Miocene incorporation of peridotite into the Hercynian basement of the Maghrebides (Edough massif, NE Algeria): Implications for the geodynamic evolution of the Western Mediterranean

Chemical Geology 261 (2009) 172–184

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Chemical Geology j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / c h e m g e o

Miocene incorporation of peridotite into the Hercynian basement of the Maghrebides (Edough massif, NE Algeria): Implications for the geodynamic evolution of the Western Mediterranean O. Bruguier a,⁎, D. Hammor b, D. Bosch a, R. Caby a a b

Equipe Manteau-Noyau, Géosciences Montpellier, Université de Montpellier II, Place E. Bataillon, 34 095 Montpellier, France Université Badji-Mokhtar, BP12, El-Hadjar, Annaba 23 000, Algeria

a r t i c l e

i n f o

Article history: Accepted 13 November 2008 Keywords: Peridotite Western Mediterranean Monazite U–Pb geochronology Laser ablation

a b s t r a c t A laser ablation ICP-MS U–Pb age of 17.84 ± 0.12 Ma (late Burdigalian) was obtained from monazites separated from a leucocratic diatexite collected in close proximity to a small peridotite massif incorporated into the lower crustal sequence of the Edough Massif (north-eastern Algeria), a southern segment of the periMediterranean Alpine Belt. Monazites extracted from a neighbouring deformed leucogranite intruding early Paleozoic phyllites yield a consistent age of 17.4 ± 1.3 Ma. Zircons occurring in the same leucogranite, with magmatic characteristics, have an age of 308 ± 7 Ma interpreted as dating magmatic crystallisation of the leucogranite and reflecting partial melting during the Hercynian orogeny. Low Th/U domains (Th/U b 0.10) from the same grains substantiate recrystallisation during a younger metamorphic event whose upper age limit is 286 ± 11 Ma. These results emphasize the polycyclic evolution of basement rocks preserved in the crystalline units of the western Mediterranean and indicate that part of their metamorphic features were inherited from older, Hercynian, events. Taken together with published Ar–Ar dates, the late Burdigalian age of monazites indicates a rapid cooling rate of c. 370 °C/Ma and is regarded as closely approximating the emplacement of the peridotites into the Hercynian basement. The monazite ages are significantly younger than those recorded for orogenic peridotites from the Betic-Rif orocline and for the timing of lithospheric extension forming the Alboran sea. It is also younger than rifting and back-arc extension opening the Liguro–Provençal basin. The late Burdigalian age is interpreted as dating the incipient rifting event that opened the Algerian basin, which is consequently not a continuation of the Liguro–Provençal basin. At the scale of the western Mediterranean, these observations concur with current models supporting slab roll-back and an eastwards migration of extension in the western Mediterranean, but suggest that the Algerian basin opened as a result of torsion and stretching of the Thethyan slab due to its steepening under the Alboran microplate. © 2008 Elsevier B.V. All rights reserved.

1. Introduction Sizeable pieces of deep mantle material crop-out worldwide within orogenic belts involving major continental collisions (Brueckner and Medaris, 2000) or within regions with rift-thinned continental margins (e.g., Nicolas et al., 1987; Schärer et al., 1995). Although minor components of most metamorphic belts, the understanding of how and when orogenic peridotites were emplaced within the continental crust and their subsequent exhumation is paramount for the knowledge of the processes operating at the Earth crustmantle interface, in subduction as well as rift environments. The question of determining the age of the peridotites, and in particular their emplacement within their host rocks, is a challenging geochro-

⁎ Corresponding author. Fax: +33 4 67 14 47 85. E-mail address: [email protected] (O. Bruguier). 0009-2541/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2008.11.016

nological problem since many of the minerals commonly used as robust chronometers (e.g., zircon) do not typically occur in peridotite. Others, such as the Sm/Nd system, commonly give rise to ambiguous interpretations as they define trends which are variously regarded either as reliable dates or on the contrary as mixing lines between different mantle components (see discussion in Brueckner et al., 1996). Some chronometers, more adequate for peridotites (e.g., Re/Os), are often subject to disturbances and may yield minimum model ages (Snow and Schmidt, 1999), which in addition relate to phases of mantle differentiation, rather than to the crustal emplacement of the peridotites (Reisberg and Lorand, 1995). The Lu/Hf system has proved to be a valuable chronometer (Blichert-Toft et al., 1999), but requires fractionating phases or large inter-sample variations in order to get a spread of data to allow for a precise age determination. This restricts the studies aimed at tackling the timing of emplacement of orogenic peridotites and thus hampers any chronology and comparison between crustal processes and mantle dynamics and in particular

O. Bruguier et al. / Chemical Geology 261 (2009) 172–184

whether emplacement of mantle material, high-grade metamorphism and crustal anatexis were coupled phenomena. On the other hand, the emplacement of hot peridotite into the crust is often accompanied by contact metamorphism (Bosch and Bruguier 1998) or partial melting of the surrounding rocks (Schersten et al., 2000), which makes it possible to date emplacement of the peridotites into crustal units. This study presents laser ablation (LA-) ICP-MS U–Pb analyses of monazites and zircons, extracted from a diatexite migmatite and a leucogranite that crop out in the Edough massif (north-eastern Algeria), a window of the crystalline basement of the Maghrebides. The diatexite was sampled close to a small peridotite body, the so-called Sidi Mohamed peridotite (c. 300 × 150 m). The aim of this paper is three fold: 1) assess the chronology of peridotite incorporation within the crustal sequences of the Edough massif by dating anatexis related to its emplacement; 2) provide time constraints on basement components surrounding the peridotite; 3) consider the regional tectonic implications of these dates for reconstructing the geodynamic evolution of this area with an emphasis on the late Cenozoic kinematic evolution of the western Mediterranean basin. 2. Geological setting The Kabylies and the Edough massif (north-eastern Algeria) constitute the internal zones of the Maghrebides. This belt, that runs from Morocco to Algeria, represents the southern segment of the periMediterranean Alpine Belt (Fig. 1a) that resulted from the Cenozoic collision between Africa and Eurasia along with a set of microplates (Iberia and Apulia) trapped in between (see Frizon de Lamotte et al., 2000 and references therein). Although an important part of the West Mediterranean orogen, this area has been studied little, thus hampering orogen-scale tectonic correlations since it constitutes the southern flank of the belt and is located midway from its extremities, i.e., between the Betic-Rif and Calabria-Sicily. The dearth of data also impedes studies aimed at attempting reconstructions of the evolution of the western Mediterranean area and the development of extensional basins in a broadly N–S convergent regime (Michard et al., 2006). The Edough massif (Fig. 1b) constitutes the easternmost crystalline basement of the Maghrebides. Its shape is that of a broad asymmetric dome approximately 50 km long and 20 km wide (Fig. 1c) that structurally underlies a tectono-metamorphic pile of lower grade rocks including Early Paleozoic metasediments (“Alternance series”), allochtonous greenschist-facies Mesozoic sediments, and the Numidian flysch nappe. The granite-gneiss core of the massif is tectonically overlain to the north by an allochtonous unit consisting of garnet amphibolites and metagabbros with associated slices of peridotites known as the Amphibolite–Peridotite Unit (Caby et al., 2001) and derived from tholeiitic magmas emplaced in a continental setting (Ahmed-Said and Leake,1992). The granite–gneiss dome of the Edough massif comprises a package of granitoids, gneiss, diatexite migmatites and high-grade metasediments and also includes a lherzolite body with alternating dunite and pyroxenite bands, first described by Bossiere et al. (1988) and known as the Sidi Mohamed peridotite. This ultramafic body forms a c. 300 × 150 m outcrop and is enclosed in diatexites and sheeted granitoids with leucocratic bands. The eastern contact between the peridotite and the diatexite is outlined by a twelve meter thick layer of phlogopitedominated rock close to the MARID (Mica–Amphibole–Rutile–Ilmenite– Diopside) suite described by Dawson (1987) and by plagioclase-bearing pyroxenite containing sparse garnet (see Fig. 2). The plagioclase-bearing pyroxenite contains sodic augite that equilibrated at 750 °C and 1.2–1.4 Gpa. Both rock types (phlogopite-dominated rock and pyroxenite) are included in an agmatitic patchwork of plagioclase rich leucocratic pegmatoid veins (see Fig. 2). Published geochronological data for rocks of the massif are sparse. U–Pb zircon dating of gneissic units yielded ages around 600 Ma (Hammor and Lancelot, 1998) suggesting the occurrence of Pan-African related material in the massif. Metamorphic monazite from a paragneiss

