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Research Article
A magmatic barcode for the São Francisco Craton: Contextual in-situSHRIMP U\\Pb baddeleyite and zircon dating of the Lavras, Pará deMinas and Formiga dyke swarms and implications for Columbia andRodinia reconstructions
Fabrício de Andrade Caxito a,b,⁎, Steffen Hagemann b, Tatiana Gonçalves Dias a,b, Vitor Barrote c,Elton Luiz Dantas d, Alexandre de Oliveira Chaves a, Marcos Santos Campello a, Filippe Couto Campos a
a Centro de Pesquisa Prof. Manoel Teixeira da Costa, Instituto de Geociências, Universidade Federal de Minas Gerais (CPMTC-IGC-UFMG), Av. Antônio Carlos 6627, CEP 31270-901 Belo Horizonte,MG, Brazilb Centre for Exploration Targeting, School of Earth and Environment, The University of Western Australia, 35, Stirling Highway, Crawley, WA 6009, Australiac Isotopia Laboratory, EAE, Monash University, Melbourne, VIC, 3800, Australiad Instituto de Geociências, Universidade de Brasília, Campus Universitário Darcy Ribeiro, Asa Norte, CEP 70910-900, Brasília, DF, Brazil
⁎ Corresponding author at: Centro de Pesquisa Prof. Made Geociências, Universidade Federal de Minas Gerais (CCarlos 6627, CEP 31270-901 Belo Horizonte, MG, Brazil.
Eoarchean to Rhyacian crust is preserved in the São Francisco Craton of eastern Brazil. To position this crustal seg-ment in paleocontinental reconstructions, precise, accurate and robust geochronological data are necessary, es-pecially for the diverse regional-scale mafic dyke swarms that crosscut the cratonic basement. Thisgeochronological database can then be used to construct a magmatic barcode and compare it to the barcode ofother cratons around the world, in search of similarities that might help to position these pieces in thepaleocontinental puzzles. New U\\Pb SHRIMP contextual in-situ (thin section) dating of baddeleyite and zirconfrom six samples of three different dyke swarms in the southern São Francisco Craton, in addition to novellithogeochemical and Nd\\Sr isotopic data, allow to pinpoint dyke emplacement at ca. 2.55 Ga (Lavras Iswarm; εNd(t) = −6 to +2; TDM not calculable), ca. 1.8–1.7 Ga (Pará de Minas I and II dyke swarms; εNd(t) = −10 to −5; TDM = 2.5–3.0 Ga) and at ca. 900 Ma (Formiga dyke swarm; εNd(t) = −7 to 0; TDM =1.4–2.3 Ga). The new geochronological data suggest a link between the regional dyke swarms and extensionalstresses during the onset of crustal rifting related to the evolution of the Minas, Espinhaço and Macaúbas basins,respectively. A barcode comparison shows strong similarity between the São Francisco and North China cratons(Lavras-Taipingzhai/Naoyumen swarms, Pará de Minas-Taihang/Miyun swarms, Formiga/Pedro Lessa-Sariwon/Dashigou swarms; and possible correlations of the poorly dated 2.2–2.0 Ga Paraopeba swarm with similar agedswarms inNorth China), suggesting proximity of those two cratonic blocks,whether theywere part or not of Pro-terozoic paleocontinents such as Columbia and Rodinia. The novel geochronological data support previous inter-pretations based on paleomagnetic data and provide further refinements of the geochronological record of thesouthern hemisphere cratonic blocks, allowing for better-tied global correlations.
Regional-scale mafic dyke swarms are key units for the reconstruc-tion of ancient geodynamic configurations (e.g., Srivastava, 2011).They are usually deeply rooted andwell preserved due to vertical exten-sion in the interior of Archean and Paleoproterozoic cratons, and oftenpreserve the orientation of Earth'smagnetic poles during crystallization,
noel Teixeira da Costa, InstitutoPMTC-IGC-UFMG), Av. Antônio
thus revealing the ancient latitude and azimuthal orientation of the cra-tonic piece (e.g. Bispo-Santos et al., 2014; D'Agrella-Filho et al., 2020;Evans et al., 2016; Pesonen et al., 2012; Salminen et al., 2016; Teixeiraet al., 2013; Xu et al., 2014). Furthermore, they carry geochemical infor-mation that can be used to infer specific geotectonic settings (e.g.Santiago et al., 2020) and in correlations with other igneous rockssuch as those found in LIPs (Large Igneous Provinces; Ernst, 2014).LIPs are very important geodynamic features andmay have contributedto major paleoclimatic and biologic events in Earth's history (e.g. Ernstand Youbi, 2017), therefore their precise dating and interpretation inthe geotectonic context is crucial in order to propose causal feedbacklinks between these geospheres.
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The Archean crust preserved in Planet Earth is today dispersed incirca 35 main fragments (Bleeker, 2003). Whether or not those frag-ments were once joined in the various proposed supercontinents(e.g., Superia, Vaalbara, Slave, Columbia, Rodinia; Zhao et al., 2004; Liet al., 2008) is yet to be resolved. The reconstruction of pre-Pangea su-percontinents is very problematic, since it cannot rely on the record ofoceanic crust spreading, thefit of continental borders and other tectonicfeatures commonly used in the reconstruction of plate motions thattook place within the last hundreds of million years.
A proposed tool to resolve this problem is the construction of mag-matic “barcodes” for each cratonic fragment (Bleeker and Ernst, 2006).This is done through precise dating of intrusive and extrusive LIProcks, dykes and sills from mafic swarms. Those barcodes are usuallypresented as a column of lines or bars, younging towards the top. Eachbar represents a regional magmatic event, which can be used to identifywhich crustal fragment was affected by it. If it coincides, at least in part,with the barcode of another cratonic fragment, then both cratons mayhave been dispersed from the same ancient paleocontinent.
The São Francisco Craton (Heilbron et al., 2017 and references therein)in east-central Brazil (Fig. 1) represents an Archean-Paleoproterozoiccrustal fragment. The São Francisco Cratonwas united to its African coun-terpart, the Congo Craton, from circa 2.05 Ga to the opening of the SouthAtlantic circa 130Ma ago (e.g. Heilbron et al., 2017; Santiago et al., 2020).
Fig. 1. The São Francisco Craton (B) in the context of Gondwana (A). In (C), simplified geologicarespective key in (D); (E) shows an Analytical Signal Magnetometric map (data from CODEMIG
Despite recent efforts (Cederberg et al., 2016; Chaves et al., 2019; Ernstet al., 2013; Silveira et al., 2013), there are still few studies that offer pre-cise and reliable ages for the mafic intrusions in both cratons, whichwould allow to better understand the evolution of these cratonic frag-ments through the comparison of their respective barcodes with thoseof other cratonic fragments around the world.
In this paper, we present the results of contextual in situ (thin sec-tion) dating of mafic dykes from the Lavras, Pará deMinas and Formigaswarms in the southern São Francisco Craton, southeastern Brazil. Theresults link the swarms to rifting and basin evolution within the SãoFrancisco paleocontinent (Lavras I swarm – Minas basin at 2.5 Ga;Pará de Minas I and II swarms – Espinhaço basin at 1.79–1.71 Ga;Formiga swarm –Macaúbas basin at 900Ma), provides robust geochro-nological information for the ongoing construction of a magmaticbarcode for the São Francisco Craton, and helps to position the cratonwithin the context of ancient supercontinents and superplumes.
2. Geological context and previous geochronological data
Dyke swarms of distinct age and orientation occur in the SouthernSão Francisco Craton (Fig. 1), in the surroundings of Belo Horizonte(Carneiro et al., 1998; Chaves, 2001; Chaves and Corrêa Neves, 2005;Corrêa da Costa et al., 2006; Oliveira, 2004; Pinese, 1997; Teixeira
l map of the southern São Francisco Craton showing themain regional dyke swarms, with, 2001, 2006) of the same area. Gr. - Group; Sg. – Supergroup; Undiff. – Undifferentiated.
