ORIGINAL PAPER

Chemical and mineralogical evidence of the occurrenceof mantle metasomatism by carbonate-rich meltsin an oceanic environment (Santiago Island, Cape Verde)

Sofia Martins & João Mata & José Munhá &

Maria Hermínia Mendes & Claude Maerschalk &

Rita Caldeira & Nadine Mattielli

Received: 29 December 2008 /Accepted: 8 September 2009# Springer-Verlag 2009

Abstract Lavas from Santiago Island attest to a complexmagmatic history, in which heterogeneous mantle source(s)and the interactions of advecting magmas with thick meta-somatised oceanic lithosphere played an important role in theobserved isotopic and trace element signatures. Young(<3.3 Ma) primitive lavas from Santiago Island are charac-

terised by pronounced negative K anomalies and traceelement systematics indicating that during partial meltingDK>DCe. These features suggest equilibration with anoceanic lithospheric mantle containing K-rich hydrousmineral assemblages, consistent with the occurrence ofamphibole + phlogopite in associated metasomatised lherzo-lite xenoliths, where orthopyroxene is partially replaced bynewly formed olivine + (CO2 + spinel + carbonate inclusion-rich) clinopyroxene. Metasomatism induced a decrease inameltSiO2

and Ti/Eu ratios, as well as an increase in fO2, Ca/Scand Sr/Sm in the Santiago magmas, suggesting a carbonatiticcomposition for the metasomatic agent. Santiago primitivelavas are highly enriched in incompatible elements and showa moderate range in isotopic compositions (87Sr/86Sr=0.70318–0.70391, 143Nd/144Nd=0.51261–0.51287,176Hf/177Hf=0.28284–0.28297). Elemental and isotopic sig-natures suggest the involvement of HIMU and EM1-typemantle end-members, in agreement with the overall isotopiccharacteristics of the southern Cape Verde Islands. Theoverall geochemical characteristics of lavas from SantiagoIsland allow us to consider the EM1-like end-member asresulting from the involvement of subcontinental lithosphericmantle in the genesis of magmas on Santiago.

Introduction

The highly variable elemental and isotopic composition ofOceanic Island Basalts (OIB) has been thought to resultfrom sampling of a heterogeneous mantle (Gast et al. 1964).Such heterogeneity is a consequence of mantle differenti-ation due to partial melting and crust formation, subductionrecycling of crustal components (oceanic crust and sedi-ments) and the possible existence of delaminated subcon-

Editorial handling: A. Möller

Electronic supplementary material The online version of this article(doi:10.1007/s00710-009-0078-x) contains supplementary material,which is available to authorized users.

S. Martins : J. Mata : J. MunháFaculdade de Ciências da Universidade de Lisboa,Departamento de Geologia (GeoFCUL),Campo Grande, C6,1749-016 Lisboa, Portugal

S. Martins (*) : J. Mata : J. Munhá :R. CaldeiraCentro de Geologia da Universidade de Lisboa (CeGUL),Campo Grande, C6,1749-016 Lisboa, Portugale-mail: smmartins@fc.ul.pt

M. H. MendesIICT - DES, Lisboa Portugal;Instituto de Investigação Científica e Tropical,Alameda D. Afonso Henriques 41-4° D.,1000-123 Lisboa, Portugal

C. Maerschalk :N. MattielliDepartment of Earth and Environmental Sciences, DSTE,CP 160/02, Université Libre de Bruxelles,Avenue F.D. Roosevelt, 50,Brussels 1050, Belgium

R. CaldeiraINETI - Geology Department,Estrada da Portela - Zambujal,2721-866 Alfragide, Portugal

Miner PetrolDOI 10.1007/s00710-009-0078-x

tinental lithospheric blocks in the oceanic mantle (e.g.White and Hofmann 1982; Hofmann 2003; Stracke et al.2005). It has also been shown, for example by a study ofthe Cape Verde archipelago that the contribution of crustalassimilation to the composition of OIBs cannot beneglected (Hoernle et al. 1991; Doucelance et al. 2003;Escrig et al. 2005; Millet et al. 2008), indicating that thesublithospheric oceanic mantle may be less heterogeneousthan suggested by the OIB compositional variability (Milletet al. 2008).

The majority of OIBs are thought to originate from theadiabatic melting of mantle plumes. The original composi-tion of plume magmas is also prone to substantialmodification by interaction with lithospheric mantle duringtheir ascent (e.g. Class and Goldstein 1997). This isespecially pronounced when relatively low solidus fertiliseddomains developed in the lithosphere as a result of previousmetasomatic events. In fact, the contribution of mantlemetasomatism to alkaline magmatism has long beenestablished, mainly through the study of mantle xenoliths(e.g. Mattielli et al. 1999; Grégoire et al. 2000; Ionov et al.2002; Bodinier et al. 2004), but also from the geochemicalcharacteristics of lavas (e.g. Class and Goldstein 1997;Späth et al. 2001; Pilet et al. 2008).

Elemental and isotopic (Sr, Nd and Pb) heterogeneitiesof the Cape Verde archipelago lavas are usually interpretedas the result of mixing, in different proportions, of HIMU-,EM1- and DMM-type mantle components, whereas the Heisotopes require the contribution of a relatively undegassedand unradiogenic reservoir, i.e., the lower mantle (Gerlachet al. 1988; Christensen et al. 2001; Doucelance et al. 2003;Holm et al. 2006). In addition, assimilation of the island’sbasaltic crustal basement and of carbonatitic rocks has beeninvoked to explain the peculiarities of São Nicolau Island(Doucelance et al. 2003; Millet et al. 2008). Nevertheless,evidence gathered from geochemical studies of the eruptedlavas (Jørgensen and Holm 2002) and their associatedultramafic xenoliths (Ryabchikov et al. 1995; Bonadiman etal. 2005) suggests a significant role of mantle metasoma-tism in this archipelago, characterised by the occurrence ofhighly SiO2-undersaturated lavas (e.g. melanephelinites andmelilitites) and by exceptional amounts of outcroppingcarbonatites (e.g. Silva et al. 1981; Hoernle et al. 2002;Mourão et al. 2009). Both the nature and the mechanism ofmantle metasomatism in Cape Verde are far from settled,and even its existence has been challenged based on thestudy of lavas (Trindade et al. 2003) and xenoliths (Shaw etal. 2006).

This paper reports new petrological and geochemicaldata for Santiago Island’s lavas and ultramafic xenoliths,demonstrating the occurrence of lithospheric metasoma-tism. The nature of the mantle’s residual mineralogy andthe nature of oceanic mantle metasomatism will be

discussed in order to address the influence of magma-lithosphere interaction on the composition of magmaserupted within an oceanic intraplate setting. In addition,the origin of the EM1-type mantle end-member alreadydescribed for Santiago (Gerlach et al. 1988; Davies et al.1989; Doucelance et al. 2003; Escrig et al. 2005) isdiscussed based on new elemental and isotopic data,including the first Hf isotopic data reported for Cape Verde.

Geological setting

The Cape Verde archipelago (15–17°N, 23–26°W) islocated in the Central Atlantic Ocean, 500 km off theSenegalese coast and 2,000 km to the east of the Mid-Atlantic Ridge. It is situated on a topographic swell risingmore than 2 km above the regional oceanic floor where thelithosphere is ≈ 80 km thick (Courtney and White 1986).The islands are located between the M0 and M16 magneticanomalies, corresponding to 120–140 Ma old oceanic crust(Williams et al. 1990). The crust is significantly thickenedunderneath the islands, and the oceanic plateau is associatedwith geoid, gravity, seismic and heat flow anomalies,interpreted as the result of a mantle plume at depth (e.g.Courtney and White 1986; Lodge and Helffrich 2006; Pimet al. 2008). Plume activity was recently imaged down tothe core-mantle boundary by seismic tomography (Montelliet al. 2006). Such a lower mantle origin for the Cape Verdeplume is supported by the unradiogenic He isotopicsignature (R/Ra up to 15), which characterises both silicateand carbonatitic Cape Verde rocks (e.g. Doucelance et al.2003, 2007; Mata et al. 2006, 2009; Mourão et al. 2007).The archipelago consists of ten major islands, sub-dividedinto two groups (northern: Santo Antão, São Vicente, SantaLuzia, São Nicolau, Sal and Boavista; southern: Brava,Fogo, Santiago and Maio) characterised by significantdifferences in degree of silica saturation, trace elementpatterns and Sr, Nd, Pb and He isotopic signatures (e.g.Gerlach et al. 1988; Doucelance et al. 2003).

Although magmatic activity in the archipelago hasoccurred since 30 Ma (Torres et al. 2002; Holm et al.2008) the islands show neither a typical linear spatialdistribution nor a simple age progression. They arearranged in a horseshoe configuration opening westwards.The volcanism seems to have started on the eastern islandsbut remained sub-aerially active in most of the archipelagountil 1 Ma (Holm et al. 2008). The lack of a clear “age toplume distance correlation” has been related to theproximity of the African plate rotation pole (e.g. Grippand Gordon 2002), as well as the existence of translitho-spheric fractures constraining the upward magma transferfrom mantle to crust and, consequently, the locations ofsome of the islands (Torres et al. 1998).

S. Martins et al.

Santiago is the largest island on the archipelago with a totalarea of 991 km2. Santiago belongs to the Cape Verdesouthern islands group, occurring between Maio and Fogoand lying on an oceanic crust about 20-km thick (Lodge andHelffrich 2006). The island has an N-S elongated shape andis characterised by rougher relief (up to 1,392 m) ascompared with the neighbouring flat island of Maio, locatedsome 20 km to the east. This attests to the tendency for awestward age decrease (Serralheiro 1976, Madeira et al.2009), which is more evident in the southern islands (forradiometric ages see Gerlach et al. 1988; Torres et al. 2002;Jørgensen and Holm 2002; Plesner et al. 2002; Madeira et al.2005; Holm et al. 2006, 2008; Duprat et al. 2007).