173

yielded an age of 18 ± 5 Ma (Hammor and Lancelot, 1998), consistent with Ar/Ar mica ages at around 16–17 Ma from gneisses and micaschists (Monié et al., 1992). The Edough massif is regarded as a Miocene metamorphic core complex in which the thermal anomaly was provided by the upward tectonic emplacement of slices of hot mantle material and crystallization of ultramafic cumulates (Caby et al., 2001). The studied samples (see Figs. 1c and 2) consist of a leucocratic diatexite (Ed322) and a deformed leucogranite (Ed325). The leucocratic diatexite Ed322 was collected in the central part of the granite gneiss dome, along the road from Annaba to Seraïdi, at about 20 m from the Sidi Mohamed peridotite body. It consists of quartz, garnet, antiperthitic plagioclase, muscovite, minor biotite and accessory monazite. The occurrence of garnet suggests that partial melting took place by dehydration melting mineral reaction (Spear et al., 1999). The leucogranite Ed 325 was collected close to the city of Annaba in the southeast part of the dome. The sample was taken from a sheet of protomylonitic tourmaline-bearing leucogranite intruding phyllites and was affected by the same low- to medium-pressure metamorphism that can be observed in the surrounding metapelites (andalusite–staurolite–garnet with temperature around 550 °C, see Caby et al., 2001) away from the granite-gneiss core. The deformed leucogranite contains stretched globular quartz, euhedral plagioclase, K-feldspar porphyroclasts, muscovite, minor biotite and accessory zircon and apatite, frequently included in blue tourmaline clasts. Except for inner relict domains from K-feldspar clasts, the igneous mineralogy was thoroughly replaced by metamorphic phases and both micas are of metamorphic origin (Caby et al., 2001). 3. Analytical techniques Zircon and monazite were hand-picked in alcohol from the least magnetic concentrates (6° tilt and 1° tilt at full amperage for monazite and zircon respectively). Selected crystals were then embedded in epoxy resin, ground and polished to expose the internal structure. They were subsequently examined using back-scattered electron (BSE) imaging with a scanning electron microscope (SEM) at the University of Montpellier II. The sample mounts were later used for U–Th–Pb microanalyses using a Lambda Physik CompEx 102 excimer laser generating 15 ns duration pulses of radiation at a wavelength of 193 nm. For analyses, the laser was coupled to a VG Plasmaquad II ICPMS and analytical procedures followed those outlined in Bruguier et al. (2001) and given in earlier reports (Dhuime et al., 2007), they are only briefly summarized below. Ablation experiments were conducted under ultrapure He, which enhances sensitivity and reduces Pb–U fractionation (Günther and Heinrich, 1999). The He gas stream and particles from the samples were then mixed with Ar before entering the plasma. The laser was fired at an energy density of 15 J/cm2 at a frequency of 4 Hz for zircon and 2 Hz for monazite. During all experiments oxide level, measured using the ThO/Th ratio, was below 1%. The spot size of the laser beam was 26 or 51 µm for zircon and 26 µm for monazite analyses. Data were acquired in the peak jumping mode using 1 point per peak and each element was measured using an equal dwell time of 10.24 ms except for 207Pb which was measured during 40.96 ms. Unknowns were bracketed by measurements of the G91500 (Wiedenbeck et al., 1995) and Manangotry (Poitrasson et al., 2000) standards for zircon and monazite, respectively, where the ratio of unknown to standard was 5:4. Standard measurements were used for mass bias and inter-element fractionation corrections. For mass bias, all standard measurements performed during one session were averaged, whereas for Pb–U fractionation, only the 4 standards bracketing the five unknowns were used for each batch of analyses. The calculated bias factors and their associated errors were then added in quadrature to the individual errors measured on each unknown. In the course of this study, 16 analyses of the Managotry monazite were performed and yielded a 207Pb/206Pb weighted average of 0.05862 ± 0.00019 (2σ), which corresponds to an age of 553±7 Ma. This is in good agreement with the EPMA (557±20 Ma,

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Fig. 2. Schematic cross-section of the southern contact between the Sidi Mohamed peridotite and gneissic units of the Edough massif at sampling site of Ed322. Top field photograph is a view of sample Ed322, a leucocratic diatexite collected c. 20 m from the peridotite. The small black arrow indicates location of the collected sample. The lower field photograph is an outcrop view of the Sidi Mohamed peridotite (on the left), surrounded by pyroxenite (pyr) and a phlogopite-rich rock (phl). The two latter are enclosed in a network of anorthiterich pegmatoid veins.

