Anatomy of the Mondoñedo Nappe basal shear zone

INTRODUCTION

THE Mondoiiedo Nappe basal shear zone is well suited for the study of problems arising in this type of structural association, and has already been the subject of some previous papers (Bastida & Pulgar 1978, Pulgar 1980, Martinez-Catalan 1985, Bastida et at. 1986). The shear zone, 3.4 km wide, is located in the northern part of the Iberian Hercynian belt (Fig. 1), presenting a continuous outcrop with strong ductile deformation all along the Cantabrian coast. It is bound at its base by a major thrust developed in brittle-ductile conditions during the last stage of the shear zone deformation history. The shear zone outcrops again, further to the east, in an area closer to the front of the Mondoiiedo Nappe (Fig. 1). Here the shear zone is characterized by a brittle-ductile defor mation zone not wider than 100 m, above the basal thrust.

GEOLOGICAL SETTING

The shear zone is the lower boundary of a large thrust sheet, the Mondoiiedo Nappe, situated in the western part of the Westasturian-Leonese zone, a major tectonic zone in the hinterland of the Iberian Hercynian belt. The Mondoiiedo Nappe unit is made up of a stack of recum bent, almost isoclinal, large amplitude E-facing D 1 folds ( Fig. 1) with an associated tectonic foliation (Sl)' These folds are affected by the nappe basal thrust and the associated shear zone. This shear zone is important in the western part of the nappe, where a large number of shear related E-facing minor folds (D z ) are developed. In this part, Sl has been flattened (Sl+z), crenulated or replaced by a new mylonitic foliation (Sz). Finally, the whole nappe has been folded into a gentle synform in the east and an antiform in the west (D 3 ), both with a subvertical axial plane and a large wavelength. D 3 minor folds and S., crenulation cleaval!e occur locallv. D., folds homoaxially overprint D 1 folds producing a Type 3 interference pattern (Ramsay 1967) that can be seen both on the map and in cross-section ( Fig. 1).

Uplift due to the D 3 major antiform has resulted in exposure of the thrust and the basal shear zone around the Xistral semi-window ( Fig. 1) where amphibolite facies metamorphism was reached (Martinez-Catalan 1979. It is in this western outcrop of the shear zone, in the section between the coastal localities of Burela and Foz (Fig. 1), that most of this work was focused.

GEOLOGY OF THE CROSS-SECTION

The rocks affected by the shear zone as well as those of the Xistral semi-window, belong mainly to the Cindana Group (Lower Cambrian) (Walter 1966(Walter , 1968). Precam brian rocks, consisting of micaschist with common sand stone intercalations and porphyroids outcrop just at the base of the nappe. The Candana Group consists of four units, only two of which are affected by the shear zone: the lower Candana formed by coarse-grained felds pathic sandstones and quartzite with intercalated schists, and the middle Candana formed by micaschist, commonly with some sandstone, and minor black slate, carbonatic and amphibolitic layers.

The macroscopic structure shown in Fig. 2 consists of several very tight D 1 folds, downward-facing due to their location in the western limb of a large gentle D 3 syn form. The basal thrust, also in a tilted position, can be seen in the western part of the limb. The ductile defor mation of the shear zone is shown by an association of minor structures that developed in the nappe footwall (shaded area in profiles of Fig. 1).

The general features of metamorphism and its chronological relationship with the deformation in the Mondoiiedo Nappe have been pointed by out Bastida & Pullmr (1978). Martinez-Catalan (19R5) and Ba!'\tida et Bastida et ai. (1986). BT, basal thrust.

al. (1986). The metamorphic zones that occur on the staurolite occurring locally as residuals inside andalusite coastal region of the Mondofiedo Nappe unit are shown porphyroblasts. in Fig. 1. The shear zone sector is situated within the Three metamorphic episodes can be differentiated in andalusite and almandine zones (Fig. 2). The andalusite the area of Fig. 2 (Bastida et al. 1986): M}> with a zone oartlv overorints the staurolite zone with relict oaral!enesis containinl! staurolite and I!amet (medium pressure); M z , with andalusite (rise in temperature and perhaps a pressure drop); and M 3, presenting retrograde metamorphism to greenschists facies. M I is pre-D z . The rise in temperature during M z is coeval with ductile deformation in the shear zone (M z early syn-D z ). M 3 is late-to post-D z (end of ductile deformation in the shear zone and development of the basal thrust) and is related to the drop in pressure and temperature during and after uplift of the nappe rocks.

