SESolid EarthSESolid Earth1869-9529Copernicus PublicationsGöttingen, Germany10.5194/se-7-579-2016Insights on high-grade deformation in quartzo-feldspathic gneisses during
the early Variscan exhumation of the
Cabo Ortegal nappe, NW IberiaFernándezFrancisco Josébrojos@geol.uniovi.esLlana-FúnezSergiohttps://orcid.org/0000-0002-8748-5623Valverde-VaqueroPabloMarcosAlbertoCastiñeirasPedroDepartamento de Geología, Universidad de Oviedo,
Jesús Arias de Velasco s/n, 33005 Oviedo, SpainÁrea de Laboratorios, Instituto Geológico y Minero
de España, La Calera 1, 28760 Tres Cantos, SpainDepartamento de Petrología y Geoquímica,
Universidad Complutense de Madrid, José Antonio Novais 12, 28040 Madrid,
SpainFrancisco José Fernández (brojos@geol.uniovi.es)13April20167257959825November20157December201521March201622March2016This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://se.copernicus.org/articles/7/579/2016/se-7-579-2016.htmlThe full text article is available as a PDF file from https://se.copernicus.org/articles/7/579/2016/se-7-579-2016.pdf
High-grade, highly deformed gneisses crop out continuously along the
Masanteo peninsula and constitute the upper part of the lower crustal
section in the Cabo Ortegal nappe (NW Spain). The rock sequence formed by
migmatitic quartzo-feldspathic (qz-fsp) gneisses and mafic rocks records the
early Ordovician (ca. 480–488 Ma) injection of felsic dioritic/granodioritic dykes at the base of
the qz-fsp gneisses, and Devonian eclogitization (ca. 390.4 ± 1.2 Ma),
prior to its exhumation. A SE-vergent ductile thrust constitutes the base of
quartzo-feldspathic gneissic unit, incorporating mafic eclogite blocks within
migmatitic gneisses. A NW-vergent detachment displaced metasedimentary
qz-fsp gneisses over the migmatites. A difference in metamorphic pressure of
ca. 0.5 GPa is estimated between both gneissic units. The
tectono-metamorphic relationships of the basal ductile thrust and the normal
detachment bounding the top of the migmatites indicate that both discrete
mechanical contacts were active before the recumbent folding affecting the
sequence of gneisses during their final emplacement. The progressive
tectonic exhumation from eclogite to greenschist facies conditions occurred
over ca. 10 Ma and involved bulk thinning of the high-grade rock sequence in
the high pressure and high temperature (HP–HT) Cabo Ortegal nappe. The necessary strain was accommodated by the
development of a widespread main foliation, dominated by flattening, that
subsequently localized to a network of anastomosing shear bands that evolved
to planar shear zones. Qz-fsp gneisses and neighbouring mafic granulites
were exhumed at > 3 mm yr-1, and the exhumation path involved a
cooling of ∼ 20 ∘C/100 MPa, These figures are
comparable to currently active subduction zones, although exhumation P–T trajectory and ascent rates are at the hotter and slower end in comparison
with currently active similar settings, suggesting an extremely ductile
deformation environment during the exhumation of qz-fsp gneisses within a
coherent Cabo Ortegal nappe.
Introduction
The processes involved in the exhumation of high-pressure (HP) and ultra-high-pressure (UHP) rocks in subduction
zones remain a hot topic in tectonics given the complexity of strain and
displacement paths that rocks follow, from the surface to great depths and
back to the surface (e.g. Gerya et al., 2008). The boundary interval between
convergent plates concentrates a large amount of strain and also
heterogeneity (e.g. Escher and Beaumont 1997). This interval in subduction
zones, named as the subduction channel, is characterized by non-parallel
planar rigid edges on either side, on a profile having a narrow,
downward-tapering triangular shape (i.e. Bird, 1978; England and Holland,
1979; Shreve and Cloos, 1986; Mancktelow, 1995). Under this configuration,
the convergence of rigid plates squeezing a non-compressible viscous
material introduces a stress gradient in the system that leads to a
non-lithostatic pressure gradient with depth (e.g. Mancktelow, 1995). If the
shearing associated with the convergence is taken into account, the result
is that particles close to the subducting plate will follow the lower
boundary, but once they reach the vertex of the triangular channel will be
entrained to return to the surface, following the upper rigid boundary (see
Fig. 4 in Shreve and Cloos, 1986). In fact, the exhumation of high pressure
rocks represents the return flow in the system, so that subduction zones
need to be active in order that high pressure rocks may reach the surface.
The rheology of rocks in the subduction channel has an impact that affects
the velocity of exhumation in comparison to subduction rates. Exhumation
rate is about 1/3 of the subduction rate if deformation approaches a
Newtonian behaviour; but is slower if deformation is dominated by more
non-linear processes such as dislocation creep, becoming about 1/6 of the
subduction rate (Gerya and Stöckhert, 2002). In addition, there is
substantial intrinsic heterogeneity in the system at the boundary between
plates, which can be now be visualised in numerical models (e.g. Gerya et
al., 2008; Burov et al., 2014a, b). The rock record does not always preserve
all the deformation stages, thus the difficulty in inferring a finite
strain/displacement pathway for rocks and rock units remains.
In continental collision, subsequent in most cases to a subduction stage,
there are some analogies with the “subduction channel” or the boundary
zone between plates, but some major differences. The first major difference
is that, as a consequence of the less-rigid plate boundaries involved, the
extent in cross-section of this idealized downward-tapering plate boundary
increases substantially (e.g. Gerya et al., 2008). In the upper part it
consists of an orogenic wedge, a Coulomb wedge or an accretionary wedge. It
has a triangular shape in cross-section, but the angles between sides can
vary. Displacement paths of particles within the system do follow the sides
of this wedge, but the dynamics are completely different from the deeper
domain. In orogenic wedges, the exhumation of subducted rocks from depths
greater than 50 km cannot be satisfactorily explained by classical collision
models, such as in the dynamics of accretionary wedges (i.e. Davis et al.,
1983; Platt, 1986) or by exhumation by extensional collapse of the orogen
(i.e. Chemenda et al., 1995). In fact, insights from numerical models of
ultra-high-pressure (UHP) exhumation at the continental collisional phase
are consistent with a multi-stage process, where exhumation seems to start
after a degree of continental subduction for most continental collision
zones (e.g. Burov et al., 2014a, b).
One aspect of significance associated with the dynamics of channel flow at a
crustal scale in collisional orogens is the upward extrusion of high-grade
rocks squeezed between colliding plates. In the case of the Himalayan-Tibet
system, the great crustal thickness beneath the Tibetan Plateau contributes
significantly to the lithostatic pressure gradient required to force the
lateral and frontal flow of a ductile lower crust (e.g. Beaumont et al.,
2004; Rutter et al., 2011). Highly sheared and migmatized rocks of the
greater Himalayan sequence between the Main Central Thrust and the South
Tibetan Detachment are effectively extruded towards the foreland. The
extrusion process involves substantial thinning of the slab or fragment of
crust involved in the subduction or collision. This is also apparent from
numerical models, where a weak accretionary wedge can be squeezed out
between the mantle portions of both colliding plates (e.g. Gerya et al.,
2008).
Relics of the plate boundary region between northern Gondwana and Laurasia,
and the accretionary complex sandwiched in between, are preserved in the
high-grade allochthonous complexes of NW Iberian peninsula (e.g. Ries and
Shackleton, 1971; Martínez-Catalán et al., 1997; Matte, 2001). The
timing, kinematics, and structure of one of these, the Cabo Ortegal complex
(COC), has been discussed over the past 45 years, and many hypotheses and
models have been proposed to explain the exhumation and final tectonic
emplacement of the high-grade, high-pressure rocks that form the greater
part of the COC as it is presently exposed. Initially the debate was focused
on whether the Cabo Ortegal complex represents evidence of a mantle plume,
as proposed by the Leiden group (e.g. van Calsteren et al., 1979) or was in
fact an allochthonous thrust sheet, as supported by many other schools since
then (e.g. Ries and Shackleton, 1971; Bayer and Matte, 1979). A second point
of contention was whether high-pressure metamorphism and the sequence of
structures represented a deep tectonic setting, i.e. the remains of the
subduction channel itself (e.g. Ábalos et al., 2003), or were due to the
superposition of structures during their exhumation from high-pressure
conditions (e.g. Marcos et al., 2002). Regarding the initial exhumation from
high-pressure conditions to mid-crustal depths, several mechanisms have been
proposed, in relation to extensional tectonics (e.g.
