Folding in orogens: a case study in the northern Iberian Variscan Belt

GEOLOGICAL JOURNAL

Geol. J. 45: 597–622 (2010)

Published online 19 January 2010 in Wiley Online Library

(wileyonlinelibrary.com) DOI: 10.1002/gj.1199

Folding in orogens: a case study in the northern Iberian Variscan Belt

F. BASTIDA 1*, J. ALLER 1, J. A. PULGAR1, N. C. TOIMIL 2, F. J. FERNANDEZ 1,N. C. BOBILLO-ARES 3 and C. O. MENENDEZ 3

1Departamento de Geologıa, Universidad de Oviedo, Oviedo (Asturias), Spain2Hulleras del Norte S.A. (HUNOSA), Oviedo (Asturias), Spain

3Departamento de Matematicas, Universidad de Oviedo, Oviedo (Asturias), Spain

Folding during the Variscan deformation in NW Iberia has been analysed from the foreland to the hinterland of the orogen using severalgeometrical techniques complemented by the numerical simulation of kinematical folding mechanisms with the aid of a computer program(‘FoldModeler’). In the foreland, folds are related to thrusts; in the inner zones, folds can be attributed to three deformation phases. D1 and D2

involved deformation with horizontal foreland-directed displacements. D1 gave rise to closed or tight folds (F1) and cleavage (S1); develop-ment of large recumbent F1 folds in the hinterland required a simple shear regime and a final coaxial strain component with sub-verticalmaximum shortening. Strain incompatibilities at deeper levels, together with high temperatures, favoured the concentration of ductiledeformation in shear zones and development of mylonites during D2. Local flow instabilities generated F2 folds that were passively amplifiedby a combination of simple shear, coaxial strain and area change. D3 involved a change to a regime with dominant coaxial deformation and asub-horizontal maximum shortening; it gave rise to upright or steeply inclined folds (F3). Development of D3 structures was heterogeneousand depended on the previous dip of the bedding and S1, the presence or emplacement of granitoids, the stacking of thrust sheets or theprevious development of large faults bringing into contact rocks with different competence. D3 structures are mainly concentrated inmetapelitic areas and appear distributed in bands on several scales. On a small scale, tectonic banding appears associated with small folds as aresult of pressure-solution processes. In areas with sub-vertical S1, later sub-horizontal kink bands were formed by a vertical compression.Copyright # 2010 John Wiley & Sons, Ltd.

Received 27 March 2009; accepted 6 October 2009

KEY WORDS structural analysis; folding; shear zones; strain; Variscan Belt; Iberia

1. INTRODUCTION

Folds are very common geological structures whose

geometrical, kinematic and dynamic analysis can illustrate

the deformation undergone by rocks and the stresses that

drove this deformation. In addition, folds form an essential

part of the framework of orogenic belts and knowledge of

their time–space patterns is a key contribution to the analysis

of these parts of the lithosphere. Despite the interest in the

investigation of these structures in the context of orogens,

this type of study is scarce in the geological literature and

most of the regional studies place little emphasis on folding

analysis. On the other hand, studies of natural folds are

common and have been made with varied methodologies,

but are related in general to specific structures or to folds

developed over a small area. Nevertheless, there are some

* Correspondence to: F. Bastida, Departamento de Geologıa, Universidad deOviedo, Jesus Arias de Velasco s/n, 33005 Oviedo (Asturias), Spain.E-mail: [email protected]

fold studies that are remarkable from the point of view of

understanding folding in orogenic belts. This is the case with

papers on the large recumbent folds that appear in most of

the orogens, such as those in the Alps (Ramsay, 1981;

Ramsay et al., 1983; Dietrich and Casey, 1989; Butler, 1992;

Rowan and Kligfield, 1992) and the Iberian Variscan belt

(Fernandez et al., 2007). The detailed studies made by

Hudleston (1973a) on folds of Loch Monar (Scotland),

Sanderson (1979) and Rattey and Sanderson (1982) on folds

from the Variscan fold belt of Southwest England and Toimil

and Fernandez (2007) on folds from the Iberian Variscan belt

are also noteworthy. In some of these studies, the role of

simple shear or sub-simple shear (combination of simple

shear and pure shear; De Paor, 1983; Simpson and De Paor,

1993) has been recognized as an essential bulk strain regime

in the development of folds. For this reason, theoretical and

experimental studies on folding in a simple or sub-simple

shear regime have been made by several authors (Ghosh,

1966; Manz and Wickham, 1978; Ez, 2000; Carreras et al.,

2005). Some authors have emphasized the influence of

Copyright # 2010 John Wiley & Sons, Ltd.

598 f. bastida ET AL.

gravity and the existence of heterogeneities in the multilayer

on the development of recumbent folds (Bucher, 1956, 1962;

Blay et al., 1977; Vacas Pena and Martınez Catalan, 2004).

Other analyses have attempted to show the variation of

attitudes or styles of folds in orogenic belts or in ductile

orogenic wedges (for example, Fyson, 1971; Chattopadhyay

and Mandal, 2002).

The aim of this paper is to analyse the time–space

development of folds through the Variscan Belt of NW

Spain. This sector includes all the parts of an orogen, from

the hinterland to the foreland, and has the advantage of

presenting a complete cross-section with continuous out-

crops along the Cantabrian coast. We will try to illustrate the

variation of the geometrical characteristics of folds and the

kinematic mechanisms that operated in their development

through this sector. Some mechanical considerations to

Figure 1. Geological map of the Northern Iberian Variscan Belt wi


explain these geometrical and kinematic characteristics will

also be presented.

2. GEOLOGICAL SETTING

The Variscan Belt in NW Iberia has been divided into several

major zones with different geological characteristics

(Julivert et al., 1972; Farias et al., 1987) (Figure 1):

Cantabrian zone, Westasturian-Leonese zone, Ollo de Sapo

domain (Central-Iberian zone) and Galicia-Tras-os-Montes

zone. The general trend of the structures in the belt describes

the Ibero-Armorican Arc (or Asturian Arc). The regional

geology of the different units of the orogen has been studied

by many authors; modern geological syntheses of the whole

belt can be found in Gibbons and Moreno (2002) and Vera

th the location of the cross-sections shown in several figures.

Geol. J. 45: 597–622 (2010)

folding in iberian variscan belt 599

(2004). An evolutionary model of the belt was provided by

Perez-Estaun et al. (1991).

The Cantabrian zone represents the external part of the

orogen in this sector of the Variscan Belt and forms the core

of the Ibero-Armorican Arc. Two tectonostratigraphic units,

one preorogenic and the other synorogenic, can be

distinguished in this zone (Julivert, 1978; Marcos and

Pulgar, 1982). The former is formed by Palaeozoic pre-

Carboniferous and early Carboniferous series with an

alternation of carbonate and siliciclastic formations and

has a wedge shape, thinning towards the foreland. The latter

is formed by sediments of mostly Carboniferous age and

consists of several clastic wedges which are a result of the

filling of foredeep basins formed in front of the major thrust

units. Late Variscan unconformable Stephanian rocks with

molassic facies appear in several areas of the Cantabrian zone.

The structure mainly consists of thrusts directed towards the

Asturian arc core and fault-related folds. Deformation

occurred in a thin-skinned regime, with low strain, and

mainly under diagenetic conditions.

The Westasturian-Leonese zone is located to the west of

the Cantabrian zone. The boundary between the two zones is

formed by a Precambrian outcrop named the ‘Narcea

antiform’ (Figure 1). The zone mainly consists of extremely

thick Lower Palaeozoic rocks, mainly siliciclastic with the

exception of a Lower-Middle Cambrian carbonate level, and

represents the transition to the hinterland of the orogen in

such a way that the orogenic metamorphic grade increases

westward up to the amphibolite facies; the granitoid

outcrops occupy a large area in the western part of the

zone. Deformation is polyphase, three phases having been

distinguished (Marcos, 1973). During the first deformation

phase (D1), closed to tight folds (F1) vergent towards the

foreland with sub-horizontal hinges and associated cleavage

(S1) were formed. During the second deformation phase

(D2), thrusts and ductile shear zones, also vergent towards

the foreland, were developed; the ductile character of the

shear zones mainly appears in the western part of the zone,

where they have associated folds (F2), mylonites and

schistosity (S2 or S1þ2). During the third phase of

deformation (D3), upright open folds with associated

crenulation cleavage (S3) were formed; these folds are

almost homoaxial with F1 folds. Post D3 structures, mainly

sub-horizontal kink bands and normal faults, were also

developed. For the folding analysis it is convenient to divide

the Westasturian-Leonese zone into two units (Marcos,

1973) (Figure 1): the Mondonedo nappe unit to the west, and

the Navia-Alto Sil unit to the east. The two units are

separated by an important thrust: the Mondonedo nappe

basal thrust. In areas with low metamorphism (chlorite

zone), this is syn-kinematic with S1 (Marcos, 1973), but in

areas with higher metamorphic grade (western part of the

Mondonedo nappe unit), where a thick ductile shear zone


exists, three metamorphic episodes can be differentiated

(Bastida et al., 1986): M1, pre-D2, with a paragenesis

containing staurolite and garnet; M2, early syn-D2, with

andalusite; and M3, late to post-D2 (end of the ductile

deformation in the shear zone and development of the basal

thrust), presenting retrograde metamorphism to greenschist

facies.