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et al., 1998) and in the Espinhaço Range, fringing the eastern cratonicborder (Chaves et al., 2019; Machado et al., 1989). Distinct generationsof swarms are proposed based on orientation, emplacement mecha-nisms, available radiometric ages, petrography and lithochemistry. Al-though some robust zircon and baddeleyite U\\Pb ages haveprogressively been acquired for some of those swarms in the last fewyears (e.g. Cederberg et al., 2016; Chaves et al., 2019), in general, theavailable geochronological dataset is still very restricted and prelimi-nary, based mainly on Rb\\Sr and Sm\\Nd whole-rock isochron datesand K\\Ar and 40Ar\\39Ar mineral dates. Unfortunately, these methods(especially Rb\\Sr and Ar-system ages) are prone to resetting, leadingto partial loss of the daughter isotope during later thermal events.Thus, the dates obtained have to be carefully interpreted as minimumcrystallization ages or ages possibly reset by secondary processes (e.g.metamorphism).
The main families of mafic dykes and sills of the southern SãoFrancisco Craton (Fig. 1) can be grouped in the following swarms:
Fig. 2. Field,macroscopic andmicroscopic characteristics of some of the studied dyke swarms. A)swarm (scale bar is 1 mm long); B) Site P7, a basaltic porphyritic dyke from the Pará de Minaroadcut near Bom Sucesso, with a basaltic dyke of the Pará de Minas swarm; E) and F) Sample
2.1. Lavras swarm(s)
Located at the extreme SWborder of the craton, intruded inMeso- toNeoarchean (3.2–2.9; 2.7–2.65 Ga) medium- to high-grade gneisses ofthe cratonic basement,TTG terranes and mafic-ultramafic complexesof ca. 2.7 Ga (Teixeira et al., 1998). Despite the fact that these dykesare commonly grouped as part of a single Lavras swarm, two distinctswarms can be identified. The first (Lavras I) constitutes circa 90% ofthe dykes and is composed of norites and gabbronorites (Fig. 2A),trending WNW-ESSE, with each individual dyke reaching up to ca.30 km length and 100 m width (Corrêa da Costa et al., 2006; Pinese,1997). No robust geochronological data has yet been presented forthese rocks, and the only radiometric date available is a Sm\\Ndwhole-rock isochron of 2658 ± 44 Ma (Pinese, 1997).
The second group (Lavras II) is composed of thin (2–10 m) amphib-olite,metabasite and diabase dykes trendingNNW-SSE, N-S andNE-SW.The only radiometric date available is an amphibolite 40Ar\\39Ar
Photomicrographyunder crossedpolarizers of sample L4 of a norite from the Lavras I dykes swarm; C) Sample P1 of a porphyritic dyke of the Pará de Minas swarm; D) Site P2 in as of sites F6 and F5, coarse diabase dyke samples of the Formiga dyke swarm.
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plateau age of 1970±7Ma, interpreted as aminimumage for dyke em-placement or metamorphism (Pinese, 1997).
2.2. Paraopeba swarm
Exposed in association to vertical shear zones in gneisses and gra-nitic rocks of 2.8–2.65 Ga of the Belo Horizonte Complex to the NW ofBelo Horizonte, at the Paraopeba river valley. The dykes are composedof amphibolite, trending NNW-SSE and are up to ca. 5 km long and2–10 m thick. The only radiometric data available are a Rb\\Sr whole-rock isochron of 2189 ± 45 Ma (Chaves, 2001) and K\\Ar amphiboleages of 2116 ± 46, 2054 ± 37, 2049 ± 61, 2005 ± 26 and 1995 ±31 Ma (Teixeira et al., 1998). As the rocks are metamorphic and boththe Rb\\Sr andK\\Ar clocks are prone to partial to total resettingduringrelatively mild tectono-thermal events, these dates represent only esti-mates of themaximum crystallization age of the dykes, and probably anestimate of the age of metamorphism. However, recently, Caxito et al.(2015) and Barros (2019) obtained similar ages for metamafic rockswhich might represent dyke swarms in the northwestern margin ofthe São Francisco Craton (Angico Farm ambhibolite and FazendaRealeza mafic dyke) using the more reliable U\\Pb clock in zircon, at1958.3 ± 16 Ma and 2070 ± 42 Ma, respectively. Whether or notthose ages will be reproduced in other dykes throughout the craton re-mains to be tested.
2.3. Pará de Minas swarm
Trending N50-60 W, this is the most extensive dyke swarm of theexposed portion of the Southern São Francisco Craton (Fig. 2B, C, D). In-dividual dykes may reach up to 100 m width and up to ca. 150 km incontinuous length. Dykes are tholeiitic, either fine- to coarse-grainedphaneritic sub-ophitic massive gabbros or porphyritic gabbros withcentimetric labradorite-andesine phenocrystals in a groundmass of au-gite, with accessory ilmenite, interstitial K-feldspar, apatite andbaddeleyite. The dykes crosscut the Archean Belo Horizonte, CampoBelo and Bonfim complexes, but also supracrustal rocks of the Rio dasVelhas Supergroup and of the Paleoproterozoic Minas Supergroup.
Geochronological data available for the Pará de Minas swarm com-prise a K\\Ar amphibole age of 1707 ± 64 Ma (Carneiro et al., 1998),a mineral Rb\\Sr age of 1740 ± 54 Ma (Chaves, 2001), an 40Ar\\39Arplateau age of 1752 ± 15 Ma (Oliveira, 2004) and recently, U\\Pb(TIMS) upper intercept baddeleyite dates clustering at ca. 1795 and1710 Ma (Cederberg et al., 2016). The U\\Pb (TIMS) dates led the au-thors to suggest two distinct generations of Statherian mafic dykes inthe Pará de Minas swarm, at ca. 1795 and 1710 Ma (Pará de Minas Iand II, respectively). Cedeberg et al. (2016) also presented a 766 ±36 Ma 207Pb/206Pb weighted mean date using three baddeleyite crystalfractions from a sample of a dyke formerly believed to be part of thePará deMinas swarm near Campo Belo. In the Quadrilátero Ferrífero re-gion, the Ibirité gabbro, trendingN-S, also presented anU\\Pb upper in-tercept baddeleyite date of 1714 ± 5 Ma (Silva et al., 1995), thuscorrelating to the youngest of the Statherian dyke emplacement eventsfrom Cedeberg et al. (2016). All of those dates are very similar tosistematically dated plutonic and volcanic mafic and felsic rift-relatedrocks on the eastern border of the São Francisco Craton that are partof the Espinhaço rift system (São João da Chapada volcanic rocks andBorrachudos granite), which continues northward to the Rio dosRemédios Group, marking a very important extensional event withinthe cratonic landmass (Chemale Jr. et al., 2012; Guadagnin andChemale Jr., 2015).
Recently, Chaves and Rezende (2019) presented a baddeleyiteU\\Pbweightedmean 207Pb/206Pb date of 1762± 2Ma for a NW-SE di-abase dyke from a swarm in the region of Januária, in the central portionof the craton, where a basement window outcrops below the BambuíGroup carbonate-siliciclastic cover. Chaves and Rezende (2019) haveproposed that the Pará de Minas swarms and Januária swarm are
related to different mantle plume conduits of a same Statheriansuperplume event below Columbia supercontinent.
2.4. Formiga and Pedro Lessa swarms
At the southern São Francisco Craton, the Formiga dyke swarm ischaracterized by dykes trending N60-70E, almost perpendicular to thePará de Minas swarms. Most basalt and diabase dykes are equigranularaphanitic or medium-grained phaneritic in the central portions, withophitic to intergranular textures (Fig. 2E and F). Subordinated porphy-ritic rocks contain up to about 5 cm long plagioclase phenocrysts.Andesine-labradorite laths constitute up to 40–50 vol%, and aresurrounded by augite crystals making up to 30 vol% of the groundmass.Ilmenite, apatite and minor sulphides are the main accessories(Campello et al., 2015; Chaves, 2001).