Martínez-Catalán et al., 1997), ductile slab breakoff (e.g.
Llana-Fúnez et al, 2004) or more recently channel flow (e.g. Albert et
al., 2012).
The Masanteo peninsula (Figs. 1 and 2), located on the eastern side of the
Cabo Ortegal Complex, offers a continuous exposure through the
quartzo-feldspathic (qz-fsp) gneisses in the structurally upper part of the
coherent rock sequence that constitutes the Cabo Ortegal nappe, the part of
the COC that registers widespread high-pressure and high-temperature
metamorphism which is also associated with generalized intense deformation.
In this paper, we present a detailed structural analysis of the high-grade
tectonic sequence at the contact between the eclogitic mafic gneisses and the
qz-fsp gneisses, on the eastern side of the Cabo Ortegal nappe (Fig. 2a).
These gneisses are quite exposed in clean exposures along the shores of the
Masanteo peninsula, covering an area of 4.5 km2.
(a) Geological map of the Variscan Belt in N Iberia,
highlighting the allochthonous complexes, based on Parga Pondal et
al. (1982). (b) Geological map with location of Fig. 2 and the
CPOs samples. (c) Lithostratigraphic sequence and (d)
cross-section of Cabo Ortegal nappe (after Marcos et al., 2002).
(a) Geological map of the Masanteo peninsula with location
of samples, figures and cross-sections. Pole figures in (b) show the
distribution of S2-foliation poles and related structures (lineation and
intrafolial fold axes). Equal area projection, lower hemisphere.
A detailed mapping of the gneisses and the reconstruction of the rock unit
geometry on the basis of the attitude of the main tectonic foliation
(S2) is therefore presented here with the aim of understanding the
mechanism of deformation in the tectonized zone between mafic and
quartzo-feldspathic rocks during the exhumation of subducted lower crust. The
deduction of the tectonic evolution in the gneisses is based on newly
observed structural relationships, recent Uranium-lead (U-Pb) ages, and is supported by
geothermobarometry data from previous studies in the gneisses and
neighbouring mafic rock units. The insights from well-preserved high-grade
structures observed in the field are believed to be the key to understanding
the processes of orogenic collision and to constrain thermo-mechanical
models. The integration of published geochronological data, metamorphic
evolution and the structural development helps to constrain some
characteristics of the process of exhumation of high-grade gneisses from
eclogite conditions to greenschist facies. In this contribution, we provide
robust figures for the bulk exhumation rate and amount of cooling during
exhumation of high-pressure rocks in the Cabo Ortegal nappe which may be
regarded as representative for a larger lower crustal section.
The geological framework: the Cabo Ortegal complex
High-grade relicts of continental collision overlie tectonically most of the
hinterland of the Variscan orogeny in NW Iberia, forming a tectonic pile of
oceanic and sedimentary material that can be recognized totally or partially
within five allochthonous complexes (Martínez-Catalán et al.,
1997). Three main units form this allochthonous tectonic pile. The upper
unit includes ultramafic, mafic, and qz-fsp gneissic rocks that recorded high-pressure and high-temperature (HP-HT) metamorphism (e.g. Vogel, 1967). Rocks
characteristic of distinct geodynamic settings, such as E-MORB basalts,
tectonic melanges and arc volcanics (Arenas, 1986; Díaz-García et
al., 1999; Pin et al., 2006; Arenas et al., 2007) form the intermediate
ophiolitic unit. The basal unit is formed by metasediments intruded by acid
and basic igneous rocks that recorded blueschist facies and low- to
intermediate-temperature eclogite facies conditions (e.g. Gil-Ibarguchi and
Ortega-Girones, 1985; Arenas et al., 1995; López-Carmona et al., 2014).
The Cabo Ortegal complex is the northernmost of the allochthonous complexes,
sitting on rocks of the Central Iberian Zone, and the closest to the foreland
of the orogen (Fig. 1a; e.g. Pérez-Estaún et al., 1991). This complex
is exposed along a coastal section with cliffs up to 600 m above sea-level,
and contains some of the best exposures of HP-HT rocks of the Variscan
orogen. Internally, it is divided into two main tectonic units: the Cabo
Ortegal nappe and the lower tectonic unit (Marcos et al., 2002). The Cabo
Ortegal nappe (Fig. 1b) is composed of ultramafic rocks, mafic granulites,
eclogites, and high-grade gneisses with HP-HT metamorphism, which are the
objects of this study. The underlying lower tectonic unit is composed of
three thrust sheets: the ophiolitic unit, the basal unit, and the
para-autochthonous rock sequence (Marcos and Farias, 1999).
There are three major lithological units that form the high-grade Cabo
Ortegal nappe (Vogel, 1967). From bottom to top, according to the
lithostratigraphic column in Fig. 1c, the base is formed by
> 600 m of mantle-derived ultramafic rocks, metaperidotites, and
garnet pyroxenites (Girardeau et al., 1989), overlain by a 400 m thick layer
of mafic HP granulites. The granulites are in sharp contact with a
100–200 m thick unit of massive, high–temperature eclogite derived from
N-MORB mafic protoliths (e.g. Peucat et al., 1990). The eclogite unit is
capped by quartzo-feldspathic gneisses (> 600 m thick), also
described as HP gneisses by Ábalos et al. (2003). The qz-fsp gneisses
along the contact with the eclogite unit contain decametric to metric scale
lenses of mafic eclogite and locally show evidences of partial anatexis and
migmatization (Vogel, 1967; Gil-Ibarguchi et al., 1990; Fernández, 1997).
These migmatitic gneisses correspond to the Chimparra and banded gneisses of
Vogel (1967). Metasedimentary, supracrustal, pelitic, and psammitic
paragneisses overlie the migmatitic qz-fsp gneisses, and mark the top of the
high-grade nappe. This higher paragneiss unit is the Cariño gneiss of
Vogel (1967). It also displays evidence of HP metamorphism (Fig. 5) but
indications of anatexis are scarce. These units are interpreted to outcrop
largely in the shape of an early recumbent synform (Fig. 1d; Marcos et al.,
2002). Overall, the whole lithostratigraphic sequence has been used as a
proxy for the continental crust-mantle transition (Brown et al., 2009); and
according to Ábalos et al. (2003), it represents a stack of crustal and
mantle units assembled in a subduction channel, i.e. with oceanic origins; it
could also represent a continental lower crustal/uppermost mantle sequence
that has been partially subducted and exhumed.
Tectonic evolution of the high-grade Cabo Ortegal nappe
The high-pressure rocks of Cabo Ortegal are affected by several phases and
stages of deformation and metamorphism. Protolith ages of the ultramafic and
mafic rocks in the Cabo Ortegal nappe are considered Cambro-Ordovician on the
basis of 520–490 Ma U-Pb zircon (Peucat et al., 1990; Ordoñez-Casado et
al., 2001) and Samarium-Neodymium (Sr-Nd) ages (Santos et al., 2002). Fernández-Suárez et
al. (2002) reported 480–490 Ma U-Pb monazite and zircon ages from anatectic
pods in the mafic granulites and the migmatitic qz-fsp gneisses, suggesting
the presence of a Cambro-Ordovician metamorphic overprint similar to that
reported by Abati et al. (1999) from equivalent units in the neighbouring
Órdenes complex (Fig. 1a).
