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ges of protolith and Neoproterozoic metamorphism of Al-P-bearinguartzites of the Veredas formation (Northern Espinhac o, Brazil):A-ICP-MS age determinations on relict and recrystallized zircon andeodynamic consequences
erhard Franza,∗, Giulio Morteanib, Axel Gerdesc,d, Dieter Rhedee
Fachgebiet Mineralogie-Petrologie, Technische Universität Berlin, ACK 9, Ackerstraße 71-76, D 13355 Berlin, GermanyGmain Nr. 1, D 84424 Isen, GermanyInstitut für Geowissenschaften, Goethe Universität Frankfurt/M, Altenhöferallee 1, D 60438 Frankfurt/M, GermanyDepartment of Earth Sciences, Stellenbosch University, Private Bag X1, Matieland 7602, South AfricaDeutsches GeoForschungsZentrum Potsdam, Telegrafenberg, Sektion 4.2, D 14473 Potsdam, Germany
r t i c l e i n f o
rticle history:eceived 11 April 2013eceived in revised form 14 March 2014ccepted 6 May 2014vailable online 27 May 2014
eywords:luminium-phosphate quartzitespinhac oão Francisco Cratoneredas formationenotime formation from zircon
a b s t r a c t
The N-S running Espinhac o fold belt, situated in the São Francisco craton in central-eastern Brazil,separates this craton in a western and an eastern part along the Paramirim Corridor, a Brasiliano-Panafrican (∼650–500 Ma) deformation zone. The (meta)sediments of the Espinhac o basin consist ofseveral synthems with dominantly siliciclastic deposits. Among them, Al- and P-rich, lazulite-bearingblue quartzite, the amphibolite-facies equivalent of a phosphatic sandstone, of the Veredas Forma-tion in the Northern Espinhac o form a conspicuous horizon. Zircon crystals from these quartzites havebeen dated using laser ablation inductively coupled plasma mass spectrometry. The zircon crystalsare both detrital and metamorphic in origin, the latter exhibiting highly unusual textures producedby dissolution–reprecipitation reaction and concomitant replacement of metamict areas: (REE, Y, Th,U) zircon + fluid = zircon + xenotime. Detrital zircons reveal relict age groups of 3.1–3.3 Ga, 2.6–2.7 Ga,2.0–2.3 Ga, and 1.5 Ga. The end of sedimentation and erosion of the magmatic and metamorphic rocks ofthe Sao Francisco craton can be estimated therefore as ≤1.5 Ga. The major metamorphic event recordedby reprecipitated zircon is Neoproterozoic and documented by a lower-intercept age of 634 ± 19 Ma,
which was calculated from analyses of individual spots on19 grains with a common relict Orosirian-Rhyacian age of 2032 ± 31 Ma, temporal equivalent to Transamazonian. The Neoproterozoic age of themetamorphosed zircons of the Vereda formation corresponds to the Brasiliano-Panafrican remobilizationof the São Francisco-West Congo craton along the eastern side of the Espinhac o fold belt in the ParamirimCorridor.
. Introduction
Detrital zircon geochronology has become a powerful tool fornravelling the history of sedimentation in siliciclastic rocks andheir metamorphic equivalents. Furthermore, recent work hashown that zircon can recrystallize in metamorphic rocks, event low grade or hydrothermal conditions (e.g., Balan et al., 2001;
eisler et al., 2003; Tomaschek et al., 2003; Rasmussen, 2005;elattre et al., 2007; Tropper et al., 2007; Hay et al., 2010; Vorhiest al., 2013). With the introduction of sensitive radiometric age
determinations working in situ in thin section, it became possi-ble to accurately determine crystallization ages of small crystals, ofreplaced areas in recrystallized grains, or overgrowths on detritalzircon. We applied LA-ICP-MS (laser ablation inductively coupledplasma mass spectrometry) on zircon crystals to a rare type ofmetapsammitic rock, a metamorphosed Al- and P-rich quartzitefrom the Veredas Formation, Northern Espinhac o fold belt, Brazil, inwhich detrital and metamorphosed zircon occurs closely together.The Espinhac o fold belt is an inverted rift basin in the São Franciscocraton in central-eastern Brazil. The Al-P-quartzite contains a con-
siderable amount of zircon with textures typical of corrosion ofold detrital zircon and crystallization of young metamorphic zirconby a fluid-driven dissolution–reprecipitation process. The associa-tion with less altered sedimentary zircons offered the possibility
G. Franz et al. / Precambrian Research 250 (2014) 6–26 7
Fig. 1. (a) Tectonic sketch map of the Paleoproterozoic São Francisco craton, surrounded by Neoproterozoic (Brasiliano) fold belts (Bfb) and crosscut by the ParamirimCorridor. Rocks of the Espinhac o Supergroup are well exposed in the Diamantina Plateau (DP) and the Northern Espinhac o (NE), which is part of the Espinhac o fold belt,subdivided into the Northern, Central (CE) and Southern Espinhac o (SE). Black square shows location of (b). Simplified after Danderfer et al. (2009), Cordani et al. (2010). (b)Geological sketch map of the Serra da Veredas, compiled from Tavora et al. (1967), Santos and Souza (1983), Loureiro and Souza (1985), Schobbenhaus (1996), Danderfer(2002), Danderfer and Dardenne (2002) and own field work; asterisks show sample localities, in bold for zircon dating.
8 ian Research 250 (2014) 6–26
trthDat1B
2
wnEs‘GS
tg1See2TtPasiWRRBE1dAtdt(
tp2dPlsccmeaBA
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Fig. 2. (a) Photograph of a cut surface of the Al-P quartzite in the outcrop ofthe quarry Conservice (see Fig. 1b). The blue layers are enriched in lazulite anddumortierite and indicate the sedimentary layering, also shown as crossbedding (leftof the hammer handle); arrows indicate sedimentary discontinuities later enriched
G. Franz et al. / Precambr
o date the age of metamorphism as well as the age of the sourceocks of the zircons, thereby constraining the age of sedimenta-ion. Until now, all age determinations for the Espinhac o fold beltave been exclusively produced from intercalated igneous rocks.irectly dated metasediments have not yet been reported and thege of metamorphism was not well constrained. We show thathe Al-P-quartzite contains detrital zircons from 3.3 Ga down to.5 Ga. The later metamorphic rejuvenation occurred during therasiliano-Panafrican at 634 ± 19 Ma.
. Geologic setting
The Serra do Espinhac o fold belt (Schobbenhaus, 1996) togetherith the Diamantina Plateau (Chapada Diamantina) are promi-ent geological and physiographical features of central Brazil. Thespinhac o fold belt extends approximately 1000 km N–S from theouthern border of the States of Piauí and Pernambuco to theQuadrilátero Ferrífero’ (‘Iron Quadrangle’) in the State of Minaserais (9◦S–20◦S). It is subdivided into the Northern, Central andouthern Espinhac o (Fig. 1a).
The geological history of the Espinhac o fold belt is closely relatedo that of the São Francisco-Congo craton. The gneisses, migmatites,ranulites and charnockites of the São Francisco craton (Almeida,977; Santos and Souza, 1983; Lopez and Souza, 1985; Loureiro andouza, 1985, see also review in Alkmim et al., 2006) record oldervents such as the Jequié-Aroan tectonomagmatic cycle (Almeidat al., 1978; Brito Neves et al., 1990; Barbosa and Sabaté, 2002,004) at about 2.6 and 2.7 Ga. The craton was consolidated in theransamazonico event (Orosirian-Rhyacian ∼2.0 Ga) together withhe West Congo craton and locally rejuvenated, for example in thearamirim Corridor, by the Neoproterozoic Brasiliano event (equiv-lent to Panafrican). The São Francisco craton (Almeida, 1977) isurrounded by several Brasiliano fold-and-thrust belts, which are:n the southeast the Arac uaí belt as part of the Arac uaí orogen-
est Congo belt (see Babinski et al., 2012), in the south the Altoio Grande belt, in the west the Brasilia belt, and in the north theio Preto and the Riacho do Pontal belts (Fig. 1a; compiled afterarbosa and Sabaté, 2004; Danderfer et al., 2009). The Northernspinhac o belt (Serra do Espinhac o Septentrional; Schobbenhaus,996) forms the western flank of the so-called Paramirim Corri-or, a north-south trending belt of Brasiliano age (650–500 Ma;lkmim et al., 1993; Schobbenhaus, 1996; Danderfer, 2002). In
he Central and Southern Espinhac o, the fold belt is also bor-ered in the west by rocks of the São Fracisco craton, but inhe east it is overthrusted by rocks of the Brasiliano Arac uaí beltFig. 1).
