Lithos 122 (2011) 137–151

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40Ar/39Ar ages and Sr–Nd–Pb–Os geochemistry of CAMP tholeiites from WesternMaranhão basin (NE Brazil)

Renaud Merle a,⁎, Andrea Marzoli a, Hervé Bertrand b, Laurie Reisberg c, Chrystèle Verati d,Catherine Zimmermann c, Massimo Chiaradia e, Giuliano Bellieni a, Marcia Ernesto f

a Dipartimento di Geoscienze, Università di Padova, via Gradenigo 6, 35100 Padova, Italyb Laboratoire des Sciences de la Terre, UMR-CNRS 5570, ENS Lyon et Université Lyon1, 46 Allée d'Italie, 69364 Lyon cedex 07, Francec Centre de Recherches Pétrographiques et Géochimiques (CRPG), Nancy Université, CNRS, BP 20, 54501 Vandoeuvre-lès-Nancy cedex, Franced OCA, Université de Nice-Sophia Antipolis, UMR Géoazur, Parc Valrose, 06108 Nice, Francee Section des Sciences de la Terre, Université de Genève, 13 rue des Maraîchers, 1205 Genève, Switzerlandf Departamento de Geofisica, Instituto Astronomico, Geofisico e Ciencias Atmosfericas, Universidade de São Paulo, Rua do Matão, 1226, São Paulo, CEP 05508-900, Brazil

⁎ Corresponding author. Tel.: +39 049 8279154; fax:E-mail address: renmerle@yahoo.fr (R. Merle).

0024-4937/$ – see front matter © 2010 Elsevier B.V. Aldoi:10.1016/j.lithos.2010.12.010

a b s t r a c t

a r t i c l e i n f o

Article history:Received 29 April 2010Accepted 20 December 2010Available online 28 December 2010

Keywords:CAMPOsmium isotopesSCLM

The Central Atlantic Magmatic Province (CAMP), emplaced at the Triassic–Jurassic (T–J) boundary (~200 Ma),is among the largest igneous provinces on Earth. The Maranhão basin in NE Brazil is located around 700 kminland and 2000 km from the site of the earliest Pangea disruption. The CAMP tholeiites occur only in thewestern part of the basin and have been described as low and high-Ti. Here we document the occurrence oftwo sub-groups among the high-Ti tholeiites in the Western Maranhão basin. The major and trace elementsand the Sr–Nd–Pb isotopic ratios define three chemical groups corresponding to the low-Ti (TiO2b1.3 wt.%),high-Ti (TiO2~2.0 wt.%) and evolved high-Ti (TiO2N3 wt.%) western Maranhão basin tholeiites (WMBT). Thenew 40Ar/39Ar plateau ages obtained on plagioclase separates for high-Ti (199.7±2.4 Ma) and evolved high-Ti WMBT (197.2±0.5 Ma and 198.2±0.6 Ma) are indistinguishable and identical to those of previouslyanalyzed low-Ti WMBT (198.5±0.8 Ma) and to the mean 40Ar/39Ar age of the CAMP (199±2.4 Ma). We alsopresent the first Re–Os isotopic data for CAMP basalts. The low and high-Ti samples display mantle-like initial(187Os/188Os)i ranging from 0.1267 to 0.1299, while the evolved high-Ti samples are more radiogenic ((187Os/188Os)i up to 0.184) We propose that the high-Ti WMBT were derived from the sub-lithosphericasthenosphere, and contaminated during ascent by interaction with the subcontinental lithospheric mantle(SCLM). The evolved high-Ti WMBT were derived from the same asthenospheric source but experiencedcrustal contamination. The chemical characteristics of the low-Ti group can be explained by partial melting ofthe most fertile portions of the SCLM metasomatized during paleo-subduction. Alternatively, the low-TiWMBT could be derived from the sub-lithospheric asthenosphere but the resulting melts may haveundergone contamination by the SCLM. The occurrences of high-Ti basalts are apparently not restricted to thearea of initial continental disruption which may bring into question previous interpretations such as thoserelating high-Ti CAMP magmatism to the initiation of Atlantic ridge spreading or as the expression of a deepmantle plume.We propose that the CAMPmagmatism in theMaranhão basinmay be attributed to local hottermantle conditions due to the combined effects of edge-driven convection and large-scale mantle warmingunder the Pangea supercontinent. The involvement of a mantle-plume with asthenosphere-like isotopiccharacteristics cannot be ruled out either as one of the main source components of the WMBT or as a heatsupplier.

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1. Introduction

Many aspects of both the genesis and consequences of continentalflood basalt (CFB) volcanism remain poorly understood and contro-versial, in particular the relationship between volcanism and

continental rifting or the development of hot spot tracks. The CentralAtlantic Magmatic Province (CAMP, Marzoli et al., 1999) is one of thelargest CFB provinces on Earth, extending more than 7500 km northto south, on both sides of the central Atlantic Ocean over a surface inexcess of 107 km2 (Fig. 1; Marzoli et al., 1999; McHone, 2000). CAMPemplacement occurred at the Triassic–Jurassic boundary with a peakat ~199 Ma (40Ar/39Ar age) and distinct pulses of volcanic activityoccurring until about 190 Ma (Deckart et al., 1997; Jourdan et al.,2009; Knight et al., 2004; Marzoli et al., 1999, 2004; Nomade et al.,

9°S

6°S

4°S

46°W 44°W50°W 48°W

8°S

Faults

Normal faults

Brazil

French Guyana Surinam

Ma ranhão basin

SOUTHAMERICA

NORTHAMERICA

WEST AFRICA

250 km

?LTiB

HTiB

LTiB

LTiB

LTiB

LTiB

Maranhão CAMP flows

Amazonianshield

Guyanashield

Leo shield

Reguibat shield

?

Cassiporedyke swarms

Grajau

Barro do Corda

Filadelfia

Porto Franco

Tocantinopolis

Araguatins

Fortaleza dosNogueiras

Araguaina

dyke

sill and lava-flowcraton (exposed)

archaean

A

B

C

Sardinha formation:124-129 Ma (Parana CFB)

Mesozoic sediments

Mosquito formation: 199 Ma (CAMP CFB)

Paleozoic sediments

Araguaia orogenic belt(Brazilian orogen : 550-600 Ma)

Amazonian craton (>2.5Ga)

Evolved high-Ti tholeiite

High-Ti tholeiite

Low-Ti tholeiite

Fig. 1. A: General map of the circum-Atlantic regions at the time of CAMP emplacement and Pangea break-up (~200 Ma, modified after Deckart et al., 2005). The area where high-TiCAMP basalts were formerly recognized is shown by a dashed contour. Dykes of unknown age (Reguibat shield and Senegal) are represented by a dotted line. B: location of theMaranhão Basin in South America. Also shown is the CAMP dyke swarm of Cassiporè. C: Schematic map of the Maranhão Basin with CAMP magmatism and sampling locations ofhigh-Ti, evolved high- and low-Ti rocks.

138 R. Merle et al. / Lithos 122 (2011) 137–151

2007; Sebai et al., 1991; Verati et al., 2007). This magmatic event hasbeen linked to the break-up of Pangea leading to the initial opening ofthe Central Atlantic Ocean in the Florida–Guyana–Liberia area (e.g.,Bertrand, 1991; Bertrand et al., 1982; Cebria et al., 2003; De Min et al.,2003; Dupuy et al., 1988; Nomade et al., 2007, Pegram, 1990, Veratiet al., 2005). Magmatismmay have been induced either by the impactof a plume head under the continental lithosphere (Cebria et al., 2003;Courtillot et al., 1999; Ernst and Buchan, 2002; Hill, 1991; May, 1971;Morgan, 1983; White and McKenzie, 1989; Wilson, 1997) or by heatincubation under thick continental lithosphere and/or edge-drivenconvection generated by thickness contrasts of different lithosphericdomains (Anderson, 1994; Coltice et al., 2007; De Min et al., 2003;McHone, 2000).

The origin of the chemical characteristics of the CAMP basalts andtherefore the nature of their source(s) have been long debated. In theCAMP, as in other CFB provinces, TiO2 concentrations are used todefine two types of tholeiites: prevailing low-Ti (TiO2b2.0 wt.%) andoccasional high-Ti tholeiites (TiO2≥2.0 wt.%). The low-Ti tholeiitesare widely distributed and differ fundamentally from normal mid-ocean ridge basalts (N-MORB), having higher concentrations of light

Rare Earth elements (LREE), large ion lithophile elements (LILE) andmore enriched isotopic signatures. These chemical characteristicswere interpreted, by some authors, in terms of crustal contaminationof asthenospheric (MORB-type) melts (Dostal and Dupuy, 1984) orderivation from an OIB-type plume-related mantle source (Janneyand Castillo, 2001). However, recent geochemical studies havechallenged this model (De Min et al., 2003; Deckart et al., 2005)suggesting alternatively derivation from an enriched (metasoma-tised) subcontinental lithospheric mantle (SCLM) source. This lattermodel involves no or only limited crustal contamination duringmagma ascent, without any interaction with OIB-type components(Alibert, 1985; Bertrand, 1991; Bertrand et al., 1982; De Min et al.,2003; Deckart et al., 2005; Dupuy et al., 1988; Heatherington andMueller, 1999, Pegram, 1990).

In contrast, high-Ti tholeiites were thought to be confined to anarrow zone in northernmost South America (Guyana, Surinam andthe Cassiporé region of Brazil) and western Africa (Liberia, SierraLeone), contiguous to the sites of the first continental disruption ofPangea (Fig. 1). Compared to the low-Ti group, these tholeiites havedistinct isotopic signatures, more akin to those of enriched MORB (E-

139R. Merle et al. / Lithos 122 (2011) 137–151

MORB) and are thought to originate in the asthenosphere (De Minet al., 2003; Deckart et al., 2005; Dupuy et al., 1988). These spatialand geochemical features have been interpreted to reflect a line ofpreferential asthenospheric upwelling during a late CAMP stage,isotopically transitional towards the forthcoming oceanic crust(Deckart et al., 2005). Alternatively, this distribution could resultfrom preferential channelling of melts from a plume-related astheno-spheric upwelling along lithospheric discontinuities.

Radiogenic isotope systematics (Sr–Nd–Pb) and 40Ar/39Ar geo-chronology help to constrain the genesis and the mantle source(s) ofthe CAMP. Improved geochronological data from the many poorlydated CAMP occurrences is particularly useful, because the variousgenetic models predict specific spatial age distributions. For instance,if continental rifting is initiated by a plume head impacting beneaththe lithosphere, a radiating dyke swarm (May, 1971) and a time-progression of ages away from the central plume head region (e.g.,Ernst et al., 1995) are expected. In addition, a contribution of OIB-typecomponentsmay be expected in the isotopic composition of the CAMPlavas if they are related to a mantle plume. Thus, new Sr, Nd and Pbisotope data can help resolve the debate concerning the geodynamicalprocesses responsible for this volcanism, i.e. plume head (Cebria et al.,2003; Courtillot et al., 1999; Morgan, 1983; Wilson, 1997) vs heatincubation-related continental rifting (Coltice et al., 2007, 2009; DeMin et al., 2003; Deckart et al., 2005).