The stratigraphy of Santiago Island, established bySerralheiro (1976) and Alves et al. (1979), comprises sixunits. Listed from base to top (Fig. 1), they are the OldEruptive Complex and the Flamengos, Orgãos, Pico daAntónia, Assomada and Monte das Vacas formations.Orgãos is the only sedimentary formation distinguished inSantiago’s stratigraphy, although some calcarenite levelswere mapped inside the Pico da Antónia Formation. After

a period of intense erosion, which exposed the plutonicbasement (Old Eruptive Complex; ≥ 20 Ma, Gerlach et al.1988), submarine volcanic activity (comprising lavas,breccias and pyroclasts) covered extensive areas, generat-ing the Flamengos Formation (4.5 to 5.5 Ma; Holm et al.2008). During a subsequent period of volcanic quiescence,an important erosion phase occurred. The existing reliefwas significantly dismantled, resulting in extensive lahar-type deposits grouped in the Orgãos Formation. The Picoda Antónia Formation testifies to the shield stage ofSantiago (3.3–2.3 Ma; Holm et al. 2008), characterised byvoluminous Hawaiian-type submarine and subaerial vol-canism. Previous studies (e.g. Martins et al. 2003) haveidentified two chemically distinct groups within the Picoda Antónia Formation, strongly suggesting the existenceof two non-coeval lithostratigraphic units. On the basis oftheir chemical characteristics and consistent field rela-tions, the present paper refers to subaerial lavas as theUpper Pico da Antónia Formation (UPA) and to submarinelavas as the Lower Pico da Antónia Formation (LPA). Thisassertion is supported by dating performed by Holm et al.

Fig. 1 Location of the Cape Verde archipelago and geological map of Santiago Island [simplified from Serralheiro (1976) and Alves et al. (1979)]

Chemical and mineralogical evidence of mantle metasomatism

(2008), who reported older dates for the submarine lavas(2.9–2.8 Ma) than for the main subaerial event (2.6–2.3 Ma) of the Pico da Antónia Formation. The AssomadaFormation occupies a limited portion of the central part ofthe island, displaying a trapezoidal configuration (10 kmacross); it is composed of extrusive subaerial lava flows andpyroclasts. The last volcanic manifestations of Santiagowere mainly strombolian, being grouped in the Monte dasVacas Formation with approximately 50 pyroclastic conesand small lava flows, dispersed along the island (Fig. 1).These two post-shield formations span an age range of0.4 m.y., lasting from 1.1 to 0.7 Ma (Holm et al. 2008).

Analytical procedures

We collected 62 lava samples from the entire volcanicsequence of Santiago in order to obtain a representativesample set of the volcanostratigraphic units. For bulkgeochemical analysis, 49 lava samples were selected onthe basis of their petrographic characteristics and freshness.In addition, 35 ultramafic xenolith samples were alsoselected for petrological and geochemical studies.

At the Departamento de Geologia da Universidade deLisboa, the lava samples were crushed with a hydraulic press,to remove all visible signs of alteration, and then reduced in sizeby a jaw crusher and powdered in an agate ring and puck mill.

The major- and trace-element contents of the sampleswere measured using inductively coupled plasma-opticalemission spectrometry (ICP-OES) and inductively coupledplasma-mass spectrometry (ICP-MS), respectively, at theActlabs laboratory, Canada. Alkaline dissolution withlithium metaborate/tetraborate followed by nitric aciddissolution was performed for all analyses, except for thedetermination of Cd, Cu, Ni and Zn, for which aciddigestion was carried out. The calculated reproducibility is≈ 1% for major-element contents (SiO2, Fe2O3, MnO, MgOand CaO), <4% for rare earth elements (REE) and ≈ 5% forhigh field strength elements (HFSE).

Mineral analyses were performed on carbon-coatedpolished thin sections using a JEOL 733 Electron Micro-probe in wavelength dispersive mode at the Centro deGeologia da Universidade de Lisboa. Minerals wereanalysed with an acceleration voltage of 15 kV and acurrent of 10 nA, using a beam 5 μm wide.

The isotopic analyses and chemical preparation of 18samples were performed in the clean laboratories of theDepartment of Earth and Environmental Sciences of the“Université Libre de Bruxelles” (Belgium). To removesecondary phases and any potential contamination, aleaching procedure was applied to rock powder samples(about 250 mg), with repeated additions of 6 N sub-boiled HCl (five or six times) followed by 30-minutes

ultrasonic baths until a clear solution was obtained.Samples were then rinsed twice with milli-Q water andtreated with ultrasound for 30 min (see Weis and Frey1996). The dissolution was performed with a mixture ofsub-boiled HF (48%) and HNO3 (14 N) acids in closedSavillex® beakers on a hot plate for 48 h at 130°C. Afterevaporation, the samples were re-dissolved in 6 N sub-boiled HCl for another 24 h at 130°C. The Sr, Nd and Hffractions were all extracted from the same initial samplesolution. Sr, REE and Hf were isolated using a columnloaded with 18.5 ml of Dowex AG50W-X8 100–200 meshcation exchange resin. Nd was recovered by loading theREE fraction onto a column with HDEHP-coated Teflon.Hafnium was separated following the slightly modifiedtwo-column procedure described in Blichert-Toft et al.(1997), involving anion and cation exchange columns(Mattielli et al. 2002; Weis et al. 2006). Measurements ofthe total procedural Nd, Sr and Hf blanks gave 28 pg,400 pg and 23 pg, respectively. High Sr blanks reflect themuch higher Sr concentration in the rock. The isotopiccompositions of Nd and Hf were determined on a Nuplasma multi-collector inductively coupled mass spec-trometer (MC-ICP-MS) at the Department of Earth andEnvironmental Sciences of the “Université Libre deBruxelles”. Standards were systematically run betweenevery two samples to monitor the instrumental mass biasduring the analysis session, resulting in 143Nd/144Nd=0.511935±20 (2σ, n=57) for the Rennes Nd standardand 176Hf/177Hf=0.282143±19 (2σ, n=30) for the JMC475 Hf standard. The Nd and Hf isotopic measurementswere internally normalised to 179Hf/177Hf=0.7325 and146Nd/144Nd=0.7219, respectively.

The Sr isotopic ratios were analysed on the FinniganTriton Thermo-Ionization Mass Spectrometer (TIMS) at thePacific Centre for Isotopic and Geochemical Research atthe University of British Columbia (Canada). The NBS 987Sr standard was run after every ten samples, and gave amean value of 87Sr/86Sr=0.710244±7 (2σ) for six measure-ments. The raw data were corrected for mass fractionationusing 86Sr/88Sr=0.1194.

The quality of chemical preparation procedures andreproducibility of isotopic measurements can be assessedby duplicate analyses of three samples (Table 4), which arebetter than 66 ppm, 51 ppm and 25 ppm for 143Nd/144Nd,176Hf/177Hf and 87Sr/86Sr, respectively.

Santiago lavas

Petrology

Considering the main objectives of our study, this paperfocuses on mafic rocks, even though evolved lithotypes

S. Martins et al.

such as trachytes and phonolites also occur on Santiago(Serralheiro 1976; Alves et al. 1979). Detailed petrographicanalysis of the Santiago lavas (Martins 2003) indicates thatfor mafic rocks there is little lithological variation amongthe different volcano-stratigraphic units. On the whole,Santiago primitive lavas range in composition frombasanites, melanephelinites and nephelinites to melilitites,with a general predominance of melanephelinites (seebelow). They are usually porphyritic, with a glassygroundmass predominating among the submarine lavas.Phenocrysts are mainly olivine and clinopyroxene; theirrelative abundance depends on the lithotype considered. Inorder to ensure that analysed samples are representative ofmagmatic liquids, only those with less than 10–15%phenocrysts were selected for analysis.

Petrographic and mineral chemistry data indicate that theSantiago mafic magmas were variably modified by crystalfractionation processes. The overall fractionation sequencewas dominated by early crystallisation of chromian spinel(Cr/ΣR+3=0.504–0.230) + olivine (Fo=88–85 mol%) ± Cr-rich diopside (in melanephelinites), followed by olivine(Fo=85–70 mol%) + Al- Ti-rich diopside + Ti-magnetite +apatite ± plagioclase/nepheline/melilite assemblages. Pla-gioclase (mostly labradorite in basanites) and nepheline (innephelinites) fractionation becomes increasingly importantat more advanced stages of fractionation, whereas (primary)globular carbonate aggregates occur sporadically in thegroundmass of melilitites, suggesting late-stage develop-ment of carbonate-silicate liquid immiscibility.

Liquidus temperatures were calculated according to both aMgO-in whole-rock (Helz and Thornber 1987) and a olivine-liquid (Putirka 2005) geothermometer, by using olivinephenocryst and whole-rock compositions (Tables 1, 2 and 3)that had previously been readjusted (by iteration with fO2

estimates) to proper Fe2O3/FeO values (Kilinic et al. 1983)and corrected for potential olivine or clinopyroxene accu-mulation. The two geothermometers gave similar results(±30°C; 2σ). For primitive samples (olivine Fo>85 mol %),the estimated liquidus temperatures were 1,159–1,212°C(MgO-in; Helz and Thornber 1987) and 1,171–1,258°C(olivine-liquid; Putirka 2005), whereas differentiated basan-ites/nephelinites indicate Mg-in liquidus temperatures, aswell as olivine- and clinopyroxene-liquid (Putirka et al.2003) equilibria, within the range of 1,080–1,160°C.

Clinopyroxene-liquid geobarometry (Putirka et al. 2003)indicated that magmatic differentiation was polybaric; themain fractionation occurred at multiple levels within themantle (average depth ∼30±5 km; P=0.89±0.15 GPa) andwas followed by temporary stagnation within crustalmagma chambers (2–15 km depth) before eruption.

Magmatic silica activity values ameltSiO2

� �were estimated

by applying Ghiorso’s silicate melt solution model (Ghiorsoet al. 1983; updated by Ghiorso and Sack 1995) to

corrected whole-rock compositions at the P-T conditionsspecified above. Primitive Santiago lavas display a widespectrum of ameltSiO2

� �values, ranging from 0.080 to 0.149 in

melilitites, through 0.213–0.297 in melanephelinites, to0.329 in basanites. These values are about 0.7 to 0.1 logunits below those defined by the olivine-orthopyroxeneequilibrium (T=1,160–1,210°C; P=1 GPa; O’Neill andWall 1987), and the corresponding corrections have beenapplied throughout the calculations for estimating fO2

(Ballhaus et al. 1991) from coexisting olivine phenocrystsand chromian-spinels. Results for primitive samples indi-cate initial fO2 conditions more oxidised than O’Neill’s(1987) QFM values (Δlog (fO2)QFM=0.3 to 1.5; averageΔlog(fO2)QFM=1.1±0.3 (2σ)), similar to those reported byBallhaus (1993) for OIBs elsewhere (see also Mata andMunhá 2004). Further crystallisation proceeded underdecreasing oxidation conditions, as indicated by low Δlog(fO2)QFM (≈−0.1) values estimated for evolved (Foolivine=71–72 mole %) basanitic magmas.

