Metasomatized Mantle Xenoliths as a Record of
the Lithospheric Mantle Evolution of the
Northern Edge of the Ahaggar Swell, In Teria
(Algeria)
M.-A. Kaczmarek1,2 *, J.-L. Bodinier1, D. Bosch1, A. Tommasi1,
J.-M. Dautria1 and S. A. Kechid3
1Geosciences Montpellier, Universite de Montpellier and CNRS, Cc 60, Place E. Bataillon, 34095 Montpellier,
France; 2Department of Applied Geology, Institute for Geoscience Research, Curtin University, GPO Box U1987,
Perth, WA 6845, Australia and 3Universite des Sciences et de la Technologie Houari Boumediene, BP 32 El Alia
16111, Bab Ezzouar, Algeria
*Corresponding author. Present address: University of Lausanne, Institute of Earth Sciences, Geopolis,
Mouline, CH-1015 Lausanne, Switzerland.
E-mail: [email protected]
Received March 6, 2013; Accepted February 17, 2016
ABSTRACT
Mantle-derived xenoliths hosted by melilitite lavas from In Teria (Ahaggar, SE Algeria) include gar-
net and spinel peridotites, pyroxenite and phlogopite megacrysts. The spinel and garnet peridotites
record an early deformation event, which formed porphyroclastic microstructures and olivine crys-
tal preferred orientations, followed by static infiltration of hydrous alkaline melts. This metasomaticstage (stage 1) is characterized by the crystallization of phlogopite in the garnet and spinel perido-
tites, amphibole in the spinel peridotites and clinopyroxene in the garnet peridotite, which record
chemical equilibration with an alkaline silicate melt. These early events were largely overprinted by
carbonatitic metasomatism (stage 2), which is observed only in the spinel peridotites. Spinel
peridotite major and trace element compositions, as well as the compositions of newly formed
minerals, are characteristic of interaction with carbonate melt, associated with strong enrichmentin incompatible trace elements in clinopyroxene. This second stage was followed by crystallization
of pyroxenites (stage 3) in vein conduits, probably segregated from alkaline melts. We propose a
scenario in which the different metasomatic imprints record successive stages of interaction be-
tween lithospheric mantle and sublithospheric melts throughout the Cenozoic. In Sr–Nd isotope
space, the host melilitites and several xenoliths are clustered and plot close to the HIMU mantle
end-member. However, some peridotite xenoliths are shifted towards more radiogenic 87Sr/86Sr
values. In 207Pb/204Pb–206Pb/204Pb and 208Pb/204Pb–06Pb/204Pb space the In Teria samples define arelatively large domain characterized by high 206Pb/204Pb and 208Pb/204Pb, consistent with a contri-
bution of an HIMU component, considered to represent a sublithospheric signature. The highest87Sr/86Sr values are comparable with those ascribed to the EM1 mantle end-member, representing
the signature of the lower continental lithosphere, and are probably inherited from the pre-meta-
somatic lithospheric mantle beneath In Teria. Numerical modelling of porous percolation of melt of
sublithospheric origin through an EM1-like lithospheric mantle protolith reproduces the In Teriaperidotite compositions, using moderately sub-chondritic Sr/Nd values for the peridotite (e.g. In
Teria garnet peridotite) and moderately super-chondritic Sr/Nd values in the melt (approximately
ocean island basalt values). A few spinel peridotites require a component characterized by
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J O U R N A L O F
P E T R O L O G Y
Journal of Petrology, 2016, 1–38
doi: 10.1093/petrology/egw009
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a 143Nd/144Nd signature higher than both the EM1 end-member and the local Ahaggar basalts; the208Pb/204Pb compositions of several samples point to a component with a depleted mid-ocean
ridge basalt (MORB) mantle (DMM) signature. Thus the lithospheric mantle beneath In Teria prob-
ably did not have a uniform EM1 signature before the onset of metasomatism; it included a DMM
peridotite component as well as some peridotites with elevated 143Nd/144Nd values recording long-
term LREE depletion.
Key words: Hoggar Swell; mantle xenoliths; metasomatism; North Africa; Sr–Nd–Pb isotopes;subcontinental lithosphere; Tuareg Shield
INTRODUCTION
North Africa is characterized by the existence of several
volcanic swells (e.g. Ahaggar, Tibesti, Darfur) associ-
ated with regional topographic uplift, which suggests
interactions between lithospheric architecture and man-
tle plume activity. However, the relationships between
volcanism and large-scale mantle circulation beneath
North Africa remain poorly understood. Several models
have proposed a plume origin for the recent volcanism
and topography in North Africa, but they differ on the
proposed location of the plumes. Other models high-
light the importance of pre-existing lithosphere archi-
tecture. Ebinger & Sleep (1998) proposed the feeding of
northern African hotspots (Ahaggar, Tibesti and Darfur)
by asthenospheric material rising from the Afar plume
and channelled along zones of thinned lithosphere.
Another view is the development of splash plumes
(Davies & Bunge, 2006); these are buoyancy instabilities
rising from the mantle Transition Zone. The formation
of splash plumes in North Africa might be due to the ac-
cumulation of subducted oceanic slabs during closure
of the Tethys Ocean in the Mesozoic (Lustrino & Wilson,
2007). Another model proposes that the volcanism is
due to local mantle upwellings linked to reactivation of
older lithospheric structures during the early stages of
collision of the African plate with Eurasia (Liegeois
et al., 2005). Edge-driven convection at craton bounda-
ries was also proposed to explain some of the recent
volcanism in North Africa (King & Anderson, 1995,
1998; Missenard & Cadoux, 2012). The absence of hot-
spot tracks and the persistence of a plume signature be-
neath North Africa has been attributed to plume-head
material emplaced during the Cretaceous beneath the
Central Atlantic province and dragged by the African
plate since the Paleocene (Piromallo & Faccenna, 2004).
This study focuses on the Ahaggar (Hoggar) Massif,
which is the most extensively studied volcanic swell in
North Africa, and more specifically on its northeastern
termination. The In Teria district, near Illizi, is located
at the northeastern edge of the Ahaggar volcanic
swell and at the western edge of the Saharan
Metacraton, to the south of the Sahara Basins (Fig. 1a
and b). The lithospheric mantle beneath In Teria
was probably extensively modified during major
lithospheric–asthenospheric interactions that produced
tholeiitic and alkaline magmatism in central Ahaggar
(Aıt-Hamou et al., 2000) (Fig. 1b). The occurrence of gar-
net peridotite among the suite of In Teria mantle xeno-
liths (Dautria et al., 1992) suggests that relics of older
thicker lithosphere may have been preserved beneath
In Teria (Beccaluva et al., 2007). Paradoxically, the In
Teria district belongs to a domain of anomalously high
heat flow (>100 mW m–2) (Takherist & Lesquer, 1989;
Lesquer et al., 1990), and the exhumed mantle xenoliths
display evidence of extensive interaction with silica-
undersaturated silicate and carbonate melts (Dautria
et al., 1992). Both the high heat flow and the metasoma-
tism might be related to marginal effects of Ahaggar
plume activity (Davies & Bunge, 2006), or to more re-
cent mantle flow beneath North Africa (e.g. Ebinger &
Sleep, 1998; Piromallo & Faccenna, 2004; Forte et al.,
2010).
To understand the above observations we have con-
ducted a detailed study of the In Teria peridotite and
pyroxenite xenoliths, phlogopite megacrysts, and their
host melilitite. Microstructural analysis based on the
measurement of the crystallographic preferred orienta-
tions (CPO) of constituent minerals allows constraints
to be made on the deformation processes affecting the
lithospheric mantle beneath In Teria and on the relative
timing of deformation and metasomatism, as well as
determination of possible topotaxial relationships dur-
ing melt–rock reaction. New isotopic data (Sr–Nd–Pb)
for the mantle xenoliths and host lavas, together with
major and trace element analyses of whole-rocks and
constituent minerals, allow definition of the compos-
ition of the Ahaggar mantle lithosphere and description
of its evolution in response to different metasomatic
events. The Sr–Nd–Pb isotopic data indicate that the
percolating melt has a mantle affinity with an EM1 sig-
nature, which has not yet been observed from the
Ahaggar, nor in Morocco or Libya (Beccaluva et al.,
2007, 2008; Raffone et al., 2009; Wittig et al., 2010;
Natali et al., 2013). Together these data allow discussion
of the nature and evolution of the lithospheric mantle at
the northern edge of the Ahaggar Swell and the rela-
tionship of this evolution to the different geodynamic
events that affected the region.
GEOLOGICAL SETTING
The main structures of the Ahaggar Massif are inherited
from the Pan-African orogeny, which resulted from
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continental collision between the West African Craton
and an East African block between 580 and 760 Ma (e.g.
Caby et al., 1981; Lesquer et al., 1984; Bertrand et al.,
1986; Liegeois et al., 1994). Most of the uplift occurred
from Late Eocene to Quaternary times (Guiraud &
Bellion, 1995; Guiraud et al., 2005), but fission-track age
data suggest that the regional uplift of the Ahaggar
basement started at least as early as the Cretaceous
(Khaldi et al., 2006). Volcanic activity in the Ahaggar
Massif began during the late Eocene (c. 34 Ma; Remy,
1959; Rossi et al., 1979; Aıt-Hamou et al., 2000), peaked
during the Miocene and continued episodically through
the late Pliocene and into the late Quaternary (Girod,
1971).The In Teria district is situated within an anomal-
ously ‘hot’ east–west-trending zone characterized by
heat flow higher than 100 mW m–2, which has been
attributed to anomalously high temperatures in the
underlying asthenosphere (Fig. 1a; Takherist & Lesquer,
1989; Lesquer et al., 1990). Significantly lower heat flow
is recorded to the south of the study area in the
Ahaggar Massif (53 mW m–2 on average), suggesting a
stable lithosphere �100 km in thickness (Lesquer et al.,
1989). However, free-air gravity data led Crough (1981)
to suggest that the Central Ahaggar is underlain by
anomalously low-density upper mantle. Similar conclu-
sions have been reached based on seismic tomography
models (Ayadi et al., 2000; Sebai et al., 2006; Zhao,
2007; Begg et al., 2009).
The In Teria volcanic district includes about 20 dia-
treme craters, each no larger than 2 km in diameter
(Megartsi, 1972). These cross-cut horizontal
Carboniferous strata and correspond to pipe-like struc-
tures. The ejecta are cemented by calcite and include
lava blocks up to 1 m in size, as well as lapilli, olivine
and phlogopite megacrysts, and various mantle and
crustal xenoliths. Ultramafic xenoliths were collected by
J.-M. Dautria from the ejecta within a 50 km2 area. The
studied samples come from two diatremes, No. 14 and
No. 17 (Fig. 1c); No. 14 is the larger structure (2 km in
diameter), formed of hydromagmatic tuff containing
pyroxenite and mafic granulite xenoliths, phlogopite
megacrysts, phlogopitite xenoliths, spinel peridotite
xenoliths and the only garnet peridotite xenolith
sampled in the In Teria district. The second diatreme
(No. 17) is 500 m in diameter with a 5 m high ejecta rim
characterized by block and ash deposits. The ejecta in-
clude granite, pyroxenite and peridotite xenoliths,
phlogopite megacrysts and pyroclastic breccia. All
xenolith samples analyzed measured between 5 and
12 cm. Pyroxenite xenoliths represent the majority of
the xenolith population. Most pyroxenite xenoliths have
a black coating and may contain centimetre-sized
amphibole crystals. The xenoliths are embedded in a
porphyritic melilitite host lava (Megartsi, 1972), which
formed at 1000–1200�C at a pressure greater than 20
kbar (Girod, 1971; Kechid & Megartsi, 2005). Although
the age of the volcanic activity is not well established,
Mediterranean Sea
Atlan
tic O
cean
West African Craton
Ahaggar
Illizi
Sahara Basins
80
100
80
(a)
N
Algiers
(b)
300km
Saharan Meta-craton
Tamanrasset
Atakor(alk)
Illizi
AdrarN'Ajjer (alk)
Manzaz (alk)
Tahalra (alk)
Taharaq (thol)
Eggere (alk)
26º48'N
Nº14
Nº15
Nº13
Nº1726º50'N
E'04º9E'83º9
(c)
InTeria
8º30'E
24º30'N
22º30'N
26º30'N
6º30'E4º30'E
(c)
Fig. 1. (a) Map of NW Africa and the Ahaggar Shield showingthe major tectonic units and north–south-trending mega-faults.Dashed lines represent heat flow isolines of 80 and 100 mWm–2 values to the north of the Ahaggar Shield (after Takherist& Lesquer, 1989). The black rectangle corresponds to the loca-tion of (b). (b) Map showing the major faults of the AhaggarSwell and the major volcanic fields, modified after Dautria &Lesquer (1989). Dark grey is the Pan-African basement, lightgrey represents the Paleozoic sedimentary cover, black areasare volcanic districts of Late Mesozoic to Miocene–Quaternaryage, and white represents Quaternary sedimentary cover. (c)The In Teria study area (near Illizi). Dotted lines indicate num-bered diatreme craters (image from Google EarthVR ).
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the freshness of some tuff rings suggests a Quaternary
age (Conrad, 1969; Megartsi, 1972).
ANALYTICAL METHODS
Crystal preferred orientationsOlivine, pyroxene, and amphibole CPO were measured
in 19 samples by indexing of electron back-scattered
diffraction (EBSD) patterns using a JEOL JSM 5600
scanning electron microscope at Geosciences
Montpellier. Data collection, indexing and analysis of
electron backscatter diffraction patterns (EBSP) were
performed with the CHANNEL 5.10 software by Oxford
Instruments. For each sample, we obtained crystallo-
graphic orientation maps covering most of the thin sec-
tion (maps are usually 20 mm� 35 mm) with a regular
grid step ranging from 35 to 55mm depending on the
mean grain size. Average raw indexation rates were
�70%. The EBSD data were noise-reduced using a
‘wildspike’ correction to remove isolated erroneous
points, followed by a six-neighbour zero solution ex-
trapolation following standard procedures (Prior et al.,
2002; Bestmann & Prior, 2003). At each of these steps,
the resulting orientation maps were compared with
band contrast maps to ensure that the data treatment
did not compromise the data. The crystallographic
orientation data are presented as lower hemisphere
pole figures prepared using the Pf_ctf software of D.
Mainprice (http://www.gm.univ-montp2.fr/PERSO/main
price/W_data/CareWare_Unicef_Programs/). Average
Euler angles for each grain (one point per grain) were
used to avoid over-representation of large grains in thin
section. (The complete dataset is available in
Supplementary Data as Fig. A1; the supplementary data
are available for downloading at http://www.petrology.
oxfordjournals.org.) In those samples in which the foli-
ation could not be identified, the CPO data are pre-
sented with the maximum concentration of [100] axes
of olivine parallel to x (east–west) and the maximum
concentration of [010] axes of olivine parallel to z
(north–south) for easy comparison between the
samples.
Whole-rock and mineral compositionsWhole-rock major element compositions were analyzed
at Nancy SARM by inductively coupled plasma optical
emission spectroscopy (ICP-OES). Whole-rock trace
element compositions were determined by inductively
coupled plasma mass spectrometry (ICP-MS) using a
Quadrupole VG-PQ2 system at Montpellier University
(France) following the procedure described by Ionov
et al. (1992). Major elements in minerals were analyzed
by electron microprobe using the CAMECA-SX100 at
the Microsonde Sud facility (Universite de Montpellier
II) equipped with five wavelength-dispersive spectrom-
eters (Supplementary Data Table A2). Operating condi-
tions comprised an acceleration voltage of 20 kV and a
10 nA beam current. K and Na were counted for 20 s
with a 10 s background and the other elements were
counted for 30 s with a 15 s background. Mineral trace
element data were obtained by laser ablation (LA)-ICP-
MS using the facility available at the AETE platform
(Supplementary Data Table A3). The GeoLas Qþ laser
system used at Geosciences Montpellier is an Excimer
(Compex 102) operating in the deep UV (193 nm).
Ablations were performed in a pure He atmosphere
(�0�6 l min–1) using a beam diameter ranging between
30 and 120mm, with an energy density of c. 15� 10–3 J
cm–2. The LA platform is linked to an extended range
Element 2 ICP-MS system operated in low-resolution
mode at 1350 W. The ICP-MS system was daily tuned to
maximum sensitivity while keeping oxide production to
its minimum level (ThO/Th� 1%). The NIST612 glass
was used as external standard (Pearce et al., 1997) and
SiO2 or CaO contents determined by electron probe for
each mineral were used as an internal standard. Data
were processed using the GLITTER software package
(Van Achterbergh et al., 2001).
Sr–Nd–Pb isotopesSr, Nd and Pb isotopic compositions were obtained
from 400 mg of powdered whole-rock samples. After a
step of leaching with 6N HCl at 80�C for 30 min and
three cycles of rinsing with purified milli-Q H2O, sam-
ples were then dissolved for 48 h on a hot plate in a mix-
ture of HF 48% (1:2) and 13 N HNO3 (1:2). After
evaporation to dryness (120�C), 2�5 ml of 13 N HNO3
was added to the residue and kept at about 100�C for
24–48 h and evaporated. A last cycle of dissolution and
evaporation was performed with 2 ml of 13 N HNO3
(1:2). For Pb separation, after complete evaporation,
�100ml of 8 N HBr was added to the sample and kept at
90�C for 5 h before another complete evaporation. The
chemical separation of Pb was carried out using 50 ml of
anion exchange resin (AG1X8, 200–400 mesh) using
0�5 N HBr and 6 N HCl as eluants. Strontium isotopes
were separated using Sr Eichrom resin (Pin et al., 1994).
For nine samples we conducted a stronger leaching to
compare Sr isotope compositions. The first step of
leaching was with 2�5 N HCl at 95�C for 1 h 30 min fol-
lowed by three cycles of rinsing with purified Milli-Q
H2O; samples were then leached with 6 N HCl at 95�C
for 1 h 30 min followed by three cycles of rinsing with
purified Milli-Q H2O. After leaching an identical dissol-
ution procedure was applied. Neodymium isotopes
were separated during three steps: first a rare earth
element (REE) separation using AG50WX12 cation ex-
change resin, followed by two steps of Nd purification
using HDEHP columns. Total blank contents for Pb, Sr
and Nd were less than 30, 60 and 30 pg, respectively.
Lead and neodymium isotopic compositions were
measured on a VG Plasma 54 and a Nu 500 multicollec-
tor (MC)-ICP-MS system at the Ecole Normale
Superieure de Lyon. Pb isotopic compositions were
measured with an external precision of c. 100–150 ppm
for 206,207,208Pb/204Pb using the Tl normalization method
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described by White et al. (2000). Every two samples
were bracketed between the NIST 981 standard. For Nd
isotopic measurements, every two samples were brack-
eted between the Lyon ‘in-house’ Nd standard (a
500 ppb dilution of the JMC commercial solution of ICP-
MS Nd standard, batch 801149A; Luais et al., 1997),
with an average 143Nd/144Nd¼ 0�512133 6 22 (2r)
(n¼ 20). Strontium isotopic compositions were meas-
ured on a Finnigan Triton TI mass spectrometer at the
‘Laboratoire de Geochimie GIS’ of Nımes. Results on
NBS 987 Sr standards yielded a mean value of87Sr/86Sr¼ 0�710254 6 11 (2r) (n¼ 18). Results on NBS
981 Pb standards yielded a mean value of208Pb/204Pb¼36�682 6 1 (2r), 207Pb/204Pb¼ 15�4874 6 4
(2r), 206Pb/204Pb¼ 16�9331 6 4 (2r) (n¼ 14). Given the
inferred Quaternary age of the In Teria magmatic dis-
trict (Conrad, 1969; Megartsi, 1972), no age correction
was performed on the measured isotopic ratios.