An early–mid Devonian (400–388 Ma) HP-HT metamorphic event is well
documented by Garnet-Clinopyroxene (Grt-Cpx) whole-rock Sm-Nd ages in the metaperidotites of the
ultramafic unit (Santos et al., 2002), U-Pb zircon ages in the eclogite unit
(e.g. Peucat et al., 1990; Santos-Zalduegui et al., 1996;
Ordóñez-Casado et al., 2001) and zircon, monazite, and titanite U-Pb
ages from the mafic granulite and the qz-fsp gneisses (Santos-Zalduegui et
al., 1996; Fernández-Suárez et al., 2002, 2007). This event reached
metamorphic conditions above 800 ∘C and 1.7 GPa (Gil-Ibarguchi et
al., 1990; Fernández, 1997; Galán and Marcos, 2000). The eclogitic
and granulitic rocks registered partial anatexis during the initiation of
decompression at peak temperature conditions, which was followed by partial
retrogression and amphibolitization. Cooling ages in amphibole
(390–380 Ma), rutile (U-Pb, ca. 383 Ma; Santos-Zalduegui et al., 1996;
Valverde-Vaquero and Fernandez, 1996) and muscovite (375 Ma; Argon-Argon (Ar-Ar) data,
Peucat et al., 1990) ages indicate a fast cooling and rapid exhumation of the
HP-HT rocks that constitute the nappe following peak HP conditions (see
Ordoñez Casado et al., 2001).
The development of structures associated with metamorphism allow the
definition of a relative chronology made of locally four deformation events
in the Cabo Ortegal nappe. It should be noted that these phases that do not
have a direct correlation with the three main regional deformation phases in
the underlying, autochthonous rock sequence (e.g. Pérez-Estaún et
al., 1991). Some authors interpret inclusion trails as representing a
distinct earliest event (D1 structures) formed during the prograde path
related to the subduction stage (i.e. Ábalos et al., 2003); even though
only the fabric elements of the retrograde Pressure–Temperature–time (P–T–t) path have been unequivocally
recognized (Gil-Ibarguchi et al., 1990; Fernández, 1997; Galán and
Marcos, 2000). All units of the high-grade nappe show a pervasive
blastomylonitic tectonic fabric, which locally is highly heterogeneously
developed (Fernández, 1997; Marcos et al, 2002). This main tectonic
fabric and associated structures define a regionally recognizable deformation
episode (labelled D2), formed during the exhumation from high-pressure
conditions. Frequently, the blastomylonitic foliation (S2) forms
networks of anastomosed shear zones, and defines lozenge-shaped bodies of
layered migmatitic gneisses that preserve even earlier-developed fabrics (Fernández
and Marcos, 1996). Whilst there is a lack of a pervasively developed mineral
lineation, the symmetry of quartz crystallographic preferred orientation
(CPO) patterns suggests a predominantly coaxial deformation during fabric
development in the gneisses (Fernández, 1997). The omphacite CPO fabrics
in neighbouring eclogite units show a similar pattern, also consistent with
flattening strain in the plane of foliation (Llana-Fúnez et al., 2005).
Overall, bulk coaxial strain dominated the D2 deformation and controlled
the bulk tectonic thinning of the rock sequence in the Cabo Ortegal nappe
(Llana-Fúnez et al, 2004).
Large-scale recumbent folding resulted in the inversion of the
lithostratigraphy along a reverse limb for more than 6 km in the direction
of tectonic transport, and determines the overall outcrop pattern in the COC.
This folding deforms the D2 fabric and is hence recognized as D3.
Later asymmetric folds of decametric size cut across D2 folds and mark
the formation of a large E-verging recumbent D3-fold, which was
overprinted by two major thrusts attributed to D4 (Fig. 1b and d; Marcos
et al., 1984, 2002). The subsequent tectonic evolution was controlled by
the progressive localization of strain and the deformation along thrusts and
the final emplacement of the HP-HT Cabo Ortegal nappe.
The D4a-thrusts imbricated the Cabo Ortegal nappe as part of its
progressive emplacement toward the ESE over the underlying ophiolitic rock
units (Marcos and Farias, 1999). This presumably implies that the emplacement
over the ophiolitic rocks was now under greenschist facies conditions. Late
D4b-upright refolding produced the elliptical final shape of the
Cabo Ortegal synformal complex that led to its preservation from erosion. The
other allochthonous complexes were similarly affected, and upright
large-scale folds of similar orientation also affect the rocks of the
underlying autochthon during the later part of the Variscan Orogeny (Matte,
1968; Marcos, 1971; Pérez-Estaún et al., 1991). The kilometric
amplitude and wavelength of upright folds reach crustal-scale and evidence
continuing shortening subsequently to the emplacement of allochthonous
complexes (Llana-Fúnez and Marcos, 2007).
The rock sequence at Masanteo peninsula
The rock sequence that crops out in the Masanteo peninsula is more than
300 m thick. It is a high-grade imbricated section of the upper part of the
mafic-banded gneisses and the qz-fsp, migmatitic, and metasedimentary gneisses
(Fig. 2a). The whole sequence shows a heterogeneous deformation defined by a
ductile pervasive S2 foliation. Anastomosed and planar D2 shear
zones are widespread, and include metric-sized boudins of eclogite,
ultramafic, and mafic rocks. The anastomosing shear zones surround
lozenge-shaped bodies within the layered migmatitic gneisses. The size of
lozenges ranges from 0.5 to 4 m. The D2-shear zones also include
symmetric- and rotational structures of centimetre-size and
rootless-intrafolial folds. The D2-planar tectonites do not show a
consistently developed stretching lineation. However, the intersection
between the S2 foliation and the compositional or migmatitic layering
forms locally a linear fabric. Also locally, the D2 high-strain zones
show garnet or amphibole lineations with scattered patterns (Fig. 2b). The
mineral assemblages (M2) associated with this dominant D2
deformation indicate retrogression from eclogite to amphibolite facies during
the imbrication of the whole stack (Gil-Ibarguchi et al., 1990;
Fernández, 1997; Mendia 2000). The following is a description of the main
lithological units within the D2 imbricate complex at the Masanteo
peninsula.
Mafic gneisses
Highly strained amphibole-bearing gneisses crop out at the base of the
Masanteo cliff (Fig. 3). These amphibolitic gneisses enclose boudins and
blocks of eclogite, partly retrogressed eclogite, and rare metagabbros. The
amphibolitic gneisses have a high-strain D2 fabric defined by a mineral
assemblage with
Qz + Pl + Hbl + Ky + Grt + Bt ± Kfs ± Czo ± Ilm ± Ttn
(mineral abbreviations following Whitney and Evans, 2010). The eclogite
boudins are composed of Omp + Grt ± Hbl, and often preserve undeformed
textures with inclusions of Rt in Grt, and locally contain mesocratic melt
pods. The mineral assemblage in the retrogressed eclogites is formed by
Qz + Grt + Omp ± Hbl ± Bt ± Pl ± Rt ± Ilm ± Ttn.
The rare metagabbros (Ol + Pl + Grt ± Ab ± Ep) preserve
relict ophitic textures and prograde pre-eclogitization coronitic garnets.
This D2 fabric developed in both mafic and migmatitic gneisses and
characterizes such contacts. Peak metamorphic conditions during the eclogite
stage have been constrained by Mendia (2000) at 800 ∘C and 2.2 GPa.
Mafic gneisses and related rocks. (a) Structures at the
exposure scale; the sketch shows the attitude of the main S2 foliation
at the contact between mafic gneisses and migmatitic gneisses.
Microphotographs: (b) retrogressed coronitic metagabbro (sample
B917); (c) Bt-Grt-bearing amphibolite gneisses within the less
deformed lozenges (sample B714). Sample location is indicated in Fig. 2a.