The rocks of the Serra do Espinhac o belong dominantly tohe Espinhac o Supergroup and consist of eight Paleo- and Meso-roterozioc synthems (Danderfer, 2002; Danderfer and Dardenne,002; Danderfer et al., 2009). The mostly clastic sediments of theifferent synthems have been deposited in rift basins during severalaleo- to Mesoproterozoic stages. By dating zircon from interca-ated anorogenic volcanics and granites (Danderfer et al., 2009),edimentation was determined to have begun at 1731 ± 5 Ma, thusonfirming previous assumptions that rifting on the São Franciscoraton occurred at ∼1.7 Ga. Basin inversion and the age of meta-orphism are constrained as Brasiliano. The so-called ‘Espinhac o
vent’ supposedly includes a metamorphic overprint, for whichges between 1.0 Ga and 1.5 Ga are given (Inda and Barbosa, 1978;rito Neves et al., 1979; Turpin et al., 1988; Cordani et al., 1992;lmeida-Abreu, 1995).
In the Northern Espinhac o Supergroup, the metamorphosedhosphatic sandstone forms a characteristic horizon within theeredas Formation, part of the Sitio Novo Synthem (Fig. 1b).hosphatic sandstones are typically interpreted as Precambrian
by fluid interaction in phosphate and borate minerals. (b) Photograph of a cut hand-specimen; white is quartzite matrix, thin blue layers are enriched in lazulite anddumortierite and show the metamorphic foliation.
shallow-platform sediments (e.g., Vallini et al., 2005; Morteaniet al., 2007). The most prominent metamorphic minerals in thesequartzites are lazulite, a (Mg, Fe)-Al phosphate, the mineral groupcrandallite-svanbergite-goyazite, which are solid solutions of (Ca,Sr, REE)-Al phosphates-sulphates, and, in the presence of B, theAl-borosilicate dumortierite. Cassedanne and Franco (1966) andCassedanne et al. (1989) were the first to describe these rocks.Fleischer (1971), Morteani et al. (2007) and Jordt-Evangelista andDanderfer (2012) discussed the petrogenetic aspects of the bluequartzites of Northern Espinhac o. The Al-P quartzites occur as anapproximately 30 m thick horizon (Fleischer, 1971) in the up to2500 m thick sequence of siliciclastic rocks of the Veredas Forma-tion, the lowermost formation of the Sitio Novo Synthem with anestimated total thickness of 12,000 m (Danderfer et al., 2009). Thesemassive and hard rocks display a beautiful colour contrast betweenthe white quartzitic matrix and blue layers rich in lazulite anddumortierite (Fig. 2). They are mined extensively all along the Serrada Vereda. Lopez and Souza (1985) and Loureiro and Souza (1985)show a total of 16 quarries in the “Carta Metalogenetica 1:250,000”.The dimension stones are sold with the commercial names “Azulde Macaúbas”®, “Azul Imperial”® or “Azul de Boquira”®. From theSouthern Espinhac o, Horn et al. (1996, 2000) and Morteani et al.(2001) described similar rocks.
3. Sample material
The samples have been collected in four quarries in the Veredasmountain ridge, located west of the towns of Boquira and Macaúbas
and part of the Northern Espinhac o fold belt: Conservice, São Marco,Sitio Novo and Vaca Morta (Fig. 1b, Table 1). The latter two aresituated south of Boquira and represent the southernmost samp-ling point of our study, but the horizon extends much further
G. Franz et al. / Precambrian Research 250 (2014) 6–26 9
Table 1Sample list and whole rock major and trace element composition.
Locality Sitio Novo Vaca Morta Conservice São Marco
Sample # 5859 5864 A 5868 B 13562 13567 13560 A 13584 BRock type lazulite-
lank = not determined; – below detection limit; Cu, Pb, U below detection limit of
h 5 ppm.
outh-southeast (Morteani and Ackermand, 2004). All samplesere taken from active quarries and do not show any sign of weath-
ring or hydrothermal overprint. Seven samples were selectedor screening age determinations by electron microprobe analy-is (EMPA) on monazite and xenotime. The results showed a wideange of ages between ∼2.3 Ga and ∼200 Ma for monazite and2.85 Ga and ∼350 Ma for xenotime (Inline Supplementary Table1, Fig. S1). The data can be divided into two age groups, an oldne with dominant ages around 2.1 Ga down to 1.5 Ga (including
few higher ages up to 2.85 Ga), and a young age group with agesetween ∼900 Ma and ∼450 Ma. The large errors are due to the gen-rally very small and irregular grain size (two examples shown innline Supplementary Fig. S2), which makes the crystals very sensi-ive to retrograde reequilibration by diffusion. Two samples with aarge number of zircon crystals and the highest Zr content (Table 1:868B from Vaca Morta, 13584B from São Marco) were thereforeelected for subsequent LA-ICP-MS dating on zircon.
Inline Supplementary Table S1 can be found online atttp://dx.doi.org/10.1016/j.precamres.2014.05.011.
Inline Supplementary Fig. S1 can be found online atttp://dx.doi.org/10.1016/j.precamres.2014.05.011.
1.93 1.67 0.73
pm; other detection limits: oxides ≈0.01 wt%; S 0.002 wt%; As, Ba 20 ppm; Mn, Rb,
Inline Supplementary Figs. S2 can be found online athttp://dx.doi.org/10.1016/j.precamres.2014.05.011.
4. Methods
Zircon grains were analyzed in situ in polished thin sec-tions for U, Th, and Pb isotopes by LA-ICP-MS techniques at theGoethe-University Frankfurt/M (GUF) using a Thermo-FinniganElement II sector field ICP-MS coupled to a New Wave UP213UV laser system (first analytical sessions in 2006) following themethod described in Gerdes and Zeh (2006, 2009) and Frei andGerdes (2009). Selection of spots for analysis was guided by internalstructures as seen in BSE (back-scattered electron) and CL (cathodo-luminescence) images while considering the 20 �m diameter and∼15 �m penetration depth of the laser. A teardrop-shaped, lowvolume laser cell (Horstwood et al., 2003) was used to enablesequential sampling of heterogeneous grains (e.g., growth zones)
during time resolved data acquisition (see Janousek et al., 2006).The second batch of analyses (Table 4c) have been acquired in 2011at GUF with a Resolution M50 193 nm ArF Excimer laser (Com-pexPro 102, Coherent) equipped with two-volume ablation cell
Fig. 3. BSE images from sample # 13584B, Sao Marco quarry. Two sites in thin sec-tion (a and b) show detrital and metamorphic zircon crystals near to each other.The crystals sit in a matrix of polygonal quartz (Qtz), with muscovite (Mu) prefer-ably along quartz grain boundaries, accompanied by hematite (Hem), rutile (Rut)and rare phosphate (berlinite Ber; black areas in (a) near berlinite and in (b) nearmuscovite are outbreaks in thin section). Notable is a high micro-porosity, espe-cially well visible in (a). The zircon crystals also show the craters produced by laserablation for age determination.
Fig. 4. (a and b) BSE images of zircon crystals from the two sites in polished thin sectiabsence of a regular zoning pattern, the irregular shape with abundant inclusions of xenoas products of dissolution and recrystallisation of a euhedral or rounded precursor zircon.form classifies the grains 104 and 108 as relict detrital zircons. Zircon outgrowths and begiand well-preserved detrital crystals are not completely immune against dissolution and
indicating mobility of Zr. Dashed circles and squares indicate position for laser ablation (a
search 250 (2014) 6–26
(Laurin Technic, Australia) using a slightly modified method (Zehand Gerdes, 2012). Spots of 10–26 �m in diameter were fired with5.5 Hz at a fluence of 4 J cm−2, which yielded at a spot size in zir-con of 26 �m and depth penetration of 0.7 �m s−1 a sensitivity ofabout 13,000 cps/�g/g for 238U. Each analysis consisted of approx-imately 20 s of background acquisition followed by 22 s of dataacquisition. Correction for common Pb was applied whenever thecorrected 207Pb/206Pb was outside the internal error of the uncor-rected ratio and was based on the interference- (e.g., 204Hg, usingthe measured 202Hg and the natural 204H/202Hg) and background-corrected 204Pb signal and the two-stage terrestrial Pb evolutionmodel of Stacey and Kramers (1975). Reported uncertainties (95%confidence level) were propagated by quadratic addition of theexternal reproducibility (2 SD; standard deviation) obtained fromthe standard zircon GJ-1 (∼1.6% and 1.2% for the 207Pb/206Pb and206Pb/238U, respectively; n = 12) during individual analytical ses-sions and the within-run precision of each analysis (2 SD; standarderror). Repeated analyses of the reference zircon Plesovice, 91500,and Felix (Slama et al., 2008; Millonig et al., 2012) during the sameanalytical session yielded an accuracy of better 1% and a repro-ducibility of <2% (2 SD). All uncertainties are reported at the 2sigmalevel. Concordia diagrams (2� error ellipses), concordia ages andupper intercept ages (95% confidence level) were calculated usingIsoplot/Ex 3.00 (Ludwig, 2003).
