for the studied rocks (SiO₂ in wt. %, and trace elements in ppm).

Rare earth elements

The three rock groups exhibit discrete differences in their REE patterns and ratios, as follows. Group 1 (Fig. 10A) is enriched in light rare earth elements (LREE) with (La/Yb)_N ratio of 3.45, has a slightly positive europium anomaly (Eu/E* = 1.32), an almost flat heavy rare earth element (HREE) with (Gd/Yb)_N of 1.3, and the lowest normalized Yb (Yb_N = 11.0). Its overall pattern is like those found in Archean enriched tholeiites (Condie 1976). Group 2 (Fig. 10B) has steeper LREE with (La/Yb)_N of 29.5 ± 18.0, and less pronounced HREE [(Gd/Yb)_N = 2.6 ± 1.0] patterns. There is no significant europium anomaly (Eu/Eu* = 0.91 ± 0.44) and the Yb_N varies around 21.2 ± 14.2. Such patterns are similar to those of Archean sanukitoids, which have (La/Yb)_N = 41, Yb_N = 7.6, and Eu/Eu* = 0.84 (Martin et al. 2005, 2010, Heilimo et al. 2010). Group 3 (Fig. 10C) presents a fractioned trend from more-abundant LREE toward less-abundant HREE with (La/Yb)N = 7.0 ± 0.9 and the highest negative Eu anomaly (Eu/Eu* = 0.65 ± 0.14), as well as the highest Yb_N (43.2 ± 5.0). The Gd_N/Yb_N ratio (1.4 ± 0.1) is slightly lower than group 1. These REE patterns share similarities with modern arc granitoids, in which (La/Yb)_N = 7.75, Yb_N = 21.8, and Eu/Eu* = 0.69 (Martin et al. 2010).

Figure 10
Chondrite-normalized (Sun and McDonough 1989) rare earth element patterns. (A) Group 1 sample ES407; average of enriched Archean Tholeiite (EATh) from Condie (1976). (B) Group 2 samples; average of late Archean sanukitoid from Martin et al. (2009). (C) Group 3 samples; average of modern arc granitoids from Martin et al. (2009).

Crystallization conditions

Pressure and temperature

Pressure estimates were obtained from the analyzed amphibole of sample ES36 (from group 3, as defined in the geochemical section) with the Amp-TB2 (Ridolfi 2021), an updated model for the amphibole-only barometer of Ridolfi and Renzulli (2012; equation P1e), and the amphibole/plagioclase Al-Si partitioning barometer of Molina et al. (2015). The first barometer yielded pressures from 497 to 601 MPa (average = 558 ± 28 MPa), whereas the formulation of Molina et al. (2015) gave slightly higher pressures in the range of 550–718 MPa (average = 636 ± 36 MPa), calculated using temperatures derived from the Amp-TB2 thermometer (see below) and an An₁₄Ab₈₄Or₁ oligoclase (within the plagioclase compositional range used in Molina’s calibration). Altogether, it is assumed that the SJCGR has an average confining pressure of 597 ± 46 MPa. However, a word of caution is necessary when Fe-rich amphiboles are used for pressure calculations. Anderson and Smith (1995) argued that amphiboles with fe# [Fe/(Fe+Mg)] > 0.65 are indicative of low oxygen fugacity and yield overestimated pressures. Yang (2017) proposed an empirical method using CIPW normative quartz (Qtz) and albite (Ab) plus orthoclase (Or) compositions to estimate emplacement pressure of granite intrusions, referred to as Qtz-geobarometer (Yang et al. 2019, 2021). Using the improved Qtz-geobarometer of Yang et al. (2021), which corrects pressure estimation for both normative anorthite (following Blundy & Cashman 2001) and redox estate, most of the SJCgr samples (with normative quartz < 35 wt. %) cluster around 335 ± 124 MPa. Therefore, considering all pressure estimates, we assume that the emplacement of the SJCgr took place at pressures from ~400 to 600 MPa that corresponds to emplacement depths between ~15 and 23 km (for an average crustal density of 2.7 g/cm³).