Migmatitic gneisses
These migmatitic gneisses correspond to the “banded-gneisses” of
Vogel (1967) and are equivalent to his Chimparra gneiss. They are the highest
grade qz-fsp gneisses in the Cabo Ortegal nappe. They consist of
Ky ± Rt ± Grt-bearing layered, biotite-rich migmatites. At the
Masanteo peninsula they are sandwiched between the mafic gneisses and the
overlying metasedimentary gneisses. Locally, they contain centimetric to
decimetric thick bands of orthogneiss
(Qz + Mc + Pl + Grt + Ms + Bt) intruded by felsic
(Qz + Pl + Kfs + Grt + Hbl ± Czo) and
tonalitic/granodioritic dykes intercalated in the migmatitic gneisses. The
total thickness of this gneissic unit ranges from 50 to 200 m. The
migmatitic gneisses have two compositional end-members: a
melanosome-dominated, biotitic qz-fsp gneiss (Fig. 4d) with
Ky + Grt + Bt ± Hbl ± Czo ± Ilm ± Ttn, with
less than 20 % of leucosome; and a banded leucocratic qz-fsp gneisses
(Fig. 4f) with
Qz + Pl + Kfs ± Ky ± Czo ± Ilm ± Ttn with
less than 20 % of melanosome. The difference in modal composition may
relate to differences in the primary composition of the metasedimentary
rocks; however, compositional differentiation may also be a consequence of
migmatization and/or subsequent deformation. The banded leucocratic gneisses
are located on the upper section of the migmatitic unit, while the
biotite-rich migmatitic gneiss occurs along the contact with the underlying
mafic gneiss. A phyllonitic fabric of the biotite gneisses, including
centimetric layers of restitic material (Fig. 4e), and its location overlying
the mafic gneisses points to deformation in high-grade conditions.
Migmatitic biotite qz-fsp gneisses and related rocks. (a) Folding affecting a felsic dioritic dyke and the S2 foliation.
(b) Microphotograph of the felsic diorite dyke showing a coarse
foliation (sample DM-2). (c) Anastomosing shear zones defined by the
S2 foliation surrounding lozenges of less deformed migmatitic qz-fsp
gneisses. (d) Microphotograph of the biotite qz-fsp gneisses (sample
B23). (e) Restite in migmatitic qz-fsp gneisses. (f) Microphotograph of the leucocratic qz-fsp gneisses (sample B12). Sample
locations are in Fig. 2a.
Peak metamorphic conditions estimated for the migmatitic gneisses in the
Masanteo peninsula have been estimated at 720 ∘C and 1.5 GPa
(Gil-Ibarguchi et al, 1990). Estimates of metamorphic conditions of
equivalent biotite qz-fsp gneisses in Punta Tarroiba (location in Fig. 1b),
the Chimparra gneisses (Vogel, 1967), show slightly higher values of
800 ∘C and 1.7 GPa (Fernández, 1997) comparable to conditions
calculated in the eclogites (Fig. 5).
P–T data calculated for metasedimentary and migmatitic qz-fsp
gneisses in Cabo Ortegal nappe, based on the available published data,
indicated in the legend. Al2SiO5 phase diagram after
Holdaway (1971). Error bars are also plotted. P–T path in the gneisses
proposed is in grey. Detailed metamorphic work by Galán and Marcos (2000)
traced the P–T-t path (thick dashed line) of the underlying Mafic
granulites (after Galán and Marcos, 2000) is also plotted. Thin dashed
lines trace sub-parallel retrograde paths for metasedimentary and migmatitic
qz-fsp gneisses
U–Pb CA-ID-TIMS data of the diorite dyke samples DM-2 and DM-3
(small white filled ellipses – zircon; grey ellipses – monazite). Locations
are in Fig. 2a.
Locally, the structural relationships between the blastomylonitic
S2 foliation and felsic dioritic/granodioritic dykes allow the relative
timing of events in these gneisses to be constrained. The felsic dykes are
buckled by metric folds that are transected by the S2 foliation (Fig. 4a
and b), demonstrate that intrusion and folding of the felsic dykes occurred
before the D2 deformation. The S2 foliation shows sub-parallelism
with the migmatitic layering and bounds concordantly the eclogite blocks
(Fig. 4c). Castiñeiras et al. (2010) sampled one of these eclogitic
boudins (sample COZ-4; Fig. 2) and obtained a U-Pb zircon age of
390 ± 2 Ma. The field relationships suggest that local anatexis and
migmatization must have occurred during the early stages of D2, after
eclogitization. Thus, the migmatitic qz-fsp gneisses apparently have recorded
an early intrusive event related to the injection of the pre-D2
dioritic/granodioritic dykes in the gneisses, and a latter anatectic melting
event that produced the migmatitic layering, which is preserved within the
less deformed lozenges bodies surrounded by anastomosing D2-shear bands.
Metasedimentary qz-fsp gneisses and related rocks. (a) Leucosome
veins (to right of the scale marker) parallel to the compositional banding.
(b) Intrafolial folds related to the S2 foliation superimposed on the
compositional banding. (c) Microphotograph of a metapelite band (sample B1427). (d) Microphotograph of a metapsammite band (sample B22). Sample
locations are in Fig. 2a.
New U-Pb ID-TIMS geochronology in the migmatitic gneiss
In order to constrain the ages of pre-D2 felsic dykes, two separate
felsic dykes (DM-2 and DM-3; Fig. 4a, b) were dated by U-Pb ID-TIMS at the
IGME geochronology laboratory in Tres Cantos (Spain). Zircon and monazite
were analysed following the procedures outlined in Rubio Ordoñez et
al. (2012). The zircon fractions were chemically abraded before final
dissolution.
In the case of sample DM-2, two zircon and three monazite fractions were
analysed (Table 1; Fig. 6). The zircon fractions are discordant, while the
three monazite fractions overlap the concordia curve, providing concordant
ages at 475 (M1), 478 (M2) and 485 Ma (M3). These three
monazite fractions are colinear and provide a lower intercept age of
384 ± 180 Ma and an upper intercept age of 479 ± 6.5 Ma. For
sample DM-3, four zircon and three monazite fractions were dated (Table 1;
Fig. 6). The monazite and zircon fractions Z1, Z4, and Z3 define a mixing line
anchored at 480 ± 8 Ma by the concordant monazite and an upper
intercept at 2.56 Ga, suggesting Proterozoic zircon inheritance. In this
sample, monazite analyses were done using single crystals. Monazites M2 and
M3 overlap each other and provide a concordant age of 480 ± 1 Ma (MSWD
0.44), while monazite M1 is concordant at 488 Ma, resembling the monazite
from sample DM-2. These data clearly demonstrate the presence of
Cambro-Ordovician (ca. 480–490 Ma) monazite in both dykes. A similar spread
of early Ordovician monazite ages, such as those in sample DM-2, was reported
by Fernández-Suárez et al. (2002) in the Cabo Ortegal nappe from
leucosomes of the Chimparra gneiss, suggesting minor Devonian (ca. 386 Ma)
overprint of Cambro-Ordovician monazite. The same authors also reported a
zircon age of 487 Ma from a leucosome in the mafic granulites. Therefore, we
consider that the monazites provide the best estimate for the intrusion age
of the felsic dykes DM-2 and DM-3, which would be bracketed by a minimum age
of 480 Ma (intercepts of the discordia lines) and a maximum age of
485–488 Ma (oldest concordant monazite fractions).
U-Pb CA-ID-TIMS geochronological data of the felsic dykes. Z:
zircon fraction, chemically abraded (CA; Mattison, 2005). M: monazite
fraction. Mnz single xtl: monazite single crystal. Zrn tips: zircon tips.
Weight estimated before CA. Number of grain in each fraction is given within
brackets. Pb (pg), total common Pb blank. * Measured ratio corrected for
blank and fractionation. Atomic ratios corrected for fractionation
(0.11 ± 0.02 % AMU Pb; 0.10 ± 0.02 % AMU U), spike
(208Pb-235U), laboratory blanks (6 pg Pb; 0.1 pg U) and initial
common Pb after Stacey and Kramers (1975). Errors are at the 2-sigma level.
Data reduced with PbMacDat (Isachsen et al., www.earth-time.org).