PETROGRAPHY AND TEXTURE OF THE
XENOLITHS
Garnet peridotiteSample Int-14-7 is a phlogopite-bearing garnet lherzo-
lite with a weakly foliated porphyroclastic texture
(Table 1, Fig. 2a). Although the petrography of this
sample was previously described by Dautria et al.
(1992), we summarize here the key characteristics of
this rock. It has a centimeter-scale compositional layer-
ing marked by alternate olivine and pyroxene enrich-
ment. Garnet has a spheroidal shape with a diameter
between 4 and 8 mm. Brown Al-spinel locally forms cor-
onas around garnet (Fig. 2a). Primary clinopyroxene
(Cpx-I) is located around garnet or is associated with
orthopyroxene; it has a grain size up to 1 mm.
Orthopyroxene porphyroclasts (Opx-I) have grain sizes
smaller than 1 mm. Olivine (Ol-I) has rounded or elon-
gated shapes (aspect ratios 4:1) with a grain size up to
1 mm. Elongated olivine grains show low-angle boun-
daries or undulose extinction perpendicular to the grain
elongation, which is in general parallel to the layering.
The primary assemblage composed of olivine, ortho-
pyroxene, clinopyroxene and garnet is overprinted by
metasomatic phlogopite (Phl-II). The occurrence of Phl-
II is irregular; it occurs either around garnet or as an
interstitial phase with a poikilitic texture (up to 2�5 mm
in size) (Fig. 2b).
Spinel peridotitesLherzolite, harzburgite and Ol-websteriteIn Teria spinel peridotites are characterized by a por-
phyroclastic texture formed by primary minerals that
Table 1: Modal composition of studied samples from In Teria
Primary phases (I) Secondary Tertiary phases (III) Melt-P
phases (II) (%)
Grt Ol Opx Cpx Sp Amph Phl Ol Cpx Sp–Chr F-K
Garnet peridotiteInt-14-7 6 49 29 10 1�5 — 4 — — 0�5 — —Spinel peridotiteLherzolite and harzburgiteErem-2 — 62 9 5 1 8 tr 6 8 1 — 15Il-17-4 — 76 6 — 1 2 tr 4�5 10 0�5 — 15Il-17-5 — 51�5 — — 0�5 8 5 15 19�5 0�5 — 35Il-17-10 — 57 7 — 1 6 4 9 15 1 — 25Il-17-13 — 68 6�5 — 0�5 5 2 7 10 1 — 18Il-17-14 — 58�5 5 — 0�5 2 9 10�5 14 0�5 — 25Il-17-18 — 67 14�5 — 0�5 3 tr 6 7 2 — 15Int-15-1 — 56 19 — 1 6 tr 6 8 2 — 16Int-17-1 — 58�5 — — 0�5 1 tr 8 30 1 1 40Int-17-6 — 79�5 — — 0�5 6 1 5 7 0�5 0�5 13Int-17-7 — 69 10 — 1 15 1 1 2 0�5 0�5 4Int-17-12 — 64�5 2 — 0�5 13 tr 5 14 0�5 0�5 20Int-17-13 — 73 2 — 1 14 tr 2 7 0�5 0�5 10Ol-websteriteInt-14-6 — 34�5 50 — 0�5 tr tr 4 7 1 3 15WehrliteIl-17-21 — 20 — — 5 1 24�5 49 0�5 — 74Int-17-14 — 7 — — 4 tr 38 50 1 — 89Ol-clinopyroxeniteIl-17-16 — — — 6 16 tr 29 48 1 — 78Pyroxenite Cpx Amph Phl SpIl-17-22 59 24 15 2Il-17-24 51 25 19 5Il-17-27 43 34 22 1Int-17-100 62 19 15 4
Modal proportions were estimated visually. Grt, garnet; Ol, olivine; Opx, orthopyroxene; Cpx, clinopyroxene; Amph, amphibole;Phl, phlogopite; Sp–Chr, spinel and chromite; F-K, alkali-feldspar; Melt-P, proportion of melt pockets; tr, trace.
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are overprinted by two metasomatic assemblages. The
primary assemblage (I) comprises olivine, orthopyrox-
ene and spinel 6 clinopyroxene. Erem-2 is the only
spinel-bearing sample containing primary clinopyrox-
ene (Cpx-I) (Table 1, Fig. 2c). Cpx-I grains from
sample Erem-2 are anhedral, up to 0�5 cm wide, and
contain thin exsolution lamellae of orthopyroxene.
The orthopyroxenes (Opx-I) in all peridotites are
2–5 mm in size and have irregular shapes (Fig. 2d and
e). Olivine (Ol-I) occurs as elongated crystals (up to
8 mm long) that define the foliation (Fig. 2f). Olivine
grains display undulose extinction and subgrains ori-
ented perpendicular to the grain elongation (Fig. 2f).
Spinel, commonly replaced by chromite, is up to
Fig. 2. Photomicrographs in plane-polarized light (PPL) and cross-polarized light (XPL) and backscattered electron (BSE) imagesillustrating the typical microtextures of the In Teria peridotites. (a) Garnet peridotite with a layer rich in orthopyroxene; PPL. (b)Garnet with a corona essentially composed of spinel and clinopyroxene; PPL. Phlogopite (Phl-II) is poikilitic. (c) Xenolith Erem-2with primary clinopyroxene (Cpx-I), orthopyroxene (Opx-I) and olivine (Ol-I), showing reaction zones with former melt pockets; PPL.(d) Primary orthopyroxene (Opx-I) in IL-17-8 is partially replaced by amphibole-II and phlogopite-II; PPL. (e) Opx-I in the IL-17-4 is‘dismembered’ and partially replaced by crystallized melt pockets; PPL. (f) Spinel peridotite IL-17-20 characterized by tabular olivinecrystals and subgrain boundaries; XPL. Dashed line indicates foliation. (g) Coarse granular spinel peridotite IL-17-10 displays foli-ation marked by spinel and amphiboleþphlogopite. The green areas correspond to former melt pockets with clinopyroxene (Cpx-III) enrichment; PPL. Dashed line indicates foliation. (h) BSE image of IL-17-10 showing a melt pocket composed of clinopyroxene(Cpx-III), olivine (Ol-III), interstitial glass and holes (black areas) generated during polishing. (i) Melt pocket with partially orientedCpx-III in IL-17-11; PPL. It should be noted that amphibole is replaced by Cpx-III. (j) BSE image of an amphibole–clinopyroxene sym-plectite in IL-17-16. grt, garnet; cpx, clinopyroxene, opx, orthopyroxene, ol, olivine, amph, amphibole, gl, glass.
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400mm in size and has a skeletal shape, but with align-
ment parallel to the foliation when the latter is observed
(Fig. 2g). The Ol-websterite, Int-14-6, has a texture and
grain size similar to those of the lherzolites and harz-
burgites, but is characterized by a higher proportion of
orthopyroxene (�50%, Table 1).The second mineral assemblage (II), previously
referred to as ‘metasomatic hydrous mineral phases’ by
Dautria et al. (1992), corresponds to the occurrence of
phlogopite (Phl-II) and amphibole (Amph-II) in variable
proportions (Table 1, Fig. 2b–d). Amphibole and phlogo-
pite often replace Opx-I or spinel (Fig. 2d and g). The
spinel–phlogopite–amphibole aggregates are aligned
with the foliation, but the elongation of single amphi-
bole and phlogopite grains (up to 1 mm) is not always
parallel to the foliation (Fig. 2g) and these minerals do
not display intracrystalline deformation features.
The third assemblage (III) consists of reaction aggre-
gates or patches (possible former melt-pockets), which
are present in variable proportions in the spinel perido-
tites (10–40% of the sample volume; Table 1). These re-
action aggregates are composed of microgranular
clinopyroxene, olivine, chromite, Al-spinel, sulphides,
K-feldspar and glass (Fig. 2c and e–i) and were inter-
preted by Dautria et al. (1992) as the products of carbo-
natitic melt percolation through the peridotite. Dark
green clinopyroxene (Cpx-III) and associated minerals
partially replace primary orthopyroxene (Fig. 2e) and
secondary amphibole (Fig. 2i); Cpx-III is euhedral and
has a grain size between 50 and 200 mm (Fig. 2h and i); it
may show elongated shapes parallel to orthopyroxene
or amphibole cleavages (Fig. 2i). Ol-III crystallizes as
neoblasts with subhedral shapes and displays a grain
size between 50 and 150mm (Fig. 2h).
Wehrlite and Ol-clinopyroxeniteWehrlite and Ol-clinopyroxenite form the Cpx-III-rich
peridotite xenoliths that are common at In Teria (5–10%
of collected xenoliths). Three samples were studied:
Il-17-21, Int-17-14 and Il-17-16. Wehrlites Int-17-14 and
Il-17-21 contain primary olivine (7 and 20%, respect-
ively; Table 1) with a grain size up to 1 mm and anhedral
shapes. The clinopyroxene is similar to Cpx-III observed
in the spinel peridotites. It has euhedral shapes with a
grain size between 50 and 200mm. Ol-III grain size varies
between 50 and 150mm. Amphibole and phlogopite are
occasionally found and have anhedral shapes with
grain sizes of up to 0�3 mm.The Ol-clinopyroxenite Il-17-16 is a peculiar sample,
characterized by the lack of primary olivine (Ol-I). This
sample contains clinopyroxene megacrysts with nu-
merous thin orthopyroxene exsolution lamellae and
clinopyroxene–amphibole symplectites (Fig. 2j). The
clinopyroxene megacrysts are embedded in a fine-
grained reaction matrix, similar to that which composes
Fig. 2. Continued.
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the reaction patches in the spinel peridotites, character-
ized by a large proportion of Cpx-III and Ol-III (Table 1).
PyroxenitesPyroxenite xenoliths are more abundant (�70%) than
peridotite xenoliths at In Teria (�20%; Kechid &
Megartsi, 2005). They are characterized by a cumulate
texture, the absence of olivine and garnet, and no reac-
tion aggregates or patches like those observed in the
spinel peridotite xenoliths. In some samples subtle
layering is defined by variations in grain size. They are
mainly clinopyroxenite, composed of clinopyroxene
and phlogopite, with or without amphibole; websterites
are rare. Clinopyroxenes are zoned xenomorphic crys-
tals, measuring between 3 and 4 mm in size. Phlogopite
has a grain size of 1–3 mm. It is generally interstitial and
associated with clinopyroxene. When present, the
amphibole is millimeter-sized and is interstitial between
clinopyroxene and phlogopite.
CRYSTAL PREFERRED ORIENTATIONS
Crystal preferred orientations (CPO) of olivine, ortho-
pyroxene, clinopyroxene and amphibole were analyzed
in the garnet peridotite and 18 spinel peridotite xeno-
liths. Results for the garnet peridotite and six spinel
peridotites are shown in Fig. 3. The complete dataset is
available as Supplementary Material Fig. A1.
The garnet peridotite (Int-14-7) has an olivine CPO
characterized by a strong maximum orientation of
[010] axes close to the pole of the foliation (z) (max-
imum density MD¼ 7�86, Fig. 3a). The [001] and [100]
axes form a girdle roughly parallel to the foliation
plane (x). Both axes have weak maxima subparallel to
each other, but the [001] maximum is stronger than
the [100] maximum. This distribution characterizes a
fiber-[010] olivine CPO. Orthopyroxene and clinopyr-
oxene have similar CPO with a strong alignment of
[001] axes in the foliation plane and subparallel to the
olivine [100] maxima. The [100] and [010] axes of both
pyroxenes form incomplete girdles in the foliation
plane. The [010] axes form a maximum at high angle
to the foliation.
Most spinel peridotites show a weak preferred orien-
tation of olivine grains with a point concentration of
[100] axes parallel to the foliation plane, probably mark-
ing the lineation. The olivine [010] axes form either a
maximum perpendicular to the foliation plane or a gir-
dle normal to the [100] maximum (Fig. 3b–d and
Supplementary Material Fig. A1). In all spinel perido-
tites olivine [001] axes are more dispersed than the
other two, but they form a weak maximum in the foli-
ation generally perpendicular to the [100] maximum.
In spinel peridotites, orthopyroxene has a weak CPO,
which is characterized nevertheless by a weak [001]
maximum subparallel to the olivine [100] maximum
(e.g. Il-17-17 and Int-17-10; Fig. 3b). The multiple max-
ima present in the orthopyroxene CPO are probably
due to over-representation of some former orthopyrox-
ene porphyroclasts, which were corroded and dislo-
cated during subsequent metasomatism, now being
present as separate grains. This effect was corrected for
as much as possible during data reduction.Clinopyroxene and especially amphibole CPO in spi-
nel peridotites are strongly dispersed, but distributions
of the [100], [010] and [001] axes are consistently paral-
lel between these minerals. They display two types of
microstructural relationships with olivine and orthopyr-
oxene from the primary mineral assemblage: (1) Type
A, where despite the much more dispersed CPO, the
clinopyroxene and amphibole [001] axes maxima are
parallel to the [001] maximum of orthopyroxene and to
the [100] axes maximum of olivine (Fig. 3b); (2) Type B,
where clinopyroxene and amphibole display similar
CPO orientations, which are discordant to orthopyrox-
ene CPO (Fig. 3d). Six samples show an intermediate or-
ganization between Type A and B (Fig. 3c and
Supplementary Data Fig. A1). The high dispersion of
the CPO, as well as the Type B and intermediate textural
relationships are consistent with petrographic observa-
tions, which imply that all clinopyroxene and amphibole
in the spinel peridotites are products of metasomatism.
WHOLE-ROCK CHEMISTRY
The previous study of Dautria et al. (1992) on In Teria
xenoliths reported major and trace element data for
peridotite and pyroxenite xenoliths, as well as host
melilitite samples. Major elements were analyzed by
atomic absorption and trace elements (Hf, Ta, Th, U,
and REE) were determined by neutron activation (sam-
ples marked with asterisks in Tables 2 and 3). In this
study we report major element analyses for 10 new
samples by ICP-OES and new trace element analyses
for all samples by solution ICP-MS. Overall, 17 perido-
tite and 11 pyroxenite xenoliths, one phlogopite mega-
cryst and five host melilitites were analysed in this
study for major and trace elements. Analytical results
are given in Tables 2 and 3 and illustrated in Figs 4
and 5.
Garnet peridotiteThe garnet peridotite has low MgO (35�7 wt %) and high
Al2O3 and Fe2O3 contents (3�3 and 11�4 wt %, respect-
ively; Fig. 4, Table 2). Such a composition overlaps the
field of off-cratonic mantle xenoliths (Fig. 4a; Canil,
2004). However, comparison with the mantle melting
array shows that this garnet peridotite does not repre-
sent a melting residue (Fig. 4b and c). This sample has a
relatively flat REE pattern with (La/Lu)N¼2�81 [the sub-
script N indicates Primitive Mantle normalized values
from McDonough & Sun (1995)], a positive Ti anomaly,
and Rb–Ba enrichment relative to the REE (Fig. 5a and
b). Except for Ti, the other high field strength elements
(HFSE: Nb, Ta, Zr and Hf) do not form clear anomalies
relative to neighbouring elements.
8 Journal of Petrology, 2016, Vol. 0, No. 0
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nloaded from
(a)
Ga
rne
t p
erid
otite
MD
= 2
.91
PfJ
= 1
.60
MD
= 7
.86
PfJ
= 3
.37
MD
= 4
.24
PfJ
= 1
.66
N =
115
8
MD
= 4
.75
PfJ
= 1
.44
MD
= 5
.39
PfJ
= 1
.77
MD
= 6
.62
PfJ
= 2
.20
MD
= 3
.15
PfJ
= 1
.34
MD
= 4
.37
PfJ
= 1
.68
MD
= 5
.83
PfJ
= 2
.35
N =
965
N =
408
ZZ
ZZ
ZZ
ZZ
Z
[100
] [0
01]
[010
]
(b)
Sp
ine
l p
erid
otite
(Typ
e A
)
Clin
opyr
oxen
eelobihp
mA
enexorypohtrO
Oliv
ine
[100
] [0
01]
[010
] [1
00]
[001
] ]100[
]001[ ]010[
[010
]
Il-17
-17
N =
405
N =
192
N =
110
5N
= 3
00
MD
= 2
.86
PfJ
= 1
.40
MD
= 3
.36
PfJ
= 1
.59
MD
= 2
.17
PfJ
= 1
.12
MD
= 3
.09
PfJ
= 1
.19
MD
= 3
.76
PfJ
= 1
.27
MD
= 2
.80
PfJ
= 1
.31
MD
= 4
.74
PfJ
= 1
.61
MD
= 5
.22
PfJ
= 1
.59
MD
= 4
.87
PfJ
= 1
.79
MD
= 2
.71
PfJ
= 1
.24
MD
= 2
.86
PfJ
= 1
.25
MD
= 2
.48
PfJ
= 1
.19
X
N =
85
N =
936
N =
178
N =
541
PfJ
= 2
.34
MD
= 2
.81
PfJ
= 1
.59
MD
= 2
.69
PfJ
= 1
.26
MD
= 2
.99
PfJ
= 1
.30
MD
= 2
.78
PfJ
= 1
.26
MD
= 3
.42
PfJ
= 1
.44
PfJ
= 1
.31
MD
= 5
.98
PfJ
= 2
.02
MD
= 9
.89
PfJ
= 2
.70
MD
= 4
.86
PfJ
= 1
.60
MD
= 5
.93
PfJ
= 1
.87
MD
= 4
.40
PfJ
= 1
.91
MD
= 5
.78
MD
= 6
.23
Int-1
7-10
X
X
ZZ
Z
Int-1
4-7
Fig
.3.
Cry
sta
lp
refe
rre
do
rie
nta
tio
ns
(CP
O)
of
oliv
ine
,o
rth
op
yro
xe
ne
,cl
ino
py
rox
en
ea
nd
am
ph
ibo
lefr
om
InT
eri
ag
arn
et
pe
rid
oti
tex
en
olith
(a)
an
dsp
ine
lp
eri
do
tite
xe
no
lith
s(b
–d).
Sp
ine
lp
eri
do
tite
sa
reo
rga
niz
ed
inth
ree
gro
up
sd
efi
ne
db
yth
ere
lati
on
sb
etw
ee
nth
eo
rth
op
yro
xe
ne
,cl
ino
py
rox
en
ea
nd
am
ph
ibo
leC
PO
.(b
)T
yp
eA
,w
he
reo
liv
ine
,o
rth
op
yro
xe
ne
,cl
ino
py
rox
en
ea
nd
am
ph
ibo
lea
reco
he
ren
t;(c
)in
term
ed
iate
typ
eb
etw
ee
nT
yp
es
Aa
nd
B,
cha
ract
eri
zed
by
ap
art
ial
corr
ela
tio
nb
etw
ee
nth
eo
rth
op
yro
xe
ne
an
dth
ecl
ino
py
rox
en
ea
nd
am
ph
ibo
leC
PO
;(d
)T
yp
eB
,w
he
reo
liv
ine
an
do
rth
op
yro
xe
ne
CP
Oa
reco
he
ren
tb
ut
un
corr
ela
ted
toth
ecl
ino
py
rox
en
ea
nd
am
ph
ibo
leC
PO
.T
wo
sam
ple
so
fe
ach
typ
ea
reillu
s-tr
ate
d.
