A major shear zone within the Nagssugtoqidian of West Greenland

Tectonophysics, 27 (1975) 191-209 0 Elsevier Scientific Publishing Company, Amsterdam - Printed in The Netherlands

A MAJOR SHEAR ZONE WITHIN THE NAGSSUGTOQIDIAN GREENLAND

JENS BAK, JOHN KORSTGI(RD Geologisk

Institut,

(Submitted

Aarhus

OF WEST

and KAI S@RENSEN

Universitet,

Arhus

(Denmark)

November 25, 1974; revised version accepted April 11, 1975)

ABSTRACT Bak, J., Korstgilrd, J. and Sorensen, K., 1975. A major shear zone within the Nagssugtoqidian of West Greenland. Tectonophysics, 27: 191-209. A regional lineament of highly parallelized planar and linear structures can be followed over a distance of 150 km from the coast to the Inland Ice at Nordre Strdmfjord in central West Greenland. This shear zone is situated within the Nagssugtoqidian mobile belt and transects an area of intricate interference structures, from which it has been formed by high shear strain. By using data for the orientation of planar structures outside and within the zone, the shear strain is calculated to be approximately 6. This value applies to the shear zone close to the coast where its width is approximately 15 km. The zone is cut by the granulite/amphibolite facies boundary, and towards lower metamorphic grade the width of the zone decreases. At the same time the shear strain increases, so that the offset across the zone could be constant irrespective of width. At the coast, where the zone has been mapped, the fanning of planar structures shows the zone to be wedge-shaped, thinning upwards. It is therefore suggested that the zone represents a deep-seated ductile part of a major, transcurrent fault with a sinistral displacement of at least 100 km.

INTRODUCTION

The Nagssugtoqidian mobile belt (Fig. 1) was established by Ramberg (1948) on the basis of deformation and metamorphism of regional swarms of dolerite dykes (Escher et al., 1975). Ramberg (op.cit.) divided the mobile belt into three metamorphic complexes, a northern amphibolite facies complex, a central granulite facies complex, and a southern amphibolite facies complex. Detailed mapping between 1966 and 1969 in the Agto Area (Fig. 1) of the central granulite facies complex (Bondesen, 1970a) has contributed information on the metamorphic and structural evolution of this part of the mobile belt (Sorensen, 1970; Skjernaa, 1973; Winter, 1974). The existence of zones with pronounced linear patterns in the Nagssugtoqidian mobile belt was recognized by Noe-Nygaard and Berthelsen (1952), who ascribed the linear pattern to a horizontal course of the fold axes. Sorensen (1970) described these belts as features of large-scale transposition alternating

192

EGEDESMINDE

NORDRE STR0MFJORD SHEAR-ZONE

HOLSTEINS

BOR

Fig. 1. Part of the west coast of Greenland, showing the southern two thirds of the Nagssug toqidian mobile belt and the northern part of the adjoining Archean block. In frame the Agto map sheet. The granulite facies areas are shaded. The Nordre Strdmfjord shear zone is shown with solid lines where its boundaries have been mapped and with broken lines where inferred from aerial photographs.

belts where fold interferences dominate the structural pattern. Escher (1970) suggested that the linear belts represented zones where a horizontal stress resulted in a tangential shear-strain component and a normal flattening component. Escher et al. (in press) explain the linear belts in the Nagssugtoqidian mobile belt in general as having originated through simple shear strain. The linear belt, on which this account concentrates, is situated in the Nordre with

Fig. 2. The Agto map area. The arrows in the lower left and upper right corners indicate the boundaries of the shear zone. Mapping in the area by: J. Bak, 0. Bek, E. Bondesen, N. Roholt Jensen, V. Jensen, E. Kirsbo, J.A. KorstgBrd, V. Linderoth, B. Pedersen, SW. Platou, L. Skjernaa, K. Sdrensen, K. Thamdrup and J. Winther. Photo interpretation by E. Bondesen (1970b).