STRUCTURES IN THE BASAL SHEAR ZONE

The strong ductile deformation associated with the shear zone gave rise to a large number of small-scale D z structures, and pre-existing structures (D I ) were modi fied. D z structures consist of folds, foliations and linea tions as well as microfabrics, which will be discussed separately.

Folds

Folds are very abundant in the western part of the section, where they are mainly concentrated in four narrow discrete zones (Fig. 2): Punta das Cabras, Punta Riomar and two others with very similar characteristics, that will be considered together (Punta Morago). All three zones are situated along the normal limbs of major D I folds. This position is likely to have favoured the development of D z folds, since the reverse limbs were likely situated in a stretching position.

The lithology of the folded zones consists of quartzite, feldspar sandstone and micaschist. The quartzite and the sandstone occur in thin layers. In Punta das Cabras and Morago these rocks are more abundant than the schists, whereas in Punta Riomar the amount of quartzite and sandstone is often equivalent to the amount of schist. These lithologies may have favoured the development of D z folds, contrasting with the massive quartzite that occurs in the lower Candana, and the Precambrian or middle Candana schistose zones.

The folds are E-facing and commonly occur as strongly asymmetric anticline-syncline couples. Small fold trains also occur locally. Folds show varying limb thickness. The D z folds are not associated with major folds and preserve the same asymmetry regardless of their location within the D 1 folds. The length of the short limbs ranges in most cases between 20 and 50 em, measured in the fold profile along the layer between two adjacent hinges.

Fold hinges are SE-to SW-plunging ( Fig. 2), with a increased dispersion of plunge directions towards the Hudleston's (1973) parameters (below) in the three main D z fold localities of the shear zone.

base of the shear zone. Curved hinges are common and Several mesoscopic Type 3 fold interference patterns give rise to conical or sheath folds and to eye-shaped (Ramsay 1967), pre-dating D 3 , have been recorded in structures. The interlimb angle distribution shows that Punta Riomar. At this locality, two superimposed gen D z folds are very tight to isoclinal (Fig. 2). The fold erations of microfolds also deform the previous 51 folia geometry has been analysed using Hudleston's (1973) tion. This suggests that progressive superimposed classification (Fig. 3). Amplitudes show a higher mode folding took place during the evolution of the shear in Punta das Cabras and Punta Morago than in Punta zone. Riomar. There is a great diversity in fold shape. The analysis of fold layers has also been carried out using Folding mechanisms Ramsay's method (1967), modified by Bastida (1993) (Fig. 3). This analysis shows that in Punta das Cabras the The geometry of the folds, together with the type of large majority of folds are subsimilar, without geometri deformation that affected the shear zone suggest that a cal differences between adjacent folded layers of differ significant part of the fold evolution resulted from pass ent lithology. Class lC folds are most common in the ive amplification of pre-existing folds or irregularities in other two localities, with alternation of class 3 and lC non-coaxial flow. Evidence for a passive mechanism folds that imply a difference in mechanical behaviour operating during at least the later stages of D z folding is between layers.

provided by the broad similarity between quartz c-axis orientation patterns in limbs and hinges of folded quartzite layers. An example is shown in Fig. 4. The presence of sheath folds provides further evidence of non-coaxial progressive deformation (Quinquis et at. 1978, Cobbold & Quinquis 1980, Ramsay 1980, Skjer naa 1989. The initiation of D 2 folding would have been enhanced by the orientation of So in the shortening field of deformation, due to obliquity between So and the shear plane (Ramsay 1980, Martinez-Catalan 1985.

The origin of pre-existing irregularities necessary for fold amplification in shear zones has been a matter of discussion (Berthe & Brun 1980, Cobbold & Quinquis 1980, Lister & Williams 1983, Platt 1983, Ghosh & Sengupta 1984. The lack of periodicity in the case of anticline-syncline couples (mainly present in Punta das Cabras) suggests an initial mechanical instability which does not necessarily derive from a buckling mechanism. Rather, it could derive from an instability in the flow, producing local increases in the shear deformation rate (Lister & Williams 1983, Platt 1983. If buckling had taken place it would have stopped operating at a very early stage since its evolution, always implying propaga tion by wave lateral aggregation, should have involved a periodicity (Biot et at. 1961, Biot 1965, Cobbold 1975.