(Th
eco
mp
lete
da
tase
tis
giv
en
inS
up
ple
me
nta
ryD
ata
,F
ig.
A1
.)E
qu
al-
are
alo
we
rh
em
isp
he
rest
ere
og
rap
hic
pro
ject
ion
sin
the
stru
ctu
ral
refe
ren
cefr
am
e.
Nis
the
nu
mb
er
of
gra
ins
an
aly
zed
.C
on
tou
rsa
rea
t0�5
mu
ltip
les
of
un
ifo
rmd
istr
ibu
tio
nin
terv
als
.O
rth
op
yro
xe
ne
an
da
mp
hib
ole
ste
reo
plo
tsw
ere
con
tou
red
wh
en
mo
reth
an
10
0g
rain
sco
uld
be
me
as-
ure
din
the
thin
sect
ion
.M
Dis
ma
xim
um
de
nsi
ty;P
fJis
asc
ala
rm
ea
sure
of
the
stre
ng
tho
fth
ea
xis
ori
en
tati
on
.
Journal of Petrology, 2016, Vol. 0, No. 0 9
at University of C
ambridge on A
pril 12, 2016http://petrology.oxfordjournals.org/
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nloaded from
(c)
Sp
ine
l p
erid
otite
(In
term
ed
iate
)
Il-17
-9
MD
= 4
.25
PfJ
= 2
.00
MD
= 7
.06
PfJ
= 2
.90
MD
= 4
.04
PfJ
= 1
.54
MD
= 3
.14
PfJ
= 1
.30
MD
= 2
.95
PfJ
= 1
.49
MD
= 2
.88
PfJ
= 1
.34
MD
= 5
.21
PfJ
= 1
.78
MD
= 3
.66
PfJ
= 1
.40
MD
= 5
.89
PfJ
= 1
.82
MD
= 2
.32
PfJ
= 1
.18
MD
= 2
.18
PfJ
= 1
.18
MD
= 1
.90
PfJ
= 1
.11
N =
791
N =
122
0N
= 3
60N
= 3
04
MD
= 4
.73
PfJ
= 1
.77
MD
= 3
.88
PfJ
= 1
.34
MD
= 4
.99
PfJ
= 1
.75
N =
305
Il-17
-11
MD
= 2
.83
PfJ
= 1
.49
MD
= 4
.68
PfJ
= 2
.01
MD
= 3
.61
PfJ
= 1
.42
MD
= 3
.01
PfJ
= 1
.35
MD
= 3
.01
PfJ
= 1
.34
MD
= 3
.67
PfJ
= 1
.40
N =
401
N =
104
3N
= 2
17
MD
= 3
.86
PfJ
= 1
.32
MD
= 3
.61
PfJ
= 1
.57
MD
= 4
.12
PfJ
= 1
.46
X X
(d)
Sp
ine
l p
erid
otite
(Typ
e B
)
Il-17
-12
Il-17
-13
MD
= 3
.89
PfJ
= 1
.75
MD
= 4
.71
PfJ
= 1
.96
MD
= 2
.82
PfJ
= 1
.29
MD
= 4
.93
PfJ
= 1
.69
MD
= 3
.61
PfJ
= 1
.32
MD
= 4
.50
PfJ
= 1
.73
MD
= 3
.01
PfJ
= 1
.51
MD
= 3
.86
PfJ
= 1
.75
MD
= 2
.87
PfJ
= 1
.18
MD
= 4
.55
PfJ
= 1
.59
MD
= 4
.99
PfJ
= 1
.41
MD
= 3
.48
PfJ
= 1
.47
MD
= 2
.53
PfJ
= 1
.25
MD
= 3
.19
PfJ
= 1
.27
MD
= 3
.04
PfJ
= 1
.34
MD
= 2
.67
PfJ
= 1
.26
MD
= 3
.21
PfJ
= 1
.34
MD
= 3
.05
PfJ
= 1
.53
MD
= 2
.64
PfJ
= 1
.30
MD
= 2
.19
PfJ
= 1
.13
MD
= 1
.98
PfJ
= 1
.16
MD
= 3
.36
PfJ
= 1
.45
MD
= 3
.42
PfJ
= 1
.47
MD
= 3
.24
PfJ
= 1
.48
N =
162
N =
404
N =
277
N =
226
N =
327
N =
135
7
N =
211
N =
952
X X
[100
] [0
01]
[010
] C
linop
yrox
ene
elobihpm
Aenexorypohtr
OO
livin
e[1
00]
[001
] [0
10]
[100
] [0
01]
]100[ ]001[
]010[[0
10]
ZZ
ZZ
ZZ
ZZ
ZZ
ZZ
Fig
.3.
Co
nti
nu
ed
.
10 Journal of Petrology, 2016, Vol. 0, No. 0
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pril 12, 2016http://petrology.oxfordjournals.org/
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nloaded from
Tab
le2:
Ma
jor
(wt
%)
wh
ole
-ro
cka
na
lyse
so
fp
eri
do
tite
,p
yro
xe
nit
ex
en
olith
sa
nd
me
liliti
tes
Grt
Sp
ine
lp
eri
do
tite
s
pe
rid
ote
lhe
rzo
lite
sa
nd
ha
rzb
urg
ite
s
Ol-
we
bst
eri
tew
eh
rlit
eO
l-cl
ino
-p
yro
xe
nit
eS
am
ple
:In
t-1
4-7
*E
rem
-2*
Il-1
7-4
Il-1
7-5
Il-1
7-1
3Il-1
7-1
4Il-1
7-1
8In
t-1
5-1
*In
t-1
7-1
*In
t-1
7-6
*In
t-1
7-7
*In
t-1
7-1
2*
Int-
17
-13
*In
t-1
4-6
*ll-1
7-2
1In
t-1
7-1
4*
Il-1
7-1
6(-
2-)
(-9
-)(-
3-)
(-4
-)(-
5-)
(-6
-)(-
7-)
(-8
-)(-
1-)
SiO
24
6�1
44�7
43�0
43�0
44�9
44�9
44�0
44�2
41�3
42�6
44�4
43�5
45�1
34
9�8
49�3
50�4
48�1
Al 2
O3
3�2
82�1
51�7
72�1
32�4
02�0
22�0
41�9
01�3
70�6
02�1
82�0
82�5
40�9
83�1
93�5
55�0
1F
eO
10�3
7�5
78�3
88�2
17�7
07�4
57�6
07�6
08�5
18�0
17�4
77�9
67�9
88�6
04�8
84�8
26�6
1M
nO
0�1
50�1
20�1
50�1
60�1
60�1
30�1
30�1
20�1
50�1
40�1
20�1
30�1
20�1
70�1
10�1
10�1
3M
gO
35�8
41�1
40�6
40�4
38�9
39�4
41�1
42�4
44�2
45�7
42�2
40�4
39�4
37�3
28�1
25�5
21�2
Ca
O2�2
32�2
62�7
42�6
92�6
32�8
92�1
81�7
31�8
21�3
71�8
52�9
32�7
91�0
28�5
09�4
61
2�3
Na
2O
0�2
70�3
31�0
81�2
11�1
01�2
10�6
90�3
90�9
00�4
90�7
31�1
30�8
60�7
22�7
63�1
42�7
2K
2O
0�3
50�0
30�1
30�1
20�3
10�1
00�0
60�0
20�1
40�0
90�1
00�3
50�1
00�2
40�1
90�2
30�2
8T
iO2
0�5
30�0
70�0
70�0
60�0
70�0
40�0
60�0
70�0
60�0
30�0
70�1
10�1
10�0
70�1
40�1
40�2
6P
2O
50�0
30�0
30�0
40�0
40�0
40�0
70�0
30�0
30�0
40�0
40�0
40�0
30�0
30�0
30�0
30�0
80�0
6T
ota
l1
00�2
99�1
99�1
98�9
99�1
99�0
98�9
99�2
99�4
10
0�0
99�9
99�2
99�1
99�8
98�8
99�0
98�8
Mg
#.
0�8
60�9
10�9
00�9
00�9
00�9
10�9
10�9
10�9
00�9
10�9
10�9
00�9
50�8
90�9
10�9
10�8
5
Py
rox
en
ite
sM
eliliti
tes
Sa
mp
le:
Il-1
7-2
2Il-1
7-2
4Il-1
7-2
7In
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2).
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Table 3: Trace element (ppm) whole-rock analyses of peridotite, pyroxenite xenoliths and melilitites by solution ICP-MS
Garnet Spinel peridotiteperidotite Lherzolite and harzburgite
Sample: Int-14-7 Erem-2 Il-17-4 Il-17-5 Il-17-13 Il-17-14 Il-17-18 Int-15-1 Int-17-1 Int-17-6 Int-17-7 Int-17-12 Int-17-13(-2-) (-9-) (-3-) (-4-) (-5-) (-6-) (-7-) (-8-)
Rb 6�47 0�17 1�58 0�96 5�03 0�68 0�76 0�23 1�21 0�67 0�81 4�67 0�82Sr 23�2 45�9 143 107 107 128 70�2 32�6 82�4 44�7 92�0 83�7 110Y 2�50 2�50 3�77 2�95 4�60 4�96 2�50 2�50 2�50 2�50 2�68 3�72 2�88Zr 12�9 1�77 6�33 13�5 6�48 2�51 3�54 1�66 8�19 16�0 4�69 6�74 1�86Nb 1�35 0�37 22�4 2�80 21�2 27�0 4�09 0�26 19�7 13�1 10�2 21�1 7�38Cs 0�06 0�01 0�03 0�02 0�22 0�02 0�03 0�01 0�02 0�03 0�03 0�21 0�02Ba 65 33 64 32 336 97 63 98 38 37 108 255 31La 1�21 2�37 7�87 4�57 7�97 10�2 4�27 1�86 3�08 1�91 5�93 5�53 5�97Ce 2�72 3�03 17�9 10�1 20�4 26�7 7�92 2�03 7�90 4�83 12�8 14�6 11�3Pr 0�35 0�25 1�95 1�24 2�38 3�12 0�81 0�14 1�05 0�62 1�47 1�85 1�08Nd 1�63 0�85 7�40 5�32 9�61 12�6 2�99 0�48 4�54 2�62 5�68 7�57 3�63Sm 0�38 0�17 1�19 1�03 1�60 2�00 0�45 0�12 0�83 0�48 0�87 1�29 0�52Eu 0�14 0�07 0�38 0�35 0�52 0�66 0�16 0�05 0�27 0�15 0�28 0�41 0�18Gd 0�46 0�25 0�98 0�96 1�30 1�55 0�42 0�19 0�70 0�39 0�71 1�04 0�50Tb 0�08 0�05 0�14 0�13 0�18 0�21 0�06 0�04 0�09 0�05 0�09 0�14 0�08Dy 0�53 0�35 0�79 0�75 1�08 1�14 0�42 0�27 0�53 0�27 0�56 0�82 0�54Ho 0�11 0�08 0�14 0�13 0�19 0�20 0�09 0�06 0�09 0�04 0�11 0�15 0�11Er 0�30 0�23 0�37 0�31 0�50 0�50 0�24 0�17 0�21 0�10 0�28 0�37 0�32Tm 0�04 0�03 0�05 0�04 0�07 0�07 0�03 0�03 0�03 0�01 0�04 0�05 0�05Yb 0�27 0�23 0�33 0�24 0�43 0�43 0�23 0�17 0�18 0�08 0�26 0�32 0�32Lu 0�05 0�04 0�05 0�04 0�07 0�07 0�04 0�03 0�03 0�01 0�04 0�05 0�05Hf 0�42 0�07 0�13 0�14 0�11 0�03 0�08 0�06 0�11 0�18 0�09 0�15 0�07Ta 0�10 0�01 0�40 0�88 0�79 0�46 0�12 0�01 0�79 0�98 0�25 0�76 0�10Pb 0�31 1�20 0�67 0�37 0�47 0�53 0�65 0�96 0�52 0�6 3�12 0�62 0�80Th 0�15 0�73 9�04 0�36 1�36 0�78 0�79 0�57 0�34 0�11 1�07 1�10 0�57U 0�04 0�23 0�23 0�07 0�58 0�18 0�20 0�16 0�08 0�03 0�34 0�18 0�23
Spinel peridotite Pyroxenite Melilitite
Ol- Ol-clino- wehrlitewebsterite pyroxenite
Sample: Int-14-6 Il-17-16 ll-17-21 Il-17-22 Il-17-24 Il-17-27 Int-17-100 8814 8815 8816(-1-)
Rb 2�76 1�60 0�56 35�0 78�0 46�0 45�0 50�0 56�0 41�0Sr 68�9 306 333 283 142 142 174 1292 1209 1106Y 2�50 10�30 7�73 8�96 5�15 7�40 6�03 24�0 23�0 22�8Zr 1�99 18�6 8�10 186 218 429 107 359 333 317Nb 21�0 41�1 34�5 23�0 26�3 28�1 20�2 117 113 107Cs 0�10 0�02 0�01 0�32 0�80 0�76 0�48 0�46 0�60 0�46Ba 131 53 140 450 683 323 467 904 853 834La 7�84 15�7 10�6 11�3 3�30 3�17 3�53 90�0 87�0 89�5Ce 12�5 39�3 26�6 31�1 9�59 10�1 10�0 182 176 180Pr 1�21 4�82 3�07 4�29 1�50 1�71 1�61 20�0 19�0 19�6Nd 4�19 20�7 12�2 21�3 8�15 9�62 8�55 80�4 78�0 80�2Sm 0�57 3�79 2�13 4�56 2�05 2�58 2�16 13�3 13 13�3Eu 0�16 1�24 0�85 1�44 0�67 0�89 0�75 3�96 3�92 4�00Gd 0�38 3�26 1�88 4�09 2�09 2�8 2�26 10�6 10�4 10�7Tb 0�05 0�44 0�28 0�50 0�27 0�38 0�3 1�26 1�25 1�26Dy 0�31 2�49 1�69 2�57 1�47 2�06 1�62 6�34 6�22 6�25Ho 0�05 0�44 0�31 0�40 0�24 0�34 0�26 1�00 0�97 0�97Er 0�14 1�06 0�79 0�85 0�51 0�74 0�57 2�19 2�16 2�12Tm 0�02 0�14 0�11 0�10 0�06 0�09 0�06 0�26 0�25 0�25Yb 0�13 0�85 0�68 0�57 0�34 0�49 0�35 1�52 1�48 1�48Lu 0�02 0�13 0�10 0�08 0�05 0�07 0�05 0�22 0�21 0�21Hf 0�03 0�33 0�20 5�29 6�07 10�6 4�19 6�37 6�21 5�98Ta 0�48 1�46 0�77 2�12 3�04 3�29 2�58 7�56 7�76 7�70Pb 0�55 1�04 1�66 0�75 0�44 0�42 0�45 4�58 4�51 4�25Th 2�17 5�43 0�84 0�99 0�25 0�21 0�28 11�6 11�1 11�4U 0�18 0�46 0�27 0�21 0�06 0�08 0�09 3�18 2�95 2�84
Sample numbering as in Table 2.
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Spinel peridotiteLherzolite, harzburgite and Ol-websteriteLherzolite, harzburgite and one orthopyroxene-rich peri-
dotite (sample Int-14-6) have Mg# varying from 88�5 to
91�0, whereas Al2O3 and CaO contents vary between
0�98 and 2�54 wt %, and 1�02 and 2�7 wt %, respectively.
The whole-rock Ca/Al ratio is highly variable and
reaches values as high as 3�1 (sample IL-17-4). The ma-
jority of the In Teria lherzolites and harzburgites form a
group within the off-cratonic mantle array (Fig. 4a) with
compositions comparable with those of the most fertile
Cap Verde peridotite xenoliths (Bonadiman et al., 2005)
(Fig. 4b and c). Only three samples show distinct com-
positions. Two samples (Int-17-1 and Int-17-6), are more
refractory and plot within or close to the field of Canary
Island peridotite xenoliths (Neumann et al., 2004; Fig.
4b and c); the Ol-websterite (sample Int-14-6) has low
Al2O3, CaO and MgO contents compared with the other
peridotites (Fig. 4).
Lu varies from 0�012 to 0�13 ppm (Table 3), and is
positively correlated with Al2O3 and negatively with
MgO. Although showing variable REE contents and
distributions, all spinel peridotites are more enriched
in light REE (LREE) than the garnet peridotite (Fig. 5a
and c). On the basis of normalized REE patterns three
subgroups can be distinguished (Fig. 5). The first
group (group 1) comprises most of the lherzolites and
harzburgites and is characterized by almost straight
positive slope from heavy REE (HREE) to LREE (Fig.
5c). The other two groups (groups 2 and 3) are distin-
guished by concave-upwards REE patterns. The se-
cond group comprises the lherzolites Erem-2 and Int-
15-1; Erem-2 is the sample containing relicts of pri-
mary clinopyroxene. These two samples have a
roughly flat to slightly middle REE (MREE)-depleted
MREE–HREE segment [0�67< (Eu/Lu)N< 0�74] and are
enriched in LREE relative to MREE [(La/Sm)N� 8]. The
third group includes the lherzolite samples IL-17-18
and Int-17-13 and the Ol-websterite sample Int-14-6.
These samples are distinguished from the second
group by positive (Eu/Lu)N ratios, between 1�5 and 3�3,
and lesser LREE enrichment [6< (La/Sm)N� 8). On
Primitive Mantle (PM)-normalized diagrams group 1
peridotites are mostly characterized by negative
anomalies of Zr, Hf, and Ti relative to the MREE.
Conversely, several samples show positive anomalies
of Nb and Ta relative to Th, U, and LREE; the anoma-
lies tend to be marked in samples with lower Th, U,
and LREE contents. Except for a positive spike in Th in
one sample, group 1 peridotites also tend to be
depleted in Th and U relative to Ba and LREE. Group 2
samples show markedly distinct trace element pat-
terns, characterized by a strong negative anomaly in
Nb and Ta, whereas Zr, Hf, and Ti do not show signifi-
cant anomalies. Group 2 samples are also distin-
guished by selective enrichment of Th, U, Pb, and Sr.
In contrast, group 3 is more akin to group 1, particu-
larly in showing negative anomalies of Zr, Hf, and Ti,
and lacking substantial anomalies of Nb and Ta.
Wehrlite and Ol-clinopyroxenitesThe wehrlite and Ol-clinopyroxenites (Il-17-21, Int-17-14
and Il-17-16) have low bulk-rock MgO contents (21�2–
28�1 wt %), relatively high Al2O3 (3�19–5�01 wt %), and
elevated CaO contents (8�50–12�3 wt %; Fig. 4a, Table 2).