The Ollo de Sapo domain is located to the west of the

Westasturian-Leonese zone. The boundary between them is

the Viveiro Fault, an important normal fault longitudinal to

the general trend of the major structures (Figure 1). This unit

is characterized by a siliciclastic Lower Palaeozoic succes-

sion with a basal porphyroid: the ‘Ollo de Sapo Formation’,

Lower Ordovician in age (Gebauer et al., 1993; Valverde-

Vaquero and Dunning, 2000; Montero et al., 2007). The

sequence of deformation phases in this unit is comparable to

that of the Westasturian-Leonese zone, although in general

F3 folds are more developed (Matte, 1968). The regional

metamorphism in this domain ranges from the chlorite to the

sillimanite - K-feldspar zone. The metamorphic peak is post-

D1 to pre- or syn-D3.

The Galicia-Tras-os-Montes zone is separated from the

Ollo de Sapo domain by a basal thrust and an important fault

(Pontedeume-Valdovino Fault). Two very different domains

have been distinguished in the Galicia-Tras-os-Montes zone:

Galicia-Tras-os-Montes Schist domain and Allochthonous

Complexes of Galicia-Tras-os-Montes.

The Galicia-Tras-os-Montes Schist domain is formed by a

thick succession of siliciclastic rocks, mainly metapelites,

probably chiefly Ordovician in age (Farias and Marcos, 2004;

Valverde-Vaquero et al., 2005). The succession of defor-

mation phases in this zone hasbeen considered the sameas that

in the Westasturian-Leonese zone and the Ollo de Sapo

domain (Marquınez, 1984; Farias, 1990; Farias and Marcos,

2004), although there are important differences. It has been

interpreted that the main regional foliation is the S2, which

presents a generalized character and obliterates to a large

extent the S1. S2 is a domainal cleavage with relict microfolds.

D1 structures are very scarce and major D2 structures are rare.

D3 structures are common and well developed. In this domain,

a progressive metamorphism exists from the chorite zone up to

a migmatitic zone, with grade increasing towards the granitoid

areas that occupy most of the domain. The metamorphic peak

occurred during the inter-phase D2–D3.

The Allochthonuous Complexes of Galicia-Tras-os-

Montes appear isolated as great klippen in the core of

synforms. They are formed by metasediments and mafic and

ultramafic rocks that underwent high-grade metamorphism.

Each of these complexes contains several units, some of

them ophiolitic, separated by thrusts. Their internal structure

is complex with a structural history longer than that of the

autochthonous rocks. The correlation and comparison of

their folds with those developed in the autochthonous rocks

Geol. J. 45: 597–622 (2010)


is difficult and they are not considered in the present

analysis, except for the late structures (D3), which affect

both the allochthonous and autochthonous rocks.

Figure 3. Diagram of t0a2 against sin2a, showing the definition of t012, t022

and angles b1 and b2 from a curve representative of a fold (after Bastidaet al., 2005).

3. METHODS

The regional analysis of folds is based on the use of several

geometrical and kinematical methods. In addition to

measurements of several standard elements, such as

interlimb angles and the orientation of axes and axial

surfaces of folds, other parameters are used to characterize

the geometry of the profiles of folded surfaces and folded

layers. Knowledge of this geometry is necessary for the

kinematical analysis of folds.

Two parameters are used to analyse the geometry of

folded surfaces: the aspect ratio (h¼ y0/x0) and the

normalized area (A¼ 2A/x0 y0) (Bastida et al., 1999,

2005; Aller et al., 2004; Srivastava and Lisle, 2004; Lisle

et al., 2006). The meaning of A, x0 and y0 is shown in

Figure 2. Parameters s1 and s2 are used for the analysis

of folded layers. These parameters were defined by

Bastida et al. (2005) to synthesize the results obtained

from Ramsay’s (1967) classification. For their definition let

us consider a curve of t0a2 vs. sin2a for a fold limb

(t0a ¼ normalized orthogonal thickness; a¼ dip) (Figure 3),

and two points, A and B, on this curve. A is the point of the

curve where the abscissa equals (sin2am)/2, and B is the final

point of the curve with abscissa sin2am (am is the maximum

dip of the folded layer). After drawing the line segments, OA

and AB, we define s1 and s2 as (Figure 3)

s1 ¼ tan b1 ¼ 2ð1 � t012Þsin2 am

(1)

s2 ¼ tan b2 ¼ 2ðt012 � t022Þsin2 am

(2)

Figure 2. Reference system and geometrical elements of a folded surfaceprofile. H, hinge point; I, inflection point; A, area characterizing the shape of

the limb profile for specific values of x0 and y0.


Thus, a simple s1 vs. s2 graph offers a means of classifying

folds, since each fold limb can be represented by a single

point.

The analysis of the kinematic folding mechanisms has

been made using a new version of the computer program

‘FoldModeler’ (Bobillo-Ares et al., 2004), which allows the

numerical modelling of folds by any simultaneous or

successive superposition of flexural flow, tangential longi-

tudinal strain with or without area change, heterogeneous

simple shear parallel to the axial trace and any type of

homogeneous strain. The application of this program to

natural folds requires at least a knowledge of the geometry of

the profiles of folded surfaces and layers and the pattern of

the tectonic foliation distribution through the folded layer

profile. This pattern can be characterized by a curve that

shows the variation of the foliation dip as a function of the

layer dip.

4. GEOMETRICAL ANALYSIS OF THE FOLDS OF

THE MAIN FOLDING PHASE (F1 FOLDS)

The best development of F1 folds occurred in the

Westasturian-Leonese zone and in the Ollo de Sapo domain.

They are very scarce in the Galicia-Tras-os-Montes zone,

and they only appear in the Cantabrian zone near Cabo Penas

(Figure 1). In the latter zone, the first structures were thrusts

Geol. J. 45: 597–622 (2010)


with associated fault-bend folds and fault-propagation folds

whose characteristics are very different to those of F1 folds.

The strong development of F3 folds in the Ollo de Sapo

domain hampers a detailed analysis of the F1 folds in this

area. Furthermore, the Viveiro Fault represents a major

structural and metamorphic discontinuity that makes it

difficult to establish the original spatial variation of the F1

structures along a traverse of the orogen. Nevertheless, this

discontinuity is less pronounced in the south-eastern part of

the domain, where a large F1 structure crops out: the Courel

recumbent syncline (Matte, 1968; Martınez-Catalan, 1985;

Fernandez et al., 2007) (Figure 1). As a consequence, we

give here a detailed analysis of the F1 folds of the

Westasturian-Leonese zone, the north-western sector of the

Cantabrian zone (Cabo Penas sector) and the south-eastern

sector of the Ollo de Sapo domain, placing special emphasis

on the traverse along the Cantabrian coast. Previous detailed

studies on F1 folds from this traverse were made by Bastida

(1980) and Toimil (2005). Most of the F1 folds analysed

were developed in comparable lithology, that is, on a

multilayer composed of an alternation of metasandstones

and metapelites of the Candana Group (Lower Cambrian) or

the Cabos Series (Middle Cambrian to Lower Ordovician).

Deviations from this multilayer type will be described in

specific cases.

Analysis of the major F1 folds along the Cantabrian coast

(Figures 4 and 5) allows some conclusions to be drawn about

Figure 4. Geological cross-section of the Westasturian-Leonese zone along the Canand axial surfaces (bottom row) of minor F1 folds fr


their geometry. They are asymmetric folds verging to the

foreland. The length of the overturned limb (measured

between two adjacent hinges) decreases towards the

foreland, from more than 10 km for some major folds of

the Mondonedo nappe unit to 2–4 km for the major folds of

the Navia-Alto Sil unit. The ratio between the lengths of

normal and overturned limbs, that is, the asymmetry of F1

folds, increases towards the foreland; this is a consequence

of the F1 folds being more abundant in the Mondonedo

nappe unit than in the Navia-Alto Sil unit. The interlimb

angle of major F1 folds increases from near isoclinal folds in

the Mondonedo nappe unit to tight or closed folds in the

Navia-Alto Sil unit. Although the superposition of F3 folds

on F1 folds probably modified the original interlimb angle of

the major F1 folds, the variation pattern observed probably

reflects an original feature of F1 folds.

The present attitude of F1 folds changes along the

Cantabrian coast traverse, mainly due to an effect of later F3

deformation (Figure 4). In the Mondonedo nappe unit, the

folds face downwards in the western part, are recumbent in

the central part and are near upright in the eastern part. In the

Navia-Alto Sil unit, the F1 folds range from recumbent in

some localities of the western part to near upright in the

eastern part. The original attitude of F1 folds is discussed

below.

The geometry of minor F1 folds can be analysed in greater

detail than that of major F1 folds (Figure 6). They have sub-

tabrian coast with stereographic plots of the orientation of the axes (top row)om several sectors. See Figure 1 for location.