Whole-rock elemental geochemistry

For this study, 49 of the collected samples were selected onthe basis of their representativeness and freshness. Table 1reports the whole-rock analyses (major and trace elements)of twelve representative samples. In order to evaluatealteration effects, the most mobile elements (K, Rb, Srand Ba) were projected against LOI (not shown); theabsence of any significant correlation indicates that theanalytical results are not biased by weathering.

All the analysed Santiago samples are ultrabasic (SiO2<45 wt.%) alkaline lavas characterised by highly SiO2-undersaturated compositions (basanites, melanephelinites,nephelinites and melilitites; Le Maitre et al. 2002) with Niranging from 54 to 390 ppm and Cr ranging from 54 to1,040 ppm (Table 1). Mg# values range from 41.8 to 75.7(Table 1), supporting petrographic/mineral chemistry obser-vations (see above) about the role of crystal fractionation onthe observed chemical variability. For example, Al2O3/CaOratios tend to increase with decreasing Mg#, showing therole of clinopyroxene on the fractionation processes whichdrove magma compositions from subchondritic to supra-chondritic ratios.

However, significant differences in Na2O/K2O and CIPWnormative olivine/diopside ratios between lavas from differ-ent complexes suggest time-dependent variations in themajor element compositions of Santiago primary magmas.For example, Flamengos Formation lavas (4.5–5.5 Ma; Holmet al. 2008) are characterised by lower normative olivine/diopside x ¼ 0:45ð Þ and Na2O/K2O ratios (0.62–3.65) thanlavas from the younger Assomada Formation (0.7–1.1 Ma),which are of a similar degree of magmatic evolution(x olivine=diopside ¼ 0:73; Na2O/K2O up to 5.31).

Chemical and mineralogical evidence of mantle metasomatism

Tab

le1

Major

andtraceanalyses

ofSantiago

lavas

Formation

Mon

tedasVacas

Assom

ada

Upp

erPicoda

Antón

iaLow

erPicoda

Antón

iaFlamengo

s

Sam

ple

ST-16

ST-24

ST-18

ST-19

ST-21

ST-27

ST-48

ST-50

ST-36

ST-37

ST-5

ST-34

Rocktype

aBas.

Bas.

Melil.

Bas.

Olv.Mel.

Olv.Mel.

Meli.Neph.

Bas.

Bas.

Olv.Mel.

Px.

Mel

Olv.Mel.

Major

elem

ents(w

t.%)

SiO

243

.39

43.47

36.49

42.32

40.66

41.27

38.96

44.21

43.63

41.39

40.98

40.81

TiO

23.69

4.05

3.43

3.18

3.41

2.90

3.06

2.64

3.04

3.26

4.18

3.33

Al 2O3

15.43

14.17

10.24

13.34

10.10

10.06

9.68

12.91

13.21

11.32

12.18

11.18

Fe 2O3

2.00

3.36

1.97

1.99

2.02

1.92

1.79

1.86

1.73

1.88

1.84

1.88

FeO

10.02

11.21

9.86

9.93

10.10

9.62

8.93

9.32

8.66

9.42

9.20

9.40

MnO

0.19

0.22

0.25

0.18

0.18

0.18

0.18

0.17

0.17

0.17

0.18

0.18

MgO

6.25

5.59

12.12

9.93

15.13

14.80

15.64

10.51

9.60

13.33

9.33

13.31

CaO

9.77

8.94

15.76

10.22

11.98

12.55

13.81

11.96

12.25

12.02

11.87

12.51

Na 2O

2.89

3.85

3.43

2.71

2.21

2.65

2.84

2.85

2.74

1.87

2.09

2.45

K2O

2.52

2.80

1.45

1.56

0.46

0.53

0.75

1.27

1.02

1.76

1.99

1.74

P2O5

0.64

0.74

1.26

0.55

0.38

0.67

0.92

0.55

0.41

0.46

0.61

0.57

LOIb

1.60

0.04

1.62

3.13

1.54

1.90

1.90

0.09

2.09

1.76

3.90

0.93

Total

98.39

98.44

97.87

99.03

98.16

99.05

98.46

98.34

98.55

98.64

98.34

98.29

Mg#

c52

.64

47.05

68.66

64.06

72.75

73.28

75.63

66.77

66.40

71.61

64.38

71.61

Trace

elem

ents(ppm

)Cr

78bd

ld48

315

596

179

385

449

054

476

252

875

9Ni

298

251

134

328

387

390

227

176

323

155

284

Ga

1923

159

1515

1711

1914

2119

Rb

5068

3046

4120

2723

2338

5140

Sr

1,22

41,23

12,34

691

062

71,05

71,16

185

475

969

81,07

982

7Zr

131

214

254

178

175

171

263

155

227

184

284

283

Nb

64.5

77.6

143

38.9

50.3

62.3

78.5

37.1

52.7

48.9

88.7

73.3

Ba

1,06

797

197

275

754

975

291

964

9511

635

866

688

Hf

3.2

4.3

7.1

7.0

6.1

5.6

6.5

3.4

5.7

5.2

6.9

6.9

Ta

5.8

6.32

11.2

5.08

4.87

4.40

5.25

3.85

3.76

4.58

7.30

5.49

Th

3.6

5.06

11.6

5.14

2.56

5.20

7.49

3.23

3.87

3.41

5.80

4.11

La

44.7

57.4

131

59.3

28.9

60.1

74.0

46.3

40.2

35.9

58.7

43.1

Ce

95.3

118

264

121

64.0

121

149

96.6

83.6

75.5

118

91.1

Nd

45.5

56.0

111

52.6

34.1

53.2

65.9

43.1

38.0

35.6

57.5

44.0

Sm

8.92

10.5

18.1

9.24

6.90

9.41

11.6

8.01

7.45

7.00

10.5

8.37

Eu

3.07

3.54

5.66

2.97

2.28

3.05

3.62

2.66

2.47

2.36

3.40

2.74

Yb

1.64

1.83

2.14

1.46

1.04

1.28

1.26

1.25

1.39

1.06

1.62

1.13

Lu

0.21

0.24

0.24

0.16

0.11

0.14

0.11

0.15

0.16

0.14

0.19

0.12

aRockTyp

e.Bas.Basanite;Melil.Melilitite;

Olv.Mel.Oliv

ineMelanephelin

ite;Px.

Mel.Pyrox

eneMelanephelin

ite;Meli.Neph.

Melilitic

Nephelin

itebLOIloss

onignitio

ncMg#

¼Mg2

þ�Mg2

þ�Fe2

þ�

��

� �100

dbelow

detectionlim

it

S. Martins et al.

Tab

le2

Representativeclinop

yrox

eneandolivinemicroprob

eanalyses

ofSantiago

lavas

Mineral

Clin

opyrox

ene

Oliv

ine

Formation

Assom

ada

Upp

erP.A.

Flamengo

sAssom

.Upp

erP.A.

Low

erP.A.

Flamengo

s

Sam

ple

ST-21

ST-23

ST-48

ST-34

ST-12

ST-18

ST-48

ST-37

ST-30

ST-34

M.a

Ph.

Coreb

Ph.

Core

Ph.

Rim

cM.

Ph.

Rim

Ph.

Core

Ph.

Core

Ph.

Rim

Ph.

Core

M.

Ph.

Core

Ph.

Rim

SiO

2(w

t.%)

48.19

46.46

46.90

43.84

42.64

50.49

SiO

2(w

t.%)

40.32

40.15

39.03

40.72

38.47

40.06

39.40

TiO

22.43

2.41

2.54

4.17

4.97

1.78

TiO

20.02

0.03

0.06

0.05

0.01

0.03

0.01

Al 2O3

5.31

6.78

7.10

7.69

9.46

3.91

Al 2O3

0.03

0.01

0.00

0.03

0.05

0.00

0.01

Cr 2O3

0.67

0.04

0.11

0.07

0.15

0.41

FeO

12.23

12.29

19.30

11.40

19.70

12.81

14.72

Fe 2O3

2.50

5.72

3.80

3.97

4.44

1.02

MnO

0.14

0.18

0.55

0.17

0.33

0.17

0.22

FeO

3.10

2.15

2.33

2.50

3.05

4.45

MgO

46.14

46.89

39.86

47.15

39.37

46.85

43.94

MnO

0.08

0.10

0.06

0.06

0.06

0.11

NiO

0.33

0.24

0.11

0.24

0.10

0.27

0.21

MgO

13.71

13.00

13.18

11.58

11.12

14.76

CaO

0.20

0.29

0.91

0.21

0.59

0.24

0.37

CaO

24.03

23.15

23.32

24.49

23.92

23.22

Total

99.41

100.08

99.82

99.97

98.62

100.44

98.88

Na 2O

0.31

0.57

0.56

0.33

0.40

0.29

K2O

0.00

0.00

0.00

0.02

0.00

0.00

Total

100.3

100.38

99.90

98.71

100.22

100.48

Si4+

1.79

1.75

1.76

1.68

1.62

1.86

Si4+

1.01

1.00

1.01

1.01

1.01

0.99

1.00

AlIV

0.21

0.25

0.24

0.32

0.38

0.14

Ti4+

0.00

0.00

0.00

0.00

0.00

0.00

0.00

AlV

I0.03

0.05

0.07

0.03

0.04

0.03

Al3+

0.00

0.00

0.00

0.00

0.00

0.00

0.00

Fe3

+0.07

0.16

0.11

0.11

0.13

0.03

Fe2

+0.26

0.26

0.42

0.24

0.43

0.27

0.31

Cr3+

0.02

0.00

0.00

0.00

0.00

0.01

Mn2

+0.00

0.00

0.01

0.00

0.01

0.00

0.00

Ti4+

0.07

0.07

0.07

0.12

0.14

0.05

Mg2

+1.72

1.74

1.54

1.74

1.54

1.73

1.67

Fe2

+0.10

0.07

0.07

0.08

0.10

0.14

Ni2+

0.0

0.00

0.00

0.00

0.00

0.01

0.00

Mn2

+0.00

0.00

0.00

0.00

0.00

0.00

Ca2

+0.01

0.01

0.03

0.01

0.02

0.01

0.01

Mg2

+0.76

0.73

0.74

0.66

0.63

0.81

Ca2

+0.96

0.93

0.93

1.01

0.97

0.92

Na2

+0.02

0.04

0.04

0.02

0.03

0.02

K+

0.00

0.00

0.00

0.00

0.00

0.00

Wo

50.20

48.26

49.43

53.30

52.43

47.82

Fo

87.06

87.19

78.64

88.05

78.08

86.70

84.18

En

39.86

37.71

38.89

35.08

33.92

42.29

Fa

12.94

12.81

21.36

11.95

21.92

13.30

15.82

Fs

8.79

11.88

9.52

10.31

12.04

8.81

Ac

1.15

2.15

2.16

1.31

1.60

1.08

aM.Matrix

bPh.