Mineral analyses were performed in order to estimate metamor-phic conditions. We applied the Zr-in-rutile (Watson et al., 2006)and the Ti-in-quartz geothermometers (Wark and Watson, 2006).The analytical conditions for the JEOL JXA-8500F (HYPERPROBE)electron microprobe at the GFZ (Deutsches GeoForschungsZen-trum, Potsdam, Germany) included an accelerating voltage of 15 kV,a beam current of 30 nA and a beam diameter of one �m. Withcounting times of 500 s on the peak and 2 × 250 s on backgroundfor both trace elements, we attained a detection limit of 20 �g/g(Zr L� PETH) and 7 �g/g (Ti K� PET). In order to check for possibleimpurities we simultaneously measured Si, Fe, Ti and Nb in rutileand Al, Fe, Mg and Zr in quartz. Simultaneous analysis of two in-house reference samples (QTiP-7; QTiP-38) was used to check the
analyses of the samples. Zircon (Zr and Si), rutile (Ti), periclase (Mg),orthoclase (Al), hematite (Fe) and Nb-metal (Nb) were used for thecalibration and for quartz the theoretical values of 53.24 wt% O and
on where relict and metamorphic grains occur next to each other (cf. Fig. 3). Thetime (bright) and matrix minerals (black) identifies the zircon grains 105 and 109
The absence of xenotime or matrix mineral inclusions and the zoning and euhedralnning replacement along the rim of the detrital grains show that also homogeneousreplacement processes. Zircons are also found as nano-sized crystals in the matrix,pparent discordant Pb-Pb ages, relict ages concordant; in Ma).
G. Franz et al. / Precambrian Research 250 (2014) 6–26 11
F pointt ent adl
4w
TSgwp
5
5
l
TR
1
ig. 5. Cathodo luminescence images of (a) a largely pristine, relict zircon; arrowshe core, also on the lower tip. (b) Metamorphic crystal with a relict core; replacemow and high luminosity.
6.76 wt% Si were included in the CITZAF corrections procedure,hich is based on the ˚(�Z) method (Armstrong, 1995).
Major and trace elements of whole rocks were analyzed at theechnical University of München (Germany) by a fully automatediemens XRF (X-ray fluorescence) instrument using Li-tertraboratelass disks. Rare earth elements (REE), U, Th and B analysesere performed on commercial basis by ICP-MS (induced coupledlasma-mass spectrometry) by the X-Ral Co. (Canada).
. Results
.1. Petrography and whole rock composition
Deep blue lazulite and, in some samples, deep blue to vio-et dumortierite, slightly greenish kyanite, and white mica are
able 2epresentative electron microprobe analyses of zircon.
) calculated on the basis of 16 oxygen 2) includes minor amounts of other cations; Al2O3
to replacement areas along oscillatory zoning and to patchy replacement areas invances along cracks; arrows point to borders of areas in metamorphic zircon with
enriched in internally folded and strongly tectonically sheared lay-ers and patches. Patches and layers of hematite, often surroundedby dumortierite and lazulite-scorzalite enriched rims, are typicalfor the quartzites (Cassedanne et al., 1989). These aggregatesare ascribed to boudinage of former hematite-rich layers. Inrare cases, cross-bedding, marked by lazulite and dumortierite,is preserved (Fig. 2a) and sedimentary bedding is recognizableunder the microscope by distinctive layers of heavy minerals suchas hematite, rutile, zircon, monazite and xenotime. There is noindication in the sampled rocks for hydrothermal alteration. Inthe outcrop, enrichment of lazulite and/or dumortierite along the
sedimentary layering could occasionally be observed (Fig. 2a).A weak foliation is developed oblique to the bedding (Fig. 2b).In addition to lazulite-scorzalite (Mg,Fe)Al2[PO4]2(OH)2, otherrock forming phosphates are present and comprise complex solid
Fig. 6. BSE-images of different textural types of metamorphic zircon; numbers atthe laser ablation spots are apparent ages in Ma, insets show the grains beforeablation. (a) Fractured crystal, with a dominantly pristine core and a metamorphicouter zone, with xenotime and mica inclusions (inset at lower right); the frag-ment right of the crystal shows the same age (∼2500 Ma) as the central grain, thefragment in the upper left a significantly lower age (∼890 Ma). (b) Pseudomorphicgrain, consisting of variably sized zircon crystals plus goyacite-svanbergite (Al-Sr-P)in quartz-muscovite, intergrown with small rutile. (c) Skeletal zircon crystal with
n = number of measurements; numbers in parenthesis are 1� standard deviationsf the last digits).
olutions of the crandallite – svanbergite – goyazite mineral groupaAl3[PO4]2(OH)5·H2O–SrAl3[PO4/SO4](OH)6–SrAl3[PO4]2(OH)5·2O, as well as augelite Al2[PO4](OH)3, trolleite Al4[PO4]3(OH)3,erlinite AlPO4, apatite, and amblygonite LiAl[PO4](F, OH)Morteani and Ackermand, 2004).
Whole rock compositions including the trace element contentsf the samples are listed in Table 1. The two samples for LA-ICP-S-dating were selected based on their high Zr content indicating
large amount of zircon crystals. The samples are B-poor with10 ppm B (Vaca Morta; sample 5868B) and with 29 ppm B (Sãoarco; sample 13584B), with negligible amounts of tourmaline
nd/or dumortierite and belong to the B-poor rock type (Morteanind Ackermand, 2004). Jordt-Evangelista and Danderfer (2012)istinguished among others a dumortierite-kyanite-muscovite-uartzite from a lazulite-trolleite-quartzite, and our dated samplesost probably belong to the second type. The concentration of P2O5
n both samples (9.37 and 4.52 wt%, respectively) is much higherhan in common clastic sediments, thus classifying them as formerhosphatic sandstones. Samples with such a high P-content showEE distribution patterns that are enriched in heavy REE (see Fig.
in Morteani and Ackermand, 2004) and slightly depleted or notepleted in light REE (normalized to North American Archaeanhale). The complexity of textures in outcrop (Fig. 2) in combi-ation with the equilibrium boundaries of the quartz grains anduartz-mucovite grain boundaries (Fig. 3; see also Inline Supple-entary Fig. S3, and Jordt-Evangelista and Danderfer, 2012, their
ig. 2D) documents a multistage deformational history with a finalervasive static recrystallisation. BSE images (Fig. 3) also show aigh amount of small pores aligned along grain boundaries but alsoithin quartz.
Inline Supplementary Fig. S3 can be found online atttp://dx.doi.org/10.1016/j.precamres.2014.05.011.
.2. Zircon
.2.1. Textures and crystal chemistryTwo major types of zircon can be distinguished based on tex-
ure: relict (detrital) and metamorphic. Fig. 3 gives an overviewf the textural position of the zircons and shows that both typesre found near each other, only a few �m apart. They are sit-
ated in a matrix of polygonal quartz, together with muscovite,ematite and rutile; phosphate minerals are present in the sam-les, but not necessarily in contact with zircon. The relict zirconsre generally rounded (Fig. 4a; grain 104) or euhedral with rounded
abundant xenotime (bright), muscovite (black, lath-shaped) and quartz inclusions.Different contrast of the small and irregularly shaped xenotime crystals is causedby variable REE contents.
