Lithos 112 (2009) 163–187
Contents lists available at ScienceDirect
Lithos
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 / l i t h o s
Jurassic back-arc and Cretaceous hot-spot series In the Armenian
ophiolites — Implications for the obduction process
Yann Rolland a,⁎, Ghazar Galoyan a,b, Delphine Bosch c, Marc Sosson a, Michel Corsini a,
Michel Fornari a, Chrystèle Verati a
a
b
c
Géosciences Azur, Université de Nice Sophia Antipolis, CNRS, IRD, Parc Valrose, 06108 Nice cedex 2, France
Institute of Geological Sciences, National Academy of Sciences of Armenia, 24a Baghramian avenue, Yerevan, 375019, Armenia
Géosciences Montpellier, CNRS UMR-5243, Université de Montpellier II, Place E. Bataillon, 34095 Montpellier Cedex 05, France
a r t i c l e
i n f o
Article history:
Received 2 July 2008
Accepted 16 February 2009
Available online 10 March 2009
Keywords:
Nd–Sr–Pb isotopes
Armenian ophiolite
Back-arc
Obduction
Oceanic plateau
Tethys
Lesser Caucasus
a b s t r a c t
The identification of a large OIB-type volcanic sequence on top of an obducted nappe in the Lesser Caucaus of
Armenia helps us explain the obduction processes in the Caucasus region that are related to dramatic change
in the global tectonics of the Tethyan region in the late Lower Cretaceous. The ophiolitic nappe preserves
three distinct magmatic series, obducted in a single tectonic slice over the South Armenian Block during the
Coniacian–Santonian (88–83 Ma), the same time as the Oman ophiolite. Similar geological, petrological,
geochemical and age features for various Armenian ophiolitic massifs (Sevan, Stepanavan, and Vedi) argue
for the presence of a single large obducted ophiolite unit. The ophiolite, shows evidence for a slow-spreading
oceanic environment in Lower to Middle Jurassic. Serpentinites, gabbros and plagiogranites were exhumed
by normal faults, and covered by radiolarites. Few pillow-lava flows have infilled the rift grabens.
The ophiolite lavas have hybrid geochemical composition intermediate between Arc and MORB signatures:
(La/Yb) N = 0.6–0.9; (Nb/Th) N = 0.17–0.57; (143Nd/144Nd) i = 0.51273–0.51291; (87Sr/86Sr) i = 0.70370–
0.70565; (207Pb/204Pb)i = 15.4587–15.5411; (208Pb/204Pb)i = 37.4053–38.2336; (206Pb/204Pb)i = 17.9195–
18.4594. These compositions suggest they were probably formed in a back-arc basin by melting of a shallow
asthenosphere source contaminated by a deeper mantle source modified by subducted slab-derived products.
87
Sr/86Sr ratios and petrological evidence show that these lavas have been intensely altered by mid-oceanic
hydrothermalism as well as by serpentinites, which are interpreted as exhumed mantle peridotites.
The gabbros have almost the same geochemical composition as related pillow-lavas: (La/Yb)N = 0.2–2.3;
(Nb/Th)N = 0.1–2.8; (143Nd/144Nd)i = 0.51264–0.51276; (87Sr/86Sr)i = 0.70386–0.70557; (207Pb/204Pb)i =
15.4888–15.5391; (208Pb/204Pb)i = 37.2729–37.8713; (206Pb/204Pb)i = 17.6296–17.9683. Plagiogranites show
major and trace element features similar to other Neo-Tethyan plagiogranites (La/Yb)N = 1.10–7.92; (Nb/
Th)N = 0.10–0.94; but display a less radiogenic Nd isotopic composition than basalts [(143Nd/144Nd)i =
0.51263] and more radiogenic (87Sr/86Sr)i ratios. This oceanic crust sequence is covered by variable thicknesses
of unaltered pillowed OIB alkaline lavas emplaced in marine conditions. 40Ar/39Ar dating of a single-grain
amphibole phenocryst provides a Lower Cretaceous age of 117.3 ± 0.9 Ma, which confirms a distinct formation
age of the OIB lavas. The geochemical composition of these alkaline lavas is similar to plateau-lavas [(La/
Yb)N = 6–14; (Nb/Th)N = 0.23–0.76; (143Nd/144Nd)i = 0.51262–0.51271; (87Sr/86Sr)i = 0.70338–0.70551;
(207Pb/204Pb)i = 15.5439–15.6158; (208Pb/204Pb)i = 38.3724–39.3623; (206Pb/204Pb)i = 18.4024–19.6744].
They have significantly more radiogenic lead isotopic compositions than ophiolitic rocks, and fit the
geochemical compositions of hot-spot derived lavas mixed with various proportions of oceanic mantle. In
addition, this oceanic + plateau sequence is covered by Upper Cretaceous calc-alkaline lavas: (La/Yb)N = 2.07–
2.31; (Nb/Th)N = 0.08–0.15; (144Nd/143Nd)i = 0.51271–0.51282; (87Sr/86Sr)i = 0.70452–0.70478), which
were likely formed in a supra-subduction zone environment. During the late Lower to early Upper Cretaceous
period, hot-spot related magmatism related to plateau events may have led to significant crustal thickening in
various zones of the Middle-eastern Neotethys. These processes have likely hindered subduction of some of the
hot and thickened oceanic crust segments, and allowed them to be obducted over small continental blocks such
as the South Armenian Block.
© 2009 Elsevier B.V. All rights reserved.
⁎ Corresponding author. Tel.: +33 4 92 07 65 86.
E-mail address: [email protected] (Y. Rolland).
0024-4937/$ – see front matter © 2009 Elsevier B.V. All rights reserved.
doi:10.1016/j.lithos.2009.02.006
164
Y. Rolland et al. / Lithos 112 (2009) 163–187
1. Introduction
The role of Oceanic Plateaus in the obduction processes of oceanic
crust has still not been clearly established. We understand that their
larger crustal thickness and buoyancy as compared to ‘standard’
oceanic crust does not allow them to subduct, in particular when they
reach subduction zones soon after their formation (e.g., Ben-Avraham
et al., 1981; Cloos, 1993; Abbot and Mooney, 1995; Abbot et al., 1997;
Kerr and Mahoney, 2007). However, the reasons for oceanic crust
obduction onto continental margins are still debated: (i) is obduction
driven by subduction of continental crust? Or (ii) does it result from
the intrinsic nature of the oceanic crust? In the first case, ophiolites are
obducted due to the mechanical coupling of continental crust with the
dense subducting slab (e.g., O'Brien et al., 2001; Guillot et al., 2003).
Continental subduction may be facilitated by the thinned margins of
continental domains following earlier phases of divergence rifting
that precede oceanic crust emplacement (Guillot and Allemand,
2002). In the second case, a lower density of oceanic lithosphere
might result from intra-oceanic hot-spot and magmatic arc events,
which will lead to crustal thickening and a decrease in lithosphere
density (Cloos, 1993; Abbot and Mooney, 1995). The emplacement of
oceanic plateaus has a great influence either on the slab dip, but also
on the cessation of subduction and on the onset of obduction as is
proposed for the Ontong–Java plateau (Petterson et al., 1997). The
ability of an oceanic plateau to resist subduction and eventually be
transported onto continental crust depends on both crustal thickness
and plateau age (Kerr and Mahoney, 2007). The older a plateau, the
cooler and thus the less buoyant it will be. Alternative hypotheses for
obduction involve rapid inversion of tectonic plate motions and rapid
continental convergence (e.g., Agard et al., 2007). Obduction is
ascribed to the presence of young oceanic crust in the hanging-wall
of the subduction zone, as a result of subduction initiation at the MidOceanic Ridge (e.g., Boudier et al., 1988; Nicolas, 1989); or to scalping
of oceanic lithosphere (e.g., Agard et al., 2007 and references therein).
The case of Armenian ophiolites (Lesser Caucasus) is peculiar as
recent investigations (Galoyan et al., 2007, 2009; Rolland et al., in
press) have shown the presence of slow-spreading ophiolites in
several locations. Further, the ophiolites were tectonically transported
above the South Armenian Bloc or SAB (Zakariadze et al., 1983).
Although some blueschists are locally found, these affect oceanic
Fig. 1. Tectonic map of the Middle East — Caucasus area, with main blocks and suture zones, after Avagyan et al. (2005), modified.
Y. Rolland et al. / Lithos 112 (2009) 163–187
crust-derived rocks which underwent intra-oceanic subduction and
exhumation within accretionary prisms (Rolland et al., 2009). In
contrast, the underthrusted Armenian continental crust appears not
to have been metamorphosed by any subduction event. Therefore, the
obduction of the Armenian ophiolites might be explained by the
intrinsic nature of the oceanic crust. However, the slow-spreading
nature of the ophiolites, and in particular the fact that exhumed
mantle forms a large part of the reconstructed ophiolite is rather in
agreement with a relatively dense oceanic lithosphere.
In this paper, we report new geochemical data, including major and
trace elements and Nd, Sr, Pb isotopes, on magmatic series from several
Armenian ophiolites (i.e. Stepanavan (NW Armenia), Sevan (N
Armenia), Vedi (central Armenia); Fig. 2). We identify three superposed
levels of lavas corresponding to three distinct environments: (1) backarc, (2) ‘OIB’-like and (3) arc. Moreover, we suggest that these ophiolite
165
windows correlate with each other and be part of a unique obducted
nappe. Tectonic transport of this nappe onto the SAB can be dated to the
Coniacian–Santonian (88–83 Ma; Sokolov, 1977; Sosson et al., in press).
Finally, the influence of oceanic plateau event in oceanic lithosphere
rheology and its role in the obduction process is discussed.
2. Geological setting
During the Mesozoic, the Southern Margin of the Eurasian
continent has been featured by closure of the Palaeo-Tethys and
opening of the Neo-Tethys Ocean (e. g.; Sengör and Yilmaz, 1981;
Tirrul et al., 1983; Ricou et al., 1985; Dercourt et al., 1986; Stampfli and
Borel, 2002, Fig. 1). Later on, subductions, obductions, micro-plate
accretions, ranging mostly from the Cretaceous to the Eocene, and
finally continent–continent collision have occurred between Eurasia
Fig. 2. Sketch geological map of Armenia, with location of the studied area: 1 — Stepanavan area; 2 — Sevan area; 3 — Vedi area.
166
Y. Rolland et al. / Lithos 112 (2009) 163–187
and Arabia. The study of Armenian ophiolites allows us to unravel part
of this complex history. The ophiolites are located in the northern part
of the Lesser Caucasus region (Fig. 1). The Lesser Caucasus lies south of
the Great Caucasus range, between the Black and Caspian seas. Here,
the ophiolitic belt separates the SAB from the active Eurasian margin.
The SAB is correlated westwards to the Taurides-Anatolide blocks,
which were separated from Gondwana in the Early Mesozoic (LowerMiddle Jurassic) and accreted to the Eurasian margin in the Late
Mesozoic periods (Upper Cretaceous). The Gondwanian nature of the
SAB is shown by the Proterozoic age of the basement crust and the age
and lithologies of the overlying sedimentary series (Aghamalyan,
2004; Sosson et al., in press). The active Eurasian margin is formed by
a thick volcanic arc sequence formed above an active margin, resting
on a Paleozoic (Caledonian to Hercynian) crystalline basement
(Adamia et al., 1981). The ophiolites are situated in three geographic
zones (Fig. 2):
(1) The Stepanavan ophiolite situated in NW Armenia.
(2) The Sevan ophiolite located in NNE Armenia.
(3) The Vedi zone, disposed in a more southerly position, in the
centre of Armenia.
The two first ophiolites are interpreted to correlate with each other
along the Sevan–Akera suture zone at the northern rim of the SAB, and
at the southern edge of the European active continental margin
(Knipper, 1975; Adamia et al., 1980). The Vedi ophiolite is diversely
interpreted as being an obducted sequence above the SAB (Knipper
and Sokolov, 1977; Zakariadze et al., 1983), or within a suture zone
correlating with Central Iran and Alborz ophiolites (Sokolov, 1977;
Adamia et al., 1981). It is generally believed that the different ophiolite
locations may represent suture zones, and thus feature several paleosubduction zones (Aslanyan and Satian, 1977; Knipper and Khain,
1980; Adamia et al., 1981; Aslanyan and Satian, 1982).
A companion paper written on the geology of the Sevan ophiolite
has already put up in details the lithologies and radiometric age of this
ophiolite (Galoyan et al., 2009). Main features are summarized below;
these include:
(i) A high level of fractional crystallisation in the series, with
cumulate olivine gabbros and two pyroxene gabbros overlain and
intruded by amphibole-bearing gabbros and more differentiated
melts (diorites to plagiogranites). These melts are maximally
differentiated and are generally emplaced in ductile extensive
shear zones cross-cutting the gabbros. This complete differentiation series suggests small percent partial melts and long-lived
cooling as is proposed to occur in slow spreading ophiolite settings
(Lagabrielle et al., 1984; Lagabrielle and Cannat, 1990). Absolute
radiometric datings indicate oceanic crust emplacement in the
Middle Jurassic, constrained at 165–160 Ma by zircon U–Pb age of
one tonalite (160 ± 4 Ma; Zakariadze et al., 1990) and by 40Ar/39Ar
amphibole age on gabbro (165.3 ± 1.7 Ma; Galoyan et al., 2009).
(ii) Rare pillow lavas are found, with compositions ranging from
tholeiitic basalts to andesites. The density of the feeding dyke
swarms is reduced, as rare dolerite dykes have been found
crosscutting the series. The slight calc-alkaline composition is also
evidenced by Nb–Ta negative anomalies, which agree with some
slab-derived contamination. These geochemical features support
slow spreading in a back-arc setting.
(iii) Peridotites are frequent and often exhumed as a result of intraoceanic extension. They are generally serpentinized, and witness
further hydrothermal alteration when exhumed at the contact with
marine water (‘listwenites’). The mineralogical nature of the mantlederived ultramafic rocks is still difficult to assess. The previous
petrographical investigations on the serpentinized ultramafics
suggest that protoliths were mantle-derived with various composi-
tions ranging from lherzolites to harzburgites and dunites (e.g.,
Melikyan et al.,1967; Harutyunyan,1967; Palandjyan,1971; Abovyan,
1981; Ghazaryan, 1987; Zakariadze et al., 1990). Undeformed
ultramafics have intrusive cross-cutting relationships and bear a
cumulative magmatic origin, shown by poikilitic texture of olivine
inclusions within large enstatite crystals (up to 10–15 mm;
Palandjyan, 1971). We have observed similar textures, together
with cumulative strata, contained in magmatic pods cross-cutting
serpentinites in the Stepanavan area (Galoyan et al., 2007). These
latter serpentinites are strongly deformed and altered. The ductile
character of deformation is in agreement with a mantle origin for
these rocks.
(iv) Radiolarites are found interlayered or disconformably overlying
the various lithologies described earlier. The fact that they overlie
gabbros, plagiogranites and serpentinites shows that these rocks were
uplifted and denuded by normal faults. Radiolarite datings undertaken in the different ophiolites all agree with oceanic accretion in the
Middle–Upper Jurassic (Danelian et al., 2007, 2008).
The ophiolitic sequences are weakly deformed with anchizonal
metamorphism. Only some outcrops show evidence of small shear
zones ascribed to the ophiolite obduction in the Coniacian–Santonian
(Sokolov, 1977; Zakariadze et al., 1983). HP metamorphism is
described in the Stepanavan region (Fig. 2), where blueschists outcrop
in small km2-size tectonic windows below the ophiolite. Timing of
metamorphism from radiometric 40Ar/39Ar phengite datings indicates
a HP metamorphic peak at ca 95 Ma, and MP–MT retrogression at 73–
71 Ma (Rolland et al., 2009).
The ophiolite series are locally overlain by (1) alkaline lavas, which
have a Lower Cretaceous age, though with very large error bars,
ranging from 120 to 95 ± 20 Ma, (Baghdasaryan et al., 1988; Satian and
Sarkisyan, 2006); and (2) Upper Cretaceous andesites and detrital
series (Dali valley; Stepanavan; Galoyan et al., 2007). The alkaline lavas
are alternatively interpreted as (1) intra-continental rifting (Satian
et al., 2005) in the Vedi area, and (2) plume-derived Ocean Island
magmatism above the oceanic crust before the obduction (Galoyan
et al., 2007, 2009). The calc-alkaline series are ascribed to intra-oceanic
arc emplacement above this oceanic crust sequence and implies the
presence of a subduction zone between the ophiolite and the SAB,
featured by the Stepanavan blueschists (Rolland et al., 2009). These
two magmatic sequences closely predate the ophiolite obduction onto
the SAB during the Coniacian–Santonian (Sokolov, 1977).