Geol. J. 45: 597–622 (2010)

Figure 5. Geological cross-section of the Westasturian-Leonese zone along the Cantabrian coast with histograms showing the frequency distribution ofinterlimb angles (f; top row), aspect ratios (h; middle row) and normalized area (A; bottom row) of minor F1 folds from several sectors. See Figure 1 for location.


horizontal axes (Figure 4) with directions following the

Ibero-Armorican Arc (NNE–SSW along the traverse of the

Cantabrian coast). Their asymmetry is variable and depends

on the location within major F1 folds. The limb length of the

analysed folds usually ranges between 3 and 15 m; this size

does not depend on the location of the folds within major

structures, but on the local properties of the multilayer.

The interlimb angle of minor F1 folds has the lowest

values in the western part of the Westasturian-Leonese zone

(Figure 5). The highest values and dispersion of data appear

in the Navia-Alto Sil unit (Nalon-Cabo Vidio section in

Figure 5). A similar variation pattern is observed for the

aspect ratio of the folds (Figure 5).

The shape of the folded surfaces, measured by the

normalized area, mainly ranges from parabolas to chevron

(Figures 5 and 6). Typical chevron folds usually appear in

regularly stratified multilayers with dominance of tabular

competent layers with a thickness lower than 30 cm.

Sometimes chevron and rounded folds coexist in the same

structure, with a transition from chevron shape in the fold

core to rounded folds in the external part of the structure.

Most of the folds developed in competent layers are class


1C, sometimes alongside class 2, or a combination of classes

1C and 3 in the same layer (Figure 7). All folds with a class

1A part and the greater part of folds with s2< 0.5 are located

outside the Mondonedo nappe unit. Class 3 folds are the

most common in the incompetent layers.

5. GEOMETRY OF FOLDS (F2) FROM DUCTILE

SHEAR ZONES DEVELOPED DURING D2

In the Westasturian-Leonese zone, the most obvious results

of the second deformation phase are the Mondonedo basal

thrust and its associated structures. This thrust crops out in

the central part of the zone with a trend that follows the

Asturian arc and separates the Navia-Alto Sil unit from the

Mondonedo nappe unit (Figure 1). However, the thrust is

folded by the wide gentle F3 synform that affects the latter

unit and appears again at the western limb of this fold, that is,

at the north-western corner of the Westasturian-Leonese

zone (Martınez-Catalan, 1985) (Figures 1, 4 and 5). At the

eastern outcrop, nearer to the foreland, the thrust has an

associated narrow corridor with development of small folds,

Geol. J. 45: 597–622 (2010)

Figure 6. Outcrops with F1 folds. (a) Recumbent chevron fold in Benquerencia beach (Lugo). (b) Chevron folds with boudinage development in the limbs(Burela coast, Lugo). (c) Minor folds in the hinge zone of the Courel recumbent syncline (to the north of Quiroga, Lugo). (d) Fold in the Tapia de Casariegocoastline (Asturias); a bulge can be observed in the inner arc of the competent layer. (e) Chevron anticlines separated by a rounded syncline (Tapia de Casariegocoast, Asturias). (f) Folds in Tapia de Casariego harbour (Asturias). (g) Syncline showing a large dilation space filled by incompetent material in the hinge zone(coast near Banugues beach, Asturias). (h) Tight anticline cut by several vertical faults (coast near Banugues beach, Asturias). (a) and (b) Lower Cambriansandstones and slates (Westasturian-Leonese zone); (c) Lower Ordovician sandstones and slates (Ollo de Sapo domain); (f) Lower Ordovician sandstones andslates (Westasturian-Leonese zone); (g) and (h), Devonian carbonate rocks (Cantabrian zone). (a), (b), (d), (g) and (h), east is to the left; (c), south is to the left;

(e) and (f), west to the left.


commonly with curved hinges and crenulation cleavage

(Marcos, 1971, 1973). In the north-western part of the

Westasturian-Leonese zone, where there is a higher

metamorphic grade, the thrust has an associated shear zone

with kilometric thickness and intense ductile deformation,

which involves development of folds, schistosity and


mylonites (Bastida and Pulgar, 1978; Pulgar, 1980; Bastida

et al., 1986; Aller and Bastida, 1993).

A domainal cleavage with relict microfolds and very

small scarce associated folds has been described as formed

during D2 in the Galicia-Tras-os-Montes Schist domain

(Marquınez, 1984; Farias, 1990; Farias and Marcos, 2004).

Geol. J. 45: 597–622 (2010)

Figure 7. Geometry of competent F1 folded layers determined by the parameters s1 and s2 (Bastida et al., 2005) for F1 folds of the Westasturian-Leonese zoneand the Cabo Penas sector (Cantabrian zone). Note that folds from the more internal zones, represented by white symbols and bounded by a discontinuous line,

are more concentrated than the rest of the folds.


Nonetheless, in this domain D2 structures have a regional

development and are not concentrated in well-defined shear

zones.

F2 folds are very abundant in the shear zone associated

with the Mondonedo basal thrust that crops out in the north-

western part of the Westasturian-Leonese zone (Figure 8).

They are minor recumbent folds that are mainly concen-

trated in a few discrete areas, tending to be found in the

normal limbs of major F1 folds, probably due to the

overturned limbs being situated in a stretching position. As

has been described in other shear zones (e.g. Carreras et al.,

2005), progressive deformation in this shear zone gave rise

to superposition between F2 folds, resulting in local type 3

interference patterns (Ramsay, 1967). F2 folds usually

deform the stretching lineation, as observed in sheath folds

in other shear zones (Alsop and Carreras, 2007). They

predominantly face towards the foreland and commonly

occur as strongly asymmetric anticline-syncline couples

with a Z asymmetry (looking to the north). They are not

associated with major folds, and the length of the short

limbs, measured between adjacent hinges, ranges in most

cases between 20 and 50 cm. Fold hinges plunge gently to

the SE to SW, with dispersion of plunge directions

increasing towards the base of the shear zone (Figure 9).

Curved hinges are common and result in conical or sheath

folds and in eye-shaped structures (Figure 8c, e and f). The

increase of the hinge curvature towards the base of the shear

zone explains the increase of the dispersion towards this base


in the measured hinge directions. F2 folds are very tight to

isoclinal (Figure 9), and the most common aspect ratio of the

short limbs ranges from 5 to 10, with the highest values in

the basal part of the shear zone. There is a great diversity in

the shape of the fold surface profile, from elliptical to

chevron folds. The layer geometry corresponds to sub-

similar folds in the basal part of the zone and to class 1C

folds in upper parts.

6. GEOMETRY OF F3 FOLDS

F3 folds are in general upright or steeply inclined and their

axial directions are sub-horizontal and nearly homoaxial

with F1 folds. Nonetheless, these structures have very

different characteristics depending on the geological unit

considered.

In the Galicia-Tras-os-Montes Schist domain, major F3

folds form the most obvious structure on the geological map

and cross-sections (Figure 10). They have a variable

geometry but are mostly gentle to open folds. Below the

allochthonous complexes of Galicia-Tras-os-Montes, they

are rounded folds, often laterally bounded by angular ones,

and major mullion structures have been described (Alonso

and Rodrıguez-Fernandez, 1981). Minor F3 folds are

common in this domain (Marquınez, 1984; Farias, 1990).

They are small folds (limb length between adjacent hinges

less than 1 m). The interlimb angle is very variable (mostly

Geol. J. 45: 597–622 (2010)

Figure 8. F2 folds located in the basal shear zone of the Mondonedo nappe unit and developed in Lower Cambrian metasandstones and micaschists (Cantabriancoast between Foz and Burela, Lugo). (a), (b) and (d), Strongly asymmetric folds. (c) Eye-like structure and other small folds. (e) Folds with curved hinges. (f)

Sheath folds.


between 20 and 808), the most frequent aspect ratio ranges

from 0.5 to 2.5 and the shape of the folded surface profiles is

parabolic or sinusoidal. The dominant geometry of the

competent folded layers corresponds to class 1C with a

tendency in some cases to class 2. Tectonic foliation

associated with F3 folds is very common in this domain and

mainly consists of crenulation cleavage or schistosity.

In the Ollo de Sapo domain, major F3 folds are very well

developed and have a great diversity of morphologies. Axial

planes generally dip to the west, but in some cases to the

east. Vergence changes give rise to a fan-like pattern in the

folds developed below the Cabo Ortegal Complex

(Figure 11). Major F3 folds are gentle to closed and in

exceptional cases have an overturned limb. The interlimb

angle of the majority of minor F3 folds ranges from 40 to 908and in the competent layers these are class 1C folds or folds

formed by a combination of classes 3 and 1C in the same

layer (Bastida et al., 1993). Crenulation cleavage S3 is well

developed.


In the Mondonedo nappe unit, only two major upright

folds, an antiform to the west and a synform to the east, exist

in the northern part (Figures 4 and 5). They are open folds

that become gentle ones towards the south. Minor F3 folds

and crenulation cleavage S3 are scarce in this unit.

In the Navia-Alto Sil unit two different parts can be

distinguished. In the western part, major F3 folds are

common (Figure 12a). They are better developed in the

normal limbs of the major F1 folds and are asymmetric, with

a subhorizontal or slightly East-dipping shorter limb and

another subvertical or strongly West-dipping limb (Pulgar,

1980) (Figure 13). The overturned limbs of the major F1

folds exhibit no or only very gentle F3 folding. Super-

imposition of F3 on F1 folds gave rise to type 3 interference

patterns with a hook shape (Figures 12b, 13 and 14).