However, Sr–Nd–Pb systematics alone do not allow clear identi-fication of the enriched component involved in the source of the low-Ti basalts of many CFB, since contamination by upper continentalcrust and metasomatised SCLM produce similar Sr–Nd–Pb signatures(Carlson, 1991). In this study we employ the Re–Os isotopic system toresolve this issue. The Re–Os compositions of crustal materials (highRe/Os, thus high 187Os/188Os) are clearly distinct from those ofmantle-derived melts (low Re/Os, therefore low 187Os/188Os), inparticular, melts from the SCLM have very low 187Os/188Os (Shireyand Walker, 1998). In consequence, the Re–Os system is a powerfultool for tracing both crustal assimilation and the contribution of SCLMto CFB. For example, Re–Os analyses of several CFB (Karoo, Ferrar,Emeishan) have provided evidence for a SCLM component in theirsources (e.g., Brauns et al., 2001; Ellam, et al., 1992; Molzahn et al.,1996; Xu et al., 2007).

This study presents the first Re–Os data for CAMP basalts, togetherwith major and trace element analyses, 40Ar/39Ar geochronology, andSr–Nd–Pb isotopes. We also report the first large scale occurrence ofhigh-Ti basalts sampled up to 2000 km south of the first continentalbreak-up area, in the westernMaranhão basin where low-Ti tholeiitesand only one high-Ti sample have been previously documented(Bellieni et al., 1990; De Min et al., 2003; Fodor et al., 1990). Theunique exposure of both high-Ti and low-Ti tholeiites in SouthAmerica far from the region of initial Pangea break-up is critical fordiscussing: (1) whether the two tholeiitic types reflect differentmantle sources, contrasting magmatic processes, crustal interaction,or some combination of these processes and (2) the relationshipbetween Pangea break-up and the genesis of the two magmaticgroups, in particular, the geodynamical process leading to theemplacement of the high-Ti tholeiites far from the early AtlanticOcean.

2. Geological framework

2.1. Geology of the Maranhão basin

The Maranhão sedimentary basin is located in north-easternBrazil, ~700 km from the continental margin and approximately250 km east of the Archean Brazilian craton (N2.5 Ga). The latter iscomposed of greenstone belts, orogenic plutons and high grademetamorphic rocks. The Maranhão basin extends over an area of6×105 km2 and is related to Palaeozoic rifting and lithospheric

thinning (Almeida, 1986; Arora et al., 1999; de Oliveira and Mohriak,2003; Petri and Fulfaro, 1983). This has affected the Late ProterozoicBrazilian mobile belt (600–550 Ma, Trompette, 1994) which iscomposed of metasediments metamorphosed in the amphibolitefacies. Rift structures occur at the NW margins of the Maranhão basin(Fig. 1). Basin rifting may have been re-activated during of the firststage of mid-Atlantic rifting during the Mesozoic (Arora et al., 1999).The sedimentary fill of the basin is mainly composed of Silurian to LatePermian rocks (conglomerates and sandstones) with a few Mesozoicsediments (fine grained sandstones and clays) (Fig. 1). Triassic–earlyJurassic aeolian sandstones are interlayered with the CAMP tholeiiteswhich are locally covered by Middle–Late Jurassic sediments (Petriand Fulfaro, 1983). The maximum thickness of the sediment fill in theMaranhão basin reaches 2000 m (Almeida, 1986).

During Mesozoic times, the Maranhão Basin underwent wide-spread basaltic magmatic activity associated with the two-stageopening of the Atlantic Ocean adjacent to South America. Twotholeiitic events are clearly distinguished in the Maranhão Basin(Baksi and Archibald, 1997; Bellieni et al., 1990; Fodor et al., 1990).Early Jurassic tholeiites crop out in the western part of the basin aslava flows of low-Ti tholeiites (Baksi and Archibald, 1997; Bellieniet al., 1990; Fodor et al., 1990) and are related to the opening of theCentral Atlantic Ocean (i.e., CAMP tholeiites; De Min et al., 2003). Thepresent-day exposures of these volcanic products are lava flows of thestratoid Mosquito Formation (~40,000 km2 Bellieni et al., 1990; Fodoret al., 1990). The CAMP basaltic pile reaches 175 m in the western sideof the basin (Almeida, 1986). Early Cretaceous tholeiites crop out inthe eastern part of the Maranhão Basin mainly as dykes and up to400 m thick sills, constituting the Sardinha Formation (Bellieni et al.,1990; Fodor et al., 1990). These rocks are mostly high-Ti tholeiitesyielding 40Ar/39Ar and K/Ar whole-rock ages of~129–124 Ma. Twoaltered samples yielded an apparent K/Ar whole-rock age of 144.1±4.7 Ma (Baksi and Archibald, 1997; Bellieni et al., 1990; Fodor et al.,1990), but this technique has been proven to be unreliable (Jourdanet al., 2007a; Merle et al., 2009). The tholeiites from the SardinhaFormation are considered as part of the Early Cretaceous Paranà-Etendeka CFB, which is related to the opening of the Southern AtlanticOcean. It is worth noting that the CAMP CFB is not associated withmafic-alkaline magmatism (i.e. lamproites, nephelinites, and kimber-lites) contrary to others CFBs such as the Paranà-Etendeka (e.g.Gibson et al., 1995), Karoo (see Jourdan et al., 2007b for a review) andSiberia CFBs (e.g. Carlson et al., 2006).

2.2. Previous data on the western Maranhão basin tholeiites (CAMPtholeiites)

The geochronology and petrology of the western Maranhão basintholeiites (WMBT) have been previously studied but few isotope dataare available. Step-heating 40Ar/39Ar ages obtained from whole-rockand separated plagioclase range from 190.5±1.6 Ma to 198.5±0.8 Ma (Baksi and Archibald, 1997; Marzoli et al., 1999) which is con-sistent with the main episode of South American CAMP magmatism,mainly centered at 198 Ma (compilation of Nomade et al., 2007).

Among the WMBT, prevailing low-Ti tholeiites (1.2NTiO2N

0.8 wt.%) have been previously documented (Bellieni et al., 1990;De Min et al., 2003; Fodor et al., 1990). These tholeiites are gen-erally evolved (basalts and basaltic andesites), having 50bSiO2b

54 wt.% and 9NMgON5 wt.% (De Min et al., 2003; Fodor et al.,1990). The low-Ti WMBT are slightly enriched in the mostincompatible elements relative to the less incompatible elements,and display strong negative Nb anomalies in primitive mantlenormalised multi-elements patterns (Fig. 4). The REE patterns aremoderately enriched in light REE (LREE) relative to middle REEand heavy REE (LaCN/YbCN=3.9–5.8; DyCN/YbCN=1.1–1.4; LaCN/SmCN=2.0–3.0; LaCN/CeCN=1.1–1.4, chondritic normalizationvalues from Sun and McDonough, 1989). The composition of the

140 R. Merle et al. / Lithos 122 (2011) 137–151

only published WMBT sample with a relatively high TiO2 content(~1.9 wt.%) is distinct from that of the other WMBT, in particularbeing less enriched in LREE (LaCN/YbCN≈2.3; DyCN/YbCN≈1.4;LaCN/SmCN≈1.2; LaCN/CeCN≈1) and having a less pronounced Nbanomaly.

Sr–Nd initial isotope ratios are available for some WMBT. The Pbisotopes measured by Fodor et al. (1990) cannot be taken intoconsideration since they are not back calculated to 199 Ma. For thelow-Ti WMBT, the initial Sr and Nd isotope ratios (calculated at199 Ma) are (87Sr/86Sr)i=0.7061–0.7073, (143Nd/144Nd)i=0.51222–0.51242, εNd=−3.2 to 0.7 (De Min et al., 2003; Fodor et al., 1990).These values are consistent with those of other low-Ti CAMP tholeiitesfrom South America (De Min et al., 2003; Deckart et al., 2005). Theseisotopic signatures, combined with incompatible element ratios, havebeen attributed to the melting of a subduction-related metasomatised(enriched) SCLM and contamination by the upper crust (De Min et al.,2003; Deckart et al., 2005) or interaction between a hypothetical hot-spot and the SCLM (Fodor et al., 1990). The initial isotopic ratios ofthe high-Ti CAMP tholeiites sampled in South America (dykes fromGuyana and Cassiporé state, NE-Brazil) are clearly different fromthe low-Ti WMBT values: (143Nd/144Nd)i=0.51260–0.51268 (εNdi=+4.2 to +5.8), (87Sr/86Sr)i=0.70317–0.70508, (206Pb/204Pb)i=17.97–18.49, (207Pb/204Pb)i=15.46–15.64, (208Pb/204Pb)i=37.66–38.10 (De Min et al., 2003; Deckart et al., 2005). They are con-sistent with a source mainly composed of depleted asthenosphericmantle and a subordinate EMII component (Zindler and Hart, 1986),possibly modified by upper crustal contamination during magmaascent (Deckart et al., 2005). For the single published high-TiWMBT sample, the initial isotopic ratios are (87Sr/86Sr)i=0.70309and (143Nd/144Nd)i=0.51269 (εNdi=+6.1; De Min et al., 2003).

3. The studied samples

All of the studied rocks were sampled in the western part of theMaranhão Basin (Fig. 1; sampling coordinates are given in Table A1 inthe annexe). In this study, together with the low-Ti basalts, wedocument the occurrence of two sub-groups among the high-Titholeiites in the Western Maranhão basin.

While low-Ti rocks (lava flows) are the most widespread in theentire western Maranhão Basin and attain a lava pile thickness ofabout 170 m (Almeida), high-Ti rocks outcrop as a few thick lavaflows near the town of Araguaina. Evolved high-Ti rocks, inferred to bedykes, were sampled in a very restricted area about 200 km to thesouth of the main CAMP lava flow outcrops. Due to sparse exposureand a generally smooth topography it was not possible to definestratigraphic relationships among the three rock groups.

Three textural groups were distinguished, corresponding to thethree chemical groups (low-Ti, high-Ti and evolved high-Ti tholeiites,respectively; see Major and trace elements section). Despite being the

Table 1Summary of Maranhão geochronological data.