0
2
4
6
20 25 30 35 40 45 50 55
Ahaggar
Canary
Cape-Verde
(a)
Int-14-6
Canary
Ahaggar
Cape-Verde
(c)PM
Int-14-6
off-craton
cratonInt-14-6
IL-17-21
Int-17-14
Int-17-16
melting-trend
(b)grt. peridotitesp. peridotiteErem-2Wehrlite & Ol-cpx
0
2
4
35 40 45 50
CaO
(wt%
)A
l2O3 (
wt%
)
MgO (wt%)
Al2O
3 (w
t%)
0
1
2
3
4
35 40 45 50
PM
Fig. 4. Whole-rock major element compositions of peridotitexenoliths from In Teria compared with (a) off-craton xenoliths(small grey dots) and on-craton mantle xenoliths (small blackdots) from worldwide localities [data from Canil (2004)], and (b,c) the residual trend after melt extraction (black) from a fertilelherzolite source (PM) (Niu et al., 1997; Beccaluva et al., 2007).Fields for Cape Verde xenolith (Bonadiman et al., 2005);Tenerife, Lanzarote and Hierro xenoliths from the CanaryIslands (Neumann et al., 2004); and xenoliths from the Ahaggardistrict (Beccaluva et al., 2007) are shown for comparison.These xenoliths are anhydrous peridotites and metasomatizedspinel lherzolites and harzburgites. PM, Primitive Mantle com-position from McDonough & Sun (1995). grt, garnet; sp, spinel;Ol-cpx, Ol-clinopyroxenite.
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Their normalized REE and trace element patterns
(Fig. 5c and d) are almost indistinguishable from perido-
tite group 1 patterns, except for higher concentrations
of most elements.
PyroxeniteThe pyroxenites contain between 6�95 and 11�6 wt %
Al2O3 and between 6�0 and 14�9 wt % CaO, and their
MgO ranges between 11�5 and 16�9 wt % (Table 2). They
PM
nor
mal
ized melilitite
grt peridotite
PM
nor
mal
ized
pyroxenite
spinel peridotite
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu Gd
TbDy
HoEr
TmYb
Lu
0.1
1
10
100
1000
0.1
1
10
100
1000
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
0.1
1
10
100
1000
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Ch
norm
aliz
edC
h no
rmal
ized
spinel peridotite
melilitite
grt peridotite
pyroxenite
(a)
(c)
(g)
(b)
(d)
(h)
(f)
PM
nor
mal
ized
PM
nor
mal
ized
Ch
norm
aliz
edC
h no
rmal
ized
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
spinel peridotitespinel peridotite
0.1
1
10
100
1000
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
0.1
1
10
100
1000
0.1
1
10
100
1000
0.1
1
10
100
1000
0.1
1
10
100
1000
Ti
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu Gd
TbDy
HoEr
TmYb
LuTi
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu Gd
TbDy
HoEr
TmYb
LuTi
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu Gd
TbDy
HoEr
TmYb
LuTi
group 1
group 2
group 3
(e) Erem-2
Ol-cpx (Il-17-16)
Il-17-22
wehrlite (Il-17-21)
Fig. 5. Whole-rock REE and trace element compositions of (a, b) host melilitite and garnet peridotite; (c–f) spinel peridotite and(g, h) pyroxenite xenoliths. Spinel peridotites are separated into three groups that have different chemical compositions (see textfor discussion). Ol-clinopyroxenites and wehrlite are shown by open triangles (c, d) and lherzolite Erem-2 (with primary clinopyrox-ene) by grey filled circles (e, f). REE and whole-rock trace element contents are normalized to chondrite (Ch) and Primitive Mantle(PM), respectively; normalizing values after McDonough & Sun (1995).
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show convex-upward REE patterns, characterized
by positive slopes from the HREE to the MREE [5�6<Eu/
Lu)N< 7�4] and a roughly flat LREE segment [(La/
Sm)N¼ 0�77–1�55; Fig. 5g]. The PM-normalized trace
element patterns are distinguished by positive HFSE
anomalies in Nb, Ta, and Ti in all samples, and Zr–Hf in
two samples. The pyroxenites are also distinctively en-
riched in Rb and Ba relative to Th and U (Fig. 5h).
MelilititeThe compositions of the five analysed host-rock melili-
tites are homogeneous (Table 2). The samples are
undersaturated in SiO2 (�33�5 wt %) and have low total
alkali (K2OþNa2O) contents varying from 3�0 to 4�7 wt
%. The high Mg (MgO¼ 14�2–15�3 wt %), high Ni and Cr
contents indicate that these lavas have not been exten-
sively modified by fractional crystallization (Dautria
et al., 1992). The chondrite-normalized REE patterns
show a smooth positive slope from HREE to LREE
[48� (La/Yb)N� 70; Fig. 5a]. The PM-normalized trace
element patterns are convex upwards, with high but
relatively unfractionated abundances of the most in-
compatible elements (Rb to LREE), except for Pb, which
shows a negative anomaly relative to the LREE [(Pb/
Ce)N¼0�28]. Among the HFSE, only Zr and Hf show a
noticeable negative anomaly.
MINERAL CHEMISTRY
Major elementsGarnet peridotiteGarnet has a homogeneous, pyrope-rich composition
(Mg#¼0�80; Cr#¼ 0�04, Supplementary Data Table A2);
TiO2 increases from core (0�19 wt %) to rim (0�35 wt %).
Clinopyroxenes have high TiO2 contents (1�16 wt %) and
show some compositional similarities to secondary
clinopyroxenes from Manzaz in the Ahaggar district
(Beccaluva et al., 2007) and primary clinopyroxenes
from Cape Verde (Bonadiman et al., 2005). The ortho-
pyroxene porphyroclasts (Opx-I) are zoned; for in-
stance, Al2O3 increases from core to rim (3�10 to 4�09 wt
%, respectively; Supplementary Data Table A2), as does
Mg# (from 0�87 to 0�89). Olivine has a homogeneous
composition with low Mg#¼ 0�86. The phlogopite (Phl-
II) has Mg# 0�86 and a high TiO2 content (5�50 wt %,
Supplementary Data Table A2).
Using garnet–pyroxene geothermobarometry (Nickel
& Green, 1985; Brey & Kohler, 1990) core equilibrium
temperature and pressure have been estimated at
1050–1100�C and 2�5–2�7 GPa, respectively (Dautria
et al. 1992). However, pyroxene and garnet rim compos-
itions yield temperatures and pressures of 1210–1240�C
and 2�6–2�7 GPa.
Spinel peridotiteLherzolite, harzburgite and Ol-websterite. Primary clino-
pyroxenes (Cpx-I) from lherzolite Erem-2 have high
Al2O3 (6�78 wt %) and Na2O (1�74 wt %; Supplementary
Data Table A2) contents and resemble primary clinopyr-
oxene from the Ahaggar district (Beccaluva et al., 2007)
(Fig. 6a). The primary orthopyroxene (Opx-I) Mg# varies
from 0�87 to 0�91, and its Al2O3 content varies from 0�53
to 4�86 wt %. Ol-I have Mg# between 0�89 and 0�94 and
low CaO contents (0�2 wt %). In the secondary mineral
assemblage (II), the amphibole (Amph-II) is pargasite to
edenite with very variable Mg# (between 0�89 and 0�96)
and Na2O (3�30–5�10 wt %) contents, and low contents
of TiO2 (0�15–0�35 wt %) and K2O (0�06–0�60 wt %). The
phlogopite (Phl-II) shows less variability (Mg#¼ 0�92–
0�93) and is TiO2-poor (0�27–0–28 wt %) and K2O-rich
(6�8–7�0 wt %). Clinopyroxenes from the third assem-
blage (Cpx-III, Fig. 6a) have Mg# between 0�92 and 0�93,
high CaO (18�5–23�5 wt %), low but variable Al2O3 (0�83–
3�09 wt %), and low TiO2 (0�08–0�25 wt %) contents.
These compositions are similar to those of finely disse-
minated secondary clinopyroxene replacing orthopyr-
oxene in Cape Verde mantle xenoliths (Bonadiman
et al., 2005) and secondary large poikilitic clinopyroxene
from the Canary Islands (Neumann et al., 2002; Fig. 6a).
Ol-III displays a large variation in Mg# (0�89–0�94), en-
compassing the composition of Ol-I; however, Ol-III has
on average a more Mg-rich composition (Mg# 0�92)
Clinopyroxene
Olivine
(a)
(b)
17
19
21
23
25
0 2 4 6 8
CaO
(wt%
)
Al2O3 (wt%)
0
0.1
0.2
0.3
0.4
0.85 0.87 0.89 0.91 0.93 0.95 0.97
CaO
(wt%
)
Mg #
Cape VerdeCpx-II
Cpx-III
grt peridotite
Cpx-I (Erem-2)
Cape Verde
spinel peridotites
Ol-III
grt peridotite
Ol-III (Il-17-16)
Ol-Ispinel peridotites
Cpx-III (Il-17-16)Cpx-I (Il-17-16)
Cpx-I (Canary)
Ahaggar
Cape Verde
Ahaggar
Canary
AhaggarCpx-II
Fig. 6. Major element composition: Al2O3, CaO (wt %) and Mg#of (a) clinopyroxene and (b) olivine. Represented data are from11 samples analyzed in this study and three samples fromJ.-M. Dautria (unpublished data). Fields of olivine and clinopyr-oxene from Cape Verde (Bonadiman et al., 2005), Manzaz,Ahaggar (Beccaluva et al., 2007) and the Canary Islands(Neumann et al., 2004) are shown for comparison. Grt, garnet.Cpx-II indicates secondary clinopyroxenes from other studies.
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than Ol-I (Mg# 0�90, Fig. 6b). Olivine neoblasts within re-
action patches have high CaO contents (0�12–0�24 wt %)
compared with Ol-I CaO contents (0�01–0�03 wt %). The
glass analysed within the melt pockets is siliceous (up
to 68 wt % SiO2) and peraluminous (up to 25 wt %
Al2O3) in composition, and has low MgO, FeO
and CaO contents, resembling feldspar in its compos-
ition. This glass may have formed at the expense of
feldspar (Dautria et al., 1992). Its composition is similar
to that observed in many mantle xenolith suites (e.g.
Frey & Green, 1974; Ionov et al., 1993b; Coltorti et al.,
2000).Using the Brey & Kohler (1990) geothermometer cali-
bration on Opx-I and Cpx-I (sample Erem-2), Dautria
et al. (1992) obtained a core equilibrium temperature
between 770 and 870�C and a rim temperature increas-
ing up to 1080�C. According to the regional geotherm
(Lesquer et al., 1990), such temperatures would yield a
pressure range between 1�6 and 2�1 GPa. However the
absence of garnet indicates that pressure should be
lower than 1�9 GPa (Green & Ringwood, 1967; Wallace
& Green, 1988).
Wehrlite and Ol-clinopyroxenite. Primary olivine in
wehrlite has the same composition as Ol-I in spinel
peridotite with Mg# 0�89–0�90 and 0�01–0�03 wt % CaO
(Supplementary Data Table A2). The amphibole (Amph-
II) has a similar composition to Amph-II in the spinel
peridotites with Mg# 0�89, Na2O of 3�87 wt %, and low
TiO2 (0�27 wt %) and K2O (0�35 wt %) contents.
0.01
0.1
1
10
100
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu
TiGd
TbDy
HoY
ErTm
YbLu
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu
TiGd
TbDy
HoY
ErTm
YbLu
Spinel peridotite
Garnet peridotite
PM
nor
mal
ized
Pyroxenite
Spinel peridotite
Garnet peridotite
Pyroxenite
etipogolhPenexoryponilC
PM
nor
mal
ized
PM
nor
mal
ized
PM
nor
mal
ized
PM
nor
mal
ized
PM
nor
mal
ized
Cpx-III
(a)
(b)
(c)
(d)
(e)
(f)
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu
TiGd
TbDy
HoY
ErTm
YbLu
0.0001
0.001
0.01
0.1
1
10
100
1000
0.0001
0.001
0.01
0.1
1
10
100
1000
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu
TiGd
TbDy
HoY
ErTm
YbLu
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu
TiGd
TbDy
HoY
ErTm
YbLu
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu
TiGd
TbDy
HoY
ErTm
YbLu
0.01
0.1
1
10
100
Cpx-I(Erem-2)
0.01
0.1
1
10
100
0.0001
0.001
0.01
0.1
1
10
100
1000
megacryst
Fig. 7. Primitive Mantle normalized trace element patterns of clinopyroxene and phlogopite in the garnet peridotite, spinel perido-tites and pyroxenite xenoliths from In Teria. Values are normalized to the Primitive Mantle of McDonough & Sun (1995). Each pat-tern represents an average of several analyses in different grains (three on average). Error bars represent one standard deviationand are reported for representative samples. In (f) the phlogopite megacryst was analysed by solution ICP-MS.
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Ol-clinopyroxenite Il-17-16 contains clinopyroxene
phenocrysts and clinopyroxene in symplectites, which
have low Al2O3 (0�80–1�13 wt %) and high CaO (22�8–
23�4 wt %) contents (Fig. 2j; Supplementary Data Table
A2). Cpx-III forming the fine-grained reaction matrix dis-
plays a similar composition to that in other spinel peri-
dotites. This sample contains olivine neoblasts (Ol-III)
with a high CaO content (0�23 wt %) as observed in Ol-III
from spinel peridotites, but with lower Mg# (0�89)
(Fig. 6b).
PyroxeniteThe clinopyroxene porphyroclast cores are homoge-
neous in composition with Mg# from 0�77 to 0�78.
Phlogopite has Mg# varying between 0�73 and 0�74, and
TiO2 between 2�77 and 6�18 wt %, which is the typical
composition for upper mantle phlogopite (e.g. Delaney
et al., 1980; Ionov et al., 1997). Two distinct phlogopite
zoning trends from core to rim were observed: the first
shows a decrease in TiO2 (6�18–5�58 wt %) associated
with an increase in Al2O, Na2O and Mg# (e.g. sample
Int-17-89-100), and the second an increase in TiO2 from
2�77 to 4�76 wt % (sample Il-17-27). The amphiboles
have Mg# 0�70–0�72 and high K2O contents from 2�34 to
2�56 wt %. Temperature estimates for crystallization of
the pyroxenites range between 1100 and 1180�C
(Boissiere & Megartsi, 1982).
Trace elementsGarnet peridotiteGarnet has a high HREE content (LuN¼ 9�27) and a
strong LREE depletion relative to the HREE and MREE
[(La/Lu)N¼ 0�002; Supplementary Data Table A3].
Clinopyroxene has a convex PM-normalized REE pat-
tern, in which the MREE are enriched relative to the
HREE [(Eu/Lu)N¼7�1; Fig. 7a] and LREE [(La/
Sm)N¼ 0�56]. It is characterized by relatively unfractio-
nated abundances of LILE except for Pb, which shows a
negative anomaly [(Pb/Ce)N¼ 0�11]. The HSFE show
barely detectable positive anomalies [e.g. (Nb/
Ce)N¼0�16]. Orthopyroxene REE patterns display a rela-
tively flat HREE to MREE segment [(Eu/Lu)N¼1�01] and
a negative slope from MREE to LREE [(La/Sm)N¼0�13]
with well-marked Zr–Hf and Ti anomalies [(Zr/
Sm)N< 3�14; Fig. 8a]. Olivine displays enrichment from
Dy to Lu [(Dy/Lu)N¼ 0�19]. Phlogopite is characterized
by low REE abundances [(Lu)N¼ 0�010; (La)N¼ 0�028],
but prominent positive HFSE anomalies [(Nb/
La)N¼422; (Zr/Sm)N¼ 27�8; Fig. 7d] and strong Cs, Rb
and Ba enrichment, but low U–Th content
(Supplementary Data Table A3).
Spinel peridotiteLherzolite, harzburgite and Ol-websterite. The primary
clinopyroxene from sample Erem-2 displays a LREE-
PM
nor
mal
ized
spinel peridotite
garnet peridotite spinel peridotite
pyroxenite
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu
TiGd
TbDy
HoY
ErTm
YbLu
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu
TiGd
TbDy
HoY
ErTm
YbLu
PM
nor
mal
ized
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu
TiGd
TbDy
HoY
ErTm
YbLu
RbBa
ThU
NbTa
LaCe
PbPr
SrNd
ZrHf
SmEu
TiGd
TbDy
HoY
ErTm
YbLu
PM
nor
mal
ized
Orthopyroxene Amphibole(a) (c)
(b) (d)
PM
nor
mal
ized
0.001
0.01
0.1
1
10
Erem-2
0.1
1
10
100
0.1
1
10
100
0.001
0.01
0.1
1
10
100
Fig. 8. Primitive Mantle normalized trace element patterns of orthopyroxene from garnet and spinel peridotites and amphibolefrom spinel peridotites and pyroxenites. Values are normalized to the Primitive Mantle (PM) values of McDonough & Sun (1995).Each pattern represents the average of several analyses (three on average). Error bars represent one standard deviation and are re-ported for representative samples.
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depleted PM-normalized trace element pattern
[1�5� (La/Yb)N�4�3; Fig. 7b]. Such REE patterns can be
interpreted in terms of melt extraction after 7–12% frac-
tional or batch partial melting, respectively [estimated
using modal Cpx: 20% initial clinopyroxene
(YbN¼ 1�03)]. Erem-2 Cpx-I is similar to that reported
from the Manzaz volcanic district (Central Hoggar) by
Beccaluva et al. (2007). The trace element patterns of
metasomatic clinopyroxenes from melt pockets (Cpx-III)
show a continuous positive slope from HREE to LREE
[1�1< (Eu/Lu)N< 4�7; 0�7< (La/Sm)N<6�4] with well-
marked negative Zr–Hf and Ti anomalies but variable
Nb–Ta abundances compared with the neighbouring
REE [0�28< (Nb/La)N<10�5]. Opx–I from sample
Erem-2 displays a general negative slope and positive
Ti and Zr–Hf anomalies [(Zr/Sm)N< 2�4] (Fig. 8b).
Orthopyroxenes from the rest of the spinel peridotites
display a concave PM-normalized pattern and negative
Zr–Hf anomalies [0�12< (Zr/Sm)N< 0�75]. They show a
positive slope from MREE to LREE [2�1< (La/
Sm)N< 11�3]. The slope of the orthopyroxene trace
element patterns and the Zr–Hf anomalies are opposite
to those of orthopyroxene from the garnet peridotite.