3. Analytical methods
Mineral compositions were determined by electron probe microanalysis (EMP). The analyses are presented in Figs. 4 and 6. They were
carried out using a Cameca Camebax SX100 electron microprobe at
15 kV and 1 nA beam current, at the Blaise Pascal University
(Clermont-Ferrand, France). Natural samples were used as standards.
For 40Ar/39Ar dating of the alkaline suite, fresh amphibole grains
were separated from the Vedi ophiolite unaltered sample AR-05-104.
Geochronology of amphiboles was performed by laser 40Ar/39Ar
dating. Results are presented in Table 1 and Fig. 7. Amphibole crystals
were separated under a binocular microscope. The samples were then
irradiated in the nuclear reactor at McMaster University in Hamilton
(Canada), in position 5c, along with Hb3gr hornblende neutron
fluence monitor, for which an age of 1072 Ma is adopted (Turner et al.,
1971). The total neutron flux density during irradiation was 9.0 × 1018
neutron cm− 2. The estimated error bar on the corresponding 40Ar⁎/
39
ArK ratio is ±0.2% (1σ) in the volume where the samples were set.
Three amphibole grains (~ 500 μm in diameter) were chosen for
analysis by the laser UV spectrometer in Géosciences Azur laboratory
at the Nice University. Analyses were done by step heating with a
Y. Rolland et al. / Lithos 112 (2009) 163–187
167
Table 1
Summary of amphibole 40Ar/39Ar dating results from the trachybasalt samples AR-05-104 and AR-05-70.
Step
Laser power (mW)
Atmospheric cont (%)
39
Ar (%)
37
ArCa/39ArK
40
Ar⁎/39ArK
Age (Ma ± 1σ)
Amphibole AR-05-104, J = 44.55, plateau age: 117.3 ± 0.9 Ma (92.4% 39Ar), isochron age: 117.5 ± 0.8 Ma (MSWD: 0.78)
1
400
99.99
0.26
8.84
2
500
95.85
0.88
9.90
3
550
92.07
0.46
2.24
4
650
24.82
6.00
4.15
5
718
8.79
6.40
5.02
6
750
4.17
22.09
5.46
7
800
0.00
9.24
5.54
8
1111
0.81
54.68
5.75
–
4.18
1.19
3.48
3.65
3.64
3.78
3.70
–
132.1
38.5
110.6
115.8
115.6
119.7
117.5
±
±
±
±
±
±
±
±
–
41.2
27.3
3.5
1.9
1.3
1.6
0.5
Amphibole AR-05-70 (1), J = 4.57, plateau age: –, isochron age: 107.8 ± 18 Ma (MSWD: 0.66)
1
340
99.99
4.44
2
460
96.12
25.09
3
560
87.34
52.41
4
640
70.78
12.68
5
640
58.83
2.15
6
1111
62.59
3.23
1.44
3.92
3.08
7.08
46.80
8.09
–
2.41
6.94
11.05
29.02
38.33
–
15
42
67
171
223
±
±
±
±
±
±
–
17
7
16
87
55
Amphibole AR-05-70 (2), J = 35.20, plateau age: –, isochron age: 114.5 ± 37 Ma (MSWD: 15)
1
380
144.64
6.23
2
500
113.84
17.22
3
772
92.47
16.80
4
1093
50.17
53.82
5
1190
123.34
0.71
6
4000
86.24
5.22
2.35
1.31
1.40
4.52
194.72
30.61
–
–
0.90
4.18
–
16.57
–
–
37
166
–
584
±
±
±
±
±
±
–
–
36
10
–
187
50 W CO2 Synrad 48–5 continuous laser beam. Measurement of
isotopic ratios was done with a VG3600 mass spectrometer, equipped
with a Daly detector system; see detailed procedures in Jourdan et al.
(2004). The typical blank values for extraction and purification of the
laser system are in the range 4.2–8.75, 1.2–3.9, and 2–6 cm3 STP for
masses 40, 39 and 36, respectively. The mass-discrimination was
monitored by analyzing air pipette volume at regular intervals. Decay
constants are those of Steiger and Jäger (1977). Uncertainties in
apparent ages in Table 1 are given at the 1σ level and do not include
the error on the 40Ar⁎/39Ark ratio of the monitor.
Thirty-seven samples of magmatic rocks from the Sevan, Stepanavan and Vedi ophiolites have been analyzed for major and trace
elements including Rare Earth Elements (REE; Table 2). Samples were
analyzed at the C.R.P.G. (Nancy, France). Analytical procedures and
analyses of standards can be found on the following website (http://
www.crpg.cnrs-nancy.fr/SARM).
For isotope measurements, powdered samples were weighed to
obtain approximately 100 to 200 ng of Sr, Nd and Pb. A leaching step with
6N HCl during 30 min at 65 °C was done before acid digestion. After
leaching, residues have been rinsed three times in purified milli-Q H2O.
Sr, Nd and Pb blanks for the total procedures were less than 50 pg, 15 pg
and 30 pg, respectively. Lead isotopes were measured by multi-collector
inductively-coupled plasma mass spectrometry (MC-ICP-MS; VG Plasma
54) at the Ecole Normale Supérieure in Lyon (ENSL). Details about isotope
chemical separations and analytical measurements including reproducibility, accuracy and standards, can be found in Bosch et al. (2008) and
references therein. The Nd and Sr isotopic data were measured on a
Finnigan MAT261 multicollector mass spectrometer at the Geochemical
Laboratory, Paul Sabatier University of Toulouse. 87Sr/86Sr was normalised to 8.3752, NBS standard was measured to 0.710250 (±15). 143Nd/
144
Nd ratio was normalised to a value of 146Nd/144Nd of 0.71219; measure
of Rennes standard was 0.511965 (±12).
4. Results
4.1. Field relationships
Synthetic logs are drawn on Fig. 3, showing the lithological
associations and the structural relationships in each of the three
studied ophiolitic zones.
4.1.1. Description of the ophiolitic units
In Stepanavan (Fig. 3A, B), ophiolite sections exhibit abundant
serpentinites, cross-cut by normal fault and shear zones in which
gabbro-norites, gabbros and plagiogranites are intrusive and
deformed (see Galoyan et al., 2007, 2009 for details and crosssections). Laterally, thick layers of pillow basalts are observed which
interlayer and are covered by radiolarites. On top of the ophiolite
section, a thin layer of alkaline lava flows is found. Above, these lavas
are eroded and unconformably covered by Upper Cretaceous
conglomerates and limestones, and calc-alkaline pillow basalts or
graywackes. The ophiolite sequence is thrusted over a blueschist facies
metamorphic sole.
In the Sevan area, sections are extremely variable laterally (Fig. 3C–E;
see Galoyan et al., 2009 for details and cross-sections). Pillow lavas are
rare, and serpentinites were frequently exhumed. Intense hydrothermal
alteration (‘listwenites’) has transformed the uppermost part of
exhumed serpentinites. The feeding doleritic dyke swarm is extremely
scarce. Large intrusive pods of amphibole-bearing gabbros and
plagiogranites are also exhumed and covered by radiolarites. Normal
faults are observed, and are interpreted as the cause of such lateral
variations, by vertical uplifting of footwall sections, and local infilling of
axial rift valleys, following the model of slow-spreading ophiolite (e.g.,
Lagabrielle et al., 1984; Lagabrielle and Cannat, 1990). Locally, thick
sequences of alkaline pillow lavas are observed. The ophiolite is locally
eroded, and unconformably covered by basal conglomerates and soils,
and an Upper Cretaceous section of limestones comprising graywackes
interlayers.
In the Vedi area, the ophiolite section is much thinner (Fig. 3F–H,
see Galoyan et al., 2009 for details and cross-sections). The basal
tectonic contact is exposed, with the top oriented to the south sense of
shear. At the base, the ophiolite rests on a serpentinite layer by a
tectonic contact. The ophiolite is intensely sheared above the basal
contact with boudins of tholeiitic basalts (Fig. 3H). Laterally, the
ophiolite consists mainly of gabbros (Fig. 3G) or serpentinites, which
suggests a similar lithology as in the Sevan and Stepanavan areas.
However, the different parts of the ophiolite are dismembered and
displaced from each other as a result of obduction deformation. Above
the ophiolites, layers of radiolarites are found below a very thick
section of alkaline pillow lavas (Fig. 3H). This section is of variable
thickness depending on the location. This may be due to lateral
168
Table 2
Representative whole-rock analyses of samples from the Sevan, Stepanavan and Vedi areas, major oxides are in wt.%, and trace elements and REE in ppm.
Groups
Sevan ophiolitic series
No.
Flaser
gabbro
Olivine
gabbro
Sevan Alkaline series
Olivine
gabbro
Gabbro
Gabbronorite
Hornblende
gabbro
Diorite
Diorite
Plagiogranite
Diabase
Trachyandesite
Basaltic
tracyandesite
Stepanavan ophiolitic series
Andesite
Basanite
Trachybasalt
Trachyandesite
Websterite
Hornblende
gabbro
Hornblende
gabbro
AR-03-25
AR-05-86
G150
AR-03-39
AR-03-24
AR-03-10
AR-04-218
AR-03-23
AR-03-19
AR-03-02
G154
AR-03-17
AR-03-34
G142
AR-05-80
AR-03-33
AR-04-03
AR-04-16
AR-04-45D
SiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
TiO2
P2O5
LOI
Total
Mg#
Rb
Sr
Y
Zr
Nb
Ba
Hf
Ta
Pb
Th
U
V
Cr
Co
Ni
Cu
Zn
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Eu/Eu⁎
(La/Sm)N
(La/Yb)N
45.36
13.32
14.91
0.14
7.44
10.89
3.15
0.22
2.55
0.31
1.46
99.8
52.1
2.67
231.7
61.81
175.9
4.51
36.77
4.25
0.35
–
0.08
0.11
440.7
35.43
26.81
77.98
18.16
58.57
14.89
34.75
5.057
24.09
7.38
3.11
9.32
1.60
10.47
2.22
6.37
0.94
6.28
0.95
1.15
1.27
1.60
48.09
16.72
5.94
0.11
10.51
14.1
1.65
–
0.27
0.02
2.8
100.2
79.3
0.48
102
6.41
5.28
0.08
4.1
0.22
0.01
–
–
–
138
802
40.2
188
102
28.3
0.33
0.94
0.18
1.23
0.58
0.34
0.87
0.16
1.14
0.24
0.65
0.10
0.65
0.09
1.45
0.36
0.34
48.39
15.63
6.44
0.12
10.48
16.65
1.16
–
0.29
0.04
0.65
99.8
77.9
–
101
7.24
4.85
–
–
0.21
–
–
–
–
190
412
41.7
131
111
29.6
0.36
1.02
0.19
1.28
0.61
0.32
0.98
0.18
1.24
0.26
0.74
0.11
0.72
0.11
1.28
0.37
0.34
49.49
14.11
11.59
0.18
6.79
9.38
3.52
0.29
1.32
0.14
2.92
99.7
56.1
3.05
189.5
30.05
75.20
1.55
120.8
2.12
0.11
–
0.29
0.09
319.9
94.50
34.77
32.35
52.59
86.18
3.39
9.57
1.57
8.27
2.93
1.07
4.01
0.71
4.89
1.03
3.02
0.46
3.03
0.48
0.95
0.73
0.76
50.60
7.20
7.77
0.17
15.29
17.65
0.48
–
0.20
0.06
0.79
100.0
81.1
–
58.2
7.84
5.24
–
14.14
0.2
–
–
–
–
195.9
810.3
41.88
136.2
142.3
47.98
0.28
0.99
0.22
1.40
0.67
0.28
1.07
0.20
1.36
0.29
0.83
0.13
0.86
0.13
1.01
0.26
0.22
50.68
18.17
9.09
0.16
6.75
9.26
3.21
0.15
0.36
0.05
1.21
99.1
61.6
0.84
303.7
11.93
22.89
0.50
34.36
0.81
0.04
–
0.14
0.07
222.2
104.3
33.71
29.23
44.81
71.99
1.60
4.51
0.74
3.80
1.34
0.48
1.70
0.30
1.92
0.41
1.22
0.19
1.32
0.21
0.97
0.75
0.82
55.09
13.45
8.51
0.15
9.84
8.76
2.59
0.17
0.24
0.04
1.87
100.7
71.6
1.42
207.4
7.82
21.08
0.32
34.9
0.76
0.03
–
0.07
0.04
158.1
562.7
39.6
148.1
15.63
73.35
1.44
3.73
0.57
2.87
0.94
0.35
1.13
0.20
1.31
0.28
0.82
0.13
0.92
0.15
1.03
0.97
1.06
57.41
14.10
8.84
0.14
2.24
4.92
6.36
0.12
0.87
0.16
4.33
99.5
35.6
1.81
212.5
25.22
75.07
1.83
55.59
2.27
0.14
2.62
0.75
0.25
122.6
136.9
15.76
10.09
21.02
50.45
4.62
11.73
1.82
9.05
2.91
0.97
3.68
0.69
4.60
0.97
2.95
0.47
3.25
0.52
0.91
1.00
0.96
74.91
12.32
3.58
0.03
0.42
3.05
4.20
0.31
0.21
0.04
0.57
99.6
20.2
2.22
145.2
27.65
71.64
2.15
65.83
2.48
0.10
1.32
1.09
0.62
50.8
1464
5.52
37.19
6.91
9.98
5.38
12.66
1.78
8.36
2.68
0.67
3.50
0.64
4.33
0.94
2.94
0.46
3.30
0.53
0.67
1.26
1.10
46.02
16.29
8.39
0.13
7.73
10.68
3.53
0.37
1.26
0.17
4.61
99.2
66.6
13.97
630.9
23.93
127.4
2.95
285.3
2.91
0.24
1.51
1.01
0.29
179.3
277.0
38.26
54.43
58.20
66.89
7.07
18.32
2.71
12.68
3.53
1.36
3.95
0.67
4.21
0.84
2.45
0.36
2.39
0.37
1.12
1.26
2.00
53.70
14.09
11.35
0.15
4.52
3.64
6.07
–
1.36
0.12
4.85
99.9
46.2
–
28.76
27.82
81.43
1.77
16.09
2.36
0.14
1.61
0.55
0.35
334.9
–
29.48
8.95
102.8
67.20
4.25
11.14
1.67
8.76
2.93
1.14
4.04
0.71
4.78
1.01
2.99
0.46
3.16
0.51
1.01
0.91
0.91
54.27
15.16
12.36
0.19
3.74
4.38
6.63
–
1.33
0.15
1.78
100
39.5
–
49.82
29.88
85.60
1.44
19.01
2.38
0.12
–
0.37
0.27
321.8
–
24.33
5.14
15.61
80.0
3.91
10.93
1.81
9.35
3.26
1.15
4.18
0.75
4.99
1.06
3.11
0.47
3.22
0.51
0.95
0.75
0.82
55.48
14.13
12.45
0.13
4.07
5.49
3.96
0.61
1.17
0.11
2.18
99.8
41.6
5.05
102.5
27.49
54.39
1.69
16.65
1.63
0.12
–
0.33
0.14
347.4
251.7
27.09
16.09
5.23
19.68
2.69
7.17
1.15
6.17
2.32
0.79
3.36
0.62
4.34
0.95
2.82
0.44
2.98
0.47
0.86
0.73
0.61
40.63
14.40
11.70
0.29
4.15
11.40
3.64
1.42
2.06
0.48
10.01
100.2
43.4
31.89
147.0
25.58
153.4
40.63
166.7
3.57
2.99
4.98
4.13
1.43
257.4
33.86
38.26
34.29
61.86
100.1
32.41
64.63
7.64
29.89
6.02
1.97
5.58
0.82
4.69
0.90
2.48
0.35
2.33
0.36
1.04
3.39
9.39
43.80
17.58
9.46
0.11
6.70
4.95
4.09
2.24
1.68
0.44
8.89
99.9
60.5
44.83
330.8
17.37
131.4
17.95
299.0
2.86
1.19
4.63
4.06
0.94
271.7
21.68
31.44
28.76
54.48
91.55
29.46
59.03
6.92
27.11
5.21
1.63
4.45
0.62
3.42
0.61
1.64
0.22
1.44
0.22
1.03
3.56
13.83
51.57
14.34
6.06
0.11
0.77
9.78
6.34
0.56
1.98
1.08
6.65
99.3
21.7
7.93
341.5
52.91
411.2
49.22
168.7
9.03
3.80
4.45
5.15
4.61
286.1
162.0
24.06
13.83
18.34
92.74
48.54
107.1
13.38
56.55
12.98
4.14
12.43
1.84
10.27
1.87
4.80
0.64
4.05
0.60
1.0
2.35
8.09
53.24
1.03
6.02
0.15
23.18
16.52
0.12
–
0.05
0.03
0.54
100.9
90.0
–
11.79
1.05
–
–
3.71
–
–
–
–
–
135.1
2804
55.82
361.3
340.3
25.68
–
0.15
0.02
0.15
0.08
0.03
0.13
0.03
0.18
0.04
0.12
0.02
0.12
0.02
0.85
0.0
0.0
47.30
14.39
12.90
0.21
9.11
10.14
2.93
0.19
1.18
0.07
1.76
100.2
60.4
1.12
125.4
20.91
42.74
1.01
32.71
1.21
0.08
–
0.18
0.05
324.7
236.4
51.46
78.0
–
60.28
2.40
6.37
1.08
5.76
2.09
0.96
2.98
0.53
3.53
0.74
2.15
0.32
2.10
0.32
1.17
0.72
0.77
53.77
14.00
8.92
0.15
7.81
6.98
3.34
2.42
0.16
0.05
2.44
100.1
65.6
30.17
213.4
5.79
21.26
2.14
228.1
0.63
0.21
3.61
1.27
0.43
94.47
324.2
31.5
101.7
189.9
60.33
3.06
6.28
0.65
2.30
0.49
0.19
0.55
0.10
0.81
0.19
0.64
0.12
0.91
0.17
1.12
3.94
2.27
Y. Rolland et al. / Lithos 112 (2009) 163–187
Sample
Table 2 (continued )
Stepanavan ophiolitic series
Stepanavan alkaline series
Stepanavan calk-alkaline
series
Vedi ophiolitic series
Vedi Alkaline series
No.