Stromfjord region and extends from the mouth of this fjord in an ENE direction towards the Inland Ice, a total distance of 150 km (Fig. 1). Its course and the general structural pattern within the Agto map sheet can be seen on the simplified structural map.(Fig. 2). Part of this Agto map sheet was mapped during the period 1966-1969 as an educational project run by professor E. Bondesen. A project is planned with the aim of mapping the remaining part of the map sheet and studying the complete shear zone. The analysis presented here is thus preliminary. AGTO MAP SHEET

Lithology

The dominant rock types in the area are hypersthene and/or biotite-bearing gneisses of granodioritic to tonalitic compositions. The principal minerals are plagioclase (An 30-40), quartz, ~kali-feldsp~, biotite and hypersthene, alkali-feldspar being mostly restricted to the hypersthene-free gneisses. Rocks of undoubted sedimentary origin occur as conformable bands and layers in

194

the hyperstheneand biotite-gneisses. These metasedimentary rocks cover a wide range in composition, mineralogy and texture from garnet-biotite-sillimanite gneisses to rusty, sulphide- and graphite-bearing schists, marbles and talc-silicate rocks, often interlayered with pyroxene amphibolites. All the rock types can be migmatitic. No unconformity between the hypersthene/biotite-gneisses and the metasedimentary rocks has been found. Structures The formation of the shear zone is believed to represent the last major structural event in the area and therefore overprints structures similar to those now found outside the zone. The structures occurring outside the shear zone have been described and interpreted by Sorensen (1970) and Skjernaa (1973). The mesoscopic folds of the map area in general have a similar profile. The interlimb angle is highly variable even for contemporaneous folds, and it is impossible to designate any single mesoscopic fold to any specific phase of deformation. In a similar way it has not been possible to relate a mineral lineation defined by sillimanite in the garnet-biotite-sillimanite gneisses and in rare cases by hornblende in the basic rocks to any specific event of deformation. At the present state of knowledge, the only difference on a mesoscopic scale between the shear zone and its surroundings appears to be the type of interference patterns found. In the shear zone, nearly all folds are co-axial, forming type-3 interference patterns (Ramsay, 1967) whereas all sorts of interference patterns are developed outside the shear zone. On the basis of the macroscopic structures, it is generally possible to distinguish at least four phases of deformation outside the shear zone, the last of which has moderately dipping ENE trending axial surfaces with a highly variable pitch. This pattern is obvious in the northwestern part of the Agto map sheet (see Fig. 2). Also within the shear zone several phases of folding can be recognized, but in contrast to the surroundings, the axial surfaces are subvertical, the folds being tight to isoclinal and with a pitch varying between horizontal and 20”. In general, the amplification of folds is greater than outside the zone. The interference patterns seen are apparently simple type-3 patterns locally modified by open NW-folds. These open folds may be regarded as former tight folds opened during the formation of the shear zone. A correlation between the structures within and outside the shear zone is at the moment not possible, neither is it possible to decide whether the shearzone formation is contemporaneous with any of the folds in or outside the zone. The transformation from the surroundings into the shear zone takes place across a transition zone in which the lithological boundaries are gradually rotated towards parallelism with the shear zone (Fig. 2). The sense of rotation is anti-clockwise towards the shear zone, indicating a sinistral “displacement” across it. As a demonstration of the geometrical effect of the shear zone on the orientation and variability of the mesoscopic structures a series of stereographic

195

Fig. 3. Stereographic projections of mesoscopic structures. S1 = fold axes, from three subareas within the shear zone, 1 = a zone zone; 2 = a transitional zone; 3 = the shear zone. Contours at 0, 1% area. Equal area, lower hemisphere. The increasing degree of evident.

projections have been (Fig. 3). The diagrams tures within the shear the data we have used ANALYSIS

planar structures, LN = well outside the shear 4 and 8% observations per parallelism from 1 to 3 is

prepared for the foliation (S,) and the fold axes (J&) show the high degree of parallelism between the struczone as opposed to the structures outside. In treating the “EDP system Agto” developed by Platou (1971).

OF THE MAP SECTION

OF THE

SHEAR

ZONE

Assumptions Our analysis of the strain within the shear zone has been based on assumptions which are mentioned below and further discussed later in the paper. Analyzing first the two-dimensional information offered by the map we have made assumptions along two lines: first on the chronologic relation between the shear zone and its surroundings, and secondly on the kinematics of the deformation. Significance of chronology The geometrical element used in the interpretation of the shear zone is planar structures, of which compilations are shown in Fig. 3. From the geometry of these structures alone we cannot rule out one in favour of the