Montel et al., 1996) or TIMS (554±4 Ma, Horstwood et al., 2003) reference ages. Reproducibility of the measured 206Pb/238U ratios was 1.2% (1σ). The zircon standard G91500 (Wiedenbeck et al., 1995) was analysed twenty-four times and gave a 207Pb/206Pb weighted average of 0.07490 ± 0.00018 (2σ), which corresponds to an age of 1066 ± 5 Ma. Reproducibility of the measured 206Pb/238U ratios was 1.4% (1σ). 4. Results 4.1. Leucocratic diatexite Ed322 Monazite extracted from this sample is translucent pale yellow, 100– 200 µm in size with a euhedral to sub-euhedral external morphology. Back-scattered electron imaging (Fig. 3a, b) shows that the crystals display broad zoning, which together with their euhedral shapes, is consistent with crystallisation from a melt. Twenty-eight laser ablation spot analyses have been performed on sixteen grains and all have indistinguishable 206Pb/238U ages from 16.7 to 19.3 Ma (see Table 1). On the Terra–Wasserburg diagram (Fig. 4) they plot along a line connecting the Stacey and Kramers common Pb composition (Stacey and Kramers, 1975) and an age of 17.84 ± 0.12 Ma (2σ, MSWD = 1.07, n = 28). 4.2. Leucogranite Ed325 Monazites from this sample are translucent pale yellow and have irregular shapes suggesting growth inhibition and sub-solidus nucleation (Fig. 3c). Four laser ablation spot analyses have been performed on three grains and they yield indistinguishable within error 206Pb/238U ages from 17.4 to 18.1 Ma (see Table 1). Pinned to the Stacey–Kramers common Pb composition (Fig. 4 inset), they define an age of 17.4± 1.3 Ma (2σ, MSWD =0.05, n =4), less precise, but consistent with the age of monazites from the leucocratic diatexite Ed322. Zircons from this sample are colourless, translucent and have subhedral to euhedral shapes. Back-scattered SEM imaging (Fig. 3d–f)

reveals an homogeneous internal structure with no or only very faint oscillatory zoning, suggesting the grains are magmatic in origin. Some grains in addition are characterized by a thin, BSE bright (high U) outer zone preserving the prismatic shape of the grains (Fig. 3d). Lastly, a few grains have rounded terminations and an ovoid morphology (Fig. 3g). They are interpreted as undissolved xenocrysts preserved in Zr saturated melt (Watson and Harrison, 1983). Grains with this morphology plot concordantly (Fig. 5) at c. 550 Ma (analyses #13), c. 1000 Ma (analyses #5 and #17) and at c. 2400 Ma (analysis #12) and reflect the ages of inherited source materials. The remaining analyses yield 206Pb/238U ages ranging from 260 Ma to 330 Ma. Spread of data along the concordia (Fig. 6) may be interpreted in terms of protracted zircon growth or differential Pb-losses from a single zircon population. Protracted zircon growth can be envisioned on the basis of thermobarometric modelling of melt-bearing systems (Kelsey et al., 2008), but would require that the leucogranitic magma did not cool significantly, nor crystallize completely during about 70 Ma. This would also require temperatures over 700 °C in the surrounding country rocks, which contrasts with their metamorphic gradient (c. 550 °C, see above). This explanation is thus considered unlikely. Following the hypothesis of differential Pblosses, all data have been combined to give a discordia line anchored at 17.4 ± 1.3 Ma and intersecting concordia at 299.9 ± 8.4 Ma (2σ, MSWD = 1.02, n = 30). At first sight, it is tempting to interpret this alignment by a simple model of U–Pb disturbances of magmatic zircons during the Miocene metamorphic overprint, the c. 300 Ma age dating crustal anatexis. U–Th–Pb analyses are consistent with this view since the youngest ages are associated with low Th/U ratios (Th/U b 0.2). These low Th/U domains correspond to internal structures where the primary zoning has been wiped out, whereas high Th/U ratios sometimes are associated with zircon domains preserving a faint oscillatory zoning. This is consistent with a solid-state recrystallisation process (Pidgeon, 1992), which in addition to blurred primary structures, is associated with expulsion of non-essential constituents and a lowering of the Th/U ratio of the protolith zircons (Pidgeon et al., 1998; Hoskin and Black,

Fig. 1. a) Tectonic sketch map of the Western Mediterranean basin showing the location of the Maghrebides within the framework of the West Mediterranean orogen (modified after Frizon de Lamotte et al., 2000). The square indicates the location of c; b) enlargement of the western part of the Mediterranean basin with location of the ultramafic complexes in the Alboran sea area. GK: Greater Kabylia Massif; PK: Lesser Kabylia Massif; EM: Edough Massif; (c) Schematic geological sketch map of the Edough massif showing location of the studied samples (★). Sampling site for sample Ed322 is shown in Fig. 2.

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Fig. 3. Scanning electron microscope (BSE) imaging of monazite and zircon from the studied samples. a) and b): monazite grains from leucocratic diatexite sample Ed322. The crystals have euhedral shapes and display broad oscillatory zoning consistent with crystallisation from a melt; c): irregular-shaped monazite grain from deformed leucogranite Ed325; d) to g): zircon grains from deformed leucogranite Ed325. The crystals have euhedral to sub-euhedral shapes and sometimes display a faint oscillatory zoning preserved in some part of the grains. When analysed, the domains with remnants of oscillatory zoning have older 206Pb/238U ages and higher Th/U ratios. The low Th/U ratios and younger ages of the structureless zircon domains are interpreted as reflecting a static recrystallisation process under metamorphic conditions; h): rounded anhedral zircon grain typical of xenocrysts preserved in the leucogranitic melt. Ages are quoted at the 1σ level and correspond to 206Pb/238U ages. Circles are the laser ablation analysis sites.

2000). In the Th vs U diagram (Fig. 7a) the low Th/U analyses define a field distinct from the high Th/U (magmatic) domains, and indicate that the recrystallisation was accompanied by Th depletion and U enrichment. These zircon domains are also characterized by high radiogenic Pb (Pb⁎) contents (see Table 1), which suggests that the U gain was ancient

enough to produce a significant accumulation of Pb⁎. A simple calculation of the Pb⁎ produced by the radioactive decay of U (see Fig. 7b) on the low Th/U recrystallized domains indicates that their Pb⁎ contents cannot be accounted for by a young recrystallisation event occurring at c. 18 Ma or that the recrystallisation process was largely

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Table 1 Laser ablation ICP-MS isotopic data for monazites and zircons from the investigated lithologies of the Edough Massif (NE Algeria). Sample

Pb⁎ (ppm)

U (ppm)

Th (ppm)

Th/

204

208

U

206

206

Pb/ Pbm

Pb⁎/ Pb⁎

207

Pb⁎/

206

Pb⁎

±

207

±

206

Pb⁎/

±

(1σ)

235

(1σ)

238

U

(1σ)

Pb⁎/ U

Rho

Apparent age (Ma) 206

Pb⁎/ U

± (1σ)

18.3 18.5 18.3 17.7 18.4 18.1 17.9 18.4 18.3 18.4 18.0 18.1 18.3 17.7 18.1 17.8 18.3 19.2 19.3 17.9 18.9 17.1 17.1 16.7 16.9 17.6 17.8 18.5

0.4 0.4 0.5 0.3 0.7 0.4 0.3 0.2 0.1 0.6 0.4 0.3 0.2 0.4 0.8 0.4 0.4 0.6 1.4 0.6 0.7 1.1 0.8 0.5 0.4 0.8 0.7 0.6