G. Franz et al. / Precambrian Research 250 (2014) 6–26 13
Fig. 7. BSE image of a metamorphic zircon grain with abundant inclusions (sample 13584B, São Marco quarry) showing two laser ablation spots, which yielded apparentdiscordant Pb-Pb ages (given in Ma) and an upper intercept age of 2092 ± 66 Ma. To better highlight the xenotime inclusions the inset at lower right shows the xenotimeinclusions at a contrast that is different in respect to that of the overview. Heterogeneities in the brightness of the xenotime grains (upper and middle inset, enlarged image)are due to differences in the content of rare earth elements. Rutile, hematite and zircon form a contact paragenesis. Rutile crystals from such parageneses were used fora
ezgctxgtraeFecicanl
tprcqtgL(odrfzosmzs
(Figs. 6b and 8) and textures such as growth on metamorphic zir-con, intergrowth with hematite and negative crystal faces towardsquartz (Fig. 8) suggest their equilibration during metamorphism.In such crystals we determined the Zr content (Table 3) in order
pplication of the Zr-in-rutile geothermometer.
dges and are characterized in BSE images by a mostly concentriconing pattern (Fig. 4b; grain 108). In some crystals the pattern sug-ests a well-rounded grain formed as a fragment of a former largerrystal (grain 104). Some large inclusions have been identified inhe relict zircon, consisting mostly of quartz, but no inclusions ofenotime or other phosphates were found. Very often, the relictrain has an inclusion-rich margin or specific parts of the crys-al are inclusion-rich; these parts are interpreted as metamorphiceplacements (Fig. 4b; grain 108). In some cases, relict crystalslso show outgrowths and the irregular shape of these outgrowthsxcludes them from being relict (enlarged part of grain 104 inig. 4a; grain 108 in Fig. 4b). Cathodo luminescence-imaging revealsssentially the same features (Fig. 5); dominantly pristine relictrystals show high luminescence along the zoning pattern, but alson patches in the core. A largely metamorphic grain with a preservedore shows cracks in the core (Fig. 5b), where the replacementdvanced, and the replacement area shows a generally homoge-eously distributed high luminescence, but also some areas with
ow luminescence.Most of the metamorphic zircon grains have shapes illus-
rated in Figs. 4 and 6, with all transitional stages of incipient toartial until complete replacement. Metamorphic areas have noegular zoning, but abundant inclusions of xenotime, quartz, mus-ovite, and other minerals, which are characteristic of the Al-Puartzite matrix, such as trolleite and svanbergite. In some cases,he amount of xenotime is very high. Xenotime is often hetero-eneous (Figs. 4 and 6) due to variable REE-contents replacing Y.ocally the metamorphic zircon grains are accompanied by minute≤1 �m) zircon crystals in the matrix (Fig. 4b). In addition, threether textural types of metamorphic zircon occur: fractured, pseu-omorphic, and skeletal. The fractured grain (Fig. 6a) shows a clearelict core, surrounded by a rim, and parts of the rim sit as smallragments in the matrix near to the original grain. Their shape andoning pattern identifies them as parts of the large grain. Clustersf anhedral zircon crystals (Figs. 6b and 7) ≤1 �m to tens of �m in
ize concentrated in round to oval shapes with a large amount ofatrix minerals are indicative of pseudomorphs after a precursor
ircon grain. Skeletal crystals (Fig. 6c) also show abundant inclu-ions, but the homogeneus appearance in BSE images (except for a
slight streaky zoning) indicates that they are single crystals. Moreexamples are shown in Inline Supplementary Fig. S4.
Inline Supplementary Fig. S4 can be found online athttp://dx.doi.org/10.1016/j.precamres.2014.05.011.
According to EMPA (Table 2) the Hf concentrations in zirconvary mainly between 1 and 2 wt% HfO2, corresponding to 0.04and 0.06 cations per formula unit (pfu), respectively. There is noobservable difference between crystals classified as metamorphicand relict. Quantitative determination of REE, U and Th contents inthe metamorphic crystals was hampered by the abundant xeno-time inclusions. From the BSE contrast it appears that most ofthe replaced areas have a lower trace element content than relictzircon.
Rutile often appears together with metamorphic zircon
Fig. 8. BSE-image of a representative rutile crystal, used for Zr-in-rutile geother-mometry. Dark circle is the trace of electron beam damage. Small bright spots arenano-sized zircon inclusions.
Data acquired in November 2006. Spot size 20 �m, depth of crater ∼15 �m.a Within run background-corrected mean 207Pb signal.b U and Pb content and Th/U ratio were calculated relative to GJ-1 reference.c Corrected for background, within-run Pb/U fractionation and common Pb using Stacey and Kramers (1975) model Pb composition and subsequently normalised to GJ-1 (ID-TIMS value/measured value); 207Pb/235U calculated
using 207Pb/206Pb/(238U/206Pb*1/137.88).d Rho is the error correlation defined as err206Pb/238U/err207Pb/235U.e Degree of concordance = 206Pb/238U-age/207Pb/206Pb-age × 100.f Mean and 2 standard deviation of primary (GJ1, n = 12; ratios are not corrected for inter-element fractionation) and secondary reference standard (Plesovice and 91,500, n = 9) during respective analytical session.
16
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Franz et
al. /
Precambrian
Research
250 (2014)
6–26Table 4cAnalytical data for LA-ICP-MS dating in zircon, São Marco, sample 13584B, second analyses series.
Data acquired in May 2011. Spot size = 10–26 �m; depth of crater ∼15 �m. 206Pb/238U error is the quadratic additions of the within run precision (2 SE) and the external reproducibility (2 SD) of the reference zircon. 207Pb/206Pberror propagation (207Pb signal dependent) following Gerdes and Zeh (2009). 207Pb/235U error is the quadratic addition of the 207Pb/206Pb and 206Pb/238U uncertainty.
a Within run background-corrected mean 207Pb signal in cps (counts per second).b U and Pb content and Th/U ratio were calculated relative to GJ-1 reference zircon.c Percentage of the common Pb on the 206Pb. b.d. = below dectection limit.d Corrected for background, within-run Pb/U fractionation (in case of 206Pb/238U) and common Pb using Stacey and Kramers (1975) model Pb composition and subsequently normalised to GJ-1 (ID-TIMS value/measured value);
207Pb/235U calculated using 207Pb/206Pb/(238U/206Pb*1/137.88).e Rho is the 206Pb/238U/207Pb/235U error correlation coefficient.f Accuracy and daily reproducibility was checked by repeated analyses (n = 8) of reference zircon Plesovice, 91500 and Felix; data given as mean with 2 standard deviation uncertainties.
18 G. Franz et al. / Precambrian Research 250 (2014) 6–26
Fig. 9. Concordia diagrams for LA-ICP-MS dating (samples 5868B and 13584B, see Tables 4a and 4b). Detrital zircon grains with metamorphic replacement are plottedtogether (a and e), for all other zircons (relict) the data are plotted in individual diagrams for each zircon crystal. In case of concordia ages the quoted MSWD, mean squareweighted deviation, is the MSWD of concordance. The data show a major Rhyacian age group (∼2 Ga), which is reset by a Brasiliano event, two ages of ∼2.3 Ga (b and f), earlyMesoproterozoic ages of Jequié (2.5–2.7 Ga; c, g and h), and Archaean ages (>3 Ga; d, i–m); for detailed discussion, see text.
ian Research 250 (2014) 6–26 19
ta(isiaTfcc2(tbw
5
dmaadmrc(c
ssmmtAtacadcTt(ar
te0icdM2mtetmsca
g
Fig. 10. Concordia diagrams for LA-ICP-MS dating (sample 13584B, see Table 4c).(a) Relict crystals with a common concordant age of 2050 ± 50 Ma. (b) metamorphicgrains with discordant data points near to the lower intercept. The lower intercept of634 ± 19 Ma indicates the age of peak metamorphic conditions for the crystallizationof metamorphic zircon, the upper intercept is a common mean age of the precursor
G. Franz et al. / Precambr
o apply the Zr-in-rutile geothermometer (Watson et al., 2006),nd in associated quartz crystals the Ti-in-quartz geothermometerWark and Watson, 2006). Observation with high-resolution BSEmages allowed detecting and avoiding nano-sized zircon inclu-ions (Fig. 8); the presence of such inclusions ensures Zr-saturationn rutile. However, inclusions below the surface can still affect thenalysis and lead to anomalously high values, which were excluded.he average value (Table 3) for sample 13584B is 97(30) ppm Zr andor sample 5868B 95(47). The average calculated temperature ofrystallization is 559(25) ◦C and 556(45) ◦C respectively. Titaniumontents in quartz in these two samples (Table 3) are 18(9) and0(10) ppm, corresponding to 576(45) and 585(50) ◦C, respectivelyWark and Watson, 2006). The large spread in calculated tempera-ures is due to the very small trace element contents. Nevertheless,oth geothermometers agree quite well and for the later discussione use an average value of 570 ± 35 ◦C.