Minor F3 folds of several sizes are common in the western

part of the Navia-Alto Sil unit (Figure 12). They do not have

a uniform distribution, but appear primarily in the gentler

dipping shorter limb of the major F3 folds (Figure 13). An

Geol. J. 45: 597–622 (2010)

Figure 9. Geological cross-section of the Mondonedo nappe basal shear zone, showing orientation of F2 folds, histograms of interlimb angle frequency, numberof folds and apical directions of conical folds along the section (direction bisecting the apical angle; Alsop and Carreras, 2007). Major folds shown in the section

are all F1 folds. Contours: 1, 2, 4 and 8% per 1% of area. After Aller and Bastida (1993). See Figure 1 for location.


analysis of these structures was made by Pulgar (1980). Near

Luarca, F3 folds and associated crenulation cleavage are

very well developed. These folds show a great variation in

the interlimb angle values, depending on the lithology,

although the more common values range between 100 and

1208. The more frequent aspect ratio is about 2 and folded

Figure 10. Cross-section showing the D3 structure through the Galicia-Tras-os-MoMartınez Catalan e


surface profiles are mainly parabolic and sinusoidal.

Nonetheless, when pelitic rocks are dominant, kink-bands

or small chevron folds are common (Figure 12e). Class 1C is

the dominant geometry of the folded layers. In general, a

very well-developed crenulation cleavage is associated with

these folds, with common development of tectonic banding

ntes Schist domain and the Ordes Complex (see Figure 1 for location) (Aftert al., 1996).

Geol. J. 45: 597–622 (2010)

Figure 11. Geological cross-section through the Ollo de Sapo domain in the area located to the south of the Cabo Ortegal Complex (see Figure 1 for location).The axial traces of F3 folds form a fan-shape below this complex. After Aller and Bastida (1996).


(Pulgar, 1981) (Figure 12g). In the eastern part of the Navia-

Alto Sil unit, the dominant folds were formed during D1 and

F3 folds are almost absent.

Some D3 structures appear in the north-western part of the

Cabo Penas sector, where an F3 antiform was superimposed

on an F1 anticline giving rise to a type 3 interference pattern.

In the shales that outcrop in the adjacent F3 synform, an

associated local sub-vertical crenulation cleavage S3

appears.

7. FOLDS IN THE CANTABRIAN ZONE

With the exception of the folds located in the Cabo Penas

sector, which represents a transitional area to the Westas-

turian-Leonese zone, the folds of the Cantabrian zone are

difficult to correlate with those in the rest of NW Spain for

several reasons. The structural style of this zone is different

because its dominant structures are thrusts and nappes and

most of the folds developed within it are related to these

structures. Most of the folds from this zone have no tectonic

foliation, and this makes it difficult to distinguish

generations of folds according to geometrical superposition

criteria. A continuous progressive deformation has been


proposed to explain the structural evolution of the

Cantabrian zone, without deformation phases (Alonso,

1987; Perez-Estaun et al., 1988).

Two fold systems can be distinguished in this zone, which,

according to their relationship to the Asturian arc, have been

named arched (or longitudinal) system and radial system

(Julivert, 1971; Julivert and Marcos, 1973) (Figure 15). The

folds of the former system are generally parallel to the trend

of the thrusts, whereas the folds of the latter system are

transversal to them. Many folds of the two systems have

been interpreted as fault-bend folds; longitudinal folds are

mainly related to frontal structures of thrusts, whereas radial

folds are mainly related to lateral ramps (Perez-Estaun et al.

1988). Nevertheless, there are also longitudinal folds that

were formed to balance totally or partially the displacement

of thrusts (Alonso, 1987). Among them, fault-propagation

folds are common in the north-western sector of the

Cantabrian zone (Alonso et al., 1991; Alonso and Marcos,

1992; Bulnes and Marcos, 2001; Bulnes and Aller, 2002),

and folds detached on footwall flats or folds associated with

hanging-wall ramps have been described in the southern part

of the Cantabrian zone (Alonso, 1987).

A N-S directed late-Variscan shortening gave rise to

amplification of previous thrust related transversal folds and

Geol. J. 45: 597–622 (2010)

Figure 12. F3 folds (a–g) and later kink bands (h). (a) Antiform developed in Lower Ordovician quartzites (El Portizuelo, Luarca; Asturias). (b) Interferencepattern F1–F3 giving a hook-like profile (Lower Ordovician sandstones near Garganta pass, Asturias). (c) Minor folds developed in Lower Ordoviciansandstones and slates (north-eastern part of the Ollo de Sapo domain). (d) and (f) minor folds in Lower Ordovician sandstones, siltstone and slates (El Portizuelo,Luarca; Asturias); (e) Small chevron folds developed in Middle Ordovician slates and siltstones (Cantabrian coast near Otur, Asturias); (g) Folds in a quartz veinand associated tectonic banding parallel to the axial planes in Middle Ordovician slates; the white bands are controlled by the hinge zones and the short limbs of

folds (Cantabrian coast, to the west of Luarca, Asturias). (h) Sub-horizontal kink-bands (Luarca beach, Asturias).

Copyright # 2010 John Wiley & Sons, Ltd. Geol. J. 45: 597–622 (2010)


Figure 13. Geological cross-section along the Cantabrian coast near Luarca (Navia-Alto Sil unit; see Figure 1 for location). A Ramsay type 3 interferencepattern originated by the superimposition of F1 and F3 folds is observed. After Pulgar (1980).


to new E-W trending folds which are more important

towards the core of the Asturian arc. An additional N-S

shortening, due to the Alpine deformation, led to the

tightening and rotation of folds, mainly in the southern part

of the Cantabrian zone (Marın et al., 1995; Alonso et al.,

1996; Marın, 1997; Pulgar et al., 1999).

Figure 14. Geological cross-sections made in several localities through the westerinterference patterns resulting from superimposition of F1 and F3 folds can be o(Middle Cambrian-Lower Ordovician); 3, Luarca slates (Middle Ordovician) (a, q

slates. After Pulg


8. KINEMATIC ANALYSIS OF F1 FOLDS

A first approach to the determination of the kinematic

folding mechanisms can be made from the analysis of the

minor structures associated with the folds. The presence of

convergent cleavage is common in F1 folds formed in

n part of the Navia-Alto Sil unit (see Figure 1 for location). Ramsay type 3bserved. 1, Vegadeo limestone (Lower-Middle Cambrian), 2. Cabos Seriesuartzite); 4, Agueira Formation (turbidites) (Upper Ordovician); 5 Silurianar (1980).

Geol. J. 45: 597–622 (2010)

Figure 15. Structural sketch of the Cantabrian zone showing the traces of folds of the arched and radial systems (after Julivert and Marcos, 1973).


competent layers and is indicative of layer shortening at the

beginning of folding.

Several structures indicative of tangential longitudinal

strain occasionally appear in folded competent layers. They

are:

- W

C

edge-shaped tension gashes opened towards the outer

arc in the hinge zone and filled with quartz or calcite; they

involve tangential extension with volume increase

(Bobillo-Ares et al., 2006; Lisle et al., 2009). Occasion-

ally, these structures are located near the hinge zone, but

outside of it, and they suggest a hinge migration later than

the formation of gashes.

- C
leavage in the inner part of the hinge zone; this indicates
tangential shortening with volume loss.

- B
ulging in the inner part of the hinge zone; this is a result
of tangential shortening and can occur without area

change on the fold profile (Figure 6d).

All these structures are common in the eastern part of the

Westasturian-Leonese zone, but they are rare in the western

part.

opyright # 2010 John Wiley & Sons, Ltd.

There is very little evidence of flexural flow and, when it

occurs, it is mainly quartz veins deformed in accordance

with simple shear parallel to bedding. Quartz fibres, or

calcite fibres in the case of the eastern part of Cabo Penas,

are common on bedding surfaces in the limbs of minor F1

folds; they are indicative of a flexural slip mechanism. The

angle between the fibres and the corresponding fold axis for

several fold limbs is shown in Figure 16. These angles are in

general high (about 68% of data corresponds to angles

higher than 608); the deviation from perpendicularity

between axis and fibres can be interpreted as being due to

oblique flexural slip (Ramsay, 1967, p. 396). The presence of

fibres in the hinge zone in a few cases suggests a late hinge

migration.

Geometry of the folded layers provides some ideas about

the folding kinematics. The fact that most F1 minor folds are

class 1C suggests that these folds have undergone super-

imposition of a homogeneous strain. Nevertheless, folds

formed by homogeneous strain superposed on parallel folds

are represented in the s1-s2 diagram by points on the

diagonal through the origin of the square that limits the field

of class 1C folds (Figure 7). Several reasons can explain why

Geol. J. 45: 597–622 (2010)

Figure 16. Histogram showing the frequency distribution of the angle abetween the F1 fold axis and the direction of mineral fibres in the limbs fromseveral F1 folds of the Westasturian-Leonese zone and the Cabo Penassector (Cantabrian zone). N, number of data; a, average value of a; S,

standard deviation.


the points located in this field are not on the diagonal. Firstly,

the theoretical model of flattened parallel folds is obviously

an idealized one, because neither is the fold previous to the

flattening perfectly parallel, nor is the superimposed strain

completely homogeneous. Class 1C folds with s1< s2 can be

formed by flattening that is simultaneous with buckling; this

process has been described by Hudleston (1973b). It is

usually admitted that superimposition of a homogeneous

strain probably occurs when the growth of parallel folds by

buckling becomes difficult (Ramsay, 1967, pp. 411–412).