196

other of the following possibilities: (1) the shear zone formed as a late event by sinistral shear of structures like those now found in the surroundings; or (2) the surroundings have formed by an equal amount of dextral shear from rocks with structures similar to the shear-zone structures. For convenience we proceed with the analysis assuming the shear zone to be a late feature superimposed on an area like its present surroundings. We then reach a measure for the strain that could be termed the incremental strain from the su~oundings into the shear zone. For the calculation of strain the chronologic distinction above is immaterial since it only has bearing on the designation of the strain. Assuming the second chronologic possibility the strain obtained by using the first would be the reciprocal strain for the surroundings. As will be shown in the following calculation of the strain, the increase in strain across the zone boundary must be extremely high. The shear strain is calculated using the reorientation of planar structures across the zone boundary. High shear strain in a zone like this one can be a result of several combinations of orientations of deformed and non-deformed planar structures to the orientation of the shear plane. The high shear strain of the Nordre Stromfjord shear zone is due to the small angle between the reoriented planar structures and the shear plane (us of Fig. da). The negligible influence of the predeformation orientations (except for highly speculative ones) on the shear-strain value means, that any variation of the orientation of the planar structures outside the shear zone contemporaneous with the deformation within the shear zone is of no consequence in this context. In case the planar structures of the surroundings were deformed during the deformation within the shear zone (the third chronologic possibility) our calculation still gives the value of incremental strain across the zone, that is the strain of the zone per se. Choice of strain model With the restricted knowledge of the shear zone so far obtained we cannot evaluate the heterogeneity of strain within the zone. Whatever this could be, the off-set of a planar structure entering the zone at one side and coming out at another side can be taken as a measure of the integral strain within the zone, assuming that no later homogeneous deformation of the zone and its environment has taken place. The integral strain of the zone is obtained by using averaged data for the orientation of planar structures outside and within the zone, as mapping has not proceeded so far that the integral strain can be evaluated by tracing planar structures through the whole of the shear zone. The concentration of fold axes within the zone is subhorizontal (Fig. 3) and the boundaries of the zone are well-defined. The facts suggest that a simple shear model is appropriate. The direction of movement is assumed to be subhorizontal and parallel to the zone boundaries, and the flow plane to be subvertical. In choosing a simple shear model we follow the strain compatibility arguments of Ramsay and Graham (1970) and Escher et al. (1975).

197

WC

1% 140 1x 1X 110

lot

s

90 80

m 60 50 40 30 20 10 0

0

10

20

30

Lo

50

60

70

60

90

100

110

120

130

iL0

150

160

170

180

"1

Fig. 4. Shear strain in a plane perpendicular to the shear plane and containing the shear direction (indicated by the set of opposing arrows). a) Shear strain (y) of rotated line (Z, to 2;) as related to the angular shear (I//). b) Chart of shear-strain values as a function of the angle between the rotated line and the shear direction (ug of Fig. 4a) and the angle of rotation (ul of Fig. 4a). The broken line, along which the values of shear strain are written, indicates the sets of values ul/uz for which the rotation brings a line from its original position to a position symmetric about the perpendicular to the shear direction. These are minimum values of shear strain for a given angle of rotation.

Strain in the horizontal section Our assumption about chronology means that we consider a deflection of planar structures crossing the zone boundaries as caused by shear strain within the zone. When specifying the sense of shear (sinistral or dextralf it is in accordance with this assumption. The calculation of the strain is straightforward assuming the simple shear model (Ramsay, 1967; Ramsay and Graham, 1970), in which case we need not assume the orientation of the principal axes of the strain ellipse known.

198

Calculation

of shear strain

With a horizontal shear direction and a vertical flow plane the horizontal plane parallels the X2-plane of the finite-strain ellipsoid, i.e. the most eccentric section (X > Y > 2). In a shear zone expressed through a reorientation of existing planar structures (as opposed to a shear zone expressed by newly formed, planar structures) the known variables are the angle of rotation (ur of Fig. 4a) and the angle between the reoriented feature and the shear direction (up). For largescale structures the boundaries of the shear zone and therefore the angle u2 is probably only known approximately due to lack of continuous exposures, the existence of transitional zones and later, cross-cutting faults. This is in most cases bound to impose a high absolute uncertainty in the strain estimate for high strains, a fact which can be read from the shear strain chart (Fig. 4b). At several localities within and outside the map area an anti-clockwise rotation of planar structures across the zone boundaries can be seen, and at localities 1 to 4 (Figs. 2 and 5) the angle of rotation (ur ) is well defined.