.2.2. LA-ICP-MS age determinationsDue to the small size of the metamorphic zircon crystals and
ue to abundant inclusions, overlap of the laser beam with matrixinerals or inclusions was unavoidable. We decided therefore to
nalyze in a first measurement cycle large inclusion-free grains (orreas in grains) compared to inclusion-bearing ones in order toetermine the influence of the inclusions. The clear grains yieldedostly concordant data; but when the laser hit a metamorphic
eplacement area the data are discordant near to the upper inter-ept. In both samples determinations yielded consistent resultsTables 4a and 4b; concordia diagrams Figs. 9 and 10) with con-ordant or slightly discordant data.
In order to obtain reliable ages for the lower intercept, weelected for a second measuring cycle metamorphic grains fromample 13584B, which has a high number of zircon crystals. Theajority of data from these crystals (Table 4c) are discordant withany points occurring near the lower intercept, thereby allowing
he calculation of the lower intercept’s value with high confidence.dditionally we selected inclusion-free relict grains in order to bet-
er constrain the age of the protoliths. Eight grains show concordantges between 2032 ± 19 Ma and 2069 ± 18 Ma (Fig. 10a). For thealculation of a discordia for the metamorphic grains, we selectedll analyses with an upper intercept around or below 2.1 Ga whichefine a common trend; the upper and lower intercepts of this dis-ordia are 2032 ± 31 Ma and 634 ± 19 Ma, respectively (Fig. 10b).he upper intercept of the metamorphic grains is within the uncer-ainty range of the concordant ages. The age of the lower interceptBrasiliano) is supported by three other grains with a higher relictge of 2.6–2.7 Ga (e.g., grain 27; Fig. 9g) and one grain with a lowerelict age near 1.5 Ga (Table 5).
In Fig. 7 we show a grain that is classified as metamorphic dueo its pseudomorphic shape. In two spots, this grain yielded appar-nt near-concordant Pb–Pb ages of ∼1800 Ma (lower intercept of). Assuming, however, a lower intercept of ∼600 Ma the upper
ntercept age is 2092 ± 66 Ma consistent with the age of the con-ordant analyses of the 8 grains mentioned before. In the discordiaiagrams (Fig. 9) similar grains from Vaca Morta (Fig. 9a) and Sãoarco (Fig. 9e) show an upper intercept age of 2020 ± 22 Ma and
033 ± 33 Ma, respectively, confirming our assumption of a com-on relict age for many (but not all) metamorphic grains. However,
he analyses on 10 grains from Vaca Morta (Fig. 9a) define a consid-rable younger lower intercept age of 347 ± 52 Ma, which suggesthat the grains were affected in addition to a late Neoproterozoic
etamorphism by some younger or recent Pb-loss. As the analy-es are relatively far away from the lower intercept the age is not
onsidered as reliable as the previously given value of 634 ± 19 Mand not interpreted in terms of a geological meaning.
The ages of the relict detrital crystals fall into the following fiveroups (Table 5): Seven zircon crystals yielded upper intercept or
zircon crystals.
concordant ages between ∼3.1 and 3.3 Ga. Three of these Archaeanages (São Marco; sample 13584B) are concordant with a small errorof ±13–20 Ma, two of them are shown in Fig. 11a. The analyzedgrains (Fig. 11b and c) have the rounded shape (Fig. 11b) typicalof detrital grains and a zoning pattern that reveals that they arefragments of larger grains (Fig. 11c). In addition to minor replace-ment textures, these grains also show small outbreaks near thegrain boundary separated by quartz and indicating minor fragmen-tation (Fig. 11d and e). Grains 12 and 25 from this sample (Fig. 9iand m) show two concordant spots in the inclusion-free centre ofthe crystals and one discordant spot in the inclusion-rich rim. Grain23 (Vaca Morta; sample 5868B; Fig. 9d) is discordant and its upperintercept also points to an old age.
Nine zircon crystals yielded ages of ∼2.6–2.7 Ga (Table 5; Fig. 9c,g, and h). Many of these ages are concordant or lie near the upperintercept. Two grains (São Marco; grains 106, 113) yielded lowerintercept ages of 651 ± 22 Ma and 638 ± 52 Ma, respectively. Twocrystals yielded ages of 2345 ± 42 Ma (13584B, grain 17; Fig. 9f)and 2385 ± 39 Ma (5868B, grain 6; Fig. 9b). These upper interceptages are well defined for both crystals by concordant analysis of
spots in inclusion-poor areas of the crystals. The discordant agesare in the inclusion-bearing rim of the crystals.
20 G. Franz et al. / Precambrian Research 250 (2014) 6–26
Table 5Age groups of zircon LA-ICP-MS data; zircon classified as relict and metamorphic (see text); R + M indicates strong replacement, however with some areas which are lessaffected. Concordant ages marked with ‘c’ (fifth column), grains with two digits = first session, with three digits second session. MSWD, mean square weighted deviation; -no MSWD quoted for two-point discordia’s; for ‘c’ MSWD of concordance quoted.
Sample Grain R = relict M = meta morphic Spot age/upper intercept Lower intercept MSWD
13584B 142 R 3373 ± 18 c 2.55868B 23 R 3361 ± 91 1467 ± 130 0.39
13584B 21 R 3284 ± 24 c 0.4713584B 133 R 3243 ± 13 c 0.7913584B 12 R 3199 ± 34 1701 ±720 0.6813584B 26 R 3127 ± 20 c 0.2113584B 25 R 3115 ± 52 1856 ± 370 0.3613584B 106 R + M 2714 ± 26 651 ± 22 0.9413584B 15 R 2695 ± 24 c 0.1513584B 131 R 2703 ± 10 c 0.4613584B 123 R 2699 ± 14 c 1.313584B 136 R + M 2685 ± 30 722 ± 120 1.9
5868B 5 R 2623 ± 40 894 ± 210 0.5913584B 113 R + M 2630 ± 26 638 ± 52 0.2613584B 150 R 2589 ± 19 467 ± 370 –13584B 27 R + M 2568 ± 39 675 ± 97 0.33
5868B 6 R 2385 ± 39 656 ± 57 0.9813584B 17 R 2345 ± 42 777 ± 310 0.3813584B 23 R 2056 ± 54 c 0.3913584B 132 R 2069 ± 18 c 1.0613584B 115 R 2059 ± 15 c 0.7813584B 102 R 2055 ± 18 c 1.813584B 155 R 2053 ± 13 c 0.4613584B 127 R 2052 ± 18 c 0.5913584B 126 R 2052 ± 11 c 0.46
5868B 4 R 2036 ± 37 324 ± 250 1.1413584B 125 R 2047 ± 15 c 0.9413584B 107 R 2032 ± 21 c 0.1713584B 100 R + M 1984 ± 85 633 ± 34 0.0213584B 103 M 2092 ± 110 670 ± 66 1.0513584B 105 M 2440 ± 260 565 ± 40 –13584B 108 R 1996 ± 33 685 ± 66 –13584B 109 M 2722 ± 1500 707 ± 74 –13584B 112 R + M 2006 ± 58 650 ± 34 0.4113584B 116 M 2210 ± 730 660 ± 140 –13584B 117 M 1963 ± 130 576 ± 120 1.713584B 122 M 1937 ± 140 636 ± 36 0.1913584B 124 M 651 ± 20 c 0.2413584B 129 M 1942 ± 130 644 ± 31 2.213584B 130 R 2069 ± 24 634 ± 47 –13584B 134 M 2154 ± 320 655 ± 58 –13584B 137 M 1874 ± 98 650 ± 32 0.2513584B 139 M 2032 ± 31 634 ± 19 3.413584B 140 M 646 ± 14 c 1.513584B 141 M 1930 ± 160 580 ± 72 3.413584B 143 R + M 2093 ± 43 703 ± 37 1.613584B 144 M 2510 ± 1100 665 ± 44 –
5868B 34 R 2013 ± 50 580 ± 150 1.513584B 1 R 1998 ± 27 c 1.01
5868B 8 R 2036 ± 78 407 ± 140 0.9413584B 16 R 2023 ± 40 533 ± 88 –
5868B 12 M 1993 ± 110 122 ± 1000 0.165868B 9 M 2036 ± 150 403 ± 240 1.6
13584B 18 R 1980 ± 39 c 0.755868B 32 R 1985 ± 64 219 ± 140 0.12
13584B 10 R 1958 ± 96 78 ± 730 –13584B 24 M 2092 ± 66 ∼600 1.513584B 28 R 2172 ± 145 ∼600 –
5868B 7 R 1645 ± 120 247 ± 640 1.95868B 33 R 2128 ± 590 441 ± 390 1.8
13584B 22 R 2019 ± 170 601 ± 130 –13584B 110 R 1520 ± 21 c 0.04
s2aom
13584B 104 R
13584B 114 R
13584B 101 M
The dominant age group, from individual, mostly concordantpot analysis and from upper/lower intercept ages, ranges from
092 ± 42 to 1958 ± 96 Ma (37 crystals; Figs. 9 and 10a), well inccordance with the known Rhyacian (2.25–2.05 Ga) consolidationf the São Francisco craton. From São Marco, 19 grains classified asetamorphic (Table 5) yielded relict ages that are identical within
1513 ± 18 c 0.621502 ± 29 c 2.31491 ± 130 670 ± 42 1.5
error between 2.0 and 2.1 Ga. This age corresponds to the majorityof concordant relict ages.