However, a transitional stage between buckling and

homogeneous strain in which both mechanisms operate

simultaneously could exist in some cases. Water-rock

interaction processes, such as pressure solution, may operate

preferentially on the competent–incompetent interface in

parts of high dips of the fold limbs. These processes could

produce thinning in the limbs and lead to a switch to a class

1C-3, where 0< s1< 1< s2 (Aller et al., 2008). As regards

the class 1C folds, represented by points with 1> s1> s2,

Toimil (2005) has concluded, through numerical modelling

of F1 folds, that they have been formed with a very important

contribution of tangential longitudinal strain.

Analysis of some cases of folded cross-bedding allows

some conclusions to be drawn about the kinematic folding

mechanisms. Bastida (1980), analysing two cases from the

westernmost part of the Westasturian-Leonese zone (Burela

sector in Figure 5), concluded that these F1 folds involved a

combination of tangential longitudinal strain plus an

oblique, slightly heterogeneous flattening. The angle

between the maximum elongation direction and the fold

axial trace could be, at most, 308. Modelling with

‘FoldModeler’ of a fold located in the Navia-Alto Sil unit

(Nalon-Cabo Vidio sector in Figure 5) (Figure 17) indicated

that the mechanism involved was tangential longitudinal

strain followed by flexural flow for layer A (green curve of

Figure 17), and flexural flow for layer B. The initial obliquity


angle d0 of the cross-beds varied between 1548 and 1608(sign convention in Figure 17b).

Kinematics of 20 competent folded layers from different

structures of the Cantabrian coast, analysed previously by

Toimil (2005) using the software ‘FoldModeler’ (chevron

folds were excluded from this analysis), has been re-

analysed using the new version of this software. As a result,

we conclude that homogeneous layer shortening at the initial

stages of folding is indispensable for explaining the

convergent character of cleavage. Tangential longitudinal

strain is also an essential mechanism. In most cases, similar

fits can be obtained when this mechanism operates with or

without area change (the differences are explained below).

Flexural flow is usually not necessary to numerically obtain

good fits of natural F1 folds and, when necessary, it operates

in lower proportion than tangential longitudinal strain. In

addition, flexural flow must usually operate after tangential

longitudinal strain. Superimposition of homogeneous strain

is usually necessary at the final stage of folding. At the

beginning of this stage the homogeneous strain may occur

simultaneously with tangential longitudinal strain or flexural

flow, but finally it must always operate as a single

mechanism. In the analysed folds, the R value (R¼ strain

ratio) for the homogeneous layer-shortening episode ranges

from 1.21 to 5.17. The aspect ratio increment of the

tangential longitudinal strain step, with or without area

change, ranges between 0.2 and 1.5, and the flexural flow

increment between 0 and 1.12. The R value for the final

homogeneous strain increment ranges between 1.04 and

2.78. We have detected some differences depending on

whether the fits were obtained using tangential longitudinal

strain with or without area change. The difference lies in the

amount of deformation in the inner and the outer arc of the

hinge zone. If we consider tangential longitudinal strain with

area change, there is a reduction in the value of R in the inner

arc of up to 4. The value of R in the outer arc increases only

slightly, generally less than 1.

A general study of the kinematics of chevron folds has

been made by Bastida et al. (2007), who have found for these

a sequence of mechanisms similar to that described above

for rounded folds, albeit with some differences. The first

stage of layer shortening is limited by the fact that chevron

folds are usually formed in very competent layers. A second

stage of equiareal tangential longitudinal strain operates

when fold curvature is small and the strain level is low. Later,

longitudinal strain with area change is necessary to avoid

high strain levels in the hinge zone. A later stage involves a

very small amount of flexural flow to avoid strain

concentration and excessive area change in the hinge zone

when the fold acquires an angular shape. Finally, super-

position of a homogeneous strain may occur; a thickness

ratio (orthogonal thickness in the limb/orthogonal thickness

in the hinge)< 1 for the F1 chevron folds of the

Geol. J. 45: 597–622 (2010)

(a) (b)

(c) (d)

Figure 17. (a) F1 anticline developed in Cambrian sandstone and shale of the Westasturian-Leonese zone near Cudillero (Asturias, Spain) showing a foldedpattern of cross-bedding. (b) Sketch of the natural folded oblique surfaces (in black) with the theoretical fold obtained with ‘FoldModeler’ superposed (layerboundaries and grid in green and oblique surfaces in red); the sign convention is shown. (c) and (d) Variation of the obliquity angle d along the folded layer forthe top of layers A and B respectively. Points correspond to natural data; the green curve is the best fit obtained for layer A; red curves correspond to a mechanismof pure tangential longitudinal strain without area change; blue curves correspond to a pure flexural flow mechanism with initial obliquity angle d0 of 1608 and

1548. The latter has been constructed to fit the points with a between 60 and 808 in Figure 17d. After Bobillo-Ares et al. (2009).


Westasturian-Leonese zone indicates that this last stage is

important in these folds. Hence, the development of chevron

folds presents more kinematic restrictions than that of

rounded folds. As the curvature increases in the hinge zone

and the fold acquires an angular shape, tangential

longitudinal strain with area change and a small amount

of flexural flow become necessary mechanisms for the fold

development.

When we try to fit a F1 fold by ‘FoldModeler’ to ascertain

its kinematic folding mechanisms, we encounter a problem

in determining the characteristics of the homogeneous strain

that it is necessary to superimpose at the last stage of folding.

This difficulty is mainly due to three reasons: (a) we do not

know the fold attitude prior to the superimposition of the

homogeneous flattening; (b) we do not know the fold

attitude at the end of the F1 folding, because this position

was modified later during D2 and D3, that is, during the


displacement of rocks on D2 thrusts and during the

development of F3 folds, which led F1 folds to their present

position, and (c) there are an infinite number of different

homogeneous strains that may be superimposed on a fold to

obtain the same final fold shape, so that the final folds will

only differ in a rigid body rotation. We will discuss now

which of these infinite possibilities seem the most likely

from a geological point of view, in order to reach some

conclusions about the attitude of F1 folds before and just

after the superimposition of the homogeneous strain.

The first question to consider is whether the homogeneous

strain was rotational or irrotational. The fold asymmetry and

the systematic fold vergence towards the foreland suggest

that this strain was rotational. In addition, in the context of

an orogenic wedge, where the deformation is driven by

horizontal lateral forces with a probable important gradient

in the vertical direction and where there are important shear

Geol. J. 45: 597–622 (2010)


stresses due to sub-horizontal friction forces, it is difficult to

admit irrotational strain superimposed on folding. In

agreement with these considerations, if we fit minor F1

folds of the Westasturian-Leonese Zone by ‘FoldModeler’

assuming superimposition on folds of a homogeneous

simple shear with a shear direction with plunge of 58 towards

the hinterland, we can obtain a first approximation of the

final dip of the axial surface at the end of the D1 and before

the subsequent deformation phases. The dips obtained in this

case range between 35 and 488 and the shear strain values

involved have a notable dispersion, but they are less than

one. Removing the rotation of the axial surface involved in

the simple shear we find that the dip of the axial surface prior

to the superimposition of the simple shear ranges between 50

and 658. Making a comparable fit for the case of

simultaneous superimposition of a simple shear (with the

same shear direction as above) and pure shear (maximum

stretch direction parallel to the shear direction) and

equivalent components (in strain ratio value) of simple

shear and pure shear, we find final dips of the axial plane of

between 17 and 278 and dips prior to the superimposition of

the axial plane of between 28 and 398. Taking into account

the mean interlimb angle of these F1 folds (Figure 5) and the

characteristics of F3 folds superimposed on them (see

below), the actual range of dips of the axial surfaces at the

end of the D1 was probably intermediate between the two

intervals cited above, that is, between 26 and 388, and

normal limbs had a very low dip towards the hinterland.

8.1. The problem of major recumbent F1 folds

Structures that deserve special consideration are the large

recumbent folds that form the Mondonedo nappe unit and

the Courel recumbent syncline. A kinematic study of the

latter was made by Fernandez et al. (2007) and most of their

conclusions are also valid for the Mondonedo nappe unit. A

set of minor F1 folds located in or close to the hinge zone of

the Courel recumbent syncline (Figure 6c) has been

analysed using ‘FoldModeler’. The bulk shortening involved

in the development of these folds ranges between 55 and

75% and the sequence of kinematic mechanisms obtained is

the same as for the other F1 folds analysed; that is:

homogeneous layer shortening (strain ratio R¼ 1.56),

equiareal tangential longitudinal strain (aspect ratio incre-

ment, Dh, between 0.3 and 0.4), tangential longitudinal

strain with area change (Dh between 0.2 and 0.6), flexural

flow (Dh between 0.1 and 0.2) and homogeneous strain (R

between 2 and 2.6). The latter R interval involves lower

values than that obtained for the major fold by the Srivastava

and Shah (2006) method (R� 5), suggesting that the minor

folds were developed in a late event during the evolution of

the major fold. The above method gives an aspect ratio of


about 2.3 for the buckling fold prior to the later

homogeneous strain. Other features that must be taken into

account to explain the evolution of the Courel recumbent

syncline are the coaxial character of the last stages of the fabric

evolution suggested by the quartz c-axis fabrics and the

coincidence between the maximum stretching direction (l1

direction) and the fold axial direction. The latter feature

represents the main difference with the recumbent folds of the

Mondonedo nappe unit, where the maximum stretching

direction is perpendicular to the axial direction of folds, at least

over wide parts of the unit. Another premise is that the

evolution of the recumbent folds requires a significant rotation

of the axial surfaces towards a sub-horizontal attitude. This

involves a non-coaxial deformation regime, at least during

part of the evolution of the folds.