At the eastern border of the São Francisco Craton, along the N-Strending Espinhaço range, a wide group of metabasic sills, plugs anddykes crosscuts the Paleo-Mesoproterozoic metasedimentary sequenceof the Espinhaço Supergroup; this is commonly named Pedro LessaSuite (e.g. Machado et al., 1989) but sometimes referred to as Diaman-tina Suite (e.g. Girardi et al., 2017). The dykes trend E-W, N-S and NE-SW and are commonly affected by the Brasiliano west-verging(cratonwards) thrust-and-fold belts of the Ediacaran/Cambrian (ca.575–530 Ma) Araçuaí Orogen. Machado et al. (1989), in a conferencepaper, obtained a U\\Pb composed baddeleyite/zircon date of 906 ±2 Ma for a metadiabase sill, but the data remain unpublished. Otherpoorly constrained K\\Ar and 40Ar\\39Ar data also suggest emplace-ment of dykes near the Quadrilátero Ferrífero region around1000–900 Ma (Carneiro et al., 1998; Oliveira, 2004). Recently, Chaveset al. (2019) presented U\\Pb baddeleyite dates for the Santa Mariade Itabira gabbroic intrusion that crosscuts the basement to theEspinhaço Supergroup in the eastern border of the homonymousrange with an upper intercept date of 940 ± 42 Ma.
Although the Formiga swarm was not formerly dated, based oncrosscutting relationswith the other dyke swarms and regional correla-tions, it is normally considered as a correlative to the Pedro Lessa suitefurther east (e.g. Chaves, 2001).
2.5. Dyke swarms of the northern São Francisco Craton
In the northern São Francisco Craton, in Bahia state, the Uauá,Curaçá, Chapada Diamantina and Salvador-Olivença swarms occur.
The Uauá swarm is composed of two families; the oldest, dated at2726 ± 3 Ma (U\\Pb baddeleyite TIMS), trends NW-SE to N-S and iscomposed of norites and tholeiitic amphibolites, and the youngest,dated at 2623 ± 7 Ma (U\\Pb zircon) trends NE-SW and is composedof diabase and metabasic dykes (dates from Oliveira et al., 2013). TheCuraçá diabase dykes trend preferentially NE-SW, dated throughU\\Pb in baddeleyite at 1507 ± 7 Ma (Silveira et al., 2013).
The Chapada Diamantina swarm comprises mafic dykes and sills in-truded in the Paleo/Mesoproterozoic sedimentary rocks of theParaguaçu and Chapada Diamantina groups, correlate to the EspinhaçoSupergroup further south. The dykes trend preferentially NNW-SSE.The U\\Pb zircon dates of 1514 ± 22 Ma (Babinski et al., 1999),1496 ± 3.2 Ma (Guimarães et al., 2005), a baddeleyite U\\Pb date of1501 ± 9.1 Ma (Silveira et al., 2013) and Ar\\Ar dates of 1512 ± 6and 1514 ± 5 Ma (Battilani et al., 2005) indicate that this swarm is co-eval to the Curaçá swarm.
Lastly, the Salvador-Olivença or Bahia dyke swarm (e.g. Corrêa-Gomes and Oliveira, 2000; Evans et al., 2016) is composed of N-S andE-W (Salvador region) and NE-SW (Olivença region) trending diabasesand amphibolites, dated at 924.2± 3.8, 921.5± 4.3 (Salvador), 926.1±4.6 and 918.2 ± 6.7 Ma (Olivença - U\\Pb baddeleyite, Evans et al.,2016). Paleomagnetic data from the Salvador dykes place the SãoFrancisco/Congo craton in moderate to high paleolatitudes at ca.920 Ma, allowing for various interpretations of the position of this
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cratonic block within or outside the Rodinia paleocontinent (Evanset al., 2016).
3. Materials and methods
Available in the Supplementary Files.
4. Results
4.1. Whole-rock geochemistry
Whole-rock geochemistry data (Supplementary Material, Table 1)are presentedmainly for comparisonwith established geochemical pat-terns for the southern São Francisco Craton dyke swarms and to assurethat the samples collected and dated are representative of each dykeswarm. Thus, geochemical results are not extensively discussed here(see Girardi et al., 2017 for a complete review). Chondrite-normalizedREE and MORB-normalized incompatible element patterns of the ana-lyzed samples are presented in Fig. 3.
4.2. Lavras swarm
Norite and gabbronorite samples of the Lavras I swarm are charac-terized by 49.54–52.06% SiO2, with Mg# = 0.40–0.65. Chondrite-normalized REE patterns are flat and with only slight LREE enrichment((La/Yb)N = 1.78–2.26), with no significant Eu anomaly (Eu/Eu* =0.72–0.94). MORB-normalized incompatible element spidergramsshow modest LILE enrichment (Ba <25 times MORB) and a Nb\\Tatrough. Those characteristics correspond to previous results for noritesand gabbronorites from the Lavras I swarm, characterized by50.05–55.49% SiO2 and Mg# = 0.50–0.72 (Girardi et al., 2017; Pinese,1997).
4.3. Pará de Minas (I and II) swarm
Geochemistry data for the Pará de Minas swarm samples are re-markably homogeneous, placing the dykes as subalkaline basalts or ba-saltic andesites (SiO2 = 53.70–55.63%) with evolved Mg# of 0.23–0.34.Chondrite-normalized REE patterns are relatively flat and showmoder-ate LREE enrichment ((La/Yb)N = 4.80–10.16), with no significant Euanomaly except for samples P8 (Eu/Eu* = 0.69) and P2 (Eu/Eu* =0.81). MORB-normalized incompatible element spidergrams showLILE enrichment (Ba >100 times MORB) and a Nb\\Ta trough. Thosecharacteristics correspond to previous results for the Pará de Minasswarm (Chaves, 2001).
4.4. Formiga swarm
Geochemistry data of the two dated samples (F5 and F6) indicates abasalt trachyandesite compositionwith 55.41 and 57.69% SiO2 andMg#of 0.28 and 0.31, respectively. Chondrite-normalized REE patterns arerelatively flat and steep, with (La/Yb)N ratios of 16.20 and 15.76, andno Eu anomalies (Eu/Eu* of 1.06 and 0.99). MORB-normalized incom-patible element spidergrams show LILE enrichment (Ba >100 timesMORB) and a Nb\\Ta through. Those characteristics correspond to themost evolved terms of the Formiga swarm (Chaves, 2001). Anothersample thatwas not dated butwas analyzed for geochemistry (F8) indi-cates an alkaline basalt compositionwith 49.78% SiO2 andMg# of 0.42, aless steep LREE-enriched chondrite-normalized pattern ((La/Yb)N =3.77 and Eu/Eu* = 0.74) but equally LILE-enriched MORB-normalizedpattern with a Nb\\Ta through. This sample corresponds to the lessevolved members of the Formiga swarm (Chaves, 2001).
Two samples from the Pedro Lessa dyke swarm, collected atMorro doJuá (JUA1) and Ipoema (IPO1) were also analyzed for comparison. Theyboth correspond to sills interleaved with metasedimentary rocks of theEspinhaço Supergroup. They show a subalkaline basalt composition
with 51.07 and 52.37% SiO2 and Mg# of 0.59 and 0.61, respectively.Chondrite-normalized REE patterns are flat and less steep than theFormiga samples, with (La/Yb)N ratios of 7.47 and 4.99, and no Eu anom-alies (Eu/Eu* of 0.97 and 1.11). MORB-normalized incompatible elementspidergrams also show minor LILE enrichment (Ba <60 times MORB)and a Nb\\Ta negative anomaly. These samples are comparable to theless evolved sample from the Formiga swarm (F8), especially due tothe very similar REE and incompatible element patterns (Fig. 3).
4.5. U\\Pb geochronology
Figs. 4 and 5 show microscopic aspects of some of the dated zirconand baddeleyite crystals, both under the optical microscope and underSEM. Below we will discuss each sample separatedly.