CorePheno

crystCore

cPh.

Rim.Pheno

crystRim

e;Upp

erP.A.Upp

erPicoda

Antón

ia;Low

erP.A.Low

erPicoda

Antón

ia

Chemical and mineralogical evidence of mantle metasomatism

Santiago lavas are strongly enriched in incompatibleelements owing to their origin from low-degree partialmelting, which also accounts for the highly SiO2-undersat-urated character of the samples (Table 1; Fig. 2a). SampleST-18 (Assomada Formation) stands out as the mostenriched and undersaturated of the total data set, reflectingits melilititic character and its probable origin from a verylow-degree of partial melting (e.g. Brey and Green 1975).This would also explain the observed difference in thedegree of incompatible element enrichment relative to theother plotted sample from the same formation.

REE are significantly fractionated with steep chondrite-normalised patterns [(La/Yb)n=14–42]. Variable Tb/Yb andLa/Yb ratios (0.64 to 0.95 and 21.2 to 61.5, respectively,Fig. 2b) for primitive samples (Ni>150 ppm; Mg#>59)suggest that (even inside the same volcanic complex)important differences in the degree of partial melting andin the amount of residual garnet are necessary in order to

explain the observed chemical variability of magmas (seeGeorge and Rogers 2002).

Primitive mantle-normalised incompatible element pat-terns for primitive Santiago lavas (Fig. 2a) show character-istics that are consistent with those previously reportedfrom other southern Cape Verde islands (Gerlach et al.1988; Doucelance et al. 2003; Escrig et al. 2005). Inparticular, we note: 1) Nb and Ta enrichment relative toLREE and LILE; 2) pronounced negative K anomalies inthe lavas from the Assomada and UPA Formations (downto 0.18 and 0.2, respectively); and 3) (Ba/Rb)n and (Ba/Th)nhigher than 1 (up to 7.44 and 2.64 , respectively).

Sr, Nd and Hf isotope composition

Isotopic data for selected Santiago lavas are reported inTable 4 and displayed in Fig. 3. In addition to the first Hfisotope data for the Cape Verde archipelago, we also report

Table 3 Representative spinel microprobe analyses of Santiago lavas

Mineral Spinel

Formation Assomada Upper P.A. Lower P.A. Flamengos

Sample ST-18 ST-21 ST-23 ST-45 ST-46B ST-48 ST-30 ST-37 ST-42 ST-12 ST-34Inc.Olv.a

Inc. Olv. Inc. Olv. Inc. Olv. Inc. Olv. Inc. Px.b Inc. Olv. Inc. Olv. Inc. Olv. Inc. Olv. Inc. Olv.

Al2O3

(wt.%)11.26 24.40 19.25 30.34 29.96 19.74 22.89 26.25 25.68 15.63 24.23

TiO2 4.61 2.56 3.41 1.64 2.08 3.23 2.37 1.93 1.95 3.31 1.90

Cr2O3 25.91 29.34 32.86 27.47 21.89 25.50 30.85 28.97 30.54 36.72 31.67

V2O3 0.00 0.00 0.00 0.00 0.00 0.20 0.00 0.00 0.08 0.00 0.00

Fe2O3 24.64 11.05 10.54 9.31 13.04 19.49 12.32 12.51 10.45 11.78 9.73

FeO 25.65 20.33 22.38 15.24 20.21 17.70 17.33 15.64 15.82 21.39 19.11

MnO 1.20 0.45 0.56 0.38 0.42 0.32 0.41 1.50 0.49 0.43 0.59

MgO 7.13 11.14 9.54 14.52 11.46 12.68 12.77 13.54 13.72 9.99 11.29

Total 100.40 99.28 98.54 98.92 99.05 98.87 98.93 100.35 98.73 99.25 98.51

Al3+ 3.59 7.17 5.88 8.54 8.62 5.90 6.72 7.48 7.42 4.81 7.16

Ti4+ 0.94 0.48 0.66 0.30 0.38 0.62 0.44 0.35 0.36 0.65 0.36

Cr3+ 5.53 5.79 6.73 5.19 4.22 5.11 6.08 5.54 5.92 7.58 6.28

V3+ 0.00 0.00 0.00 0.00 0.00 0.04 0.00 0.00 0.02 0.00 0.00

Fe3+ 5.01 2.07 2.06 1.67 2.40 3.72 2.31 2.28 1.93 2.31 1.84

Fe2+ 5.79 4.24 4.85 3.05 4.12 3.75 3.61 3.16 3.24 4.67 4.01

Mn2+ 0.28 0.10 0.12 0.08 0.09 0.07 0.09 0.31 0.10 0.09 0.12

Mg2+ 2.87 4.14 3.69 5.17 4.17 4.79 4.74 4.88 5.01 3.89 4.22

Fo (olv) 0.88 0.86 0.84 0.86 0.84 0.87 0.86 0.88 0.86 0.88 0.88

T (Mg-in)°C

1,185 1,176 1,162 1,195 1,142 1,172 1,182 1,204 1,189 1,184 1,212

ameltSiO20.08 0.25 0.27 0.21 0.22 0.15 0.27 0.25 0.23 0.29 0.22

ΔQFM 1.37 0.99 0.83 0.27 0.99 1.28 0.83 1.17 0.65 1.43 0.8

a Incl. Olv. Inclusion in olivineb Incl. Px. Inclusion in pyroxene

S. Martins et al.

new Sr and Nd isotopic results. The isotopic data coverthe following ranges: 87Sr/86Sr=0.70318–0.70391,143Nd/144Nd=0.51261–0.51287, 176Hf/177Hf=0.28284–0.28297. The 87Sr/86Sr and 143Nd/144Nd ratios are consis-tent with those previously reported for the southern CapeVerde islands (e.g. Gerlach et al. 1988; Davies et al. 1989;Doucelance et al. 2003; Barker et al. 2009), overlapping86% of the isotopic range of this group of islands, whichrepresent ≈ 20 Ma of volcanic activity (cf. Holm et al.2008). However, we emphasise that the observed176Hf/177Hf isotopic ratios variations (Δ=0.00013) aresignificantly smaller than those reported from some otherislands (e.g. Kerguelen, Δ=0.00040; Mattielli et al. 2002).The Sr isotope ratios are negatively correlated with143Nd/144Nd and 176Hf/177Hf. Most of the samples plot in

the depleted quadrant of the Sr-Nd and Nd-Hf diagrams(Fig. 3). As already noticed for major elements (see above),the isotopic composition of the Santiago lavas also changedover time. In general, the 143Nd/144Nd ratios decrease whilethe 87Sr/86Sr ratios increase over time, as previouslyreported by Gerlach et al. (1988) (see also Doucelance etal. 2003).

The heterogeneous isotopic signature of the Santiagomagmas is especially obvious within the range observedamong those characterised by lower degrees of partialmelting (i.e., higher Th contents). The initial low meltfractions (higher Th and La concentrations or La/Yb ratios)tend to sample more restricted portions of the mantle, andso are more likely to image several distinct domains. Withthe gradual increase of partial melting (lower Th and Laconcentrations) such diversity tends to be dissipatedtowards a homogeneous composition, better reflecting thepredominant Santiago mantle signature (Fig. 4).

Ultramafic xenoliths

Santiago lavas from the LPA, UPA and FlamengosFormations frequently carry ultramafic xenoliths rangingfrom spinel lherzolites to dunites and wehrlites (Mendes1995; Tables S1 and S2, supplementary material). Dunitesand wehrlites often display igneous textures and comprisevariable amounts of olivine, clinopyroxene and spinelwhose crystal chemistry is typical of ultramafic cumulatescrystallising from alkaline magmas in ocean islandselsewhere (e.g. Munhá et al. 1990). Spinel lherzolitesdisplay allotriomorphic-granular to porphyroclastic texturesin which large grains of olivine + clinopyroxene +orthopyroxene occur within a finer matrix of annealedolivine + clinopyroxene neoblasts. Lherzolite clinopyrox-enes (Cr-diopside, in porphyroclasts and neoblasts) havelower Al2O3 (2.76±0.66 wt.%) and TiO2 (0.11±0.09wt.%), but higher Na2O (up to 3.32 wt.%) and Cr2O3 (upto 4.45 wt.%) content than those from many other spinelperidotites, being similar to clinopyroxenes in carbonatite-metasomatised lithospheric xenoliths reported by Yaxley etal. (1998).

A notable feature of several lherzolite samples is thepresence of vein/vesicular patches of siliceous, aluminousand alkali-rich glass (SiO2=58–64 wt. %; Al2O3=19–23 wt.%; Na2O=4.1–8.4 wt.%; K2O=0.36–9.52 wt.%)containing secondary olivine + clinopyroxene + spinelassemblages. The glassy material partially replaces primaryclinopyroxene and orthopyroxene, being frequently associ-ated with the occurrence of kaersutite, phlogopite andcalcite. In addition to calcite, CO2 is also abundant as(gaseous-liquid CO2 ± spinel ± carbonate) inclusions inolivine and, particularly, in clinopyroxene. Large, anhedral,interstitial spinel grains occur either at the contacts between

Fig. 2 a Incompatible element patterns for representative Santiagolavas normalised relative to primitive mantle values (Palme andO’Neill 2003). b Variation of La/Yb and Tb/Yb ratios illustratingvariable degrees of partial melting (F) and different amounts ofresidual garnet in the mantle source. Samples from the Monte dasVacas formation were not included due to their more differentiatedcharacter (Ni<29 ppm; Mg#<52)