The apatite (TsatAp) and zircon (TsatZr) saturation thermometric expressions after Watson and Harrison (1983) and Harrison and Watson (1984), along with the Amp-TB2 amphibole-only thermometer (Ridolfi 2021), were used to estimate magmatic temperatures for the SJCgr. The TsatAp thermometer yielded the highest estimates of 958–868°C (average = 904 ± 28°C) that should approach the liquidus temperature. The TsatZr was applied to a group of ten samples with SiO₂ ranging from 69.19 to 75.12 wt. %, in which Zr (648–212 ppm) behaves as a compatible element (see whole rock composition in Suppl. Data Table 2). Results are slightly lower than TsatAp, varying from 812 to 913 (average = 870 ± 32°C). Amphibole crystallization temperatures are between 803 and 835°C, with an average of 819 ± 7°C. Perthitic exsolution in K-feldspar point to temperature lower than 600°C and P_H2O > 200 MPa (Bowen & Tuttle 1950) for the subsolidus stage. Furthermore, a temperature of ~588°C obtained for the magnetite-ilmenite pair (Carmichael 1967) may reflect late- to post-magmatic reequilibration.

Oxygen fugacity (fO ₂)

Redox conditions during magma crystallization can be qualitatively monitored by the fe# number [ = Fe/(Fe+Mg)] of amphibole and biotite (e.g., Anderson & Smith 1995, Anderson et al. 2008). The analyzed amphibole and biotite crystals present fe# values between 0.90–0.92 and 0.89–0.9 respectively, which are indicative of relatively reducing crystallization conditions. Quantitative f O₂ estimation was done with the Amp-TB2 oxybarometer (Ridolfi 2021) and the improved Qtz-geobarometer of Yang et al. (2021) using, respectively, amphibole and whole rock compositions. The Amp-TB2 oxybarometer yielded oxygen fugacity values between –14.7 and –14.1, whereas higher values were obtained with the Qtz-geobarometer (–9.9 to –13.7). These f _O2 values are slightly under the NNO (Ni-NiO) buffer (–1.1 < Δ_NNO < –0.5) and confirm the relatively moderate to reducing crystallization environments for the SJCgr magmas. Such a conclusion is also supported by the scattering of samples through the reduced and slightly oxidized A-type granite fields in Fig. 8B.

It is worth mentioning that magnetite is a common accessory phase in some SCJgr varieties, and this may pose a contradiction with the above-inferred redox conditions. Although magnetite-bearing granites are generally interpreted as being “oxidized” (cf. Ishihara 1981), many authors have argued that, in fact, the presence of magnetite in granites is not incompatible with a more reduced character (cf. Dall’Agnol and Oliveira 2007, Campos et al. 2016).

U-Pb zircon geochronology

Sample ES471 (Latitude 6.236^oS/Longitude 35.710^oW; see location in Fig. 3 and whole composition in Suppl. Data Table 2) of the geochemically defined group 3 was selected for in-situ zircon U-Pb dating. The sample corresponds to a slightly deformed magnetite-apatite-amphibole-bearing medium- to coarse-grained granite (Figs. 4A and 4B) with euhedral zircon grains (length up to 0.9 mm), often with triangular terminations, crosscutting amphibole (Fig. 4B).

U-Pb analyses were performed on 37 zircon grains, totaling 44 spot analyses. The results are in Suppl. Data Table 3. Figures 11A and 11B display cathodoluminescence images of the analyzed grains. They usually have oscillatory zoning and prismatic habits and a bipyramidal shape, with an average length (L) of 243.5 ± 98.9 μm, width (W) of 92.6 ± 21.9 μm, and aspect ratios of L/W = 2.7 ± 1.1; Th/U ratio varies around 0.61 ± 0.17 (Th = 126 ± 145 ppm, U = 189 ± 158 ppm). Some zircon grains clearly present two growth generations with distinct age populations (Fig. 11B; spots z25c and z25r with ²⁰⁷Pb/²⁰⁶Pb ages of 2670 ± 24 Ma and 2613 ± 23 Ma, respectively). A group of 44 single analyses with < 5% of discordance yields intercepts at 2683 ± 22 and 724 ± 370 Ma (MSWD = 0.23; Fig. 11B). A total of 19 spots gave a concordia age of 2664 ± 13 Ma (MSWD = 2.8; inset in Fig. 11C). The age of 2664 ± 13 Ma is interpreted as the magmatic age. Slightly younger ages at 2605 ± 23 and 2582 ± 25 Ma (e.g., spots z26c and z26r in Fig. 11B) may represent a new generation of igneous zircon or post-magmatic lead loss. The results here obtained agree within the analytical errors with previous ages obtained by TIMS U-Pb zircon of 2683 ± 7 Ma (sample CE116B) and 2655 ± 4 Ma (sample EC80) by Dantas (1996).