In those cases where there is a periodicity and where the behaviour of competent and incompetent rocks is differ ent (mainly Punta Riomar), buckling seems to be the fold initiation mechanism. However, the geometry of the folded layers suggests a remarkable kinematic ampli fication. Therefore, it seems unlikely for buckling to have given rise, in this zone, to very tight or isoclinal folds as described by Ghosh & Sengupta (1984) in their study of folds of the Kolar Gold Field (India). These observations suggest slightly different folding evolution patterns for the different localities of the shear zone. Passive evolution seems to have played an important role in Punta das Cabras, while in Punta Riomar.buck ling must have been more important. Punta Morago appears to be intermediate between the two. The differ ences in evolution may be conditioned by lithological differences and, perhaps, by small differences in press ure and temperature, which are higher towards the base of the nappe and decrease as the deformation continues.

As regards fold hinge orientations, the asymmetric character of most projections (Fig. 2) with respect to the direction of shear suggests a slight initial obliquity be tween So and the shear plane (Martinez-Catalan 1985). The gradual approach of the projection maxima in sectors IV, III and I (Fig. 2) to the shear direction and the increase in dispersion suggest stronger deformation of the plane containing the hinges toward the basal part of the shear zone (Sanderson 1973, Williams 1978.

Lineations

A well developed mineral lineation (L m ) commonly occurs on the mylonitic foliation along most of the shear zone, mainly in its lower part (Fig. 5). The lineation is defined by the dimensional preferred orientation of micas, especially biotite, and quartz aggregates. The lineation plunges gently towards the east, and is usually folded by later D 2 structures. This folding is especially evident in Punta Riomar. In the upper part of the shear zone, the most frequent lineation is an intersection lineation between the regional foliation S2 and So, which is parallel to D 2 fold hinges.

Quartzite microstructures

With quartz microstructures it is possible to dis tinguish several areas of mylonitic development along the shear zone. There is also a marked change in micro structure development upwards, away from the basal thrust.

In quartzitic layers near to the basal thrust (Fig. 6) the foliation is defined by small grains (Fig. 7a) with straight grain boundaries and a strong shape preferred orien tation. Intracrystalline deformation microstructures are scarce. These features seem to indicate a total and fast dynamic recrystallization resulting in an overall de crease in grain size. The microstructure of these rocks is that of an ultramylonite.

A mylonitic foliation is also developed in Punta Mor ago. Here, small equidimensional recrystallized grains constitu~e a dominant matrix in which elongate old grains with a strong preferred shape orientation define the foliation. To the west of Punta Morago, the tran sition to a mylonitic foliation can be seen as weakly oriented old grains become strongly oriented. Quartz ribbons are the cores of the old grains that remain unrecrystallized.

Above the ultramylonites of the basal part, intracrys talline deformation features such as deformation bands and subgrains are well developed. On the whole, the grain size is larger than in the more basal quartzite and the foliation is less well developed (Fig. 7b). The larger grain size together with the less well developed pre ferred grain shape orientation seem to indicate a lower stress, at least during the final stages of microstructural development, than in quartzite near the basal thrust. Further to the east (higher in the shear zone), quartz grains show a weak to absent preferred grain shape orientation. The main foliation is defined by a small amount of long mica crystals with a grain shape orien tation (Fig. 7c). Usually, these samples are predomi nantly coarse grained with irregular grain boundaries, but not strongly lobate or serrated. Grain size decreases towards the eastern part of the section. Subgrains and deformation bands are abundant and a minority of small recrystallized grains with very few deformation micro structures also occurs. The quartzite commonly exhibits a weak foliation oblique to the main one, which formed by local dynamic recrystallization (Fig. 7c). Similar microstructures, although finer-grained, have been pre sented by Lister & Snoke (1984, figs. 4e and 10d), who interpreted them as recrystallized type II S-C mylonites. Well-oriented biotite ribbons (long and thin mica fish in Lister & Snoke 1984) suggest intense deformation. All these rocks on top of the basal ultramylonites can be considered as blastomylonites, in the sense defined by Sibson (1977).

Non-mylonitic tectonites occur predominantly in the eastern half, and also above the shear zone (Fig. 6). The boundary with the preceding group is at the andalusite isograd. Non-mylonitic tectonites also appear in the footwall. The quartz preferred grain shape orientation is Fig. 6. Quartz c-axis fabrics (lower-hemisphere, equal-area projections; contours: 0.5, 1,2,4 and 8%), showing locality and fabric type (foliation is the vertical E-W plane and L m is horizontal), isograds, quartzite tectonite types and schist foliation types along the shear zone. Abbreviations as in Fig. 2. ,... ,... ,... weak and crystal-plastic deformation features in the grains vary, though in most cases deformation bands and undulose extinction are common (Fig. 7d). These rocks differ from the blastomylonitic rocks in that they do not show the microstructural features of S-C mylonites.