With the above arguments it is possible to outline an

evolution for the F1 recumbent folds (Figure 18). After a

short initial episode of layer shortening, the folds evolved

through tangential longitudinal strain and minor flexural

flow in the context of a general regime with dominant simple

shear deformation that produced the axial plane rotation

(Figure 18a, b, c). The vergence of folds suggests a shear

direction towards the foreland of the belt. The contribution

of a simple shear regime to the formation of major

recumbent folds requires an initial obliquity between the

shear direction and the layers. For this reason, the bedding

should dip moderately towards the foreland and the shear

direction should have a small plunge towards the hinterland

at the beginning of folding. This suggests an uplift of more

internal parts of the belt during the first episodes of the D1

(Figure 18b, c). This agrees with a progression of the D1

deformation from the hinterland to the foreland of the belt,

as has been shown by Perez-Estaun et al. (1991) and

Dallmeyer et al. (1997). Assuming simple shear with a shear

direction plunging 108 and constant arc-length during active

folding, a bulk shear strain of about 2 is necessary to

generate a rounded fold with the aspect ratio predicted by the

Srivastava and Shah (2006) method (h� 2.3).

After active folding, a homogeneous strain was super-

imposed on folds. Assuming no volume change in the

quartzite layers accompanying the strain, a l1 direction

along the fold axis, and the strain ratio deduced above for the

homogeneous strain on the fold profile ðR ¼ffiffiffiffiffiffiffiffiffiffiffiffil2=l3

p� 5Þ,

an area decrease> 42% is required on the fold profile, as

R0 ¼ffiffiffiffiffiffiffiffiffiffiffiffil1=l3

p> 5. Hence, in the Courel recumbent syn-

cline, the superimposed strain can probably be considered as

a simultaneous combination of simple shear and coaxial

strain (maximum shortening direction sub-vertical). The

latter component was very important and probably increased

with the fold development until becoming clearly dominant

(Figure 18d). This component is required to explain the

recumbent attitude of the fold and agrees with the coaxial

deformation suggested by the quartz c-axis fabrics and with

Geol. J. 45: 597–622 (2010)

Figure 18. Synthetic kinematic model for the Courel recumbent syncline (after Fernandez et al., 2007). For explanation, see text.


the stretching lineation following the fold axial direction.

The latter is clearly incompatible with solely simple shear.

F3 folding and rotation placed the fold in the present position

(Figure 18e).

The evolution described above involves a transition from

a non-coaxial deformation with dominant simple shear to a

homogeneous strain with a dominant component of coaxial

strain with area change. This changeover could be due to the

increasing influence of gravity during folding as a con-

sequence of tectonic superimposition. In fact, gravity has

been considered as an important factor for the development

of recumbent folds (Bucher, 1956, 1962; Hudleston, 1977;

Vacas Pena and Martınez Catalan, 2004).

A difference in the deformation regime between the

Courel recumbent syncline and the recumbent folds of the

Mondonedo nappe unit is the stretching direction. This

difference involves a relative major extension along the axial

direction in the former fold that could be related to the

kinematics of the Ibero-Armorican Arc, which could involve

a more pronounced longitudinal extension in the area of the

Courel recumbent syncline. This suggests that the devel-

opment of this arc was already active during the Variscan D1

deformation (Ribeiro et al., 2007).

8.2. Discrepancies in strain measurements on F1 folds

An interesting problem is posed by the strain measurements

in F1 folds. All measurements made from microstructural

analyses (Fry or Rf/f methods) give strain ratio (R) values

much lower than those deduced from numerical modelling

by ‘FoldModeler’ for the corresponding folds. This


difference was detected in minor F1 folds from the

Navia-Alto Sil unit and the Courel recumbent syncline

(Toimil and Fernandez, 2007; Fernandez et al., 2007). The

very low R values obtained from samples collected in

different parts of very tight folds are surprising in the latter

case. They range between 1.2 and 1.7, whereas the values

obtained by ‘FoldModeler’ usually range between 1.8 (outer

arc of the hinge zone) and 6.3 (inner arc). It is equally

surprising that microscopic observation of a sample

collected in the hinge zone of a minor F1 chevron fold

from the Cape Penas sector suggests no or very low strain

(Bastida et al., 2007). This paradox has been observed by

several authors (Borradaile, 1981; Engelder and Marshak,

1985; Narahara and Wiltschko, 1986), and the only known

mechanism that can explain it is grain boundary sliding. This

mechanism has been considered by some authors as typical

in very fine-grained rocks deformed at high temperatures,

such as mylonites (for example, Passchier and Trouw, 2005).

Although direct evidence of grain boundary sliding is

difficult to find in deformed rocks, it is possible that this

mechanism is usual under conditions of low-grade meta-

morphism, associated with other mechanisms, as for example

pressure solution, in which fluid presence can favour sliding

(Borradaile, 1981; Engelder and Marshak, 1985).

9. KINEMATICS OF F2 FOLDS

The development of D2 structures involved a change in the

deformation regime, so that F2 folds are a result of the

ductile deformation within shear zones. However, their

Geol. J. 45: 597–622 (2010)


vergence is also towards the foreland. This change could be

due to instabilities generated by strain incompatibility in the

D1 that, together with a softening produced by an increase of

temperature with depth, permitted the development of shear

zones with intense ductile deformation involving myloni-

tization. The asymmetry in limb shape (adjacent limbs with

different thickness) and in limb length suggest development

in a strongly rotational deformation regime; this agrees with

the existence of sheath folds and the asymmetry of the quartz

c-axes fabrics (Aller and Bastida, 1993). According to these

authors, the general deformation regime that gave rise to the

shear zone probably involved an important simple shear

component with the shear direction towards the foreland and

parallel to the bedding of the normal F1 limbs. The high

anisotropy due to the bedding controlled the shear plane in

these limbs; in fact, detached F2 folds, with the bedding as

decollement level, can be seen within the shear zone.

Development of F2 folds involved a first amplification as a

result of flow instabilities (Lister and Williams, 1983; Platt,

1983; Carreras et al., 2005); it could be active or passive. The

existence of abundant, strongly asymmetric anticline-syncline

couples without periodicity (Figure 8b and d) suggests an

initial mechanical instability that did not necessarily derive

from a buckling mechanism. In the higher part of the shear

zone, folding with several waves and a certain periodicity can

be observed in some outcrops, suggesting that here the

multilayer presented a high competence contrast and the

initial mechanism came closer to buckling. After this initial

stage, a passive amplification of folds occurred. The

deformation superimposed on the folds may have been

heterogeneous, resulting in folds with curved hinges and

sheath folds. Nevertheless, sheath folds can also be developed

by superimposition of homogeneous simple shear (Cobbold

and Quinquis, 1980; Ramsay, 1980). In the present case, it is

very difficult to explain recumbent folds solely by simple

shear with the shear direction forming a low angle to the

layers. It requires enormous amounts of shear strain that

would result in a huge thinning of the overturned limb that is

not usually observed. A combination with pure shear to

generate sub-simple shear facilitates the development of

recumbent folds, although it generates a problem of strain

compatibility at the boundaries of the shear zone that must be

solved by a heterogeneous deformation gradient or by the

development of fractures (Ramsay and Huber, 1987, pp. 612–

613). A way to mitigate or even solve this problem is through

an area decrease associated with the strain (Ramsay, 1980,

Figure 2c). Nevertheless, since the stretching lineation is

almost coincident with the shear direction, the area change

would require a volume decrease that is too large for strain

compatibility. It is possible that the strain varied progressively

upwards towards simple shear to make the compatibility

possible, but downwards the incompatibility gave rise to the

initiation of the basal thrust of the unit.


10. KINEMATIC MECHANISMS OF F3 FOLDS

The development of D3 structures involves an important

change in the deformation conditions. The common upright

position of F3 folds, the sub-vertical attitude of the

crenulation cleavage S3 and the fact that minor D3 structures

are better developed when S0 and S1 are sub-horizontal, the

case in which folds tend to be symmetrical, suggest a

structural regime involving a dominantly coaxial bulk strain

with a sub-horizontal direction of maximum shortening.

Therefore, the deformation style changes from a regime with

a component of simple shear with sub-horizontal shear

direction and development of thrusts (first and second

deformation phases) to a compressive, dominantly coaxial

regime with sub-horizontal maximum shortening (third

deformation phase). The reasons for this change are difficult

to evaluate; it could be controlled by a change in the

dynamics of the plates involved in the orogenic processes or

it could be triggered when the orogenic wedge reached a

critical length that inhibited the ulterior horizontal

translation of rocks in its front section and wedge

compression to readjust its shape became necessary.