Concentration Isotopic ratios AgeWeightUPbPb206Pb*208Pb206Pb%207Pb%207Pb%206Pb207Pb207PbFractions(mg)(ppm)(ppm)(pg)204Pb206Pb238Uerr235Uerr206Pberr238U235U206PbRhoDM-2Z1: 2 prims + 2–3 tips0.071741560752.70.01570.083100.470.716490.910.0625330.785155496920.52Z2: 2 Zrn tips0.0223516131398.30.00120.074810.130.586580.280.0568690.244654694870.53M1(Mnz single xtl)0.021131116812614.20.45790.078260.200.612950.240.0568070.144864854840.83M2 (Mnz – 2 small xtls)0.031672225365646.00.97120.076460.170.596630.230.0565910.154754754760.75M3 (Mnz – 3 small xtls)0.04517072141981635.30.78810.077090.440.602460.450.0566760.054794794790.99DM-3Z1: Single grain prism0.0265460441098.70.01550.092900.481.028610.680.0803020.4457371812040.76Z2: Single grain0.08891237931.50.07120.123990.281.547840.420.0905370.3175495014370.,69Z3: Single grain prism0.0812523233453.80.05670.179320.143.304950.220.1336710.171063148221470.65Z4: 2 Zrn tips0.022061919853.20.02810.093550.161.058680.280.0820770.2257673312470.61M1 (Mnz single xtl)0.0260367441104.70.51920.0786800.220.61660.430.056840.374884884850.52M2 (Mnz single xtl)0.0491588783334440.60.37240.0774160.350.60480.350.056660.044814804780.99M3 (Mnz single xtl)0.04767675410711010.40.41610.0771050.540.60240.540.056670.544794794790.99Metasedimentary gneisses
The upper unit of the tectonic imbricate in Masanteo peninsula is composed of
high-pressure qz-fsp paragneisses with a paragenesis containing ± St ± Ky ± Rt ± Grt. These metasedimentary gneisses preserve
a compositional layering formed by alternations of psammitic and pelitic
layers and are also known as Cariño gneiss (Vogel, 1967). The
paragneisses also show occasional leucosomes and have been strongly deformed
during D2 (Fig. 7a and b). The metapelitic layers are composed of Ky + Grt + Bt +
Ms ± St ± Hbl ± Czo ± Ilm ± Ttn, while the metapsammitic layers lack Ky and the other alumina-rich phases
(Fig. 7c and d). Albert et al. (2015) reported a maximum depositional age
of ca. 510 Ma from this paragneiss, where the source for siliciclastic
detritus was mostly continental. North of our area of study, the paragneisses
are intruded by amphibolitized flaser gabbro (Fig. 2a),
Fernández-Suárez et al. (2002) reported 400 Ma (Ttn) and 386 Ma
(Mnz) U-Pb ages from these gneisses. Peucat et al. (1990) provided a
well-defined muscovite Ar-Ar plateau age of 376 ± 2 Ma from this
gneiss in the Masanteo peninsula, marking the cooling of the whole nappe
below 350 ∘C (muscovite blocking temperature; see discussion in
Ordoñez-Casado et al., 2001).
Relation of deformation structures inside and outside lozenge
migmatitic bodies: (a) sketch showing the trace of the S2
foliation in bounding shear zones and within the lozenge (location of
observations in Fig. 4c); and (b) pole figure of main S2
foliation, intersection lineation and intrafolial hinge lines within the
lozenge in (a). Equal area projection, lower hemisphere projection
also shows the V1 eigenvector and the mean S2-foliation plane. The
arrows indicate the orientation of the horizontal maximum extension
inferred.
Metamorphic peak conditions in metasedimentary gneisses in the Masanteo
peninsula are 700 ∘C and 1.2 GPa (Fig. 5; Castiñeiras, 2005),
consistent with the presence of St and the absence of eclogite or
retroeclogite blocks in them. Peak T conditions are comparable to those
recorded in the underlying migmatitic qz-fsp gneisses, however there is a
difference of 0.5 GPa in peak pressure conditions. This difference in peak
lithostatic pressure conditions represents a difference in depth of
∼ 13 to 15 km between the migmatitic and the metasedimentary gneisses.
Most outcrops examined show a gradual transition between migmatitic and
metasedimentary gneisses, which is accommodated by the intense development of
the blastomylonitic S2 foliation. This contact is quite exposed in the
Serrón beach (Figs. 2a and 12c), where a sub-horizontal shear zone
deflects the S2 foliation, indicating an extensional sense of shear
relative to lithological layering.
Crystallographic preferred orientation (CPO) patterns in quartz and
plagioclase, in relation to the main S2 foliation in the qz-fsp
gneisses. Sample locations are in Fig. 2. Contouring is in multiples of
uniform distribution (Gaussian half width 15). Items indicated in the stereo
plots are the following: bottom left, the J-index; right, the values of
contours. Equal area projection, lower hemisphere. S2 foliation is
plotted E–W vertical and the L2-lineation, if sufficiently developed,
is plotted E–W horizontal. Crystallographic preferred orientation (CPO)
pattern in garnet formed in the S2-tectonic fabric in sample CO4 is
also plotted.
StructureThe main tectonic fabric
Evidence of the structural evolution prior to eclogite facies deformation is
rarely observed in Cabo Ortegal nappe rocks because the main tectonic fabric,
S2, (Figs. 3a, 4c, 7a), is so pervasive. The most common tectonites
formed in relation to shear zones are planar (S-tectonite) or plano-linear
(LS-tectonite). S2 foliation involved the formation of decompressive
textures, such as the growth of large Phg bounded by Bt flakes that enclose
small Grt (Fig. 4d and f), evidencing a fast isothermal decompression
during D2-deformation (Fernández, 1997).
The lozenge-shaped bodies bounded by anastomosing shear zones, which preserve
migmatitic layering within less deformed qz-fsp gneisses. The lozenges
include rootless intrafolial fold hinges and an intersection lineation
between the migmatitic layering and the lozenge shear walls. The orientation
of the intersection lineation can be useful to infer kinematics during
deformation. Eigenvector v1 orientation for the intersection and
intrafolial hinge lines lie sub-parallel to N20E direction (Fig. 8), and the
overall geometry is consistent with bulk strain controlled by flattening
(Ponce et al., 2013).
Coastal sections of the basal thrust (see Figs. 2a and 12a for
locations). (a) Continuous section showing the contact between the
mafic gneisses and the migmatitic gneisses. The white line represents the sea
level. (b) Elliptical sections in sheath folds of decametric size in
the lower domain of the basal thrust. The arrow points to an fisherman for
scale, also used as reference in the sketch outlining the S2 foliation
underneath. (c) Phyllonitic domain in the basal thrust. Structures
related to this domain are outlined in the sketch below the picture.
Crystallographic preferred orientation (CPO) patterns of Qz, Pl, and Grt have
low intensity during the development of LS- and S-tectonites in D2 and
are comparable in metasedimentary and migmatitic gneisses (Fig. 9). The lack
of a well-developed stretching lineation in samples CO4 and CO5 makes its
kinematic interpretation difficult. The preferred orientation of Qz c axes
is characterized by a single girdle of c axes normal to the foliation plane
in sample CO16; and by a single girdle in samples CO4 and CO5 dominated by a
strong maximum within the girdle and parallel to the foliation. Such CPO
patterns are usually found in fabrics formed at medium and high T, in
relation to the dominant activity of the prism <a> and rhomb <a> slip
systems (e.g. Law, 1990).
The basal ductile thrust
The normally pervasive blastomylonitic S2 foliation is disrupted by a
discrete high-strain shear zone. The layering and foliations intersect
relative to the sense of shear indicate that is a basal ductile thrust,
between the mafic gneisses and the migmatitic qz-fsp gneisses (Figs. 3 and
14a). The shear zone has a thickness < 100 m. Three deformation
domains can be differentiated. The associated structures decrease in size and
the domains decrease in thickness towards the upper boundary of the ductile
thrust, indicating the progressive localization of deformation. The lower
domain affects the underlying mafic gneisses along a band ca. 50 m in
thickness. It contains metric- and decametric-sized sheath folds. The
orientation of angular spread of the fold hinges indicates NW–SE stretching
(Fig. 10b).