Main foliation (Sl +z or Sz) in micaschists

Micaschists in the shear zone have a high quartz content which usually gives a heterogeneous domainal character to the foliation. It is possible to distinguish two microstructural types: phyllonitic foliation and schisto sity. Phyllonitic foliation (Sz) occurs at the base of the shear zone (Fig. 6) and is formed by very fine-grained minerals; biotite has a strong grain-shape preferred orientation and quartz domains have a mylonitic charac ter. Schistosity presents scarce microfolds mainly in two zones (Fig. 6). In Punta Riomar, the main foliation associated with D z microfolds is crenulated by late D z microfolds.

Shear band foliation

Small extensional shear bands within the schists are common in the shear zone (Fig. 5). Their spacing ranges from several decimetres to less than 1 em. Locally they appear as very short cleavage domains associated with gentle folding. They also occur as mesoscopic bands, which locally become small-scale normal faults that can also be interpreted as riedels. The angle between bed ding and shear bands ranges in most cases between 20° and 40°.

Conjugate sets of shear bands are locally present. In such cases, one of the sets is clearly dominant. Shear bands commonly dip east or southeastwards (Fig. 5). However, there are also cases where the shear bands dip westwards, presenting a reverse sense of movement. Field observations suggest a good correlation between the shear band orientation and their position on differ ent limbs of the major D 1 folds. In the normal limb the bands dip east or southeastwards, while in the reverse limbs they dip westwards. This means that the orien tation of the anisotropy controls the shear band orien tation. If shear bands are formed in some geometric relationship with the shear direction as several authors have proposed (Platt 1984, fig. 1, Weijermars & Ron deel 1984, the orientation of the shear bands can be explained by deformation partitioning (Lister & Wil liams 1983) (Fig. 8). In reverse limbs, the anisotropy represented by bedding is oblique to the shear direction and the shear bands can be related to slip or flow parallel to the anisotropy due to strain partitioning. If this is so, major D 1 folds would have gone on tightening during the late stages of D z owing to a flexural-flow or flexural-slip component generated at the reverse limbs. At the same time, the reverse limbs would have become thinner because of a coaxial component of flow.

Shear bands give a lenticular or sigmoidal microstruc tural aspect to the main foliation. During shear band development biotite and chlorite, and less commonly quartz, crystallized along the shear plane. The former have a well-developed preferred grain-shape orien tation. Quartz commonly forms fine-grained aggregates that define very narrow bands parallel to the shear bands, enhancing the mylonitic character of the rock. The presence of chlorite implies that the shear bands developed during retrograde metamorphism that can be correlated to the M 3 episode. These shear bands also affect D z folds and late D z biotite porphyroblasts. According to this, shear band formation represents a very late episode in the formation of the Mondoiiedo Nappe basal shear zone.

Quartz c-axis fabrics

The quartz c-axis orientation in quartzite samples from the shear zone and from adjacent zones has been analysed using a universal stage. Figures 4, 6 and 10 show the diagrams obtained for the shear zone and for the rocks that immediately overlie it, while Fig. 9 illus trates the results for samples from the footwall. Accord ing to the position of maxima and the symmetry relative to the foliation (Lister 1977, Schmid & Casey 1986, the following types of fabrics can be observed in these figures.

(1) A random fabric (R) (sample 301).

(2) Fabrics with orthorhombic symmetry (0). One of the samples (302) presents a maximum subperpendicu lar to the foliation and two other samples (B-8 and B-9) present maxima that are symmetrical with respect to the vertical N-S plane.

(3) Monoclinic small-circle girdle fabrics (SG) (samples 8, 43, 108, 125, 142 and B-2B). The best developed maximum (or pair of maxima) forms a smaller angle with the foliation than that of the other maxima.

(4) Monoclinic type I crossed girdle fabrics (CG). The line linking the inclined maxima can be orthogonal to the foliation (samples 20, 76, 95, 162 and B-4) or oblique (samples 11,300 and B-2). In some cases, one of the branches of the crossed girdle is poorly developed.