16-

51 LOCALITIES STRAIN HAS

WHERE BEEN

SHEARCALCULATED

Fig. 5. Shear strain at four localities (synonymous with u2 of Fig. 4).

53 STRIKE

as dependent

on the choice

55 OF

SHEAR

57

59

PLANE

of shear-plane

direction

199

From the map (Fig. 2) the approximate strike of the shear zone is seen to be 55”) this also being the shear direction. At the four localities the angle u2 varies between 4 and 8” and the angle of rotation between 24 and 63”. From Fig. 4b these numbers are seen to give shear strains of between 6 and 12. But as seen from the chart the value of shear strain is almost exclusively determined by the value of the angle u2, the angle u1 being only significant for values smaller or very much larger than the ones found here. Thus an uncertainty in the determination of the angle u2, i.e. in the determination of the zone-boundary strike results in large absolute uncertainty in the value of shear strain (Fig. 5). Therefore the spread in shear strain determined for the four localities close to the zone boundaries does not mean that the spread is real; it might indicate that our choice of shear direction is wrong. Indeed, the general strike of the shear zone increases when moving from the coast to the ice and the actual strike of the shear-zone boundary, and thus the angle u2 could be smaller for locality 1 than for 2, 3 and 4 lying close together both geographically and with respect to shear-strain values (Fig. 5). Averaging planar structures we obtain a shear strain of 6 for the zone as a whole. The strain ellipse corresponding to this shear strain has an axial ratio 40/l (X/Z) and the strike of the long axis is close to 65”, i.e. very close to the maximum concentration of strike orientations (Fig. 3).

\

d =T-W.SOKM da5 =(T+ TAN 35)W.lOO

KM

0

1 IO

20 KI

Fig. 6. Off-set (d) across the shear zone for a shear strain of 6 (7) and a width of the zone (w) equal to 15 km. The majority of planar structures entering the shear zone at its southeastern boundary do so with an EW strike. The “apparent displacement” of this direction is d35.

200

Off-set across the shear zone The horizontal off-set of a planar structure by the shear zone assuming the strain to be homogeneous within the zone is the shear strain multiplied by the width of the zone (Fig. 6). I

II

m

IP

1.2.75

Shear

direction

Fig. 7. Redeformation of the “Kangilerssua structure” (B in fig. 2 and I of this figure). The strfcture is progressively redeformed by simple shear along vertical shear planes striking 55 . The shear strain (7) is given at each step and visualized through the deformation of a square. Open folds with highly variable axial trends within the shear zone (a, b, and c) have been interpreted as separate, late folds. By redeformation they become tight folds with mutually parallel axial surfaces parallel to the most prominent axial surface trends of the structures of the surroundings (compare step V with the inset, which is a map of the nearby “Qinqaq structure”, A in Fig. 2). Both structures have been mapped by SW. Platou (1970).

201

The strain is certainly not homogeneous, and as we define the zone we have not considered a transition zone approximately 2 km wide in which the shear strain rises from zero to 6-7. Within the shear zone the value of shear strain attains at least 6 over a distance of 15 km. So the minimum off-set across the zone and its transition zone is 100 km. Geologically this means that not one lithologic unit within the map sheet can possibly occur on both sides of the shear zone. Until mapping of a larger part of the shear zone has been carried out we do not want to consider the possible heterogeneities in strain across the zone. Redeforma

tion

Following the simple shear model a test on structures outside and within the zone has been performed. This test does not prove the correctness of the assumptions made, but it does show that structures like those seen outside the shear zone can be produced by redeforming the shear zone structures according to the chosen model. As the shear zone has developed through sinistral shear we redeform a shear-zone structure by dextral, simple shear with a subvertical shear plane striking 55”, the shear direction being subhorizontal. The test was carried out on a deck of paper and the results are shown in Fig. 7. Step V should be compared with the inset (Fig. 7) which is a detailed map of an area outside the shear zone. As will be shown later, the shear zone represents the last major structural event in the area, and the following discussion will take the deformation (in nature) as progressing from step V to step I (Fig. 7). From the figure the following may be seen: (1) The shear-zone structure (I) as now found can be produced by deforming a structure (V), which is very similar to the kind of structures found outside the shear zone (Fig. 7 inset, and Fig. 2). (2) One of the effects of the shear-zone formation is an opening up of tight NE-folds (e.g. a in Fig. 7), and a simultaneous anti-clockwise rotation of their axial surfaces into a NNW orientation. The opening up of some folds and closing of others may result in a pattern which can erroneously be interpreted as a fold interference pattern, in this case a refolding of the closed structures by the open structures. Thus folds of very different style and orientation could erroneously be designated to different phases of deformation, such as the folds b and c of Fig. 7. As now found in the shear zone (b and c of step I, Fig. 7) they differ widely in both respects, but may have attained this difference by the shear-zone deformation (compare with their orientation and style in step V, Fig. 7). THREE-DIMENSIONAL