Sample Geochemical group Isochron age±2σ MSWD (40Ar/36Ar)ia±2σ

M10 Evolved high-Ti 196.9±1.0 2.1 314±60

M12 Evolved high-Ti 198.0±1.2 1.3 284±62

M15 High-Ti 199.2±1.0 0.02 308±13

The irradiation standard is the Hb3gr hornblende, with an age of 1072 Ma (Jourdan et al., 2Plateau age calculation uses the mean weighted by the inverse variance of the percent of 3

a (40/36Ar)i intercept in the inverse isochron diagram.b Range of 37ArCa/39ArK for the plateau fraction. Related mean Ca/K=37ArCa/39ArK×1.83c Age uncertainty for plateau age includes the error on the 40Ar/39Ar ratio of the monitord Gas fraction which defines the plateau age.e Temperature steps used for isochron and plateau age calculations.

most widespread, the low-Ti tholeiites are fairly uniform petrographi-cally (see mineral analyses in Table A2). They are fine-grained rocks,having a slightly porphyritic texture carrying pheno- to micropheno-crystals of clinopyroxene, plagioclase and very rare olivine. Thegroundmass is composed of plagioclase, clinopyroxene and the Fe–Tioxides magnetite and ilmenite. The high-Ti basalts are coarse-grainedrocks, displaying a doleritic texture. They contain augite, zonedplagioclase and very rare olivine phenocrysts. Oxides and pigeoniteoccur as microphenocrysts. The evolved high-Ti rocks display aporphyritic texture with mm-sized, strongly zoned plagioclasephenocrysts. Oxides are rather large (up to about 200 μm) and haveTi-magnetite composition. Augite and pigeonite are, except for a fewmicrophenocrysts, limited to the groundmass.

Regardless of the chemical group, the altered samples displaythe same features such as olivine transformation to iddingsite andclays, chloritisation of pyroxene, sericitization of plagioclase andcrystallisation of zeolites, carbonates and oxides-hydroxides in thegroundmass.

4. Results

Details of all the analytical techniques are available in theelectronic annexes.

4.1. 40Ar/39Ar ages

Three of the freshest samples, containing sufficiently largeplagioclase crystals, were selected for 40Ar/39Ar geochronology; twoevolved high-Ti tholeiites (M10 and M12) and one high-Ti tholeiite(M15). Results of individual analyses are summarized in Table 1.

The high-Ti sample M10 yields a well-defined plateau age at197.2±0.5 Ma (2σ), concordant with the isochron age of 196.9±1.0 Ma (Table 1, Fig. 2), and representing more than 96% of totalreleased 39Ar. The 37ArCa/39ArK spectrum associated with the plateauage displays a relatively flat pattern, with values ranging from 14.9to 17.0, concordant with microprobe Ca/K analyses (restricted peakat 37ArCa/39ArK=18 in the probability density distribution, Fig. 2).The inverse isochron age is concordant with this plateau age, andyields atmospheric initial 40Ar/36Ar (295.5). Therefore, this plateauage is considered to be a crystallization age.

SampleM12 yields a plateau age of 198.2±0.6 Ma, with 84% of thetotal 39Ar. One younger step at 1200 °C is not concordant with theplateau age within a 2σ confidence level. The statistically robustisochron age for M12 (198.0±1.2 Ma; MSWD=1.3) is concordantwith the plateau age and yields atmospheric initial 40ArCa/36Ar.Although the statistical distribution of the microprobe Ca/K (trans-lated to 37ArCa/39ArK) is more scattered than that of M10, with threedifferent peaks at 13, 16 and 19 (Fig. 2B), the related 37ArCa/39ArKspectrum (values from 14.5 to 17.5) remains concordant with

37ArCa/39ArKb Ca/K Plateau agec±2σ %39Ard T° stepse

14.9–17.0 ~30 197.2±0.5MSWD=1.6

96.5 800 °C-Fuse

14.5–17.1 ~30 198.2±0.6MSWD=1.3

91.0 880 °C-Fuse (one stepat 1200 excepted)

71.2–91.3 ~145 199.7±2.4MSWD=0.0

69.0 750 °C–950 °C

006; Turner et al., 1971).9Ar of each step.

.s, but not the error on the age of the monitor.

1000 10001000

A

B

Fig. 2. A: 40Ar/39Ar age spectra and 37ArCa/39ArK spectra, as a function of %39Ar released. The error boxes of each step are at the 1σ level. Arrows show the plateau age fraction. Error ofthe plateau ages is given at the 2σ level. For samples M10 andM12, the 37ArCa/39ArK ranges of the plagioclase deduced from electron microprobe Ca/K data are reported (grey fields),as well as related peak values from B shown by dotted lines. Microprobe Ca/K peak value from the plagioclase of the high-Ti basalt M13 is reported on the M15 37ArCa/39ArK spectrum(see text for explanation). B: Probability density distribution of 37ArCa/39ArK for samples M10 and M12 calculated from electron microprobe Ca/K data. The Ca/K ratio is proportionalto the 37ArCa/39ArK ratio (with the relationship Ca/K=1.83 x 37ArCa/39ArK). n=number of electron microprobe analyses. Analytical details and complete data set are available in theonline data set.

141R. Merle et al. / Lithos 122 (2011) 137–151

microprobe values. This dispersal of microprobe Ca/K is due to thezoning of the plagioclase (An71 to An55) rather than to post-magmaticalteration (sericitization), since the decrease of 37ArCa/39ArK withdecreasing apparent age that would be expected in case of alterationis not observed. Therefore, the plateau age of 198.2±0.6 Ma isconsidered to be valid.

Sample M15 shows a mini-plateau age at 199.7±2.4 Ma (69% of39Ar; according to the adopted definition, a plateau age requires atleast 70% of 39Ar), concordant with the isochron age (Fig. 2, Table 1).The 37ArCa/39ArK spectrum displays high values (71–92 for the plateaufraction), consistent with the high Ca/K of the anorthite-rich (An88-60)plagioclase of the high-Ti WMBT. The mini-plateau age of M15probably corresponds to unaltered plagioclase because it is related tothe highest 37ArCa/39ArK. Lower apparent ages for the high tempera-ture steps are linked to (1) K-rich alteration phases, i.e. sericite, and(2) high atmospheric contamination (up to 10%).

These three plagioclase bulk samples of evolved high- and high-TiWMBT yield concordant plateau and mini-plateau ages, which areinterpreted as emplacement ages.

4.2. Major and trace elements

Samples from the western Maranhão basin are fresh to slightlyaltered (Table 2). Three low-Ti samples displaying more than 4%LOI (loss on ignition) are not considered further. Compositions ofsamples with LOIb4 wt.% range from basalt to basaltic andesite and

plot in the tholeiitic field on the total alkali-silica diagram (Fig. A1). Noprimitive samples were found, as shown by the relatively low MgO(2.6bMgOb7.9 wt.%) and compatible trace element contents (e.g.,Ni=18–103 ppm).

On a TiO2 vsMg# diagram (Fig. 3), three groups are distinguished:low-Ti (TiO2b1.3 wt.%; 5.6bMgOb7.9 wt.%), high-Ti (TiO2≈2.0 wt.%;6.6bMgOb7.2 wt.%) and evolved high-Ti tholeiites (TiO2=3.4–3.7 wt.%; 2.6bMgOb2.7 wt.%). The tholeiites of the evolved high-Tigroup are the most differentiated among the sampled rocks and theother high-Ti basalts from the rest of the CAMP (e.g., lower Mg#;Fig. 3).

The three chemical groups are also clearly identified in terms ofother major and trace element contents and display differentevolutionary trends (Figs. A2, A3, A4). In the low-Ti group, TiO2,FeOt, Na2O and K2O increase whereas CaO and Al2O3 decrease withdecreasing MgO, suggesting low-pressure clinopyroxene and plagio-clase fractionation (Figs. A2 and A3). The chemical evolution of theevolved high- and high-Ti groups is difficult to constrain, given thesmall number of samples and the apparent homogeneity of eachgroup. However, evolved high-Ti basalts are strongly depleted inAl2O3 and CaO compared to the other groups, which can be related todominant plagioclase and clinopyroxene fractionation. The high-Tibasalts are characterized by lower SiO2 and higher FeO relative to thelow-Ti samples with similar MgO (Fig. A2). Olivine fractionation likelyoccurred in order to explain the lowNi and Co contents in the samplesfrom the three chemical groups.

Table 2Major and trace elements analyses.

Sample M10 M11A M11B M12 M13 M15 M2 M3 M4 M5 M6 M16 M17 M18 M19 M19A 21 M22 M23 M24 M25 M26 M27

Group Evolvedhigh-Ti

Evolvedhigh-Ti

Evolvedhigh-Ti

Evolvedhigh-Ti

High-Ti High-Ti Low-Ti Low-Ti Low-Ti Low-Ti Low-Ti Low-Ti Low-Ti Low-Ti Low-Ti Low-Ti w-Ti Low-Ti Low-Ti Low-Ti Low-Ti Low-Ti Low-Ti

SiO2 (%Wt) 51.33 50.67 51.70 51.53 47.30 47.43 51.88 51.07 51.98 52.37 51.55 50.68 48.74 49.14 50.49 51.13 3.90 51.47 53.53 50.85 50.57 51.02 50.02TiO2 3.24 3.51 3.28 3.39 2.03 2.05 1.15 1.08 1.23 1.15 0.99 1.09 1.03 0.99 1.07 1.07 1.21 1.06 1.11 1.12 1.05 1.13 1.02Al2O3 12.34 12.40 12.43 12.29 13.73 13.87 14.04 14.00 13.60 14.01 14.31 14.35 14.03 13.86 14.68 14.70 3.88 13.98 13.81 13.88 14.37 13.68 13.70FeO* 14.98 15.09 14.41 14.82 12.65 13.50 9.88 9.87 10.15 9.76 9.31 9.46 9.25 8.95 9.23 9.28 9.88 9.65 9.71 9.69 9.61 9.85 9.63MnO 0.24 0.24 0.25 0.24 0.27 0.23 0.17 0.18 0.15 0.16 0.17 0.14 0.16 0.13 0.14 0.15 0.15 0.17 0.16 0.16 0.15 0.16 0.16MgO 2.56 2.64 2.52 2.55 6.89 6.37 7.03 7.42 6.17 6.31 7.06 7.14 7.52 7.20 7.68 7.68 5.53 6.84 6.09 6.86 6.67 6.96 7.63CaO 6.71 6.93 6.63 6.76 9.69 9.90 9.46 9.92 8.90 9.91 10.21 10.58 10.87 10.98 10.94 10.95 9.25 10.17 9.77 10.18 10.12 10.29 10.17Na2O 2.56 2.55 2.58 2.54 2.26 2.30 2.08 2.11 2.11 2.13 2.03 2.05 1.55 1.79 1.95 1.98 2.46 1.95 2.25 2.28 2.47 2.40 2.39K2O 1.37 1.21 1.43 1.29 0.28 0.30 1.01 0.86 0.98 0.94 0.79 0.75 0.36 1.00 0.70 0.63 1.49 0.91 1.05 0.97 0.93 0.93 0.78P2O5 0.80 0.76 0.82 0.80 0.15 0.17 0.14 0.13 0.15 0.15 0.12 0.13 0.11 0.11 0.13 0.13 0.17 0.13 0.16 0.12 0.13 0.12 0.11L.O.I. 2.34 2.42 2.30 2.32 3.52 2.61 2.61 2.52 3.27 2.07 2.38 2.66 5.39 5.12 2.59 2.36 2.00 3.58 1.97 3.42 3.02 3.12 4.40Tot 99.08 99.09 98.95 99.11 99.22 99.22 99.82 99.38 99.05 99.28 99.16 99.44 99.37 99.66 99.84 100.28 0.30 100.06 99.99 99.78 99.50 99.94 100.38Sc (ppm) 16 13 15 16 23 23 27 19 22 22 17 16 19 9 20 17 2 18 23 20 19 25 23V 241 273 242 252 440 440 268 258 286 266 243 276 248 270 258 244 6 252 254 268 262 267 263Cr 9 7 b6 9 65 67 170 188 96 150 202 271 267 263 318 314 8 188 162 263 163 303 318Co 39 41 37 42 53 54 47 46 47 46 45 51 43 43 56 51 2 52 46 54 47 46 50Ni 18 22 19 21 104 93 73 77 50 67 79 82 77 76 95 86 6 76 66 64 69 66 72Rb 42 – 46 42 5 6 – 29 35 – 27 – – – 19 – – 28 – – 31 –