4.6. Lavras Swarm
4.6.1. Sample L4Both polycrystalline zircon aggregates with irregular borders, min-
ute prismatic baddeleyite, and composite grains with baddeleyitecores and zircon rims are observed in this sample. As this sampleshows no sign of metamorphic assemblages, the polycrystalline zirconaggregates are interpreted as the result of late-stagemagmaticfluid per-colation that caused progressive silicification of the original baddeleyitecrystals, with all stages of this transformation observed in thin section.Unfortunately, single baddeleyite crystals and composed grains are toosmall (<5 μm) for reliable SHRIMP analysis. Polycrystalline aggregatesof pure zircon can reach up to ca. 25 μm width (Fig. 5A) and thereforewere analyzed. The analyzed spots produced Th/U ratios between 0.28and 0.70. Eighteen spots in 11 crystals yielded a discordia line with anupper intercept at 2550 ± 17 Ma and a lower intercept at 617 ±270 Ma (Fig. 6A); the high uncertainty on the lower intercept is due tothe fact that most crystals are only slightly discordant, with up to 20%discordance on spot ZR16.2. Thefivemost concordant spots in four crys-tals yielded a Concordia age of 2551.1±9.8Ma (MSWD=0.78; Prob. offit = 0.38), interpreted as the crystallization age of the zircon aggre-gates (Fig. 6A).
4.7. Pará de Minas (I and II) swarm
4.7.1. Sample P6Baddeleyite crystals occur as small blades and wafers, mostly
20–30 μm long x 10–15 μm wide, with some isolated crystals reachingup to 80 μm long (B6; Fig. 4B and 5C). Crystals are normally clearunder BSE images, but discrete zircon rims occur locally. All of the 10spots in five crystals produced concordant data and a Concordia age of1807 ± 32 Ma. Due to the orientation effects of baddeleyite on206Pb/238U ratios during ion microprobe analysis (Wingate andCompston, 2000), the overlapping mean 207Pb/206Pb date of 1799 ±37 Ma (MSWD= 1.07; Prob. of fit = 0.38) is considered as the best es-timate for crystallization age (Fig. 6B).
4.7.2. Sample P1Baddeleyite crystals occur as blades and thin wafers disseminated
through the groundmass, up to 60 μm long x 10–15 μm wide (Fig. 4Aand C, 5B). Irregular zircon rims are common (Fig. 4C). Most of the ana-lyzed spots are concordant, but three discordant spots allow for the cal-culation of a discordia regression with an upper intercept at 1739 ±69 Ma and a lower intercept with no geological meaning at 178 ±280 Ma. Six concordant spots in four crystals yielded a Concordia ageof 1741 ± 52 Ma, and the mean 207Pb/206Pb date using all nine spotsin six crystals is 1736±63Ma (MSWD=0.74; Prob. offit=0.66), con-sidered as the best estimate for crystallization (Fig. 6C).
Fig. 3. Chondrite-normalized (Sun andMcDonough, 1989) REE andMORB-normalized incompatible element (Thompson et al., 1984) spidergrams of samples from the three studied dykeswarms.
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4.7.3. Sample P8Baddeleyite crystals in this sample form thin blades or wafers dis-
persed throughout the groundmass (Figs. 4D and 5D), mostly ca.10 μm wide to 20 μm long; some crystals are exceptionally long, suchas ca. 200 μm long B11 (Fig. 5D), but invariably thin. Zirconolite is an-other common Zr-bearing phase in this sample, and is locally associatedwith baddeleyite (Fig. 4D). Baddeleyite crystals are mostly clear and nozircon rimwas detected. Seven spotswere analyzed in seven crystals, allof them yielding concordant data and producing a Concordia age of1716 ± 53 Ma. Due to the large errors, the mean 207Pb/206Pb age is1743 ± 180 Ma. As the two ages overlap and all of the data points pro-duced concordant dates, the Concordia age of 1716± 53Ma (MSWD=
0.01; Prob. of fit = 0.93) is interpreted as the best estimate for crystal-lization of baddeleyite in this sample (Fig. 6D).
4.8. Formiga swarm
4.8.1. Sample F5Skeletal zircon crystals occur throughout the sample groundmass
(Figs. 4E, F, G and H; 5E), indicating late-stage crystallization. Crys-tals are highly irregular with exquisite shapes, most are cloudy inBSE images, and show an intricated network of sutures. Nobaddeleyite was found in this sample. The 21 analyzed spots in 14crystals yielded a discordia line with an upper intercept at 892 ±
Fig. 4.Microscopic aspects of someof the zirconium-rich phases studied. A) andB) baddeleyite from the Pará deMinas swarm samples under crossed polarizers (opticalmicroscope). C) toH) are BSE images (SEM) of baddeleyitewith a thin,micrometric zircon rim (C), baddeleyite (brighter) associatedwith zirconolite (D), and exquisite late-stage skeletal zircon crystals fromsample F5 (E to H).
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Fig. 5. SHRIMP spots on SE images of some of the baddeleyite and zircon crystals studied in the Lavras I (A), Pará de Minas (B, C, D) and Formiga (E, F) dyke samples.
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38 Ma and a meaningless lower intercept at 293 ± 140 Ma. The fourmost concordant spots in four different crystals yielded a Concordiaage of 897 ± 20 Ma (MSWD= 0.18; Prob. of fit = 0.67), interpretedas the best estimate for crystallization age (Fig. 6E). The Th/U ratiosare high, from 0.47 to 1.20.
4.8.2. Sample F6Zircon crystals occur in this sample as euhedral parallelepipeds,
squares, rhomboids, or prisms (Fig. 5F), indicating early-stage crystalli-zation from the magma batch. Square crystals are up to 20 × 20 μm andelongated prisms can reach up to ca. 40 μm × 10–15 μm width. Most
Fig. 6.Concordia diagrams showing the results of SHRIMPgeochronology of the studied samples. In A, C and E the coloured spots indicate data used for the Concordia Ages calculation. In B,D and F all spots were used for the Concordia Ages calculation.
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Fig. 7.Nd isotope evolution diagram for samples from the studied swarms. CHUR=Chondritic Uniform Reservoir; DM=DepletedMantle (DM curve from DePaolo, 1981). São FranciscoCraton basement field after data from Teixeira et al. (1996), Noce et al. (2000) and Barbosa and Sabate (2004).
Fig. 8. Initial Sr x Nd diagram for samples from the studied swarms, compared withcommon reservoirs such as DMM, EMI, EMII and the continental crust (Zindler and Hart,1986).
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crystals are clear and show oscillatory zoning in CL images. All of theanalyzed spots cluster in the Concordia curve (Fig. 6F), yielding aConcordia age of 896 ± 11 Ma (MSWD = 0.04; Prob. of fit = 0.85).Th/U ratios are high, from 0.43 to 1.29.
4.9. Nd\\Sr isotope geochemistry
Nd and Sr isotope data are presented in the Supplementary Material(Table 3) and in the diagrams of Figs. 7 and 8.
4.10. Lavras I swarm
The Nd isotope data for the analyzed samples of the Lavras I swarmyield variable 143Nd/144Nd ratios of 0.5118–0.5120, corresponding toεNd(2.55 Ga) of −6 to +2. The high 147Sm/144Nd ratios of 0.17 precludecalculation of depletedmantle model ages. Initial 87Sr/86Sr ratios, calcu-lated for 2.55 Ga, are in the range of 0.70181–0.70442, corresponding toεSr(2.55 Ga) = 4 to 42. Lavras I norite samples previously presented Ndisotope data with 143Nd/144Nd ratios in the 0.5116–5118 range, corre-sponding to εNd(t) of−3 to+5, and εSr(t) =−16 to 36 (Pinese, 1997).
4.11. Pará de Minas (I and II) swarm
The Nd\\Sr isotope data are presented for the first time for samplesof the Pará de Minas dyke swarms, yielding quite homogeneous143Nd/144Nd of 0.5113–0.5116, corresponding to εNd(1.75 Ga) of −4 to−10 and TDM = 2.5 to 3.0 Ga. Initial 87Sr/86Sr ratios, calculated for1.75 Ga, are in the range of 0.70902–0.71111, corresponding to εSr(1.75Ga) = 93 to 123.