238

Leucocratic diatexite Ed322 [7°42'50.6q E; 36°55'31.2q N] Monazite #Mo1-1 – 4763 – – 0.00004 #Mo1-2 – 4495 – – 0.00004 #Mo2 – 6124 – – 0.00003 #Mo3 – 7123 – – 0.00003 #Mo4 – 4554 – – 0.00004 #Mo5-1 – 3817 – – 0.00012 #Mo5-2 – 4763 – – 0.00011 #Mo5-3 – 4193 – – 0.00012 #Mo6 – 4260 – – 0.00012 #Mo7-1 – 4185 – – 0.00011 #Mo7-2 – 3851 – – 0.00013 #Mo7-3 – 3818 – – 0.00014 #Mo7-4 – 4110 – – 0.00013 #Mo8 – 4685 – – 0.00012 #Mo9-1 – 4912 – – 0.00011 #Mo9-2 – 4448 – – 0.00013 #Mo9-3 – 4183 – – 0.00013 #Mo10 – 4045 – – 0.00013 #Mo11-1 – 4204 – – 0.00013 #Mo11-2 – 4103 – – 0.00014 #Mo12-1 – 3512 – – 0.00015 #Mo12-2 – 6205 – – 0.00009 #Mo13 – 6359 – – 0.00009 #Mo14-1 – 6972 – – 0.00009 #Mo14-2 – 4941 – – 0.00014 #Mo15 – 4421 – – 0.00015 #Mo16-1 – 4166 – – 0.00016 #Mo16-2 – 4303 – – 0.00015

– – – – – – – – – – – – – – – – – – – – – – – – – – – –

0.0783 0.0618 0.0519 0.0513 0.0696 0.0590 0.0599 0.0611 0.0631 0.0617 0.0596 0.0621 0.0656 0.0577 0.0697 0.0619 0.1302 0.0598 0.0853 0.0620 0.0557 0.0503 0.0510 0.0556 0.0551 0.0605 0.0638 0.0925

0.0093 0.0026 0.0026 0.0010 0.0092 0.0006 0.0009 0.0005 0.0017 0.0007 0.0014 0.0018 0.0118 0.0016 0.0209 0.0019 0.1066 0.0005 0.0347 0.0009 0.0015 0.0004 0.0024 0.0018 0.0008 0.0010 0.0013 0.0190

0.0307 0.0245 0.0203 0.0195 0.0274 0.0229 0.0229 0.0241 0.0247 0.0243 0.0229 0.0241 0.0257 0.0219 0.0271 0.0236 0.0511 0.0246 0.0352 0.0237 0.0225 0.0184 0.0187 0.0199 0.0199 0.0228 0.0243 0.0367

0.0037 0.0011 0.0012 0.0005 0.0038 0.0006 0.0006 0.0003 0.0007 0.0008 0.0008 0.0008 0.0046 0.0008 0.0082 0.0009 0.0418 0.0008 0.0145 0.0008 0.0010 0.0012 0.0012 0.0009 0.0005 0.0011 0.0011 0.0076

0.0028 0.0029 0.0028 0.0028 0.0029 0.0028 0.0028 0.0029 0.0028 0.0029 0.0028 0.0028 0.0028 0.0028 0.0028 0.0028 0.0028 0.0030 0.0030 0.0028 0.0029 0.0027 0.0027 0.0026 0.0026 0.0027 0.0028 0.0029

0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001 0.0002 0.0001 0.0001 0.0002 0.0001 0.0001 0.0001 0.0001 0.0001 0.0001

0.16 0.41 0.50 0.59 0.28 0.93 0.78 0.80 0.16 0.93 0.68 0.45 0.07 0.67 0.14 0.62 0.03 0.97 0.17 0.91 0.80 0.99 0.69 0.69 0.84 0.94 0.90 0.16

Leucogranite Ed325 [7°45'55.0q E, 36°54'20.8q N] Zircon #Zr1-1 , mag. 52 995 446 0.45 #Zr1-2 , mag. 15 272 89 0.33 #Zr2-1 , recr. 29 656 57 0.09 #Zr2-2 , mag. 20 408 102 0.25 #Zr3 , mag. 20 394 115 0.29 #Zr4 , mag. 21 419 190 0.45 #Zr5, inh. 44 232 156 0.67 #Zr6 , recr. 51 1106 86 0.08 #Zr7 , mag. 22 464 223 0.48 #Zr8 , mag. 9 166 93 0.56 #Zr9 , recr. 27 613 53 0.09 #Zr10, p. recr. 19 439 58 0.13 #Zr11-1, p. recr. 19 355 57 0.16 #Zr11-2, p. recr. 23 432 54 0.12 #Zr12, inh. 73 137 55 0.4 #Zr13-1, inh. 11 128 32 0.25 #Zr13-2, inh. 15 124 82 0.66 #Zr14, p. recr. 26 493 57 0.12 #Zr15, p. recr. 17 320 37 0.12 #Zr16, inh. 21 285 291 1.02 #Zr17, inh. 36 191 86 0.45 #Zr18 , mag. 15 325 65 0.20 #Zr19, p. recr. 18 401 42 0.10 #Zr20 , recr. 59 1306 98 0.08 #Zr21-1, inh. 33 105 69 0.65 #Zr21-2, inh. 37 862 44 0.05 #Zr22, p. recr. 18 415 58 0.14 #zr23-1 , recr. 24 634 54 0.08 #zr23-2 , recr. 57 1774 66 0.04 #Zr24-1, p. recr. 16 387 56 0.15 #Zr24-2 , mag. 10 206 118 0.57 #Zr25, p. recr. 24 520 59 0.11 #Zr26 , mag. 22 387 107 0.28 #Zr27 , recr. 51 1252 112 0.09 #Zr28, recr. 87 2257 66 0.03 #Zr29, inh. 12 81 27 0.33 #Zr30 , mag. 73 1435 975 0.68 #Zr31-1 , mag. 73 1360 834 0.61 #Zr31-2, p. recr. 25 587 78 0.13

0.12 0.10 0.03 0.09 0.09 0.14 0.19 0.04 0.15 0.18 0.03 0.04 0.09 0.04 0.12 0.11 0.23 0.06 0.04 0.31 0.19 0.06 0.04 0.03 0.16 0.02 0.04 0.03 0.01 0.04 0.18 0.03 0.06 0.02 0.06 0.13 0.17 0.17 0.04

0.0518 0.0528 0.0524 0.0529 0.0533 0.0520 0.0724 0.0514 0.0518 0.0521 0.0512 0.0528 0.0529 0.0525 0.1557 0.0587 0.0583 0.0523 0.0516 0.0626 0.0729 0.0534 0.0521 0.0516 0.1270 0.0574 0.0531 0.0526 0.0524 0.0528 0.0530 0.0528 0.0520 0.0516 0.0523 0.0720 0.0527 0.0527 0.0530

0.0004 0.0006 0.0007 0.0007 0.0005 0.0006 0.0007 0.0010 0.0007 0.0012 0.0005 0.0031 0.0006 0.0009 0.0007 0.0010 0.0007 0.0013 0.0017 0.0021 0.0009 0.0014 0.0002 0.0010 0.0012 0.0016 0.0013 0.0006 0.0005 0.0010 0.0004 0.0007 0.0004 0.0003 0.0004 0.0013 0.0003 0.0003 0.0005