ANALYSIS

Metamorphism

The Agto area is situated in the northernmost part of the central granulite facies complex of the Nagssugtoqidian mobile belt (Fig. 1). This peripheral

202

position of the area is reflected in the coexistence of parageneses with and without orthopyroxene. The transition from granulite facies in the central complex to amphibolite facies in the northern complex is gradual, with a considerable amount of interfingering (Fig. 1). There seems to be no reason to assume a time difference as an explanation for the difference in metamorphic grade in the two complexes (Escher et al., in press). In the Agto area granulite facies parageneses are found within the shear zone as well as outside, and the distribution of parageneses is only affected by the shear zone to the extent that the shear zone affects the lithological boundaries. Following the shear zone to the NE from the Agto area (Fig. 1) it is cut by the granulite-amphibolite facies boundary and the metamorphic event responsible for the present distribution of metamorphic facies in the area therefore outlasted the formation of the shear zone. The position of the granulite-amphibolite facies transition is known from reconnaissance work only, but considering the magnitude of the off-set of the shear zone the chronologic relation between metamorphism and shear-zone deformation seems clear enough. The orthopyroxene-bearing parageneses in the area belong to the hornblendegranulite facies (De Waard, 1966) and orthopyroxene-free parageneses to the sillimanite-almandine-orthoclase subfacies of the amphibolite facies (Winkler, 1967). The metamorphic

transition

and a model

Tracing the shear zone eastwards towards the ice it is seen to decrease in width. Several features indicate that this thinning is brought out by a difference in erosion level. The metamorphic transition from the coast to the ice gives a first indication of the nature of crustal level difference. The highest metamorphic grade is found at the coastal end of the shear zone, the lowest near to the ice. This transition from well inside the granulite facies to low amphibolite facies indicates that the difference is from lower to higher erosional level from coast towards the ice. The thinning of the shear zone in the same direction then indicates an upwards thinning of the zone. In our model of the shear zone we want to connect it with a known high-level feature, and we suggest a (major) transcurrent fault. In Fig. 8 we have depicted our model, splitting it into two parts. Fig. 8a reflects the upwards thinning and for this we can present unequivocal evidence. Fig. 8b represents the fault-part of the model, for which evidence can be found in the strain associated with the shear zone in its narrow part at the ice. Dip variation at the coast: wedge shape of the zone For the shear zone to be contained within a crust of reasonable thickness and vanish to zero width well below the crustal surface its two boundaries

203

Fig. 8. Three-dimensional model and interpretation of the shear zone. The shear zone is wedge-shaped (a) and is interpreted as a ductile, transcurrent fault feature (b). With constant off-set (d) across the shear zone irrespective of width, the relation between shear strains pg;;zd 72) and widths (WI and luz) to be fulfilled, is derived according to part c of the

1

.

have to make an appreciable angle. If the present-day coast exposure originally formed at, say, a depth of 30 km and the shear zone had transformed into a fault or fault zone of negligible width at a depth of 10 km, the angle between the boundaries has to be 45”. Thus the interfacial angle between the shear-zone boundaries has to be expressed through a fanning of planar structures following the proposed model. In Fig. 9 stereographic projections of planar structures from five strips parallel to the shear zone in the mapped coastal part are shown. From the northwestern strip to the most southeastern there is an angular difference of about 40” judging from the maximum pole concentrations and a regular increase in dip towards SE in the shear zone. As the SE boundary is not exposed in this part of the zone, it can be safely assumed that the planar structures along the SE boundary have shallower dips towards the SE than those of subarea 3.5 in Fig. 9 (dip: 80” towards SE). Thus the expected fanning of planar structures does occur, but may not be symmetrically dispersed around the vertical. This does not imply that the shear zone is asymmetric, i.e. that the shear-zone boundaries are not of approximately equal dip. Most likely the planar structures have rotated in the same sense all through the zone during the shear-zone deformation, rotating towards the Y-axis of the strain ellipsoid (Fig. 9). In this way the shear zone may be symmetrical (see also Fig. 10). Assuming the deformation of the shear zone to have proceeded with constant volume ( an assumption contained in our use of the simple shear model) the dip of the Y-axis might be calculated.