4.12. Formiga swarm
The dated Formiga samples (F5 and F6) yieldedNd isotope datawithεNd(0.9 Ga) of−0.30 and− 0.09 and TDM= 1.4 and 1.3 Ga, respectively.Initial 87Sr/86Sr ratios, calculated for 900Ma, are of 0.70941 and 0.71396,
corresponding to εSr(900 Ma) = 85 and 150, respectively. The F8 sampleis quite different from the others, with negative εNd(0.9 Ga) of−6.59 andTDM = 2.4 Ga, andmuch higher initial 87Sr/86Sr ratio of 0.73089 and εSr(900 Ma) = 390.28.
The IPO1 and JUA1 samples of the Pedro Lessa swarm yieldedεNd(0.9 Ga) of 0.86 and − 1.25 and TDM = 1.4 and 1.6 Ga, respectively,with initial 87Sr/86Sr ratios of 0.70563 and 0.70642 correspondingto εSr(900 Ma) = 31 and 42, respectively. A previous study reports sam-ples from the Pedro Lessa swarm with εNd(0.9 Ga) of −1.9 to 0.5 andεSr(0.9 Ga) of 35 to 49 (Mazuchelli et al., 2000), consistentwith the valuespresented here. Chaves and Corrêa Neves (2005) have found εNd(0.9 Ga)
of −0.3 to −1.0 and TDM = 1.4–1.6 Ga for Formiga and Pedro Lessaswarms, within the range of the values presented here.
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5. Discussion
5.1. Ages of emplacement and correlation to regional rifting events
The southern São Francisco Craton basement shows a protracted Ar-chean tectonic evolution from ca. 3.2 to 2.7 Ga (Teixeira et al., 2017).Upon this basement, important volcano-sedimentary basins developed,startingwith 2.8–2.7 Ga greenstone belts such as the Rio das Velhas andPitangui, followed by the evolution of the typical rift-to-passive marginPaleoproterozoic Minas basin from 2.5 Ga onwards (Alkmim andTeixeira, 2017). These units are superseded by syn-orogenic depositsof the Sabará Group at 2.2 Ga, during the Rhyacian Orogeny that amal-gamated the Archean cores of the São Francisco and Congopaleocontinents into one single coherent paleocontinental mass(Alkmim and Teixeira, 2017 and references therein). ThisPaleoproterozoic continent then underwent distinct attempts ofbreakup and rifting during the Proterozoic, with deposition of extensivecontinental rift successions of the Espinhaço Supergroup (e.g.Guadagnin and Chemale Jr., 2015). Finally, rift- to passive margin- sed-imentary rocks of the Macaúbas Group indicate complete break-up inthe Tonian to Cryogenian, when the paleocontinent was finally sepa-rated carving a partially oceanized gulf-shaped basin that would laterbecome the site of building up of the Araçuaí-West Congo Orogen dur-ing West Gondwana assembly (Amaral et al., 2020). All of these riftingevents were accompanied by extensive development of mafic dykeswarms due to fracturing of the cratonic landmasses providing conduitsfor deep-seated magmas to ascend, and can be traced using the geo-chronological data acquired in this project.
The Concordia age of 2551 ± 9.8 Ma obtained for a noritic dyke ofthe Lavras I swarm suggests that the previously available Sm\\Ndwhole-rock isochron of 2658±44Ma (Pinese, 1997) either does not in-dicate the real crystallization age of this swarm, or that distinct swarmsof Archean age are present in this dyke family. The age of 2.55 Ga couldrepresent crustal stretching related to the opening of thePaleoproterozoic Minas basin over an Archean paleocontinent. In effect,the basal unit of the Minas Supergroup, the Moeda Formation, is com-posed of gold-bearing metaconglomerates and quartzites deposited ina continental rift setting, with younger detrital zircons at ca.2.5–2.6 Ga (Dopico et al., 2017).
The age of the basal units of the Minas Supergroup, including theiron ore deposits within the Cauê Iron Formation is subject of intensedebate (e.g. Cabral et al., 2012). The 2.55 Ga U\\Pb date obtained for amafic dyke of the Lavras I swarm could represent a first robust geochro-nological result to pinpoint the age of crustal riftingwithin the southernSão Francisco basement, thus providing important constraints on theevolution of theMinas Supergroup and chronostratigraphic positioningof the iron deposits within the global framework.
The Pará de Minas dyke swarm is now relatively well constrainedchronologically. Two pulses may be represented at ca. 1795 and1710 Ma (Cederberg et al., 2016;). In either case, these correspond tothe age of mafic and felsic volcanic intercalations at the base of thethick rift-related Espinhaço Supergroup, both in the Serra do Espinhaçoin Minas Gerais and at the coeval Rio dos Remédios Group in Bahia,dated by various workers (Chemale Jr. et al., 2012; Guadagnin andChemale Jr., 2015; Machado et al., 1989). Granitic rocks of theBorrachudos Suite in Minas Gerais and of the Lagoa Real Suite inBahia, all dated at the same age range, represent the plutonic counter-part of this well-developed rift system (Lobato et al., 2015; Silva et al.,2002).
The U\\Pb zircon ages obtained for the Formiga dyke swarm, of ca.900 Ma, confirm a chrono-correlation with the Pedro Lessa dykeswarm, dated at the same age range (Machado et al., 1989). Thesedyke swarms are related to the E5 Tonian rifting event proposed byPedrosa-Soares and Alkmim (2011) and to the Bahia LIP (Chaves et al.,2019). This event seems to have affected the whole São Francisco-Congo paleocontinent, with corresponding dyke swarms on the
African side (e.g. Corrêa-Gomes and Oliveira, 2000) and, in effect, mayhave evolved to a drift phase which essentially delineated theNeoproterozoic paleocontinental margins (Amaral et al., 2020;Salgado et al., 2016). A-type granite intrusions (Silva et al., 2008) andkm-thick bimodal volcanics of the West Congo belt (Tack et al., 2001)are also dated in the same age range, reinforcing the importance ofthis rifting event upon the paleocontinental mass. In addition, thisevent is also marked by Ni-Cu-PGE mineralized mafic-ultramafic intru-sions such as the Brejo Seco Complex in the northern cratonic margin(Salgado et al., 2016), superseded by rift-related basalts of thePaulistana Complex at 888 Ma and finally to new oceanic crust at ca.820 Ma (Caxito et al., 2020). Salgado et al. (2016) interpreted theBrejo Seco layered mafic-ultramafic intrusion as marking the locationof a plume head in the northern cratonic margin, whichwould separatethe São Francisco-Congo paleocontinent from the Borborema-Nigeriaprovince basement to the north during the Tonian (Caxito et al.,2020). A similar evolution is proposed for the other São Francisco-Congo paleocontinental margins (e.g. Amaral et al., 2020; Pedrosa-Soares and Alkmim, 2011).
5.2. Considerations on the elemental and isotope geochemistry data
The Lavras I norites are the most primitive samples of the presenteddataset, both geochemically and isotopically, with higher Mg# and onlyslightly enriched LREE and LILE. The Nd\\Sr isotopic characteristics plotbetween the chondritic values and a composition similar to EMI (Fig. 8),which may suggest variable mixing between depleted mantle (DMM)and an enriched mantle source with similar composition to EMI (seealso discussion in Girardi et al., 2017)..Alternatively, mixing betweenDMM and the Archean continental crust is also a possibility.
In the context of opening of the Minas Basin, which transitionedfrom a rift stage (Moeda and Batatal Formations) to a drift stagewith in-stallment of a fully developed continental margin (Cauê BIFs andGandarela dolostones), the interpretation of installment of a plumehead under a previous Archean paleocontinent, which would generatemantlemagmas that during ascentwould interactwith the thinned lith-osphere of the cratonic margin, is here preferred.