Chemical and mineralogical evidence of mantle metasomatism

Table 4 Radiogenic isotopic data of Santiago lavas

Formationa Sample 87Sr/86Srb 143Nd/144Ndb 176Hf/177Hfb εNdc εHf

c

M.V. ST-16 0.703623±7 0.512680±15 0.282932±10 0.83 5.68

ST-24 0.703767±8 0.512631±11 0.282892±12 −0.12 4.25

DR1 0.282884±8 3.96

DD1 0.703775±6 0.512660±9 0.282895±6 0.43 4.36

DD2 0.703743±7 0.512610±15 0.282876±9 −0.53 3.67

Assomada ST-18 0.703424±9 0.512802±12 0.282929 ± 9 3.21 5.58

DR1 0.512773±11 2.65

DR2 0.512768±11 2.54

ST-21 0.703422±6 0.512773±12 0.282960±7 2.64 6.65

ST-23 0.703822±7 0.512613±10 0.282878±10 −0.47 3.78

DR 0.282886±7 4.02

Upper P.A. ST-27 0.703611±7 0.512696±13 0.282902±9 1.14 4.60

ST-40 0.703527±7 0.512749±10 0.282949±9 2.18 6.28

ST-44 0.703529±8 0.512728±12 0.282933±11 1.77 5.71

ST-48 0.703362±8 0.512792±13 0.282940±13 3.01 5.96

DR 0.512754±19 0.282950±14 2.27 6.29

ST-50 0.703365±8 0.512789±19 0.282956±5 2.96 6.52

DR 0.512789±25 2.96

DD1 0.703370±7 0.512781±13 0.282952±8 2.79 6.35

DD2 0.703356±7 0.512765±26 0.282946±8 2.49 6.14

ST-59 0.703907±8 0.512634±11 0.282841±11 −0.07 2.46

DR1 0.282853±9 2.85

DR2 0.282851±5 2.79

Lower P.A. ST-37 0.703549±7 0.512702±13 0.282956±8 1.26 6.54

ST-42 0.703538±10 0.512721±11 0.282944±10 1.63 6.10

DR 0.703538±9

ST-30 0.703547±7 0.512735±12 0.282926±9 1.91 5.45

DR 0.282954±9 6.43

DD1 0.703541±7 0.512738±11 0.282943±8 1.95 6.03

DD2 0.703542±6 0.512712±15 0.282945±7 1.45 6.12

ST-74 0.703521±8 0.512671±20 0.282897±9 0.66 4.42

DR1 0.703561±8 0.512686±12 0.282939±15 0.94 5.90

DR2 0.282903±6 4.63

Flamengos ST-5 0.703177±8 0.512868±9 0.282975±8 4.50 7.20

DR 0.282963±9 6.77

ST-12 0.703400±9 0.512772±10 0.282975±17 2.62 7.19

DR 0.282965±7 6.82

ST-34 0.703352±13 0.512705±11 0.282937±9 1.32 5.84

DR 0.703534±9

aMV Monte das Vacas; Upper PA Upper Pico da Antónia; Lower PA Lower Pico da Antóniab Reported 2σ applies to the last decimal placec εNd calculated according to 143Nd

�144Nd

� �m

.143Nd

�144Nd

� �CHUR

h i� 1

n o� 104, where CHUR (CHondritic Uniform Reservoir) =

0.512638 (DePaolo and Wasserburg 1976); εHf calculated according to 176Hf�177Hf

� �m

.176Hf

�177Hf

� �CHUR

h i� 1

n o� 104, where CHUR

(CHondritic Uniform Reservoir) = 0.282772 (Blichert-Toft and Albarède 1997).

DD stands for duplicate dissolution and DR duplicate run

S. Martins et al.

embayed clinopyroxene, olivine and orthopyroxene or asfracture fillings, which cut across those minerals. Thespinels extend to unusually high Cr# [Cr/(Cr+Al)], rangingfrom 0.033 to 0.802. Due to these mineralogical andgeochemical characteristics, we interpret these lherzolitesas modally metasomatised.

The spinel-olivine Fe-Mg exchange (Ballhaus et al.1991) temperatures of most dunites and wehrlites arewithin the range of 1,036–1,136°C (except for two dunitesamples at 965°C and 989°C), whereas in lherzolites,olivine + spinel equilibrated at lower temperatures,varying from 858°C to 1,108°C (a similar temperaturerange was obtained from lherzolite two-pyroxene geo-thermometry (Wells 1977; Andersen et al. 1993), withx(T2px-Tsp-ol) ≈ −30°C). Single clinopyroxene geobarom-

etry (Nimis 1999) applied to dunites and wehrlitessuggests an average P=0.9±0.3 GPa; this pressure rangeis identical (within error) to that previously estimated forthe main magma differentiation level(s) below Santiagoisland (P=0.89±0.15 GPa), supporting the hypothesis thatdunites/wehrlites represent former cumulates crystallisedfrom the same type of magmas that are now representedby effusive rocks at the surface. In contrast, lherzolitexenoliths seem to be derived from greater depths (1.6±0.6 GPa—Andersen et al. 1993; P=1.8±0.3 GPa—Nimis1999), probably representing lithospheric mantle close tothe transition between spinel- and garnet lherzolite (seeO’Neill 1981). Oxygen geobarometry (Ballhaus et al.1991) indicates that fO2 conditions within the lithosphericmantle below Santiago Island were oxidised relative to

Fig. 3 Isotopic variability ofSantiago lavas in Sr-Nd-Hfspace. a 176Hf/177Hf—143Nd/144Nd isotope data;mantle array (MA) (see Vervoortet al. 1999) and several islandgroups, considered to be the mostrepresentative of the commonmantle components (HIMU: StªHelena, Tubuai; EM 1: Tristan daCunha, Walvis Ridge, Pitcairn;EM 2: Samoa) are projectedfor reference (data from theGEOROC database: http://georoc.mpch-mainz.gwdg.de).b 143Nd/144Nd–87Sr/86Sr isotopedata; fields of the northern andsouthern Cape Verde Islands(Gerlach et al. 1988; Jørgensenand Holm 2002; Doucelance etal. 2003; Holm et al. 2006) andCape Verde carbonatites(Hoernle et al. 2002) are repre-sented in comparison with thenew isotopic data from SantiagoIsland (2σ deviation is withinsymbol size). The Bulk SilicateEarth (BSE) value is indicatedby a star

Chemical and mineralogical evidence of mantle metasomatism

those of the QFM buffer, with most lherzolite Δlog(fO2)QFM values clustering within a narrow range from+0.4 to +0.8.

Representative whole-rock geochemical data from thestudied ultramafic xenoliths are presented in Table S1(supplementary material). Santiago lherzolites have rela-tively high CaO, P2O5, K2O and Na2O, and low Al2O3 andTiO2 contents, displaying striking similarities to peridotitesreported by Yaxley et al. (1998). As for the metasomatisedperidotite suite from the southeastern Australian lithospher-ic mantle (Yaxley et al. 1998), the unusual geochemistry ofthe Santiago lherzolites reflects their particular mineralogy.The unusually low Al2O3/CaO (0.13 to 1.26 with average0.74, compared with 1.23 for primordial mantle values;Palme and O’Neill 2003) is linked to high modalclinopyroxene, while the high K2O/Al2O3 (average 0.055compared with the typical mantle value of 0.007; Palmeand O’Neill 2003) is related to metasomatic phlogopite.Santiago’s modally metasomatised lherzolites also displaysignificant enrichment in incompatible elements, withP2O5/TiO2, LREE/HREE and Ti/Eu mean values (P/Ti=0.60; La/Yb=14; Ti/Eu=3,800) clearly distinct from thoseestimated for primitive mantle (P/Ti=0.06; La/Yb=1.5; Ti/Eu=7,914) (e.g. Palme and O’Neill 2003).

Discussion

The origin of the EM1-like signature of the Santiago lavas

Although the existence of mantle components, reservoirsand end-members has recently been challenged byseveral authors (Meibom and Anderson 2003; Willboldand Stracke 2006; Armienti and Gasperini 2007), their useto explain mantle variability is still accepted, being a

handy way to discuss the origins of different contributionsto magma mantle sources (e.g. Hofmann 2003; Stracke etal. 2005).

The isotopic signatures of magmas from Santiago, as oneof the southern Cape Verde islands, have been thought to bedominated by end-members with EM1 and HIMU affinities(e.g. Gerlach et al. 1988; Doucelance et al. 2003).

Despite the “multi-component” character of the Santiagomantle sources, the HIMU contribution is evident from: 1)relatively radiogenic lead isotopic signatures (206Pb/204Pbup to 19.43; see Gerlach et al. 1988, Doucelance et al.2003, Barker et al. 2009) and 2) Nb and Ta enrichmentrelative to LILE and LREE, with Rb/Nb, K/Nb and La/Nbratios (0.55±0.25, 166±106 and 0.79±0.18, respectively)significantly below primitive mantle values (Rb/Nb=1.0,K/Nb=433 and La/Nb=1.14; Palme and O’Neill 2003) (seeWeaver 1991; Willbold and Stracke 2006). The predomi-nance of the HIMU-like end-member relative to the EM1 isalso supported by our data showing that the Santiago lavasplot in the depleted quadrant of the Sr-Nd and Nd-Hfspaces. This indicates that they were derived from source(s)with time-integrated depletion in the more incompatibletrace elements. However, such depletion is less pronouncedthan that observed for the northern Cape Verde islands,which are characterised by more unradiogenic Sr andradiogenic Nd signatures (Gerlach et al. 1988; Jørgensenand Holm 2002; Doucelance et al. 2003; Holm et al. 2006).

On the other hand, the high 176Hf/177Hf for a given143Nd/144Nd, the somewhat high 87Sr/86Sr and low143Nd/144Nd and 176Hf/177Hf ratios, as well as the occur-rence of some (unaltered) samples with Ba/Nb>10 (Tables 1and 4) imply a contribution from an enriched end-member,which has been considered akin to the EM1 in previousstudies on the southern Cape Verde islands (see Gerlach etal. 1988; Doucelance et al. 2003; Escrig et al. 2005).

Fig. 4 143Nd/144Nd—Th rela-tionships for Santiago lavas.Lavas characterised by smallerdegrees of partial melting (highTh content) display highly het-erogeneous isotopic signatures(143Nd/144Nd) that tend to ho-mogenise as the melting degreeincreases (represented byarrows). Similar trends areobtained by using La concen-trations or La/Yb ratios asproxies of the degree of partialmelting

S. Martins et al.

The nature of the EM1 mantle component has been thesubject of intense debate and several hypotheses have beensuggested including recycled ancient pelagic and/or metal-liferous sediments, recycled oceanic plateaus, ancientsubcontinental lithospheric mantle (SCLM), ancient delami-nated lower continental crust, interaction with deep meta-somatised recycled oceanic lithosphere and the involvementof pyroxenitic restites (see Weaver 1991; Hofmann 1997;Gasperini et al. 2000; Tatsumi 2000; Niu and O’Hara 2003;Willbold and Stracke 2006). For the southern Cape Verdeislands, ancient recycled pelagic sediments, subcontinentallithospheric mantle (SCLM) and ancient delaminated lowercontinental crust have been discussed as possible origins forthe EM1-like end-member (Gerlach et al. 1988; Doucelanceet al. 2003; Escrig et al. 2005).

