Evidence for 1.82 Ga transpressive shearing in a 1.85 Ga
granitoid in central Sweden: implications for the regional
evolution
Karin Ho¨gdahl
a,*, Ha˚kan Sjo¨stro¨m
baSwedish Museum of Natural History,Laboratory for Isotope Geology,Box50007,SE-10405Stockholm, Sweden bDepartment of Earth Sciences,Uppsala Uni6ersity,Villa6a¨gen16,SE-752 36Uppsala,Sweden
Received 28 February 2000; accepted 15 July 2000
Abstract
Two crustal-scale shear zone systems, the Storsjo¨n – Edsbyn deformation zone and the Hassela shear zone join in the western part of central Sweden and form an anastomosing pattern of ductile shear zones. Several of these, up to 1 km wide, zones truncate an intrusion previously defined as a component of the major, ca. 1.80 – 1.77 Ga magmatic suite (the Revsund granitoids) in north central Sweden. However, within the southern part of the intrusion, titanites that are components of the fabric in a pervasively deformed shear zone (Forsaa˚n) have a U – Pb TIMS age of 181692 Ma. Zircon porphyroclasts in the same fabric yield a U – Pb SIMS age of 1849914 Ma, overlapping with the U – Pb TIMS age of 185192 Ma for titanites from an undeformed granitoid. These results are interpreted to define the ages of ductile shearing and magma emplacement, respectively. This interpretation is supported by microstructures as well as the regional structural pattern characterised by shear zones enveloping virtually undeformed pods of the granitoid. Previously reported and locally existing magmatic flow structures indicate some syn-emplacement deformation. The kinematic results, characterised by a high proportion of pure shear within the granitoid and oblique dextral shearing along its margin, reflect strain partitioning under transpressive conditions. The kinematic and geochronological results fit into the regional framework of late-orogenic localised deformation along crustal scale shear zones. Recent, absolute datings of shear fabrics defines ca. 1.82 – 1.80 Ga as a period of shear zone activity in central Sweden. Temporally, this period overlaps with the emplacement of S-type granites and granitoids derived from deeper crustal levels. A similar situation exists in southern Finland. Structural, geochronological and geophysical data suggest that the crustal scale shear zones in central Sweden and southern Finland may be related temporally and spatially. The age and the post-tectonic nature (on a regional scale) of the dated 1.85 Ga granitoid have important regional implications. As both the age of the intrusion and the ductile shearing are older than the established age range (1.80 – 1.77 Ga) of the Revsund granitoids, either an extended intrusive history for the suite is indicated, or the granitoids in this part of the type areas should be excluded from the Revsund suite. Their post-tectonic nature implies that the pervasive Svecokarelian deformation in the area occurred earlier than generally assumed (1.85 – 1.80 Ga). On the other hand, approximately coeval, calc-alkaline granitoids to the south-east of the investigated area are folded and foliated
www.elsevier.com/locate/precamres
* Corresponding author. Fax: +46-8-51954031.
E-mail addresses:[email protected] (K. Ho¨gdahl), [email protected] (H. Sjo¨stro¨m).
K.Ho¨gdahl,H.Sjo¨stro¨m/Precambrian Research105 (2001) 37 – 56 38
pervasively, indicating a later Svecokarelian evolution in that area, i.e. that the granitoids are located in domains with separate orogenic evolutions. © 2001 Elsevier Science B.V. All rights reserved.
Keywords:Palaeoproterozic; Shear zone; Svecofennium; Transpression; U – Pb geochronology
1. Introduction
1.1. Background and aim
A spatial connection between granitic intru-sions and large-scale deformation zones is com-mon and well-established, while evidence for coeval intrusive and tectonic activities is a more recent finding (Hutton, 1988; D’Lemos et al., 1992; Saint Blanquat et al., 1998; Brown and Solar, 1998a,b). The study of such conditions has led to refined models concerning e.g. diagnostic internal structures with respect to tectonic envi-ronment and feed-back relations between melt transfer and shear zone activity (e.g. Brown and Solar, 1998a,b).
In the central part of the Fennoscandian shield, the relationship between shear zones and granitic intrusions is fundamental for the understanding of the Palaeoproterozoic crustal evolution, granitic intrusions of various ages and origins are frequent (Stephens et al., 1994, 1997; Lundqvist, 1995) and the number of identified shear zones is increasing (e.g. Ehlers et al., 1993; Ka¨rki et al., 1993; Bergman and Sjo¨stro¨m, 1994; Stephens et al., 1994, 1997; Wijbrans et al., 1995; Beunk et al., 1996).
So far, the few published ages of tectonic fabric in ductile shear zones, as well as indirect evidence, show that shearing overlapped the periods of granitic plutonism temporally (Ho¨gdahl and Sjo¨stro¨m, 1999; Ho¨gdahl et al., 1995, 1996; Korja and Heikkinen, 1995; Sjo¨stro¨m and Bergman, 1995; Stephens and Wahlgren, 1995, 1996; Wahlgren and Stephens, 1996; Lindroos et al., 1996). This general condition is valid both for intrusions derived from the deeper levels of the crust as well as S-type granites derived from shal-lower crustal levels. However, in the Svecofennian domain, the relationship between specific intru-sives and shear zones is not well understood, which emphasises the need for combined geochronological and structural investigations.
Our study focuses on such a relationship by defining the magmatic age of a granitic rock as well as the age of a ductile shear zone within the intrusion. Microstructures are studied to find out the physical conditions during deformation, and the role of minerals in the deformational fabric, which are suitable for U – Pb analyses, e.g. titanite and zircon. Combined with structural analysis in the field, the microstructures are used to sort out the kinematic conditions along the shear zone and along one of the margins of the granitoid. Ages and characteristic features of adjacent shear zones within the granitoid are presented briefly and the local conditions are integrated into the regional picture.