Sr 291 – 306 310 207 210 – 171 169 – 170 – – – 174 – – 186 – – 261 –

Y 85 – 83 85 29 31 – 23 26 – 22 – – – 22 – – 23 – – 22 –

Zr 535 – 551 438 112 107 – 98 109 – 89 – – – 87 – – 101 – – 84 –

Nb 28 – 27 26 6 6 – 8 9 – 7 – – – 7 – – 7 – – 5 –

Ba 426 – 467 460 76 87 – 214 233 – 191 – – – 172 – – 195 – – 222 –

La 41.02 – 40.51 35.22 6.63 7.33 – 11.12 13.84 – 10.71 – – – 9.49 – – 10.70 – – 9.28 –

Ce 100.24 – 101.25 90.46 17.59 19.29 – 24.70 30.12 – 22.89 – – – 21.13 – – 24.58 – – 21.52 –

Pr 14.06 – 14.42 12.21 2.76 3.05 – 3.07 3.80 – 2.97 – – – 2.81 – – 3.02 – – 2.70 –

Nd 66.62 – 65.93 55.49 13.50 14.95 – 13.46 16.75 – 13.01 – – – 12.23 – – 12.84 – – 11.57 –

Sm 16.66 – 16.45 13.58 3.90 4.31 – 3.41 4.15 – 3.22 – – – 3.13 – – 3.15 – – 2.94 –

Eu 4.95 – 5.00 4.03 1.44 1.55 – 1.07 1.23 – 1.05 – – – 1.04 – – 1.02 – – 1.00 –

Gd 17.43 – 18.26 15.11 4.55 5.10 – 3.81 4.67 – 3.61 – – – 4.20 – – 4.40 – – 4.06 –

Tb 2.72 – 2.64 2.12 0.75 0.86 – 0.63 0.78 – 0.62 – – – 0.60 – – 0.59 – – 0.55 –

Dy 15.67 – 15.16 12.14 4.62 5.23 – 3.92 4.81 – 3.87 – – – 3.74 – – 3.75 – – 3.43 –

Ho 2.92 – 2.81 2.24 0.90 1.03 – 0.79 0.98 – 0.78 – – – 0.75 – – 0.74 – – 0.69 –

Er 7.98 – 7.67 6.08 2.56 2.88 – 2.29 2.86 – 2.24 – – – 2.16 – – 2.16 – – 2.01 –

Tm 1.07 – 1.02 0.80 0.34 0.39 – 0.33 0.41 – 0.31 – – – 0.30 – – 0.30 – – 0.28 –

Yb 6.67 – 6.32 4.97 2.17 2.43 – 2.13 2.67 – 2.02 – – – 1.97 – – 1.91 – – 1.80 –

Lu 0.95 – 0.90 0.69 0.31 0.35 – 0.31 0.40 – 0.30 – – – 0.28 – – 0.28 – – 0.26 –

Hf 13.41 – 13.71 11.61 3.05 3.17 – 2.59 3.12 – 2.42 – – – 2.34 – – 2.59 – – 2.32 –

Pb 6.09 – 5.67 5.62 0.91 1.01 – 3.59 4.38 – 3.34 – – – 2.68 – – 3.56 – – 3.43 –

Th 4.31 – 4.31 2.65 0.55 0.59 – 2.42 3.05 – 2.37 – – – 1.87 – – 2.22 – – 1.76 –

U 1.39 – 1.40 1.30 0.17 0.18 – 0.60 0.71 – 0.56 – – – 0.39 – – 1.16 – – 0.34 –

Major elements, compatible trace elements and Y, Nb, Zr, Rb, Sr, Ba analyzed by XRF; REE, U, Pb, Th and Hf analyzed by ICP–MS. FeO* refers to total iron.

142R.M

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137–151

M

Lo

5

1

102

27845

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

CAMP high-Ti

0.0

1.0

2.0

3.0

4.0

5.0

6.0

0

CAMP low-Ti

Evolved high-Ti WMBT

Low-Ti WMBTHigh-Ti WMBT

Mg #

TiO

2 (W

t%)

908070605040302010

Fig. 3. TiO2 vs Mg# diagram for WMBT. Fields of low-, high-CAMP tholeiites from DeMin et al. (2003), Deckart et al. (2005), Dupuy et al. (1988), Jourdan et al. (2003),Nomade et al. (2002), Verati et al. (2005). The data have been recalculated to volatilefree values; cumulative rocks have been excluded. Mg#=Mg2+/(Mg2++Fe2+). Thickgrey arrows: differentiation trends.

143R. Merle et al. / Lithos 122 (2011) 137–151

The three chemical groups show common features on primitivemantle normalized trace elements diagrams (Fig. 4A), such asmoderate enrichment of themost incompatible elements and positivePb and negative Nb anomalies. However, each of the three groups alsodisplays distinctive characteristics, for example in terms of LILE vs REEratios. The low-Ti group shows a particularly strong enrichment in

10

100

La

Roc

k/C

hond

rite

high-Ti WMBT field(previous studies) low-Ti WMBT field

(previous studies)

Evolved high-Ti WMBT

Low-Ti WMBT

High-Ti WMBT

10

100

Rb

Roc

k/Pr

imiti

ve m

antle

Average E-MORB

A

B

YbErYDyGdTiEuSmZrPNdSrPbCeLaNbKUThBa

LuYbTmErHoDyTbGdEuSmPmNdPrCe

Fig. 4. A: Incompatible element patterns of the WMBT. Normalisation values for theprimitive mantle, chondrite and average value for E-MORB are from Sun andMcDonough (1989). B: Rare Earth element patterns of the WMBT. Light grey field:previously studied low-Ti WMBT. Data from De Min et al., 2003 and Fodor et al., 1990.Thick grey line: a previously studied high-Ti WMBT (De Min et al., 2003).

LILE compared to the other incompatible elements leading to a steeppattern. The high-Ti basalts exhibit incompatible element patternscomparable to those of enriched MORBs (e.g., nearly flat pattern) andthe weakest Nb and Pb anomalies among WMBT rocks. The evolvedhigh-Ti tholeiites show the highest incompatible element contentsbut display only a moderate enrichment in the most incompatibleelements relative to the less incompatible ones. These patterns alsoshow clear negative Sr and Ti anomalies, indicative of plagioclase andFe–Ti oxide fractionation, respectively.

The REE patterns of the low-Ti and high-Ti tholeiites overlap therespective fields of low- and high-Ti WMBT previously documented(De Min et al. 2003; Fodor et al., 1990). The high-Ti group presentsnearly flat REE patterns (Fig. 4B), particularly for the light REEs (LaCN/CeCN=0.97–0.98; LaCN/SmCN=1.1; LaCN/YbCN≈2.2). The evolvedhigh-Ti tholeiites are the most enriched in REE and display patternswith higher LREE/HREE ratios (LaCN/CeCN=1.01–1.06; LaCN/SmCN=1.6–1.7; LaCN/YbCN=4.4–5.1) except for the slight negative Euanomalies related to plagioclase fractionation. The patterns of thelow-Ti basalts are more enriched in light REE (LaCN/CeCN=1.11–1.21;LaCN/SmCN=2.0–2.2; LaCN/YbCN=3.5–4.0) and show minor negativeEu anomalies. The three chemical groups display slightly fractionatedheavy REE patterns with an increasing slope from the low-Ti (DyCN/YbCN=1.2–1.3) to the high-Ti (DyCN/YbCN≈1.4) and the high-Tigroup (DyCN/YbCN≈1.6).

4.3. Sr–Nd–Pb isotopic ratios

Table 3 presents results from six samples selected for Sr–Nd–Pbisotopic analyses: two low-Ti (M19 and M26), two high-Ti (M13 andM15) and two evolved high-Ti tholeiites (M10 and M12). All theratios discussed below are initial values calculated at 199 Ma. As forthe major and trace elements, the isotopic ratios of the three chemicalgroups are distinct, although those of the high- and the evolved high-Ti groups display similarities. The low-Ti group displays moreradiogenic initial Pb and Sr isotopic ratios and less radiogenic Ndsignatures than the two other groups (Figs. 5 and 6). The initialisotopic ratios of these samples are similar to those of previouslyanalyzed low-TiWMBT (DeMin et al., 2003) as well as to other low-TiCAMP basalts (Figs. 5 and 6). The low-Ti WMBT display radiogenic207Pb/204Pb (15.61–15.62) for relatively low 206Pb/204Pb (18.11–18.17), moderate 208Pb/204Pb (38.02–38.30), high Sr and low Ndinitial ratios (87Sr/86Sr=0.70616–0.70713, εNd=−1.92 and−1.93).

The high-Ti basalts yield substantially lower Sr and Pb initial ratios(87Sr/86Sr=0.70300–0.70306, 206Pb/204Pb=17.85–17.94, 207Pb/204Pb=15.50–15.52, 208Pb/204Pb=37.75–37.80) and higher Nd ra-tios (εNd=+6.19 and +6.43). These isotopic compositions aresimilar to that of the other high-Ti WMBT (De Min et al., 2003, Figs. 5and 6). The high-TiWMBT also yield positiveΔ7/4 (distance above theNHRL, Fig. 5), which are however lower than for the low-Ti group andclose to the field of the high-Ti CAMP tholeiites.