204

NNW

SSE

Fig. 9. Stereographic projections of planar structures within the southwestern, mapped part of the shear zone. From the NW towards the SE there is a fanning of planar structures illustrated in the lower right of the figure, which is a vertical “profile” through the shear zone perpendicular to the X-axis of the strain ellipsoid. The profile shows in a qualitative way the variation in dip of the planar structures (thin, full lines) in relation to their original positions (thin, broken lines). The Y- and Z-axes of the strain ellipsoid are shown in heavy lines. The Y-axis variation indicates qualitatively how the shear zone can be symmetrical about the vertical without the planar structures being so.

As the angle of rotation of the planar structures is not known well enough Fig. 9 only indicates qualitatively how the Y-strain ellipsoid axis is dispersed relative to the planar structures of the shear zone. The Y-axis is contained within the shear plane and thus gives almost exactly (not exactly because its dip direction with the given strain is 10” from the dip direction of the zone boundaries) the dip variation of the shear plane, and thus to the NW and SE the dip of the zone boundaries. With reference to Fig. 9 we can say that the shear zone could be symmetrical around the vertical. If we assume: (1) a vertical shear plane; (2) a shear strain of 6 centrally in the zone; and (3) constant volume deformation, then the complete solution to the finite strain would be: x: Y:Z = 6:1:0.16 with the X-axis horizontal and with izing” the strain-axis values so as to With the present lack of mapping calculation will not be reconsidered shear plane.

a trend of approximately 65” (“normal+ make Y = 1). in a greater part of the zone, the strain to accord with the varying position of the

205

The shear zone at the ice: Constant off-set? From the NE corner of the Agto map area the shear zone can be traced to the Inland Ice on the reconnaissance maps covering the inner part of the Arfersiorfik Fjord (Fig. 1). These maps incorporate photo interpretations and coast reconnaissance mapping by G. Henderson (Geological Survey of Greenland). These observations are incorporated in the 1: 500,000 S$ndre Stromfjord-Nugssuaq map (G.G.U., 1971). The shear zone is probably faulted in places in the fjords, but from Arfersiorfik to the ice its boundaries are well-defined. Returning to the model (Fig. 8b) we would expect - as a consequence of the interpretation of the shear zone as a deep-seated part of a fault - the offset across the zone to be near constant irrespective of width. For this condition to be fulfilled the shear strain has to vary inversely with width (Fig. 8~). At the Inland Ice the shear zone is approximately half as wide as the coast, and thus the shear strain has to be close to 12 to fulfill the condition for constant off-set (Fig. 8~). From the reconnaissance maps the most reasonable choice of relevant angles (compare G.G.U., 1971) corresponds to a shear strain of about 15. Considering the necessarily high uncertainty (see Fig. 4b) we only want to conclude that probable values of shear strain both at the ice and at the coast lie within certain limits such that the condition for constant off-set is fulfilled. The shear strain at the Inland Ice is certainly much higher than at the coast. DISCUSSION

OF ASSUMPTIONS

Having evaluated the strain of the zone, a discussion of the assumptions on which this analysis is based is now relevant. The assumptions made were about chronology and kinematics.

Chronology We assumed the shear zone to be relatively younger than its surroundings. The material on which calculation of the incremental strain across the shearzone boundaries has been based cannot be used to decide whether this assumption is correct or not. A consideration of the spatial consequences of the chronologic models makes a choice between them possible. If we assume the surroundings to have formed as a late event the magnitude of strain associated with this deformation would be the same as that found for the shear zone in the preceding section. As the shear zone is wedge-shaped and thins upwards, the neighbouring blocks must thicken upwards. A reversal of the chronology then faces the problem of accounting for an increasing or constant strain in a zone of increasing width and decreasing depth. This consequence makes a reversal of the assumed chronology impossible. The solution to the question about the chronology will probably be clearly demonstrated

206

close to the ice where the shear zone cuts the northern part of a large diorite body. Outside the shear zone this body has been seen (on a short reconnaissance trip) to preserve its original magmatic texture, while according to the reconnaissance maps of G. Henderson within the shear zone it is a schistose rock. This relation indicates that the diorite intruded prior to the formation of the shear zone and after the formation of the bulk of the structures of the surroundings, and thus it probably will, when studied in detail, confirm the chronology presented above. Kinematics.