Figure 11
Zircon U-Pb results for sample ES471, southern portion of the São José do Campestre granite complex. (A and B) Zircon cathodoluminescence images. (C) Concordia diagram and weighted ²⁰⁶Pb/²³⁸U age of concordant zircons (concordance > 95%). Ages for displayed spots are ²⁰⁷Pb/²⁰⁶Pb date in million years (Suppl. Data Table 3).

DISCUSSION

Petrologic evolution

Differentiation mechanism and quantification of fractionation

Considering the curved trends in some Harker-type plots (e.g., CaO and MgO in Fig. 9), quantitative modeling of fractional crystallization was done for groups 2 (diorite to granite in Fig. 7A; SiO₂ = 58.49–67.8 wt. %) and 3 (granite to alkali-granite in Fig. 7A; 69.19–75.12 wt. %). Group 1 was not considered in modeling calculations, since it includes just one sample (ES407).

Compatible (Ni, Zr) vs. incompatible (Rb) trace element correlation (Fig. 12) suggests fractional crystallization mechanism as the main process of evolution for group 3 samples (cf. Cocherie 1986, Martin 1987). Accordingly, major element least-square mass-balance modeling was done with the spreadsheet PetroMode (Christiansen 2022). Mineral compositions are those reported herein (feldspars, amphiboles, biotite, magnetite, and ilmenite), besides apatite and pyroxenes from Deer et al. (2013). Further trace element modeling followed the classical fractional crystallization equation (Rayleigh 1896): C_L = C₀F^(D-1), in which C_L = concentration of the trace element in the most evolved sample, C₀ = concentration of the trace element in the least evolved sample, F = (1-FC) [FC is the degree of fractional crystallization, with FC < 1], and D = bulk partition coefficient. The partition coefficients are those provided by the database from Nielsen (2022), as well as values reported for zircon by Martin (1987), Thomas et al. (2002), Rubatto and Hermann (2007), Burnham and Berry (2012), and Gudelius et al. (2020) and apatite by Prowatke and Klemme (2006). The precision of the modeling is monitored by the parameter Σr². The results for both groups 2 and 3 are shown in Table 1 and Figs. 13A and 13B.

For group 2, we considered L₀ = ES467 (SiO₂ = 59.7 wt. %, MgO = 4.3 wt. %, and Fe₂O₃t = 6.96 wt. %; anhydrous base) as the least differentiated sample, and L₁ = ES572 (SiO₂ = 66.19 wt. %, MgO = 1.39 wt. %, and Fe₂O₃t = 3.09 wt. %; anhydrous base) as the most differentiated one. The calculated liquid (C_L2’) and cumulate (C₂) are in Table 1. Calculations indicate a FC degree of 40% with an excellent statistical error (Σr² = 0.31). The cumulate (C₂) has mainly plagioclase, hortonolite, and clinopyroxene that correspond to a gabbro-norite modal composition (Table 1). There is a very good adjustment for trace elements (compare C_L2 and C_L2’ in Table 1) that is also observed for the REE (Fig. 13A).

For group 3, we assumed L₀ = ES471 (SiO₂ = 69.89 wt. %, MgO = 0.31 wt. %, Fe₂O₃t = 5.38 wt. %; anhydrous base) and CE116B (SiO₂ = 74.1 wt. %, MgO = 0.19 wt. %, Fe₂O₃t = 3.16 wt. %; anhydrous base) as the most differentiated sample (L₃). Modeling resulted in an excellent adjustment for major (Σr² = 0.02), trace elements (Table 1; compare C_L3 and C_L3’ in Table 1), and REE (Fig. 13B). The cumulate C₃ contains mainly plagioclase, K-feldspar, magnetite, and hornblende, corresponding to a hornblende-magnetite-bearing monzonite modal composition (Table 1; compare with Figs. 4C and 4D). In cumulate C₃, traces of zircon were added to the model to adjust zirconium. It is worth mention the absence of biotite in both cumulates and the abundance of K-feldspar, hornblende, and magnetite in cumulate C₃.