The evolved high-Ti basalts have initial isotopic ratios similar tothose of the high-Ti group (εNd=+5.63 and +6.02, 206Pb/204Pb=18.00–18.03; 208Pb/204Pb=37.73–37.85) except for slightlylower 207Pb/204Pb (207Pb/204Pb=15.48–15.49) and slightly higher87Sr/86Sr (87Sr/86Sr=0.70341–0.70356). These signatures are similarto those of other high-Ti CAMP tholeiites (Figs. 5 and 6).

The Sr–Nd–Pb isotopic compositions of the three WMBT chemicalgroups are distinct from those of Central Atlantic OIB (Cape Verde,Fernando de Noronha and Ascension) that could have been in therequired paleogeographic position at 200 Ma to initiate the CAMPmagmatism (e.g., Hill, 1991; Wilson, 1997). The high-Ti and evolvedhigh-Ti WMBT plot in the field of the pre-120 Ma Atlantic oceaniccrust (Janney and Castillo, 2001) in the Pb–Pb and Pb–Sr isotopicdiagrams, but have slightly less radiogenic Nd isotopic ratios (Figs. 5and 6).

Table 3Sr–Nd–Pb isotope ratios.

87Sr/86Sr Error 87Sr/86Sr 143Nd/144Nd

Error 143Nd/144Nd

(εNd)i 206Pb/204Pb

Error 206Pb/204Pb

207Pb/204Pb

Error 207Pb/204Pb

208Pb/204Pb

Error 208Pb/204Pb

Measured (±2σ) Initial Measured (±2σ) Initial Measured (±2σ) Initial Measured (±2σ) Initial Measured (±2σ) Initial

M10 0.704440 0.000008 0.703409 0.512867 0.000008 0.512670 5.63 18.476 0.002 18.031 15.511 0.002 15.488 38.183 0.006 37.729M12 0.704414 0.000006 0.703556 0.512883 0.000008 0.512690 6.02 18.455 0.002 18.003 15.500 0.002 15.477 38.158 0.004 37.855M13 0.703184 0.000002 0.703000 0.512939 0.000010 0.512711 6.43 18.209 0.026 17.847 15.515 0.022 15.496 38.138 0.054 37.752M15 0.703242 0.000006 0.703062 0.512926 0.000008 0.512699 6.19 18.284 0.026 17.940 15.535 0.022 15.518 38.183 0.056 37.805M19 0.706943 0.000004 0.706155 0.512485 0.000010 0.512284 −1.92 18.400 0.0006 18.112 15.628 0.0004 15.613 38.471 0.001 38.021M26 0.707985 0.000004 0.707129 0.512483 0.000010 0.512283 −1.93 18.366 0.006 18.168 15.631 0.004 15.621 38.632 0.012 38.300

(εNd)i were calculated related to the CHUR at 199 Ma using the present day values for CHUR:(143Nd/144Nd)chur=0.512638, 147Sm/144Nd=0.1967, (Jacobsen and Wasserburg, 1980).

HIMU0.702

17.5

HIMU

NHRL

EM I

37

38

39

CV

A

FdN

15.2

15.3

15.4

15.5

15.6

15.7

EM I

HIMUEM II

NHRL

Geo

chro

n

(207 Pb

/204 Pb

) i(20

8 Pb/20

4 Pb) i

(206Pb/204Pb)i

(87Sr

/86Sr

) i

FdN

CV

A

EM II

0.703

0.704

0.705

0.706

0.707

0.708

EM I

A

CV

FdN

EM II

DMM

DMM

DMM

Low-Ti CAMP

High-Ti CAMP

Low-Ti WMBT

High-Ti WMBT

Evolved high-Ti WMBT

Pre-120 Ma Atlantic crust

Post-120 Ma Atlantic crust

19.518.5

Fig. 5. Initial Pb and Sr isotopic compositions of WMBT. Fields of pre-120 Ma and post-120 Ma Atlantic MORBs from Janney and Castillo (2001), data for CAMP low- and high-Ti tholeiites from Cebria et al. (2003), De Min et al. (2003), Deckart et al. (2005), Dupuyet al. (1988), Jourdan et al. (2003), Verati et al. (2005). Also shown are fields of present-day OIB lavas from Ascension (A), Fernando de Noronha (FdN) and Cape Verde (CV);data of Ascension, Fernando de Noronha and Cape Verde lavas from the Georocdatabase. Mantle end-member values from Zindler and Hart (1986) back-calculated to199 Ma.

144 R. Merle et al. / Lithos 122 (2011) 137–151

4.4. Os isotopic results

All the samples analyzed for Sr–Nd–Pb initial isotope ratios werealso selected for Re–Os isotopic analyses. Measured Os isotopic ratios,related uncertainties, Os and Re concentrations and initial ratios arereported in Table 4. All the Os ratios discussed below are initial valuescalculated at 199 Ma.

187Os/188Os ratios have been plotted against 187Re/188Os in awhole-rock isochron diagram (Fig. 7). Despite the fact that the threegroups of WMBT were probably not cogenetic (see discussion below),the data from samples of the low-Ti and high-Ti groups define anapparent isochron yielding an age of 198.9±7.6 Ma (MSWD=0.81;Probability=0.44) with initial 187Os/188Os=0.1284±0.0039. Thisage is indistinguishable from the 40Ar/39Ar ages of the samples(199 Ma), suggesting that the Re–Os system was not disturbed afterthe eruption of the CAMP basalts, and that the initial ratios back-calculated to 199 Ma have a real geological significance. The twoevolved high-Ti samples have higher 187Re/188Os ratios and moreradiogenic measured 187Os/188Os ratios, and plot within error of(sample M12) or slightly above (sample M10) the apparent isochrondefined by the other samples.

Initial 187Os/188Os of the low-Ti and high-Ti WMBT displayrestricted variations. Samples M13, M15, M19 and M26 haveindistinguishable values, ranging from 0.1267±0.0044 to 0.1299±0.0033. The Os concentrations are lower for the low-Ti (26.5–57.7 ppt) compared to the high-Ti basalts (70.5–104.6 ppt). Thismay be related to distinct stages of magmatic differentiation since thehigh-Ti group have higher Ni contents (Table 2). There is nocorrelation between initial 187Os/188Os and Os concentration. Theinitial Os isotope ratios of the low- and high-Ti basalts are close to theestimated value of the hypothetical Primitive Upper Mantle (Fig. 8;PUM199Ma=0.1281±0.0008; based on present-day PUM ratio of0.1296 from Meisel et al., 2001), but are slightly less radiogenic thanthose of most modern OIB (e.g., 187Os/188Os=0.13–0.15; Shirey andWalker, 1998). In Pb–Os, Sr–Os and Nd–Os isotopic diagrams (Fig. 8),the low- and high-Ti WMBT do not match modern OIB compositions.In particular, the initial Os isotopic compositions of the low and high-Ti WMBT are less radiogenic than those of the lavas from the CapeVerde archipelago, which has been suggested to be a present-dayremnant of a postulated CAMP mantle-plume (e.g., Wilson, 1997).Even if the Cape Verde lavas are slightly contaminated by oceaniclithosphere (Escrig et al., 2005) even though their compositions maynot exactly represent the melts from the mantle plume, the estimatedplume composition (187Os/188Os=0.1325, Escrig et al., 2005) is stillmore radiogenic than those of the low and high-Ti WMBT.

The evolved high-Ti samples M10 and M12 have low Osconcentrations (19 and 6 ppt), consistent with their differentiatedcharacter. These samples have initial 187Os/188Os ratios of 0.184±0.012 and 0.16±0.1 respectively, with the large uncertaintiesresulting from the very large radiogenic Os corrections. The initialratio of M10 is more radiogenic than those of the low- and high-Ti

17.5

EM I EM II

HIMU

19.00.702

EM I

HIMU

0.5120

0.5125

0.5130

0.5135

FdN

A

CV

A

FdNCV

EM II

DMM DMM

High-Ti WMBT (previous studies)

Low-Ti CAMP Low-Ti WMBT

High-Ti WMBT

Evolved high-Ti WMBT

High-Ti CAMP

Low-Ti WMBT (previous studies) Pre-120 Ma Atlantic crust

Post-120 Ma Atlantic crust

PUM

0.7080.7070.7060.7050.7040.703

(87Sr/86Sr)i (206Pb/204Pb)i

(143 N

d/14

4 Nd)

i

20.019.518.518.0

Fig. 6.Nd–Sr and Nd–Pb diagrams of initial isotopic ratios fromWMBT. Same references as in Fig. 5 for fields of pre-120 Ma and post-120 Ma Atlantic MORB, Ascension (A), Fernandode Noronha (FdN) and Capo Verde (CV) OIB lavas, low-Ti and high-Ti CAMP tholeiites. Sr–Nd isotopic compositions of previously analyzedWMBT from DeMin et al., 2003 and Fodoret al., 1990. Oceanic mantle end-members values from Zindler and Hart (1986) back-calculated to 199 Ma. PUM: primitive upper mantle (values from Zindler and Hart, 1986).

145R. Merle et al. / Lithos 122 (2011) 137–151

samples, of PUM and of most OIB. The large error on the initial 187Os/188Os of M12 precludes interpretation of the Os signature of thissample.

5. Discussion

5.1. Timing of igneous activity in the Western Maranhão basin and itsrelationship with other CAMP areas

The new 40Ar/39Ar plateau ages obtained for high-Ti and evolvedhigh-Ti basalts overlap the previously published 40Ar/39Ar plateauage obtained on plagioclase from a low-Ti WMBT (198.5±0.8 Ma;Marzoli et al., 1999) which belongs to themainmagmatic phase of theCAMP in South America (198 Ma, Nomade et al., 2007). Thisdemonstrates that (1) the newly documented high-Ti basalts foundin the Western Maranhão basin belong to the CAMP magmatic eventand (2) the three chemical groups were emplaced synchronously (i.e.,within a time-span≤analytical uncertainty). CAMPmagmatic activityoccurred in South America from 199.0 until 190.5 Ma and spread overthree phases (Baksi and Archibald, 1997; Marzoli et al., 1999; Nomadeet al., 2007). Using the age data selection criteria outlined by Nomadeet al. (2007), the only previously published ages that can beconsidered as reliable for South America are those of the low-Ti lavaflows and dykes. Nevertheless, the few available data for high-Ti

Table 4Re–Os isotope data of basalts from the western Maranhao Basin.