Graphical solution

Calculating the strain we assumed the shear plane to be steep and the shear direction subhorizontal. These quantities can be found graphically. The problem is equal to that of manipulating simple shear as a model for the formation of similar folds and similar refolding (Weiss, 1955). The shear plane of simple shear is a plane of no finite deformation and can be determined when the orientation of two non-parallel lines within it is known, in a way similar to finding the axial plane of a similar fold (the shear-zone boundary is, in a way, a similar fold). One line is the strike of the shear-zone boundary. Another line is defined by the intersection of a deformed plane and its undeformed equivalent or by the intersection of progressively deformed planes. This line of intersection has to be contained within the shear plane to fulfill the condition of no finite deformation, and together with the strike line of the shear-zone boundary it defines the shear plane. To obtain the shear direction we have to locate a linear direction which has been reoriented. This could be the line of intersection of two planes (e.g. dykes) or it could be a fold axis. It is crucial to the strain evaluation to obtain the shear direction, as the strain in the plane X2 of the strain ellipsoid will vary from a finite minimum to infinite with a variation in choice of the shear direction. For the Nordre Str$mfjord shear zone the choice of shear direction as horizontal is close to that giving the minimum X2-ratio. We have no girdle of reoriented linear structures to rely on in finding the shear direction. The girdle “g” of Fig. 3 (left diagram, area 3) is defined by very few points and furthermore this girdle represents fold axes undeformed by the shear-zone deformation, because of parallelism to the shear plane rather than differentially deformed axes. At this high strain the subhorizontal maximum of fold axes has to be close to the shear direction. For a major shear zone the shear direction can be found in still another way as the line of intersection between the different oriented shear planes, and for the Nordre Str$mfjord shear zone this method can be applied (see Fig. lo), the line of intersection between the shear-zone boundaries being the shear direction. (In Fig. 10 the method of graphical solution has been applied to the Nordre Stromfjord shear zone.) In conclusion we can summarize that the reorientation of one planar structure is sufficient to determine the shear plane given the strike of the shear-

207

Fig. 10. Finding the shear plane and shear direction ~aphi~ally. 5’ = the strike direction of the shear-zone boundaries, considered parallel within the map area. The line projections (PI and Pz) are found as the intersections between deformed (iI and 12) and undeformed (I; and Ii) planar structures. The shear planes are then SP, and SP2. SP1 is the shear-plane orientation close to the NW boundary of the shear zone in the mapped area close to the coast. SP, is the shear-plane orientation along the northern shore of Nordre Strqjmfjord close to the coast. The actual SE boundary of the shear zone lies further to the southeast, and may have a shallower dip than SP,. The SE boundary might be close to SP; corresponding to an orientation 13 o,f the planar structures along the southeast boundary. The angle between SP1 and SP2 is 30 , between SP, and SP2 40 . The ruled areas are those of maximum concentrations of linear and planar structures of Fig. 3 (diagram 3).

zone boundary, and that a reoriented line is necessary and sufficient to determine the position of the shear direction. For major shear zones with a high degree of variability of shear-plane orientations, the shear direction comes out directly as the line of intersection between the shear planes. Having obtained the shear plane and the shear direction the angles used in the determination of the strain (u, and u2 of Fig. 4) can be obtained graphically if not directly obtainable from the map, i.e. if the shear direction is not horizontal. SUMMARY

OF RESULTS

The results of the analysis of the Nordre Str$mfjord shear zone can be summarized as follows: (1) The strain. Assuming the shear zone to have formed by simple, sinistral shear of structures like those now found in its surroundings, the incremental strain across the shear-zone boundaries can be calculated applying the deflec-