Olivines display a large range of variation and enrich-
ment from Dy to Lu [0�08< (Dy/Lu)N< 0�45]. Amphibole
(Amph-II) is enriched in LREE relative to MREE and
HREE [3�3< (La/Sm)N< 11�9] and its PM-normalized REE
patterns mostly mimic those of Cpx-III. Amphiboles dis-
play negative Ti and Zr–Hf anomalies [(Zr/Sm)N¼ 0�10–
0�34], but positive Nb–Ta anomalies [2�1< (Nb/
La)N<6�5]. Damp/cpx(Zr) and Damp/cpx(Hf) are around
1�4 6 0�4 and 1�1 6 0�3, respectively, whereas Damp/
cpx(Nb) and Damp/cpx(Ta) are highly variable between 0�7and 40 and between 0�8 and 34, respectively. Previous
studies (e.g. Ionov et al., 1997) have reported a similar
variability and range of Damp/cpx for the HFSE. However,
LILE partitioning between Amph-II and Cpx-III is ex-
tremely variable and overall lower than estimated in
previous studies. For instance, despite high Ba contents
(53–519 ppm), Damp/cpx(Ba) varies between 0�8 and 456
and barely overlaps with the literature range of between
200 and 3000 (Adam et al., 1995; Brenan et al., 1995;
Table 4: Pb–Sr–Nd isotopic compositions of whole-rock garnet and spinel peridotites, pyroxenites, melilitites and phlogopitemegacrysts
206Pb/204Pb 207Pb/204Pb 208Pb/204Pb 87Sr/86Sr eSr(i) Classical leaching Stronger leaching
87Sr/86Sr eSr(i)143Nd/144Nd eNd(i)
Garnet peridotiteInt-14-7 18�5241 6 19 15�6101 6 19 38�3263 6 45 0�703303 6 4 –20 0�703049 6 5 –20 0�512946 6 5 6�2Int-14-7r 18�5496 6 6 15�6054 6 6 38�3364 6 15 0�703551 6 7 –13 0�512929 6 4 6�1Spinel peridotiteErem-2 18�5893 6 4 15�6070 6 4 38�3779 6 14 0�704299 6 6 –2�9 0�513091 6 6 9Il-17-4 19�7623 6 4 15�6782 6 4 39�4682 6 13 0�704419 6 8 –0�9 0�512958 6 7 6�5Il-17-5 19�6090 6 6 15�6727 6 6 39�3158 6 17 0�704472 6 19 –0�2 0�512924 6 5 5�7Il-17-13 19�8158 6 4 15�6712 6 3 39�4382 6 11 0�703694 6 10 –11�7 0�703359 6 4 –21 0�512928 6 3 5�9Il-17-14 19�8840 6 4 15�6751 6 3 39�5007 6 10 0�703753 6 5 –7 0�512916 6 2 5�6Il-17-14r 19�8150 6 3 15�6737 6 3 39�4453 6 9 0�703725 6 19 –10�7 0�512915 6 3 5�9Il-17-18 19�9803 6 5 15�6659 6 5 39�5092 6 17 0�703396 6 3 –15�5 0�512979 6 4 6�9Int-15-1 19�1522 6 4 15�6788 6 5 39�1135 6 10 0�704925 6 9 6 0�704063 6 3 –19 0�513100 6 10 9�1Int-17-1 18�6650 6 4 15�6091 6 4 38�4934 6 11 0�705307 6 3 11�6 0�703104 6 5 –19 0�512927 6 4 5�8Int-17-1r 18�6503 6 2 15�6080 6 2 38�4879 6 52 0�705452 6 12 14 0�512926 6 4 6�1Int-17-6 19�0202 6 26 15�6373 6 21 38�7406 6 56 0�704476 6 11 –0�4 0�703868 6 9 –20 0�512889 6 8 5�1Int-17-7 18�3595 6 4 15�6003 6 4 38�1725 6 13 0�703364 6 4 –16 0�703008 6 9 –21 0�512960 6 3 6�5Int-17-12 19�8433 6 5 15�6650 6 5 39�4140 6 13 0�703600 6 6 –13 0�703037 6 8 –20 0�512964 6 4 6�6Int-17-13 19�7701 6 6 15�6561 6 4 39�3444 6 12 0�705026 6 5 8 0�512986 6 3 7Int-14-6 18�7054 6 10 15�6138 6 9 38�4810 6 28 0�703836 6 18 –9�8 0�703093 6 6 –20 0�512912 6 4 5�6Ol-clinopyroxeniteIl-17-16 19�8029 6 5 15�6631 6 4 39�3954 6 16 0�704696 6 50 3�1 0�512936 6 4 6Il-17-16r 19�8137 6 5 15�6617 6 4 39�3286 6 13 0�704550 6 8 1 0�512940 6 4 6�4WehrliteIl-17-21 19�8922 6 5 15�6686 6 5 39�4490 6 14 0�703391 6 6 –15 0�512939 6 4 6�1PyroxeniteIl-17-22 19�6459 6 5 15�6416 6 4 39�2549 6 14 0�703267 6 4 –18 0�512955 6 2 6�3Il-17-24 19�7378 6 5 15�6550 6 4 39�3305 6 11 0�703448 6 6 –20�9 0�512936 6 3 5�9Il-17-27 19�4690 6 3 15�6662 6 3 39�2097 6 12 0�703495 6 12 –17�5 0�512917 6 4 5�5Int-17-100 19�5189 6 4 15�6695 6 4 39�2885 6 93 0�704991 6 6 4�3 0�705019 6 5 –21 0�512965 6 3 6�5Melilitite8814 19�6014 6 9 15�6417 6 8 39�2337 6 21 0�703164 6 6 –19�1 0�512950 6 4 6�38815 19�6408 6 6 15�6410 6 5 39�2496 6 13 0�703227 6 4 –18�6 0�512931 6 6 5�98816 19�6569 6 6 15�6374 6 6 39�2458 6 16 0�703148 6 3 –19�3 0�512953 6 4 6�4PhlogopiteInt-4 18�9286 6 3 15�5530 6 3 38�8550 6 9 0�704234 6 6 –40 0�512783 6 10 3Int-14 19�0874 6 5 15�6878 6 5 39�5024 6 14 0�704213 6 12 –40 0�512715 6 3 1�7Int-17 18�6836 6 6 15�6333 6 5 38�5971 6 16 0�703820 6 10 –46 0�512699 6 8 1�4
Errors on isotopic ratios are 2r. Italic lines represent the duplicates.
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Ionov et al., 1997; Gregoire et al., 2000; Powell et al.,
2004). Similarly, Damp/cpx(Rb) is always lower than three
for the In Teria peridotites, significantly below natural
and experimental data (100–800; Adam et al., 1995;
Brenan et al., 1995; Ionov et al., 1997; Gregoire et al.,
2000; Powell et al., 2004). In contrast, we obtain a Damp/
cpx(Sr) of about 1�8 6 0�7, which compares well with pre-
vious data (e.g. Ionov et al., 1997; Gregoire et al., 2000;
Powell et al., 2004) and experimental determinations
(e.g. Adam et al., 1995; Brenan et al., 1995). Phlogopite
is characterized by low REE abundances (Lu< 0�7 ppb).
PM-normalized patterns show an overall positive slope
from HREE to LREE–Th [7< (La/Sm)N� 22], marked by
pronounced positive HFSE anomalies [2< (Zr/
Sm))N� 27; 280< (Nb/La)N�14684] and strong LILE en-
richment relative to REE [e.g. 157< (Sr/Nd)N� 607;
7583< (Ba/La)N� 74080]. However, positive Zr and Hf
anomalies are mainly due to low HREE–MREE contents,
as Zr and Hf concentrations in phlogopite remain low
compared with Cpx, and yield a Dphl/cpx(Zr,Hf)�0�05,
lower than reported in the literature (e.g. Adam et al.,
1995; La Tourette et al., 1995; Ionov et al., 1997;
Gregoire et al., 2000). This is not the case for Nb and Ta,
which significantly partition into phlogopite relative to
clinopyroxene [e.g. 1�3<Dphl/cpx(Nb)� 16�5]. These are
within the spread of values reported from previous nat-
ural and experimental studies (e.g. Adam et al., 1995; La
Tourette et al., 1995; Ionov et al., 1997; Gregoire et al.,
2000). Whereas Cs and Ba partition strongly into
phlogopite relative to Cpx, Sr and Rb prefer Cpx, and U,
Th and Pb partitioning is variable [e.g. 0�4�Dphl/
cpx(U)�3�6].
Wehrlite and Ol-clinopyroxenite. Clinopyroxene trace
element patterns show a positive slope from the HREE
to the LREE relative to the MREE [(La/Sm)N� 1�9 6 0�5;
(Eu/Lu)N�3�3 6 0�2], similar to the patterns observed
for Cpx-III in other peridotites. The HFSE systematics
are also comparable with those of peridotite Cpx-III,
with negative Zr–Hf and Ti anomalies relative to the
MREE [e.g. (Zr/Sm)N�0�07] and positive Nb–Ta anoma-
lies relative to the LREE (e.g. (Nb/La)N� 2�28). The in-
compatible trace element patterns of amphibole
(Fig. 8c) are almost indistinguishable from those of
clinopyroxene (Fig. 7b).
PyroxeniteClinopyroxenes in pyroxenites have a general convex-
downward PM-normalized trace element pattern with a
positive slope from the HREE to MREE and a negative
slope from the MREE to LREE, with positive Zr–Hf and
both negative and positive Ti anomalies [(Zr/
Sm)N¼ 1�35–3�12] (Fig. 7c). These anomalies are the op-
posite of those observed in clinopyroxenes from spinel
peridotite. The Nb–Ta positive anomaly is well marked
only in sample Il-17-104 and is absent or weak for the
three other pyroxenites. Amphibole REE contents and
patterns are similar to those of clinopyroxene (Fig. 8d).
Like clinopyroxene, amphibole exhibits positive Zr–Hf
[(Zr/Sm)N¼ 2�33–4�03] and Ti anomalies. The opposite
was observed for the peridotite amphibole and clino-
pyroxene. Amphibole showing both Rb–Ba enrichment
and depletion has been observed in the same pyroxen-
ite sample (Fig. 8d). Phlogopites have low REE contents
and show an overall positive slope from MREE to LREE.
Their extended trace elements patterns are similar to
those described for peridotite xenoliths (especially for
the garnet peridotite; Fig. 7d and f) and are marked by
prominent positive HFSE anomalies [e.g. 160� (Nb/
La)N<6300), 27< (Zr/Sm)N< 237] and high Cs, Rb, Ba
and Sr abundances (e.g. Sr and Ba up to�170 ppm
and�2287 ppm, respectively). In contrast to phlogopite
in spinel peridotites, that in the pyroxenites shows low
U and Th abundances, as in the garnet peridotites. A
phlogopite megacryst analyzed by solution ICP-MS
shows a similar pattern (Fig. 7f; Supplementary Data
Table A2), except for the HREE. We note that DPhl/
Amp(Nb,Ta) and DPhl/Amp(Zr,Hf) are in the range of 0�6–2
and 0�04–0�09, respectively, as reported by Gregoire
et al. (2000) and Moine et al. (2001) (and references
therein). Similarly, the partitioning of Rb (9–22), Ba
(7–20), and Sr (0�1–0�5) between these two phases is
also broadly similar to that reported in previous studies.
Sr, Nd AND Pb ISOTOPIC DATA
Whole-rock isotopic compositions of mantle xenoliths,
phlogopite megacrysts and host melilitites are reported
in Table 4. The 87Sr/86Sr of the In Teria xenoliths varies
widely from 0�70327 to 0�70503 and does not correlate
with 143Nd/144Nd (þ5�5< eNd<þ6�6; Fig. 9a). Two spinel
peridotites (samples Erem-2 and Int-15-1) have higher143Nd/144Nd values of 0�51309 and 0�51310, respectively
(Table 4). In Sr–Nd isotope space, 11 mantle xenoliths
samples (one garnet peridotite, three pyroxenites and
seven spinel peridotites) and the melilitites are closely
grouped and plot close to the HIMU mantle end-mem-
ber. This group partly overlaps the fields for Ahaggar al-
kali basalts (Allegre et al., 1981), Morocco peridotites
(Raffone et al., 2009; Wittig et al., 2010; Natali et al.,
2013) and is close to the field for Libyan alkali basalts
(Beccaluva et al., 2008). Whereas the existing data for
northern Africa mantle xenoliths define an array charac-
terized by variable 143Nd/144Nd at relative low87Sr/86Sr—extending up to high 143Nd/144Nd values re-
cording time-integrated mantle depletion—the In Teria
xenoliths show higher and more variable 87Sr/86Sr val-
ues at nearly constant 143Nd/144Nd (Fig. 9a). Another
group of samples (eight spinel peridotites and one pyr-
oxenite) from In Teria have higher 87Sr/86Sr values in
the range 0�7045–0�7055 (Fig. 9a). The field for Canary
Islands peridotites and pyroxenites shows the same
characteristics, with variable 87Sr/86Sr values at con-
stant 143Nd/144Nd values (Whitehouse & Neumann,
1995; Neumann et al., 2002, 2015). For all In Teria
samples, the 87Sr/86Sr values are not correlated with
whole-rock Sr contents, suggesting that the high
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Canary
islands
(peridotites &
pyroxenites)
Canary
islands
(peridotites &
pyroxenites)
Canary
islands
(peridotites &
pyroxenites)
NH
RL
NHRL
207 P
b/20
4 Pb
208 P
b/20
4 Pb
143 N
d/14
4 Nd
HIMU
HIMU
HIMU
DMM
DMM
EM1
DMM
EM2
EM1
EM2
garnet peridotite
phlogopitemelilititespyroxenites
spinel peridotites
EM10.5124
0.5126
0.5128
0.5130
0.5132
0.5134
0.5136
0.5138
0.702 0.703 0.704 0.705 0.706
37.0
37.5
38.0
38.5
39.0
39.5
40.0
40.5
17 18 19 20 21 22
15.4
15.5
15.6
15.7
15.8
17 18 19 20 21 22
Ahaggar
(alkali basalts)
Ahaggar
(tholeiites)
Ahaggar
(Manzaz
peridotites)
Morocco
(peridotites)
206Pb/204Pb
87Sr/86Sr
Libya
(peridotites)
Libya
(alkali basalts)(a)
(b)
(c)
Ol-cpxlherz + harzb
Ahaggar
(tholeiites)
Morocco
(peridotites)
Morocco
(peridotites)
Ahaggar
(Manzaz
peridotites)
Ahaggar
(Manzaz
peridotites)
Libya
(alkali basalts)
Libya
(peridotites)
Libya
(peridotites)
Ahaggar
(tholeiites)
Ahaggar
(alkali basalts)
Ahaggar
(alkali basalts)
Erem-2 Int-15-1
Libya
(alkali basalts)
lherz Erem-2
Sicily
(peridotites &
pyroxenites)
gg
gg
Fig. 9. Variation of 143Nd/144Nd vs 87Sr/86Sr (a), 207Pb/204Pb vs 206Pb/204Pb (b) and 208Pb/204Pb vs 206Pb/204Pb (c) for in Teria melili-tites, pyroxenites, spinel peridotites, the garnet peridotite (Int-14-7) and phlogopite megacrysts. The following fields are shown forcomparison: Ahaggar alkali basalts (Allegre et al., 1981; Dupuy et al., 1993); Ahaggar tholeiitic basalts (Aıt-Hamou, 2000; Aıt-Hamou
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87Sr/86Sr values are not a secondary weathering effect.
The three phlogopites analysed yield low 143Nd/144Nd
values (þ1�4< eNd<þ3) at relatively high 87Sr/86Sr
(0�7038–0�7042) and plot in the field of tholeiitic basalts
from central Ahaggar (Aıt-Hamou et al., 2000).
Compared with In Teria whole-rock xenoliths, some of
the separated clinopyroxenes from the xenoliths from
Libya and Ahaggar (Manzaz district) show significantly
higher 143Nd/144Nd values (Beccaluva et al., 2007, 2008).
The Pb isotopic compositions of the In Terra xeno-
liths are in the range of 18�36–19�98 for 206Pb/204Pb,
15�60–15�69 for 207Pb/204Pb, and 38�17–39�51 for208Pb/204Pb (Table 4). In 207Pb/204Pb–206Pb/204Pb and208Pb/204Pb–206Pb/204Pb space, the samples define a
relatively large range (Fig. 9b and c). Seven spinel peri-
dotites, all the pyroxenites and the melilitites are char-
acterized by high 206Pb/204Pb and 208Pb/204Pb ratios
trending towards the HIMU end-member. This group of
samples overlaps the fields defined by alkali basalts
from Ahaggar and Libya (Allegre et al., 1981; Dupuy
et al., 1993; Beccaluva et al., 2008) and the fields of
Manzaz and Libyan peridotites (Beccaluva et al., 2007,
2008). The garnet-bearing peridotite and four spinel
peridotites show significantly lower 206Pb/204Pb and208Pb/204Pb values (<18�76 and <38�50, respectively)
associated with high 207Pb/204Pb ratios (>15�60) (Fig. 9b
and c) and partly overlap the field of Libyan peridotites
(Beccaluva et al., 2008). The three phlogopite mega-
crysts display intermediate 206Pb/204Pb values (18�68–
19�09), a large variation in 208Pb/204Pb and 207Pb/204Pb,
and plot between the two xenolith groups identified
previously (Fig. 9b and c).
DISCUSSION
Mantle deformationThe association of porphyroclastic microstructures,
marked by elongated olivine crystals with common sub-
grains, and the well-developed olivine and orthopyroxene
CPO indicate that the studied xenoliths are deformed. In
the garnet peridotite, CPO are characterized by the paral-
lelism of clinopyroxene and orthopyroxene [001] maxima
and the olivine [100] maxima (Fig. 3a). This suggests that
all the minerals accommodated the same deformation,
consistent with petrographic observations indicating that
the clinopyroxene is primary (Fig. 2a). In all spinel perido-
tites, the parallelism of olivine [100] and orthopyroxene
[001] maxima implies that the primary phases (Ol-I and
Opx-I) have recorded the same deformation event. In con-
trast to garnet peridotite, in most spinel lherzolites, the
lack of correlation between clinopyroxene and amphibole
CPO, together with petrographic observations, suggests
that clinopyroxene and amphibole are secondary and that
this metasomatic addition post-dated the main deform-
ation episode. An orthopyroxene crystallographic control
on the growth of clinopyroxene and amphibole might ex-
plain the partial correlation between clinopyroxene and
amphibole CPO and the olivine and orthopyroxene CPO in
the Type A spinel peridotites (Fig. 3b). An implication of
these observations is that prior to the metasomatic event,
the shallow lithospheric mantle beneath In Teria had a re-
fractory harzburgitic composition, whereas deeper levels
had more fertile compositions.
Analysis of the olivine CPO patterns also singles out
the garnet peridotite. The latter has a fiber-[010] olivine
(Fig. 3a), which implies simultaneous activation of
the [100](010) and [001](010) slip systems or a
Clinopyroxene
Orthopyroxene
pyroxenites
spinel peridotitesgarnet peridotite
Cpx-III
Cpx-I
(Ti/Eu)N
(b)
(a)
Erem 2
Cpx-I
0
2
4
6
8
10
12
14
16
18
20
0 0.5 1 1.5 2 2.5 3 3.5
0.001
0.01
0.1
1
10
0 10 20 30 40
(Ce/
Yb) N
100
(Ce/
Yb) N
Fig. 10. Chondrite-normalized (Ce/Yb)N vs. (Ti/Eu)N for (a)clinopyroxene and (b) orthopyroxene (Opx-I) from In Teria gar-net peridotite, spinel peridotites and pyroxenites. These dia-grams are based on those proposed by Coltorti et al. (1999).Chondrite normalizing values are from McDonough & Sun(1995). Cpx-I, primary clinopyroxene; Cpx-III, clinopyroxenefrom the third metasomatic assemblage in the spinel perido-tites (stage 3).
et al., 2000); alkali basalts from Libya (Beccaluva et al., 2008),peridotite xenoliths from Manzaz, Ahaggar (Beccaluva et al.,2007); peridotite xenoliths from the Middle Atlas, Morocco(Raffone et al., 2009; Wittig et al., 2010; Natali et al., 2013); peri-dotite xenoliths from Libya (Beccaluva et al., 2008); peridotitexenoliths from Sicily (Bianchini et al., 2010); mantle xenolithsfrom Tenerife, Hierro and Fuerteventura, Canary Islands(Whitehouse & Neumann, 1995; Neumann et al., 2002, 2015).HIMU, DMM, EM2 and EM1 mantle end-member compositionsare from Zindler & Hart (1986) and Eisele et al. (2002). TheNorthern Hemisphere Reference Line (NHRL) is from Hart(1984).