Beside some near-concordant ages of grains clearly affected bythe metamorphic overprint, the only concordant ages below 2.0 Gaare 1502 ± 29 Ma and 1513 ± 18 Ma (Fig. 11d; São Marco; nos. 104;114). Grain 114 (Fig. 11e) is completely homogeneous in BSE; it
G. Franz et al. / Precambrian Research 250 (2014) 6–26 21
Fig. 11. Examples for oldest (a–c; Archaean) and youngest (d and e; ∼1500 Ma) detrital zircon grains in BSE images and concordia diagrams. All grains are relicts withoutinclusions, except for minor replacement in the outermost part (in b: grain 133) and metamorphic zircon near to the relict grain 114 (in e). Another grain (104) of the younga
snbibo
aF
ge group is shown in Fig. 4a.
hows an area of metamorphic zircon near to the rim, which wasot hit by the laser beam. One grain (grain 101, strongly affectedy the Neoproterozoic overprint, see Table 4c) yielded an upper
ntercept age of 1491 Ma (Table 5), a value that is near to these data,ut with a high error of ±130 Ma reflecting the enhanced resetting
f the U–Pb system.
In addition to Figs. 4–11 all zircon crystals with their results ofge determinations are documented in the Inline Supplementaryig. S4.
6. Discussion
6.1. Zircon formation
6.1.1. P-T-conditions and protolith
Previous P-T estimates for the Al-P quartzites (Morteani and
Ackermand, 2004) based on oxygen isotope geothermometry onhematite-quartz mineral separates and the presence of kyaniteyielded a minimum pressure of ∼0.4 GPa at 480 ◦C. Here we
2 ian Re
daWnaBrepuamchmbfmacar
a(qsaDamiwrortTstewB
6
cddm(optAdipbln(ad
z
2 G. Franz et al. / Precambr
etermined a temperature of ∼570 ± 35 ◦C using the Zr-in-rutilend Ti-in-quartz geothermometers (Watson et al., 2006; Wark andatson, 2006). The higher temperature is interpreted to be the
ear-peak metamorphic temperature and, in combination with thege data of the associated zircon crystals, the temperature of therasiliano metamorphism. The lower temperature of 480 ◦C likelyesults from resetting during the retrograde path of the Brasilianovent. The temperature of ∼570 ◦C is consistent with the dominanthase assemblage in the rocks, muscovite-quartz, which rules outpper amphibolite to granulite facies conditions. Assuming that thebundant kyanite present in these rocks also belongs to the peaketamorphic assemblage, minimum pressure conditions can be
onstrained to 0.6–0.8 GPa. To the best of our knowledge, eclogitesave not been reported in the Espinhac o and own investigations onetabasaltic rocks from the area only showed the common amphi-
olite facies assemblages, and therefore there are no indicationsor high-pressure metamorphic conditions. The given P-T esti-
ates are in line with the phosphate mineral assemblage (Morteanind Ackermand, 2004; Jordt-Evangelista and Danderfer, 2012), butomplicated textural relationships between the phosphate miner-ls and the lack of knowledge about the phase relations of theseare phosphates hinder the construction of a reliable P-T path.
In a phosphatic sandstone, kyanite probably formed fromn Al-rich clay mineral, and in rocks poor in Mg, Na and KTable 1) a likely candidate is kaolinite. The protolith of the Al-Puartzite was probably a mature quartz-sandstone, deposited in ahallow-marine or littoral facies, as indicated by the very low Mg-nd the high B-concentrations (Danderfer, 2002; Danderfer andardenne, 2002; Morteani and Ackermand, 2004; Jordt-Evangelistand Danderfer, 2012). The presence of the Al-Sr-phosphate-sulfateinerals goyacite-svanbergite together with kyanite strongly
ndicates former diagenetically or hydrothermally altered andeathered kaolinite-bearing profiles. Hematite layers in these
ocks (Cassedanne et al., 1989) point to weathering with generationf lateritic beds. The Al-P quartzites are possibly similar to rockseported by Gaboreau et al. (2005) from Proterozoic unconformity-ype U deposits in the east Alligator River Uranium Field, Australia.hese authors speculated about originally monazite-bearing sand-tone, in which the Al-phosphate-sulfate minerals resulted fromhe interaction of P- and B-bearing fluids with Al-bearing clay min-rals. Deposits containing similar Al-Sr-phosphate-sulfate mineralsith a Late Precambrian diagenetic age were described from thearker basin in Argentina (Dristas et al., 2003; Martínez et al., 2006).
.1.2. Zircon reactionAs frequently documented in the literature detrital zircon
rystals are normally resistant to weathering and hardly reacturing metamorphic overprinting, even at granulite facies con-itions or during melting. In common pelitic metasedimentsetamorphic overgrowths on detrital zircon grains are <1 �m
Vorhies et al., 2013) or <3 �m (Rasmussen, 2005), and only atr above the sillimanite zone beginning dissolution–precipitationroduces rims of ∼10 �m width (Vorhies et al., 2013). However,he unusual textures of the metamorphic zircon crystals in thel-P quartzite point to a much stronger, in many cases completeissolution–recrystallisation process. There are no indications for
nvolvement of other Zr-minerals in the reaction, such as the exam-les shown by Tropper et al. (2007), zircon formation involvingaddeleyite, or by Pan (1997), zircon formation involving zircono-
ite. In the Al-P-quartzites, a high porosity (Fig. 3), the presence ofanoscale zircon crystals in the matrix around the pseudomorphsFig. 4b) and in rutile (Fig. 8) indicates transport of Zr in a fluid
nd thus a high mobility of the high-field strength elements. Theissolution–reprecipitation reaction is
ircon1 + P-bearingfluid = zircon2 + xenotime.
search 250 (2014) 6–26
In the metasediments, zircon 1 comes from a variety of possi-ble sources and its chemical composition must be very variable.Because most of the crystals come from igneous or high-temperature rocks, they contain a large amount of trace elements.During reprecipitation, Y, the middle to heavy REE, U and Thbecome concentrated in xenotime, and the necessary amount ofP is supplied by the fluid. Xenotime is very heterogeneous withstrong variation in REE-contents, indicating incomplete equilibra-tion (Fig. 7). Similar breakdown reactions for zircon were given byPidgeon (1992), Pan (1997), Hoskin and Black (2000), Tomascheket al. (2003) and Rubatto et al. (2008). Tomaschek et al. (2003)described metamorphic zircon from high-pressure metagabbrosand meta-plagiogranites from Syros, Greece (with textures of pris-tine magmatic zircon, topotactically replaced, and skeletal crystals)and argued for a dissolution–reprecipitation reaction. In additionto Y-phosphate (possibly xenotime) they also identified Y-REE-Th-silicate in the metamorphic zircon, which was not found in theAl-P-quartzites. Another example for metamorphic zircon, pro-duced by a dissolution–reprecipitation process from the LanzoMassif, Italy (Rubatto et al., 2008) shows pseudomorphs of zir-con microcrystals associated with Na-pyroxene and epidote fromhigh-pressure/low-temperature meta-plagiogranites. These tex-tures, especially fractured zircon crystals, are also very similar tothe metamorphic zircons shown here.