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tion of planar structures. The validity of the assumptions concerning the orientation of the shear plane and the shear direction can be checked graphically. The shear strain of the zone gives an off-set across the zone close to 100 km. The uncertainty in the calculation of the shear strain is high, and any variation of shear strain along the strike of alike structures, like that reported for the northern boundary zone of the Limpopo mobile belt (Coward et al., 1973) has to be carefully considered before given any significance. And furthermore, differences in erosion level in a wedge-shaped shear zone will affect the shear-strain value. (2) Morphology of the shear zone. It can be concluded for the Nordre Str$mfjord shear zone from the variation of dip of planar structures across the shear zone in the minor part so far mapped, that the shear zone is wedgeshaped, closing upwards. At the coast the angle between the boundaries of the shear zone is approximately 40”. (3) Interpretation of the shear zone. Its wedge shape confirmed, it is proposed that the shear zone changes upwards into a fault. As a consequence of this model, we assume that the shear strain of the zone has to vary with varying width of the zone so as to give a constant off-set. This assumption is checked towards the shear zone at the ice where its width is about half that at the coast. The shear strain at the ice can fulfill the requirement needed for constant off-set but the evidence cannot be considered conclusive. Though our model of the Nordre Stromfjord shear zone must be characterized as built on multiple assumptions, it satisfactorily explains the fanning in dip across the zone, its variation in width and the variation in shear strain. The interpretation is strengthened by the decrease in metamorphic grade in the same direction as the decrease in width of the shear zone. ACKNOWLEDGEMENTS

The field work has been supported by a grant from the Carlsberg Foundation to professor E. Bondesen and financially and logistically by the Geological Survey of Greenland, whose director, K. Ellitsgaard-Rasmussen we thank for permission to publish this paper. WC furthermore thank E. Bondesen for his guidance during the field work, Lissie Jans, Inger Tofte, and Else Moltke Nielsen for their efforts at the drawing board and typewriter.

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De Waard, D., 1966. On water-vapor pressure in zones of regional metamorphism and the nature of the hornblende-granulite facies. Verh. K. Ned. Akad. Wet., Afd. Natuurk., Reeks 2, 69: 453-458. Escher, A., 1970. The general structural pattern of the Nagssugtoqidian erogenic complex between S$ndre Stromfjord and Jakobshavn, West Greenland, In: Colloquium on Nagssugtoqidian Geology. Geological Institute, Aarhus University, pp. 4-S. Escher, A., Escher, J. and Watterson, J., 1975. The reorientation of the Kangamiut dyke swarm, West Greenland. Can. J. Earth Sci., 12: 158-173. Escher, A., Sdrensen, K. and Zeck, H.P., in press. The Nagssugtoqidian Mobile Belt. In: Geology of Greenland. Geological Survey of Greenland, Copenhagen. G.G.U. (Gronlands Geol. Unders.), 1971. Geologisk Kort over Gronland. S$ndre Stromfjord- Nugssuaq 1: 500 000. Noe-Nygaard, A. and Berthelsen, A., 1952. On the structure of a high-metamorphic gneiss complex in West Greenland, with a general discussion on related problems. Medd. Dansk Geol. Foren., 12: 250-265. Platou, S.W., 1970. Note on the structure and metamorphism in the area between Ataneq Fjord and Gieseckes S& In: Colloquium on Nagssugtoqidian Geology. Geological Institute, Aarhus University, pp. 28-40. Platou, S.W., 1971. An electronic data processing system for the geological field and laboratory data. The E.D.P. system Agto. Rapp. Gr@rlands Geol. Unders., 39, 42 pp. Ramberg, H., 1948. On the petrogenesis of the gneiss complexes between Sukkertoppen and ChristianshHb, West Greenland. Medd. Dansk Geol. Foren., 11: 312-327. Ramsay, J.G., 1967. Folding and Fracturing of Rocks. McGraw-Hill, New York, 568 pp. Ramsay, J.G. and Graham, R.H., 1970. Strain variation in shear belts. Can. J. Earth Sci., 7: 786-813. Skjernaa, L., 1973. Precambrian structures of the Ikorfat peninsula, Agto region, West Greenland. Rapp. Gronlands Geol. Unders., 52, 22 pp. Sorensen, K., 1970. Some observations on the structural and metamorphic chronology on Agto and surrounding islands, central West Greenland. Rapp. Gronlands Geol. LJnders., 27, 32 pp. Weiss, L.E., 1955. Fabric analysis of a triclinic tectonite and its bearing on the geometry of flow of rocks. Am. J. Sci., 253: 225-236, Winkler, H.G.F., 1967. Die Genese der metamorphen Gesteine. Springer, Berlin, 2nd edit., 237 pp. Winter, J., 1974. The Precambrian geology of the Tungarnit area, outer Nordre StrQmfjord, central West Greenland. Rapp. Gronlands Geol. Unders., 61, 17 pp.