[Re] (ppt) [Os] (ppt) 188Os (mol g−1) (187Os/188Os)m

M10 841 18.9 1.19E−14 0.97763M12 990 5.9 2.68E−15 4.30101M13 219 70.5 4.89E−14 0.17725M15 408 104.6 7.24E−14 0.19122M19 433 57.7 3.96E−14 0.25195M26 282 26.5 1.81E−14 0.30134

Uncertainties listed for the measured 187Os/188Os ratios correspond to 2 standard errors froIn-house standard reproducibility was b0.2%.All data are blank corrected, using blank values given in the analytical procedures.Uncertainties on initial ratios include in-run errors and uncertainties on blank corrections aInitial ratios were calculated using a decay constant λ=1.666×10E−11 (Smoliar et al., 199

basaltic dykes (Deckart et al., 1997; Marzoli et al., 1999) suggest thatthey were intruded during the CAMP peak activity in South America.Our new high-quality age data confirm that the high-Ti magmas wereemplaced during the peak activity at ~198 Ma. More ages meeting thehigh quality criteria from high-Ti basalts from South America and inparticular Brazil are needed to better constrain this activity.

When compared to ages from elsewhere in the CAMP, our age datado not show evidence of a centripetal time-relatedmigration from thelocation of the hypothetical plume head (Blake Plateau; Ernst andBuchan, 2002; Ernst et al., 1995; Hill, 1991; May, 1971; Wilson, 1997)towards marginal CAMP areas such as the western Maranhão basin.Instead, and although the mean ages of CAMP basalts from thedifferent circum-Atlantic areas overlap within uncertainty, our newages seem to be consistent with a slightly younger peak event forCAMP in South America than in North America and Africa (Jourdanet al., 2009; Nomade et al., 2007).

5.2. Nature of the mantle source(s) of CAMP magmatism in the WesternMaranhão basin

As previously stated, the initial 187Os/188Os of both low-Ti andhigh-Ti WMBT are similar within uncertainties, and overlap withtypical mantle values, i.e. those of the primitive upper mantle (PUM).This suggests that (1) the amount of crustal contamination was

es error(±2σ) 187Re/188Os (187Os/188Os)199 Ma (±2σ)

0.00115 239 0.184 0.0120.00769 1247 0.16 0.100.00061 15.1 0.1273 0.00530.00017 18.9 0.1283 0.00240.00028 36.7 0.1299 0.00350.00046 52.6 0.1267 0.0050

m in-run statistics.

nd on 187Re/188Os ratios and ages used for radiogenic corrections (all 2σ).6).

data-point error crosses are 2σ0.34

0.30

0.26

0.22

0.18

0.1425 35 45 55 65155

187 O

s/18

8 Os

187Re/188Os

Model 1 Solution (±95%-conf.) on 4 pointsAge = 198.9±7.6 Ma

MSWD = 0.81, Probability = 0.44

1

2

3

4

5

400

Initial 187Os/188Os = 0.1284±0.0039

1200800

187Re/188Os

187 O

s/18

8 Os

Fig. 7. Re–Os isochron plot of WMBT. Age and uncertainties were calculated using theIsoplot software (Ludwig, 2003) for high- and low-Ti rocks. Line is regressed throughthe low- and high- Ti samples. Inset shows positions of evolved high-Ti samples relativeto this regression line. Uncertainties on each data point are 2σ and include in-run errorsand blank and weighing uncertainties.

Average SCLM

EMI

EMII

HIMU

DMM

0.5120

PUM

SCL

M

worldwide OIB

Low-Ti WMBT

High-Ti WMBT

Evolved high-Ti WMBT

CV

HIMUEMI

EMII

DMM

Average SCLM 0.11

0.12

0.13

0.14

0.15

0.16

0.17

0.18

0.19

PUM

SCL

M

CV

ACC

Average SCLM

EMII

EMI

PUM

HIMU

DMM

SCL

M

CV

ACC

ACC

(187 O

s/18

8 Os)

i

0.11

0.12

0.13

0.14

0.15

0.16

0.17

0.18

0.19(18

7 Os/

188 O

s)i

0.11

0.12

0.13

0.14

0.15

0.16

0.17

0.18

0.19

(187 O

s/18

8 Os)

i 206Pb/204Pb

87Sr/86Sr

143Nd/144Nd

2221191817

0.7080.7070.7060.7040.7030.702

0.51350.51300.5125

Fig. 8. Initial Os vs Sr–Nd–Pb isotopic ratios of WMBT. Os mantle pole values fromShirey andWalker (1998) and primitive upper mantle (PUM) fromMeisel et al. (2001).Sr, Nd and Pb mantle end-members values from Zindler and Hart (1986). The mantlepoles as well as the PUM in the diagrams were back-calculated to 199 Ma (present-dayPUM values: 187Os/188Os=0.1296±0.0008; 187Re/188Os=0.4353; Meisel et al., 2001).Data points of modern worldwide OIB and lavas from Cape Verde (CV) from Georocdatabases and Escrig et al. (2005). The OIB data have been filtered for lithospherecontamination: data with [Os]b30 ppt and 187Os/188OsN0.16 have been discarded. ACC:Average continental crust.

146 R. Merle et al. / Lithos 122 (2011) 137–151

negligible for these magmas and (2) derivation from a commonparental magma (implying the same mantle source) that wouldhave been subsequently contaminated by the same component but indifferent ways is unlikely. This latter aspect is also shown by the Sr–Nd–Pb isotopic plots (Figs. 5–6). Indeed, the WMBT plot in twodistinct fields, corresponding to the low-Ti and high- plus evolvedhigh-Ti groups. These two groups are separated by a significant gap asshown also for low- and high-Ti CAMP intrusive rocks from Guyanaand Guinea (Deckart et al., 2005). As is true for other CFB provinces inwhich low- and high-Ti basalts are present, this lack of compositionalcontinuum between the low-Ti and high- plus evolved high-Ti groupssuggests involvement of at least two distinct sources, as previouslyproposed for the Guyana CAMP basalts (Deckart et al., 2005).

5.2.1. Mantle source of the low-Ti WMBTThe low-Ti WMBT are characterized by (1) LILE enrichment,

(2) strong negative Nb and positive Pb anomalies, (3) radiogenic 87Sr/86Sr and unradiogenic 143Nd/144Nd signatures, (4) high 207Pb/204Pb forrelatively low 206Pb/204Pb and (5) mantle-like initial 187Os/188Os closeto the PUM value at 199 Ma. We will examine several hypothesesthat could explain this combination of “enriched,” characteristics (highLILE, high 207Pb/204Pb and 87Sr/86Sr, low 143Nd/144Nd) with low Osisotopic composition in the low-Ti basalts. These include: (1) deriva-tion from a deep mantle plume with the required trace elementand isotopic characteristics; (2) contamination of asthenospheric orplume-related magmas by the continental lithosphere or (3) partialmelting of enriched, metasomatized portions of the SCLM.

High 207Pb/204Pb for relatively low 206Pb/204Pb is usually docu-mented either in old (N1 Ga) continental crust material or in mantleplumes incorporating recycled continental crust or sediments (EM-typemantle plumes; Zindler and Hart, 1986). However the hypothesisof a mantle plume as the sole component or possibly contaminatedby continental lithosphere (SCLM and/or crust) presents severalproblems:

(1) Ocean Island Basalts with trace element and Sr–Nd–Pb–Osisotopic compositions similar to those of the Maranhão low-Tibasalts are absent in the Central Atlantic OIB lavas (see Figs. 5, 6and 8). In particular those ocean islands suggested to be thepresent-day expression of the mantle-plume that may havegenerated the CAMP, i.e. Cape Verde, Ascension and Fernandode Noronha, yield clearly higher 206Pb/204Pb, 187Os/188Os and

143Nd/144Nd, and lower 87Sr/86Sr isotopic ratios than the rocksstudied here. However, as suggested for the Paraná-EtendekaLIP (Thompson et al., 2001), the tholeiitic basalts of the CFBevent could sample a distinct portion of the mantle plume andthus have a geochemical signature distinct from that of therelated OIBs. In this case, the CAMP magmas might represent

147R. Merle et al. / Lithos 122 (2011) 137–151

high temperature melts, while the cited Atlantic OIBs would allbe low-temperature alkaline magmas that probably sampledthe most fertile and enriched portions of a heterogeneousmantle-plume. However, this hypothesis is in contradictionwith the Sr, Nd and 207Pb/204Pb isotopic ratios of the CAMPlow-Ti basalts which display more enriched compositions thanthe OIBs.

(2) It is widely accepted that plume-related melts (i.e. OIBs) do notshow negative Nb and positive Pb anomalies (Weaver, 1991).Exceptions are magmas issued from the extreme end-memberEM-II mantle plume from Samoa which display Sr–Nd–Pbisotopic compositions and negative Nb (but not positive Pb)anomalies broadly similar to those observed in the low-TiWMBT (Jackson et al., 2007). However, such extreme composi-tions are absent in other oceanic islands worldwide and rareeven at Samoa. A similar small-scale heterogeneity wouldproduce isotopic zoning as observed at Samoa, contrasting withthe uniform composition of the low-Ti CAMP over the entireprovince (N 10 million km2). In consequence, an extinct plumewith the required trace element and isotopic characteristics(extreme EMII composition such as documented for theSamoan lavas) seems unlikely. Nevertheless we cannotcompletely rule out this possibility.

(3) In the hypothesis of a plume interacting with the SCLM, trendsbetween OIB and SCLM compositions might be expected in theOs–Nd, Os–Pb and Os–Sr isotopic diagrams (Fig. 8). On thecontrary, the compositions of the low-Ti WMBT plot far fromany OIB composition in the 143Nd/144Nd vs 187Os/188Os diagram(Fig. 8), suggesting that the SCLM signature is dominant and ishiding possible sub-lithospheric components.

(4) The enriched signature of low-Ti WMBT could alternatively beattributed to large amounts of continental crust assimilation ofplume-related magmas during ascent towards the surface.Regardless of the composition of this plume (present CentralAtlantic OIBs or a distinct hypothetical plume), the initial Osisotopic ratios of the samples argue against this scenarioespecially considering the low Os concentration of sample M26which would make it sensitive to crustal contamination. Anopen magma chamber in the lower crust, undergoing AFC andmagma replenishment at the same time might producemagmas with unradiogenic Os compositions coupled withenriched Sr–Nd–Pb compositions (Molzahn et al. 1996).However, such magma chamber behaviour tends to yieldhigh incompatible elements contents (e.g. Aitcheson andForrest, 1994) which are not observed in the low-Ti WMBT.