Crustal contamination

It has long been demonstrated that crustal assimilation coupled with fractional crystallization (AFC) is a relevant process during evolution of mantle-derived magmas through their ascent and emplacement into the lower and upper continental crust (DePaolo 1981, Huppert et al. 1985, Wilson 1989). For geochemical modeling, we have applied mixing equations of DePaolo (1981) and trace element composition of the lower continental crust (LCC; Wedepohl 1995). Calculations for different fractions of crustal material reveal the possibility of contamination of the dioritic (group 2) magma by ~5% of LCC, followed by 30–40% of fractional crystallization (Fig. 13C). Some amount of crustal contamination seems to account for group 3 set as well (Fig. 13D), which is corroborated by negative ε_Nd(2.7 Ga) of –4.02 to –6.18 (Suppl. Data Table 2).

Magma genesis and evolution

The geochemical characteristics described above suggest an upper mantle peridotite as the most probable source for the SCJgr complex. Any earlier crustal material would be impressed as a slab component (low- and high-silica adakites) hybridized with the upper mantle, leading to late Archean sanukitoids (or the ~2.6 to 2.5 Ga Closepet-type) and modern arc granitoids (Stern and Hanson 1991, Rapp et al. 1999, Smithies and Champion 2000, Moyen et al. 2001, 2003, Lobach-Zhuchenko et al. 2005, Martin et al. 2005, Nebel et al. 2018). This hybridized source should be LREE-enriched (e.g., Fig. 10). In this sense, the available ε_Nd(2.7 Ga) of –4.1 to –6.2 and Nd model age (T_DM) of 3.9–3.3 Ga for our samples (Dantas 1996, Dantas et al. 2013) support such an interpretation of an earlier crustal component that provided, by partial melting, the mixing component to the upper mantle peridotite (e.g., Jayananda et al. 2000, 2018, Nebel et al. 2018).

According to Wilson (1989), primary magmas derived directly from the upper mantle peridotite would have mg# > 70, Ni > 400–500 ppm, Cr > 1,000 ppm, and SiO₂ < 50 wt. %. The least evolved samples (ES407 of group 1 and ES467 of group 2) have SiO₂ = 53–58 wt. %, mg# = 53–58, Ni = 81–96 ppm, and Cr = 41–164 ppm and, therefore, may not represent primary melts. However, olivine ( ± chromite) fractionation at mantle depths from a primary basaltic melt would generate silica-enriched and Ni- and Cr-depleted liquids (Nicholls and Ringwood 1972) like the ones described here.

Moderate-to-low contents of Sc (25–15 ppm) in the least evolved sample (ES407, ES467) suggest the presence of minor garnet and amphibole in the source as these minerals have K_D^Sc >> 1 (Nielsen 2022). Through investigation of high-pressure experiments that produced saturated tholeiites, andesitic, and dacitic melts, Nicholls and Ringwood (1972) and Green (1973) suggested that they were generated by partial melting of hydrous mantle peridotite (with pargasitic amphibole) followed by different degrees of fractional crystallization of olivine at mantle or crustal depths. Moreover, petrological constraints from high-pressure experimental runs discussed by Pilet (2015) led to assume a lithosphere metasomatized by amphibole-bearing veins as the most probable source of sodic (Na₂O/K₂O > 1), mantle-derived magmas. This agrees with the less evolved samples having Na₂O/K₂O = 3.3–2.0 wt. %.

Figure 14 summarizes the polybaric evolution of the studied rocks. The first stage is the partial melting of a hydrous metasomatized mantle peridotite (probably with garnet and/or spinel, pargasitic amphibole, and phlogopite), forming a basaltic to andesitic magma at pressures of 2.5–3.0 GPa (~85 to 102 km) and temperatures of 1,000–1,200°C. In the second stage, these magmas underwent olivine fractionation at mantle and/or crustal depths giving origin to our less evolved samples (ES407 and ES467). The third stage comprises fractional crystallization (± crustal assimilation) under pressures of ~400 to 600 MPa (~15 to 23 km) and temperatures > 800°C, under reducing to slightly oxidizing conditions that has left olivine gabbro-norite or olivine monzonite cumulates (Table 1).