The fact that zircon in the Al-P quartzites reacts easily isexplained by the nature of the fluid in the quartzite. As indicatedby the presence of phosphates (e.g., berlinite, trolleite and augelite)and phosphate-sulphate minerals (e.g., goyacite-svanbergite), anyfluid generated within the rock during prograde metamorphism viathe breakdown of kaolinite (or other clay minerals) and hydratedphosphates such as variscite, producing berlinite (Drüppel et al.,2007) must be rich in P and S and thus very acidic. Fluid supply fromexternal sources as in the case of an igneous protolith is not nec-essary. Tomaschek et al. (2003) and Rubatto et al. (2008) describedhigh-pressure rocks and therefore argue for a strong influence ofpressure on the solubility of Zr. In case of the Al-P-quartzites,a minimum pressure of 0.6 to 0.8 GPa probably also enhancedsolubility. Rubatto et al. (2008) argued for a strong influence ofNa-rich solutions, but here the amount of alkalies is rather low(see Table 1).
Another factor that influences the dissolution is the extent ofmetamictization of the zircon crystals prior to the onset of meta-morphism and related fluid generation. Grains or parts of grainswith a high degree of metamictization will dissolve and recrys-tallize easier than non-metamict ones. Differences in the degreeof metamictization can be the reason why in some zircon crystalszones or patches are selectively replaced, e.g., by xenotime. Geisleret al. (2003) have shown that even low-temperature hydrothermalfluids can react selectively with amorphous regions of metam-ict zircons. The solubility of radiation damaged zircon and itsgrowth at low temperatures (∼250 ◦C) was described by Balanet al. (2001), Rasmussen (2005), Delattre et al. (2007) and Hayet al. (2010). Experimental treatment of natural metamict zirconat 450 ◦C and 1.3 kbar also confirmed that metamict zircon eas-ily recrystallizes by water infiltration and ion exchange (Geisleret al., 2001). Rubatto et al. (2008) and Tomaschek et al. (2003)argue that metamictization is not a necessary requirement as adriving force for the dissolution–reprecipitation, because of theyoung Paleocene-Eocene age of the rocks they described. In con-trast, the Paleoproterozoic, partly even Archaean zircons of theAl-P-quartzites are old enough to accumulate the necessary radi-ation dosage. Metamictization is accompanied by hydration and
a considerable volume increase. This might be the cause for thefracturing, observed in a number of grains (Fig. 6a and b). Theselectivity of the dissolution and recrystallization as a result of vari-able degree of metamictization is demonstrated by the fact that
ian Re
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6
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t∼ttifbpttfsett
waeeetnahF
tdtteBoaEieob
G. Franz et al. / Precambr
ompletely recrystallized zircons can be found at a microscopicistance from completely unaltered detrital zircons (Figs. 3 and 4).
.2. Geological evolution of the Espinhac o fold belt
.2.1. Ages of source areas and sedimentationThe zircon-bearing clastic sediments of the Veredas Fm. orig-
nated from the erosion of gneisses, migmatites, granulites andharnockites of the Archaean to Paleoproterozoic São Franciscoraton (Almeida, 1977; Almeida et al., 1981; Santos and Souza,983; Lopez and Souza, 1985; Loureiro and Souza, 1985). Theão Francisco craton consists of the Gaviao, Serrinha and Jequiélocks, and the Itabuna-Salvador-Coruc a belt, which collided inhe Transamazonian orogenesis around 2.0 Ga (e.g., Barbosa andabaté, 2002). The basin that originated the Northern Espinhacoold belt was flanked on both sides by rocks of the Gaviao blockith ages predominantly between 2.8 and 2.9 Ga. From the tonalite-
rondhjemite-granodiorite massifs within this block ages between.1 and 3.6 Ga are recorded that belong to the Gurian cycle, the old-st tectonomagmatic cycle identified in South America (Teixeirat al., 1996). In the amphibolite facies Serrinha block, ages areround 2.9 Ga, and in the granulitic Jequié block ages around.7–2.8 Ga are reported. Both ages belong to the Jequié-Aroanectonomagmatic cycle (Almeida et al., 1978; Brito Neves et al.,990; Barbosa and Sabaté, 2002, 2004, and literature therein).
The four age groups of the relict zircons of the Vereda Forma-ion correlate with the ages in the craton. The oldest age group from3.1 to 3.3 Ga corresponds to the ages found in the TTG-gneisses of
he Gaviao block, the oldest crustal rocks of the São Francisco cra-on (Teixeira et al., 1996) that have been eroded and transportednto the Espinhac o rift. Valeriano et al. (2004) reported similar agesrom the southern Brasilia belt. The age group around 2.7 Ga cane attributed to the Jequié-Aroan tectonomagmatic cycle as it isresent in the Serrinha and Jequie blocks. The frequent ages ofhe detrital zircons around 2.07–1.95 Ga put the vast majority ofhe eroded source rocks in the Rhyacian, the time span postulatedor the plate tectonic assembly of the Transamazonian-Eburneanupercontinent (e.g., Alkmim and Marshak, 1989; Martins-Netot al., 2001; Eriksson et al., 2001) and the major crustal consolida-ion of the São Francisco craton. This time span is slightly youngerhan the time span of 2.25–2.05 Ma given by Danderfer et al. (2009).
The two ages near 2.3 Ga (Table 5) do not correlate directlyith a known major tectono-thermal event. Valeriano et al. (2004)
lso found several ages between 2.2 and 2.4 Ga in the south-rn Brasilia belt, which were attributed to the Transamazonianvent, and together with the age spectrum given by Danderfert al. (2009) this might indicate a relatively long time span forhe amalgamation of the Transamazonian-Eburnean superconti-ent. Alternatively, these crystals might come from igneous rocksssociated with intraplate magmatism in the Jequié craton or per-aps from a minor orogenic event not detected until now in the Sãorancisco craton.
The spectrum of relict ages goes down to 1.5 Ga (Table 5; see alsohe age spectrum obtained on monazite and xenotime, by EMPAating, Inline Supplementary Fig. S2), indicating minor activity onhe newly assembled supercontinent. Similar ages are present inhe age spectrum of detrital zircons from the Brasilia belt (Valerianot al., 2004). The age spectrum is also similar to that determined byabinski et al. (2012) for the Macaúbas formation in the Arac uaírogen further south, however they also found ages between 1.5nd 1.0 Ga. The age of the youngest detrital zircons in the Northernspinhac o determined here is similar to the youngest zircon ages
n the Tiradentes quartzites (1536 ± 33 Ma; 1540 ± 45 Ma; Ribeirot al., 2013) in the São João del Rei basin from the southern endf the São Francisco craton (Tiradentes Formation in the Ribeiraelt), which further underlines the correlation of these rocks with
search 250 (2014) 6–26 23
the Northern Espinhac o (Ribeiro et al., 2013). In the coeval Cha-pada Diamantina metasedimentary sequence Silveira et al. (2013)dated crosscutting mafic dikes and sills with an age of 1501 ± 9 Ma(U-Pb badelleyite), indicating an important intraplate event in theNorthern Espinhac o, which continues into the southern range.