(5) Monoclinic incomplete single girdle fabrics (IG) (samples 25, 41, 47 and 74). These have a sigmoidal trend with two inclined maxima, except for sample 74 which has a single centred maximum. There is a com plete transition between this fabric and type CG.

Quartz c-axis fabrics along the shear zone display a variable degree of c-axes dispersion, with Rand 0 fabrics having the largest amount of dispersion and IG the least. The amount of dispersion increases away from the core of the shear zone (Figs. 6 and 9) and corre sponds to a gradation from higher to lower strain, and higher to lower metamorphic grade.

Some of the quartz c-axis fabrics from outside the shear zone (R and 0 types), and therefore not affected by D z , show a symmetry consistent with a coaxial defor mation (cf. Raleigh 1965, Tullis et al. 1973, Etchecopar 1977, Lister & Hobbs 1980. c-axis fabrics from inside the shear zone have a monoclinic symmetry (SG, CG and IG types) consistent with rotational strain due to the ....., simple shear component (cf. Bouchez & Pecher 1976, 1981, Etchecopar 1977, Lister & Hobbs 1980, Simpson & Schmid 1983. However, there are some samples outside the shear zone that also have a monoclinic symmetry (SG and CG types). Samples B2B and B4 present fabrics very similar to those in the shear zone and can thus be interpreted as local manifestations of D z· Peripheral maxima in the position found in our samples are indicative of low-temperature basal slip, ~h~reas peripheral and inclined or centred maxima may IndI~ate basal and rhombohedral or prismatic slip at medIUm temperature, all along an <a>-axis (Nicolas & Poirier 1976, Vernon 1976, Bouchez 1977, Nicolas et ai. 1977, Bouchez & Pecher 1981, Hobbs 1985. Inclined or centred maxima suggest higher temperature <a> slip on rhombohedral or prismatic planes (Bouchez & Pecher 1981, Hobbs 1985. The different slip systems active in quartz across the shear zone agree with the temperature changes recorded by the metamorphic assemblages (Fig. 6).

The distribution of c-axis maxima in CG fabrics, in the context of a retrograde metamorphism, suggests that these fabrics may be associated to a drop in temperature. This agrees with c-axis orientations of old and new grains in a mylonite from Punta Morago (Fig. 10). New recrys tallized grains have an increased dispersion of c-axes, associated with the activation of basal slip.

Although there is no one-to-one correspondence be tween microstructural and c-axis orientation types, a correlation of the two features of the rock fabric has been established (Fig. 6). Rand 0 fabric always corre spond to non-mylonitic tectonites. SG fabric commonly occurs in non-mylonitic tectonites, although it is also found in basal ultramylonites and, rarely, in blastomylo n.ites..CG fabri~ presents a wider microstructural variety SInce It occurs In all the microstructural types described here; this probably means that it has a transitional character between SG and IG types. Finally, IG type is recorded only in blastomylonites.

SHEAR ZONE MOVEMENT DIRECTION

Two types of criteria have been used: those that do not provide information on the movement directional sense like the mineral lineation, shear band foliation fold hinges and conical fold apical directions, and thos~ that do supply data about the movement directional sense like fold asymmetry and facing, oblique foliation by dynamic recrystallization in quartzite and quartz c-axis distributions. In order to visualize the movement direc tion, the direction of the first eigenvector of Bingham's distribution (Whitten 1966, Cheeney 1983) has been projected in Fig. 11 for several criteria in different sectors along the shear zone.

Movement direction indicators

The mineral lineation (L m ) is folded in some cases by D z folds. Hence, it does not record all the D z defor mation. In addition, the folding of L m causes its disper sion, which can be mainly observed in Punta Riomar (Fig. 5, localities 40-47). However, quartz c-axis fabrics show symmetries that generally agree with the position of this lineation, indicating that L m agrees with the movement direction in late stages of the deformation. The first eigenvector of L m distributions has a direction between N75°E and N115°E (Fig. 11), that is, with a dominant eastward component.

The relationship between the position of the shear bands and the type of major fold limbs (D I ) on which they develop implies that it is difficult to use the shear bands as a shear-sense indicator. However, shear band formation is related to shear zone deformation and therefore provides information about its movement direction during the late stage at which shear bands formed. Since So approximately coincides with the shear plane, the direction contained in So which is perpendicu lar to the intersection between So and the shear bands was used as a direction movement indicator. The first eigenvectors of the corresponding distributions have a fan-like distribution whose directions range between N65°E and N1400E (Fig. 11).