Plagiogranite Basaltic trachy- Basaltic trachy- Basaltic trachy- Basaltic
Diabase
andesite
andesite
andesite
trachy-andesite
Olivine
basalt
Basaltic
trachyandesite
Basaltic Hornblende
trachygabrro
andesite
Diorite
Plagiogranite Basalt
Basaltic
andesite
Basalt
Trachybasalt
Basaltic
Trachydacite
trachyandesite
Sample
AR-04-44
AR-04-20
AR-04-30
AR-06-02
AR-03-53
AR-0405
AR-0432
AR-0440A
AR-0431
AR-05-113
AR-05110
AR-05-111
AR-05114
AR-05106
AR-0578
AR-05-104
AR-05-102
AR-04-75
SiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
TiO2
P2O5
LOI
Total
Mg#
Rb
Sr
Y
Zr
Nb
Ba
Hf
Ta
Pb
Th
U
V
Cr
Co
Ni
Cu
Zn
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Eu/Eu⁎
(La/Sm)N
(La/Yb)N
75.35
12.20
2.71
0.03
0.77
2.05
5.03
–
0.11
0.02
1.07
99.3
38.1
0.58
91.04
1.23
6.48
0.35
20.58
0.15
–
–
0.02
0.01
30.74
421.5
7.71
24.86
189.5
23.13
2.42
3.90
0.42
1.58
0.29
0.35
0.24
0.04
0.22
0.05
0.15
0.03
0.21
0.04
4.05
5.25
7.92
51.53
14.69
14.81
0.23
4.15
4.86
5.74
0.18
1.62
0.13
1.89
99.8
37.7
1.4
61.01
35.91
86.0
1.62
19.58
2.54
0.13
2.29
0.43
0.12
459.8
99.23
38.85
22.76
64.58
130.6
4.23
11.08
1.88
9.88
3.44
1.28
4.69
0.87
5.93
1.28
3.79
0.58
3.89
0.61
0.98
0.77
0.73
48.55
13.29
8.67
0.15
6.86
10.49
4.74
0.24
1.08
0.11
6.01
100.2
63.1
7.94
95.7
26.52
68.43
1.9
21.73
1.83
0.15
1.93
0.19
0.09
305.4
316.7
42.98
109.1
132.8
81.13
2.53
7.37
1.31
7.04
2.55
0.99
3.49
0.65
4.34
0.93
2.75
0.42
2.84
0.44
1.02
0.63
0.60
45.37
14.27
13.52
0.32
6.22
4.09
1.53
5.37
3.12
1.31
5.11
100.2
49.8
45.3
157
40.6
268
52.8
608
5.97
3.97
7.04
5.39
1.34
172
–
26.3
6.5
20.4
177
50.3
96.1
12.1
53.3
11.2
3.87
10.5
1.50
8.11
1.49
3.75
0.51
3.18
0.45
1.09
2.83
10.68
48.54
15.01
12.65
0.27
4.25
5.33
3.93
2.69
2.64
1.08
3.16
99.5
42.3
33.17
322.6
44.41
294.4
57.95
578.3
6.51
4.20
2.54
5.98
1.46
94.35
25.77
16.75
–
9.48
137.1
50.59
107.0
12.97
53.32
11.26
4.08
10.41
1.52
8.69
1.57
4.16
0.57
3.64
0.56
1.15
2.83
9.38
50.19
13.91
13.73
0.24
3.27
5.85
5.11
0.42
3.39
0.67
2.94
99.7
34.0
7.62
198.8
51.24
373.5
42.33
156.6
8.01
3.24
2.24
4.65
1.20
201.6
–
31.93
–
14.94
152.7
40.02
85.12
10.85
45.27
10.35
3.39
10.3
1.63
9.50
1.80
4.93
0.70
4.51
0.69
1.0
2.43
5.99
49.15
18.53
10.19
0.16
5.25
8.25
4.36
0.52
0.86
0.14
3.12
100.5
52.7
9.61
520.3
16.0
44.4
2.14
133.9
1.25
0.17
7.22
0.72
0.19
241.5
21.05
29.51
15.55
12.6
150.9
4.93
11.46
1.74
8.35
2.36
0.94
2.60
0.44
2.80
0.57
1.62
0.24
1.60
0.25
1.15
1.31
2.07
49.79
15.80
8.82
0.15
3.54
9.12
3.54
1.24
1.07
0.20
7.27
100.6
48.4
18.12
303.8
24.4
99.18
2.29
239.1
2.69
0.18
3.42
1.46
0.67
279.1
31.91
29.05
22.38
188.1
86.53
7.87
18.51
2.68
12.58
3.47
1.13
3.89
0.65
4.09
0.85
2.46
0.38
2.54
0.39
0.94
1.43
2.09
52.20
17.05
9.28
0.16
3.59
6.56
4.61
1.00
0.94
0.18
5.36
100.9
44.5
18.45
282.3
24.19
95.3
3.32
213.6
2.61
0.26
5.66
1.67
0.58
263.7
73.83
27.36
18.95
170.9
100.0
8.69
18.05
2.53
11.65
3.15
1.04
3.57
0.61
3.96
0.83
2.46
0.38
2.54
0.40
0.95
1.74
2.31
45.09
21.24
4.32
0.07
7.88
14.29
2.12
0.17
0.16
–
5.08
100.4
79.7
2.41
492.7
3.70
4.30
0.08
131.6
0.16
–
–
–
–
89.5
785.2
30.98
130.7
92.4
24.06
0.231
0.71
0.14
0.78
0.36
0.24
0.53
0.10
0.64
0.14
0.38
0.06
0.35
0.06
1.64
0.40
0.44
58.57
16.03
6.66
0.11
5.18
7.08
3.94
0.47
0.33
0.04
2.08
100.5
62.7
4.12
254
10.1
42.3
0.49
57.5
1.3
0.04
1.1
0.42
0.12
185
123
22.8
42.4
18.3
50.7
2.09
4.84
0.72
3.74
1.20
0.42
1.46
0.25
1.65
0.34
1.01
0.16
1.14
0.18
0.97
1.10
1.24
70.45
14.69
4.44
0.08
1.06
3.87
4.47
0.2
0.43
0.09
1.04
100.8
34
1.11
161
13.1
86.5
0.59
33.5
2.35
0.05
1.12
0.48
0.13
44.5
9.1
6.6
6.4
–
39
2.91
6.15
0.81
4.48
1.45
0.69
1.81
0.32
2.11
0.46
1.35
0.21
1.52
0.25
1.31
1.26
1.29
47.5
16.17
8.76
0.15
8.46
8.57
3.99
0.72
0.93
0.09
4.78
100.1
67.6
3.51
134.6
20.78
54.9
0.69
141.7
1.48
0.07
–
0.15
0.06
211.3
416.7
42.44
195.5
14.05
65.52
1.96
6.25
1.09
5.9
2.15
0.90
3.0
0.53
3.48
0.74
2.08
0.32
2.15
0.33
1.08
0.57
0.62
48.3
15.03
10.14
0.16
5.97
8.01
3.96
0.18
1.2
0.12
7.14
100.2
56.0
4.14
110.1
26.91
73.98
2.46
16.2
2.04
0.20
–
0.22
0.10
238.3
324.7
49.08
122.6
81.78
94.28
3.07
8.51
1.47
7.97
2.88
1.12
4.02
0.70
4.67
0.97
2.78
0.42
2.83
0.45
1.0
0.67
0.73
44.58
12.52
9.36
0.12
2.63
15.53
3.83
–
2.35
0.33
9.17
100.4
37.8
0.63
153.5
24.94
160.8
23.22
1097
3.91
1.76
1.53
2.085
0.643
240.7
50.2
28.9
26.28
28.41
106.8
18.44
39.11
5.02
21.85
5.50
1.93
5.72
0.85
4.85
0.88
2.30
0.31
1.95
0.29
1.05
2.11
6.39
44.64
15.41
11.99
0.14
4.85
7.85
4.24
0.96
3.67
0.85
5.04
99.6
46.6
10.48
926
36.26
318.8
67.52
444.9
6.81
4.88
1.25
4.60
1.18
219.8
4.22
31.4
21.81
50.82
146.8
49.35
109.7
14.15
59.18
12.52
4.15
10.96
1.48
7.84
1.32
3.26
0.42
2.49
0.35
1.08
2.48
13.39
50.39
16.2
7.76
0.13
5.05
7.16
3.24
1.96
2.36
0.64
5.07
99.9
58.4
27.1
643
22.4
260
43.3
659
6.03
3.29
6.71
8.39
1.8
135
136
64.3
126
42.2
134
50.7
91.8
9.83
42.9
8.47
2.67
7.36
0.99
5.01
0.81
1.92
0.25
1.46
0.22
1.03
3.77
23.44
59.61
17.48
7.64
0.12
1.11
1.89
6.37
2.4
0.72
0.25
2.14
99.7
23.16
64.89
260.4
58.18
680.5
82.24
422
14.76
5.99
4.30
12.28
2.78
5.26
66.9
4.98
–
9.93
167.9
74.8
142.6
15.84
59.79
12.64
3.68
11.69
1.877
10.97
2.08
5.77
0.86
5.84
0.88
0.93
3.73
8.64
Y. Rolland et al. / Lithos 112 (2009) 163–187
Groups
169
170
Y. Rolland et al. / Lithos 112 (2009) 163–187
Fig. 3. Representative geological logs of the Stepanavan, Sevan and Vedi ophiolites.
Y. Rolland et al. / Lithos 112 (2009) 163–187
variations in the amount of erupted lavas, or may be explained by
tectonic scalping of the ophiolite upper part during obduction. These
alkaline lava flows are well exposed and preserved in the Vedi ophiolite.
They are made of amphibole-bearing basaltic pillows. The pillows are
larger (metre scale) than the ophiolites ones (several decimetre scale),
and interlayer with thin pink limestones. At the front of the obduction in
the Vedi zone, an olistolith formation exhibits conglomerates and slided
blocks in a muddy matrix (Fig. 3F). The olistostrom age is Coniacian–
Santonian (nanofossils, Carla Muller, com. pers.), and it connects
progressively below and above with Lower and Upper Coniacian reef
limestones, respectively. Therefore, the obduction age can be bracketed
to the Coniacian–Santonian, which agrees with former estimates
(Sokolov, 1977). Laterally, always in the Vedi zone, the upper part of
the ophiolite is made of kilometre scale slided blocks, mainly comprised
of alkaline pillow basalts and calc-schists. These blocks slide on a
greenish mudstone rock, probably originated from the ophiolite
alteration. The Upper Coniacian uncomformity is variably marked by
conglomerates, marls and reef limestones.
4.1.2. General features of the Armenian ophiolites: evidence for LOT
features
As emphasized in Galoyan et al. (2009) in the Sevan area, and by
Galoyan et al. (2007) in the Stepanavan area, the lithologies found in
all the exposed Armenian ophiolites are in good agreement with the
hypothesis of a slowly expanding spreading centre, as described for
the western Alps ophiolites and exposed earlier in Section 2 (Nicolas
and Jackson, 1972; Nicolas, 1989).
The similar lithological and age features found in the several
Armenian ophiolites suggest that they were part of the same oceanic
crust section. This has to be confirmed by the comparison of
geochemical data from each zone. The presence of three magmatic
series: ophiolite s.s. (tholeiitic), ‘OIB’ (alkaline) and arc (calc-alkaline)
in the same structural position (from bottom to top, respectively) has
been evidenced in the three zones. Isotopic geochemistry on the three
series will allow us to constrain the nature of sources and to identify
the magmatic processes that existed prior to ophiolite obduction.
4.2. Petrography and mineral chemistry
The field and microscopic analyses of Armenian ophiolite magmatic rocks show a continuous magmatic succession from ultramafic
cumulates (wherlites, websterites) to gabbros and plagiogranites,
cross-cutting intensely altered serpentinites in each of the three
studied ophiolites. All these lithologies were exhumed in the footwall
below normal faults and covered by pillow-basalts.
4.2.1. Serpentinites
The study of serpentinite mineralogy is difficult due to intense
serpentinization. However, EMP analysis of chromiferous spinel relicts
from Sevan Tsapatagh area (sample AR-05-80 in Galoyan et al., 2009;
Fig. 3) reflects still unaltered mineral compositions. EMP analyses
show a very narrow compositional range in Cr# (Cr / Cr + Al = 0.71–
0.73) and Mg# (Mg / Mg + Fe = 0.58–0.59) ratios (Fig. 4). These Cr#
compositions are more elevated than those of abyssal peridotites
(Brynzia and Wood, 1990) and agree with a fore-arc peridotites
composition (Parkinson and Pearce, 1998). However, Mg# values are
slightly lower than those of Parkinson and Pearce (1998), which is
ascribed to high partial melting in such context. However, we cannot
exclude any hydrothermal process, as the primary nature of chromites
is uncertain.
4.2.2. Ophiolite plutonic rocks
Wehrlites are found in the Stepanavan ophiolite (see Galoyan et al.,
2007). They have a poikilitic texture showing numerous clinopyroxene crystals with diopside composition (Wo45–47En48–50Fs2–4),
included in large olivine Fo87–88 (N60–65%) porphyric grains.
171
Fig. 4. Chemistry of Cr-spinel from Armenia with respect to the compositional fields of
Abyssal peridotites (1; Brynzia and Wood, 1990) and arc-related peridotites (Mariana
seamount peridotites (2) and dunites (3); Parkinson and Pearce, 1998).