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transpressional deformation (Tommasi et al., 1999,
2000). However, the strong [001] maximum of the pyr-
oxene CPO is not consistent with transpressional de-
formation. We suggest therefore that this olivine CPO
records deformation at high stress or under high-pres-
sure conditions (e.g. Jung & Karato, 2001; Mainprice
et al., 2005; Demouchy et al., 2013). The spinel perido-
tites, on the other hand, show olivine CPO with ortho-
rhombic or fiber-[100] symmetries, consistent with
dominant activation of the (010)[100] slip system, sug-
gesting deformation at low-stress, high-temperature,
low-pressure conditions (Tommasi et al., 2000).
Metasomatic imprintsBased on petrographic evidence and geochemical data,
three metasomatic imprints are recognized in the In
Teria xenoliths, as follows.
1. The garnet and spinel peridotites were first affected
by crystallization of amphibole and/or phlogopite at
the expense of orthopyroxene, a feature that was
ascribed by Dautria et al. (1992) to reactive porous
percolation of a hydrous alkaline melt. Although
it shows a lesser extent of modal metasomatism
(crystallization of phlogopite; Fig. 2b), the garnet
peridotite (sample Int-14-7) is characterized by a con-
vex-upward normalized REE pattern in clinopyroxene
(Fig. 7a), which suggests equilibration with a LREE-
enriched melt (Irving, 1980; Irving & Frey, 1984;
Bodinier et al., 1987). Moderate LREE enrichment
without significant negative HFSE anomalies (or with
positive anomalies) in clinopyroxene (or whole-rocks)
is widely considered as evidence for mantle meta-
somatism involving alkaline silicate melts (e.g.
Downes, 2001; Bodinier et al., 2004; Powell et al.,
2004; Kaeser et al., 2006). Together the observations
on the garnet and spinel peridotites suggest that the
lithospheric mantle beneath In Teria was pervasively
metasomatized by hydrous alkaline melts.
2. The second stage (stage 2) of metasomatism is akin
to the percolation of carbonatitic melt and has been
well documented by Dautria et al. (1992). It is re-
stricted to the spinel peridotites, where it is respon-
sible for the crystallization of Cr-rich diopside at the
expense of orthopyroxene and amphibole. It results
in the formation of microgranular aggregates (for-
mer melt pockets) composed of secondary clino-
pyroxene, olivine, Al- and Cr-spinel, sulphides,
K-feldspar and glass, which extensively replace theprimary phase assemblages (Fig. 2h and i). Most of
the major and trace element characteristics of these
samples and their minerals are typical of mantle
metasomatism by carbonate melt (Woolley &
Kempe, 1989; Yaxley et al., 1991; Ionov et al., 1993a;
Hauri & Hart, 1994; Blundy & Dalton, 2000;
Chakhmouradian, 2006). These characteristics not-
ably include a strong enrichment of highly incompat-
ible trace elements in whole-rocks and pyroxenes
(Opx-I and Cpx-III), negative HFSE (Ta, Zr, Hf and Ti)
anomalies in Cpx-III, and negative Zr and Hf anoma-
lies in Opx-I (Figs 5 and 7). Ol-III have high Mg# and
are enriched in CaO compared with Ol-I, and Cpx-III
are enriched in Cr (Fig. 6). Moreover, secondary
clinopyroxene (Cpx-III) and coexisting orthopyrox-
ene (Opx-I) show evolutionary trends in Fig. 10
marked by a strong increase in (Ce/Nb)N at rela-
tively low (Ti/Eu)N, confirming interaction with a car-
bonatitic melt. The possibility that the melt pockets
were formed by interaction with the host lava during
the exhumation of the peridotite xenoliths is
excluded because only spinel peridotites show this
interaction (i.e. it is not observed in the garnet peri-
dotite and pyroxenites).
3. The metasomatic history of the In Teria xenolithsuite concludes with the crystallization of pyroxen-
ites (stage 3). These rocks, which are predominant
among the xenolith suite, crystallized in the spinel
stability field, lack the carbonate-melt imprint (stage
2) observed in all spinel peridotites and resemble
the amphibole-pyroxenite veins found in Pyrenean
peridotites (Bodinier et al., 1987, 2004) with respect
to their major and trace element characteristics
(Figs 6 and 11). Trace element similarities between
the In Teria and Pyrenean pyroxenites include
0
5
10
15
20
25
30 35 40 45 50 55 60
pyroxenite xenoliths (this study)literature data for mantle pyroxenites
to carbonatite
SiO2 (wt%)
CaO
(wt%
)
melilitites (this study)
(3) (2)
(1)
amphibole pyroxenites veins from Lherz
Fig. 11. Variation of CaO vs SiO2 for the In Teria pyroxenitexenoliths and host melilitites, compared with literature data forpyroxenites and basaltic to carbonatitic mantle melts (shadedfields). The pyroxenite dataset includes samples from the fol-lowing xenolith suites: Hawai (Frey, 1980); Australia (Griffin &O’Reilly, 1986; O’Reilly & Griffin, 1987; O’Reilly et al., 1988);Sidamo, Ethiopia (Bedini, 1994); orogenic lherzolite massifs(Lherz, France, Bodinier et al., 1987, 1990; McPhail et al., 1990;Ronda, Spain, Garrido & Bodinier, 1999); ophiolites (Thailand,Orberger et al., 1995; New Caledonia, J. L. Bodinier, unpub-lished data); the intrusive pyroxenite–carbonatite body ofTamazeght, Morocco (Marks et al., 2008). The overlap betweensome of the In Teria pyroxenites and the amphibole-bearingpyroxenite veins from Lherz (Bodinier et al., 1987; Bodinier,1989) should be noted. Shaded lava fields: (1) South African al-kaline basalts and olivine melilitites (Janney et al., 2002); (2)tholeiites, transitional basalts, basanites and nephelinites fromSicily, Italy (Beccaluva et al., 1998); (3) undersaturated alkalineto carbonatitic melts from Morocco (Wagner et al., 2003).
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convex-upward normalized REE patterns and posi-
tive HFSE anomalies in whole-rocks and amphibole
(Nb, Ta, Zr, Hf, Ti) and, to a lesser degree, in clino-
pyroxene (Zr, Hf, 6 Ta). In Fig. 10a, the pyroxenites
show a trend defined by a strong increase in the (Ti/
Eu)N ratio of clinopyroxene correlated with a moder-
ate increase in (Ce/Yb)N. This evolutionary trend di-
verges markedly from that defined by the secondary
clinopyroxene in the spinel peridotites. The In Teria
pyroxenites probably represent crystal segregates
from hydrous alkaline melts crystallized in vein
conduits—rather than in large intrusions.
Metasomatism: several events or multi-stagemelt–rock interactionMicrostructural observations indicate that all metasom-
atic events postdate the last deformation event, re-
corded by olivine in both garnet and spinel peridotites,
that affected the lithospheric mantle in this region. The
alignment of the phlogopite–amphibole aggregates in
the foliation, but the undeformed character of the crys-
tals and the dispersion of crystallographic axes of
Amph-II and Cpx-III (e.g. Type B in Fig. 3d) indicate that
although melt infiltration postdated the deformation, it
was at least partially controlled by the pre-existing
structure of the peridotites. A stronger structural control
of the pre-existing microstructure on the reaction prod-
ucts is also indicated by the similarity of Amph-II and
Cpx-III CPO in Type A spinel lherzolites (Fig. 3b), which
suggests topotaxial growth of at least part of the meta-
somatic products.
Our geochemical data show striking similarities be-
tween the mineral compositions of peridotites affected
solely by stage 1 metasomatism and those of stage 3
pyroxenites. This may be exemplified by comparing the
trace element signatures of clinopyroxene and phlogo-
pite in the garnet peridotite sample and in the pyroxen-
ites (Figs 7 and 10a). Moreover, our data show the clear
geochemical affinity of the pyroxenite xenoliths (stage 3)
with the melilitite host lavas. As illustrated in Fig. 11, the
In Teria pyroxenites differ from other mantle pyroxenites
(including, to some degree, the Pyrenean veins) in their
lower SiO2 contents and as such are compositionally
akin to strongly undersaturated and carbonatitic mag-
mas. The equilibrium melt calculated from the trace
element composition of clinopyroxene in pyroxenite is
comparable with the host melilitites—as well as with the
garnet peridotite equilibrium melt [Fig. 12; partition coef-
ficients (KD) from Hart & Dunn (1993)]. The only signifi-
cant differences are the positive HFSE anomalies
(particularly for Zr and Hf) that are observed in the pyrox-
enite equilibrium melt (but not in the garnet peridotite
equilibrium melt). These anomalies probably reflect the
inadequacy of available experimental clinopyroxene/
melt KD values for clinopyroxenes crystallized from
strongly SiO2-undersaturated alkaline melts.
The microstructural observations and the geochem-
ical affinity of the xenoliths with the melilitite host lavas
favour a scenario whereby the different metasomatic im-
prints record successive stages of interaction between
lithospheric mantle and sublithospheric melts. The evo-
lution from stage 1 to stage 2 (silicate/hydrous to carbon-
ate melt) may result from a differentiation process
involving melt ingress down a lithospheric thermal gradi-
ent associated with melt–rock reactions at decreasing
melt mass (Bedini et al., 1997; Bodinier et al., 2004).
mel
t / P
rimiti
ve M
antle
Ba Th U Nb Ta La Ce Pr Sr Nd Zr Hf SmEu Ti Gd Tb Dy Ho Y Er TmYb Lu
melt: Cpx-III sp peridotites
melt: Cpx pyroxenite (Il-17-27)
melt: Cpx-Isp peridotite(Erem-2)
melilitite (8814)
melt: Cpx-I grt peridotite (Int-14-7)
0.1
1
10
100
1000
10000
Fig. 12. Calculated compositions of theoretical melts in equilibrium with Cpx-I (Erem-2) and Cpx-III in spinel peridotites, the garnetperidotite and a pyroxenite sample. We used the partition coefficients between Cpx and basaltic melt of Hart & Dunn (1993) for thespinel peridotite, the garnet peridotite and the pyroxenite, and the partition coefficients between Cpx and carbonatitic melt ofKlemme et al. (1995) for secondary Cpx (Cpx-III) in the other spinel peridotites. A host melilitite (sample 8814) is shown for compari-son. sp, spinel; grt, garnet.
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Gradual solidification of the melt, also referred to as ‘per-
colative fractional crystallization’ (Harte et al., 1993),
would be responsible for marked changes in its compos-
ition, culminating in the individualization of volatile-rich
and/or carbonated small melt fractions that are expected
to be strongly enriched in highly incompatible elements
(McKenzie, 1989). Equilibrium melts calculated from the
trace element composition of In Teria Cpx-III are enriched
in incompatible elements and display negative HFSE
anomalies in their PM-normalized patterns [Fig. 12; KD
from Klemme et al. (1995)]. This signature is similar to
that obtained by numerical simulations of reactive por-
ous flow at decreasing melt mass (Bedini et al., 1997;
Ionov et al., 2002a).
Trace element modelling predicts that the combin-
ation of chromatographic effects related to melt trans-
port and ‘source’ effects related to melt–rock reactions
results in spatial decoupling between melt–rock reac-
tions and their trace element expression (Godard et al.,
1995). This may be responsible for transient depletion
of several incompatible elements in porous-flow sys-
tems, owing to buffering of the percolating melt by the
depleted peridotite protolith (see, e.g. fig. 13 of Bodinier
et al., 2008). In contrast to other pyroxenes from the In
Teria xenolith suite, Cpx-I and Opx-I of the spinel peri-
dotite Erem-2 are depleted in LREE (Figs 7, 8, 10 and
12). The unique trace element signature of this sample
may represent a remnant of the pre-metasomatic,
depleted composition of the lithospheric mantle be-
neath In Teria. Alternatively, it may also be explained
by transient buffering of the percolating melt by the
peridotite protolith.
Melt compositions appear to have varied with perco-
lative differentiation from siliceous and hydrous, but
locally impoverished in highly incompatible trace elem-
ents relative to the infiltrated melt (e.g. sample Erem-2
equilibrium melt with cpx-I; Fig. 12), to carbonated and
strongly enriched in highly incompatible trace elem-
ents, but with negative HFSE anomalies (other spinel
peridotites equilibrium melt in Fig. 12). Clinopyroxene
trace element compositions in the garnet peridotite and
pyroxenites (Fig. 7) and the similarity between the equi-
librium melt for the garnet peridotite and the host meli-
litite (Fig. 12) indicate that the garnet peridotite has
interacted with a slightly evolved silicate melt. This may
have a deeper origin than the spinel peridotites and/or a
100 mWm-2
100 mWm-2
60 mWm-2 80 mWm-2 80 mWm-260 mWm-2
Pre
ssur
e G
Pa
Temperature ºC
garnet
4.0
spinel
3.0
2.0
1.0
0.0
700 800 900 1000 1100 1200 1300 1400 1500
700 800 900 1000 1100 1200 1300 1400 1500
plagioclase
spinel
parg out
solidus
hydrous
peridotite
solidus
anhydrous
peridotite
spinel peridotites (T from mineral cores)
grt peridotite (Int-14-7) mineral cores
grt peridotite (Int-14-7) mineral rims
spinel peridotites (T from mineral rims)
Fig. 13. Pressure and temperature estimates for the In Teria mantle xenoliths. For the spinel peridotites the temperatures were ob-tained from mineral cores and rims using the Na-in-Opx/Cpx geothermometer of Brey & Kohler (1990). The pressure conditions forthe xenoliths were constrained by the absence of plagioclase and garnet in the peridotites. For the garnet peridotite, both pressureand temperature estimates were obtained from mineral cores and rims, using the Nickel & Green (1985) Al in Opx/Grt geobarome-ter and the Brey & Kohler (1990) Na-in-Opx/Cpx geothermometer. The plagioclase–spinel and spinel–garnet phase transitions arefrom Gasparik (1987) and Gasparik (1984), respectively. The pargasite stability field (parg. out) is from Niida & Green (1999), andthe anhydrous and hydrous peridotite solidus is from Taylor & Green (1983). The geotherms for the Ahaggar shield and the In-Salah–Illizi zone are from Lesquer et al. (1990). They were calculated for surface heat flows in the range 80–100 mW m–2, mean crus-tal heat productions of 0�8mW m–3 (continuous lines) and 1�2 mW m–3 (dashed lines), a 30–35 km thick crust and no contributionfrom the mantle.
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possible origin in the wall-rock of a magma conduit.
The garnet peridotite records higher pressure–
temperature (P–T) conditions than the spinel peridotites
and is devoid of any carbonatitic metasomatic over-
print, suggesting a deeper origin.The multi-stage metasomatic event that affected the
In Teria xenoliths was probably coupled with a transient
thermal event; that is, heating of subcontinental litho-
sphere followed by thermal relaxation. Heating is re-
corded by the compositional zoning of the primary
minerals, indicating a core to rim increase of equilib-
rium temperatures from 1050–1100� to 1210–1240�C for
the garnet peridotite sample Int-14-7 and from 770–
870�C to about 1080�C for the spinel peridotite sample
Erem-2 (Fig. 13). Combined with pressure estimates,
these rim temperatures define a rather steep thermal
gradient roughly consistent with the high heat flux
measured in the In Salah–Illizi district (>100 mW m–2;
Lesquer et al., 1990). The rim P–T estimates lie in the
interval between the hydrous and anhydrous peridotite
solidus (Fig. 13) and may therefore record lithospheric
P–T conditions during the first metasomatic stage. The
intergranular porous flow of hydrous alkaline melt
inferred for this stage requires ambient temperatures in
excess of (at least) the hydrous peridotite solidus. The
observation that phlogopite is the only metasomatic
mineral formed in the garnet peridotite is consistent
with the high ‘rim’ temperature of this sample, well be-
yond the stability field of pargasite (Niida & Green,
1999; Fig. 13). Compared with amphibole, phlogopite is
stable at higher temperatures (Wendlandt & Eggler,
1980; Mengel & Green, 1989). Conversely, the predom-
inance of pargasite over phlogopite in spinel peridotites
would reflect the lower temperatures of the shallower
lithosphere during stage 1, overlapping the amphibole
stability field (Fig. 13).
In a scenario in which carbonate melt is produced by
reactive differentiation of percolating silicate melt down
a thermal gradient, the silicate and carbonate melt
metasomatism are coeval but disconnected in space
(Bedini et al., 1997; Bodinier et al., 2004). During con-
ductive heating of the lithosphere, of an asthenospheric
origin, silicate melt metasomatism would occur in
deeper and/or hotter lithosphere whereas carbonate
melt metasomatism would affect shallower and/or
colder domains (Fig. 14a). Such an arrangement of
metasomatic aureoles related to a thermal gradient has
been observed on a small scale in vein wall-rocks within
the Lherz peridotite (Bodinier et al., 2004). In the In Teria
suite, the lack of a carbonate melt imprint in the deeper
garnet peridotite, whereas it is pervasive in the shal-
lower spinel peridotites, provides evidence for a similar,
thermally controlled arrangement of metasomatic do-
mains at lithospheric scale.
In the proposed scenario, overprinting of stage 1
metasomatism in spinel peridotites by stage 2 meta-
somatism is explained by downwards subsidence of
the metasomatic domain owing to lithospheric thermal
relaxation (Fig. 14b). The lack of any carbonate-melt im-
print in the In Teria pyroxenites, in contrast to the spinel
garnetspinel60
30
35
40
45
carbonate melt metasomatism
50
55
(a) (b) (c)
(km)
silicate melt metasomatism (reactive percolation of hydrous alkaline melt at decreasing melt mass)
continental
crust
silicate melt metasomatism
carbonate melt metasomatism
(reactive percolation of evolved carbonate
melt fractions)
pyroxenites
~1100˚C
~1100˚C
the
rma
l re
laxa
tio
n
~1100˚
Fig. 14. Schematic illustration of the three stages of metasomatism associated with the lithospheric thermal evolution proposed forthe In Teria mantle xenoliths: (a) upon conductive heating of the lithosphere by an asthenospheric source, silicate melt metasoma-tism occurs in the deeper and/or hotter parts of the lithosphere whereas carbonate melt metasomatism is restricted to shallowerand/or colder domains; (b) during subsequent thermal relaxation and downwards subsidence of the metasomatic domains of stage1, silicate melt metasomatism in the spinel peridotites is overprinted by carbonate melt metasomatism (stage 2); (c) upon furtherlithospheric cooling extensive hydraulic fracturing allows the injection of silicate melts forming a dense network of pyroxenite veins(stage 3)
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peridotites, which are extensively metasomatized, sug-
gests that the pyroxenite veins were late stage. The em-
placement of a dense vein network—judging from the
predominance of pyroxenites within the xenolith
suite—might be associated with further lithospheric
cooling favouring extensive hydraulic fracturing
(Fig. 14c). Based on their similarity to phlogopite within
the pyroxenites, the phlogopite megacrysts may repre-
sent deep crystal segregates from the infiltrating alka-
line silicate melt (Fig. 7).
Based on the time constraints necessary to allow
successive heating and thermal relaxation of the litho-
spheric mantle, the metasomatic evolution proposed in
our model probably lasted for the whole Cenozoic era.