The Pará de Minas samples are distinct from the others for plottingexclusively within the enriched quadrant in the Nd\\Sr diagram(Fig. 8) and presenting the most enriched LREE and LILE patterns. Infact, they plot entirely within the field for typical continental crust,reflecting the importance of crustal contamination in the developmentof this dyke swarm in the area. This indicates that by ca. 1.75 Ga, thecontinental crust was thick enough in this area to produce significantcontamination during magma ascent, or even to remelt and producenew magma batches. This comes as no surprise as the main orogenicevents that formed the basement of the South American platform areconstrained between ca. 2.2 and 2.0 Ga, in a series of Rhyacian orogeniesthatwould contribute to the formation of the SouthAmerican continen-tal crust to a large extent (Brito Neves and Fuck, 2014).
Finally, the Formiga dyke swarm is part of a Tonian rift event whichis represented in the São Francisco Craton margins and within the cra-ton, either as A-type intrusions (Silva et al., 2008), mafic-ultramafic lay-ered complexes (Salgado et al., 2016), dyke swarms (Machado et al.,1989) and gabbroic plutons (Chaves et al., 2019). In effect, this Tonianrift event apparently delineated the São Francisco paleocontinentalmargins as they would interact with the orogens that surround the cra-ton during the Ediacaran-Cambrian Brasiliano Orogeny. Evidence foremplacement of plume heads during the Tonian below the SãoFrancisco lithosphere is discussed by Salgado et al. (2016) on the basisof Nd isotope and trace elements geochemistry data of the Brejo Secomafic-ultramafic complex, and Chaves et al. (2019) interpret thewhole set of Tonian rocks surrounding the craton as representing theBahia-Gangila LIP of continental proportions. Our own trace elementsand Nd\\Sr isotope data are compatible with this interpretation, withplotting of the Formiga/Pedro Lessa samples either close to zero or
Fig. 9. Magmatic barcode for the São Francisco Craton and comparison to the barcodes of other possible adjunct cratons. Black bars show precise U\\Pb dates with the correspondinguncertainties (data sources for the São Francisco Craton and surrounding basement interpreted as part of the Central African Block: 1 - Oliveira et al., 2013; 2 – Barros, 2019; 3 – Caxitoet al., 2015; 4 – Cederberg et al., 2016; Chaves and Rezende, 2019; 5 - Silva et al., 1995; Babinski et al., 1999; Guimarães et al., 2005; Silveira et al., 2013; 6 - Evans et al., 2016; Chaveset al., 2019; 7 – Cederberg et al., 2016; 8 – Santiago et al., 2020), and dates from this work are shown color-coded according to the key in Fig. 1. Grey bars show less precise dates withcorresponding uncertainties based on Sm\\Nd, Rb\\Sr, K\\Ar and 40Ar\\39Ar geochronology (data sources: Pinese, 1997; Carneiro et al., 1998; Teixeira et al., 1998; Chaves, 2001; Oliveira,2004; Battilani et al., 2015; Coelho and Chaves, 2017). Data sources for barcodes from the other cratons: Ernst et al. (2013), Söderlund et al. (2010), Olsson et al. (2010), Peng (2015) andreferences therein. Geological features for the São Francisco Craton from Heilbron et al. (2017) and references therein, and for the North China Craton from Kusky et al. (2007), Lu et al.(2008) and references therein.
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close to the EMII field, probably representing a mixture between DMMand an enrichedmantle source (Zindler and Hart, 1986). The Tonian riftevent would lead to the development of new Cryogenian oceans sur-rounding the São Francisco Craton from all sides (Amaral et al., 2020;Caxito et al., 2020; Pedrosa-Soares and Alkmim, 2011).
Interestingly, the mafic dyke swarms related to extensional eventsconnected to a fully complete continental breakup, leading to the for-mation of new thinned continental passive margins and new oceans,such as the Archean Lavras I and Tonian Formiga swarms, bear the sig-natures characteristic of strong input frommantle sources, with variable
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amount ofmixing between depleted and enrichedmantle. This could bedue to overthinning of the hyperextendedmargins during crustal riftingleading to continental breakup. On the other hand, the 1.75 Ga Pará deMinas I and II dyke swarms are related to the Espinhaço rift system,one of the most extensive and well developed rift-sag systems of east-ern Brazil which would not evolve to a drift stage, bear the typical char-acteristics of mafic dykes generated in a thick continental crust setting,becoming importantly contaminated during ascent.
5.3. Barcode for the São Francisco Craton and comparison to other cratons
Fig. 9 shows an updated barcode for the São Francisco Craton, andtentative correlationswith other cratons. So far, the best-correlated cra-tonic piece is theNorth China Craton,which could have been adjacent tothe São Francisco-Congo paleocontinent from ca. 2.5 Ga to ca. 900 Ma(Cederberg et al., 2016; Chaves et al., 2019; D'Agrella-Filho et al.,2020; Peng, 2015; Peng et al., 2011; Xu et al., 2020).
In the nearby Congo Craton, despite the close proximity to the SãoFrancisco Craton and well documented link, a correspondent to the dis-tinct 1.7–1.8 Ga dyke swarm that characterizes the widespreadEspinhaço rifting has not yet been described. The newly obtained2.55 Ga age for a norite dyke of the Lavras I swarm could suggest a cor-relation with the Zimbabwe Craton due to the age proximity with theGreat Dyke and Cristal Springs dyke swarm emplaced at ca. 2575 Ma(Oberthür et al., 2002), but the absence of the other bars prevent a con-clusive answer.
Fig. 10 shows possible reconstructions for the position of the NorthChina craton relative to the São Francisco-Congo craton as proposedby different authors, both within Columbia (Fig. 10A, B) and Rodinia(Fig. 10C, D, E). Unfortunately, most reconstructions do not show
Fig. 10. Proposed reconstructions for the position of the São Francisco-Congo (SFC) and North CRhyacian Mineiro/Itabuna-Salvador-Curaçá belt of the SFC and Inner Mongolian-Northern HebeStatherian and Tonian rift successions of both cratons. Geological features based on data compiXiong'er plume head position is evidenced by AMS (Anisotropy of Magnetic Susceptibility) studmargin of the North China craton (Xu et al., 2020).
geologic features. Thus, in these figures, the regional geology back-ground was added, highlighting the main areas of location of Archeanbasement and the ubiquitous Rhyacian (ca. 2.2–2.0 Ga) collisionalorogens that join the ancient Archean continental fragments (Zhaoet al., 2002). We choose here the interpretation of Kusky et al. (2007)that the south-verging Inner Mongolian-Northern Hebei orogen thruststhe northern North China craton margin. This orogeny started as an ac-cretionary margin at 2.3 Ga and evolved to an Andean-type and then toa collisional orogeny between 2.2 and 1.9 Ga. This is roughly coeval tothe similar Mineiro belt crosscutting the southern São Francisco Cratonmargin, which also started its evolution at ca. 2.3 Ga as an accretionarybelt and then evolved to a collisional belt at ca. 2.2–1.9 Ga (Teixeiraet al., 2015). This belt then connects to the N-S trending Itabuna-Salvador-Curaçá Orogen of the northern São Francisco Craton(Barbosa and Sabate, 2004) through the reworked basement of theNeoproterozoic Araçuaí belt to the east of the craton (Alkmim andTeixeira, 2017).