Three main factors outline regional implica-tions of this work, (1) The area represents the junction of the two largest shear zones in central Sweden, one of which may continue to southern Finland (Sjo¨stro¨m et al., 2000). (2) The investi-gated intrusion has previously been referred to as a part of the Revsund granitoids (Ho¨gbom, 1894; Lundega˚rdh et al., 1984). These granitoids make up a major Palaeoproterozoic intrusive suite of north central Sweden and compared with their areal extent, existing age determinations are scarce. Our results have consequences for the definition of a Revsund granitoid in one of the type areas. (3) The results also indicate that per-vasive Svecokarelian deformation in the area may have been earlier than generally assumed for the Palaeoproterozoic evolution of north central Sweden.
1.2. The Re6sund granitoids: regional occurrence
and local character
Claes-son and Lundqvist, 1995; Delin, 1996; Delin and Aaro, 2000), are interpreted to have a deep crustal origin (Claesson and Lundqvist, 1995) with input from the mantle (Gorbatschev et al., 1997). Their emplacement is assumed to have occurred after Svecokarelian peak metamorphism and deforma-tion and they have, therefore, been classified as post-orogenic (Gaa´l and Gorbatschev, 1987), or post-kinematic. It has been suggested that the Revsund suite belongs to the Transscandinavian Igneous Belt (TIB) (Gorbatschev and Bogdanova, 1993) extending from southern Sweden to northern Norway, partly below the Scandinavian Cale-donides (Fig. 1).
The Revsund granitoids in Ja¨mtland county constitute the southernmost component of this intrusive suite. Based on differences in major ele-ment composition, the granitoids in the two areas first described by Ho¨gbom (1894), were treated subsequently as separate massifs (Gorbatschev et al., 1997), a northern (Fja¨llsjo¨) and a southern massif (Fig. 2). Both intrude older Svecofennian
granitoids and metasupracrustal rocks.
The southern of these massifs is divided into an eastern and a western part (Gorbatschev et al., 1997) by a boundary more or less coinciding with the extension of the Hassela shear zone (HSZ), while the western margin is affected by the Storsjo¨n – Edsbyn deformation zone (SEDZ, Bergman and Sjo¨stro¨m, 1994) (Fig. 2). Along these zones, the granitoid has been affected by ductile deformation that has been overprinted partly by brittle structures.
The composition of the Revsund granitoids ranges regionally from granite sensu stricto to granodiorite with a metaluminous affinity (Gor-batschev, 1990; Claesson and Lundqvist, 1995). In the northern Fja¨llsjo¨ massif, there are also varieties with quartz-monzodioritic composition (Persson, 1978). The colour varies successively from pale grey to red and a typical feature is a coarse porphyritic texture (cf. Fig. 4a). In contrast to other areas, the Revsund granitoids in the southern massif of Ja¨mtland are associated with
K.Ho¨gdahl,H.Sjo¨stro¨m/Precambrian Research105 (2001) 37 – 56 40
minor pegmatites and fine-grained uranium- and thorium-rich granitic dykes (Ho¨gdahl et al., 1998).
2. Regional deformation zones: the
Storsjo¨n – Edsbyn deformation zone and the Hassela shear zone
On aeromagnetic anomaly maps, the SEDZ (Bergman and Sjo¨stro¨m, 1994) appears as a 10 – 20-km wide and 300-km long zone between Eds-byn in the south and Lake Storsjo¨n in the north (Fig. 2). There is independent evidence that the SEDZ continues to the north-northwest beneath the Caledonides (Bergman and Sjo¨stro¨m, 1994). The SEDZ essentially separates \1.8 Ga rocks (mainly Ljusdal and Revsund granitoids and older Svecofennian supracrustal rocks) to the east from a younger, ca. 1.7 Ga TIB intrusion to the west. Structural data and the recognition of various kinds of mylonites show that the SEDZ has been active repeatedly (Bergman and Sjo¨stro¨m, 1994). Ductile, retrograde and brittle – ductile mylonites are the most common types along this deforma-tion zone. Dextral, transpressive shearing result-ing in steep stretchresult-ing lineations, has been suggested either to be connected to the emplace-ment of the Ra¨tan intrusion, or a post-emplace-ment phenomenon. In the Revsund granitoid affected by SEDZ-deformation, a coarse, dextral C/S-fabric is the dominating structure. Sinistral shear zones occur as conjugate sets, or occasion-ally as later overprinting structures.
The HSZ is localised along the boundary be-tween the Ljusdal Batholith and the older Sve-cofennian metasedimentary rocks (greywacke-schist) of the Bothnian Basin to the north (Bergman and Sjo¨stro¨m, 1994; Sjo¨stro¨m and Bergman, 1996; Fig. 2). Previously, this boundary has been referred to as a primary fea-ture. It is a steep, WNW- to NW-striking domi-nantly dextral shear zone formed under wrench conditions. Narrow sinistral zones, conjugate to the dextral pattern, were probably formed during progressive bulk dextral deformation. Some of the former display a retrograde character reflecting shear during a late stage of the deformation.
The dextral rotation of the HSZ into the SEDZ indicates either that the HSZ is older, or that they formed simultaneously (Bergman and Sjo¨stro¨m, 1994). The timing of the main ductile deformation along HSZ is bracketed by its imprint on the 1.85 – 1.84 Ga Ljusdal Batholith (Delin, 1993; Welin et al., 1993), and its syn-metamorphic rela-tion to the regional low-pressure metamorphism (LPM) at \ca. 1.82 Ga (Claesson and Lundqvist, 1995).
3. Local shear zone network
In the area where the HSZ joins the SEDZ, there are several local, ductile deformation zones dividing the Revsund granitoid into internally un-deformed, or only slightly deformed lenses (Figs. 2 and 3). This pattern is apparent on various scales (regional to outcrop), and the deformation zones are characterised by a coarse C/S-fabric. The existence of anastomosing faults (Lun-dega˚rdh et al., 1984) coinciding partly with ductile shear zones, indicates that brittle reactivation was favoured by ductile structures.