Santiago samples plot close to the εNd-εHf mantle array(Vervoort et al. 1999; Chauvel et al. 2008) but with adistinctly shallower slope (Fig. 3a). The observed Santiagoslope (0.76, r2=0.64) is significantly shallower than themantle (1.33) and OIB (1.42) arrays (Vervoort et al. 1999),indicating a higher time-integrated Lu/Hf with respect to agiven Sm/Nd. Similar “shallow” trends in εNd vs. εHfspace have also been reported for Pitcairn lavas (1.0; Eiseleet al. 2002) as well as for specific Hawaiian volcanoes (0.8;Koolau and Haleakala) and are considered to be indicativeof recycled pelagic sediments in the source (see also,Blichert-Toft et al. 1999). In addition to the Hf isotopicevidence, some Santiago primitive lavas also show signif-icant Ba enrichment (up to 1,150 ppm), as well as high La/Nb (0.8±0.2) and Ba/Nb (12.5±3.2), in agreement with theputative influence of sediments in the mantle source(Kahoolawe: La/Nb=1.0±0.1, Ba/Nb=8.6±2.5, Huang etal. 2005; Koolau: La/Nb=1.3±0.1, Ba/Nb=9.2±1.5, Huangand Frey 2005). However, Santiago lavas are clearly lessradiogenic in Sr (87Sr/86Sr≤0.70390) than Pitcairn(87Sr/86Sr≤0.7052; Eisele et al. 2002) and Gough lavas(87Sr/86Sr≤0.7053, Class and le Roex 2008), and they lackNb and Ce negative anomalies (Nb/Nb* up to 2.2; Ce/Ce*≥1.0), which have been identified as reflecting thecontribution of ancient recycled sediments (Pitcairn: Nb/Nb*=0.9–1.2, Eisele et al. 2002; Gough: Ce/Ce*∼0.92,Class and le Roex 2008). Therefore, we conclude that, ifthe presence of sediments as a cause for the EM1 signaturecannot be ruled out completely, its role was probablysubordinate.

On the other hand, Santiago’s primitive lavas arecharacterised by somewhat high (Ba/Th)n (up to 4)contrasting with the signature of the lower continentalcrust [(Ba/Th)n=2.7; Rudnick and Gao 2003] and makingit an unsuitable explanation for the EM1-type signatures.Indeed, considering Santiago compositions and the factthat the HIMU-type end-member is presumably charac-terised by (Ba/Th)n≤1 (Chauvel et al. 1992; Willbold and

Stracke 2006), the Santiago EM1-type end-member wouldbe characterised by much higher (Ba/Th)n values thanthose typical of lower continental crustal compositions.Such a characteristic is matched by the Leucite Hillslamproites ((Ba/Th)n≤5.6), considered to be proxies forSCLM (Mirnejad and Bell 2006). Low mean (Th/Nb)n,(La/Nb)n and (Th/U)n ratios (0.53±0.11, 0.79±0.18 and1.05±0.2 average values, respectively, normalised toprimitive mantle; Palme and O’Neill 2003) of the Santiagolavas support the notion of an insignificant role of thelower continental crust [(Th/Nb)n=1.73, (La/Nb)n=1.48and (Th/U)n=1.57; Rudnick and Gao 2003)] in Santiagomagma mantle source.

In Sr-Nd isotope space (Fig. 5), Santiago samplesdisplay a trend that deviates from those of Pitcairn,Gough, Tristan da Cunha and the (recently described)Godzilla seamounts (Geldmacher et al. 2008), which areconsidered the best representatives of the EM1 end-member (Willbold and Stracke 2006). The deviationtowards lower values of 143Nd/144Nd supports a contribu-tion from an enriched end-member with isotopic affinitiessuch as those displayed by the Leucite Hills lamproites.These lamproites are thought to be the result of preferentialmelting of a metasomatic vein assemblage (phlogopite ±richterite ± clinopyroxene ± apatite ± titanite) within thecontinental lithospheric mantle (Mirnejad and Bell 2006).Interestingly, the majority of Lages silicate rocks (EasternBrazil), which belong to the continental alkaline-carbonatiticcomplex of the Late Cretaceous Paraná-Angola-EtendekaProvince (Comin-Chiaramonti et al. 2002), are also charac-terised by similar relatively unradiogenic Nd isotopesignatures plotting in the same trend between the Santiagosamples and the Leucite Hills lamproites, thus indirectlysupporting the role of SCLM for magma genesis in the CapeVerde islands.

The contribution of SCLM to the Cape Verdemagmas had already been proposed for the silicatelavas from the southern Cape Verde islands as one ofthe possible explanations for the enriched component(Gerlach et al. 1988; Davies et al. 1989; Kokfelt et al.1998; Doucelance et al. 2003) and for the generation ofCape Verde magnesio-carbonatites (Hoernle et al. 2002).However, Escrig et al. (2005) proposed, on the basis of theradiogenic Os isotope signatures of the neighbouring FogoIsland lavas, that the EM1-like signatures testify to thecontribution of lower continental crust to such magmasand extended this explanation to the other southern CapeVerde islands. As stated above, our data favour an SCLMorigin for this component in Santiago, suggesting that thecontribution of continental lithosphere to the Cape Verdemagmatism was variable, including its crustal (e.g. Fogo,Escrig et al. 2005) and mantle portions (e.g. Santiago, thispaper).

Chemical and mineralogical evidence of mantle metasomatism

Previous studies on Cape Verde demonstrated that thearchipelago is characterised by significant inter-islandheterogeneity and that the enriched end-member, akin tothe EM1 mantle component, was sampled only by southernislands’ magmatism (e.g. Gerlach et al. 1988; Doucelanceet al. 2003). Farnetani et al. (2002) showed that hetero-geneities in mantle plumes tend to be preserved duringascent. However, the apparent restriction of SCLM signa-tures to the southern Cape Verde islands is betterunderstood if the component is not present in the plumebut resides in the upper mantle as a passive heterogeneity;this is in agreement with seismic data from Griffin et al.(2004), who described the occurrence of a high velocityzone as a restricted mantle body below the southern CapeVerde islands. This may represent delaminated fragments ofthe African continental lithospheric mantle left behindduring the opening of the Central Atlantic Ocean (see alsoBonadiman et al. 2005).

Such a shallow residence locus for the continentallithosphere end-member(s) in Cape Verde was previouslyproposed on the basis of temporal isotopic variationsdescribed for Fogo (Escrig et al. 2005) and the correlationbetween isotopic signatures and differentiation indices(Hoernle et al. 1991; Millet et al. 2008).

Metasomatism in the oceanic lithosphere

Spinel lherzolitic xenoliths carried by Santiago lavas showevidence of metasomatism. Noticeable mineralogical fea-tures include the occurrence of late kaersutite, phlogopiteand calcite associated with alkali-rich glasses and the

development of secondary olivine + clinopyroxene + spinelassemblages. Replacement of orthopyroxene by newlyformed clinopyroxene + olivine implies the interaction ofa lherzolite/harzburgite host rock and a metasomatic agentwith ameltSiO2

too low to be in equilibrium with orthopyroxeneunder lithospheric P-T-conditions (e.g. Kogarko et al.2001). Furthermore, the high Cr# of spinels may indicatemantle-melt interactions as well, producing the dissolutionof clinopyroxene + orthopyroxene + low Cr/Al spinel andthe formation of Cr-enriched spinels. Kelemen and Dick(1995) found that variations in Cr-numbers relative to Ticontents in spinels constitute reliable indicators to distin-guish metasomatic products (characterised by enrichmentsin both Cr# and Ti) from residual melting products(characterised by decreasing Ti and increasing Cr#).Without an unequivocal Cr# - Ti trend, spinels from modalmetasomatised Santiago lherzolites indicate a simultaneousincrease in both Cr# and Ti. These features, together with 1)the negative correlation between spinel Cr# and Mg#, 2) thepositive correlation between spinel-olivine temperaturesand spinel Cr# and 3) relatively high Δlog(fO2)QFM values,suggest the reaction of the Santiago lithospheric mantlewith an infiltrating, oxidising, metasomatic agent.

It is well known that under the low degrees of meltingrequired to generate the silica-undersaturated magmas suchas those from Santiago, residual phases other than the majorupper-mantle mineral components may be present.

The negative K anomalies displayed by primitivesamples from the Assomada and UPA Formations (seeFig. 2a), coupled with K-Ce relations indicating DK>DCe atodds with the partition coefficients usually reported for

Fig. 5 Sr-Nd isotopic range of Santiago lavas. Note that the trenddefined by Santiago samples contrasts with those of ocean islandscommonly ascribed to represent the presence of the EM1 end-member(Pitcairn, Gough and Tristan da Cunha), but it is similar to thosereported for Gran Canaria, Walvis Ridge, Lages silicate rocks andLeucite Hills (LH) lamproites (data from the GEOROC database:

http://georoc.mpch-mainz.gwdg.de). The Bulk Silicate Earth (BSE)value is indicated by a star. For the definition of the Gran Canaria fieldonly the analyses of the less evolved rocks were used to avoid theeffects of oceanic crust assimilation by AFC processes (see Hoernle etal. 1991)

S. Martins et al.

anhydrous mantle peridotites, strongly suggest that someSantiago magmas equilibrated with mantle mineral para-genesis containing a K-bearing phase. The K-Rb relationsimply that DK>DRb, indicating that amphibole ratherthan phlogopite (Damph

K � 1:0, DamphRb � 0:3; Dphlo

K � 4:0,Dphlo

Rb � 6:0; Francis and Ludden 1995) is the predominantmineral responsible for the K retention. It should be notedthat the K negative anomalies are not ubiquitous in theSantiago magmas, being absent in lavas from both theFlamengos and the LPA formations. This explains the twotrends observed on the Th vs. K diagram (Fig. 6); samplesfrom the Flamengos and LPA formations show a positiveTh-K trend (indicating the incompatibility of K), whereas inthe Assomada and UPA formations K is almost invariantregardless of the degree of melting. Ti shows a similarbehaviour, indicating that mantle partial melting eventsgiving rise to magmas of the Assomada and the UPAFormations were buffered by the presence of mineralscapable of retaining both K and Ti. Equilibration with Ti-pargasite is consistent with the behaviour of K and Ti, andcould also explain the sub-chondritic Nb/Ta ratios (<16;Table 1) characterising those magmas (see Green 1995).Although the predominant K-retaining phase has beenidentified as amphibole, the deduced somewhat compatiblecharacter of Ba (Damph

Ba � 0:5, DphloBa � 2:1; Späth et al.