0.3577 0.3772 0.3296 0.3595 0.3743 0.3442 1.6647 0.3188 0.3582 0.3420 0.3218 0.3237 0.3762 0.3779 10.2182 0.6458 0.7718 0.3637 0.3627 0.4902 1.8082 0.3555 0.3347 0.3263 5.0671 0.3979 0.3282 0.3262 0.3163 0.3265 0.3495 0.3708 0.3438 0.3112 0.2966 1.4561 0.3663 0.3816 0.3254

0.0299 0.0270 0.0240 0.0268 0.0369 0.0182 0.0615 0.0168 0.0397 0.0129 0.0139 0.0210 0.0157 0.0258 0.4154 0.0384 0.0777 0.0239 0.0263 0.0356 0.1372 0.0171 0.0161 0.0228 0.2428 0.0428 0.0299 0.0192 0.0246 0.0138 0.0070 0.0096 0.0326 0.0174 0.0097 0.0632 0.0182 0.0139 0.0094

0.0501 0.0518 0.0456 0.0493 0.0509 0.0480 0.1667 0.0450 0.0501 0.0476 0.0456 0.0444 0.0516 0.0523 0.4761 0.0798 0.0960 0.0504 0.0509 0.0568 0.1798 0.0483 0.0466 0.0459 0.2893 0.0503 0.0449 0.0450 0.0438 0.0449 0.0479 0.0510 0.0480 0.0437 0.0412 0.1468 0.0504 0.0525 0.0446

0.0042 0.0037 0.0033 0.0036 0.0050 0.0025 0.0059 0.0022 0.0055 0.0014 0.0019 0.0013 0.0021 0.0034 0.0192 0.0045 0.0096 0.0031 0.0033 0.0037 0.0135 0.0020 0.0022 0.0031 0.0136 0.0052 0.0039 0.0026 0.0034 0.0017 0.0009 0.0012 0.0045 0.0024 0.0013 0.0058 0.0025 0.0019 0.0012

0.99 0.99 0.98 0.99 0.99 0.98 0.96 0.92 0.99 0.80 0.98 0.44 0.97 0.97 0.99 0.96 0.99 0.93 0.89 0.89 0.99 0.85 0.99 0.96 0.98 0.96 0.96 0.98 0.99 0.90 0.94 0.87 0.99 0.99 0.97 0.91 0.99 0.98 0.94

0.00001 0.00003 0.00002 0.00002 0.00003 0.00003 0.00001 0.00001 0.00004 0.00010 0.00002 0.00003 0.00003 0.00002 0.00001 0.00005 0.00004 0.00002 0.00003 0.00003 0.00002 0.00004 0.00004 0.00001 0.00002 0.00002 0.00004 0.00004 0.00002 0.00006 0.00010 0.00004 0.00004 0.00001 0.00003 0.00005 0.00002 0.00002 0.00003

315 326 288 310 320 302 994 284 315 300 287 280 325 328 2510 495 591 317 320 356 1066 304 294 289 1638 316 283 284 276 283 301 320 302 276 260 883 317 330 281

26 22 20 22 31 15 33 14 34 9 12 8 13 21 83 27 56 19 20 22 73 12 14 19 68 32 24 16 21 11 5 7 28 15 8 32 15 11 7

207

206

Pb⁎/ Pb⁎

± (1σ)

1155 666 280 254 918 566 599 643 711 665 589 677 792 520 919 672 2101 598 1323 674 442 207 241 438 417 620 736 1478

237 90 117 46 272 21 33 17 58 25 53 61 379 60 616 66 1437 17 787 30 59 20 107 72 33 36 43 388

275 319 303 326 343 284 998 258 277 291 250 322 322 305 2409 556 541 299 270 695 1012 347 289 268 2057 508 331 312 303 320 327 318 285 270 297 985 315 315 327

20 24 32 29 21 24 20 46 31 51 21 127 24 40 8 37 26 55 75 70 24 57 9 45 16 62 55 27 21 41 16 28 19 15 17 36 15 15 22

(continued on next page) page) (continued

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Table 1 (continued) Sample

Pb⁎ (ppm)

U (ppm)

Th (ppm)

Th/

204

208

U

206

206

Pb/ Pbm

Pb⁎/ Pb⁎

207

Pb⁎/

206

Pb⁎

±

207

±

206

Pb⁎/

±

(1σ)

235

(1σ)

238

U

(1σ)

Pb⁎/ U

Rho

Apparent age (Ma) 206

Pb⁎/ U

± (1σ)

17.4 18.1 17.4 17.5

1.3 1.3 1.3 1.5

238

Monazite #Mo1-1 #Mo1-2 #Mo2 #Mo3

– – – –

2161 1896 1443 2309

– – – –

– – – –

0.00008 0.00014 0.00018 0.00007

– – – –

0.0534 0.0578 0.0555 0.0500

0.0043 0.0018 0.0071 0.0069

0.0199 0.0224 0.0207 0.0187

0.0022 0.0018 0.0031 0.0030

0.0027 0.0028 0.0027 0.0027

0.0002 0.0002 0.0002 0.0002

0.68 0.92 0.50 0.52

207

Pb⁎/ Pb⁎

206

346 521 433 196

± (1σ) 181 68 287 323

: analyses used in the age calculation of the magmatic event (see text); : analyses used in the age calculation of the recrystallisation event (see text); mag.: magmatic domain; p. recr.: partly recrystallised domain; recr.: recrystallized domain; inh.: inherited grains. For monazites, Th and 208Pb were not measured as this resulted in a high signal and tripping of the detector. The asterisk (⁎) indicates radiogenic Pb. Rho is the error correlation between the 206Pb/238U and 207Pb/235U ratio. GPS coordinates have been estimated using Google Earth™.

inefficient, leaving most of the radiogenic Pb in the crystal lattice, while at the same time significantly lowering the Th content. This is clearly at odds with diffusion of Pb and Th in the zircon lattice (Lee et al., 1997; Cherniak and Watson, 2001). Accumulation of Pb⁎ thus occurred after the Th depletion and implies that zircon recrystallisation occurred during an event older than the Miocene. Pb loss enhanced by radiation damage to the zircon lattice (Silver and Deustch, 1963) and annealing during a thermal pulse can also be ruled out to explain the spread of the analyses as the correlation between apparent 206Pb/238U age and U content (Fig. 7c) produces a sub-horizontal trend. Although the low Th/U analyses may be consistent with a Pb loss model by radiation damage (Trend 1), they also fit a U gain model (Trend 2). Moreover, analyses with high-U contents, above the theoretical metamictization threshold (Williams, 1992), do not show a significant lowering of the 206Pb/238U ages, except for analysis #28, which yields the highest U content (2260 ppm) and the youngest age (260±8 Ma). Pb loss controlled by radiation damage can be considered for this analysis, but is not obvious for others. The subset of data with low Th/U ratios (b0.2) has a large spread in age (from 260 to 320 Ma) interpreted as an incomplete resetting of the U–Pb system during the recrystallisation process. Hence the analyses with the youngest ages and the lowest Th/U domains reflect a more pronounced recrystallisation on which to base our best age approximation. The weighted mean of these analyses is 286 ± 11 Ma (2σ,