The deformation zone along River Forsaa˚n at the southern end of Lake Locknesjo¨n is the most pervasively deformed of the local shear zones (Fig. 3). It will be described in detail below, after a brief summary of the characteristic features of other zones in the area.
dif-K.Ho¨gdahl,H.Sjo¨stro¨m/Precambrian Research105 (2001) 37 – 56 42
Fig. 3. Geological map of the western part of the Southern Massif in Fig. 2 showing the shear zones described in the paper (modified after Lundega˚rdh et al., 1984; Lundqvist, 1996; Lundqvist and Korja, 1997; Sturkell et al., 1998), (1), dextral, west-side-up shear zone at the boundary between metasedimentary rocks and the Revsund granitoid; (2), discontinuous mylonite zone; (3), Forsaa˚n zone; and (4) a belt of anastomosing deformations zones dominated by L\S-mylonites.
ferent lithologies have been juxtaposed, the my-lonites are typically banded. Such a mylonite, affecting early Svecofennian rocks in the central part of the deformation zone, has been dated at 180292 Ma (Ho¨gdahl et al., 1996). A prominent shear zone bounds the deformation zone to the east. In its northern part, this shear zone separates
the Revsund granitoid (to the east) from metased-imentary rocks.
personal communication, 1997). Within the coarse Revsund granitoid east of Lake Bo¨rjesjo¨n (Fig. 3(1)) steep, metre-wide ultramylonites occur, which strike ca. 330° and have a strong, oblique stretching lineation plunging ca. 45° to the south-east. Kinematic indicators (shear bands, rotation of gneissosity) verify oblique dextral- and south-west-side-up displacement.
Such kinematic conditions have been recorded also in intensely deformed, very planar, steeply dipping metagreywackes ca. 200 m from the con-tact to the granitoid. In this case, the kinematics are verified by asymmetric boudinage and shear bands (C%), combined with a pronounced stretch-ing lineation plungstretch-ing ca. 45° to the southeast. A weak, shallow plunging lineation is developed lo-cally on platy quartz. The differences in plunge between the lineations may either be the result of two distinct episodes of shearing or represent two stages of progressive shearing in a transpressive environment (cf. Tikoff and Teyssier, 1994).
3.1. The Forsaa˚n shear zone
The deformation zone along Forsaa˚n (Fig. 3, (3)) was first recorded by Ho¨gbom (1894) and interpreted as a penetrative foliation in the Rev-sund granitoid. It was later re-interpreted as an enclave dominated by felsic metavolcanic rocks with a minor proportion of early-orogenic grani-toids and some lenses of Revsund granitoid (Lun-dega˚rdh et al., 1984; Gorbatschev et al., 1997). The re-interpretation was based on the assump-tion that ductile structures should not exist in rocks classified as post-orogenic. Our study sup-ports Ho¨gbom%s interpretation, i.e. the protolith is a coarsely porphyritic granitoid (cf. Fig. 4A and B). Although strain variations exist, the Forsaa˚n section is very homogeneous compositionally, lacking lithological variations.
The deformation zone is ca. 1-km wide and can be traced for more than 10 km in a NNW – SSE to NW – SE direction. The deformational fabric is more or less continuous through the width of the zone and the boundaries to undeformed rocks are distinct (Fig. 4B).
Towards the northwest, the deformation zone follows Lake Locknesjo¨n where it affects early
Svecofennian metavolcanic rocks (Mansfeld et al., 1998). This part is recorded on Bouguer anomaly maps as a local gravity low (Sturkell et al., 1998). Phanerozoic rocks and Caledonian thrust sheets cover the continuation farther northwest.
The dominating structure in the granitoid along the deformation zone at Forsaa˚n is a coarse, penetrative, subvertical C/S-fabric (Fig. 4C), sometimes grading into pervasively deformed gneiss zones without C/S-fabric (Fig. 4D). With increasing strain, the C/S-fabric is transformed into millimetre- to metre-wide mylonites. Locally, the C/S-fabric is cut by mylonites indicating that the latter are slightly younger. Subordinate ultra-mylonites have been recorded, in which the folia-tion is defined mainly by platy quartz (i.e. recrystallised ribbons) and thin mica-rich bands.
The bulk sense of shear is not obvious in the steep, NE-dipping Forsaa˚n shear zone and kine-matic indicators are contradictory. Dextral shear zones truncating a sinistral C/S-fabric exist (Fig. 4E), as well as sinistral shear zones truncating a pervasive gneissic foliation, or tensile quartz veins indicating sinistral rotation (Fig. 4F, Fig. 5A and D). Altogether, these examples indicate a sequen-tial formation of sinistral and dextral kinematic patterns.
In pervasively deformed gneissic parts, there is, at least locally, a faint asymmetric pattern indicat-ing dextral sense of shear in sub-horizontal sec-tions. Still, sinistral shear bands and minor shear zones dominate among the data collected (Fig. 5A). However, more important is that dextral and sinistral shear bands (C%) and minor shear zones are symmetrically arranged with respect to the pervasive, partly mylonitic (C) foliation (Fig. 4E and Fig. 5A).
K.Ho¨gdahl,H.Sjo¨stro¨m/Precambrian Research105 (2001) 37 – 56 44
The orientations of strain axes, indicated by two, conjugate dextral and sinistral shear zones, are NNW-trending (close to horizontal) X-axes (stretching), WSW-trending (close to horizontal) Z-axes (shortening) and steepY-axes (Fig. 5C).X and Y thus plot in the fields of stretching- and constructed intersection lineations, respectively, and Zclose to the field of poles to the mylonitic foliation, i.e. Z is more or less perpendicular to that foliation (Fig. 5B and C).
In terms of strain axes, the distribution of stretching and intersection lineations indicates thatXand Yrotate within the shear plane, while Z is less variable and more or less orthogonal to that plane. Apparently, pure shear predominated during the development of the deformation zone and resulted in the development of the S\L tectonites reflected by the poorly developed stretching lineations. Such conditions would also explain a sequential development of minor shear
K.Ho¨gdahl,H.Sjo¨stro¨m/Precambrian Research105 (2001) 37 – 56 46
zones and the symmetric kinematic pattern shown in Fig. 5A – C. An important inference is that the kinematic data from the Forsaa˚n section (within the granitoid) deviate considerably from those recorded along the margin of the granitoid. This difference is significant for the interpretation that strain was partitioned.