Low-Ti basalts represent the vast majority of preserved CAMProcks and their origin has been previously investigated based on Sr–Nd–Pb isotopic data (De Min et al., 2003; Deckart et al., 2005; Pegram,1990). These isotope studies argued in favour of a dominant SCLMcomponent in the low-Ti magmas (De Min et al., 2003; Deckart et al.,2005; Pegram, 1990) but were unable to dismiss the possibility ofcontamination by the continental crust. However, the contrastingchemical characteristics of the low-Ti WMBT, i.e. low 187Os/188Os and“enriched” crustal-like incompatible trace element and Sr–Nd–Pbisotopic signatures, match those expected for continental flood basaltsderived from the SCLM (e.g., Brauns et al., 2001). As the off-cratonicSCLM beneath the WMBT is probably more fertile than cratonic SCLM(Griffin et al., 2009) and may also have been hydrated by subductionderived fluids, partial melting of this metasomatised lithosphere is aplausible process (Gallagher and Hawkesworth, 1992). Indeed, partialmelting of a heterogeneous and fluid-metasomatised SCLM has beenproposed to explain the isotopic characteristics of the low-Ti basaltsfrom the Karoo CFB (Jourdan et al., 2007b). The resulting basalticmagmas would be characterised by enrichment in LILE, LREE and Pband by high 207Pb/204Pb at moderate 206Pb/204Pb. This “continental-

like” geochemical signature would be carried by subduction-relatedfluids into the mantle wedge and the overlying lithospheric mantle.The initial 187Os/188Os ratios of the low-Ti group are at the high end ofthe range of typical off-cratonic SCLM values (~0.118–0.129; Carlson,2005). This could be explained by very minor amounts of crustalcontamination (see the AFC modelling below) or by partial melting ofSCLM components with slightly more radiogenic Os compositions,similar to those measured in mantle xenoliths from arc settings(187Os/188Os=0.1182–0.1585; Brandon et al., 1996; Saha et al., 2005;Widom et al., 2003).

However, SCLM melting (Gallagher and Hawkesworth, 1992)requires a thick lithosphere which is unlikely underneath theMaranhão Basin. Alternatively, the low-Ti basalts might have beenproduced by mixing between magmas derived from an astheno-sphere-like source (sub-lithospheric asthenosphere or astheno-sphere-like mantle plume) and liquids that originated from meltingof metasomatic veins in the SCLM (e.g. Arndt and Christiansen, 1992).The geochemical signature of the asthenosphere-derived magmasmay have been swamped by the SCLM-derived melts such asdemonstrated for the Paraná-Etendeka CFB (Gibson et al., 1995).

Both hypotheses (i.e. partial melting of the SCLM contaminated byupper continental crust or asthenosphere-derived magmas mixedwith melts from metasomatised parts of the SCLM) will be tested bysimple numerical modelling.

5.2.2. Source of the high- and evolved high-Ti WMBTDue to the broad similarities of Sr–Nd–Pb isotopic signatures

between the high- and evolved high-Ti WMBT, a common source issuggested for these rocks. Nonetheless, small but significant differ-ences between the isotopic signatures of the two groups indicateeither heterogeneities within a single source or involvement of one orseveral other components. Since the moderately radiogenic initial187Os/188Os ratios of the evolved high-Ti WMBT argue for acontribution from the continental crust (at least for Os isotopes),only the high-Ti samples were used to constrain the origin of themantle source of the magmas.

The mantle source characteristics of the high-Ti WMBT may beinferred from their main geochemical features which are: (1) moder-ately enriched REE patterns, similar to those of E-MORBs, (2) weaknegative Nb and positive Pb anomalies, (3) 143Nd/144Nd slightly lowerthan those of N-MORBs, (4) Sr–Pb isotopic compositions similar tothose of pre-120 Ma Atlantic MORBs (Janney and Castillo, 2001) and(5) Os isotopic compositions close to that of the PUM.

These data suggest a dominant involvement of a depletedasthenosphere-type mantle source (sub-lithospheric asthenosphereor asthenosphere-like mantle plume), largely similar to that of thepre-120 Ma Atlantic MORBs. As pointed out also by Janney andCastillo (2001) for the Jurassic Atlantic MORBs, a minor input of anenriched component is required to explain the fact that the high-TiWMBT are slightly enriched compared to present-day Atlantic N-MORBs having slightly lower 143Nd/144Nd, slightly higher 87Sr/86Srand incompatible element contents, and Pb isotopic compositionsslightly above the NHRL. Since the isotopic compositions of thepresent-day hot-spot related Atlantic OIBs do not match those of thepre-120 Ma Atlantic MORBs (and of the CAMP high-Ti basalts) (Figs. 5and 6), Janney and Castillo (2001) interpreted the enriched signatureof the pre-120 Ma MORBs as related to a short-lived and now extinctmantle-plume with EM-I characteristics. According to this interpre-tation, such an “extinct plume” would have generated the CAMP andcould explain the slightly enriched composition of the high-Ti CAMPbasalts. However, unlike the pre-120 Ma MORBs, high-Ti WMBT arecharacterized by negative Nb and positive Pb elemental anomalies,features that are not typical of plume-related OIBs. A significantcontamination by the upper continental crust could produce theseelemental anomalies, but is in contradiction with the PUM-like Ossignature. Alternatively, the SCLM is a good candidate for the required

148 R. Merle et al. / Lithos 122 (2011) 137–151

enriched component since its composition matches EMI-type compo-sitions and is consistent with low initial 187Os/188Os (Hofmann, 1997).Moreover, the CAMP high-Ti magmas plot between a depletedasthenospheric end-member (e.g., DMM) and the enriched composi-tions of the low-Ti CAMP basalts, which we interpret to be dominatedby a SCLM contribution. Under this scenario, the enriched signaturemay have been acquired by assimilation of the SCLM en route to thesurface by magmas originating from either the sub-continentalasthenosphere or an asthenosphere-like mantle plume.

In summary, based on the isotopic and trace element data weconclude that the low-Ti basalts may be dominated by a SCLMcontribution which could reflect either partial melting of thesubcontinental lithosphere or contamination during ascent of as-thenosphere-derived magmas. The high and evolved high-Ti WMBTwere produced by melting of an asthenospheric source but werecontaminated by distinct components of the continental lithosphere.The plausibility of these scenarios will be tested by a numericalmodelling.

5.3. Petrogenesis of the western Maranhão basin tholeiites

5.3.1. Fractional crystallization of the WMBTThe generally low MgO, Cr and Ni concentrations show that

fractional crystallization strongly influenced the chemical evolution oftheWMBT. However, the major element trends, the REE patterns, andmost conclusively, the initial Sr, Nd and Pb isotopic compositions,preclude derivation of one group from another through fractionalcrystallization (see Figs. A2, A3 and A4 in annexes). On the other hand,the major element trends and the concentrations of compatible traceelements, as well as the Sr, Eu and Ti anomalies in multi-elementdiagrams (Fig. 4) argue for fractionation of plagioclase, clinopyroxene,Fe–Ti oxides and olivine within each group in different proportions.The more pronounced negative Sr and Eu anomalies observed in theevolved high-Ti basalts suggest that plagioclase fractionation has beenmore extensive in this group than in the other groups. Even if olivinephenocrysts are always very rare in CAMP basalts (but not in CAMPintrusions; e.g. Deckart et al., 2005), the evolved character of thebasalts suggests that this phase probably crystallized from the mostprimitive magmas. Olivine phenocrysts were probably subsequentlyresorbed by reaction with the more Fe-rich evolved magma.

5.3.2. Modelling lithospheric contamination of the WMBT

5.3.2.1. Modelling of crustal contamination of the evolved high-Ti WMBTthrough an AFC process. Considering that the evolved high-Ti WMBThave Sr–Nd–Pb isotopic compositions broadly similar to those of thehigh-TiWMBT (Figs. 5 and 6), we suggest that theywere derived froma similar sub-continental asthenospheric source. However, assumingthat sample M10 is representative, the evolved high-Ti WMBT haverelatively high 187Os/188Osi suggestive of a significant continentalcrustal contribution. This interpretation, which will be modelledbelow, is consistent with (1) higher enrichment in LREE and LILE,(2) slightly stronger negative Nb anomalies, and (3) higher initial87Sr/86Sr and slightly lower εNd of the evolved high-Ti WMBT (Figs. 4,6 and 8).

Here we try to quantitatively evaluate the extent of crustalassimilation. This cannot be constrained from Sr–Nd–Pb isotopiccompositions alone because: 1) we do not know the initial primarycomposition of the mantle melts; 2) data are available for only a fewWMBT samples (high-Ti and evolved high-Ti rocks in particular arerare); 3) the only isotopic data published for the local basement arefrom a section of the Amazonian craton (Andorinhas greenstone belt)that has very high Sr and very low Pb isotopic ratios (measured 206Pb/204Pb=13.9–17.1; 87Sr/86Sr back calculated to 199 Ma=0.710568–2.577866, de Souza et al., 2001; Gibbs et al., 1986) and which thuscannot plausibly represent the crustal contaminant involved in the

genesis of theWMBT basalts. In contrast, Os isotopesmaymore tightlyconstrain crustal contamination processes due to the large contrastbetween typical mantle and crustal compositions.

The effect of continental crust contamination has been modelledassuming an Assimilation-Fractional-Crystallization (AFC) process(DePaolo, 1981). We considered as a target the Os compositions of thesample M10. For the initial primary melt we chose a typicalasthenospheric (PUM) Os isotopic composition for the evolved high-Timagmas (187Os/188Os=0.129). Since no data are available for the localcrust, we used the average continental crustal Os parameters (187Os/188Os=1.05 and [Os]=31 ppt; Peucker-Ehrenbrink and Jahn, 2001).Further parameters are given in Table A5.

The evolved high-Ti magmas may be obtained from a primarymagma with picritic Os content (~1600 ppt) by 10.9% crustalassimilation. This degree of assimilation by an AFC process (para-meters: Dsr=0.6, F and r as in Table A5)may explain the shift of the Srisotopic ratios from an asthenosphere-derived parental magma (87Sr/88Sr=0.70220, [Sr]=188 ppm) to those measured for the evolvedhigh-Ti WMBT (87Sr/88Sr=0.70341) if the contaminant is typicalcontinental crust (87Sr/88Srcont crust=0.7100, [Sr]=235 ppm). How-ever, derivation of the evolved high-Ti basalts from the high-Ti basaltsvia a simple AFC process is not easy to reconcile with the Pb isotopiccompositions, since the evolved high-Ti WMBT have lower 207Pb/204Pb, higher 206Pb/204Pb and similar 208Pb/204Pb, and are thus closerto the NHRL (Fig. 5) than the high-Ti basalts. Therefore, a slightdifference in the composition of the primary magmas and thus in themantle source, or in the relative contributions of the asthenosphere vsthe SCLM may be envisaged for the two groups.