Gabbros are the most abundant rocks in the crustal complex, and
are found in each ophiolite zone (Galoyan et al., 2007, 2009). Their
petrography evolves from cumulate-banded olivine gabbros in their
lower part towards more leucocratic plagioclase-rich gabbros in the
upper part (Abovyan, 1981; Ghazaryan, 1987, 1994). The cumulative
banded olivine gabbros and websterites are found locally, only in the
Sevan and Stepanavan areas, while more leucocratic gabbros are
widespread in the three zones.
Olivine gabbros found in Stepanavan and Sevan ophiolites (Galoyan et
al., 2007, 2009) are fresh, massive, and fine- to medium-grained (0.5 to
2 mm). They have cumulate, ophitic textures and consist of plagioclase
(~60–65%; An68–74, An80–89), olivine (~5–10%; Fo72–76), and clinopyroxene (~25–35%). Clinopyroxene is of augite (Wo39–44En45–48Fs11–13)
and diopside (Wo45En44Fs11) types. Some enstatite orthopyroxenes
(Wo2En75Fs23) are also found rimming olivine porphyrocrysts.
Websterites found in Stepanavan and Sevan ophiolites (Galoyan
et al., 2007, 2009) have a granular texture with large 2–8 mm
porphyrocrysts of orthopyroxene (30–70%), clinopyroxene (70–30%)
and olivine grains (0–35%; Fig. 5A). Orthopyroxenes are enstatite-rich
(Wo1–5En59–84Fs11–37) and clinopyroxenes are augites (Wo35–42En36–
40Fs15–19), olivine is relatively rich in forsterite (Fo84–88). Gabbronorites (from Stepanavan; Galoyan et al., 2007) have a gabbroic texture,
with plagioclase (10–60%, 1–3 mm), clino- and ortho-pyroxene.
Plagioclase is of bytownite type (An80–85), while orthopyroxenes (1–
4 mm) are enstatites (Wo2–5En59–61Fs34–37), and clinopyroxenes are
augites (Wo35–42En36–40Fs15–19).
Mesocratic to leucocratic gabbros of the upper section found in the
three ophiolites (Galoyan et al., 2007, 2009; Rolland et al., in press) are
massive, fine- to medium-grained and have gabbroic (or gabbro-ophitic),
xenomorphic granular texture (0.5–4 mm), with plagioclase (~40–65%;
An50–75, An72–93), clinopyroxene (8–45%; augite) and hornblende (0–
40%), without any olivine. Accessory minerals (1–10%) are apatite,
titanomagnetite, ilmenite and rarely quartz. The hornblende-rich gabbros
(Galoyan et al., 2009; Rolland et al., in press) have coarse granular
textures (Fig. 5B), with ~50–65% euhedral to subhedral plagioclase
(An54–58) and (~35–50%) anhedral to subhedral amphibole. Some brown
Ti-rich euhedral hornblende is presumed to be a primary mineral; while a
Ti-poor subhedral to xenomorphic green magnesio-hornblende (Leake et
al., 1997) is thought to be a secondary phase formed by hydrothermal alteration as it replaces generally the brown type. The augite
(Wo40–42En39–47Fs11–14), diopside (Wo45–48En40–44Fs8–15), and enstatite
(Wo2En57Fs41) relicts (5–10%) are found in the crystals of magnesiohornblendes that replace the pyroxenes. However, it is not related to
shear zones and fractures, and is thus a late magmatic mineral. In
leucocratic gabbros (Galoyan et al., 2009), the clinopyroxene (augite
172
Y. Rolland et al. / Lithos 112 (2009) 163–187
Wo40–41En33–35Fs18–19) content does not exceed 25%. Normal zoning is
observed in plagioclase (from An85 to An60), which is frequently
altered. Clinopyroxenes have alkaline to slightly tholeiitic compositions
(0.8 b Na+ Cab 0.9; Leterrier et al., 1982, Fig. 6). Pegmatitic gabbros
crosscutting the plutonic sequence (Vedi zone; Rolland et al., accepted)
are composed of plagioclase and hornblende, and are mainly altered
into chlorite, carbonate, sericite, albite, quartz, actinolite, etc.
Diorites occur as small intrusive bodies within the gabbro units in
Sevan and Vedi zones (Palandjyan, 1971; Abovyan, 1981; Ghazaryan,
1987, 1994). They have a porphyritic to subhedral granular (1–4 mm)
Y. Rolland et al. / Lithos 112 (2009) 163–187
173
Fig. 6. Chemical compositions of studied clinopyroxenes plotted in the Ti vs. (Na + Ca) diagram of Leterrier et al. (1982). Note that a majority of data plot in the Alkaline compositional
field, and a minority is in the Toleiitic part.
texture and have relatively similar hornblende contents (5–30%) as
gabbros. Plagioclase (~65–70%) is albite-rich (An34–38) and accessory
minerals (quartz, opaque oxides) are rare. Amphibole grains are
magnesio-hornblendes in composition, sometimes rimmed by actinolite fringes and epidote aggregates. In Sevan zone, diorites grade
into quartz-diorites (quartz 5–10%), laterally and upwards in the
series.
Plagiogranites are found in the three zones (Galoyan et al., 2007,
2009; Rolland et al., in press). They appear to be dioritic intrusives
most differentiated components, forming diffuse segregations or
discontinuous networks of veins. Plagiogranites have local coarse
pegmatic, or hypidiomorphic to xenomorph granular (0.5–4 mm)
textures. They are formed by 40–65% plagioclase (An15–30), 25–45%
quartz, minor biotite (b5%), ortho-amphibole (b5%; Stepanavan), Kfeldspar (0–10%, microcline; Fig. 5C) and accessory phases (titanomagnetite, hematite, sphene and apatite).
4.2.3. Ophiolite volcanic and subvolcanic rocks
Diabases are present in several locations (Sevan and Stepanavan
areas; Galoyan et al., 2007, 2009) as isolated dikes, crosscutting the
layered gabbros. They are generally altered (chlorite, epidote, carbonates) and have a subdoleritic texture composed of plagioclase (60–65%;
An65–75) and two clinopyroxenes (augite Wo41–44En44–47Fs11–13 and
diopside Wo45En37Fs18).
The volcanic rocks of the Armenian ophiolites we studied are
present as pillowed and massive lava flows and pillowed breccias. In
general, they show signs of hydrothermal alteration but relict igneous
textures are preserved. In the three locations basalts and basaltic
andesites are vesicular (1–5 mm, filled with carbonate-calcite, chlorite
and quartz) and largely aphyric (intersertal, spilitic, microdoleritic
and variolitic, up to 1.5–2 mm in diameter), composed mainly of
albitized plagioclase and/or plagioclase–clinopyroxene microlites, Timagnetite and hematite microlites, in a devitrified (calcite + chlorite)
groundmass (Fig. 5D).
4.2.4. Alkaline lavas
The alkaline basalts are found in the three zones on top of the
ophiolite section as large massive pillow-lavas or as diabase dykes, but
their relationships with the ophiolite (s.s.) pillow lavas remain unclear.
The first group of alkaline rocks displays large vesicles (0.5–3 mm), filled
with carbonates and rarely chlorites, and have both phyric and aphyric
(Fig. 5E). They have intersertal textures, with plagioclase megacrysts
(~5%; 0.5–2 mm), microliths and opaque minerals (3–10%), surrounded
by a calcite–chlorite mesostase. The second group (Vedi and Stepanavan
zones, e.g., samples AR-05-70 and AR-05-104 dated by 40Ar/39Ar) have
doleritic (Fig. 5F) to ophitic textures. They are mainly composed of
plagioclase (~40–55%; 1–3 mm), clinopyroxene (10–30%; 1–4 mm),
amphibole (~25%; 1–3 mm) and accessory Ti-magnetite (N5–10%),
apatite (~3%; prismatic, acicular, 0.5–1.5 mm) and rarely biotite. Apatites
are present in the plagioclase crystals and in the vitreous interstices,
which are filled by carbonates or carbonates-chlorites. The tabular
plagioclase laths show a transitional zoning with bytownite to labrador
(An72–60) or labrador to andesine (An55–32) compositions. Thin rims of
pure albite (Ab — 98%) are also present. Clinopyroxenes are generally
chloritized, but still preserve diopside compositions (Wo49En35Fs16).
The amphibole is a kaersutite (Leake et al., 1997), with zoning from
kaersutite to ferro-kaersutite from core to rim. Some samples show
abundant calcite-filled veins and pockets.
A few dacitic sills and dikes occur among the basaltic pillow lava
flows in the Vedi valley. As in the pillow basalts, plagioclase is the
main mineral phase and Fe-oxides are present (~5–10%; Fig. 5G).
Some 1–2 mm large unzoned plagioclase phenocrysts of oligoclase-
Fig. 5. Microphotograph of representative magmatic rock types from the Armenian ophiolite complex. Plutonic and volcanic ophiolite series: (A) subautomorph granular texture of a cumulate
banded websterite (sample AR-04-36, Stepanavan area, Cheqnagh valley; see Galoyan et al., 2007); (B) coarse-grained hornblende gabbro with normally zoned plagioclases (sample AR-05110, Vedi area, massif of Qarakert, see Rolland et al., in press-b); (C) xenomorph granular texture of a microcline (Mc) bearing plagioclase rich leucogranite (sample AR-05-109, in the same
massif, see Rolland et al., accepted); (D) aphyric, intersertal (spilitic) and variolitic basalt composed of mainly albitized plagioclase, Ti-magnetite and hematite microlites, in a devitrified
groundmass (sample AR-05-106, Vedi area, Khosrov valley, see Rolland et al., accepted). Alkaline series: (E) aphyric, intersertal basalt, totally devoid of phenocrysts, and composed of
carbonatized plagioclase microlites and opaque minerals (~5%) in a chlorite–carbonate groundmass (sample AR-05-80, Sevan area, Tsapatagh valley, see Galoyan et al., 2009); (F) doleritic
texture in a trachybasalt composed of plagioclase, chloritized clinopyroxene, kaersutite (Krs), Ti-magnetite and apatite (sample AR-05-104, Vedi area, see Rolland et al., accepted); (G) phyric
trachydacite with a hyalopilitic to cryptocrystalline texture (sample AR-04-75, Vedi valley, see Rolland et al., accepted). Calc-alkaline series: olivine-bearing, plagioclase phyric (15–40%) basalt
with a microcrystalline (plagioclase, quartz, opaque minerals) texture from pillow lavas suit (sample AR-04-32, Stepanavan area, Herher valley, see Galoyan et al., 2007), in which the olivine
phenocrysts are entirely pseudomorphosed to quartz and rims of iron oxides. From (A) to (C) under crossed nichols, and (D) to (H) under parallel nichols. Scale bar is for all photographs.
174
Y. Rolland et al. / Lithos 112 (2009) 163–187
andesine compositions are distributed in the fine-grained devitrified
groundmass made of albitic plagioclase, opaque microlites, and
carbonate-quartz-chlorite aggregates.
4.2.5. Calc-alkaline lavas of Stepanavan zone
They consist of large pillow-lavas of basaltic and basaltic andesitic
compositions with micro-cryptocrystalline (Fig. 5H) to intersertal textures
formed of large phenocrysts (2–7 mm) and microliths of andesineoligoclase plagioclase, and minor augite (Wo36–38En42–43Fs13–15) clinopyroxenes. These lavas overlie Upper Cretaceous limestones, unconformably lying on the ophiolite stricto sensu.
4.3.
40
Ar/39Ar dating
Complementary to previously published datings (Galoyan et al.,
2009) obtained from the ophiolite, which span the Middle Jurassic, we
provide here the first unambiguous dating of the alkaline suite. In the
Stepanavan and Sevan regions, the alkaline lavas have been deformed
and altered in the late collisional evolution, so that preserved and
unaltered amphibole-bearing lavas were only sampled in the Vedi
area.
Three analyses have been done on amphibole single grains from
two trachybasalt samples (AR-05-70 and AR-05-104) from the Vedi
ophiolite, which is described in Section 4.2 though only one is
considered fully successful. These datings are listed in Table 1 and the
successful one is presented in Fig. 7.
In the two datings for sample AR-05-70 (Table 1), the 39Ar content
was very low, so no plateau age can be calculated. However, from the
36
Ar/40Ar versus 39Ar/40Ar plots, it was possible to estimate isochron
ages, with large errors, of 108 ± 18 Ma (MSWD: 0.66) and 115 ± 37 Ma
(MSWD: 15). The very low 39Ar content is interpreted as a
consequence of the low K content of amphibole, which likely resulted
from pyroxene destabilisation in this sample.
In the dating of sample AR-05-104, a well-constrained plateau of
117.3 ± 0.9 Ma (2σ) was obtained, with 92% of released 39Ar (Fig. 7A).
The average 37ArCa/39ArK ratio is similar as the EPM value of the
amphibole from ~40 in low temperature steps, decreasing steadily to
~ 30 in high temperature steps (Fig. 7B). An isochron age of 117.5 ±
0.8 Ma (MSWD: 0.77) is obtained using the five steps of the plateau
age estimate (steps 3–7), with an initial 40Ar/36Ar ratio close to the
atmospheric value [(40Ar/36Ar)0 = 238 ± 4%; Fig. 7C]. Including the
step 1 of lower temperature we calculate a similar within-error
isochron age of 117.5 ± 0.8 Ma (2 σ).
The above age of 117.5 ± 0.8 Ma is more precise than the previous
ages determined by the whole-rock K–Ar method (Baghdasaryan
et al., 1988) which ranged between 114 and 97 Ma. We interpret
these younger K–Ar ages as resulting from alteration of the vitreous
matrix. These ~ 118 Ma Albian ages are also in agreement with age
estimates undertaken by Satian and Sarkisyan (2006) who provided
whole-rock Rb/Sr errorchron ages between 120 and 95 Ma. The
Alkaline sequence is thus undoubtedly younger than the ophiolite by
about 50 Ma, unlike some generally admitted views that all the
alkaline, calc-alkaline and tholeiitic series part of the obducted
sequence were formed in the same mid-oceanic context (Sokolov
1977; Knipper and Khain, 1980).
4.4. Major-trace-REE geochemistry
The geochemical analyses of the ophiolitic rocks from the Sevan
ophiolite are of relatively alkaline composition in comparison to
MORB. Major element data of pillow — lavas and plutonic rocks show
that they have predominantly basalt to trachybasalt compositions.
4.4.1. Major elements
Major element analysis of plutonic rocks ranges from gabbros to
granites (plagiogranites) with intermediate dioritic compositions
Fig. 7. 40Ar/39Ar amphibole dating results of trachybasalt sample AR-05-104 from the
Vedi alkaline suite.
(Fig. 8A). These magmatic rocks plot in a large domain comprised
between alkaline and tholeiitic tendencies of the TAS diagram (Le
Maitre et al., 1989). In the AFM diagram (Fig. 8B) most rocks lie close
to the limit between the tholeiitic and calc-alkaline fields.
1. Overall, the rocks of the ophiolitic suite are enriched in MgO and
more depleted in TiO2, K2O and P2O5 relative to the alkaline suite
(Figs. 8–10; Table 1). Compared to the plutonic rocks of the same
series, the volcanic rocks from the different areas plot in the same
compositional range (from basalts to andesites and trachyandesites) and are slightly Na2O richer.
2. The alkaline lavas from different zones plot in the same range,
varying compositionally from basanite-trachybasalt to basaltic
Y. Rolland et al. / Lithos 112 (2009) 163–187
175
Fig. 8. Plots of magmatic rocks (ophiolitic, alkaline and calc-alkaline series) in the (A) (Na2O + K2O) vs. SiO2 (Le Maitre et al., 1989) and (B) AFM (Irvine and Baragar, 1971) diagrams.
trachyandesite and trachyandesite, and are clearly in the calcalkaline/alkaline domain of the AFM diagram (Fig. 8A, B). One of
the most significant features of the alkaline lavas is their higher
TiO2, K2O and P2O5 contents.
3 The arc-type calk-alkaline lavas, have trachybasalt and basaltic
trachyandesite compositions in TAS diagram (Fig. 8A). They occupy
a transitional position between ophiolitic and alkaline domains in
Harker's diagram (Fig. 9), except lower TiO2 and higher Al2O3,
which depend on the abundance of plagioclase in such rocks.