Another difference between the first and the third

deformation phases is that, whereas D1 structures are

widespread, with the exception of the Galicia-Tras-os-

Montes domain, the development of D3 structures is more

selective, and these vary in frequency and geometry from

one unit to another. This is a result of D3 structures having

been formed at a late stage of the orogenic belt development.

When the D3 deformation begins, several factors control its

development. The bedding and the S1 foliation are not

horizontal; they dip towards the hinterland and the bedding

dip is higher in the overturned limbs than in the normal limbs

of F1 folds, so that the D3 structures are better developed in

the latter limbs. The previous development of thrusts and

other major faults brought into contact rocks with very

different mechanical properties, determining different

development of D3 structures to both sides of the fractures.

In addition, the structure formed by the thrust systems gave

rise to major irregularities, due, for example, to the stacking

of sheets, which controlled the development of D3

structures. The existence of massive plutonic rocks, which

intruded prior to the development of the third deformation

phase and modified the previous structure, was another

factor that influenced the development of D3 structures. The

intrusion of granitic magmas coeval with the third

deformation phase gave rise to an interesting interrelation-

ship between magmatism and deformation in some areas.

Some of these factors could be the cause of vertical forces

that superimposed on the generalized sub-horizontal

compression that gave rise to D3 structures.

Assuming a sub-horizontal compression during D3, we

can make some considerations about the position of the F1

Geol. J. 45: 597–622 (2010)


folds at the moment previous to the development of D3.

Anthony and Wickham (1978) analysed with finite-element

models the development of asymmetric folds in viscous

inclined single layers. They observed that the fold

amplification decreases as the initial inclination of the

layer increases, so that when the initial dip is 208, the

amplification is very small. In addition, folds with the two

limbs dipping in opposite directions are only produced for

inclinations of less than 158. Although these models are

obviously too simple to be directly applied to natural cases,

this numerical result agrees with the conclusion obtained

above about the very low dip of the normal limbs of F1 folds

at the end of the D1. The lack of D3 folds in the eastern part

of the Navia-Alto Sil unit suggests that the relatively high

dip of the bedding at the beginning of this deformation phase

prevented the development of folds. In the western part of

the same unit only the normal limbs of F1 folds appear

significantly folded by the D3 deformation; this suggests that

the overturned limbs had too high a dip to suffer F3 folding.

In contrast, F1 folds with the two limbs folded by the D3

deformation appear in the Ollo de Sapo domain, suggesting

that the F1 folds involved were near isoclinal and recumbent.

In the Galicia-Tras-os-Montes zone, F3 folds have been

interpreted as a result of buckling plus flattening (Marquı-

nez, 1984; Farias, 1990). The minimum bulk shortening

estimated by these authors varies between 42 and 74% with

gradients mainly controlled by the lithology. The role of the

lithology must be analysed in each case, but the influence of

bodies of two-mica granitoids in the development of major

F3 folds is particularly noteworthy (Figure 10). The age of

most of these bodies ranges between the interphase D2–D3

and late-D3. They operated as buttresses, enhancing the

amplification of D3 folds developed in adjacent rocks.

F3 folds developed in the Ollo de Sapo domain have also

been formed in a similar way to F3 folds of the Galician-

Tras-os-Montes zone, that is, by buckling plus flattening,

with a total bulk shortening of between 43 and 70%,

measured by the Aller et al. (2008) method. The large-scale

disharmonic F3 folding to the south of the Cabo Ortegal

Complex is particularly interesting (Aller and Bastida, 1996)

(Figure 11). It is a result of the buckling of a complex

multilayer formed by three parts: the lower part is formed by

an alternation of competent and incompetent layers consti-

tuted by the Ollo de Sapo porphyroid and siliciclastic rocks,

with a minimum thickness of 1000 m; the middle part is

formed by mainly incompetent rocks, slates with one

porphyroid and one quartzite level in its upper part, with a

total thickness of about 4000 m, and the upper part is formed

by the mainly massive competent mafic and ultramafic rocks

of the Cabo Ortegal Complex, with a minimum thickness of

3500 m. The last of these exerts the main control on buckling

and developed a gentle rounded synform with large span; the

lower part developed closed or tight F3 folds with lower


wavelength and more angular geometry. They were super-

imposed on F1 folds and, as a whole, resulted in a

synclinorium that was accommodated to the shape of the

Cabo Ortegal Complex. The incompetent middle part

accommodated its geometry to the competent parts and

the resulting fold developed a dilation space in the hinge

zone below the Cabo Ortegal Complex. Located in this space

is a rounded granite body late- to post- F3 folds. The magma

probably used this space to intrude and subsequent

ballooning gave the intrusion its final shape (Aller and

Bastida, 1996).

The notable development of D3 structures in the Ollo de

Sapo domain contrasts with their limited development in the

Mondonedo nappe unit. The large span and low amplitude of

the two major folds and the shortage of minor folds can be

explained by the dominantly competent nature of the

Cambrian-Ordovician succession which forms the thick

multilayer that controlled the folding in this unit. The

Viveiro Fault was developed before the third deformation

phase and gave rise to an abrupt contact in several parts of its

trace between the incompetent rocks of the Ollo de Sapo

domain and the more competent rocks of the Mondonedo

nappe unit. Pre- or early-D3 granitoids that intruded in this

unit near the Viveiro Fault in several areas contributed to an

increase in the competence of the latter unit. Hence, the

footwall of the fault probably operated as a near rigid block

that enhanced the compressive stresses that generated the D3

structures in the Ollo de Sapo domain.

Perez-Estaun et al. (1991), on the basis of seismic

refraction data, suggested that the formation of the huge F3

antiform that appears in the western part of the Mondonedo

nappe (Lugo antiform) could have been driven by the

development of an antiformal stack made up of slivers of

crystalline basement. On the other hand, the intrusion of

granitic magma in this unit could have given rise to a vertical

push upward that favoured the development of this antiform.

In fact, the two processes are compatible.

In the western part of the Navia-Alto Sil, the influence of

the lithology and the attitude of the rock anisotropy are very

clear. The asymmetry of the major F3 folds is a result of the

folding of a multilayer gently dipping towards the hinterland

due to sub-horizontal compressive stresses. Development of

minor F3 folds is controlled by the geometry of the major F3

folds, so that the minor folds are concentrated in the less

dipping short limbs of the major folds (Figure 13), that is,

where the compressive stress forms a low angle to the

bedding. This selective development of the minor F3 folds is

indicative of these structures starting to develop after the

major F3 folds. This is the reason why the minor folds appear

concentrated in bands separated by other bands without

minor F3 structures. The width of the bands is variable, from

map scale to microscopic scale, and depends on the

wavelength of the corresponding major folds. An example of

Geol. J. 45: 597–622 (2010)

Figure 19. Geological cross-section and map showing the distribution in bands of D3 structures and later near horizontal kink bands in slates of Luarca beachand the surrounding area (after Pulgar, 1980). See Figure 1 for location.


this arrangement in bands, developed in slate outcrops, is

shown in Figure 19 and a scheme of its progressive

development is illustrated in Figure 20. The bands without

D3 correspond to long limbs of major F3 folds where S1

presents a very steep or sub-vertical position. In these limbs

sub-horizontal kink-bands were developed (Figure 12h). In

some cases, these folds appear affecting the S3 crenulation

cleavage, indicating that they are post-D3. They can be

interpreted as a result of a vertical compressive late-Variscan


stress caused by gravitational forces (Matte, 1969; Pulgar,

1980; Julivert and Soldevila, 1998). A similar pattern has

been described near to the northern boundary of the Truchas

syncline area (Ollo de Sapo domain; Figure 1) (Fernandez,

2001).

Minor F3 folds mainly tended to develop in the lower part

of the Luarca Slates (Lower to Middle Ordovician)

(Figure 12d and f), stratigraphically located on top of the

Cabos Series (Middle Cambrian to Lower Ordovician),

Geol. J. 45: 597–622 (2010)

(a)(b)

(c) (d)

Figure 20. Scheme to explain the evolution and distribution of D3 structures and late near horizontal kink bands in the area of Figure 19. (a) Position of the S1

cleavage before D3; (b) development of F3 folds with a sub-horizontal limb and a steep limb; (c) Development of minor F3 folds and associated S3 cleavage inthe sub-horizontal limbs and rotation of the other limbs towards the vertical direction; (d) development of later near horizontal kink bands (K-B) in the sub-

vertical limbs.


whose upper part is formed by massive competent quartzite.

The lower part of the Luarca Slates is a multilayer formed by

sandstone layers with different amounts of quartz that

alternate with slates. In this multilayer, the folds commonly

exhibit a wavelength that indicates a buckling process.

Kinematic mechanisms in competent layers depend on the

quartz content of the sandstone. When this content is low, a

disjunctive convergent cleavage S3 appears associated with

F3 folds and a very low ratio wavelength/thickness is found

for F3 folds in the competent layers. This indicates an

important layer shortening at the beginning of folding that is

corroborated by the use of the ‘FoldModeler’ program,

which shows a mechanism sequence with significant layer

shortening, tangential longitudinal strain without area

change, tangential longitudinal strain with area change

and flexural flow, with a greater amount of flexural flow than

tangential longitudinal strain mechanisms. The bulk short-

ening for these folds ranges from 48 to 55%. As the content

in quartz of the competent layers increases, the amplification

degree also increases, suggesting less layer shortening and

more flattening in the later stages of folding. The latter

mechanism is clear when the geometry of folded layers is

class 1C. In some cases, it is possible to find in the same

sequence folded layers with different amplitude and

wavelength, indicating different proportions of layer short-

ening and buckle shortening, as a consequence of the

different competence contrast (Ramberg, 1964).