Some narrow, sinistral shear zones truncating the pervasive shear fabric, have a sub-horizontal lin-eation and tend to indicate lower magnetic suscep-tibility values than the gneissic zones. These shear zones may, therefore, represent a later phase of deformation under more oxidised conditions. Fur-ther diagnostic patterns will be described in the following section.
Late brittle deformation is widespread in the area. It resulted in the formation of cataclasites in zones ranging from millimetres to several centime-tres in width, occasionally accompanied by pseudo-tachylite melts. Revsund granitoids that are deformed by brittle deformation have typically a deep red colour. Open cavities sporadically host small crystals of quartz, locally together with epidote and fluorite and in places calcite.
3.1.1. Microstructures in the Forsaa˚n shear zone Microcline porphyroclasts in the pervasive C/ S-fabric at Forsaa˚n locally show rather coarse core-and-mantle structures (Fig. 6A), formed by dynamic recrystallisation along the margins of the microcline megacrysts. Such fabrics are common in feldspars affected by deformation at temperatures of 400 – 500°C (Passchier and Trouw, 1996), i.e. low- to medium-grade conditions. This tempera-ture range is also indicated by other feldspar microstructures, e.g. the porphyroclasts generally lack fracturing and micro-kink-bands typical of low-grade conditions (300 – 400°C). They also con-tain flame perthites typical of low- to medium-grade conditions (400 – 500°C), while myrmekites characteristic of medium- to high-grade conditions are absent (Passchier and Trouw, 1996). These temperature estimates are, thus, comparable to the 500°C suggested for the development of C/S-fabric in granites affected by shearing (Gapais, 1989).
In more deformed parts, the porphyroclasts are polygonised entirely with well-developed triple points between the crystals (Fig. 6B). Bands of
epidote with allanite cores wrap around some polygonised augen. Coexisting epidote and acces-sory amounts of chlorite and muscovite probably reflects saussuritisation of plagioclase; most chlor-ite appears to be a late replacement.
The existence of titanites along C and C%shows that they are part of the deformational fabric (Fig. 6C). They occur preferably in biotite and in the saussuritised bands, minor amounts are found at quartz- and feldspar grain boundaries. Conse-quently, the titanites used for dating are part of a thoroughly recrystallised fabric showing some neomineralisation. Both the high degree of recrys-tallisation and the indicated temperature range of ca. 400 – 500°C are fundamental conditions when the relationship between the obtained age and the evolution of the shear zone is considered (see below).
The narrow sinistral shear zones that truncate the pervasive core-and-mantle fabrics and are char-acterised by intense grain-size reduction. Biotite (partly chloritised), chlorite and dynamically re-crystallised quartz define a C/S-fabric; kinked and bent muscovite occurs locally (Fig. 6D). Large quartz grains show elongate subgrains and contain bands that are recrystallised dynamically. There is evidence of both grain-boundary migration recrys-tallisation (most common) by interlobate, highly irregular grain boundaries and subgrain rotation recrystallisation. The latter is displayed by the existence of small grains with slightly different optical orientation along grain boundaries. These recrystallisation mechanisms indicate temperatures probably exceeding ca. 400°C. The microstructures and the mineralogy indicate somewhat lower tem-peratures compared with that of the pervasive pattern. The microstructures of the sinistral zones also appear less evolved, and combined with the recognition in the field that they may be more oxidised (lower magnetic susceptibility values) than the gneissic zones, support the inference of a late development.
3.2. Microstructures along the eastern margin of
the granitoid
Fig. 6. Microstructures from the Forsaa˚n section and the sheared margin to the metagreywacke in the east. (A) In the C/S-fabric, K-feldspar porphyroclasts (one in the central part) are surrounded partly by a mantle of dynamically recrystallised K-feldspar (to the left of the large grain). Forsaa˚n shear zone. Length of photograph corresponds to 5.8 mm. X nicols. (B) With higher strain (cf. Fig. 4D) the K-feldspar porphyroclast are often completely polygonised. Locally well-developed triple points exist between newly formed crystals. Forsaa˚n zone. Length of photograph corresponds to 5.8 mm. X nicols. (C) Titanites occur along C- (subhorizontal) and C%- planes (SW – NE) preferably in biotite and bands of saussurite that occasionally wrap around quartz and feldspar. Bt, biotite; Ti, titanite; Sau, saussuritised plagioclase. The sample is from a high strain gneiss zone within the Forsaa˚n zone. Length of photograph corresponds to 3 mm. (D) The fabric in thin sinistral shear zones along Forsaa˚n is defined by partly chloritised biotite, chlorite and dynamically recrystallised quartz. Length of photograph corresponds to 3 mm. X nicols. (E) In intensely mylonitised granitoid,s-porphyroclasts (recrystallised to asymmetric ribbons in the upper central part of the picture) and C%at low angle to the
K.Ho¨gdahl,H.Sjo¨stro¨m/Precambrian Research105 (2001) 37 – 56 48
Fig. 3, (1)) have a pronounced foliation made up by alternating platy quartz and bands containing fine-grained, brown biotite. The quartz plates have internal C/S-fabrics defined by optically ori-ented, dynamically recrystallised quartz crystals, i.e. microstructures typical of dislocation creep. Kinematic indicators are fairly common, includ-ing composite asymmetric augen, s-porphyro-clasts and locally existing C%at a low angle to the mylonitic foliation (C, Fig. 6E). K-feldspar por-phyroclasts generally show intense (static?) sericitisation. An important condition is that the ultramylonites lack chlorite and epidote minerals, i.e. exhibit no records of retrograde, low tempera-ture deformation.
In the metagreywackes ca. 200 m from the contact to the granitoid, a pseudoclastic texture is still visible in spite of strong shearing. The defor-mational fabric contains dynamically recrys-tallised quartz plates in which the crystals indicate grain boundary migration recrystallisation ( cf. above). Fine-grained, brown biotite dominates over white mica, but larger muscovite fish with internal strain (undulose extinction and rare kink-bands) exist (Fig. 6F).