Regarding the spread of compositional variations in major
elements within the series, it appears that only rough correlations
can be seen in the plots of SiO2 vs. other oxides (Fig. 9). Even the most
immobile elements during alteration processes, such as Al2O3, MgO
and TiO2 (e.g., Staudigel et al., 1996) do not show any clear correlations
(Fig. 9). In particular, Large Ion Lithophile Elements (LILE) such as Na
and K have scattered compositions, even in individual magmatic
suites. Such variations are ascribed to a combination of alteration and
magmatic processes. The occurrence of a long-lasting hydrothermal
event, ascribed to the slow-spreading oceanic environment is
indicated by scattered 40Ar/39Ar ages within individual gabbro
samples (Galoyan et al., 2009; Rolland et al., in press). Thus, we
ascribe the most important part of the elemental variability in the
ophiolitic suite (gabbros, diorites, plagiogranites and ophiolitic lavas)
to spilitization process in an oceanic environment while variations in
the alkaline and calc-alkaline volcanic rocks is ascribed to some
magmatic cause (variations in the source components, as is highlighted by the isotopic compositions, Section 4.5).
Thin section observations and previous studies of the Armenian
ophiolites (e.g., Palandjyan, 1971; Abovyan, 1981; Ghazaryan, 1994)
have shown that the whole ophiolitic sequence apart from the
alkaline lavas has been affected by oceanic low-temperature
hydrothermal alteration events. These processes induced modification of the whole-rock chemistry, as revealed by the increase of LOI
(Table 1).
4.4.2. Trace elements
High field strength elements (HFSE) are not mobilized during
alteration and their contents reflect, without ambiguity, those of their
parental magma (Staudigel et al., 1996). Contents in these trace
elements confirm the presence of three clearly distinct magmatic
suites, as defined in the previous section.
1. Basalts and gabbros of the ophiolite suite show strong enrichments
in LILE (Large Ion Lithophile Elements: Ba, Rb, K and Th), up to ten
times MORB values. They have negative anomalies in Nb–Ta and Ti
(Fig. 11A, B), which is generally indicative of volcanic island arc
environments (e.g. Taylor and McLennan 1985; Plank and Langmuir, 1998). However, trace element contents remain low relative
to volcanic arc lavas, which indicates a setting with little fractional
crystallization, in agreement with a back-arc setting (e.g., Galoyan
et al., 2009).
2. Overall, the concentrations of each element in the alkaline basalts
largely exceed the concentrations in the basalts from ophiolitic
series (Fig. 11C). Moreover, alkaline series basalts are characterized
by high abundances of LILE, high field strength elements (Nb, Ta, Zr
and Ti) and light rare-earth elements (LREE).
3 The calc-alkaline suite rocks show strong depletions in Nb and Ta,
relative to Th and La, and slight Ti negative anomaly (Fig. 11D). They
globally show slightly stronger enrichments in LREE and LILE than
the ophiolite suite rocks.
These differences in normalized element patterns support that
these basalts are not petrogenetically related and were most likely
derived from melts formed in different tectonic settings: (1) A backarc setting with slow-spreading rates, (2) Ocean-island within-plate
setting and (3) volcanic island arc.
Differences within the different suites can be related to magmatic
processes such as fractional crystallization and magma mixing or to
alteration processes. To analyse the importance of alteration, some
trace and two major elements are plotted versus Zr and Th (Fig. 10). Zr
is an incompatible element (for basaltic to andesitic lavas) wellknown to remain stable during alteration or weathering processes, so
it was used as reference to test the mobility of the other trace elements
(Fig. 10). The trace element composition of both ophiolitic series and
arc-type lavas plots in a restricted range of values for Zr (0–127 ppm),
while alkaline lavas with higher Zr are characterized by a large
compositional range (131–411 ppm; Table 2). This compositional
spread is ascribed to various levels of fractional crystallization with
one trachydacite sample having very high Zr content (681 ppm).
The concentration of zirconium normally increases in response to
magmatic processes such as fractional crystallisation, except for the
most differentiated lavas in which it has fractionated. Enrichment in Zr
is positively correlated with that of major elements as Ti, but there is
no clear correlation with SiO2, which are ascribed to a slight
176
Y. Rolland et al. / Lithos 112 (2009) 163–187
Fig. 9. Harker variation diagrams showing the compositions of the three (ophiolitic, alkaline and calc-alkaline) series.
fractionation of Si within each series. Compositional variations in SiO2
and TiO2 in alkaline lavas correlate well with variations in Zr (Fig. 10),
while traces and REE such as Nd and Th contents show slight positive
correlation with Zr contents, which are ascribed to fractionation of
these elements. Such process is also shown by the composition of
ultramafic cumulative plutonic rocks (websterite from Stepanavan
and gabbronorite from Sevan), which have lower trace element
contents than associated lavas due to their cumulative origin. In
contrast, Sr does not correlate well with Zr, and part of scattering of Sr
data may be due to its mobility due to alteration processes of
plagioclase. This is confirmed by a similar mobility of Ca, as shown in
the Ti vs. Ca diagram (Fig. 9).
4.4.3. REE geochemistry
In the chondrite-normalized rare earth element (REE) diagrams,
analysed ophiolite basalts and gabbros have flat and parallel REE
spectra in chondrite-normalized plots [(La/Yb)N = 0.6–0.9], showing
some slight depletions in LREE and a slight enrichment in MREE
(Fig. 11E, F). No extensive Eu anomalies were observed (Eu/
Eu⁎ = 0.95–1.15), which show that plagioclase has remained almost
stagnant, and is enriched in the final liquid. The concentration of REE
for volcanic rocks varies from 8 to 30 times chondrite and for gabbros
varies between 1 and 15 times chondrite, for exception a flaser gabbro
— 60 times (sample AR-03-25). These features are interpreted as a
result of extreme crystal fractionation involving plagioclase, clinopyroxene, orthopyroxene and, to a lesser extent, to olivine accumulation
(Pallister and Knight, 1981).
The websterite and gabbronorite have the lowest concentrations of
REE (0.1–0.9 and 1–5 times chondrite respectively) with patterns
characterized by depletion in LREE (Fig. 11F). One hornblende gabbro
(sample AR-04-45D from Stepanavan ophiolite) is characterized by
LREE enrichment ((La/Yb)N = 2.27) and some depletion in MREE (a
convex downward pattern) with smaller positive Eu anomalies (Eu/
Eu⁎ = 1.12).
Y. Rolland et al. / Lithos 112 (2009) 163–187
The diorites REE patterns (6–20 times chondrite) and plagiogranites are parallel to those of the gabbros, with smaller enrichment in
LREE ((La/Yb)N = 1.1). The most differentiated plagiogranite (sample
AR-04-44 from Stepanavan) is characterized by more depletion in the
middle to heavy REE compared to other plagiogranites, and strongly
positive Eu anomalies (Eu/Eu⁎ = 4.05) ascribed to high plagioclase
contents. Such features indicate a cumulative effect of plagioclase.
In contrast, chondrite-normalized REE patterns of alkaline lavas
(Fig. 11G) show huge LREE enrichments and HREE depletions [(La/
Yb)N = 6–14], being representative of intraplate continental basalts, as
compared to ophiolite lavas. Meanwhile, no extensive Eu anomalies are
observed (Eu/Eu⁎ = 0.95–1.15). One trachydacite sample (AR-04-75) is
featured by significant enrichments in trace elements, which is
explained by a high degree of fractional crystallization, as its REE
pattern is parallel to those of the basanite–trachyandesite series.
177
Chondrite-normalized REE patterns of calc-alkaline lavas are
strongly parallel and form a narrow domain (Fig. 11H). They have
similar HREE contents as volcanics of previous series with significantly
more depleted LREE contents than alkaline series rocks [(La/Yb)N =
2.1–2.3].
These differences in trace elements contents between the three
studied series further support that these basalts are petrogenetically
unrelated and, most likely derived from melts formed in different
tectonic settings.
4.5. Nd, Sr, Pb isotope geochemistry
4.5.1. Ophiolite series
Initial ɛNdi values of the ophiolitic lavas from the different studied
zones range from + 5.9 to +9.5 (Table 3) intermediate between
Fig. 10. Major and trace elements vs. Zr and Th diagrams. Major and trace elements are chosen to investigate the effects of alteration on the Sr, Nd and Pb isotopic systems;
explanations in the text.
178
Y. Rolland et al. / Lithos 112 (2009) 163–187
The initial Pb isotopic ratios in ophiolitic rocks range from 37.273 to
38.234 for 208Pb/204Pb, from 15.459 to 15.541 for 207Pb/204Pb and
from 17.630 to 18.459 for 206Pb/204Pb. In the Pb–Pb isotope diagrams
both ophiolitic volcanic and plutonic rocks plot on or close the MORB
domain (Fig. 12B).
4.5.2. Alkaline series
Overall, alkaline lavas show, in comparison to ophiolite rocks,
lower initial ɛNdi values ranging from +2.1 to + 4.0 but have a similar
range of (87Sr/86Sr)i ratios. In (143Nd/144Nd)i vs. (87Sr/86Sr)i diagram
(Fig. 12A) most of these rocks are located in the OIB field. In the
(208Pb/204Pb)i vs. (206Pb/204Pb)i, and (207Pb/204Pb)i vs. (206Pb/204Pb)i
diagrams these samples overlap various specific OIB provinces such
as Kerguelen, Samoa and Society and Marquises (Fig. 12B).
4.5.3. Calc-alkaline series
For the calc-alkaline lavas from Stepanavan, the initial ɛNdi values
and the (87Sr/86Sr)i ratios range from +3.8 to + 5.9 and from 0.7045
to 0.7048, respectively. Thus ɛNdi ratios appear to be intermediate
between those of ophiolitic and alkaline rocks, and Sr isotopic ratios
plot in the same range as these two series (Fig. 12A).
5. Discussion
Fig. 10 (continued).
typical MORBs and OIBs values and indicate a source region that
experienced long-term depletion in LREE (Fig. 11E). The initial (87Sr/
86
Sr)i ratios range from 0.7037 to 0.70565 for these lavas, which are
significantly too high for typical tholeiitic MORB lavas and thus, not
considered as the primary magmatic Sr isotope signature of these
rocks. The gabbros have ɛNdi (+4.3 to + 6.5) and initial Sr isotopic
values (0.70386 to 0.70557) in the same range as volcanic rocks.
In the Nd–Sr isotope diagram both ophiolitic volcanic and
plutonic rocks exhibit a significant increase of the initial Sr ratios
relatively to MORB (Fig. 12A). Moreover, this increase of (87Sr/86Sr) i
ratios positively correlates with Sr, Ba, Rb and K2O contents. This
shift towards 87Sr/86Sr radiogenic ratios is commonly attributed to
exchange between rocks and seawater during oceanic crust
hydrothermal alteration (e.g. McCulloch et al., 1981; Kawahata
et al., 2001).
Ophiolites of the Armenian Lesser Caucasus region are generally
separated into three distinct zones: (1) The Sevan-Akera zone in the
North (Knipper, 1975; Adamia et al., 1980), (2) The Zangezur zone in
the centre (Aslanyan and Satian, 1977; Knipper and Khain, 1980;
Adamia et al., 1981) and (3) The Vedi zone in the south (Knipper and
Sokolov, 1977; Zakariadze et al., 1983).
Due to the importance of Cenozoic volcanism spread over most of
Armenia (Fig. 2), it is still difficult to conclude only from geological
mapping whether the different ophiolites correlate with each other,
or if they represent various suture zones delimitating several
continental micro-blocks. For this reason, we have undertaken field
investigations in various ophiolites: Stepanavan (Galoyan et al.,
2007) and Sevan (Galoyan et al., 2009), along the northern rim of
Armenia; and the Vedi ophiolite, in the centre of Armenia. As
emphasized in the following discussion, the use of Nd, Sr and Pb
isotopes in complement to conventional major and trace element
data allow us to correlate the ophiolites with each other. These
ophiolites show some similarities and differences in their structure
and lithological successions, but these features remain compatible
with a single oceanic domain origin. This domain opened in the
Lower-Middle Jurassic and has undergone several phases of magmatic emplacement, for which we find evidence in each of the
geographic zones investigated. These correlations provide insight
into the evolution of the Tethyan domain, and in particular allow us
to propose a geodynamical model for the obduction of the ophiolite
over the Armenian block.
In the following discussion, we will evaluate the following points:
1. Petrographically and geochemically, the Armenian ophiolites are
similar to island-arc tholeiites. Such geochemical features are
typical for oceanic crust, formed in a back-arc setting with melting
of a shallow asthenospheric source contaminated by slab-derived
fluids (Saunders and Tarney, 1984). Such a hypothesis has already
been proposed for ophiolitic gabbros from Turkey (Kocak et al.,
2005), but has to be evaluated considering isotopic compositions
and partial melting constraints.
2. Alkaline lavas of variable thickness have covered this ophiolitic
sequence. Their origin has to be considered. (i) Do they also derive
from the same ophiolitic series? (ii) Did they form in an island-arc
setting or (iii) in an oceanic island/plateau environment? The
source of alkaline lavas will be discussed below regarding the Sr–
Nd–Pb isotopic data. The occurrence of alkaline magmatism prior
Y. Rolland et al. / Lithos 112 (2009) 163–187
179
Fig. 11. Trace and REE plots of the three studied magmatic suites. The multi-element spider diagrams are normalized to the N-MORB values of Sun and McDonough (1989), and REE
plots are normalized to the Chondrite values of Evensen et al. (1978). Patterns for the studied magmatic rocks: ophiolitic volcanic (A, E) and plutonic (B, F) series; OIB type alkaline
series (C, G), and arc type calk-alkaline series (D, H).
180
Table 3
Sr, Nd and Pb isotopic analyses of samples from ophiolitic complexes of Armenia.
Sevan
(206Pb/204Pb)
2σ
(238U/204Pb)
(206Pb/204Pb)i
(207Pb/204Pb)
2σ
(235U/204Pb)
(207Pb/204Pb)i
(208Pb/204Pb)
2σ
(232Th/204Pb)
(208Pb/204Pb)i
143
Nd/144Nd
2σ
(147Sm/144Nd)
(143Nd/144Nd)i
ɛNdi
(87Sr/86Sr)
2σ
(87Rb/86Sr)
(87Sr/86Sr)i
ɛSri
Stepanavan
165
165
AR-03-25
18.2577
0.00049
11.1328
17.9683
15.5033
0.00055
0.0818
15.4888
37.9444
0.0018
8.8411
37.8713
0.512957
0.000011
0.18522
0.51276
6.47
0.704433
0.000011
0.1326
0.70412
− 2.61
165
Vedi
165
165
165
117
117
117
165
117
117
95
95
165
165
165
117
117
117
AR-03-24 AR-03-10
AR-04-218
AR-03-02
G154
G142
AR-05-80
AR-04-44
AR-04-30
AR-03-53
AR-04-05
AR-04-32
AR-04-32
AR-05-113
AR-05-114
AR-05-106
AR-05-104
AR-05-78
AR-04-75
17.9321
0.00084
10.2463
17.6658
15.5344
0.00084
0.0753
15.5211
37.8028
0.0025
8.4704
37.7328
0.512957
0
0.29064
0.51264
4.25
0.70401
0.00001
0.0058
0.70399
−4.39
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
0.512911
0.000044
0.21337
0.51268
4.98
0.703875
0.000053
0.008
0.70386
− 6.39
0.512848
0.000039
1.80575
0.5109
− 29.80
0.703134
0.000048
1.2024
0.70031
− 56.69
0.512929
0.000039
0.16855
0.51275
6.27
0.704589
0.000048
0.064
0.70444
1.88
0.51275
3.4E− 05
0.11069
0.51263
4.04
0.70471
4.6E− 05
0.0184
0.70466
5.12
0.51315
0.00003
0.21883
0.51291
9.53
0.706076
0.000039
0.3765
0.70519
12.59
0.51282
3.7E− 05
0.17125
0.51271
3.8
0.70485
4.8E− 05
0.0534
0.70478
5.54
0.5129
4E−05
0.1668
0.5128
5.87
0.7048
4E−05
0.1725
0.7045
1.91
17.8966
0.00057
10.274
17.6296
15.5524
0.00065
0.0755
15.5391
37.7996
0.0024
63.7004
37.2729
0.512989
0.000042
0.22032
0.51275
6.35
0.705745
0.000013
0.0754
0.70557
17.92
18.0686
0.00065
5.7363
17.9195
15.4661
0.00064
0.0421
15.4587
37.9304
0.0017
63.5101
37.4053
0.512965
0.000007
0.22032
0.51273
5.88
0.70509
0.00001
0.0754
0.70491
8.62
18.4101
0.00051
10.6564
18.1073
15.497
0.00049
0.0783
15.4897
38.1116
0.0015
64.1466
37.5546
0.513007
0.00006
0.21848
0.51277
6.74
0.70591
0.00001
0.1087
0.70565
19.15
20.6925
0.00076
62.4439
19.6744
15.6653
0.0005
0.4587
15.6158
40.6747
0.0016
252.2401
39.3623
0.512709
0.00008
0.1279
0.51262
2.3
0.70453
0.000009
0.0327
0.70448
1.47
19.2867
0.00047
26.8011
18.8497
15.5651
0.00055
0.1968
15.5439
39.2226
0.0021
89.804
38.7553
0.512813
0.000011
0.15218
0.51271
4.01
0.704272
0.00001
0.6009
0.70338
−14.11
n.d.
n.d.