Stage 1 heating was possibly related to Eocene tholei-
itic magmatism, for which a lithospheric contribution
has been suggested (Maza et al., 1998), implying sub-
stantial heating. The resemblance of In Teria phlogopite
megacrysts to Ahaggar tholeiitic volcanism in terms of
their Nd–Sr isotope composition (Fig. 9) lends support
to this hypothesis and suggests that the phlogopites are
related to stage 1. The suggested subsequent evolution
would have lasted until the recent melilitite volcanism,
based on the observation that stage 3 pyroxenites prob-
ably represent mantle segregates from melilitite mag-
mas. The relationship between the suggested
lithospheric thermal evolution and the higher heat flow
observed in the In Teria region, compared with the rest
of the Ahaggar Massif (Lesquer et al., 1989), is poorly
understood. The presence of garnet peridotites in a re-
gion of higher heat flow is paradoxical. The origin of the
high heat flux is probably related to the situation of In
Teria at the southern border the Sahara Basins (Fig. 1),
characterized by a regional-scale heat flow anomaly
(Takherist & Lesquer, 1989; Lesquer et al., 1990), rather
than to the mantle upwelling beneath Ahaggar.
Origin of 87Sr enrichmentAmong the In Teria xenoliths, the enrichment in 87Sr with-
out any concomitant decrease of 143Nd/144Nd is restricted
to the spinel peridotites and one pyroxenite (sample Int-
17-100). This enrichment is not observed in the host meli-
litites, which show a narrow range of isotopic compos-
itions comparable with those of the Ahaggar alkaline
basalts (Fig. 9). Together with the lack of correlation be-
tween 87Sr/86Sr and Sr concentration, this does not not
support a hypothesis of late-stage introduction of radio-
genic Sr into the studied samples. The enrichment in 87Sr
is not related to the degree of metasomatism either, nor
to the metasomatic enrichment in incompatible trace
elements. Although extensively re-equilibrated with alka-
line-silicate and carbonate melts, respectively, the garnet
peridotite and the Ol-clinopyroxenite (sample Il-17-21) are
not significantly enriched in radiogenic Sr
(87Sr/86Sr< 0�7035; Fig. 9a). Conversely, the spinel perido-
tite Erem-2, which is the least metasomatized sample
with LREE-depleted pyroxenes (Fig. 10), and the phlogo-
pite megacrysts are relatively enriched in radiogenic Sr
0.5124
0.5125
0.5126
0.5127
0.5128
0.5129
0.5130
0.5131
0.5132
0.7030 0.7035 0.7040 0.7045 0.7050 0.7055
(b+d)
EMI
143 N
d/14
4 Nd
HIMU
(a)
(c)
(f)(g)
(h)
(e)
mixing curves
87Sr/86Sr
Ahaggar
(tholeiites)
Ahaggar
(alkali basalts)
Erem-2 Int-15-1
Fig. 15. Variation of 87Sr/86Sr vs 143Nd/144Nd for the In Teria melilitites, pyroxenite and peridotite xenoliths and phlogopite mega-crysts, compared with results of one-dimensional percolation–diffusion modelling aiming to explore the effects of different com-positional parameter values retrieved from the studied xenoliths and their host lavas (see text). The continuous lines a–h illustratemodel runs performed with compositional parameter values reflecting the range of Sr–Nd contents observed in the studied perido-tites for the protolith, and in the host melilitites and clinopyroxene equilibrium melts for the infiltrated melt. The complete set ofparameter values is given in Supplementary Data Table A5. The dashed lines show the range of mixing curves calculated with thesame compositional parameters as the one-dimensional percolation–diffusion modelling. The symbols for the In Teria xenolithsamples, and the references for the Ahaggar basalt fields and the mantle EM1 and HIMU end-members, are the same as in Fig. 9.
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(87Sr/86Sr> 0�704 and > 0�7038, respectively). Elevated87Sr/86Sr values found in a few limestone-hosted, Ca-rich
carbonatites from Tamazert (Morocco) are clearly related
to crustal contamination (Bouabdellah et al., 2010). To
test the origin of high 87Sr/86Sr values in the In Teria xeno-
liths we performed a stronger leaching on seven spinel
peridotites, the garnet peridotite and one pyroxenite
(Table 4). The new 87Sr/86Sr compositions for most sam-
ples are less radiogenic and closer to HIMU than previ-
ously (Supplementary Data Fig. A4). These results
suggest that the high 87Sr/86Sr values could be related to
the crystallization of interstitial micro-phases from small-
volume melts during the waning stages of metasoma-
tism. Detailed studies of trace element distribution in
metasomatized mantle xenoliths have shown that a sig-
nificant proportion of the whole-rock budget of large ion
lithophile elements (LILE, including Sr) is hosted by such
intergranular components (Bedini & Bodinier, 1999).
Recently, Kourim et al. (2014) have revealed the existence
of such a LILE-enriched interstitial component in mantle
xenoliths from southwestern Ahaggar. Chromatographic
theory predicts that downstream in a reactive percolation
column (i.e. at very low melt/rock ratio), the isotopic sig-
nature of the percolating melt may be decoupled from its
trace element signature: the melt may be isotopically
equilibrated with the peridotite protolith although
strongly enriched in highly incompatible elements as a
result of reactions at decreasing melt mass (Bodinier
et al., 2004). Owing to their ability to percolate as small-
volume melts through relatively cold peridotite
(McKenzie, 1989) and precipitate LILE-enriched micro-
phases, carbonate melts such as those invoked for the In
Teria stage 2 metasomatism are probably the best candi-
dates to generate such enrichments. This may explain
why 87Sr enrichment is observed in xenolith suites that
have experienced carbonate-melt metasomatism (e.g. In
Teria and the Canary Islands; Neumann et al., 2002,
2015).
0.5124
0.5125
0.5126
0.5127
0.5128
0.5129
0.5130
0.5131
0.5132
0.7030 0.7035 0.7040 0.7045 0.7050 0.7055
0.5124
0.5125
0.5126
0.5127
0.5128
0.5129
0.5130
0.5131
0.5132
0.7030 0.7035 0.7040 0.7045 0.7050 0.7055
0.5124
0.5125
0.5126
0.5127
0.5128
0.5129
0.5130
0.5131
0.5132
0.7030 0.7035 0.7040 0.7045 0.7050 0.70550.5124
0.5125
0.5126
0.5127
0.5128
0.5129
0.5130
0.5131
0.5132
0.7030 0.7035 0.7040 0.7045 0.7050 0.7055
10
EMIEMI
EMI EMI
143 N
d/14
4 Nd
87Sr/86Sr 87Sr/86Sr
HIMU HIMU
HIMUHIMU 14
PM 18
20
12
10
2018
1412
17
Melilitite(15.22)
Sr/Nd ratio in meltSr/Nd ratio in protolith
143 N
d/14
4 Nd
Sr-Nd concentration in protolith
PM (19.9-1.25)
(d)(c)
(b)(a)
PMx0.5PMx0.33
PMx2PMx3
Melx0.33Melx0.5
Melx3Melx2
Melilitite(1202-79)
Sr-Nd concentration in melt
(15.92)
PM
PMPM(1(15
U
Melilitite(1 22)(15.22)
Mixing curve U
MU
Ahaggar
(alkali basalts)
Ahaggar
(tholeiites)
Ahaggar
(alkali basalts)
Ahaggar
(tholeiites)
Ahaggar
(alkali basalts)
Ahaggar
(tholeiites)
Ahaggar
(alkali basalts)
Ahaggar
(tholeiites)
Erem-2 Int-15-1
Fig. 16. Variation of 87Sr/86Sr vs 143Nd/144Nd for the studied in Teria melilitites, pyroxenites and peridotite xenoliths, and phlogopitemegacrysts, compared with results of one-dimensional percolation–diffusion modelling showing the effect of compositional vari-ations in Sr and Nd (see text). The short-dashed line represents a reference experiment calculated with Primitive Mantle (PM) Sr–Nd contents as the protolith (McDonough & Sun, 1995) and the average Sr–Nd contents of the studied melilitites for the infiltratedmelt. The long-dashed line represents a mixing curve calculated with the same Nd–Sr contents. (a–d) illustrate the effects of varyingthe following: (a) the Sr–Nd concentrations in the peridotite protolith, from 0�33� to 3�PM; (b) the Sr–Nd concentrations in themelt, from 0�33�to 3� the average Sr–Nd contents in the melilitites; (c) the Sr/Nd ratios in the peridotite protolith and (d) in themelt, from 10 to 20. The other parameter values are the same as in Fig. 15 and are given in Supplementary Data Table A5. The sym-bols for the In Teria samples and the data sources for the Ahaggar basalt fields and mantle EM1 and HIMU end-members are thesame as in Fig. 9.
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The high 87Sr/86Sr ratios recorded by the In Teria
xenoliths are most probably inherited from the pre-
metasomatic composition of the lithospheric mantle.
The highest values are comparable with the 87Sr/86Sr
ratios ascribed to the EM1 mantle end-member, a com-
position that is considered to represent the signature of
the lower continental lithosphere (Hawkesworth et al.,
1986; Mahoney et al., 1991; Milner & Le Roex, 1996).
(a)(b)
(c)
(d)
EMI
87Sr/86Sr
HIMU(b)
(a)
207 P
b/20
4 Pb
208 P
b/20
4 Pb
HIMU
DMM
DMM
EMI
Pb/Sr varies in protolith and melt
Ahaggar
(alkali
basalts)
Ahaggar
(tholeiites)
Libya
(peridotites)
Ahaggar
(Manzaz
peridotites) Libya
(alkali basalts)
Ahaggar
(Manzaz
peridotites)
Ahaggar
(alkali basalts) Ahaggar
(tholeiites)
Libya
(alkali basalts)
Libya
(peridotites)
37.5
38.0
38.5
39.0
39.5
40.0
0.702 0.703 0.704 0.705 0.706
15.45
15.50
15.55
15.60
15.65
15.70
15.75
15.80
15.85
0.702 0.703 0.704 0.705 0.706
(a)(b)
(c)
(d)
(e)(f)
(g)
Erem-2
Erem-2
Int-14-7
Fig. 17. Variation of 207Pb/204Pb and 208Pb/204Pb vs 87Sr/86Sr for the studied In Teria melilitites, pyroxenite and peridotite xenoliths,and phlogopite megacrysts, compared with results of the one-dimensional percolation–diffusion modelling showing the effect ofvarying the (Pb/Sr)melt/(Pb/Sr)peridotite ratios (see text). The model runs a–g in (a) were calculated with (Pb/Sr)melt/(Pb/Sr)peridotite val-ues in the range 1�4–0�6; the model runs a–d in (b) were obtained with a more restricted range of (Pb/Sr)melt/(Pb/Sr)peridotite values,from 1�1 to 0�7. The other parameter values are given in Supplementary Data Table A5. The symbols for the In Teria samples, andthe references for the fields of North African mantle xenoliths and Ahaggar Cenozoic volcanism, as well as for the mantle DMM,EM1 and HIMU end-members, are the same as in Fig. 9.
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The Ahaggar intra-plate tholeiites contain an EM1-like
mantle component, generally ascribed to the involve-
ment of lower lithosphere during their genesis (Maza
et al., 1998; Aıt-Hamou et al., 2000). Conversely, the
HIMU signature found in North African alkali basalts,
and in In Teria melilitites, pyroxenites and several spi-
nel peridotites (Fig. 9), is widely considered to represent
the signature of upwelling asthenosphere or a mantle
plume (e.g. Maza et al., 1998; Aıt-Hamou et al., 2000;
Beccaluva et al., 2008). In this respect, it is worth noting
that the three analysed phlogopite megacrysts plot be-
tween HIMU–DMM and EM1 compositions, within the
field of the Ahaggar tholeiites, hence providing further
evidence for the existence of an EM1 mantle signature
beneath In Teria (Fig. 9).
However, although enriched in radiogenic Sr, the
isotopic composition of the In Teria spinel peridotites
differs markedly from the end-member EM1 signature.
Their 143Nd/144Nd ratios are too high and their208Pb/204Pb ratios too low at a given 206Pb/204Pb value
(Fig. 9). Among the analysed phlogopites, one sample
shows Pb isotopic compositions consistent with a con-
tribution from EM2 (i.e. a high 208Pb/204Pb value and a
composition comparable with that of the Ahaggar tho-
leiites) whereas the two other phlogopites are isotopic-
ally comparable with the In Teria peridotites (Fig. 9c).
These features are reminiscent of previous observa-
tions of isotopic decoupling in mantle rocks that have
been ascribed to daughter element fractionation as a re-
sult of melt–rock interactions (Ionov et al., 2002b;
Bodinier et al., 2004; Le Roux et al., 2009). The ‘instant-
aneous’ effects of daughter element fractionation as a
result of melt–rock interactions (as opposed to the
‘time-integrated’ effects of parent–daughter element
fractionation) might account for a significant part of the
isotopic variability observed in mantle rocks. Porous
melt flow and related diffusional or reaction processes
may generate isotopic covariation trends between peri-
dotite and melt end-members that diverge markedly
from mixing lines.
Percolation modelWe used the one-dimensional percolation–diffusion
model of Vasseur et al. (1991), modified to include iso-
topic homogenization (Bodinier et al., 2004), to evaluate
a process involving percolation of melt through a litho-
spheric mantle column to generate the Nd–Sr–Pb iso-
topic compositions observed in the In Teria mantle
xenolith suite. We consider in the model a protolith with
an EM1 signature, even if the mantle was probably het-
erogeneous in composition, varying between DMM and
EM1. The model incorporates the effects of melt infiltra-
tion velocity, distance of percolation, critical distance of
isotopic homogenization along the column, chemical
diffusional exchange between melt and minerals,
and mineral grain sizes. It assumes instantaneous
solid–liquid equilibrium at the surface of mineral grains,
considered to be spherical, and chemical diffusion
within the grains. Isotopic equilibrium between melt
and minerals is governed by the mass-balance equation
of isotopic homogenization at the scale of critical vol-
umes. The critical volume of isotopic homogenization is
defined as the volume of percolated peridotite in which
melt and minerals have reached isotopic equilibriumduring the critical time of percolation (i.e. the time it
takes for the melt to reach the top of the column).
The parameters used for modelling are compiled in
Supplementary Data Table A5 and the results are
shown in Figs 15–17. The only parameters that are com-
mon to all model runs are the isotopic compositions of
the melt and the peridotite protolith. The melt isotopic
composition was fixed to the composition of In Teriamelilitites for Nd and Sr. For Pb, we took slightly higher
values in the range of the Ahaggar and other North
African alkali basalts (Allegre et al., 1980; Dupuy et al.,
1993; Beccaluva et al., 2008). For the protolith we took
the EM1 values of Eisele et al. (2002) for Nd and Pb, and
a slightly higher value for Sr, although still within the
range of literature values (e.g. Hofmann, 2002).
The results of models to evaluate the influence of
compositional parameters (daughter element contentsin peridotite and melt, and mineral/melt partition coeffi-
cients) on the 143Nd/144Nd vs 87Sr/86Sr covariation are
shown in Figs 15 and 16. We first explored parameter
values retrieved from the composition of the peridotite
xenoliths, their host lavas and clinopyroxene equilib-
rium melts (Fig. 15). For the peridotite protolith, we tried
three Sr–Nd compositions reflecting the compositional
range observed in the studied peridotites: (1) the lessmetasomatized spinel peridotite Erem-2 (run a); (2) the
garnet peridotite Int-14-7, metasomatized by alkaline
silicate melt but devoid of carbonate-melt imprint (runs
b–d); (3) the spinel peridotites (average composition),
metasomatized by carbonate melt (runs e–h). For the
melt, we used either the average composition of the
analysed melilitites (runs a–g) or the theoretical melt in
equilibrium with clinopyroxene from the spinel perido-tites (run h). Because the composition of the percolating
melt evolved along with metasomatism, from silicate
(stage 1) to carbonate melt (stage 2), we tried two differ-
ent sets of KD. We used the experimental values of Hart
& Dunn (1993) for minerals/silicate melt (runs a–c) and
minerals/clinopyroxene partitioning, and those of
Klemme et al. (1995) for clinopyroxene/carbonate melt
partitioning (runs d–h). Among the other parameters,
only the protolith modal composition and mineral grainsizes are slightly variable (Supplementary Data Table
A5). The modes vary from that of an amphibole-free spi-
nel lherzolite (runs a and b) to the amphibole-bearing,
average composition of the In Teria spinel peridotites
(runs f–h). Mineral grain size is constant and moderately
coarse in most of the runs (Cpx radius¼ 0�15 mm), ex-
cept for runs g and h, for which we used the average
grain size of the spinel peridotites, characterized byvery fine clinopyroxene grains (Cpx radius¼ 0�01 mm).
The other parameters (length of the percolation column,
porosity, melt velocity, critical distance of isotopic
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homogenization, and chemical diffusivity in minerals
and melt) are constant in all runs. Their effects were as-
sessed separately (Supplementary Data Fig. A6).
The results in Fig. 15 show a wide range of143Nd/144Nd vs 87Sr/86Sr covariation patterns, from
deeply concave to convex upward. All of these patternsdiffer markedly from the mixing curves between perido-
tite protoliths and melts, which are almost linear. The
shape of the covariation patterns and the degree of
decoupling are correlated with the difference in Sr/Nd
between melt and peridotite. Low Sr/Nd in the melt and/
or high Sr/Nd in the protolith [i.e. (Sr/Nd)melt/(Sr/
Nd)peridotite< 1] generate concave variation patterns
indicating that 87Sr/86Sr in the protolith is re-equili-brated with incoming melt at lower melt/rock ratio than143Nd/144Nd. This is well illustrated by the strongly con-
cave covariation patterns obtained in run h [(Sr/Nd)melt/
(Sr/Nd)peridotite¼ 0�805] and run a [(Sr/Nd)melt/(Sr/
Nd)peridotite¼ 0�285). Conversely, high Sr/Nd ratio in the
melt and/or low Sr/Nd ratio in the protolith [i.e. (Sr/
Nd)melt/(Sr/Nd)peridotite> 1] generate convex patterns
indicating that 143Nd/144Nd re-equilibrates at a lower
melt/rock ratio than 87Sr/86Sr. This pattern was obtainedwith the experiments b–d where (Sr/Nd)melt/(Sr/
Nd)peridotite¼ 1�072. This confirms that the Sr/Nd ratios
in the protolith and melt play a major role in the decou-
pling of Nd–Sr isotopic compositions during porous
melt flow (Ionov et al., 2002b; Le Roux et al., 2009).