5.3.2.2. Modelling mixing between asthenosphere-derived high-Timagmas and SCLM-derived melts. Since plagioclase cannot crystallizeunder the high pressures that exist in the lithosphericmantle, it seemslikely that most of the fractional crystallisation occurred after, ratherthan during the mixing process. This would suggest that the observedisotopic compositions reflect mixing between two primitive melts,one consisting of Mg-rich melts derived from the asthenosphere andthe other of ultra-alkaline liquids from the SCLM.

To model the possible mixing between asthenosphere-derivedmagmas and ultra-alkaline liquids from the SCLM, we consider thesample with the lowest Os initial ratio (M13) which therefore may bethe least contaminated by the continental crust and the most by theSCLM. Note that the initial 187Os/188Os ratios of the two high-Tisamples are indistinguishable considering the uncertainties, thus weuse the nominal isotopic composition of M13 only as an example toillustrate the possible effects of mixing.

Wemodelled amixing process affecting an asthenosphere-derivedprimitive magma (187Os/188Os=0.129; [Os]≈1000 ppt; 87Sr/88Sr=0.70220; [Sr]=180 ppm; Carlson, 2005; Sun and McDonough, 1989;Workman and Hart, 2005). The contaminant is assumed to be aprimitive ultra-alkaline mafic melt (e.g. lamproites, kimberlites, andkamafugites) which has been proven to best represent a liquidderived from partial melting of a metasomatised off-craton SCLM (e.g.Karoo CFB, Heinonen et al., 2010; Paraná-Etendeka CFB; Gibson et al.,1995). Considering that no ultra-alkaline mafic melts are associatedwith the CAMP, we used extreme values of the ultra-alkaline maficmelts associated with the Paraná-Etendeka CFB to define the geo-chemical characteristics of the contaminant (187Os/188Os=0.122;[Os]≈3000 ppt; 87Sr/88Sr=0.70455; [Sr]≈1000 ppm; Araujo et al.,2001; Carlson et al., 1996, 2007; Comin-Chiaramonti et al., 1997;Gibson et al., 1995) in order to estimate the maximum contribution ofthe SCLMmelts in themixing. The Os concentration of this componentis rather high, but is comparable to those of highly magnesian dykesfrom the Karoo province of Antarctica (Heinonen et al., 2010). TheM13 characteristics could be matched by mixing of asthenosphere-derived magmas with 9.4% of ultra-alkaline silicate melts from theSCLM.

149R. Merle et al. / Lithos 122 (2011) 137–151

5.3.2.3. Modelling of the contamination of the low-Ti magmas by thecrust and the SCLM. We have also tried to evaluate also whether thelow-Ti WMBT magmas originated either from partial melting of ametasomatised SCLM, with possible contamination by the continentalcrust through an AFC process or from mixing of an asthenosphere-derived magma mixed with ultra-alkaline melts from a metasoma-tised SCLM.

In order to model the AFC process, we used the composition of thesample with the highest Os initial ratio (M19), which was thereforepotentially the most contaminated by the upper continental crust.However, as was the case with the high Ti samples, the initial 187Os/188Os ratios of the low Ti samples are indistinguishable within theuncertainties, so the nominal isotopic composition of M19 is used toillustrate the effects of the AFC process. We assumed an off-cratonSCLM-derived melt (187Os/188Os=0.125 according to Carlson, 2005;[Os]≈1100 ppt; 87Sr/88Sr=0.706; [Sr]=160 ppm) assimilating atypical upper continental crust (187Os/188Os=1.05; [Os]=31 ppt;87Sr/88Sr=0.710; [Sr]=230 ppm). The compositions of basalt M19 ismatched for only a negligible degree of 2.6% of crustal assimilation(DOs=8.5; DSr=0.8; F=0.7; r=0.08; Table A5). Even lower amountsof assimilation would be required if the non-contaminated low-Timagmas were assumed to be derived from a source with an Oscomposition similar to the PUM value (187Os/188Os~0.128 at 199 Ma).

Alternatively, we have considered the possibility that the low-TiWMBT formed from mixing of asthenosphere-derived magmas withultra-alkaline mafic liquids from the SCLM. As for the high-Ti group,we tried to model the composition of the sample suspected to be theleast contaminated by the continental crust and the most by the SCLMamong the low-Ti WMBT (sample M26) and we assume that themixing process involved primitive, high MgO melts, which would beexpected to have high Os contents. We assumed an asthenosphere-derived parental magma (187Os/188Os=0.129; [Os]≈1500 ppt; 87Sr/88Sr=0.70220; [Sr]≈160 ppm; Carlson, 2005; Sun and McDonough,1989; Workman and Hart, 2005) and a contaminant slightly distinctfrom those for the high-Ti group (metasomatised off-craton SCLM-derived mafic ultra-alkaline melts: 187Os/188Os= 0.125;[Os]≈2000 ppt; 87Sr/88Sr=0.7076; [Sr]≈1500 ppm). The isotopiccharacteristics of sample M26 can be modelled by a contribution ofSCLM-derived ultra-alkaline melts of~52%.

6. Geodynamic significance of CAMP magmatism from theMaranhão basin

Our geochemical data suggest that the high-Ti group is derivedfrom magmas that originated from an asthenosphere-like sourcecontaminated by the SCLM while the evolved high-Ti basalts werecontaminated by the continental crust. The low-Ti group yieldsinstead a dominant lithospheric signature which may be derived frommetasomatically enriched portions of the SCLM. Due to the slight(high-Ti magmas) or dominant (low-Ti magmas) SCLM component,any sub-lithospheric geochemical signature is either diluted orswamped out, rendering the recognition of a clear mantle-plumegeochemical signature problematic. Nonetheless, we cannot exclude acontribution from a deep mantle plume, in particular as a purveyor ofheat and as a triggering mechanism for the mantle melting. However,since our data do not explicitly point to the presence of a plume, otherheat sources that could potentially generate WMBT magmatism,should also be examined.

Two mechanisms are considered, in particular: 1) edge-drivenconvectionwhich consists of small-scale convecting cells developed atthe edges of cratonic roots by strong density and viscosity contrastsbetween a cold lithospheric mantle keel and the hotter asthenosphere(King and Anderson, 1998; King and Ritsema, 2000); and 2) large-scale mantle warming under supercontinents such as Pangaea, whichoccurs by accumulation of the internal heat beneath the insulatinglithosphere (Coltice et al., 2007, 2009). This process may lead to an

increase of temperature up to 100 °C (Coltice et al., 2007). Edge-driven convection is likely to develop in the geological context of theMaranhão basin since Pangaean rifting occurred along a mobile beltbordering the Amazonian craton (Fig. 1), thought to be underlain by adeep lithospheric keel. Indeed, such a mechanism has already beenproposed for the genesis of the other Brazilian CAMP occurrences (DeMin et al., 2003). Heat incubation is also likely to occur under themega-continent Pangea, particularly considering that in the Maran-hão region no significant basic magmatism occurred during thePaleozoic and the Triassic, before eruption of the CAMP basalts.Therefore, we propose that at the time of Pangea break-up, both edge-driven convection and heat-incubation acted as heat purveyors.Considering that a mantle-plume may have further contributed heatand that an extensional tectonic regime was active and possiblyinduced lithospheric thinning and thus general decompression, it isreasonable to envisage partial melting of the asthenospheric mantleunderneath the Maranhão basin. The increased heat flux may havepossibly also affected the lower SCLM and induced melting of its mostenriched and thus most fertile portions. These metasomaticallyenriched portions were likely formed during one of the severalorogenic events that affected this region from the Proterozoic to theearly Paleozoic (e.g., Brasiliano orogen). Since the Brazilian SCLMprobably did not experience any significant melting episodes beforethe CAMP event, these enriched components were probably pre-served. Indeed, there is no evidence in this region of SCLM-derivedmagmatic rocks such as lamproites and carbonatites of Paleozoic orTriassic age. As a consequence, the SCLM underneath the Maranhãobasin might contain large portions of metasomatic veins with meltingtemperatures more than 150 °C lower than the peridotitic mantle(e.g., Kogiso et al., 2003). Therefore, near-synchronous melting ofthese low-temperature SCLM materials and of the sub-lithosphericasthenosphere may have occurred due to the combined effects of heatincubation, edge-driven convection and possibly mantle-plumeimpingement at the base of the SCLM.

In summary, our data do not rule out the occurrence of a mantleplume of nearly asthenospheric composition as a supplier of heat andmaterial but we suggest rather that the CAMP event was triggered bythe combined effects of shallow mantle processes.

7. Summary and conclusions

Major and trace element and isotopic analyses as well as 40Ar/39Arage determinations were performed on newly sampled CAMP basaltsfrom the Maranhão basin of Brazil. Three chemical groups wereidentified: low-Ti, high-Ti and evolved high-Ti tholeiites, the latterbeing the first high-Ti CAMP basalts sampled up to 700 km inland. Themain results of this study are:

(1) The new 40Ar/39Ar plateau ages obtained from plagioclaseseparates of high-Ti (199.7±2.4 Ma) and evolved high-TiWMBT (197.2±0.5 Ma and 198.2 ± 0.6 Ma) are indistinguish-able from those of previously analyzed low-Ti WMBT (198.5±0.8 Ma) and from the mean 40Ar/39Ar age of the CAMP (mean199±2.4 Ma).

(2) The low-Ti WMBT derived either from partial melting of themost fusible parts of the metasomatised continental litho-spheric mantle or from contamination of an asthenosphericmelt by ultra-alkaline melts from the metasomatised SCLM.The high-Ti WMBT yield a dominant asthenospheric signaturewith a minor contribution from the SCLM. The evolved high-Tibasalts are the only analyzed samples which yield clearevidence of continental crust assimilation.

(3) Indistinguishable ages of the three chemical groups imply sub-contemporaneous melting of lithosphere and asthenosphere.

(4) CAMP magmatism in the Maranhão basin may be attributed tolocal hotter mantle conditions due to the combined effects of

150 R. Merle et al. / Lithos 122 (2011) 137–151

edge-driven convection and large-scale mantle warming underthe Pangea supercontinent. The mantle melting might wellhave been triggered by a mantle-plume with asthenosphere-like isotopic compositions.

Supplementarymaterials related to this article can be found onlineat doi:10.1016/j.lithos.2010.12.010.

Acknowledgments

We acknowledge P. Capiez for XRF analyses, R. Carampin forElectron Microprobe analyses, C. Douchet for the ICP-MS analyses andD. Fontignie for the TIMS analyses. Dr. S.A. Gibson and an anonymousreviewer and Dr. A. Kerr as editor are thanked for comments on themanuscript. Financial support from the Fondo Ateneo-Università diPadova, CARIPARO (to AM) and PRIN-2005 (to GB) is acknowledged.

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