19.2791
0.00051
34.3464
18.719
15.5926
0.00035
0.2523
15.5654
39.6069
0.0013
138.024
38.8887
0.51275
6E− 06
0.13822
0.51265
2.86
0.70568
9E− 06
0.1109
0.70551
16.15
n.d.
n.d.
19.2349
0.00068
36.8567
18.6339
15.5901
0.00061
0.2707
15.5609
39.5224
0.002
155.9984
38.7107
0.512696
0.00001
0.12767
0.51261
2.05
0.705421
0.000011
0.2974
0.70498
8.57
n.d.
n.d.
18.61
0.00092
12.7305
18.4024
15.5803
0.00123
0.0935
15.5702
38.6686
0.0034
56.9259
38.3724
0.512773
0.000008
0.11619
0.51269
3.7
0.704072
0.000011
0.392
0.70349
− 12.56
n.d.
n.d.
19.1457
0.00046
18.2699
18.8478
15.5681
0.00061
0.1342
15.5536
39.1236
0.002
54.4859
38.8401
0.512754
0.000008
0.12176
0.51267
3.26
0.70636
0.0001
0.6276
0.70543
14.97
n.d.
n.d.
18.8172
0.00062
13.7665
18.4594
15.559
0.00061
0.1011
15.5411
38.4167
0.0025
22.1422
38.2336
0.512966
0.000038
0.20223
0.51275
6.28
0.705061
0.000012
0.5785
0.7037
−8.55
n.d.
n.d.
0.512765
0.000037
0.12781
0.51268
3.39
0.705331
0.00005
0.7208
0.70427
− 1.59
Notes: isotopic data (2σ error) are corrected for in situ decay assuming a mean age of 165 Ma of the ophiolite (Galoyan et al., 2009), 117 Ma for the alkaline series (this paper) and 95 Ma for the calc-alkaline series from palaeontogical and Ar-Ar
dating (Rolland et al., in press).
i: initial ratios calculated at 165, 104 and 95 Ma respectively. εNdi calculated with actual (143Nd/144Nd)CHUR = 0.512638 and (147Sm/144Nd)CHUR = 0.1967 (Wasserburg et al., 1981). εSri calculated with actual (87Sr/86Sr)CHUR = 0.7045 and (87Rb/86Sr)
CHUR = 0.0827 (Wasserburg et al., 1981).
Pb isotopic ratios measured with external precision of ca. 250-300 ppm for the 206, 207, 208Pb/204Pb ratios.
Y. Rolland et al. / Lithos 112 (2009) 163–187
Locality
age (Ma)
sample
Y. Rolland et al. / Lithos 112 (2009) 163–187
Fig. 12. Plots of the magmatic rocks in (A) Sr–Nd; (B)
Northern Hemisphere Reference Line (Hart, 1984).
206
Pb/204Pb vs.
208
Pb/204Pb and
to obduction in the late Lower Cretaceous may be of importance for
the obduction model of the ophiolite crustal sequence.
3. Finally, the calk-alkaline lavas are Upper Cretaceous in age. These
volcanic arc-related series likely formed during closure of the NeoTethys ocean. Their geochemical features will be considered to
evaluate this hypothesis.
5.1. Significance of Armenian ophiolites: MOR or back-arc setting?
Armenian ophiolitic series are shown to be of slight alkaline to
tholeiitic character, ranging from basalts to basaltic andesites and basaltic
trachyandesites. Spider diagrams show clear Nb–Ta negative anomalies
(Fig. 11A, B), LILE enrichments and flat to slightly LREE-enriched spectra.
Their isotopic compositions are significantly more radiogenic in 87Sr/86Sr
and slightly less radiogenic in Nd isotopes than typical MORB compositions (Fig. 12A). These observations do not support a geochemical
“normal” ophiolitic crust and are more probably in agreement with
typical volcanic arc settings, in which enrichments in LILE, LREE result
from slab fluids/melts contaminations (Pearce et al., 1984).
5.1.1. Source components
The isotopic compositions of the ophiolitic magmatic rocks lie at
the limit of the MORB domain and overlap the OIB field (Fig. 12A, B).
Nevertheless, their flat REE spectra together with their “enriched”
isotopic character suggest partial melting from a spinel-bearing
mantle with small percent partial melts similar as to those known
for slow spreading ridges (Lagabrielle, 1987). This hypothesis will be
tested by a non-modal batch partial melting model in the following
section. Further, the measured Nb–Ta negative anomalies combined
with slight LILE enrichments are indicative of a volcanic arc setting.
Finally, the emplacement depth of pillow lava flows was clearly
abyssal as shown by the deposition of radiolarite interlayers. There-
181
207
Pb/204Pb isotopic diagrams, with the fields of MORB and of various OIB contexts. NHRL,
fore, it is likely that the ophiolitic rocks were produced by melting of a
depleted mantle source contaminated by hydrothermal slab-derived
fluids in a back-arc basin environment.
To estimate the level of slab-derived contamination in the
formation of the ophiolitic rocks, mixing curves have been drawn on
Fig. 13 between different components: depleted mantle pole (MORB),
Enriched Mantle 1 and Enriched Mantle 2 [EM1 and EM2, respectively; Zindler and Hart (1986), Salters and White (1998) and Hanan
et al. (2000)]. This isotopic modelling suggests that basaltic ophiolite
lava composition results from contamination of a typical MORB by a
mixed source composed of 1–4% EM2 and 2–5% EM1. Such degrees of
contamination appear to be relatively high in a back-arc setting, and
suggest the participation of subducted slab sediments in the source
(EM1) and a possible fertile E-MORB type source (EM2).
5.1.2. Partial melting estimates
Fig. 14 shows calculated REE composition for melts produced by a
depleted mantle type source (MORB) partial melting (Fig. 14A) and by
an enriched mantle source (EM2) partial melting (Fig. 14B). The REE
patterns of the ophiolitic suite fit those of melts produced by the partial
non-modal batch melting from 4% to 10% of a spinel-bearing mantle,
which mineralogical compositions spinel 5%, olivine 55%, orthopyroxene 20% and clinopyroxene 20%, using the partition coefficient factors
of McKenzie and O'Nions (1991), Johnson (1994), and Nikogosian and
Sobolev (1997). Such composition is rather similar to that of a slightly
depleted mantle source (e.g., Juteau and Maury, 1997).
5.2. Origin of alkaline lavas: source components and geodynamic
significance?
Mineral chemistry and geochemistry of the alkaline volcanic series
of Sevan, Stepanavan and Vedi ophiolites is similar to that of OIBs. As
182
Y. Rolland et al. / Lithos 112 (2009) 163–187
Fig. 13. Isotopic diagrams showing mixing curves between the different mantle end-members. 143Nd/144Nd vs. 206Pb/204Pb (A) and 87Sr/86Sr (B) isotopic diagrams; and 206Pb/204Pb
vs. 87Sr/86Sr (C) and 207Pb/204Pb (B). Compositions of end-members used in the calculation of mixing curves (after Hart, 1984; Zindler and Hart, 1986; Sun and McDonough, 1989;
Eisele et al., 2002) are the following. HIMU: (87Sr/86Sr) = 0,703; (143Nd/144Nd) = 0,51285; (206Pb/204Pb) = 21,5; (207Pb/204Pb) = 15,82; (208Pb/204Pb) = 40; [Sr] = 120; [Nd]A = 6,5;
[Pb] = 0,4. EM2 pole: (87Sr/86Sr) = 0,71682; (143Nd/144Nd) = 0,51216; (206Pb/204Pb) = 18,99; (207Pb/204Pb) = 15,65; (208Pb/204Pb) = 39,5; [Sr] = 218; [Nd] = 34; [Pb] = 25. EM1
pole: (87Sr/86Sr) = 0,705; (143Nd/144Nd) = 0,5122; (206Pb/204Pb) = 16,8; (207Pb/204Pb) = 15,45; [Sr] = 513; [Nd] = 33; [Pb] = 3,5; DMM: (87Sr/86Sr) = 0,7022; (143Nd/144Nd) =
0,513075; (206Pb/204Pb) = 17,3; (207Pb/204Pb) = 15,4; (208Pb/204Pb) = 37,5; [Sr] = 11,3; [Nd] = 1,12; [Pb] = 0,0489. The mixing curves equations are from Faure (1986).
Fig. 14. Modelling effect of non-modal batch melting of several mantle sources, and comparison with ophiolite and alkaline REE spectra compositional domains obtained in this study.
Calculated composition (A) of the depleted peridotite ‘ophiolite’ source: olivine 55%, orthopyroxene 20%, clinopyroxene 20% and spinel 5%; and (B) of the enriched peridotite
‘alkaline’ source: olivine 54%, orthopyroxene 20%, clinopyroxene 20%, garnet 2% and spinel 4% (Salters and Stracke, 2004).
Y. Rolland et al. / Lithos 112 (2009) 163–187
Fig. 15. Geodynamic reconstitution of the Lesser Caucasus in the Middle Jurassic to Upper Cretaceous periods.
183
184
Y. Rolland et al. / Lithos 112 (2009) 163–187
shown in the Mineral Chemistry section, pyroxenes are slightly
alkaline. The alkaline lava samples (Fig. 14) show strong enrichments in incompatible elements (up to 100 times chondrite
values). Partial melting degree calculation suggests that they may
be derived from ~ 20% non-modal batch melting of an enriched
spinel-garnet-bearing mantle source characterized by La and Lu
concentrations 3.3 times and 1.8 times the chondrite mantle,
respectively. The isotopic composition of these lavas plots in the
field of OIBs in agreement with an enriched alkaline mantle source.
In the isotopic plots of Fig. 13, the obtained isotopic data plots
mainly on the DMM-EM2 mixing curve, with ~ 2–3% of EM2, ~ 5–
15% HIMU and almost no EM1 contamination which suggests that
subduction-derived contamination may not be envisaged. The
isotopic compositions of the studied Armenian alkaline series are
thus rather in agreement with an OIB-type source. In the Vedi area,
Satian et al. (2005) already pointed out the alkaline character of
the lava series, which they interpreted as volcanic series formed in
an intra-continental rift. These lavas were emplaced above, and
formed ~ 50 Ma after the ophiolites. Moreover, they are interstratified and overlain by shallow marine reef limestone. Thus, all
these features are in agreement with a plume event that occurred
in an intra-oceanic setting.
Such alkaline magmatism is widely documented in the MiddleEast region, along the Arabian and Indian platforms, in relationship
with the formation of the Neo-Tethys ocean (e.g. Lapierre et al., 2004).
Similar Cretaceous alkaline series are found above the Iranian
ophiolite (Ghazi and Hassanipak, 1999), and in Turkey (Norman,
1984; Tüysüz et al., 1995; Tankut et al., 1998; Okay, 2000). However, it
is still difficult to relate these alkaline events due to their geographical
and temporal distance and to paucity of radio-chronological and Sr,
Nd, Pb isotopic data.
5.3. Reconstruction of the ‘ophiolite’ history
From all the available geological data, we propose the following
model for the evolution of the Armenian Ophiolite (Fig. 15):
1. The SAB is of Gondwanian origin according to lithological
associations found in central and SE Armenia (Knipper and
Khain, 1980; Kazmin et al., 1987; Aghamalyan, 2004). Therefore,
it is likely that the Sevan oceanic basin opened in response to the Ndipping subduction of Neotethys to the south of Eurasia (Fig. 15—
stage 1). The continuation of Paleotethys in the area is a matter of
debate, as the westward continuation of the Cimmeride orogenic
system is not identified in the Lesser Caucasus (e.g., Sengör, 1984,
1990). Emplacement of the ophiolite occurred in the Lower-to
Middle Jurassic (Galoyan et al., 2009). The older age of the Vedi
ophiolite (178.7 ±2.6 Ma; Rolland et al., accepted), with respect to
that of Sevan (160–165 Ma, Zakariadze et al., 1990; Galoyan et al.,
2009) implies that it was at the southern rim of the back-arc
system. The structural setting of the ophiolite obduction indicates
clearly that oceanic crust of the back-arc basin was formed between
the SAB and the Eurasian active margin.
2. Emplacement of an Oceanic Island/Plateau above the back-arc oceanic
crust during the late Lower Cretaceous (40Ar/39Ar age of 117.5±0.8 Ma
in this paper; Fig. 15—Stage 2).
3. The calc-alkaline lavas disconformably overlie the ophiolite and
related alkaline series (Galoyan et al., 2007). These lavas have
similar geochemical features as volcanic arc series, including the
isotopic Sr–Nd composition. Their emplacement is bracketed in the
Upper Cretaceous, as for the high pressure metamorphism
constrained in the Stepanavan area (Meliksetyan et al., 1984 and
references therein), constrained at about 95–90 Ma (Rolland et al.,
2009). Therefore this magmatic event can be related to the
subduction of Neo-Tethys ocean prior to the obduction of the
Armenian Ophiolites onto the SAB.
4. After this, the SAB enters the subduction zone in the Turonian (95–
88 Ma), which triggers a “collision” with the thickened oceanic
crust. During this process, part of the volcanic arc has probably
been subducted below the obducted oceanic sequence and metamorphosed in the blueschist facies (Rolland et al., 2009). The
large variety of lithologies comprising metabasites, marls and
conglomerates in a pelitic matrix, within the Stepanavan blueschists, is in agreement with such a scenario.
5. The obduction of the ‘ophiolite’ section over the SAB is further
constrained by the Lower Coniacian frontal flysch sequences,
found below and in front of the Vedi obducted sequence. The calcalkaline series found above the Stepanavan ophiolite show that
a volcanic arc was active during this time above the obducted
sequence.
6. The end of the obduction is constrained by Upper Coniacian fauna
in sediments unconformably overlying the ophiolite. Blocking of
the subduction below the Eurasian margin may stop at 73–71 Ma,
as shown by Ar–Ar age of MT-LP metamorphism in the Stepanavan
blueschists and the general tectonic uplifting of the region,
witnessed by erosion and absence of sedimentary record during
the Upper Cretaceous–Paleocene (Rolland et al., 2009). This 73–
71 Ma event is thus interpreted as the insight of ‘collision’.
5.4. Implications of a plateau/OIB event on the ophiolite obduction?
The alkaline series show features of Plume-related magmas. What
is the significance of such plume magmatism, and what is its
consequence with the obduction and preservation of the Armenian
ophiolites?
Only 3% of the current oceanic floor is composed of plumerelated crust, of which oceanic plateaus are the largest part
(Petterson et al., 1997). In Armenia, this alkaline event seems to
be relatively large due to the presence of alkaline lavas over the
ophiolite in all the studied sections. This large size is thus in
agreement with an oceanic plateau event, in which large volumes of
lavas are erupted during volcanic emplacement in a small time
range.