Paleogeographic reconstructions of the Columbia (or Nuna; Hoffman,1997) paleocontinent (Zhao et al., 2004) suggest the proximity of theNorth China and São Francisco-Congo cratons due to the location of thesimilar-aged Statherian (ca. 1.8–1.7 Ga) Pará de Minas I and II dykeswarms and Xiong'er volcanic rocks and related dyke swarms (Fig. 10A;Xu et al., 2017, 2020; Chaves and Rezende, 2019). Based on paleomag-netic data, Xu et al. (2017, 2020) suggested that the North China cratonwas positioned between the São Francisco-Congo, Siberia and Rio de LaPlata cratons during the Statherian, and that the Xiong'er LIP representsthe plume head (represented by a red star in Figs. 10, 11 and 12) due tothe radial magmatic flow evidenced by anisotropy of magnetic suscepti-bility (AMS) (Xu et al., 2020). D'Agrella-Filho et al. (2020) provided a pa-leomagnetic pole at 39.8°S, 196.8°E (A95 = 17.0°) for ca. 1790 Ma dykes
hina (NC) cratons during the Statherian (A and B) and during the Tonian (C, D and E). Thei belt of the NCwere added in the paleocontinental settings in this contribution, as well asled from Kusky et al. (2007), Peng et al. (2011) and Heilbron et al. (2017). The ca. 1.78 Gaies that propose a sustained magma flow from amagmatic source centred a the southern
Fig. 11. Alternative scenarios using the positions of the São Francisco-Congo (SFC) and North China (NC) cratons and a former Central African Block or Greater São Francisco-Congo-Sa-haran paleocontinent (Caxito et al., 2020; Cordani et al., 2013; D'Agrella-Filho and Cordani, 2017). Both the ca. 1.75 Ga and ca. 900 Ma plume heads and dyke swarms are represented inorder to test if the positioning of the cratonic pieces as proposed by (A) Chaves and Rezende (2019) and (B) D'Agrella-Filho et al. (2020) would work for both Columbia and Rodinia, thussuggesting a long-lived relationship.
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of the Pará de Minas I swarm, positioning the São Francisco Craton closeto both the North China and Rio de la Plata cratons (the latter crosscutby the ca. 1.79 Ga Florida dyke swarm, for which paleomagnetic dataare available; Teixeira et al., 2013) in the Statherian (Fig. 10B).
Similarly, a North China-São Francisco connection during theTonian (whether within Rodinia or outside of it) was proposed byPeng et al. (2011) and Peng (2015) based on the similar920–900 Ma age range of the Dashigou-Chulan and Bahia-GangilaLIPs (Fig. 10C), and supported by geochemical similarity of maficrocks from those two LIPS (Chaves et al., 2019). However, paleomag-netic constraints at 900Ma from Fu et al. (2015) and at 1.8 and 1.4 Gafrom Xu et al. (2017) suggest slightly different reconstructions(Fig. 10D). Based on the paleomagnetic constraints presentedby Evans et al. (2016) for the Salvador dyke swarm of the northernSão Francisco Craton and by Fu et al. (2015) for the North China Cra-ton at 925–920 Ma, another reconstruction with southeastern NorthChina placed side-by-side with northeastern São Francisco-Congo isused by Chaves et al. (2019). The latter authors proposed two sepa-rated but coeval plume centers, one close to Salvador and the otherto the Xu-Huai rift system, suggesting that a single plume spreadalong the base of the lithosphere and ascended to more than onethinspot producing separated but coeval LIPS.
A different reconstruction (Fig. 10E) was proposed by Cederberget al. (2016), who studied the 1.7–1.8 Ga Pará deMinas I and II swarms.In this reconstruction, the fanning Ilhéus-Salvador-Olivença or Bahiadyke swarm and the Dashigou dyke swarm converge towards a ca.920–900 Ma plume centre located between the São Francisco andCongo Cratons. This reconstruction fits well with the trend of theFormiga dykes, which would also converge to this centre, and is consis-tent with paleomagnetic data at 1.8–1.4 Ga (Xu et al., 2017), but doesnot fit the proposed paleomagnetic comparison of data from Fu et al.(2015) and Evans et al. (2016).
Salminen et al. (2016) presented a paleomagnetic pole for the SãoFrancisco Craton at ca. 1.5 Ga and suggested that the North China-SãoFrancisco connection endured at least until those times surroundingthe core of the Columbia continent. As there is no known shared LIPevent in both cratons during the 1.4–0.9 Ga interval, breakup, rift anddrift and then rejoining of the two paleocontinental masses in differentpositions that would explain the different fits proposed within Colum-bia and within Rodinia (Fig. 10) is unlikely. Additionally, the novelU\\Pb data indicating a 2.5 Ga mafic dyke event in the southern SãoFrancisco Craton fit well with the Taipingzhai-Naouymen dyke swarmof North China, suggesting that these two paleocontinents were joinedfrom the Neoarchean to the Neoproterozoic in a single paleocontinentalmass, whether it was part of Columbia and Rodinia, of only one of thosesupercontinents or none of them.
One problem noted in both Columbia and Rodinia reconstructions isthat generally the cratonic pieces are drawn with their present-dayshapes and borders. Alternative reconstructions, however, suggest vasttracts of Archean-Paleoproterozoic basement stabilized at ca. 2.0 Gaonce surrounded the São Francisco-Congo paleocontinent (Caxitoet al., 2020; D'Agrella-Filho and Cordani, 2017). These vast basementtracts would later be hyperextended and separated from the cratonicmargins during widespread Tonian rifting that occurred in the cratonicborders. Hyperextension removed the tectosphere keels of the individ-ualized blocks allowing for their reworking and intrusion by multipleorogenic plutons during the Neoproterozoic Brasiliano/Pan-AfricanOrogeny (630–500 Ma), but their general affiliation to a former greaterSão Francisco-Congo-Saharan paleocontinent (or Central African Block,according to the nomenclature of Cordani et al., 2013 and D'Agrella-Filho and Cordani, 2017) is evidenced by the identical age distributionof igneous rocks in the craton and in the reworked basement withinthe Borborema-Nigeria-Cameroon, Saharan, Brasília (Tocantins) andMantiqueira provinces of West Gondwana (Caxito et al., 2020). The
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vast tracts of marginal Archean-Paleoproterozoic basement (that wouldlater be decratonized and reworked within the orogenic belts margingthe cratonic borders) have to be represented in paleocontinentalrecontructions (Figs. 11 and 12). We argue that extensional processes,probably related to the emplacement of the ca. 900 Ma plume heads(e.g. Chaves et al., 2019; Salgado et al., 2016) of the Bahia-Gangila LIP,caused the loss of the cratonic keel and disaggregation of the CentralAfrican Block basement in minor microcontinents and ribbon conti-nents, delineating the São Francisco-Congo cratonic core as the survivalpiece of this process.
Alternative scenarios for the paleogeographic positioning of the SãoFrancisco-Congo and North China cratons using the notion of a formerCentral African Block (Caxito et al., 2020; Cordani et al., 2013;D'Agrella-Filho and Cordani, 2017) are presented in Fig. 11. In those sce-narios, the continuation of ca. 2.2–2.0 Ga orogens is also represented inthe basement of the Borborema-Nigeria-Cameroon Province of theBrasília Belt in the Tocantins Province and in the Mantiqueira Province,as part of the extensive network of worldwide Rhyacian orogens (Zhaoet al., 2002) represented in Fig. 12.
In the reconstructions of Fig. 11, both the proposed plume heads forca. 1.79 Ga dyke swarms (red stars) and for ca. 900 Ma dyke swarms(green stars) are represented, in order to test if the proposed recon-structionswouldwork for both the Statherian and Tonian, thus evidenc-ing a long-lived link between the two cratonic pieces without driftingand dispersal of the individual blocks between the Statherian andTonian (as suggested by the reconstructions of Salminen et al., 2016 atca. 1.5 Ga). In general, the radial patterns of dyke emplacement wouldwork both for the reconstruction of Chaves and Rezende (2019) andfor that of D'Agrella-Filho et al. (2020), with the North China craton ad-joining the future basement to the Brasília Belt that fringes the westernmargin of the craton. In the first case (Fig. 11A), the Inner Mongolian-Northern Hebei belt would represent a continuation of the Cristalândiado Piauí and Almas-Dianópolis blocks of central Brazil, fringing thenorthwestern São Francisco craton margin and presenting Archeanbasement highly reworked by the ca. 2.2–2.0 Ga Orogeny (Heilbronet al., 2017). In the second case (Fig. 11B), the Inner Mongolian-North-ern Hebei belt would be the continuation of the Mineiro Belt in thesouthern São Francisco Craton margin, which presents very similarSiderian accretionary processes and Rhyacian to Orosirian collisionalprocesses (Teixeira et al., 2015), although differential uplift has exposedthe granulitic roots of the Inner Mongolian-Northern Hebei orogen(Kusky et al., 2007). The thrusting direction (cratonwards) and the geo-chronological fit would work for both reconstructions.