The current model of the evolution of the Espinhac o basinassumes sedimentation in a rift that opened during the Paleopro-terozoic Statherian taphrogenesis between 1.8 and 1.7 Ga in theSão Francisco craton (Brito Neves et al., 1979; Brito Neves, 1992;Schobbenhaus, 1996; Chaves et al., 2000; Martins-Neto et al., 2001;Eriksson et al., 2001; Almeida-Abreu and Renger, 2002; Valerianoet al., 2004; Knauer, 2007; Danderfer et al., 2009; Pedrosa-Soaresand Alkmim, 2011, and references therein). The beginning ofrifting and sedimentation is based until now on a combinationof stratigraphic-tectonic relations, U/Pb ages of about 1750 and1748 Ma for synsedimentary alkaline volcanic rocks of tholeiiticaffinity (Machado et al., 1989; Schobbenhaus et al., 1994; Dussinand Dussin, 1995), an age of 1746 Ma for granitoids of the LagoaReal unit (Pimental et al., 1994), and SHRIMP and evaporationzircon ages of ∼1.7 Ga from intercalated metavolcanics (Chaveset al., 2000; Danderfer et al., 2009). Danderfer et al. (2009) andPedrosa-Soares and Alkmim (2011) presented a detailed modelon the rift evolution of the Northern Espinhac o and showed thatrifting occurred during several stages. They separated the wholesedimentary sequence into a number of synthems (see their Fig. 3)based on major unconformities. The first synthems, called Algo-dão, São Simao (with felsic igneous rocks, dated as >Statherianwith 1731 ± 5 Ma) and Sapiranga, mark the beginning of rifting. Thefollowing Pajeú Synthem has volcanic intercalations, dated withzircon to be Calymnian with 1570 ± 30 Ma. It is overlain by theSão Marcos Synthem, which contains mafic intrusive dykes thatintruded the lower part of the sequence (see Fig. 1b). The mini-mum age of sedimentation given by zircon dating of these dykesis Tonian with 854 ± 23 Ma. Danderfer et al. (2009) therefore spec-ulated about a possibly younger age of the upper sequence. TheAl-P-quartzite from the Veredas Formation belongs to the base ofthe next overlying Sitio Novo Synthem. The ages of ∼1500 ± 30 Maof the youngest zircons from this formation are well in accordancewith the ages given for the Pajeú and São Marcos Synthems.
On top of the Sitio Novo Synthem with the Veredas Formationfollows the Santo Onofre Synthem, deposited in the Santo Onofrerift (or Macaúbas rift; Alkmim et al., 2006) in the time intervalbetween 1000 and 900 Ma (Schobbenhaus, 1996). This rift borderedthe already closed Espinhac o rift to the west. The deposition ofthe pelites and turbidites, overlain by breccias, conglomerates andcarbonate rocks of the Santo Onofre-Bambuí Synthems (Danderferet al., 2009), indicates a continental margin environment ratherthan intracratonic rifting. In the Macaúbas (Santor Onofre) rift sedi-mentation started at 900 Ma and continued through the Cryogenian(Babinski et al., 2012).
Closely related to the Espinhac o Supergroup is the assumptionof an Upper Proterozoic tectonometamorphic ‘Espinhac o event’,for which ages between 1.3 and 1.0 Ga (Inda and Barbosa, 1978;Brito Neves et al., 1979) or between 1.4 and 1.5 Ga (Turpin et al.,1988; Cordani et al., 1992; Almeida-Abreu, 1995) were assumed.These ages correspond to that of the two major Mesoproterozoictectonic provinces in South America, the Rondonian-San Ignacioand the Sunsas ones (Cordani et al., 2010), and are temporal equiv-alent with the North American Grenvillian province. However, inour dataset from the metamorphic zircons, there is no indicationfor such an Espinhac o event, in contrast to the investigation ondetrital zircons from the Brasilia (Valeriano et al., 2004) and from
the Arac uaí belt (Babinski et al., 2012). This would be consistentwith the assumption that the sedimentation ended earlier than∼1.3–1.4 Ga (depending on the exact age of the Espinhac o event),possibly due to closure of the rift.
2 ian Re
6
tTttoaqtns
Potgwcm5isis6∼aoiBefiTepgmort
7
Pr(mTcEorbtDgraOsiq
to the subdivision of the Precambrian in South America. Revista Brasileira deGeosciências 20, 267–276.
4 G. Franz et al. / Precambr
.2.2. Brasiliano metamorphismThe lower intercept age of 634 ± 19 Ma (Fig. 10) is interpreted as
he age of the metamorphic overprint during the Brasiliano event.he large number of spots that are near-concordant or lie near tohe lower intercept underline the reliability of the age determina-ion, even though many analyses were influenced by the presencef xenotime (and other) inclusions. The similarity of the mineralssemblages and textures in the Al-P quartzite sampled in the fouruarries (Fig. 1b) suggests that the here determined ages of the pro-olith and metamorphism are representative at least for the wholeorthern part of the Espinhaco, but possibly also for the central andouthern Espinhaco fold belt.
Alkmim et al. (2006) analyzed the structures of the Brasiliano-anafrican deformation produced by the Arac uaí-Westcongorogeny in the south of the Sao Francisco-Congo craton wherehe Paramirim Corridor (in their terminology Paramirim aulaco-en) ends. They concluded that the closure of the Macaúbas rift,hich had opened as an aborted rift in the São Francisco-Congo
raton at ∼900 Ma, started to close as a result of differential move-ents in the large craton at ∼600 Ma. Full closure was achieved at
85–560 Ma. Our age data of 634 ± 19 Ma for the metamorphismn the northern Espinhaco would therefore correspond to the earlytage of the Panafrican compressive tectonics in this orogeny. Sim-lar ages for metamorphism were found in other Brasiliano beltsurrounding the São Francisco craton: e.g., for the Brasilia belt50–640 Ma (Piuzana et al., 2003; granulite facies metamorphism),630 Ma Valeriano et al., 2004; synmetamorphic granites and mon-zite ages from migmatized rocks); from the Ribeira belt, southf the São Francisco craton, Heilbron et al. (2010) reported lowerntercept ages of 590 ± 5 Ma for granulites. In the Pernambucoelt north of the São Francisco craton Neves et al. (2006), Nevest al. (2007) also presented ages of 626 ± 15 Ma and 619 ± 36 Maor high-T metamorphism, however with later (<592 Ma) shear-ng, indicating a similar development of the crust, with a dominantransamazonian crust formation and Brasiliano overprint. de Souzat al. (2006) inferred that north of the São Franciscon craton theeak of the high-T metamorphism (from upper amphibolite toranulite facies) and ductile extensional or transtensional defor-ation for this area may be placed at about 578–574 Ma (Sm–Nd
n garnet and U–Pb on zircon). Monazites ages of 540–550 Maecorded another phase under slightly lower temperatures beforehe ending of the Brasiliano event.
. Conclusions
The age spectrum of the relict zircons found in the Al--quartzites of the Vereda Formation (Espinhaco Supergroup)anging from the oldest ages as determined in the Gaviao Block3.1–3.3 Ga) to ages of around 1.5 Ga show that the basin accu-
ulated in that time debris from the whole São Francisco craton.he youngest age of about 1.5 Ga determined on detrital zirconsan be considered as the time where the sediment input into thespinhac o rift ended, but not necessarily as the time of the closuref the Espinhac o rift. The time of the opening of the Santo Onofreift at the western side of the already sediment-filled Espinhac oasin is under debate as long as age data of the zircon popula-ion of the sediments of the Santo Onofre supergroup are missing.uring the opening of the Santo Onofre rift, also the western mar-in of the São Francisco-W Congo craton along the Espinhac o wasemobilized leading finally to the Paramirim Corridor of Brasilianoge that finds its southern continuation into the Arac uaí fold belt.
ur radiometric data do not confirm the debated existence of a
pecific Espinhac o tectonometamorphic event. Instead, the Brasil-ano age of the amphibolite-facies metamorphism provokes theuestion, if there is a connection between the Arac uaí fold belt
search 250 (2014) 6–26
and the northern fold belts, which surround the São Franciscocraton.
Remarkable is the widespread remobilization and recrystalliza-tion of zircon and replacement by xenotime and inclusion of otherminerals during recrystallisation. The dissolution–reprecipitationof zircon was driven by the P- (and B-)rich fluids produced by thehigh-pressure amphibolite facies Brasiliano metamorphism. Suchdifference in reactivity in combination with in situ age determina-tion makes the zircon population of Al-P-bearing quartzites, such asthat of the Vereda formation, prime candidates for unravelling thehistory of psammitic sequences. Ongoing investigations on an Al-P-quartzite from the southern Espinhac o fold belt near Diamantina(see Fig. 1) supports our model of selective dissolution and repre-cipitation of zircon under amphibolite facies conditions.
Acknowledgements
We thank O. Appelt (Potsdam) for her help with themicroprobe analyses, F. Galbert and J. Nissen (Berlin) for theBSE-documentation, B. Dunker (Berlin) for art work, A. Liebscher(Berlin) for helpful discussions, C. Schobbenhaus (Brasilia), and H.D.Schorscher (Sao Paolo) for critical comments on an earlier versionof the manuscript. C. Preinfalk (München) is acknowledged for herhelp in the field and E. Berryman (Berlin) for polishing the English.Thanks also to J.B. Thomas who has made the two QTiP samples.This work was supported by DFG grant number FR 557/19-1. Threeanonymous reviewers are thanked for thorough reviews and help-ful suggestions.
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