Only several examples of obducted plateau series have been
claimed worldwide, amongst which the Wrangellia terrane of Alaska
and British Columbia (e.g. Richards et al., 1991), and Gorgona Island in
Columbia (Duncan and Hargreaves, 1984; Storey et al., 1991), but such
examples remain relatively uncommon. The paucity of obducted
plateau sequences may be explained by the fact that they are not easily
recognized in the geological records. However, their potential in the
blocking or reversion of polarity of subduction zones has been noted
in numerous cases (e.g. Petterson et al., 1997; Kerr et al., 2003; Kerr
and Mahoney, 2007). For instance large oceanic plateaus can cause the
reversal of subduction polarity, as did the Ontong Java Plateau
(Coleman and Kroenke, 1981). Cloos (1993) calculated that basaltdominated oceanic plateau crust must exceed 17 km thickness to
survive subduction, and about 30 km to cause any significant
‘collisional’-type deformation.
The Armenian ophiolites show evidence for the obduction of a
single oceanic crust sequence above the SAB, as similar geological,
petrological, geochemical and age features are found in the three
studied Armenian ophiolitic massifs (Sevan, Stepanavan, and Vedi).
The oceanic crust s.s. corresponds to a slow-spreading ophiolite
formed in the Lower-Middle Jurassic in a back-arc basin by 4–10%
melting of a shallow asthenosphere spinel-bearing source contaminated by subducted slab-derived products. Alkaline volcanic
series with OIB-type geochemical features are found above the
ophiolite sequence in each of the studied areas, which late Lower
Cretaceous age has been constrained above by the 40Ar/39Ar
method on amphibole at 117.3 ± 0.9 Ma. Therefore, this Armenian
‘plume event’ shortly predates the Coniacian–Santonian (88–
83 Ma) obduction of the Armenian ophiolitic sequence. Therefore,
Y. Rolland et al. / Lithos 112 (2009) 163–187
the thickened and hot oceanic plate had a low density when it
overrode the SAB continental margin, which suggests that the
plume event has likely played an important role in the obduction
process.
The original width and thickness of the Armenian plume related
series are difficult to assess. However, this volcanic series covered
each of the studied Armenian ophiolites, which suggests a N104 km2
surface regarding the initial ophiolite surface prior to horizontal
shortening, and N105–106 km2 if they can be correlated to similar
settings in Turkey and Iran. Further, these lavas have significantly
more radiogenic lead isotopic compositions than ophiolitic rocks,
and are related to ~ 20% melting of an enriched garnet-spinelbearing source. These high partial melting estimates are in the range
of plateau events melting estimates (14–26%, Hauff et al., 1997;
Révillon et al., 2000; Herzberg and O'Hara, 2002; Kerr et al., 2002).
The melting of a garnet-bearing source suggests the contribution of
a deep mantle source below a thick lithosphere sequence or the
melting of a relatively thick oceanic crust (N30 km, as is constrained
by the stability of garnet in metabasites; Rapp et al., 1991). Such
thickness has been estimated for the Ontong Java plateau (e.g.,
Gladczenko et al., 1997; Richardson et al., 2000; Miura et al., 2004).
Expected mantle lithosphere “root”, below such thick oceanic crust
could be N300 km (Richardson et al., 2000). Therefore, the
Armenian hot-spot related magmatism likely features the formation
of an Oceanic Plateau or a large oceanic island, with significant
crustal thickening. Further, the age of the Armenian plume-related
series is within range of a period of major oceanic plateau formation
in the late Lower Cretaceous; fitting precisely the age of formation of
one of the largest plateaus, the Ontong Java plateau (Tarnudo et al.,
1991). These alkaline series are also locally covered by an arcderived calc-alkaline volcanic sequence, which was likely formed in
a supra-subduction zone environment. Further evidence of this
subduction is provided by blueschists series dated at 95–90 Ma
(Rolland et al., 2009).
Therefore both oceanic Plateau and volcanic arc formations shortly
pre-dated the obduction, which occurred in the Coniacian–Santonian
(88–83 Ma; Sokolov, 1977). Crustal thickening related to plateau and
arc events are thought to have increased crustal buoyancy (e.g., Cloos,
1993; Abbot and Mooney, 1995; Abbot et al., 1997; Kerr and Mahoney,
2007). Such low buoyancy likely hindered subduction of the oceanic
crust and allowed it to be obducted over the SAB continental crust.
Such process is not unlikely in other obduction contexts of the PeriTethyan region, especially in the Caucasus–Middle East segment, but
the obducted ophiolite sections have to be analysed in detail to find if
whether alkaline series of similar age and geochemical signatures may
be present.
Acknowledgements
This work was supported by the Middle East Basins Evolution
project jointly supported by a consortium including oil companies
and the CNRS. Many thanks to the MEBE program coordinators E.
Barrier and M. Gaetani for their support and encouragements, and M.
F. Brunet for coordinating the project. Analytical data were acquired
with the help of the Geosciences Azur Laboratory, in which we thank
L. Vacher and J.P. Goudour for their involvement during data
acquisition. We also thank the support of the French Embassy at
Yerevan for the MAE PhD grant of G. Galoyan. The paper was
significantly enhanced by the work of two anonymous reviewers and
the editor N. Eby, while English language was improved by G. Nolet.
References
Abbot, D., Mooney, W., 1995. The structural and geochemical evolution of the
continental crust: support for the oceanic plateau model of continental growth.
Reviews of Geophysics 33, 231–242.
185
Abbot, D., Drury, R., Mooney, W., 1997. Continents as lithological icebergs: the
importance of buoyant lithospheric roots. Earth and Planetary Science Letters
149, 15–27.
Abovyan, S.B., 1981. The mafic–ultramafic complexes of the ophiolitic zones in
Armenian SSR. Izvestia Academy of Science Armenian SSR. 306 pp. (in Russian).
Adamia, S., Bergougnan, H., Fourquin, C., Haghipour, A., Lordkipanidze, M., Ozgül, N.,
Ricou, L., Zakariadze, G., 1980. The Alpine Middle East between the Aegean and the
Oman traverses. 26th International Geological Congress Paris C5, 122–136.
Adamia, S., Bergougnan, H., Fourquin, C., Haghipour, A., Lordkipanidze, M., Ozgül, N.,
Adamia, S., Chkhotua, T., Kekelia, M., Lordkipanidze, Shavishili, I., Zakariadze, G.,
1981. Tectonics of the Caucasus and the adjoining regions: implications for the
evolution of the Tethys ocean. Journal of Structural Geology 3, 437–447.
Agard, P., Jolivet, L., Vrielynck, B., Burov, E., Monié, P., 2007. Plate acceleration: the
obduction trigger? Earth and Planetary Science Letters 258, 428–441.
Aghamalyan, V.A., 2004. The Lesser Caucasus earth crust formation and evolution in the
collision zone of Palaeo-Tethys. 5th International Symposium on Eastern Mediterranean Geology, Proceedings, vol. 1, pp. 17–20. Thessaloniki, Greece.
Aslanyan, A.T., Satian, M.A., 1977. On the geological features of Transcaucasian ophiolitic
zones. Izvestia Academy of Science Armenian SSR 4–5, 13–26 (in Russian).
Aslanyan, A.T., Satian, M.A., 1982. Middle Cretaceous ophiolite zones of Transcaucasus
and tectonic reconstructions. Ofioliti 7, 131.
Avagyan, A., Sosson, M., Philip, M.H., Karakhanian, A., Rolland, Y., Melkonyan, R., Rebai,
S., Davtyan, V., 2005. Neogene to Quaternary stress field evolution in Lesser
Caucasus and adjacent regions using fault kinematics analysis and volcanic cluster
data. Geodinamica Acta 18, 401–416.
Baghdasaryan, G.P., Vardanyan, A.V., Satian, M.A., 1988. On the age of the volcanic rocks of
ophiolitic association. Izvestia Academy of Science Armenian SSR 6, 11–18 (in Russian).
Ben-Avraham, Z., Nur, A., Jones, D., Cox, A., 1981. Continental accretion: from oceanic
plateaus to allochtonous terranes. Science 213, 47–54.
Bosch, D., Blichert-Toft, J., Moynier, F., Nelson, B.K., Telouk, P., Gillot, P.Y., Albarède, F.,
2008. Pb, Hf and Nd isotope compositions of the two Réunion volcanoes, Indian
Ocean: a tale of two small-scale mantle “blobs”. Earth and Planetary Science Letters
265, 748–768.
Boudier, F., Ceuleneer, G., Nicolas, A., 1988. Shear zones, thrusts and related magmatism
in the Oman ophiolite: initiation of thrusting on an oceanic ridge. Tectonophysics
151, 275–296.
Brynzia, L.T., Wood, B.J., 1990. Oxygen thermobarometry of abyssal spinel peridotites;
the redox state and C–O–H volatile composition of the Earth's sub-oceanic upper
mantle. American Journal of Science 290, 1093–1116.
Cloos, M., 1993. Lithospheric buoyancy and collisional orogenesis: subduction of oceanic
plateaus, continental margins, island arcs, spreading ridges, and seamounts.
Geological Society of America Bulletin 105, 715–737.
Coleman, P.J., Kroenke, L.W., 1981. Subduction without volcanism in the Solomon
Islands Arc. Geo-Marine Letters 1, 129–134.
Danelian, T., Galoyan, G., Rolland, Y., Sosson, M., 2007. Palaeontological, (Radiolarian)
Late Jurassic age constraint for the Stepanavan ophiolite, Lesser Caucasus, Armenia.
Proceedings of the 11th International Congress, Athens, May 2007. Bulletin of the
Geological Society of Greece, vol. 37.
Danelian, T., Asatryan, G., Sosson, M., Person, A., Sahakyan, L., Galoyan, G., 2008. Discovery
of two distinct Middle Jurassic Radiolarian assemblages in the sedimentary cover of
the Vedi ophiolite, Lesser Caucasus, Armenia. Comptes Rendus Palevol 7, 327–334.
Dercourt, J., Zonenschain, L.P., Ricou, L.E., Kazmin, V.G., Le Pichon, X., Knipper, A.L.,
Grandjacquet, C., Sbortshikov, I.M., Geyssant, J., Lepvrier, C., Pechersky, D.H., Boulin,
J., Sibuet, J.C., Savostin, L.A., Sorokhtin, O., Westphal, M., Bazhenov, M.L., Lauer, J.P.,
Biju-Duval, B., 1986. Geological evolution of the tethys belt from the Atlantic to the
pamirs since the LIAS. Tectonophysics 123, 241–315.
Duncan, R.A., Hargreaves, R.B.,1984. Plate Tectonic Evolution of the Caribbean Region in the
Mantle Reference Frame. Geological Society of America Memoir 162, 81–93.
Eisele, J., Sharma, M., Galer, S.J.G., Blichert-Toft, J., Devey, C.W., Hofmann, A.W., 2002. The role
of sediment recycling in EM-1 inferred from Os, Pb, Hf, Nd, Sr isotope and trace element
systematics of the Pitcairn hotspot. Earth Planetary Science Letters 196, 197–212.
Evensen, N.M., Hamilton, P.J., O'Nios, R.K., 1978. Rare earth abundances in chondritic
meteorites. Geochimica et Cosmochimica Acta 42, 1199–1212.
Faure, G., 1986. Principles of isotope geology, 2nd edition. Wiley, New York.
Galoyan, G., Rolland, Y., Sosson, M., Corsini, M., Melkonyan, R., 2007. Evidence for
superposed MORB, oceanic plateau and volcanic arc series in the Lesser Caucasus,
Stepanavan, Armenia. Comptes Rendus Geosciences 339, 482–492.
Galoyan, G., Rolland, Y., Sosson, M., Corsini, M., Billo, S., Verati, C., Melkonyan, R., 2009.
Geochemistry and 40Ar/39Ar dating of Sevan Ophiolites, Lesser Caucasus, Armenia):
evidences for Jurassic Back-arc opening and hot spot event between the South
Armenian Block and Eurasia. Journal of Asian Earth Sciences 34, 135–153.
doi:10.1016/j.jseaes.2008.04.002.
Ghazaryan, H.A., 1987. Stratified gabbros of ophiolitic series of the south-eastern part of
the Sevan mountain range. Tipomorfizm and the Parageneses of the Minerals of
Armenian SSR. Izvestia Academy of Science Armenian SSR, pp. 122–139 (in Russian).
Ghazaryan, H.A., 1994. Pecularities of the geological structures and petrogenesis of
ophiolite gabbroids, on the examples of the Sevan and Vedi ophiolite zones of
Armenia. Izvestia Academy of Science Armenian SSR 3, 19–31 (in Russian).
Ghazi, A.M., Hassanipak, A.A., 1999. Geochemistry of subalkaline and alkaline extrusives
from the Kermanshah ophiolite, Zagros Suture Zone, Western Iran: implications for
Tethyan plate tectonics. Journal of Asian Earth Sciences 17, 319–332.
Gladczenko, T.P., Coffin, M.F., Eldholm, O., 1997. Crustal structure of the Ontong Java
Plateau: modeling of new gravity and existing seismic data. Journal of Geophysical
Research— Solid Earth 102, 22711–22729.
Guillot, S., Allemand, P., 2002. 2D thermal modelling of the early evolution of the
Himalayan belt. Journal of Geodynamics 34, 77–98.
186
Y. Rolland et al. / Lithos 112 (2009) 163–187
Guillot, S., Garzanti, E., Baratoux, D., Marquer, D., Mahéo, G., de Sigoyer, J., 2003.
Reconstructing the total shortening history of the NW Himalaya. Geochemistry
Geophysics Geosystems 4 (7), 1064. doi:10.1029/2002GC000484.
Hanan, B.D., Blichert-Toft, J., Kingsley, R., Schilling, J.G., 2000. Depleted Iceland mantle
plume geochemical signature: artifact of multi-component mixing? Geochemistry
Geophysics Geosystems 1, 1999GC000009.
Hart, S.R., 1984. A large-scale isotope anomaly in the Southern Hemisphere mantle.
Nature 309, 753–757.
Harutyunyan, G.S., 1967. The breakdown of the intrusions of the north-western part of the
Sevan ridge dependent on age. Izvestia Academy of Science Armenian SSR vol. 1–2,
42–52 (in Russian).
Hauff, F., Hoernle, K., Schmincke, H.U., Werner, R., 1997. A mid Cretaceous origin fort he
Galapagos hot-spot: volcanological, petrological and geochemical evidence from
Costa Rican oceanic crustal segments. Geologische Rundschau 86, 141–155.
Herzberg, C., O'Hara, M.J., 2002. Plume-associated ultramafic magmas of Phanerozoic
age. Journal of Petrology 43, 1857–1883.
Irvine, T.N., Baragar, W.R.A., 1971. A guide to the chemical classification of the common
volcanic rocks. Canadian Journal of Earth Sciences 8, 523–548.
Johnson, K.T.M., 1994. Experimental cpx/ and garnet/melt partitioning of REE and other
trace elements at high pressures; petrogenetic implications. Mineralogical
Magazine 58, 454–455.
Jourdan, F., Féraud, G., Bertrand, H., Kampunzu, A.B., Tshoso, G., Le Gall, B., Tiercelin, J.J.,
Capiez, P., 2004. The Karoo triple junction questioned: evidence for Jurassic and
Proterozoic 40Ar/39Ar ages and geochemistry of the giant Okavango dyke swarn,
Botswana. Earth and Planetary Science Letters 222, 989–2006.
Juteau, T., Maury, R., 1997. Géologie de la croûte océanique. Masson, Paris. 367 pp.
Kawahata, H., Nohara, M., Ishizuka, H., Hasebe, S., Chiba, H., 2001. Sr isotope
geochemistry and hydrothermal alteration of the Oman ophiolite. Journal of
Geophysical Research 106 (B6), 11083–11099.
Kazmin, V.G., Sbortshikov, I.M., Ricou, L.-E., Zonenshain, L.P., Boulin, J., Knipper, A.L., 1987.
Volcanic belt-indicators of the Mesozoic–Cenozoic active outskirts of Eurasia. In:
Monin, A.S., Zonenshain, L.P. (Eds.), History of the Tethys Ocean. Academy of Sciences
of the USSR, P.P. Shirshov Institute of Oceanology, Moscow, pp. 58–74 (in Russian).
Kerr, A.C., Mahoney, J.J., 2007. Oceanic plateaus: problematic plumes, potential
paradigms. Chemical Geology 241, 332–353.