Finally, in Fig. 12, we present two alternative reconstructions for Co-lumbia based on the proposals of Chaves and Rezende (2019) (Fig. 12A)and D'Agrella-Filho et al. (2020) (Fig. 12B). The Columbia reconstruc-tions of these authors use basically the same positioning for the otherconstituent cratons of other proposals (Bispo-Santos et al., 2014;Evans and Mitchell, 2011; Johansson, 2009; Pesonen et al., 2012; Xuet al., 2014, 2017; Zhang et al., 2012), but modifying the positioning ofthe North China, São Francisco-Congo and Rio de La Plata cratons (forother reconstructions see Fig. 4 of Chaves and Rezende, 2019).
Unlike all of the proposed reconstructions for Columbia, inFig. 12A and B we use the concept of a Central African Block withthe São Francisco-Congo Craton at its core (Caxito et al., 2020;Cordani et al., 2013; D'Agrella-Filho and Cordani, 2017). Thus, thetwo reconstructions were modified in order to produce the space
Fig. 12. Two alternative reconstructions of Columbiawith the position of the São Francisco-CongFilho et al. (2020). Unlike all of the proposed reconstructions for Columbia, we use the conceCordani, 2017) with the São Francisco-Congo Craton at its core, which encompassed the furtheTocantins and Mantiqueira orogenic provinces of Brazil (Caxito et al., 2020). This Central Africastructions have beenmodified in order to produce the space necessary to fit the Central Africanto the east of Australia to eastern paleolongitudes in the case of (A) and ofWest Australia, Nortcraton in the southern hemisphere was also considered in order to simulate its positioning in abased on Zhao et al. (2002) with our own additions.
needed to fit the large Central African Block. This has been donethrough displacement of the other cratonic blocks only in the lon-gitudinal coordinate, respecting the proposed paleolatitudeswhich are generally based on paleomagnetic data. In Fig. 12A, thecratonic blocks west of Siberia were displaced to westernpaleolongitudes, and those east of Australia were displaced to east-ern paleolongitudes. In Fig. 12B, only West Australia, NorthAustralia and the Rio de La Plata craton had to be slightly displacedlongitudinally. In Fig. 12A, the alternative hemisphere positioningof the Kalahari craton (i.e., in the same paleolatitude as proposed,but under a magnetic field with flipped polarity, thus plotting inthe southern hemisphere instead of the original proposition inthe northern hemisphere) is considered in order to simulate itsprobable inclusion in a greater Central African Block. As shown bythe match between the various Rhyacian-Orosirian orogenic belts(Zhao et al., 2002) and radial patterns of dyke swarms, both recon-structions would work well at this point. Refined geochronologicaland paleomagnetic data, especially for the southern hemispherecratons, is necessary in order to further interpret the best-fit prop-ositions. For the moment, alternative reconstructions, such asthose presented by D'Agrella-Filho and Cordani (2017) where theCentral African Block was neither part of Columbia nor Rodinia,are equally feasible.
6. Conclusion
The presented U\\Pb SHRIMP data on zircon and baddeleyite pin-point emplacement of dyke swarms in the southern São Francisco Cra-ton at ca. 2551 Ma (Lavras I swarm), 1795 and 1710 Ma (Pará deMinas I and II swarms) and ca. 896 Ma (Formiga swarm).
Previously dated samples through geochronological methods thatare more prone to secondary resetting (e.g. Sm\\Nd, Rb\\Sr and40Ar\\39Ar) suggest the presence of aminor Lavras II dyke swarm of un-certain age and the Paraopeba swarm, presumably at ca. 2.2–2.0 Ga. Ro-bust, precise and accurate U\\Pb geochronological data are necessary tofurther constrain the true age of emplacement of these dyke swarms,and this is clearly a target for future studies.
The obtained dates can be linked to important extensional (rift)events generating major sedimentary basins over the precursor of theSão Francisco paleocontinent, such as the Minas Basin at ca. 2.55 Ga,the Espinhaço rift basin system at ca. 1.75 Ga, and the Macaúbas I riftat ca. 900 Ma. Thus, the obtained ages serve as markers of importantevents such as deposition of the major Lake Superior-type BIFs of theCauê Formation in the Minas Basin and to the east of the EspinhaçoRange. In addition, all of these dyke swarmsmay have disturbed impor-tant Au deposits in the Quadrilátero Ferrífero region.
Lithogeochemical and Nd\\Sr isotope data suggest distinct petro-genesis and sources for the distinct dyke swarms. The Archean Lavras Iand the Tonian Formiga dyke swarms show importantmantle contribu-tions, perhaps indicating proximity of a plume head and mixing of de-pleted mantle (DMM) and enriched mantle (EMI and EMII) sources.The Sthaterian Pará de Minas I and II swarms clearly show importantcontribution from the ancient continental crust of the southern SãoFrancisco Craton.
When comparing the presented barcode for the São Francisco Cratonwith that of other cratonic blocks, striking similarities with the NorthChina Craton suggest those two cratonic blocks travelled together
o andNorth China cratons as proposed byA) Chaves and Rezende (2019) andB)D'Agrella-pt of a Central African Block (Caxito et al., 2020; Cordani et al., 2013; D'Agrella-Filho andr decratonized Saharan lithosphere and reworked basement blocks within the Borborema,n Block may also have involved the Kalahari craton of southern Africa (A). The two recon-Block, with displacement of cratonic blocks west of Siberia to western paleolongitudes andh Australia and Rio de la Plata in the case of (B). The alternative positioning of the Kalaharigreater Central African Block in (A). The probable positioning of Rhyacian orogenic belts is
17F.A. Caxito et al. / Lithos 374–375 (2020) 105708
throughout most of the Proterozoic, perhaps participating in majorpaleocontinental arrangements such as Columbia and Rodinia. Compar-ison with other cratonic blocks of the southern hemisphere, especiallythose in Africa should be possible, following the development of a re-fined geochronological record.
Last but not least, the vast tracts of reworked Archean-Paleoproterozoic basement which are now embedded as terraneswithin the Neoproterozoic orogenic belts that surround the cratonsshould be considered in paleogeographical reconstructions, before abest fit is attempted using directional dyke information, paleomagneticdata, and continuation of features such as orogenic belts and rift basins.The São Francisco-Congo craton, for example, was probably surroundedby vast basement tracts (the Borborema-Nigeria-Cameroon, Tocantins/Brasília and Mantiqueira provinces basement, and probably the“metacratonized” basement of the Saharan region), forming a CentralAfrican Block or Greater São Francisco-Congo-Saharan paleocontinent.This is the paleocontinental piece that should be compared to the posi-tioning of other pieces such as the North China and Rio de la Plata cra-tons in Proterozoic paleocontinental reconstructions.
Declaration of Competing Interest
None.
Acknowledgments
This project was supported by CNPq (Brazilian Research Council)through a post-doctoral research grant (nb. 201355/2018-3) to FAC atthe University of Western Australia. FAC is also a Research Fellow ofCNPq and thanks for the ongoing support. TheU\\Pb analyseswere per-formed using a SHRIMP II probe at the John DeLaeter Centre of the Cur-tin University, Perth, Western Australia, enabled by NCRIS via AuScope.We thank Neal McNaughton for guidance and support during the prep-aration and analysis of the SHRIMPmounts and Allen Kennedy and HaoGao for support during the analysis, and the UWA/CET staff and alumni,especially Andreas Petterson, Steve Denyszyn and Luis Parra-Avila forthe help and discussions. We also thank Cristina Burgos and all of thestaff of the Brazilian Geological Survey in Salvador, Brazil, for the tipsand help in baddeleyite separation. FAC thanks João Orestes SchneiderSantos for the support and incentive to conduct a post-doctoral researchterm in the geochronology facilities of UWA and Curtin University. Anearlier version was greatly improved after comments and suggestionsby R. Ernst and an anonymous reviewer.
Appendix A. Supplementary data
Supplementary data to this article can be found online at https://doi.org/10.1016/j.lithos.2020.105708.
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