The microstructures in the adjacent metasedi-mentary rocks are, thus, comparable with those in the mylonitised granitoid, suggesting that they formed simultaneously along a major shear zone. The comparatively low metamorphic grade of the greywacke close to the intrusion, including a pseu-doclastic texture, indicates that the greywacke has been juxtaposed by post-emplacement shearing.
4. U – Pb geochronology
4.1. Analytical method
A sample of the recrystallised C/S-mylonite from the Forsaa˚n shear zone (Swedish national grid, 697661/145656) was collected for U – Pb ti-tanite thermal ionisation mass spectrometry (TIMS) and zircon secondary ion mass spec-trometry (SIMS) analyses to investigate both the timing of shearing and the protolith age. As a reference sample, the undeformed protolith was also collected for U – Pb titanite analysis. The
sampling site of the latter is ca. 1 km west of the Forsaa˚n zone (Swedish national grid, 697210/
145676) within one of the type areas for Revsund granitoids as defined by Ho¨gbom (1894) (Fig. 3). The majority of the titanites from the C/ S-my-lonite are smaller than 100mm and oblate shaped. Due to crushing and milling, many of the crystals were fragmented. Approximately 30 transparent crystals and fragments, free from inclusions and cracks and with a total weight of 460 mg were selected for U – Pb analyses. They were dissolved in HF:HNO3 in a Savillex
®
beaker on a hot plate for ca. 70 h. The sample was evaporated and HCl added before it was aliquoted. One aliquot was spiked with a208Pb –233U –235U tracer, and U and Pb were separated using a HBr and HNO3 ion exchange technique.
The titanites from the undeformed granitoid are distinctly different from those of the mylonite. They are generally larger (200 – 250mm) and euhe-dral with well-developed crystal faces with sharp edges. Due to their large size, the crystals are almost opaque and to avoid problems with hidden impurities and inhomogeneous parts, only clear and inclusion free, ca. 150 mm large fragments ( ca. 15 corresponding to ca. 1 mg) were selected for analysis. In this case, they were dissolved in an autoclave in 205°C for 50 h and a 205Pb –233U – 235U spike were used as a tracer, while the rest of
the procedure was as above.
The U – Pb analyses were carried out on a Finnigan MAT 261 solid source mass spectrome-ter at the Swedish Museum of Natural History in Stockholm. Corrected isotope values, U/Pb, Pb/
Pb ratios and intercept ages were calculated using the program by Ludwig (1993, 1995). The initial lead correction was made according to Stacey and Kramers (1975) and the decay constants recom-mended by Steiger and Ja¨ger (1977) were used.
Fig. 7. (A) U – Pb concordia diagram from TIMS titanite analyses showing that the undeformed granitoid and the For-saa˚n shear zone yield ages of 185292 and 181692 Ma, respectively. The titanites from the former are \250mm and
euhedral with sharp crystal faces, whereas the latter areB100
mm and oblate shaped. Microstructures indicate that the
titan-ites from the Forsaa˚n zone grew during deformation and the age is therefore interpreted to represent the ductile shearing. Size of the error ellipsoid and uncertainties are given with 95% confidence. Isotopic data are presented in Table 1. (B) U – Pb concordia diagram for SIMS zircon data from the Forsaa˚n zone. A regression through ten analyses in the central part of the crystals yields an age of 1849914 Ma. The four shaded ellipses show analyses at the margin of the zircons and have been omitted from the regression. The relative probability plot (inserted diagram) shows a slight uneven distribution, indicat-ing a younger overprintindicat-ing/lead loss. Sizes of the error ellip-soids are given in 1s and uncertainties with 95% confidence.
Isotopic data are presented in Table 1.
The zircons selected for U – Pb SIMS analyses were mounted in transparent epoxy resin together with chips of reference zircon 91500 with an age of 1065 Ma (Wiedenbeck et al., 1995). The sample was polished to reveal as much of the mounted zircons as possible and coated with ca. 25 nm of gold. The analyses were performed at the Swedish Museum of Natural History, Stockholm, using the NORD-SIM Cameca IMS 1270 ion microprobe. A 4 nA O2− primary beam producing an ellipsoid analysis spot size of approximately 25mm was used, and a single electron multiplier in an ion counting mode measuring following masses, 90
Zr2 16
O (196), 204 Pb (204), background (204.2),206
Pb (206),207
Pb (207), 208Pb (208), 238U (238), 232Th16O (248), 238U16O
2 (270). Detailed descriptions of the analytical and calibration procedures are given by Whitehouse et al. (1997, 1999). Corrected isotope values, using modern common lead composition (Stacey and Kramers, 1975) and measured 204Pb, U/Pb and Pb/Pb ratios, and intercept ages were calculated using the Isoplot/Ex ver. 2.00 (Ludwig, 1999). The ages are calculated using the decay constants rec-ommended by Steiger and Ja¨ger (1977). For the regression, different line fitting models are recom-mended by Ludwig (1999). In this case, Model 1 has been used as the MSWD (mean square of weighted deviates) value is relatively low. This means that the scatter is assigned to analytical errors and error correlation only.
4.2. Results
The titanites from the C/S-mylonite in the For-saa˚n shear zone yield an almost concordant result of 181692 Ma (Fig. 7A). The best estimate of the age is given by the207Pb –206Pb data on which the diminutive discordance has an insignificant influ-ence.
The age is interpreted to be that of the deforma-tion. There are two specific reasons for this inter-pretation, which are both based on the microstructures recorded.
K.Ho¨gdahl,H.Sjo¨stro¨m/Precambrian Research105 (2001) 37 – 56 50
2. The microstructures indicate that deformation took place below the 600 – 700°C closure tem-perature for Pb diffusion in titanite of this size (Scott and St-Onge, 1995), which means that no substantial diffusional Pb-loss occurred af-ter the titanite formation.
The titanites from the undeformed reference sample of the granitoid yield a considerably higher age than the titanites from the mylonite. The U – Pb analysis is slightly discordant with a 207
Pb/206
Pb age of 185192 Ma (Fig. 7A). Cathodoluminescence images on zircon from the Forsaa˚n C/S-mylonite show a large variation of internal structures including inherited cores. The latter were not analysed and no systematic age variations with respect to different cathodolu-minescence images were found among the others. However, analyses (four points) made close to crystal edges tend to yield somewhat lower ages than analyses from the central parts. This trend is visible in the asymmetrical relative probability plot (Fig. 7B).