Kerr, A.C., Tarney, J., Kempton, P.D., Spadea, P., Nivia, A., Marriner, G.F., Duncan, R.A.,
2002. Pervasive mantle plume head heterogeneity: evidence from the Late
Cretaceous Caribbean–Colombian oceanic plateau. Journal of Geophysical
Research–Solid Earth 107. doi:10.1029/2001JB000790.
Kerr, A.C., White, R.V., Thompson, P.M.E., Tarney, J., Saunders, A.D., 2003. No oceanic
plateau — no Caribbean Plate? The seminal role of an oceanic plateau. In: Car, C., Buffler,
R.T., Blickwede, J. (Eds.), The Gulf of Mexico and Caribbean Region: Hydrocarbon
habitats, Basin formation and Plate tectonics. AAPG Memoir, vol. 79, pp. 126–268.
Knipper, A.L., 1975. The oceanic crust in the alpine belt. Tr. GIN NAS USSR 267, 207
(in Russian).
Knipper, A.L., Khain, E.V., 1980. Structural position of ophiolites of the Caucasus. Ofioliti
Special Issue 2, 297–314.
Knipper, A.L., Sokolov, S.D., 1977. Vedi ophiolites (Armenia) autochthon or allochthon?
Geotektonics 10, 261–269.
Kocak, K., Isık, F., Arslan, M., Zedef, V., 2005. Petrological and source region
characteristics of ophiolitic hornblende gabbros from the Aksaray and Kayseri
regions, central Anatolian crystalline complex, Turkey. Journal of Asian Earth
Sciences 25, 883–891.
Lagabrielle, Y., 1987. Les ophiolites: marqueurs de l'histoire tectonique des domaines
océaniques. Thèse Doctorat d'Etat, Univ. Bretagne Occidentale, Brest, 350 p.
Lagabrielle, Y., Cannat, M., 1990. Alpine Jurassic ophiolites resemble the modern central
Atlantic basement. Geology 18, 319–322.
Lagabrielle, Y., Polino, R., Auzende, J.M., Blanchet, R., Caby, R., Fudral, S., Lemoine, M.,
Mevel, C., Ohnenstetter, M., Robert, D., Tricart, P., 1984. Les témoins d'une
tectonique intra-océanique dans le domaine téthysien: analyse des rapports entre
les ophiolites et leurs couvertures métasédimentaires dans la zone piémontaise des
Alpes franco-italiennes. Ofioliti 9, 67–88.
Lapierre, H., Samper, A., Bosch, D., Maury, R.C., Béchennec, F., Cotton, J., Demant, A.,
Brunet, P., Keller, F., Marcoux, J., 2004. The Tethyan plume: geochemical diversity of
Middle Permian basalts from the Oman rifted margin. Lithos 74, 167–198.
Leake, B.E., Woolley, A.R., Arps, C.E.S., Birch, W.D., Gilbert, M.C., Grice, J.D., Hawthorne, F.C.,
Kato, A., Kisch, H.J., Krivovichev, V.G., Linthout, K., Laird, J., Mandarino, J.A., Maresch, W.V.,
Nickel, E.H., Rock, N.M.S., Schumacher, J.C., Smith, D.C., Stephenson, N.C.N., Ungaretti, L.,
Whittaker, E.J.W., Youzhi, G., 1997. Nomenclature of amphiboles: report of the
subcommittee on amphiboles of the international mineralogical association, commission
on new minerals and mineral names. American Mineralogist 82, 1019–1037.
Le Maitre, R.W., Bateman, P., Dudek, A., Keller, J., Lameyre Le Bas, M.J., Sabine, P.A.,
Schmid, R., Sorensen, H., Streckeisen, A., Woolley, A.R., Zanettin, B., 1989. A
classification of igneous rocks and glossary of terms. Blackwell, Oxford.
Leterrier, J., Maury, R., Thonon, P., Girard, D., Marchal, M., 1982. Clinopyroxene
composition as a method of identification of the magmatic affinities of palaeovolcanic series. Earth and Planetary Science Letters 59, 139–154.
McCulloch, M.T., Gregory, R.T., Wasserburg, G.J., Taylor, J., 1981. Sm–Nd, Rb–Sr, and 18O/
16
O systematics in an oceanic crustal section: evidence from the Samail ophiolite.
Journal of Geophysical Research 86, 2721–2736.
McKenzie, D., O'Nions, R.K., 1991. Partial melt distribution from inversion of rare earth
element concentrations. Journal of Petrology 32, 1021–1091.
Meliksetyan, B.M., Baghdasaryan, G.P., Ghukasyan, R.Kh., 1984. Isotopic-geochemical
and geochronological investigations of eclogite–amphibolites associated with
ophiolites of Sevan–Amasian belt, Amasian Massif. Izvestia Academy of Science
Armenian SSR 1, 3–22 in Russian.
Melikyan, L.S., Palandjyan, S.A., Chibukhchyan, Z.H., Vardazaryan, J.S., 1967. To a
question about the geological position and the age of an ophiolitic series of the zone
Shirak–Sevan–Akera of the Lesser Caucasus. Izvestia Academy of Science Armenian
SSR 1–2, 21–41 in Russian.
Miura, S., Suyehiro, K., Shinohara, M., Takahashib, N., Araki, E., Taira, A., 2004.
Seismological structure and implications of collision between the Ontong Java
Plateau and Solomon Island Arc from ocean bottom seismometer airgun data.
Tectonophysics 389, 191–220.
Nicolas, A., 1989. Structures of ophiolites and dynamics of the oceanic lithosphere.
Kluwer Academy Publisher, Dordrecht, The Netherlands. 367 pp.
Nicolas, A., Jackson, E.D., 1972. Répartition en deux provinces des péridotites des
chaînes alpines longeant la Méditerranée: implications géotectoniques. Bulletin
Suisse de Minéralogie et Pétrologie 53, 385–401.
Nikogosian, I.K., Sobolev, A.V., 1997. Ion microprobe analysis of melt Inclusions in
olivine: experience in estimating the olivine-melt partition coefficients of trace
elements. Geochemistry of Interiors 35, 119–126.
Norman, T.N., 1984. The role of the Ankara Melange in the development of Anatolia,
Turkey. Geological Society Special Publications 17, 441–447.
O'Brien, P., Zotov, N., Law, R., Khan, A.M., Jan, M.Q., 2001. Coesite in Himalayan eclogite
and implication for models of India–Asia collision. Geology 29, 435–438.
Okay, A.I., 2000. Was the Late Triassic orogeny in Turkey caused by the collision of an
oceanic plateau? Geological Society Special Publications 173, 25–41.
Palandjyan, S.A., 1971. The petrology of ultrabasites and gabbroic rocks of the Sevan
mountain chain. Izvestia Academy of Science Armenian SSR. 201 pp. (in Russian).
Pallister, J.S., Knight, R.J., 1981. Rare earth element geochemistry of the Samail ophiolite
near Ibra, Oman. Journal of Geophysical Research 86, 2673–2697.
Parkinson, I.J., Pearce, J.A., 1998. Peridotites from the Izu–Bonin–Mariana forearc, ODP
Leg 125): evidence for mantle melting and melt–mantle interaction in a suprasubductive zone setting. Journal of Petrology 39, 1577–1618.
Pearce, J.A., Lippard, S.J., Roberts, S., 1984. Characteristics and tectonic significance of
supra-subduction zone ophiolites. In: Kokelaar, B.P., Howells, M.F. (Eds.), Marginal
Basin Geology. Geological Society Special Publications, vol. 15, pp. 77–94.
Petterson, M.G., Neal, C.R., Mahoney, J.J., Kroenke, L.W., Saunders, A.D., Babbs, T.L.,
Duncan, R.A., Tolia, D., McGrail, B., 1997. Structure and deformation of north and
central Malaita, Solomon Islands: tectonic implications for the Ontong Java Plateau–
Solomon arc collision, and for the fate of oceanic plateaus. Tectonophysics 283,
1–33.
Plank, T., Langmuir, C.H., 1998. The chemical composition of subducting sediment and
its consequences for the crust and mantle. Chemical Geology 145, 325–394.
Rapp, R.P., Watson, E.B., Miller, C.F., 1991. Partial melting of amphibolite/eclogite and the
origin of Archean trondhjemites and tonalites. Precambrian Research 94,
4619–4633.
Révillon, S., Arndt, N.T., Chauvel, C., Hallot, E., 2000. Geochemical study of ultramafic
volcanic and plutonic rocks from Gorgona Island Colombia: the plumbing system of
an oceanic plateau. Journal of Petrology 41, 1127–1153.
Richards, M.A., Jones, D.L., Duncan, R.A., DePaolo, D.J., 1991. A mantle plume initiation
model for the Wrangellia flood basalt and other oceanic plateus. Science 252,
263–267.
Richardson, W.P., Okal, E.A., VanderLee, S., 2000. Rayleigh-wave tomography of the
Ontong Java Plateau. Physics of the Earth and Planetary Interiors 118, 29–51.
Ricou, L.E., Zonenshain, L.P., Dercourt, J., Kazmin, V.G., Le Pichon, X., Knipper, A.L.,
Grandjacquet, C., Sborshchikov, I.M., Geyssant, J., Lepvrier, C., Pechersky, D.M.,
Boulin, J., Sibuet, J.C., Savostin, L.A., Sorokhtin, O., Westphal, M., Bazhenov, M.L.,
Lauer, J.P., Biju-Duval, B., 1985. Méthodes pour l'établissement de neuf cartes
paléogéographiques de l'Atlantique au Pamir depuis le Lias. Bulletin de la Société
Géologique de France 8, 625–635.
Rolland, Y., Billo, S., Corsini, M., Sosson, M., Galoyan, G., 2009. Blueschists of the
Amassia–Stepanavan Suture Zone, Armenia): linking Tethys subduction history
from E-Turkey to W-Iran. International Journal of Earth Sciences 98, 533–550.
Rolland, Y., Billo, S., Corsini, M., Sosson, M., Galoyan, G. in press. The Armenian Ophiolite:
insights for Jurassic Back-arc formation, Lower Cretaceous hot-spot magmatism, and
Upper Cretaceous obduction over the South Armenian Block. Geological Society
Special Publication.
Salters, V.J.M., White, W.M., 1998. Hf isotope constraints on mantle evolution. Chemical
Geology 145, 447–460.
Salters, V., Stracke, A., 2004. Composition of the depleted mantle. Geochemistry
Geophysics Geosystems 5. doi:10.1029/2003GC000597.
Saunders, A.D., Tarney, J., 1984. Geochemical characteristics of basaltic volcanism within
back-arc basins. In: Kokelaar, B.P., Howells, M.F. (Eds.), Marginal Basin Geology.
Geological Society Special Publications, vol. 16, pp. 59–76.
Satian, M.A., Sarkisyan, E.A., 2006. On lithodinamics of Mesozoic volcanogenesedimentary complex of the Sevan range, Sevan–Akera ophiolite zone, Armenia.
Izvestia Academy of Science Armenian SSR 2, 19–26 (in Russian).
Satian, M.A., Stepanyan, J.H., Sahakyan, L.H., Mnatsakanyan, A.Kh., Ghukasyan, R.Kh.,
2005. The Mesozoic lamprophyric diatremes of Vedi zone, Armenia. Izdatelstvo
‘Nairi’ Yerevan. 148 pp. (in Russian).
Sengör, A.M.C., 1984. The Cimmeride Orogenic System and the Tectonics of Eurasia.
Geological Society of America Special Paper, vol. 195, p. 82.
Sengör, A.M.C., 1990. A new model for the late Paleozoic–Mesozoic tectonic evolution of
Iran and implications for Oman. In: Robertson, A.H., Searle, M.P., Ries, A.C. (Eds.),
The Geology and tectonics of the Oman region. Geological Society Special
Publication, vol. 49, pp. 797–831.
Sengör, A.M.C., Yılmaz, Y., 1981. Tethyan evolution of Turkey: a plate tectonic approach.
Tectonophysics 75, 181–241.
Sokolov, S.D., 1977. The olistostroms and ophiolitic nappes of the Lesser Caucasus.
Izdatelstvo Nauka, Moscow. 92 pp. (in Russian).
Y. Rolland et al. / Lithos 112 (2009) 163–187
Sosson, M., Rolland, Y., Danelian, T., Muller, C., Melkonyan, R., Adamia, S., Kangarli, T.,
Avagyan, A., Galoyan, G., in press. Subductions, obduction and collision in the Lesser
Caucasus (Armenia, Azerbaijan, Georgia), new insights. Geological Society Special
Publication.
Stampfli, G.M., Borel, G.D., 2002. A plate tectonic model for the Palaeozoic and Mesozoic
constrained by dynamic plate boundaries and restored synthetic oceanic isochrons.
Earth and Planetary Science Letters 196, 17–33.
Staudigel, H., Plank, T., White, B., Schmincke, H.-U., 1996. Geochemical fluxes during
seafloor alteration of the basaltic upper oceanic crust: DSDP Sites 417 and 418. In:
Bebout, G.E., Scholl, D.W., Kirby, S.H., Platt, J.P. (Eds.), Subduction: Top to Bottom.
AGU Monograph, pp. 19–38.
Steiger, R.H., Jäger, E., 1977. Subcomission on geochronology: convention of the use of
decay constants in geo- and cosmochronology. Earth Planetary Science Letters 36,
359–362.
Storey, M., Mahoney, J.J., Kroenke, L.W., Saunders, A.D., 1991. Are oceanic plateaus sites
of komatiite formation? Geology 19, 376–379.
Sun, S.S., McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts:
implications for mantle composition and processes. In: Saudners, A.D., Norry, M.J.
(Eds.), Magmatism in Ocean Basins. Geological Society Special Publications, vol. 42,
pp. 313–345.
Tankut, A., Dilek, Y., Önen, P., 1998. Petrology and geochemistry of the Neo-Tethyan
volcanism as revealed in the Ankara melange, Turkey. Journal of Volcanology and
Geothermal Research 85, 265–284.
Tarnudo, J.A., Sliter, W.V., Kroenke, L.W., Leckie, M., Mahoney, J.J., Musgrave, R.J., Storey,
M., Winterer, E.L., 1991. Rapid formation of the Ontong Java Plateau by Aptian plume
volcanism. Science 254, 399–403.
187
Taylor, S.R., McLennan, S.M., 1985. The continental crust: its composition and evolution.
Blackwell. 312 pp.
Tirrul, R., Bell, R., Griffis, R.J., Camp, V.E., 1983. The Sistan Suture zone of eastern Iran.
Geological Society of America Bulletin 94, 134–150.
Turner, G., Huneke, J.C., Podose, F.A., Wasserburg, G.J., 1971. 40Ar/39Ar ages and cosmic
ray exposure ages of Apollo 14 samples. Earth and Planetary Science Letters 12,
15–19.
Tüysüz, O., Dellaloğlu, A.A., Terzioğlu, N., 1995. A magmatic belt within the Neo-Tethyan
suture zone and its role in the tectonic evolution of northern Turkey. Tectonophysics 243, 173–191.
Wasserburg, B.J., Jacobsen, S.B., De Paolo, D.J., McCulloch, M.T., Wen, T., 1981. Precise
determination of Sm/Nd ratios, Sm and Nd isotopic abundances in standard
solutions. Geochimica et Cosmochimica Acta 45, 2311–2323.
Zakariadze, G.S., Knipper, A.L., Sobolev, A.V., Tsameryan, O.P., Dimitriev, L.V.,
Vishnevskaya, V.S., Kolesov, G.M., 1983. The ophiolite volcanic series of the Lesser
Caucasus. Ofioliti 8, 439–466.
Zakariadze, G.S., Knipper, A.L., Bibikova, E.V., Silantiev, S.A., Zlobin, S.K., Gracheva, T.V.,
Makarov, S.A., Kolesov, T.M., 1990. The setting and age of the plutonic part of the NE
Sevan ophiolite complex. Academy of sciences of the USSR Geological series 3,
17–30 (in Russian).
Zindler, A., Hart, S.R., 1986. Chemical geodynamics. Annual Reviews in Earth and
Planetary Sciences 14, 493–571.