2001) and higher Ba/Rb values (32–83, for 0.08 < ameltSiO2<

0.24 vs. 11–29, for 0.29 < ameltSiO2< 0.33; Table 1), as well as

the increasing Dsource=meltK towards more undersaturated and

incompatible enriched melilitite and melilitic nephelinitesamples, indicate the influence of phlogopite on theirmagma genesis.

The presence of phlogopite in the asthenosphere is notprecluded by experimental work (e.g. Harlow and Davies2004), but it has been shown (e.g. Wallace and Green 1991)that amphibole is not stable at temperatures above 1,150°C,i.e., at temperatures prevailing in the asthenosphere or inmantle plumes, thus pointing to the equilibration of some ofthe Santiago magmas with amphibole present within thelithospheric mantle (see Class and Goldstein 1997).

The presence of these hydrous minerals in the depletedoceanic lithosphere is considered to be the result of theircrystallisation from infiltrating fluids and/or melts generatedby plume influence (see below). Modal metasomatic eventscreated low solidus and highly enriched lithospheric domains,particularly prone to melting in conditions of continued re-heating (e.g. Mata et al. 1998; Späth et al. 2001). ForSantiago we can thus envisage a model involving thecontamination of plume magmas by melts generated inlow-solidus metasomatised lithosphere domains and leavingthose hydrous minerals as melting residues.

The effects of metasomatism are more evident for thepost-3.3 Ma volcanics of Santiago, which suggests thatlithospheric metasomatism occurred during the earlierstages of mantle melting in the area (Martins et al. 2007).

The carbonatitic nature of metasomatism

The physical properties of fluids and melts led to theconclusion that melts (carbonatites or alkali silicates) aremore likely to be mantle metasomatic agents than CO2-H2Orich-fluids (e.g. Coltorti et al. 2000). According toexperimental work (Gudfinnsson and Presnall 2005), acontinuous gradation from carbonatitic to silicate (kimber-litic/melilitic/melanephelinic/basanitic) melts can be gener-ated from CO2-enriched mantle sources by a progressiveincrease in the degree of partial melting. All these melts arecharacterised by low viscosity, having a strong capability topercolate in mantle rocks, reaching a maximum forcarbonatites (e.g. Hammouda and Laporte 2000). Suchmagmas, which are produced by very low degrees of partialmelting, are highly enriched in incompatible trace elements,leaving a strong chemical imprint on the metasomatisedmantle domains; upon melting, these will transfer suchenrichment characteristics to the magmas.

Significant differences in D values experimentallydetermined between pyroxene and carbonatite or silicatemelts result from distinct melt structures occurring with thechange from O2− to CO3

2− as the principal coordinating orcomplexing anion (Klemme et al. 1995; Blundy and Dalton2000). This enables us to determine the nature of themetasomatic agent. Provided that a continuous series existsfrom carbonatite to silicate melts generated from CO2

enriched sources, a gradation would be expected for themetasomatic effects (Grégoire et al. 2000). In this context,

Fig. 6 Plot of K (ppm) vs. Th (ppm) showing the distinct behaviourof K in primitive Santiago lavas. The trends indicate the compatiblebehaviour of K in the Assomada and UPA formations and incompat-ible behaviour in the LPA and Flamengos formations. Samples fromthe Monte das Vacas formation were not included due to theirdifferentiated character (Ni<29 ppm; Mg#<52)

Chemical and mineralogical evidence of mantle metasomatism

we will consider as carbonatitic melts those with enoughCO2 to have a behaviour mimicking that experimentallydetermined for carbonatitic magmas.

These experimental studies indicate that carbonatitemelts thus produced will have marked depletions in Al,Ga, HREE, Ti and Zr relative to Nb, LREE and alkali/alkali earth elements, with Klemme et al. (1995)suggesting that low Ti/Eu is the most powerful tool todiscriminate between these two types of metasomatism.Rudnick et al. (1993) proposed that high Ca/Sc and lowAl/Ca can be used to evaluate the occurrence ofcarbonatite metasomatism in the upper mantle, whereasIonov et al. (1993) argued that high Sr/Sm can be asignature of carbonate-related metasomatism.

All these geochemical features are characteristic of someof the Santiago magmas (Fig. 7). Indeed, primitive lavascharacterised by more pronounced negative K anomaliesdisplay significant sub-chondritic Ti/Eu ratios (down to

3,300), which correlates negatively with Ca/Sc (Fig. 7) andSr/Sm (not shown). Ti/Eu also correlates negatively withthe degree of melting, as measured by Th contents (Table 1),reflecting the melting of a carbonatitic metasomatisedsource, which (owing to their lower solidus) was preferen-tially sampled during lower degrees of melting.

Recently, White (2007) reported that among the OIBsthere are trace element ratios that are remarkably uniform(e.g. Pb/Ce=0.036, Nb/U=48.8) (see also Hofmann et al.1986). However there are specific islands that showstatistically significant differences. The Cape Verde archi-pelago was identified as one of these, being characterisedby higher Nb/U (56.8 ± 16.8 in Santiago, this paper) andlower Pb/Ce (0.030 ± 0.003 in Santiago, this paper), alsoindicating the action of carbonatitic metasomatism in themantle (White, pers. comm. 2007).

Carbonatite melts are ephemeral (Yaxley and Green1996). Thus, when leaving their stability field (by decom-

Fig. 7 Chemical variability of Santiago primitive lavas (Ni≥150 ppm). a The positive correlation of Ti/Eu vs. ameltSiO2

and; b goodnegative correlation between Ca/Sc and Ti/Eu ratios support thecarbonatitic nature of the metasomatic agent (primitive mantle fromPalme and O’Neill 2003); c Al/Ca (mole ratios)—Th/K variations,suggesting that Th/K ratio is also an indicator of metasomatism owingthe stabilisation of K-bearing residual phases; d Ga/Nb vs. Th/K

diagram showing the decrease of ameltSiO2with increasing Th/K. Variation

trends displayed by the UPA and Assomada formations indicate thatDsource/melt(Al) > Dsource/melt(Ca), Dsource/melt(Ga) > Dsource/melt(Nb) andDsource/melt(Ti) > Dsource/melt(Eu), consistent with enhanced amounts ofmodal clinopyroxene expected from CO3

2− enrichment and ameltSiO2

decrease induced by metasomatic processes

S. Martins et al.

pression and/or by infiltrating depleted mantle domains),carbonatite melts react with lithospheric orthopyroxene,producing sodic Cr-diopsidic clinopyroxene and forsteriticolivine and releasing a CO2-rich fluid, temporarily impos-ing a low- ameltSiO2

environment at the reaction/melting front(Matveev et al. 2001). Accordingly, the relations betweenthe Al/Ca, Ti/Eu, Ga/Nb, La/Yb element ratios and ameltSiO2

(Fig. 7; Table 1) in the studied lavas are consistent with thecarbonatite nature of the metasomatic agent. Santiagoprimitive lavas also display complex covariance amongameltSiO2

, fO2 and key element and isotope ratios (Table 1); fordistinct ranges of ameltSiO2

, Ti/Eu decreases whereas La/Ybincreases with increasing oxidation. According to Ballhaus(1993), carbonatitic melts may inherit large amounts of“excess” oxygen when they segregate from their sources;upon infiltration into the shallower upper mantle thesemelts are capable of causing both the metasomatism andrelated oxidation that apparently accompanied lithosphericenrichment processes at the Santiago Island magmasources.

Santiago peridotite xenoliths give further support tothe interaction of lithospheric mantle with carbonatitemelts. The transformation of refractory harzburgite/lherzolite into secondary clinopyroxene-rich lherzoliteassemblages observed in Santiago metasomatised lherzo-lites is consistent with experimental data (e.g. Dalton andWood 1993) on carbonatite melt-peridotite equilibriumwithin the mantle. From those experiments it is knownthat carbonatite melt stability (in equilibrium with mantlelherzolites) is restricted to high pressures; upon decom-pression, decarbonation reactions (enstatite + dolomite =diopside + forsterite + CO2) between components in thecarbonatite and peridotite wall-rocks induce the consump-tion of primary orthopyroxene with the loss of CO2-richfluids (e.g. Yaxley and Green 1996) that were partiallyretained as carbonic fluid inclusions in the newly formedolivine + clinopyroxene of the Santiago lherzolites. Thesefeatures, coupled with the development of metasomaticcarbonate-bearing + hydrous phases (calcite, kaersutite,phlogopite) and low Al2O3 and TiO2 contents in orthopyr-oxene and clinopyroxene, strongly suggest that themetasomatic agent in the Santiago lherzolites was rich incarbonate and H2O (Yaxley et al. 1998). Moreover,incompatible element enrichment with strong fractionationof elements (usually) not affected by silicate magmadifferentiation provides further support for the involve-ment of carbonatites in the metasomatic process. Indeed,the very low whole-rock Ti/Eu values exhibited by theSantiago (modally metasomatised) lherzolites (Ti/Eu=892to 1,874 compared with 7,914 of primitive mantle; Palmeand O´Neill 2003) are inconsistent with silicate metaso-matism (e.g. McPherson et al. 1996), whereas naturalcarbonatites almost always have Ti/Eu values significantly

below those of primordial mantle (e.g. Hoernle et al.2002). The observed decoupling of REE, P and Tiabundances was indeed predicted by experimental studies(Green and Wallace 1988) as a characteristic of carbo-natite metasomatised lithospheric mantle, and is a featurereported for carbonatite metasomatised peridotites else-where (e.g. Rudnick et al. 1993).

Ascending from their sources, carbonate melts becomeunstable and react at depth with mantle harzburgite,inducing metasomatism and (re-)crystallisation of garnet-clinopyroxenites; early incipient melting of these garnet-pyroxenites produces high-SiO2, alkali-rich melts(Klemme et al. 2002)—often preserved as frozen xenolithglass in Santiago modal metasomatised lherzolites nodules(and elsewhere; e.g. Schiano and Clocchiatti 1994). Thistype of glass can alternatively be considered the partialmelting product of low solidus phases (amphibole, apatite,phlogopite, carbonate) by heating and decompression ofthe xenoliths during the ascending route with the hostmagma (Shaw 1999; Ban et al. 2005). Considering themetasomatic origin of these phases, a link between theoccurrence of such glasses and metasomatism may beinferred.