Regression through ten points analysed in cen-tral parts of the crystals yield an age of 1849914 Ma with an MSWD value of 0.58 (Fig. 7B). Some spots are reversed discordant, a phenomenon, which is not uncommon in analyses made by the SIMS technique. This could be an instrumental artefact, but micro-scale heterogeneities, with lead gain or uranium loss in the analysed part of the crystals, cannot be excluded.
The zircon age overlaps with the titanite age of the undeformed granitoid. It is interpreted to reflect the minimum magmatic age of the pro-tolith, since some lead loss could have occurred during the subsequent shearing and later reactiva-tions. These events are most likely the reason for the younger overprinting/lead loss in the outer parts of the crystals.
5. Discussion
5.1. Regional implications of geochronological and
structural data
The magmatic age of the granitoids studied, i.e. in the western part of the southern Revsund
mas-sif in the Ja¨mtland county, is constrained to ca. 1850 Ma by independent and overlapping zircon and titanite analyses. In spite of being collected from a C/S-mylonite, the zircons partly show typical magmatic zonation patterns and the titan-ites from the reference sample are part of an ‘isotropic’ magmatic fabric. The titanite age is also in accordance with an 185899 Ma U – Pb SIMS result on zircon from the same undeformed granitoid, and supported by 185498 and 18599
11 Ma obtained from other K-feldspar megacryst bearing granitoids in the vicinity (Ho¨gdahl and Sjo¨stro¨m, 2000).
Blanquat et al., 1998). Consequently, the trans-pressive conditions at 1816 Ma cannot be applied to describe the syn-magmatic evolution of the ca. 1850 Ma granitoid.
On a regional scale, the Revsund granitoids have been interpreted to truncate the structures in the older, pervasively deformed gneisses (Stephens et al., 1994). This pattern has also been observed in the western part of the southern massif (Gor-batschev et al., 1997). Apparently, much of the regional deformation must have preceded ca. 1.85 Ga. Consequently, the peak orogenic deformation in the investigated area appears to be earlier than the generally assumed interval of ca. 1.85 – 1.80 Ga (e.g. Stephens et al., 1997). However, the inferred age of the regional low-pressure meta-morphism (\ ca. 1.82 Ga, Claesson and Lundqvist, 1995) and the approximately coeval Forsaa˚n shear zone would represent a second tectonometamorphic episode.
The 1849914 Ma age of the protolith in the Forsaa˚n C/S-mylonite and the 185192 Ma age of the reference sample are both anomalous com-pared with the previously dated 1.80 – 1.77 Ga age range of Revsund granitoids. As this time interval is based on a few precise age determinations scattered over ca. 6000 km2
, our results may reflect that the emplacement period for the suite was more extended than assumed generally. How-ever, the unexpected ages add to other anomalous features of the southern Revsund massif; the un-typical association of pegmatites and U- and Th-rich dykes and a deviating geochemical signature compared with ‘normal’ Revsund granitoids (M. Ahl, Stockholm, personal communication, 1997; Gorbatschev et al., 1997). Altogether these anomalies challenge the interpretation that this part of the massif consists of Revsund granitoid according to the presently applied definition, in spite of its location within a type area originally defined by Ho¨gbom (1894).
Compared with other granitic rocks in the re-gion, our results partly overlap with the 1.85 – 1.84-Ga age of the granitoids within and to the west of the Ljusdal Batholith (Fig. 2; Delin, 1993; Welin et al., 1993; Delin and Persson, 1999). Contrary to the rocks in the southern Revsund massif in the Ja¨mtland county, the Ljusdal
Batholith is foliated and folded in most parts. The structural difference between the approximately 1850 Ma ‘Revsund granitoid’ and the Ljusdal granitoid may reflect tectonic histories substan-tially different. If so, the major Svecokarelian deformation pre-dates ca. 1.85 Ga north of the Ljusdal Batholith and post-dates this age within the batholith, and consequently, an important domain boundary would exist between these ig-neous provinces. This possible difference in timing of the Svecokarelian evolution may be similar to the situation in southern Finland with a metamor-phic peak at 1.89 – 1.88 Ga and a subsequent metamorphic event at 1.84 – 1.83 Ga in the south (Nironen, 1997; Korsman et al., 1997).
The U – Pb titanite result from the Forsaa˚n C/S-mylonite is in accordance with inferred early ( ca. 1.85 – 1.80 Ga) deformation along the SEDZ and HSZ (Bergman and Sjo¨stro¨m, 1994). It over-laps partly with the age of other deformation zones occurring in the region. A shear zone ca. 10 km to the SW (Fig. 34) has been dated at 18029
2 Ma (Ho¨gdahl et al., 1996) and titanites from the Hagsta and Ljusne high-T shear zones ca. 400 km to the SSE have been dated at 180996 and 179892 Ma, respectively (Ho¨gdahl et al., 1995; Ho¨gdahl and Sjo¨stro¨m, 1999). These data fall in the same age range as U – Pb columbite – tantalite 1807 – 1803 Ma ages for late-kinematic pegmatites in southern Finland, which intruded during trans-pressive, semi-ductile conditions (Lindroos et al., 1996). This transpressive zone is, tentatively, the eastern continuation of the HSZ, which has re-cently been suggested to be a coherent structure across the Fennoscandian shield (Sjo¨stro¨m et al., 2000) based on structural (Ehlers et al., 1993; Bergman and Sjo¨stro¨m, 1994; Lindroos et al., 1996; Sta˚lfors and Ehlers, 2000), geochronological (Lindroos et al., 1996; Ho¨gdahl and Sjo¨stro¨m, 1999) and geophysical data (Korhonen et al., 1999).