Figure 12
Incompatible (Rb) vs. compatible (Ni and Zr) trace element plots for groups (A) 2 and (B) 3.

Figure 13
Rare earth element spectra for quantitative modeling of (A and B) fractional crystallization and (C and D) crustal assimilation and fractional crystallization (AFC) modeling of the studied rocks in a Sr vs. Rb/Sr plot. (A) Group 2, with L₀ = sample ES467 and L₁ = sample ES572; (B) Group 3, with L₀ = sample ES471 and L₁ = sample CE116B (complete data in Suppl. Data Table 1). Cumulate compositions are in Table 1. Explanation for a and b: L₀ less evolved sample; L₁ more differentiated sample; C₂, C₃ cumulates for groups 2 and 3, respectively; Σr² statistical error. Abbreviations for minerals (Whitney and Evans 2010): Pl: plagioclase: Kfs: K-feldspar; Ol: olivine; Opx: orthopyroxene; Cpx: clinopyroxene; Ilm: ilmenite; Mag: magnetite; Ap: apatite; Zrn: zircon. In (C) and (D), the calculated AFC curves with a mass assimilation/fractionation ratio of 0.05 (= 5% crustal contamination) of the LCC (Wedepohl 1995) for groups 2 (ES467 evolves to ES572) and 3 (ES471 evolves to CE116B) are shown in (C) and (D), respectively. Mixing the calculation equation for trace elements after DePaolo (1981).

Figure 14
Schematic evolution stages for the studied rocks (see Table 1 and Figs. 13A and 13B).

Tectonic setting and geodynamic implication

Figure 15 shows the investigated samples in ocean ridge granite (ORG)-normalized multi-element diagram. Group 1 sample ES407 (Fig. 15A) displays a spectrum with an inclination (Rb_N/Y_N) of 29.4, Yb_N = 0.02, and Y_N = 0.28, and no significant anomalies. These signatures share similarities with basaltic andesites of the Andean Central Volcanic Zone (Thorpe et al. 1984).

Figure 15
Ocean-ridge-granite (ORG; Pearce et al. 1984) normalized multielement plot for the studied rocks. (A) Group 1 compared to basaltic andesite from the Andean Central Volcanic Zone (ACVZ; Thorpe et al. 1984). (B and C) Groups 2 and 3, respectively, compared to granodiorite from the Northern Andean Coastal Batholith (NACB; Petford and Atherton 1996) and volcanic arc granite (VAG; Pearce et al. 1984).

Figure 16
SiO₂ vs. mg# for the studied rocks and experimental results from the literature. References: Mg-andesites, slab melt and hybridized melt from Rapp et al. (1999 and references therein); low-silica and high-silica adakites from Martin et al. (2005); sanukitoid from Martin et al. (2005, 2009); modern arc granitoids from Martin et al. (2009); mantle melts from Baker and Stolper (1994), Hirose (1997), Hirschmann et al. (1998), Robinson et al. (1998), Wasylenki et al. (2003), Condamine and Médard (2014), and Condamine et al. (2016). Abbreviations for minerals (according to Whitney and Evans 2010): Grt: Garnet; Am: amphibole; Cpx: clinopyroxene.

Figures 15B and 15C present four representative samples of group 2 and three of group 3. Both sets have steeper patterns and similar Rb_N/Y_N ratios (averages of 35.4 ± 29.2 and 35.9 ± 12.6, respectively). Group 2 may display positive anomalies for Ba (3.3 ± 2.9) and Ce (7.2 ± 5.3). These anomalies are slightly lower in group 3. Group 2 patterns can be compared to those found in the Andean Costal Batholiths (Petford and Atherton 1996) and Volcanic Arc Granites (Pearce et al. 1984), whereas group 3 has slightly higher normalized contents of Ce to Yb.