Compared with the Sr/Nd values, the other parameters
that were explored through the numerical experiments
shown in Fig. 15 have only limited effects. Runs c and dillustrate the subtle effect of mineral/melt partition coef-
ficients as they were calculated with the same set of par-
ameters, except for distinct mineral/melt KD values (for
silicate- and carbonate-melt, respectively). Experiments
b and c illustrate the limited effect of adding 6% amphi-
bole to the peridotite mineralogy, whereas experiments
f and g show the effect of reducing the mean radius of
clinopyroxene grains from 0�15 to 0�01 mm.Despite their capability to decouple the 143Nd/144Nd
and 87Sr/86Sr ratios, the experiments shown in Fig. 15
do not fit the whole range of data for the In Teria xeno-
liths. The samples that are moderately enriched in 87Sr
(87Sr/86Sr<0�704) can be accounted for by the experi-
ments b–d (run with a relatively low Sr/Nd value for the
protolith, based on the composition of the garnet peri-
dotite). However, the other samples are not fitted by
these experiments and would require even lower Sr/Ndratio in the protolith, or a higher Sr/Nd value in the per-
colating melt. As a consequence, we ran further experi-
ments to explore the effect of compositional variations
in Sr and Nd (Fig. 16). For this approach, we first calcu-
lated a reference model (short-dashed lines) using the
Sr–Nd composition of Primitive Mantle (McDonough &
Sun, 1995) for the peridotite protolith and that of
the In Teria melilitites for melt [(Sr/Nd)melt/(Sr/Nd)peridotite¼ 0�956]. Then we simulated variations of
the Sr–Nd contents from 0�33 to three times the concen-
trations of the reference model (Fig. 16a and b), and
variations of the Sr/Nd ratios from 10 to 20 (Fig. 16c and
d), by varying the Nd contents (Supplementary Data
Table A5). The other parameters are the same as for ex-
periment g in Fig. 15, except for the minerals/melt parti-
tion coefficients for which we used the silicate-melt
values, instead of the carbonate-melt values.As shown by the similarity between the reference
model and the mixing line, these experiments confirm
that the decoupling of Nd and Sr isotopic compositions
is virtually non-existent when the peridotite protolith and
the incoming melt have similar Sr/Nd ratios, and this re-
sult remains true within a wide range of Sr and Nd abso-
lute concentrations (Fig. 16a and b). Conversely, the
results shown in Fig. 16b and c indicate that even moder-ate variations of the (Sr/Nd)melt/(Sr/Nd)peridotite ratio gen-
erate significant isotopic decoupling. Almost the whole
range of the Nd–Sr isotopic variations observed in the In
Teria mantle xenoliths can be accounted for by experi-
ments involving (Sr/Nd)melt/(Sr/Nd)peridotite values in the
range 1�15–1�30. This value may result from a sub-chon-
dritic Sr/Nd ratio (10–12) in the protolith or from a super-
chondritic ratio (�20) in the percolating melt. The latter
alternative is more likely, as mantle rocks rarely showsuch low Sr/Nd values, whereas elevated Sr/Nd values
can be found in carbonate melts (e.g. �58 in carbonatites
from Morocco; Wagner et al., 2003). Alternatively, the In
Teria Nd–Sr isotopic compositions can be explained by a
combination of moderately sub-chondritic Sr/Nd values
in the peridotite (e.g. �14�2 as in the garnet peridotite Int-
14-7) and moderately super-chondritic Sr/Nd values in
melt [e.g. �17�14, as in ocean island basalt (OIB);McDonough & Sun, 1995]. Moreover, the decoupling can
be amplified by a short distance of isotopic homogeniza-
tion, thus requiring lower (Sr/Nd)melt/(Sr/Nd)peridotite val-
ues. The influence of parameters on Nd–Sr decoupling is
discussed in the Supplementary Material (Fig. A6).
It is worth noting that the peculiar Sr–Nd isotopic sig-
nature of the In Teria phlogopite megacrysts can be ex-
plained by the same two-component porous-flow modelas for the spinel peridotites, but requires either lower (Sr/
Nd)melt/(Sr/Nd)peridotite values (Fig. 16c and d) or a longer
distance of isotopic homogenization (Supplementary
Data Fig. A6f), or a combination of these. This result
lends support to the hypothesis that the phlogopites and
the metasomatized spinel peridotites are related to a sin-
gle event of interaction between the lithospheric mantle
and sublithospheric melts, but the former record inter-
action at relatively high temperature involving silicatemelt, whereas the spinel peridotites record lower tem-
perature interaction with more evolved, carbonate melt
(Fig. 14). However, two spinel peridotite samples (Erem-2
and In-15-1) remain unexplained by the model, showing
elevated 143Nd/144Nd at a given 87Sr/86Sr value. The iso-
topic composition of these samples requires a compo-
nent characterized by a 143Nd/144Nd signature higher that
both the EM1 signature and the composition of theAhaggar basalts. Mantle xenoliths with Nd isotopic com-
positions recording long-term depletion in LREE have
been reported from Ahaggar and other North African
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localities (Beccaluva et al., 2007, 2008). In addition, the208Pb/204Pb compositions of several samples also point
to the involvement of a depleted component with a
DMM-like signature (Fig. 9c).
To evaluate how much of the Pb isotopic systematics
can be explained by the two-component porous-flow
model, and how much requires a third, DMM compo-
nent, we carried out experiments simulating covari-
ations of 207Pb/204Pb and 208Pb/204Pb vs 87Sr/86Sr (Fig.
17). We ran several experiments involving various com-
binations of Pb/Sr ratios in the protolith and melt
(Supplementary Data Table A5) by varying the Pb con-
tent in the protolith and melt (207Pb/204Pb vs 87Sr/86Sr,
Fig. 17a), or in the melt only (208Pb/204Pb vs 87Sr/86Sr,
Fig. 17b). The other parameters, including the Sr con-
tent in the protolith [¼ PM after McDonough & Sun
(1995)] and melt (¼ In Teria melilitite), are the same as
in the reference model in Fig. 16. The results for the207Pb/204Pb vs 87Sr/86Sr covariation (Fig. 17a) show that
the whole In Teria array can be accounted for by the
two-component model within a narrow range of (Pb/
Sr)melt/(Pb/Sr)peridotite values between 0�7 and 0�9 (runs
c–f). For Pb/Sr values in melt consistent with the In Teria
melilitites (Pb/Sr¼ 0�00371 6 0�00015) and OIB values
(0�0048; McDonough & Sun, 1995), this constrains the
Pb/Sr values of the peridotite protolith in the range
0�004–0�007, consistent with the lower range observed
in the In Teria spinel peridotites.
On the 208Pb/204Pb vs 87Sr/86Sr diagram (Fig. 17b),
however, although the majority of the spinel peridotites
may be accounted for by the two-component model
within the same range of (Pb/Sr)melt/(Pb/Sr)peridotite val-
ues as for 208Pb/204Pb (runs c and d), several samples
characterized by lower 208Pb/204Pb values than EM1 re-
quire a third component with a DMM signature. These
samples notably include the garnet peridotite, the
poorly metasomatized spinel peridotite Erem-2, four
other spinel peridotites and a phlogopite megacryst. As
the 208Pb/204Pb composition of these samples is distinct
from that of the volcanism in In Teria and in the whole
Ahaggar region, the third, low 208Pb/204Pb component
probably records isotopic heterogeneities in the mantle
protolith, rather than in the incoming melt. The litho-
spheric mantle beneath In Teria was probably not uni-
formly EM1-like before the onset of metasomatism; it
probably included a DMM peridotite component, as
well as some peridotites with elevated 143Nd/144Nd val-
ues recording long-term LREE depletion.
Implications for the origin and composition ofthe lithospheric mantle and for the nature oflithosphere–asthenosphere interactions beneaththe Ahaggar SwellThe In Teria mantle xenoliths differ from their counter-
parts from the Ahaggar Swell and the majority of North
African mantle xenoliths in three respects: (1) the pres-
ence of rare garnet peridotites; (2) an extensive carbon-
ate-melt metasomatic imprint; (3) an enriched 87Sr/86Sr
signature. The elevated 87Sr/86Sr compositions are con-
sidered to be inherited from an EM1 lithospheric proto-
lith providing the source of the metasomatic melt. The
existence of an EM1 signature in the lithospheric mantle
of the Tuareg Shield was first inferred from the Pb–Sr–
Nd isotopic systematics of the Ahaggar tholeiiticvolcanism (Allegre et al., 1981; Maza et al., 1998; Aıt-
Hamou et al., 2000).
In contrast to other xenolith localities from the
Ahaggar domain (Manzaz, Beccaluva et al., 2007;
Tahalgha, Kourim et al., 2014), In Teria is not situated
on the Tuareg Shield but stands on the western margin
of the ‘Saharan Metacraton’ (Abdelsalam et al., 2002,
and references herein) (Fig. 1a). The latter is consideredas a former Archaean to Paleoproterozoic craton that
was remobilized (‘decratonized’) during the
Neoproterozoic, but behaved as a relatively rigid block
before this time. Hence the possibility cannot be
excluded that the lithospheric mantle beneath In Teria
was originally different (thicker and possibly more en-
riched) than that sampled by the Ahaggar mantle xeno-
liths. However, the existence of an EM1 signature both
in the Ahaggar continental tholeiites and in the In Teriaxenoliths and megacrysts suggests that this signature
was widely distributed in the lithospheric mantle be-
neath the Tuareg Shield and adjacent areas. For
Beccaluva et al. (2007), the absence of this signature in
the Ahaggar (Manzaz) xenoliths is thus ‘perplexing’.
Those researchers interpreted the mantle beneath the
Ahaggar Swell as newly accreted material, owing to the
lateral displacement and replacement of old litho-spheric mantle by upwelling asthenosphere. They also
suggested that the In Teria xenoliths could ‘plausibly
represent the (EM1-metasomatized) older cratonic litho-
spheric mantle’.
In fact, the In Teria mantle xenoliths differ from true
cratonic mantle in several respects (e.g. their relatively
low Mg# values compared with kimberlite-hosted peri-
dotite xenoliths and their dominantly fertile, lherzolitecompositions) and are much more similar to mantle
xenoliths from younger lithospheric domains (Menzies,
1990; O’Reilly et al., 2001, and references herein). This
lends support to the hypothesis that the destabilization
of the Saharan Metacraton in the Neoproterozoic was
associated with extensive rejuvenation of the litho-
sphere, which occurred either regionally as a result of
its delamination and thermo-mechanical erosion after
thickening (Ashwal & Burke, 1989; Black & Liegeois,1993; Abdelsalam et al., 2011) or more locally along
mega-shear zones (Liegeois et al., 2003; Fezaa et al.,
2010). Despite their strong metasomatic imprint, the In
Teria xenoliths have nevertheless preserved some min-
eralogical and geochemical features that may be
ascribed to the original Pan-African lithosphere, prior to
the Cenozoic events. These include remnants of the
EM1 lithospheric isotopic signature, as predicted byBeccaluva et al. (2007). The EM1 signature was possibly
inherited from subduction processes that occurred dur-
ing the Pan-African orogeny in conjunction with the
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amalgamation of lithospheric terranes (Black et al.,
1994; Liegeois et al., 1994; Caby, 2003).
Apart from the probable existence of a widespread
lithospheric EM1 signature, the nature of the litho-
spheric mantle beneath the Tuareg Shield and adjacent
areas prior to the Cenozoic asthenosphere–lithosphereinteractions is otherwise virtually unknown. In particu-
lar, the lateral variations in lithospheric thickness and
composition that would be expected from the juxtapos-
ition of lithospheric blocks of different provenances and
ages (e.g. Liegeois et al., 2005) are poorly constrained.
In addition to the scarcity of data, the lack of informa-
tion also reflects the modifications that have affected
the lithospheric mantle during the Cenozoic. Thesemodifications were particularly severe in the central
part of the Ahaggar Swell where the lithosphere was
strongly modified by extensive interaction with plume
melts (Dautria & Lesquer, 1989) or, possibly, rejuven-
ated by removal of old lithospheric material by upwell-
ing asthenosphere (Beccaluva et al., 2007). The
modification of the lithosphere was less intense in the
In Teria district where melt–rock interactions culmi-
nated in the pervasive infiltration of small-volume car-bonate melts. The persistence of a relatively thick
thermal boundary layer beneath In Teria would have
facilitated the formation of carbonate melts by incipient
lithospheric melting during transient heating and/or
through the evolution of primary silicate melts via melt–
rock reactions (Wallace & Green, 1988; Dalton & Wood,
1993; Sweeney, 1994; Yaxley & Green, 1996).
In scenarios involving the impingement of a mantleplume head centred on the Ahaggar Swell, or any alter-
native involving upper mantle diapiric instabilities
(Davies & Bunge, 2006; Lustrino et al., 2007) or shallow
asthenospheric upwelling (Beccaluva et al., 2007), the
evolution recorded by the In Teria mantle xenolith suite
may simply reflect the attenuated effects of this process
owing to the distal location of In Teria, to the NE of the
Ahaggar Swell (Fig. 1). As noted by Aıt Hamou &Dautria (1994), this scheme is supported for the
Ahaggar Swell by the spatio-temporal evolution of the
volcanism, varying from volumetrically predominant
Eocene to Oligocene plateau tholeiites in the central
part of the Swell (Taharaq district) to smaller volumes
of Miocene to Quaternary undersaturated alkaline lavas
in the outer parts. The isotopic signature of the In Teria
melilitites is virtually indistinguishable from that of the
Ahaggar alkaline lavas (Fig. 9) and thus this recent vol-canic district might be viewed as the most distal mani-
festation of the asthenosphere–lithosphere interactions
related to the Ahaggar Swell. However, available heat
flow data (Lesquer et al., 1989, 1990; Takherist &
Lesquer, 1989) do not support this model. In Teria is sit-
uated on an east–west high heat flow axis that trends
from Libya towards the Canary Islands, between the
Sahara Basins to the north and the Ahaggar Swell andthe West African Craton to the south (Lesquer et al.,
1990). This anomaly is a prominent feature of the heat
flow pattern in northwestern Africa, showing its
maximum values (100–120 mW m–2) in southern
Algeria. As suggested by Lesquer et al. (1990), the In
Teria volcanism might be related to the mantle proc-
esses that are responsible for this anomaly. In this re-
spect, it may be worth noting that mantle xenoliths
comparable with those from In Teria—and also brought
to the surface by dominantly undersaturated lavas—are
found along the western extension of the anomaly, in
southern Morocco (Jbel Saghro, Ibhi et al., 2002) and in
the Canary Islands (e.g. Neumann et al., 2002). The simi-
larities notably include the presence of abundant reac-
tion aggregates containing microgranular secondary
phases and glass, and evidence for interaction with car-
bonate melts. Alkaline pyroxenites and wehrlites are
common, sometimes accompanied by carbonated peri-
dotites or carbonatites. The significance of the North
Saharan thermal anomaly is poorly understood, how-
ever. Edge-driven convection (Missenard & Cadoux,
2012) was recently proposed for southern Morocco,
where the western termination of the anomaly runs
along the northern border of the West African Craton. In
Algeria, however, this model is consistent with the ob-
servation that the anomaly cross-cuts at a high angle
the north–south lithospheric structures resulting from
terrane amalgamation and shear-zone activity during
the Pan-African orogeny (Black et al., 1994).
CONCLUSIONS
The In Teria xenolith suite records three different meta-
somatic imprints, which all postdate the last deform-
ation episode in the lithospheric mantle of this region.
Stage 1 is a diffuse infiltration of hydrous alkaline melts
affecting the spinel and garnet peridotites. Melt percola-
tion and the crystallization of the reaction products dur-
ing this episode are controlled by the pre-existing
deformation microstructure of the peridotite. This first
metasomatic stage was probably coupled with a heat-
ing of the deep subcontinental lithosphere by melts of
an asthenospheric origin. Stage 2 is a carbonate melt
metasomatic imprint restricted to the spinel peridotites,
which largely overprints stage 1. Finally, stage 3 corres-
ponds to the crystallization of pyroxenites, which are
probably crystal segregates from hydrous alkaline
melts crystallized in vein conduits. The three metasom-
atic imprints represent successive stages of interaction
between the lithosphere and sublithospheric melts.
The In Teria peridotites with enriched 87Sr/86Sr com-
positions, as well as phlogopite megacrysts, plotting
within the field of the Ahaggar tholeiites, show evidence
for the involvement of an EM1-like component, prob-
ably inherited from the lithospheric mantle beneath In
Teria. However, the isotopic composition of the In Teria
spinel peridotites differs from the end-member EM1 sig-
nature, which might be accounted for by isotopic
decoupling in mantle rocks. Assuming that the Ahaggar
alkali basalts and the In Teria melilitites represent melts
of essentially sublithospheric origin, numerical model-
ling simulating porous melt percolation through an
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EM1 lithospheric mantle can reproduce the high87Sr/86Sr values observed in some peridotite xenoliths.
These high 87Sr/86Sr values may be related to the crys-
tallization of interstitial micro-phases from small-vol-
ume melts and chromatographic effects. Conversely,
the HIMU signature observed in the In Teria melilitites,
pyroxenites and several spinel peridotites, as well as
North African alkali basalts, is widely considered to rep-
resent a mantle plume signature. Additionally, the208Pb/204Pb composition of several samples points to
the involvement of a depleted component with a DMM
signature. The lithospheric mantle beneath In Teria was
probably not uniformly EM1-like; it probably included a
DMM peridotite component as well as some peridotites
with elevated 143Nd/144Nd values recording long-term
LREE depletion.
The In Teria mantle xenoliths differ from North
African and Ahaggar Swell mantle xenoliths by the
presence of garnet peridotite, the extensive carbonate
melt metasomatic imprint, and their high 87Sr/86Sr
ratios. The last are considered to be inherited from an
EM1-like lithospheric mantle protolith. The similar EM1
signature in Ahaggar continental tholeiites suggests
that this signature extends over the entire Tuareg
Shield and adjacent areas and that the generation of the
tholeiites involved melting of the base of the litho-
sphere. The EM1 signature was possibly inherited from
subduction processes that occurred during the Pan-
African orogeny. In Teria mantle xenoliths therefore
have geochemical characteristics similar to those of
young lithospheric mantle rather than cratonic mantle
domains, supporting the hypothesis that the destabil-
ization of the Saharan Metacraton in the Neoproterozoic
was associated with extensive lithosphere rejuvenation.
ACKNOWLEDGEMENTS
C. Nevado and D. Delmas supplied high-quality pol-
ished thin sections for EBSD measurements. D.
Mainprice is thanked for providing software for analyz-
ing and plotting CPO data. We are also grateful to P.
Verdou for his help during thermal ionization mass
spectrometry analyses at the GIS laboratory (Nımes)
and to P. Telouk for his help during MC-ICP-MS ana-
lyses at the Service Commun National of the Ecole
Normale Superieure de Lyon. We thank editors M.
Wilson and D. Weiss for constructive comments and we
also gratefully acknowledge G. Bianchini, J. Berger and
an anonymous reviewer for providing constructive
reviews.
FUNDING
This study was performed as part of a collaborative
multi-disciplinary research project on the Hoggar region
chiefly involving the Faculte des Sciences de la Terre,
de Geographie et d’Amenagement du Territoire
(FSTGAT, USTHB, Algiers) and Geosciences
Montpellier, funded by the Centre National de la
Recherche Scientifique (CNRS, France: INSU and
DERCI), the Direction de la Post-Graduation Recherche-
Formation (DPGRF, Algeria), and the French Ministry of
Foreign Affairs (MAE), through a CNRS–DPGRF PICS
co-operation project (2008–2010), an MAE bilateral co-
operation project in the frame of the Tassili Hubert
Curien programme (2009–2012), and several INSU re-
search grants (SEDIT 2007, Actions Coordonnees 2008,
and SYSTER 2010–2011 programmes). M.-A.K.’s post-
doctoral research in Montpellier was funded by the
Institut National des Sciences de l’Univers–Centre
National de la Recherche Scientifique (INSU–CNRS,
France). M.A.K. also thanks the Australian Research
Council (Grant DP0878453) and the Swiss National
Science Foundation (SNSF) from Ambizione fund NSF-
26083906 for funding part of this research. The EBSD–
SEM national facility in Montpellier is supported by the
INSU–CNRS and by the Conseil Regional Languedoc–
Roussillon, France.
SUPPLEMENTARY DATA
Supplementary data for this paper are available at
Journal of Petrology online.
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