The middle domain forms in biotite qz-fsp gneisses and includes eclogite
blocks within the migmatites. Migmatitic leucosomic and restitic layers are
interbedded and deformed ductilely. Metric asymmetrical folds develop
vergence to the SE (Fig. 12a and c).
The upper domain contains ∼ 10 m thick phyllonites, frequently
including mafic eclogite pods and boudins. The phyllonites are affected by
shear bands, decimetric sheath folds, superposed folds and rotational
complex structures (Figs. 10c and 11). Superposed shear folds show the type
3 (coaxial) interference pattern of folding (Ramsay, 1967) (Figs. 11 and 12).
The apical axes of sheath folds point towards N20E, indicating maximum
ductile extension along this direction.
Non-cylindrical minor fold associated with the basal thrust: (a)
sheath folds with apical axes perpendicular to the section view; and (b)
type 3 fold interference pattern (after Ramsay, 1967) in the
phyllonitic
domain (see Fig. 10 for location).
Geological sections showing the internal structure of the migmatitic
qz-fsp gneisses and the locations of larger eclogite-blocks. Cross sections
are located in Fig. 2a. The white lines represent the sea level. (a)
In the northern section, the internal structure is characterized by
asymmetrical folding and the presence of eclogite block-in-matrix close to
the thrust. (b) The internal structure of the migmatitic qz-fsp
gneisses is dominated by the presence of polyclinal folds bounded by the
basal thrust and the upper normal detachment. (c) Photograph and
sketch in the cliff of the Serrón beach showing an extensional detachment
placing the metasedimentary gneisses on top of the migmatitic gneisses.
Location of the sections is in Fig. 2a.
The internal structure of the migmatitic gneisses
A group of decametric asymmetric folds, affecting the planar blastomylonitic
S2 foliation, dominates the internal structure. The folds are tight,
overturned and vergent to the SE along the lower part of the migmatitic
gneisses (Fig. 12a). They often have associated parasitic folds, and
non-cylindrical horizontal hinges. Occasionally, minor folds relate to small
thrust surfaces that imbricate eclogite pods parallel to the blastomylonitic
S2 foliation.
The shape of eclogite blocks and boudins was measured in exposed faces
within the gneisses. The representation in a Flinn diagram using the shape
of pods according to block sizes shows in Fig. 13 that most of the large
eclogite blocks plot near to the plane strain field, while smaller eclogite
bodies plot either in the constrictional or flattening fields. However, the
results of this analysis are not a very strong argument because the eclogite
bodies do not show a strongly dominant shape. The shapes were measured in 2-D
sections, and the original shapes of these blocks is unknown. The long axis
of eclogite bodies does not show a preferred orientation (to the right in
Fig. 13).
The ca. 488 Ma felsic dykes can be regarded as passive deformation markers
during D2-deformation. A complex structure has been observed in the
coastal section at the Serrón beach (Fig. 12b). In this section, the
thickness of the migmatitic qz-fsp gneisses is less than 100 m, and both
bottom and top boundaries of this unit are quite exposed. Their thickness
decreases progressively towards the SE. Migmatitic gneisses are affected by a
shear zone in which the sense of the shear changes between the top and the
bottom, producing rootless folds of opposite vergence in the felsic
dioritic/granodioritic dykes and in the migmatitic banding. The larger
structure reconstructed from both markers (the felsic dykes and the
migmatitic banding) consists of an opposite vergence recumbent hinge defined
by the competent dioritic dykes. The limbs are disrupted and boudinaged
towards the horizontal high-strain zones located at the boundaries of the
unit. This sandwiched structure indicates orthogonal stretching with the
transport direction for the migmatitic gneisses towards the SE (Fig. 12b).
The top detachment
A horizontal discrete shear zone separating the metasedimentary and the
migmatitic gneisses is exposed at the Serrón beach (Fig. 12b and c). A
gradual transition between both types of gneisses is observed along the base
of the cliffs. Deformation partitions into anastomosing D2-shear bands
preserving evidences of previous melting episodes (Figs. 4e and 7a).
The horizontal shear zone is 20 m in thickness and strongly deflects the
migmatitic layering in extensional manner. Migmatitic layering and diorite
dykes are disrupted and boudinaged progressively towards the upper
high-strain surface (Fig. 12c). The deflection of the migmatitic layering,
parasitic “drag” folds and the boudinage of the dioritic dykes indicate top
to NW shear sense. Despite subsequent re-equilibration under
greenschist-facies conditions, evidencing a late reactivation, the mineral
assemblages in the progressively less deformed bands within the detachment
are basically the same as the high-grade qz-fsp gneisses described
previously (Fig. 5).
The shapes of the eclogite pods show a range of geometries in a
Flinn diagram from prolate to oblate. The size of the symbols is
proportional to size of the blocks.
The upper D3-recumbent fold
The metasedimentary qz-fsp gneisses lie in the core of a km-scale
D3-recumbent/overturned-synformal structure, outcropping towards the
east side of the Masanteo peninsula (Fig. 1b and d shows the location of the
fold within the general cross-section of Cabo Ortegal nappe, after Marcos et
al., 2002). This large-scale fold opens to the SE and has several parasitic
cylindrical-folds and an associated crenulation cleavage. Intrafolial and
sheath folds formed during the development of the S2 foliation (Fig. 15c and d) are refolded by parasitic D3-folds related to the recumbent
fold (Fig. 14b). Detailed cross-sections of the recumbent structure have been
constructed using the asymmetry of small-scale parasitic folds and the main
S2 foliation (Fig. 14). The fold axis plunges 5–30∘
towards N20E. The fold attitude requires that the reverse limb outcrops in
the northeastern cliffs of Cabo Ortegal and only partially along the
southeast shoreline. A late upright antiform refolds the recumbent synform
(Fig. 15b and e). This late folding affects the crenulation cleavage (Fig. 14c), which equilibrated under greenschist-facies conditions.
Geological sections of the D3-recumbent syncline reconstructed
from the small-scale parasitic folds that are folding the main S2
foliation. Location of the sections is in Fig. 2a. White lines represent the
sea level. (a) Northern outcrop section. (b) Southern
outcrop section and the structural detail with location of Fig. 15. Note that
the recumbent synform is affected by open-upright D4b-folds.
(c) Crenulation cleavage S4b, D3 fold axes and L2-4b
intersection lineation is plotted in an equal area, lower hemisphere
projection.
Small-scale parasitic folds related to the recumbent synform folding
prior to D2-isoclinal folds. Locations of outcrops are indicated in
Figs. 3 and 14. (a) Sketch of the outcrop section. (b) W–E
view of a monocline, with location of the photographs (c),
(d), and (e). (c) The apical-section of a
D2-sheath-fold, behind the hammer, indicates a N–S stretching line.
(d)D2-intrafolial folds folded by a “Z” parasitic
D3-fold (reverse limb of the D3-recumbent fold). (e) “Z”
parasitic D3-fold rotated by the D4b-monocline.
Metamorphic evolution in the gneisses
The gneissic sequence in the Masanteo peninsula shows a complex tectonic
record, including pre-Variscan and Eo-Variscan events. HP-HT metamorphism
followed by rapid decompression leading to the formation of the D2
tectonic fabric, based on the M2 metamorphic assemblages defining the
main foliation in the qz-fsp migmatites, eclogites and mafic granulites
(Gil-Ibarguchi et al., 1990; Fernández, 1997; Galán and Marcos, 2000;
Ábalos et al., 2003).
Synthetic tectonic evolution of qz-fsp gneisses within the Cabo
Ortegal nappe. (a) Simplified geological section of the Cabo Ortegal
nappe at present, showing the detail of the structure in the Masanteo
peninsula (cross section in Fig. 1d, modified from Marcos et al., 2002).