Whole rock compositions of the studied SJCgr samples indicate that these rocks correspond to slightly oxidizing to reduced A₂-type granites (Figs. 7A–7C) with some degree of crustal contamination that is reflected in their normative corundum <0.5 wt. %, and A/NK and A/CNK ratios of > 1.1 and < 1, respectively. Their trace element contents are akin to post-collisional granites, with some samples showing geochemical characteristics of volcanic arc granites (Fig. 7D). This agrees with the preservation of igneous features and rock relationships (Fig. 3), as well as microscopic textures (Fig. 4). Furthermore, the REE patterns point to a subduction tectonic setting for their genesis and evolution. In this tectonic environment, the subducting oceanic crust and the overlying mantle wedge are the main magma reservoirs. We explored this hypothesis in a SiO₂ vs. mg# binary diagram by plotting both our samples and experimental results from the literature (Fig. 16). Group 1 (sample ES407; SiO₂ 53.9 wt. %, mg# 54.2) is the closest to the field of mantle-derived melts. The less evolved group 2 sample ES467 (SiO₂ 58.5 wt. %, mg# 55) lies in between low-silica adakites and high-silica adakites and almost overlaps the average of sanukitoids (SiO₂ 58.8 wt. %, mg# 56.8). Group 2 samples ES543A and ES572 have SiO₂ and mg# of ~65 wt. % and 46, respectively and plot between high-silica adakites and modern arc granitoids (SiO₂ ~65 to 68 wt. %, mg# ~48 to 81). Group 1 sample ES407 and the least evolved sample from group 2 (ES467) scatter among the field of hybridized melts (slab-derived melts mixed in different proportions with mantle-derived melts; Rapp et al. 1999). The remaining samples have SiO₂ > 67 wt. % and mg# < 20, falling out of the field of slab melts. This means that this kind of source cannot explain the compositional variation of the studied rocks. Hence, they may be explained as late differentiates from basic to intermediate parental liquids, as highlighted in Fig. 16.

The timing of the magma generation and the subduction/collisional process are difficult to assess (Condie 1989, Davidson and Arculus 2006). Nevertheless, delamination of the continental lithospheric mantle (Rudnick 1995), underplating of basaltic magma at the base of the lower continental crust (Huppert et al. 1988), and even the action of mantle plumes (Abbott 1996, Condie 2001) could not be discarded.

Nebel et al. (2018) consider that heat from ascending hot, high-Mg sanukitoid magmas, formed earlier through interaction of TTG melts with the asthenospheric mantle, triggers melting of lower crust material to generate late Archean high-K granites, the entire evolution extending for > 200 Ma. Based on the Nd model age (T_DM) of clinopyroxene from peridotite xenoliths in alkaline and kimberlitic magmas, Menzies et al. (1987) stated that the metasomatism of upper mantle peridotite by silicate melts and hydrous fluids is a very early event. According to Menzies et al. (1987), this event could take place ~0.5 to 1 Ga before the crystallization age, and even much older than 1 Ga beneath cratons. This time lapse estimation lies within the SJCgr intrusion age interval of ~2.7 Ga (Fig. 12C) and within the Nd model ages of 3.9–3.3 Ga (Dantas 1996, Dantas et al. 2013).

Several episodes of juvenile accretion and crustal recycling have been suggested based on zircon U-Pb data for the São José do Campestre Massif (SJCM) by Dantas et al. (2013) and Souza et al. (2016): 3.5–3.4 Ga, 3.36 Ga, 3.25 Ga, 3.18 Ga, 3.12 Ga, 3.03 Ga, and 2.69 Ga. In an exhaustive compilation of zircon U-Pb and Nd model ages, Ganade et al. (2021) stated that the SJCM is one smaller cratonic blocks in northeastern Brazil that resulted from fragmentation of the São Francisco — Congo — West Africa cratons followed by successive collisions during the Proterozoic. In this regard, somewhat analogous events of crust formation are depicted for the Mairi complex, northern São Francisco Craton, by Oliveira et al. (2020) and Moreira et al. (2022), where banded gneisses (TTG-like felsic layers alternating with metamorphosed gabbro-diorite bands) have zircon U-Pb ages of 3.7–3.6 Ga. For the same complex, Moreira et al. (2022) described successive younger events of crustal growth at 3.55–3.52, 3.49–3.33, and 2.74–2.58 Ga, with main crustal recycling events in the Neoarchean. Similar magmatic events and post-collisional granite emplacement are also documented in southeastern Brazil (Marimon et al. 2022) and in the Dharwar craton (e.g., Jayananda et al. 2006, Sebastian et al. 2021, Kumar et al. 2022). In the latter case, however, they assumed that the last Neoarchean episode was marked by juvenile, mantle-derived magmatism with some degree of crustal component. This could be accommodated by convergent tectonic settings (island arcs or continental margin-types) and the widespread amalgamation of earlier Archean units. Therefore, the data here obtained for the SJCgr complex show that global plate tectonics was the main process involved in the generation of continental crust in the SJCM, as also proposed for the Southern Brasília Orogen and the Southern São Francisco Craton (Marimon et al. 2022).