Low Zr/Hf ratios induced by carbonatite metasomatism

The long-standing statement that Zr/Hf ratios are con-stant in basalts as a consequence of the coherentbehaviour of those two HFSEs has been proven to bewrong (e.g. Dupuy et al. 1992; Pfänder et al. 2007).Dupuy et al. (1992) proposed that the supra-chondriticZr/Hf ratios characterising some basaltic provincesreflect their provenance from carbonatite metasomatisedsources (see also Rudnick et al. 1993). Interestingly,Santiago’s primitive lavas originating from carbonatite-metasomatised sources are characterised by near-chondritic Zr/Hf ratios x ¼ 35:02� 5:79ð Þ, suggesting arelatively low Zr/Hf metasomatic agent.

High-P, high-T experimental studies (van Westrenen etal. 2001) show that Zr and Hf are incompatible ingrossular-poor (<19 mol% Ca) garnets, with DZr ≤ DHf <1, while at higher Ca levels both elements becomecompatible with DZr > DHf > 1; in contrast, DHREE arelargely insensitive to garnet composition, such thatsignificant differences should be expected for DZr/DHf

and (particularly) DZr/DHREE between natural garnets fromCa-poor and Ca-richer environments. Preliminary analysisof phase equilibria under melting conditions in (CO2-bearing) garnet-spinel lherzolite (see Fig. 8) indicates thatameltSiO2

may have significant and contrasting influences ondifferent garnet-producing reactions; specifically, asshown in Fig. 8, decreasing ameltSiO2

will expand the stabilityfield of the grossular end-member relative to pyrope in

Chemical and mineralogical evidence of mantle metasomatism

peridotitic garnet. Carbonatite metasomatism clearly indu-ces Ca enrichment and generates (transient) low ameltSiO2

environments in mantle magma sources; both geochemicaleffects may have converged to induce stabilisation ofgrossular-richer garnets that, remaining as a residualmelting phase, constrain Zr/Hf ratios to relatively lowvalues (see also Pfänder et al. 2007). We conclude that thesame metasomatic processes responsible for the stabilisa-tion of amphibole and phlogopite also have enhanced thegarnet grossular content and the amount of modalclinopyroxene and garnet in the mantle sources ofSantiago’s magmas (see Fig. 8), helping to explain theobserved trace element peculiarities (such as relativelylow Zr/Hf ratios).

Concerning the Zr/Hf, we emphasize that there is asignificant difference between the carbonatitic metasomaticagent x ¼ 35:02� 5:79ð Þ and carbonatites cropping out atSantiago Island (x ¼ 89:88� 39:01, Hoernle et al 2002).This is similar to the differences described betweencalculated near-solidus primary compositions and eruptedcarbonatites (Dasgupta et al. 2009). Dasgupta et al. (2009)considered these differences to reflect distinct sources and/orstemming from severe compositional evolution (e.g. wall-rock reaction; fractionation of accessory phases) suffered bycarbonatite magmas before their final cooling at high crustallevels. As an example, extrusive carbonatites from BravaIsland (Cape Verde) were recently interpreted to result fromimmiscibility processes that also produced conjugate meltsof nephelinitic composition (Mourão et al. 2009).

On the origin of the carbonatitic metasomatic agent

Edge-driven convection has been invoked as a mechanismfor the origin of the Cape Verde islands (e.g. King 2007).However, the existence of a mantle plume at Cape Verdeseems indisputable, both from tomographic studies (e.g.Montelli et al. 2006) and from He isotopic signatures (e.g.Christensen et al. 2001; Doucelance et al. 2003). Theseauthors identified an unradiogenic He component in CapeVerde magmas, which they assigned to lower mantlematerial entrained by the plume, thus showing that thecontribution of the mantle plume was not restricted to aheat source. This plume has a moderate HIMU character,having acquired the EM1-type signature by interaction withSCLM preserved in the oceanic mantle beneath CapeVerde.

Bonadiman et al. (2005) studied spinel-peridotite xen-oliths from Sal Island (Cape Verde) and found evidence forthe occurrence of kimberlite-like metasomatism, which wasinterpreted as caused by the melting of a SCLM domain.Taking into account that a continuous gradation fromcarbonatitic to silicate (kimberlitic/melilitic/melanepheli-nitc/basanitic) melts is generated from CO2-enriched mantlesources (e.g. Gudfinnsson and Presnall 2005), severaldifferent types of mantle metasomatic agents can beenvisaged as the result of different degrees of melting ofsuch source. In accordance with Bonadiman et al. (2005),the variety of metasomatic agents in Cape Verde (carbo-natitic: Jørgensen and Holm 2002; this paper; kimberlitic:

Fig. 8 Calculated phase equilibria illustrating the effect of ameltSiO2

(numbers within circumferences/squares in reaction curves) variationson garnet-forming reactions during mantle peridotite partial melting.Phase relations and solidus of pyrolite + 0.5 wt. % H2O and 0.5–2.5 wt. % CO2 from Green and Wallace (1988) and Yaxley and Green(1996). Olivine, clinopyroxene and orthopyroxene mineral composi-

tions representative of Santiago lherzolites were used in calculationstogether with a grossular-rich (XCa=0.19; XMg=0.68) garnet. Calcu-lations were performed using thermodynamic data and solutionactivity models included in the THERMOCALC software (Powell etal. 1998)

S. Martins et al.

Ryabchikov et al. 1995; Bonadiman et al. 2005) could stemfrom sub-continental lithospheric domains preserved be-neath Cape Verde as a result of the thermal influence of amantle plume. For Santiago, this hypothesis is supported bythe decrease of 176Hf/177Hf with the decrease of Ti/Eu andincrease of Sr/Sm (not shown), requiring a relativelyunradiogenic Hf source for the metasomatic agent. Weemphasise that the outcropping carbonatitic rocks arecharacterised by strongly depleted Sr and Nd isotopicsignatures (εNd=2.93 to 5.77; 87Sr/86Sr=0.703101 to0.70363), which, precluding such an SCLM origin, arealso not compatible with the isotopic characteristics inferredfor the metasomatising agent. We conclude that carbonatiticmetasomatic melts originated from a relatively enrichedsource compared to the depleted time-integrated mantledomain from which the Cape Verde outcropping carbona-tites were ultimately derived.

Because the dihedral angles between carbonate melts andsilicate minerals are very low, carbonatites are characterisedby high infiltration diffusibility (1.8×10−9 m2.s−1; Hammoudaand Laporte 2000), being particularly prone to penetrativeinfiltration/reactions (see also Yaxley et al. 1998). Itsconsumption by reactions and its dissemination through themantle inhibits magma pooling and segregation, whichtogether with low CO2 mantle content could explain thescarcity of carbonatites (Hammouda and Laporte 2000) andprobably also the lack of outcropping carbonatites withisotopic signatures similar to those inferred for the Santiagometasomatic agent.

Concluding remarks

The characteristics of the Santiago lavas and xenoliths arethought to reflect (in part) the occurrence of metasomatismin the oceanic lithospheric mantle in a clear example ofplume-lithosphere interaction. This clearly shows that OIBcomposition should be used cautiously when inferring thegeochemistry of plumes and sublithospheric mantle reser-voirs/components (Millet et al. 2008, 2009).

The inferred hydrous residual paragenesis, specificelement ratios (e.g. subchondritic Al/Ca, Ti/Eu and Ga/Nb, chondritic Zr/Hf; high Ca/Sc and Sr/Sm), decrease inameltSiO2

and increase in fO2 and the presence of carbonate-bearing lherzolitic xenoliths in the Santiago lavas thatdisplay orthopyroxene replacement by olivine and clino-pyroxene demonstrate the influence of a carbonatitic-typemetasomatic agent. However, at odds with what has beenrecently suggested for alkaline magmas elsewhere (Pilet etal. 2008), the variability of Santiago lavas cannot beexclusively explained by metasomatic processes. It mainlyreveals the heterogeneous character of the mantle sources,which are characterised by discernible HIMU, lower mantle

and EM1 contributions. The effects of carbonatitic metaso-matism are more evident for the post-3.3 Ma volcanics ofSantiago, suggesting that oceanic lithospheric metasoma-tism occurred during the initial stages of plume activity inthe area.

The EM1 signature identified in this study supports theinterpretation that a delaminated SCLM domain may existin the region. In the nearby North Atlantic basin, the onlyocean island developing a trend towards an EM1-typecomponent is Gran Canaria (Hoernle et al. 1991). EM1-type signatures were also described for samples dredged atthe Oceanographer fracture zone (Shirey et al. 1987; Dossoet al. 1999) and for the Godzilla Seamount (Geldmacher etal. 2008). However, it should be noted that the Sr-Ndisotope trend developed by the Santiago and Gran Canarialavas is distinct from that defined by the Oceanographerfracture zone and Godzilla Seamount (Fig. 5). Indeed, theSantiago-Gran Canaria trend is characterised by a lower143Nd/144Nd for a given 87Sr/86Sr, being more similar tothat reported for the Walvis Ridge in the southern Atlanticand to the Paraná magmatic province in Brazil (Comin-Chiaramonti et al. 2002). This affinity of the Santiago lavaEM1 signature with those present in the southern Atlantic isendorsed by the 143Nd/144Nd vs. 176Hf/177Hf diagram(Fig. 3a), in which there is a trend with a shallower slopethan the mantle array, directed to the Walvis Ridge field.

The existence of several magmatic provinces with EM1-type signatures in the Atlantic probably results from thecontamination of oceanic mantle by megaliths of delami-nated SCLM left behind during the opening of the oceanbasin. The variability of the EM1 signatures is expected ifwe take into account that they are proxies of the SCLM,which is highly heterogeneous given its long history andprolonged isolation from the underlying asthenosphericconvection (e.g. Pearson and Nowell 2002).

Acknowledgements This work was funded by FCT through theresearch projects PLINT (POCTI/CTA/45802/2002) and POCA/PETROLOG (POCTI/ISLF-5-263). S. Martins benefited from a“Fundação para a Ciência e Tecnologia” grant (SFRH/BD/17453/2004), which is gratefully acknowledged. We thank T. Palácios and O.Chaveiro for keeping the electron microprobe at peak runningconditions. The authors also thank the journal editor Andreas Moellerand Michael Abratis and an anonymous reviewer for helpful andconstructive comments that greatly improved the manuscript.

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Chemical and mineralogical evidence of mantle metasomatism