K.Ho¨gdahl,H.Sjo¨stro¨m/Precambrian Research105 (2001) 37 – 56 52
internally in the Forsaa˚n section (dominantly pure shear conditions) demonstrates partitioning of strain (Fig. 4F) as well as the variation in kine-matic character of adjacent deformation zones. Most, if not all, described zones have a compo-nent of orthogonal pure shear indicating trans-pressive deformation.
In the case of the Forsaa˚n and related zones, the shearing post-dates the magma emplacement. However, in a regional perspective, the shear zone activity dated at 1816 – 1798 Ma temporally over-laps with the 1.82 – 1.80 Ga emplacements of S-type granites (Claesson and Lundqvist, 1995) and lithium – caesium – tantalum (LCT) pegmatites (Romer and Smeds, 1994, 1997), and partly with the previously established age range of the 1.80 – 1.77 Ga Revsund granitoids in the central part of the Svecofennian domain.
6. Conclusions
Two crustal-scale shear zone systems, SEDZ and the HSZ, truncate the investigated granitoid of the southern Revsund massif in Ja¨mtland county, central Sweden. They form an anastomos-ing ca. 50-km wide pattern of steep, ductile shear zones and shear pods of various scales.
The age of a ‘Revsund’ granitoid located in the western part of the Southern massif is ca. 1850 Ma, constrained by U – Pb SIMS analyses on zir-cons (1849914 Ma) from a sheared granitoid and by U – Pb TIMS analysis on titanite (185192 Ma) from an undeformed granitoid. As these samples were collected within a type area origi-nally referred to as Revsund granitoids, either the previously known age range ( ca. 1.80 – 1.77 Ga) or the definition of these granitoids as part of the Revsund suite must be revised (Table 1).
The granitoid in the southern massif is mainly isotropic outside shear zones indicating that the regional, Svecokarelian deformational pattern, truncated by the granitoid, have formed earlier than 1.85 Ga in the area. This is supported by the condition that titanites from the isotropic grani-toid give the magmatic and not the metamorphic age. By contrast, the more pervasively deformed ca. 1.85 – 1.84 Ga Ljusdal granitoid to the south
indicates that major deformation in that area post-dates 1.85 Ga.
In the Forsaa˚n section, the deformational (pre-dominantly C/S-) fabric is arrested at ca. 400 – 500°C, i.e. at sub-solidus conditions. Titanite is found in C and C%planes and yields an U – Pb age of 181692 Ma. This age is considerably younger than the U – Pb titanite age of the granitoid out-side the shear zone, and is interpreted to reflect the age of the deformation. The ductile pattern has partly been overprinted by later deformation at brittle conditions, which could explain the lead loss and, thus, younger ages found in the outer parts of the zircons. The local occurrence of a magmatic foliations in the surrounding granitoids indicates that also ca. 1.85 Ga syn-emplacement deformation took place.
Kinematic data from the investigated local shear zones indicate strain partitioning, resulting in dominantly dextral simple shear in the mar-ginal parts of the granitoid massif and pure shear conditions in the internal part (Forsaa˚n section). The integrated picture indicates that the deforma-tion was transpressive.
The Forsaa˚n and associated shear zones are components of the late orogenic, crustal scale, ductile shear zone system of central Sweden. The Forsaa˚n zone is the oldest hitherto dated ductile shear zone in the region, that together with other data define periods of shearing in the time interval 1816 – 1798 Ma. This interval overlaps with activ-ity along structures kinematically similar in south-ern Finland. Possibly, the shear zones described here are part of a coherent structure extending across the Fennoscandian shield, from central Sweden to southern Finland.
Acknowledgements
Pro-K
number (ppm) (%) (%) measured age (Ma)
SIMS,zircon Forsaa˚n zone
5.129 1.85 0.3290 1.75 0.95 11 836 0.16 1849 11
0.60 1841
111 16 1834 28
1 Centre 54 45 0.1131 0.61
228 129 96 0.1128 5.104 1.87 0.3282 1.77 0.95 21 395 0.09 1844 11 1837 16 1830 28 2 Centre
5.699 1.80 0.3589 1.75 0.97 16 353 0.11 1883 7 1931 16 1977 30 3 Centre 348 173 155 0.1152 0.40
4.695 1.97 0.2943 1.69 0.85 98 135 0.02 1891 18
1.03 1840
43 0.1157 17 1821 25
5 Centre 115 48
0.71
247 104 90 0.1147 4.611 1.76 0.2915 1.61 0.91 13 512 0.14 1875 13 1824 15 1806 23 7 Centre
5.128 1.73 0.3204 1.61 0.93 12 132 0.15 1897 11
8 Centre 233 82 91 0.1161 0.63 1915 15 1960 25 4.902 1.91 0.3064 1.65 0.87 5297 0.35 1896 17
0.96 1877
262 16 1885 25
9 Centre 103 100 0.1161 1.13
144 76 55 0.1142 4.722 1.96 0.2998 1.60 0.82 17 721 0.11 1868 20 1845 17 1850 24 10 Centre
0.1148 0.94 4.963 1.83 0.3135 1.57 0.86 35 063 0.05 1877 17 1887 16 1923 24 143
11 Centre 237 96
4.683 1.73 0.2984 1.57 0.91 13 796 0.14 1861 13
0.72 1838
203 15 1843 23
13 Centre 122 79 0.1138 0.37
257 119 101 0.1111 4.887 1.79 0.3190 1.75 0.98 9681 0.19 1817 7 1800 15 1785 27 4 Edgea
1.05
145 63 55 0.1114 4.622 2.08 0.3009 1.79 0.86 14 932 0.13 1822 19 1826 17 1856 27 6 Edgea
4.356 2.07 0.2841 1.88 0.91 4699 0.40 1819 16
0.87 1777
12 Edgea 442 1435 159 0.1112 17 1765 27
4.594 1.88 0.2947 1.58 0.84 10 908 0.17 1849 18 1821 16 1823 23 14 Edgea 212 139 79 0.1131 1.02
0.1 4.951 0.6 0.3235 0.6 0.99 1556
Forsaa˚n 0.0464 127 47 0.11102 0.7 1816 0.1 1811 0.6 1807 0.6 zone
0.11317 0.1 5.157 0.8 0.3305 0.8 0.99 3234 0.9 1851 0.1 1856 0.8 1841
166 62 0.8
Undeformed 0.1061 protolith
aOmitted from the regression. bError is observed.
cIncludes external variation in standard Pb/U. dError correlation.