MSWD = 0.15, n = 7) and is a maximum estimate of the age of recrystallisation of magmatic zircons (Fig. 6c). The subset of data with high Th/U ratios (excluding inherited grains) yield consistently older apparent ages and a mean 206Pb/238U age of 308 ± 7 Ma (2σ, MSWD = 0.74, n = 11) which is interpreted as the age of crystallisation of the zircons (Fig. 6b) and thus that of crustal anatexis from which the leucogranitic magma formed. 5. Discussion The results presented above define the temporal evolution of the uplifted lower continental crust exposed in this part of the Maghrebides, allowing a comparison with other crustal sections preserved in the periMediterranean realm and thus provide important constraints on the Cenozoic evolution of the western Mediterranean. 5.1. Hercynian evolution As outlined above, the geological significance of the U–Th–Pb behaviour of zircons from the leucogranite is paramount in evaluating the various events that have affected this crustal section during the Hercynian orogeny. This is particularly important since there is still no clear consensus on the age of the HP–HT metamorphic event recognized

Fig. 4. Terra–Wasserburg concordia diagram for laser-ablation ICP-MS analyses of monazites from leucocratic diatexite Ed322. The inset shows the results of monazite analyses from leucogranite Ed325. Data have been anchored to a common Pb composition of 0.837 ± 0.015 as given by the model of Stacey and Kramers (1975) at 18 Ma. Error crosses are 1σ.

O. Bruguier et al. / Chemical Geology 261 (2009) 172–184

179

Fig. 5. Concordia diagram displaying laser-ablation ICP-MS analyses of inherited grains from leucogranite Ed325. Quoted values are 207Pb/206Pb ages. Error ellipses and ages are 1σ.

in basement rocks preserved in the western Mediterranean orogen either in the Kabylies (e.g. Peucat et al., 1996) or in the Betics (e.g. Zeck and Whitehouse, 2002). Since many grains dated in this study have

magmatic characteristics, the late Carboniferous zircon age of 308 ±7 Ma is attributed to the age of emplacement of the granitic magma and thus to that of crustal anatexis at depth. This age is in good agreement with

Fig. 6. a) Concordia diagram displaying laser-ablation ICP-MS analyses of zircons from leucogranite Ed325. Error ellipses are 1σ; b) and c) Terra–Wasserburg diagrams for the high Th/ U and low Th/U domains analysed. Zircon domains with intermediate Th/U ratios (0.1 b Th/U b 0.2) not shown. Data have been anchored to a common Pb composition of 0.856 ± 0.015 as given by the model of Stacey and Kramers (1975) at 300 Ma. Error crosses are 1σ.

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O. Bruguier et al. / Chemical Geology 261 (2009) 172–184

Fig. 7. a) Th and U concentrations of the zircon domains analyzed. White circles: magmatic domains (Th/U N 0.2); grey circles: partly recrystallised domains (0.1 b Th/U b 0.2); black circles: recrystallised domains (Th/U b 0.1). MMD is the mean of the analysed magmatic domains (Th = 280 ppm and U = 606 ppm); b) Radiogenic lead (Pb⁎) and U concentrations diagram showing the Pb⁎ content of the recrystallised and partly recrystallised domains, either measured (□, ×), or calculated back from the measured U content at 286 Ma (□, ×) and at 18 Ma (■, +). The good agreement between the values calculated at 286 Ma and the measured values on one hand and the discrepancy with values calculated assuming a recrystallisation process at 18 Ma on the other hand is evidence for an old recrystallisation-inducing metamorphic event; c) 206Pb/238U ages and U concentration of magmatic and metamorphic domains for zircons from sample Ed325. The dashed line is the theoretical metamictization threshold given by Williams (1992).

the occurrence of Achritarchs described in Ordovician to Devonian black cherty layers from the “Alternance series” (Ilavsky and Snopkova, 1987), which are intruded by the dated sample. In addition, it compares well with ages from late- to post-tectonic granitoids (290–315 Ma) from other segments of the Variscan orogen. In the French Massif Central, South Bohemian Massif and Iberian Massif, leucogranite emplacement and crustal melting occurred during this time interval (FernandezSuarez et al., 2000; Gerdes et al., 2000; Bruguier et al., 2003) and are also well documented within the intra-Alpine massifs (Vavra et al., 1999). It is widely accepted that this period reflects delamination of the lithospheric mantle (Pin and Duthou, 1990; Schaltegger, 1997). Conversely, the late Carboniferous age of 308 ± 7 Ma is at odds with the c. 606 Ma U– Pb zircon age from a similar sample, which was interpreted as reflecting the imprint of the Pan-African orogeny in this area (Hammor and Lancelot, 1998). However, the data presented by Hammor and Lancelot (1998) are strongly discordant and scattered along a calculated discordia. More likely they reflect various degree of inheritance in the multigrain fractions analysed and a combination of Pb loss. In addition, the Pan-African age contrasts with geological evidence attributing an Ordovician to Devonian age for the whole metasedimentary pile intruded by these leucogranites (Ilavsky and Snopkova, 1987). This casts serious doubts on the influence of the Pan-African event as an orogeny in this part of the Kabylies, although inherited grains with ages around 550 Ma (this study) are evidence for an Early Cambrian

component. However, these grains are best interpreted as having been stripped off from deep-seated material, either of detrital origin (as suggested by their rounded shapes) or related to the Cambro– Ordovician rifting event that affected the northern margin of Gondwana (Nance and Murphy,1994). In the light of the limited data, it is difficult to sustain that the Pan-African event played an important role in the tectonometamorphic architecture of the Edough massif. The U–Pb system of the studied zircons was subsequently disturbed and partly rejuvenated by recrystallisation processes. From the present data set, it appears that this event occurred during the Permian and is dated at 286 ± 11 Ma on the most recrystallised domains. Although this is taken as the upper age limit for the recrystallisation-inducing metamorphic event, it is intriguing to note that this age is identical to that of the granulite facies metamorphism in the Ivrea–Verbano zone (Pin, 1986) and similar to the c. 280 Ma age peak of zircons from lower crustal xenoliths from the French Massif Central and from the granulitic metasedimentary basement from Corsica (Rossi et al., 2006). Highgrade metamorphism and magmatism at around 275–285 Ma has also been recognized further west in the Kabylies (Peucat et al., 1996; Hammor et al., 2006). In the Central and West Alpine Belt (African paleomargin of the Tethys), and in Corsica (Ligurian branch of the Tethys Ocean) the 280–290 Ma early Permian time interval was also characterized by underplating of mafic-ultramafic