Partially restored section to the left shows the relationships between
D3-recumbent folds and D4a-thrust. (b)D2-stage
(390–380 Ma) summarising the effect of progressive localization of
deformation and thinning during exhumation. The P–T diagram (based on
Fig. 5) is showing the retrograde P–T-t paths for migmatites,
metasedimentary qz-fsp gneisses and mafic granulites compared to current
subduction zones. (c) Slab breakoff could eclogitize the top of the
mafic granulites during a widespread thinning of the tectonic sequence
(Llana-Fúnez et al., 2004). (d) A Cambro-Ordovician melting
event led to the injection of felsic dykes at the base of the qz-fsp gneisses
coeval with leucosome in the mafic granulites (Fernandez-Suarez et al.,
2002). Arrows indicate the inferred finite strain principal displacements.
The Cambro-Ordovician intrusion of the felsic dykes at ca. 488 Ma
demonstrates a pre-Variscan tectonothermal event (Table 1; Fig. 6). The
488–486 Ma monazites from these discrete felsic dykes coincide in age with
the monazite ages reported by Fernández-Suárez et al. (2002) also in
migmatitic gneisses from other localities in the Cabo Ortegal nappe. These
monazites have survived the high-pressure Devonian overprint. This could
point to the presence of an early Ordovician metamorphic event similar to the
one described by Abati et al. (1999) in equivalent units of the Órdenes complex. Such a pre-Variscan event would be consistent with the presence of
coronitic textures in the Cabo Ortegal metagabbros (Galán and Marcos,
1997; Marcos et al., 2002). However, to be able to tie these metastable early
Ordovician monazites from Cabo Ortegal to a particular set of metamorphic
conditions remains to be conclusively demonstrated. The high-grade eclogitic
event is well-established by the ca. 390 Ma U-Pb Zrn age from the eclogitic
boudin in the migmatitic gneiss, (sample COZ4 located in Figs. 2a and 12a;
Castiñeiras et al., 2010). In addition, the ca. 376 Ma muscovite Ar-Ar
date from the paragneisses in the Masanteo peninsula provides a cooling age,
which would be consistent with rapid cooling (ca. 90 ∘C Ma-1) and a
relatively fast exhumation rate (5.8 mm yr-1) associated with the D2
deformation, following the eclogitic event (see discussion of U-Pb and Ar-Ar
data in Ordoñez-Casado et al., 2001).
DiscussionTectonic evolution of gneisses at Masanteo
The tectono-metamorphic and geochronological data reported in this paper
based on the observations around the Masanteo peninsula are grouped into
three stages (Fig. 16). The first stage is characterized by the assembly of
a high-grade tectonic sequence composed of mafic granulites overlain by
qz-fsp gneisses. The partial melting of the mafic granulites led to the
injection of the felsic dykes into the base of the gneisses during the early
Ordovician (ca. 490 Ma).
A Devonian subduction, producing the eclogite facies metamorphism dated at
390 Ma (Peucat et al., 1990; Santos-Zalduegui et al., 1996;
Ordóñez-Casado
et al., 2001; Fernández-Suárez et al., 2002, 2007; Castiñeiras et al.,
2010), occurred 20 Ma prior to the main Variscan subduction at ca. 370–360 Ma (i.e. Martínez-Catalán et al., 1997; Arenas et al., 2014;
López-Carmona et al. 2014) and suggests the development of a complex
collision during the assembly of Pangea.
The qz-fsp gneisses underwent an episode of partial melting after
eclogitization (at ca. 390 Ma). Migmatitic qz-fsp layers are heterogeneously
mylonitized along anastomosing shear bands that progressed to planar shear
zones, incorporating eclogite blocks during D2 (Figs. 3a and 10a).
D2-tectonic fabrics have similar high temperature CPO patterns in
migmatitic and metasedimentary qz-fsp gneisses consistent with flattening
strain (Fig. 9). In addition, bulk flattening strain is also supported from
the lozenged overall structure and the scattered orientation of the kinematic
markers. Most of the tectonic pile thinning occurred during the development
of the blastomylonitic S2 foliation (i.e. Fernández, 1997;
Llana-Fúnez et al., 2004) and before the progressive localization of
strain. Lack of a persistent stretching lineation suggests non-plane strain,
with orogeny parallel stretching accompanying shearing displacements.
Additional thinning could have progressed through the reactivation of the
NW-vergent top detachment (Fig. 12b and c). The common metamorphic and
structural evolution for most of the rock units within the Cabo Ortegal nappe
indicates a tectonic setting of coherent slab subduction in which extensive
eclogitization and significant ductile thinning of the tectonic sequence are
coeval. A scenario of downdip extension during ductile breakoff was suggested
to produce observed structures and metamorphic imprint (Llana-Fúnez et
al., 2004).
The tectono-metamorphic relationships of the basal ductile thrust and the
normal detachment mapped in the Masanteo peninsula indicate that both
discrete mechanical contacts were active before the development of the
recumbent folding that affects the sequence of gneisses. These mechanical
contacts upon their development became in fact the boundaries of the
migmatitic qz-fsp gneisses (Figs. 1b and 10c). The arrangement of the
bounding shear zones defines an inclined E-dipping wedge with the migmatitic
qz-fsp gneisses in the middle.
The internal structure of the migmatites consists of a rootless double
recumbent hinge (Fig. 12). The metric sheath folds belonging to the mafic
gneisses of the basal ductile thrust are consistent with an extension
direction towards the SE (Fig. 10b). In addition, progressive localization of
strain occurred during exhumation. Frequently, Phg phenoblasts are aligned
parallel to the S2 foliation of the migmatitic qz-fsp gneisses, and are
bounded by Bt flakes that enclose small prismatic shaped Grt (Fernández,
1997). These microstructures evidence the instability of Phg during
decompression and indirectly point to exhumation at high rates. The starting
point for the exhumation of the migmatitic qz-fsp gneisses and the
metasedimentary qz-fsp gneisses differs by 0.5 GPa (Fig. 5) and indicates
that they have been initially juxtaposed by shearing displacements. The
metamorphic pressure difference between both types of gneisses could be
indicative that metasedimentary qz-fsp gneisses exhumed from maximum burial
depths of ∼ 13 to 15 km lower than the migmatitic qz-fsp gneisses if we
consider metamorphic pressure as lithostatic. However, in the context of
subduction and exhumation of high-pressure rocks, the architecture of the
qz-fsp migmatites, with a basal thrust and a top detachment, in addition to a
tectonic regime dominated by flattening may suggest the development of a
non-lithostatic pressure gradients leading to tectonic (vertical) extrusion.
In this scenario, the reported difference in metamorphic pressure may contain
a component of non-lithostatic stress (taking lithostatic stress as simply
due to depth of burial), which will affect the conversion to depths. For
instance, an overpressure (elevated mean stress) of 1.1 or 2 times the
lithostatic pressure (e.g. Mancktelow, 1995, 2008; Moulas et al., 2013) would
imply a difference in burial depths for the gneisses between 11 and 5.5 km, respectively.
The migmatite wedge within the qz-fsp gneisses described here in the Cabo
Ortegal nappe, is similar to the gneiss wedge reported within the
high-pressure terranes of the Sambagawa HP rocks, sharing exhumation
characteristics from lower to upper crustal levels (Osozawa and Wakabayashi,
2015). Another example comparable to what is reported here occurs in the
exhumation of blueschist facies rocks of Leti Island in Indonesia
(Kadarusman et al., 2010).
The final stage of large-scale structure development is dominated by the
progressive deformation associated with the Variscan convergence,
corresponding to the formation of kilometric-scale recumbent folds, thrusts
and finally folding into upright or SE verging folds, described here for Cabo
Ortegal rocks in the Masanteo peninsula as D3, D4a, and D4b,
respectively (Fig. 16a). The late evolution of the Cabo Ortegal nappe and
its kinematics (Marcos et al., 2002) is consistent and coeval with the
deformation recorded in the underlying autochthonous rock sequence in
relation to the Variscan belt (Matte, 1968; Pérez-Estaún et al.,
1991), and led to the crustal thickening necessary to drive the deformation
and metamorphism of the autochthon beneath.