CONCLUSION

Integration of the data reported here and those in the existing literature led to the main conclusions listed below.

• Field relations, petrographic, textural, and geochemical data, and in-situ U-Pb dating reveal that the plutonic rocks forming the São José do Campestre granite complex, northeastern Brazil, vary from gabbro to granite with preserved igneous textures and fabrics (e.g., Dantas et al. 2013, Souza et al. 2016).

• Laser ablation zircon U-Pb dating indicates a Concordia age of 2664 ± 13 Ma, which is considered the emplacement age (e.g., Dantas et al. 2004).

• The rocks are metaluminous, calc-alkaline to transitional to alkaline, LILE-, and LREE-enriched, with chemical signatures akin to slightly oxidized to reduced post-collisional A₂-type granites.

• Based on geochemical data, three rock groups are defined: Group 1, one sample, SiO₂ = 59.7 wt. %, MgO = 5.46 wt. %; Group 2, five samples, SiO₂ = 59.7–68.23 wt. %, MgO = 4.3–0.6 wt. %; Group 3, ten samples, SiO₂ = 69.9–75.1 wt. %, MgO = 0.1–0.3 wt. %.

• The petrogenetic evolution of the SCJgr occurred as follows: 1^st — partial melting of a metasomatized (with minor amounts of spinel ± garnet, amphibole, phlogopite) upper mantle peridotite (~83–100 km, 1,000–1,200°C), generating a basaltic to basaltic andesitic magma; 2^nd — fractional crystallization of olivine at mantle or subcrustal depths, leading to the less evolved samples of each group; 3^rd — 40–15% fractional crystallization giving rise to the magmatic series and olivine gabbro-norite and olivine monzonite cumulates (15–23 km, > 800°C).

• This magmatism has major and trace element contents analogous to late Archean sanukitoid and modern arc granitoids.

• The SJCgr complex represents the last Neoarchean episode of mantle-derived magma (and is related to some kind of plate tectonics) and marks the amalgamation of all earlier Archean blocks of the studied cratonic block.

• Supplementary data Supplementary data associated with this article can be found in the online version:
  Supplementary Tables A1–A3 (http://sfbjg.siteoficial.ws/Sf/2023/4889202320220079.pdf).

ACKNOWLEDGMENTS

This article is a tribute to Professor Hervé Martin, who sadly passed away 1 year ago. ZSS is particularly indebted to Professor Hervé as his supervisor of the DESS (Diplôme d´Études en Sciences, Université de Rennes, 1991) and a post-doctoral sabbatical stage (Université Blaise Pascal, Clermont-Ferrand, 1997). In both situations, the focus was on the acquisition and interpretation of whole rock chemistry and isotopic data from Archean to Early Proterozoic juvenile granitoids of NE Brazil. Professor Hervé always made complex topics simpler with joy, good humor, and fine intelligence. He was really a great scholar and scientist. Thank you very much for all the learning you gave us. The authors acknowledge the Programa de Pos-Graduação em Geodinâmica e Geofísica and the Departamento de Geologia of the Universidade Federal do Rio Grande do Norte (PPGG/UFRN and DGeo/UFRN) for supporting field and laboratory works. The Conselho Nacional de Desenvolvimento Científico e Tecnológico (CNPq) is acknowledged for financial support (grant number 408607/2018-1). The authors also thank Nilson Francisquini Botelho (Instituto de Geociências, Universidade de Brasília) for helping with mineral chemical data acquisition and fruitful discussions and Alexandre H. Santos Filho for microscopic image acquisition.

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Publication Dates

• Publication in this collection
  19 June 2023
• Date of issue
  2023

History

• Received
  11 Oct 2022
• Accepted
  15 Mar 2023