O. Bruguier et al. / Chemical Geology 261 (2009) 172–184

complexes at middle and lower crustal levels (Hansmann et al., 2001; Paquette et al., 2003; Hermann and Rubatto, 2003; Tribuzio et al., 1999). In the Betic Cordilleras, Zeck and Whitehouse (1999) provided an age of 285 ± 5 Ma for the anatectic climax of the metamorphic event and monazite inclusions in garnet from the Beni–Bousera kinzigites of Morocco yielded an age of 284 ± 27 Ma (Montel et al., 2000). The coincidence in age between these widely distributed crustal slices preserved in the Alpine–Apennine and west Mediterranean mountain belts supports an orogen-wide event during the early Permian. It is also consistent with all these fragments being part of the same Hercynian basement section that shared a common history before the Permian rifting that preceded opening of the Tethys Ocean (Stampfli et al., 2001). The recognition of a Hercynian metamorphic event indicates that care must be taken when reconstructing the geodynamic evolution of the Alpine belts of North Africa, since part of the metamorphic features preserved in the exposed lower crust have been inherited from older orogenic cycles and should not be de facto attributed to the Alpine evolution. 5.2. Cenozoic evolution The analyses of metamorphic monazites from the leucogranite Ed325 and of magmatic monazites from the leucocratic diatexite Ed322 yield similar ages of 17.4–17.8 Ma. Both rocks have been sampled in the Edough dome, the latter in close proximity to the Sidi Mohamed peridotite body. Since the deformation and thermal overprint increase towards the contact with the peridotite (Caby et al., 2001) it is concluded that these ages are related to the emplacement of hot mantle material into the lower crustal section of the Edough massif, which was responsible for a heat supply generating a thermal aureole and local anatexis. The leucocratic diatexite Ed322, located about 20 m from the contact with the peridotites, was generated by partial melting of the

181

surrounding gneiss and migmatites. In the deformed leucogranite Ed325, temperature conditions deduced from the staurolite–andalusite assemblages from metapelites near Annaba (c. 550 °C after Caby et al., 2001) account for the metamorphic growth of monazite, but were not associated with a new zircon growth episode, recrystallisation of preexisting zircons or Pb diffusion. The temperature conditions experienced by the leucogranite were too low to be responsible for Pb diffusion in zircon (Lee et al., 1997; Cherniak and Watson, 2001). This may explain the lack of “zircon response” of the studied sample to the emplacement of the peridotite. Moreover recrystallised zircons tend to be more resistant to subsequent disturbances and more stable, since they have already expelled contaminant elements, and this is likely to be the case for the investigated zircons. Lastly, except for zircon, the investigated leucogranite contains very few Zr-bearing phases succeptible to breakdown and to liberate Zr (just biotite, see Fraser et al.,1997) available for a new growth episode. Metamorphic temperature conditions prevailing during emplacement of the peridotites (750 °C and 1.2–1.4 GPa after Caby et al., 2001) are within the classically accepted nominal closure temperature for Pb diffusion in monazite (≥725 ± 25 °C after Copeland et al., 1988) but preservation of radiogenic lead under higher metamorphic conditions clearly calls for values in excess of 800 °C (Bosch et al., 2002; Kelsey et al., 2003; Cherniak et al., 2004) or even higher (N850 °C) when crystals are shielded by host minerals (e.g. Montel et al., 2000). Combined with published 40Ar/39Ar data from rocks of the Edough massif (Monié et al., 1992), and assuming a closure temperature corresponding to peak metamorphic conditions (750 °C), the late Burdigalian age of monazite from the leucocratic diatexite indicates a fast cooling (and uplift rates) of 369 °C/Ma (Fig. 8). Such high cooling rates are known in the Betic Cordilleras (Zeck et al., 1992; Monié et al., 1994; Platt and Witehouse, 1999) and are typical of large-scale extensional tectonics (Platt and Vissers, 1989). They indicate rapid upward tectonic transport of the peridotite and surrounding parts of the crust through tens of kilometres. Given this rapid cooling rate, and the

Fig. 8. Cooling path for rocks of the Edough massif. Ar/Ar mineral ages have been calculated using the Ludwig program (Ludwig, 2000) as a weighted mean of all ages reported in Monié et al. (1992). This yields ages of 17.20 ± 0.20 Ma (n = 1); 16.35 ± 0.94 Ma (MSWD = 6.5; n = 3) and 16.70 ± 1.30 (MSWD = 17; n = 4) Ma for phlogopite, muscovite and biotite respectively (95% confidence interval). The corresponding closure temperatures are 460 ± 50 °C (Giletti and Tullis, 1977); 445 ± 50 °C (Hames and Bowring, 1994) and 360 ± 50 °C (Harrison et al., 1985). For monazite, the closure temperature has been taken at 750 °C, i.e., the estimated peak metamorphic conditions. Error crosses are 1σ.

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O. Bruguier et al. / Chemical Geology 261 (2009) 172–184

assumed similarity between the monazite closure temperature and the metamorphic conditions experienced by the country rocks, the late Burdigalian age is regarded as closely approximating emplacement of the peridotite within the Hercynian basement. Existing geochronological information from other occurrences of orogenic peridotites exposed in the western Mediterranean (see Fig. 9) yield early Miocene ages in the range 20–25 Ma, slightly, but consistently, older than the age of emplacement of the Edough peridotite. In the Betic Cordilleras, emplacement of the Ronda peridotite (see Fig. 1) is dated at 21.5± 1.8 Ma (Zindler et al., 1983) and was coeval with high-T metamorphism as dated by the U–Pb method on metamorphic rims of zircons from gneisses at between 19.3± 0.3 and 21.2± 0.7 Ma (Platt and Whitehouse, 1999) and anatexis of surrounding rocks at 21.8 ± 0.5 Ma (Esteban et al., 2007). The same scheme applies also for the Sebtides, on the south side of the Straits of Gibraltar, where emplacement of the Beni–Bousera peridotite has been dated at 25 ± 1 Ma by the Lu–Hf method on garnet pyroxenite embedded in the peridotite. This age is slightly older, but comparable to the 22.7 ± 0.3 Ma SHRIMP age of zircon rim from granulite-facies rocks (Platt et al., 2003). Fast cooling rate following emplacement is implied by a 40Ar/39Ar biotite age of 22.5 ± 0.5 Ma from a granulite sample around the Beni–Bousera peridotite (Michard et al., 2006). These 22–25 Ma ages are indistinguishable from those of retrograde monazite from low-T/high-P schists from Beni–Mezala, 100 km north of Beni–Bousera, that gave ages of about 21 Ma (Janots et al., 2006). Thus, we propose that emplacement of the Edough peridotite is younger by about 2–8 Ma. 5.3. Geodynamic implications In the Betic Cordilleras, emplacement of the peridotites into the crustal sections has been ascribed to slab detachment following subduction of the Mesozoic Tethys lithosphere (Platt and Vissers, 1989; Zeck, 1996; Platt and Whitehouse, 1999), which was responsible for fast uplift and a centrifugal displacement of mantle material, radiating from the central part of the Alboran basin at about 20– 25 Ma (Zeck, 1997). Crustal thinning in the Valencia trough (Maillard and Mauffret, 1999), clockwise rotation of the Corsica–Sardinia block (Doglioni et al., 1997) and the onset of opening of the Liguro–Provençal