Fold hinge directions cannot be considered here as a good criterion to determine the shear direction, because the corresponding eigenvectors do not agree with those of the rest of the criteria, except for sector 1 (Punta das Cabras) (Fig. 11). The cause of the hinge directions dispersion has been discussed above.

Geometric and experimental models indicate that apical directions of sheath folds can be used to deduce shear direction regardless of means of fold nucleation (Cobbold & Quinquis 1980, Ramsay 1980. The direc tions of the first eigenvector of conical folds apical directions for different domains of the shear zone define a fan of directions between N700E and N115°E (Fig. 11).

Shear-sense indicators

.The strongly asymmetric D z folds always have a facing WIth an eastward component. This provides strong evi dence for eastward shear sense for conical folds apical directions. Oblique foliations by quartzite dynamic rec rystallization are common in our samples (Fig. 7c). This criterion (Means 1981) mostly indicates an eastward Fig. 9. Quartz c-axis fabrics (lower-hemisphere, equal-area projections; contours: 0.5, 1 and 2%) in the footwall of the Mondonedo Nappe basal shear zone. Numbers refer to locality and distance in metres below the basal thrust. movement directional sense. Asymmetry of the main Simpson & Schmid 1983) usually indicate the same shear girdle is the most suitable criterion in quartz c-axis direction. distribution diagrams in this shear zone (criterion 2 in The kinematic criteria are, in general, in good agree Simpson & Schmid 1983). In most cases, the asymmetry ment with a movement direction between N75°E and indicates a shear sense with a dominant eastward motion N115°E for the shear zone. Figure 11 shows that only of the hanging wall. Other criteria, such as those using two points corresponding to folds from two sectors are the fabric skeleton or the leading edge (criterion 3 in deviant. However, the distribution of kinematic criteria to a simple shear component approximately parallel to + 2 the normal limbs of earlier major D} folds. Evidence for this deformation is given by the development of several types of structures that occur at meso-and microscopic scales and that imply a general movement towards the foreland (eastwards). On the whole, these structures display some spatial variation that suggests a decrease in D 2 strain towards the upper part of the shear zone. However, this decrease presents some remarkable anomalies due to the existence of bands with a large number of folds (Fig. 12). Apart from ductile defor mation, it is worth mentioning that the shear zone is cut at the base by a major brittle thrust nearly parallel to the shear plane. The shear zone developed in an area with a pre existing metamorphic gradient (M}). During an eaFly upper. The basal part (Fig. 12) is characterized by the presence of phyllonites and some ultramylonites with in Fig. 11 does not allow deduction of temporal or spatial SG type quartz c-axis fabrics. This implies strong ductile systematic changes in the shear direction. deformation in a late stage at low temperature near to the basal thrust related to retrograde metamorphism. The middle part, whose eastern boundary coincides with SHEAR ZONE STRUCTURAL MODEL the andalusite isograd (Fig. 12), is characterized by the presence of blastomylonites and dominant IG and CG The Mondonedo Nappe basal shear zone shows evi quartz c-axis fabrics, which suggest a deformation in dence for strong ductile deformation which is mainly due amphibolite facies. Another notable feature is the pres- Abbreviations as in Fig. 2. ence of three bands with peculiar structural character istics. The first band is located in Punta das Cabras, where a large amount of subsimilar folds appear. The second band (Punta Riomar) contains a large'number of D z folds with variable geometries (Fig. 4) and inter ference patterns of Type 3; L m is frequently folded and quartz c-axis fabrics are SG here. The deformation of this band appears to have a long history that includes late movement during retrograde metamorphism. The third band (Punta Morago) contains a large number of folds and mylonites with quartz ribbons; mylonitization here took place during retrograde metamorphism although earlier than in the previous band. The upper part of the shear zone (Fig. 12) is featured by the presence of non mylonitic quartzite and by SG and CG c-axis fabrics, indicating deformation at a temperature that was on the whole lower than in the middle part. This is consistent with data on the metamorphism. There is no clear upper boundary to the shear zone, but a gradual transition into rocks that do not display shear zone textures. The boundary has been located at Ria de Foz (Fig. 12). However, some of the typical characteristics of the shear zone, like the blastomylonitic microstructures and the main foliation with microfolds in the schists disappear further to the west. Some other features still exist more to the east of Ria de Foz, like some quartz c-axis fabrics.

METAMORPHISM