In decimetric F3 folds, especially when these are

developed in quartz veins parallel to S1, it is common to

find mesoscopic tectonic banding parallel to the axial

plane (Pulgar, 1980, 1981) (Figure 12g). This is formed


by white bands, rich in quartz, which cross the hinge zone

in symmetric folds or the short, less dipping limbs in

asymmetric folds and alternate with dark bands, poor in

quartz and rich in phyllosilicates, which cross the limbs

in symmetric folds or the long limbs in asymmetric folds

(Figure 12g). This banding shows the heterogeneous

development of the D3 structures on this minor scale and

is the result of a pressure solution mechanism involving

quartz dissolution in the strongly dipping limbs which face

the compression, and quartz precipitation in the parts of the

folded veins with lower dip. This structure is a mesoscopic

example of the heterogeneous distribution in bands of the D3

structures.

In the cases with dominant pelitic rocks, kink-bands

frequently formed, and when high shortening took place,

small chevron folds were developed (Figure 12e) in the way

described by Paterson and Weiss (1966).

11. CONCLUSIONS

The Iberian Belt in NW Spain is an arcuate sector of the

Variscan Orogen that shows a complete cross-section from

the foreland to the hinterland. With some restrictions,

deformation in the belt can be considered as being the result

of three main phases (D1, D2 and D3).

D1 gave rise to closed or tight folds (F1) with associated

cleavage (S1). These structures are widespread in the

hinterland, mainly in the Westasturian-Leonese zone and in

the Ollo de Sapo Domain. F1 folds have sub-horizontal

hinges, with a trend that follows the Ibero-Armorican Arc,

Geol. J. 45: 597–622 (2010)


and verge to the foreland. They are recumbent folds in the

Ollo de Sapo domain and in the Mondonedo nappe unit,

where they are greater in size and have a lower interlimb

angle than in the Navia-Alto Sil unit. Parabolas to chevron

are the most common shapes for minor F1 folds, and class 1C,

sometimes next to class 2, and combinations of classes 1C and

3, are the most common geometries of the folded layers.

During D2, foreland-directed thrusts and shear zones

formed. The most noteworthy thrust is the Mondonedo basal

thrust, which developed in its most internal part a thick shear

zone with intense ductile deformation giving rise to folds,

schistosity and mylonites. F2 folds in this zone are small,

strongly asymmetric recumbent folds that often appear as

anticline-syncline couples facing towards the foreland. They

are very tight and frequently have curved hinges. They range

in profile from elliptic to chevron shapes and the layer

geometry is sub-similar or class 1C.

F3 folds are in general upright or steeply inclined and their

axial directions are sub-horizontal and nearly homoaxial

with F1 folds. These structures have very different

characteristics depending on the geological unit considered.

They are frequent in the Galicia-Tras-os-Montes Schist

domain, Ollo de Sapo domain and in the western part of the

Navia-Alto Sil unit, but are scarce in the Mondonedo nappe

unit and in the eastern part of the Navia-Alto Sil unit.

D1 and D2 involve deformation with important horizontal

displacements, either as ductile deformation in a regime with

a significant component of simple shear or as rock

translation over thrusts. Both phases are reflected in the

foreland fold and thrust belt (Cantabrian zone) as thin-

skinned tectonics, with thrust nappes and folds related to

them as dominant structures. D3, on the other hand,

represents a deformation due to a sub-horizontal com-

pression that involved a dominant coaxial deformation with

sub-horizontal maximum shortening.

In individual competent layers, the development of F1

folds involved a kinematic sequence of layer shortening,

tangential longitudinal strain, flexural flow and homo-

geneous strain. The amount of layer shortening increases as

the layer competence decreases. Tangential longitudinal

strain without area change occurs in the first stages of

folding with low strain ratios and curvatures in the folded

layer, whereas tangential longitudinal strain with area

change occurs as the curvature increases, this being strictly

necessary in chevron folds. Flexural flow is a late

mechanism of buckling. It takes place usually to a small

degree and is not always necessary. A small component of

flexural flow is necessary in chevron folds. The strain

superimposed in the later stages involved a combination of

simple shear, coaxial strain and area decrease on the fold

profile. The coaxial component is necessary in F1 recumbent

folds. In addition, large recumbent F1 folds, such as the

Mondonedo nappe unit or the Courel recumbent syncline,


required a previous uplift of the more internal areas of the

belt. Interaction between layers during folding involved

flexural slip in many cases; evidence of this mechanism

increases towards the foreland.

From a methodological point of view, the strain

measurement made from microstructural analysis in F1

folds give values of the strain ratio (R) that are much lower

than those deduced from numerical modelling by ‘FoldMo-

deler’. This discrepancy may be due to the operation of grain

boundary sliding as a deformation mechanism, which is not

taken into account in strain measurements. This mechanism

could operate in combination with pressure solution. This

discrepancy reflects the difficulty of applying continuum

mechanics concepts to the analysis of the deformation of

particle aggregates on the microscopic scale.

Increase of the deformation during the D1 probably gave

rise to strain incompatibilities at deep levels, and this,

together with high temperatures, favoured the concentration

of the ductile deformation in shear zones and the

development of mylonites that in turn contributed to the

rock softening. Local flow instabilities generated F2 folds

that were passively amplified by a combination of simple

shear, coaxial strain and some area change. This strain

regime could increment the strain incompatibilities at the

bottom of the shear zone and favor the development of basal

thrusts.

The development of D3 structures depended on the

previous dip of the bedding and the S1 foliation. In the Ollo

de Sapo domain, where the dips of the two limbs of F1 folds

were probably lower than 158, both limbs were amplified

during the D3, and a significant development of F3 folds took

place. In the western part of the Navia-Alto Sil unit, only the

normal limbs of the F1 folds had low dip and only in these

limbs are F3 folds common. In the eastern part of the Navia-

Alto Sil unit, the previous dips of bedding were probably

higher than 208 and D3 structures did not develop; only a

rotation of bedding and S1 took place.

Granitoid bodies prior to D3 played the role of buttresses

that permitted the amplification of F3 folds in some areas, as,

for example, near the eastern boundary of the Ollo de Sapo

domain. On the other hand, the intrusion of syn-D3 granitoids

and the stacking of thrust sheets favoured the development of

major F3 antiforms (the Lugo antiform, for example).

The characteristics of the multilayer control the span and

the amplitude of the F3 folds. Thick multilayers with

dominant competent layers, such as the Mondonedo nappe

unit, allowed the development of large open or gentle F3

folds. The competent massive dense character of the Cabo

Ortegal Complex permitted the development of a major F3

synform with a large dilation space in the hinge zone of the

adjacent incompetent rocks, where granitic magma intruded.

When pelitic incompetent rocks are dominant, chevron folds

or kink bands are the most common F3 folds.

Geol. J. 45: 597–622 (2010)


In the western part of the Navia-Alto Sil unit minor D3

structures appear distributed in bands at several scales. Deve-

lopment of minor F3 folds began after that of the corres-

ponding major F3 folds, and minor folds formed only in the

sub-horizontal short limbs of major folds. In the case of

slates, post-D3 sub-horizontal kink bands developed on the

sub-vertical long limbs as a result of vertical compression.

On a small scale, tectonic banding appears associated with

small folds in quartz veins as a result of pressure-solution

processes. In such cases, bands with higher quartz content

formed parallel to the axial planes; these bands cross-cut the

hinge zones of symmetric folds or the sub-horizontal short

limbs of asymmetric folds.

In general, bulk shortening involved in F3 folds is lower

than that involved in F1 folds, but the sequence of kinematic

mechanisms is similar. The main difference lies in the

character of the strain superimposed on folds, which in the

case of F3 folds is a coaxial flattening, whereas in F1 folds it

is a non-coaxial strain.

The present study is an example of the application of

geometrical analysis of folds to their kinematical study. The

knowledge of the geometry of folded surfaces and folded

layers is necessary to model and simulate folds by computer

using a combination of kinematical folding mechanisms.

Despite the limitations of the theoretical models, since they

are 2D and obtained from idealized mechanisms, the results

obtained from numerically developed folds can in general be

very useful for the understanding of the kinematics of natural

folds of all sizes and the evolution of the orogenic belts.

ACKNOWLEDGEMENTS

The present work was supported by the CGL2008-03786-

BTE project funded by the Spanish Ministerio de Ciencia e

Innovacion and the Fondo Europeo de Desarrollo Regional

(FEDER) and the project ‘Topo-Iberia’ (CSD2006-0041) of

the Spanish CONSOLIDER-INGENIO 2010 Program. The

authors are grateful to Ian Alsop, Richard J. Lisle and

Andres Perez-Estaun for many valuable suggestions that

notably improved the manuscript.

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Documents

Folding in orogens: a case study in the northern Iberian Variscan Belt