basin as a back-arc basin (Speranza et al., 2002) occurred at the same period. From their original position, i.e., along the eastern Iberian margin and north of the Betic-Rif (see Carminatti et al., 1998; Jolivet and Faccenna, 2000), crustal fragments including the Alboran block and the Kabylies drifted southwestward and southeastward, respectively, during this rifting event. In Greater Kabylia, 40Ar/39Ar mineral ages and biotite Rb–Sr cooling ages in the range 20–25 Ma (Monié et al., 1988; Peucat et al., 1996; Hammor et al., 2006) indicate that basement rocks from this part of the Kabylies were exhumed to mid-crustal levels by the late Oligocene–early Miocene. During this period, rifting strongly thinned the crust but, except in the Liguro–Provençal basin, did not evolved to the development of oceanic crust. The Burdigalian ages of monazites from the Edough massif are slightly but significantly younger, and together with 40Ar/39Ar ages (Monié et al., 1992), substantiate that exhumation of deep crustal rocks in this area was delayed by comparison with basement rocks from the Greater Kabylia and from the Betic-Rif orocline. Incorporation of the peridotites into the lower crustal sections of the Edough massif and their fast rate of exhumation are thus attributed to a second extensional event of late Burdigalian age. This event coincides with opening of the Algerian basin (Mauffret et al., 2004) and is matched by compressional tectonics (south-vergent thrusting and folding) in the Rif and Tell orogens (Frizon de Lamotte et al., 2000). Since the oldest oceanic crust in the Algerian basin is c. 16 Ma old (Mauffret et al., 2004), we propose that incorporation of the Sidi Mohamed peridotites into continental crustal material occurred during the rifting phase and incipient continental break-up. In this view, the late Burdigalian monazite ages from the present study are evidence that opening of the Algerian basin was not coeval with opening of the Liguro– Provençal basin, thus suggesting a North–South diachronism for extension similar to the Red Sea area, which displays a northward younging of spreading events (Nicolas et al., 1987). However, since the rifting phase in the Algerian basin began when spreading finished in the Liguro–Provençal basin (Rollet et al., 2002), there is no continuum between the two areas and we conclude that the Algerian basin is not a continuation of the Liguro–Provençal basin. This is consistent with the synthesis of data from seismic surveys by Mauffret et al. (2004). At the scale of the West Mediterranean basin, the Burdigalian ages fit a general model of eastward younging of extensional events and widening of the

Fig. 9. Summary of available geochronological data for emplacement of orogenic peridotites and for major geodynamic events in the Western Mediterranean. References to geochronological data are as follow: (1) Zindler et al.,1983; (2) and (3) Platt and Whitehouse,1999; (4) Esteban et al., 2007; (5) Blichert-Toft et al.,1999; (6) Platt et al., 2003; (7) Michard et al., 2006; (8) and (9) Janots et al., 2006; (10), (11) and (12) Monié et al., 1988. The dashed line corresponds to the age of the oldest oceanic crust known in the Algerian basin (after Mauffret et al., 2004). See text for references to the main geodynamic events and for explanation.

O. Bruguier et al. / Chemical Geology 261 (2009) 172–184

so-called AlKaPeCa domain (Alboran–Kabylia–Peloritan–Calabria of Bouillin, 1986) related to slab roll-back and back-arc extension. Although back-arc extension could be episodic (Schellart, 2005), the time-lag between opening of the Liguro–Provençal basin and opening of the Algerian basin is puzzling and difficult to reconcile with a simple model. It suggests that other processes were responsible for the Burdigalian extensional phase. Duggen et al. (2004) proposed that the soutwestward migration of the Alboran microplate and its collision with the North African and Iberian margins in the Early Miocene was responsible for steepening of the Tethyan slab under this area. Steepening of the Tethys lithosphere overidden by the Alboran microplate had two direct consequences. Firstly, it induced a torsion of the slab and a traction, which could have slowed down slab retreat and extension further east in the Liguro–Provençal basin. Secondly, it resulted in upwelling of asthenospheric material in the void left by the steepened slab and a significant elevation of the thermal regime as recognized in the Alboran block (Duggen et al., 2004). Upwelling of asthenospheric material may also have been responsible for emplacement of the peridotites and their fast exhumation at about 18 Ma during the incipient rifting phase affecting the already thinned northern margin of Africa. Continental break-up was followed at about 16 Ma by spreading in the Algerian basin (Mauffret et al., 2004), which matches the end of extension in the Liguro–Provençal basin. 6. Conclusions The incorporation of peridotites into the lower crustal units of the Edough massif was responsible for melting and metamorphism of the surrounding Hercynian crust (286–308 Ma) and is dated at 17.84 ± 0.12 Ma. This event is associated with very fast cooling (in excess of 360 °C/Ma) and we propose it occurred during the incipient rifting phase of the opening of the Algerian basin. It marks a step in the Cenozoic evolution of the Western Mediterranean basin and provides a link between the late Oligocene-early Miocene extension in the western part (Alboran basin, Valencia trough and Liguro–Provençal basin) and the upper Miocene extension in the eastern part (Thyrrenian basin). Acknowledgment The authors thank C. Grill for SEM imaging, and C. Nevado and D. Delmas for careful polishing of the zircon and monazite mounts. We (D. B. and O.B.) would like to use the opportunity of this Special Issue in honour of Bob Pidgeon to warmly thank him for all he learned us (geochronology and field geology) and for the great time we had together during our stay at Curtin University of Technology from 1993 to 1994. References Ahmed-Said, Y., Leake, B.E., 1992. The composition and origin of Kef Lakhal amphibolites and associated amphibolite and olivine-rich enclaves, Edough, Annaba, NE Algeria. Mineralogical Magazine 56, 459–468. Blichert-Toft, J., Albarede, F., Kornprobst, J., 1999. Lu–Hf isotope systematics of garnet pyroxenites from Beni Bousera, Morocco. Implications for basalt origin. Science 283, 1303–1306. Bosch, D., Bruguier, O., 1998. An Early Miocene age for a high temperature event in gneisses from Zabargad Island (Red Sea, Egypt): mantle diapirism. Terra Nova 10, 274–279. Bosch, D., Hammor, D., Bruguier, O., Caby, R., Luck, J.M., 2002. Monazite “in situ” 207Pb/ 206 Pb geochronology using a small geometry high-resolution ion probe: application to Archean and Proterozoic rocks. Chemical Geology 184, 151–165. Bossiere, G., Collomb, P., Madjoub, Y., 1988. Sur un gisement de peridotite découvert dans le massif de l'Edough (Annaba, Algérie). Comptes Rendus de l'Académie des Sciences de Paris 306, 1039–1045. Bouillin, J.P., 1986. Le bassin Maghrébin: une ancienne limite entre l'Europe et l'Afrique à l'ouest des Alpes. Bulletin de la Société Geologique de France 8, 547–558. Brueckner, H.K., Medaris, L.G., 2000. A general model for the intrusion and evolution of 'mantle' garnet peridotites in high-pressure and ultra-high-pressure metamorphic terranes. Journal of Metamorphic Geology 18, 123–133. Brueckner, H.K., Blusztajn, J., Bakun-Czubarow, N., 1996. Trace element and Sm–Nd ‘age’ zoning in garnets from peridotites of the Caledonian and Variscan Mountains and tectonic implications. Journal of Metamorphic Geology 14, 61–73.

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