Characteristics of the exhumation of Cabo Ortegal nappe rocks
The development of the main tectonic foliation in Cabo Ortegal nappe rocks,
defined by mineral assemblages from eclogite facies to greenschist facies,
indicates that the structure formed during the progressive exhumation from
high-pressure conditions to mid-crustal depths (Galán and Marcos, 2000;
Marcos et al., 2002) (Fig. 5). In this contribution, we would like to
highlight that the metamorphic evolution of two major rock units in the Cabo
Ortegal nappe, the mafic granulites and the qz-fsp gneisses, have parallel
P–T trajectories during their tectonic exhumation from eclogite to
greenschist facies conditions (Fig. 16b). In the first instance, this
suggests a common process and scenario for the exhumation of high-pressure
rocks to mid-crustal depths prior to their final emplacement. In fact, it is
possible to characterize the exhumation path in terms of cooling during
decompression (the common slope of P–T paths in Fig. 16b), which may be
regarded as indicative of this particular tectonic setting. Given that ages
are available at peak pressure conditions and at mid-crustal depths,
exhumation rates can also be calculated from the data summarized before.
The exhumation of either the mafic granulites or the gneisses in Fig. 16b
involves a cooling of > 200 ∘C for a
decompression of 1 GPa, most of it associated with the vertical component of
movement in the crust. For a similar depth range, the amount of cooling
during decompression is comparable to present-day subduction zones, in
particular to the trajectory in the south-west Japan subduction zone
(trajectory SWJ from Fig. 4 in Yamasaki and Seno, 2003 in included in Fig. 16b). The positive P–T slopes are characteristic of subducting slabs in
active subduction zones (e.g. Hacker et al., 2003; Yamasaki and Seno, 2003).
Therefore, based on this observation, we infer that the exhumation of Cabo
Ortegal nappe rocks began while the subduction zone was still during the
initial stage of exhumation. This is in contrast to P–T paths from equivalent
high-pressure rocks in the Órdenes Complex, where adiabatic
decompressive paths were inferred during exhumation from eclogite facies
conditions (Arenas et al., 1995; Martínez-Catalán et al., 1997).
Adiabatic decompressive paths are related to the very rapid exhumation of
small fragments, not allowing cooling during ascent, however this is
unlikely in the case of coherent slabs (Kylander-Clark et al., 2012).
The available ages at peak metamorphic conditions and at mid-crustal depths
provide time constraints for the exhumation process, which might involve at
least 140 km of horizontal displacement onto the continental margin. A rate
of > 3 mm yr-1 vertical component for migmatitic gneisses can be
estimated, if all metamorphic pressure is regarded as corresponding to
lithostatic pressure. These rates are half of those previously reported
(Ordóñez-Casado et al., 2001), and are comparable to those reported
in other parts of the Variscan orogeny, the Alps or the Himalayas. When
compared with present-day subduction zones and numerical models the
exhumation rates fall at the slow end member of scenarios (Gerya et al.,
2002; Kylander-Clark et al., 2012; Burov et al., 2014). If the dislocation
creep deformation mechanism observed (from tectonic fabrics and
microstructures) to control internal deformation in part of the gneissic
units (Fig. 9) were to be dominant across all of the Cabo Ortegal nappe it
would imply, according to numerical models by Gerya and Stöckhert (2002), subduction rates 6 times higher than exhumation rate, still at the
slower end for other subduction zones. The metamorphic evolution during the
exhumation of Cabo Ortegal nappe rocks shows a significant difference with
the inferred P–T path in other ancient and current subduction zones compiled
by Kylander-Clark et al. (2012): the exhumation path is overall 100–200 ∘C hotter than the warmest currently active (e.g. SW Japan
or the Pacific Northwest), and most of the exhumation took place under conditions of slab
melting. Melting occurring during a pervasive deformation episode will
impart extreme weakening and ductile behaviour to rocks (e.g. Rosenberg and
Handy 2005; Rutter et al., 2006), ultimately to the whole crustal section.
This environment should promote widespread thinning in the subducting slab
in which the Cabo Ortegal nappe was involved during plate convergence prior
to the Variscan collision.
Conclusions
The structural relationships of the high-grade mafic and qz-fsp gneissic
bodies along the Masanteo peninsula provide valuable insights to understand
the polyorogenic origin of the HP-HT Cabo Ortegal nappe. The
Cambro-Ordovician (ca. 480–490 Ma) intrusion of felsic dykes within the
migmatitic qz-fsp gneisses was coeval with the formation of leucosome in the
underlying mafic granulites. A second melting event partially affected the
qz-fsp gneisses coevally with the eclogitization of the whole lower crustal
section during an early Variscan high-pressure episode (ca. 400–390 Ma).
During the subsequent fast exhumation (ca. 380 Ma) the original rock
sequence was largely thinned by dominant bulk flattening associated to the
development of a main blastomylonitic foliation. Immediately after, at ca. 360 Ma, Cabo Ortegal nappe rocks, already at mid-crustal depths, were
involved in the final Variscan collision that emplaced the ensemble of
allochthonous complexes over the Iberian microplate, at the edge of
Gondwana.
Progressive strain localization during exhumation triggered the development
of anastomosing shear bands, enclosing lenticular or lozenge-shaped bodies.
Strain weakening associated with hydration and retrogression during
deformation in bounding shear zones prevented further pervasive deformation
and retrogression into the lozenges. The geometric arrangement of ductile
shear zones bounding the gneisses at separate tectonic stages during the
exhumation, forming a basal ductile thrust and a top detachment, gave way to
the development of an internal migmatite wedge within the qz-fsp gneisses,
consequence of the localization of strain during bulk thinning.
The exhumation path from eclogite facies conditions to mid-crustal depths in
qz-fsp gneisses, estimated in this contribution around
2 ∘C/100 MPa, is parallel to the trajectory in adjacent mafic
high-pressure granulites and to the calculated distribution of temperature
with depth in currently active subduction zones. This suggests that the
metamorphic and also the structural record associated with the process of
exhumation is comparable to present-day tectonic scenarios. It must be
highlighted that the absolute temperatures for the exhumation path are
substantially higher with respect to current settings, by approximately
100–200 ∘C. Such high temperatures can put the exhumation path into
slab melting conditions, which ultimately would favour an extreme ductile
behaviour of the whole rock sequence during deformation. This may have been
the case during the stretching of the lower crustal rock sequence now
preserved in the upper unit of the Cabo Ortegal nappe.
Francisco José Fernández and Alberto Marcos carried out the fieldwork and mapping. Sergio Llana-Fúnez measured
and plotted the crystallographic preferred orientation of qz-fsp gneisses.
Pablo Valverde-Vaquero and Pedro Castiñeiras determined the ages of the diorite dykes and the eclogite
sample, respectively. Francisco José Fernández prepared the manuscript with contributions
from all co-authors.
Acknowledgements
In 1988, Francisco José Fernández initiated his research career in the Masanteo peninsula
under the supervision of Professor Alberto Marcos and aimed by
Andrés Pérez-Estaún. Revisiting the area 25 years later brings
new light and some understanding to Cabo Ortegal geology. Authors thank
their colleagues for continuing discussion about the tectonic evolution of
Cabo Ortegal. Research funds from grants CGL2011-22728, CGL2011-23628/BTE,
CGL2012-38786 and CGL2014-53388-P by the Spanish government are
acknowledged. Pedro Castiñeiras stay at Stanford University was funded by CSIC grant
PA1002435. This paper benefitted from the general and specific comments of
the reviewers Neil S. Mancktelow and Thomas Blenkinsop. We thank the editor
Ernie Rutter for extensive feedback and criticism on the earlier versions
of the manuscript.
Edited by: E. H. Rutter
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