K.Ho¨gdahl,H.Sjo¨stro¨m/Precambrian Research105 (2001) 37 – 56 54
fessor Stefan Claesson at the Swedish Museum of Natural History, Stockholm, are thanked for use-ful discussions concerning the TIMS work. The staff at the Nordsim laboratory is greatly ac-knowledged. Dr Stefan Bergman at the Geologi-cal Survey of Sweden provided the original to Fig. 2 and is also thanked for valuable comments on an early version of the manuscript. We also thank Professor Thomas Lundqvist, Dr C. Friend and Dr C.-H. Wahlgren for constructive comments, which improved the manuscript. The study was financed by the Swedish Natural Science Research Council (NFR), contract number 04901-318 and Stockholm University (KH), and a research grant from the Geological Survey of Sweden (HS) con-tract number 03-854/93. Nordsim contribution no. 31.
References
Bergman, S., Sjo¨stro¨m, H., 1994. The Storsjo¨n-Edsbyn defor-mation zone, central Sweden. Research report, Sver. Geol. Unders., p. 46.
Beunk, F.F., Page, L.M., Wijbrans, J.R., Barling, J., 1996. Deformational, metamorphic and geochronological con-straints from the Loftahammar – Linko¨ping deformation zone (LLDZ) in SE Sweden: implications for the develop-ment of the Svecofennian Orogen. Geol. Fo¨ren. Stockholm Fo¨rh. 118 (Jubilee issue), A9.
Brown, M., Solar, G., 1998a. Shear-zone systems and melts: feedback relations and self-organization in orogenic belts. J. Struct. Geol. 20, 211 – 227.
Brown, M., Solar, G., 1998b. Granite ascent and emplacement during contractional deformation in convergent orogens. J. Struct. Geol. 20, 1365 – 1393.
Claesson, S., Lundqvist, T., 1995. Origins and ages of Protero-zoic granitoids in the Bothnian basin, central Sweden; isotopic and geochemical constraints. Lithos 36, 115 – 140. Delin, H., 1993. The radiometric age of the Ljusdal granodior-ite of central Sweden. In: Lundqvist, T., (Ed.), Radiometric Dating Results. Sver. Geol. Unders., Ser. C. 823, pp. 13 – 16.
Delin, H., 1996. U – Pb zircon ages of granodiorites in the Ka˚rbo¨le region, central Sweden. In: Lundqvist, T. (Ed.), Radiometric Dating Results 2. Sver. Geol. Unders., Ser. C. 828, pp. 6 – 14.
Delin, H., Persson, P.-O., 1999. U – Pb zircon ages of three Palaeoproterozoic igneous rocks in the Loos-Hamra area, central Sweden. In: Bergman, S. (Ed.), Radiometric dating results 4, Sver. Geol. Unders., C 831, pp. 21 – 31. Delin, H., Aaro, S., 2000. Kartbladen 17F A,nge NV och NO
(in Swedish). In: Delin, H. (Ed.), Regional
berggrundsgeol-ogisk underso¨kning sammanfattning av pa˚ga˚ende under-so¨kningar 1999 Sver. Geol. Unders., Rapporter och meddelanden nr. 102, pp. 40 – 42.
D’Lemos, R.S., Brown, M., Strachan, R.A., 1992. Granite magma generation, ascent and emplacement within a trans-pressional orogen. J. Geol. Soc. London 149, 487 – 490. Ehlers, C., Lindroos, A., Selonen, O., 1993. The late
Svecofen-nian granite – migmatite zone of southern Finland — a belt of transpressive deformation and granite emplacement. Precamb. Res. 64, 295 – 309.
Gaa´l, G., Gorbatschev, R., 1987. An outline of the Precam-brian evolution of the Baltic shield. Precamb. Res. 35, 15 – 52.
Gapais, D., 1989. Shear structures within deformed granites: mechanical and thermal indicators. Geology 17, 1144 – 1147.
Gorbatschev, R., 1990. Granitic magmatism in the western part of the Central Svecofennian province. Abstract, Sec-ond Symposium. Baltic shield with Workshop on Correla-tion with Laurentia, Lund, Sweden, pp. 39 – 40.
Gorbatschev, R., Bogdanova, S., 1993. Frontiers in the Baltic shield. Precamb. Res. 64, 3 – 21.
Gorbatschev, R., Kornfa¨lt, K.-A., Lundega˚rdh, P.H., 1997. Beskrivning till berggrundskartan o¨ver Ja¨mtland la¨n, del 1: urberget. Sver. Geol. Unders. Ser. Ca 53 (1), 250 In Swedish with English Abstract.
Ho¨gbom, A., 1894. Geologisk beskrifning o¨fver Jemtlands La¨n. Sver. Geol. Unders., Ser. C 140, 106 In Swedish. Ho¨gdahl, K., Sjo¨stro¨m, H., 1999. Timing and kinematics of
paleoproterozoic shear zones in central Sweden. J. Conf. Abs. EUG 10 Abs. 4, 828.
Ho¨gdahl, K., Sjo¨stro¨m, H., 2000. Timing of emplacement and shearing of a post-orogenic Revsund (?) granitoid in cen-tral Sweden. In: Eide, E. (Ed.), Abstracts, The 24 Nordiske Geologiske Vintermøte, p. 153.
Ho¨gdahl, K., Sjo¨stro¨m, H., Bergman, S., 1995. Geochronol-ogy and tectonic evolution of the Hagsta gneiss zone, east central Sweden. In: Kohonen, T., Lindberg, B. (Eds.), Abstracts, The Twenty-Second Nordic Geological Winter Meeting, p. 73.
Ho¨gdahl, K., Gromet, P.L., Claesson, S., 1996. The geochronology of a ductile shear zone in Ja¨mtland, central Sweden. Geol. Fo¨ren. Stockholm Fo¨rh. 118 (Jubilee issue), A14.
Ho¨gdahl, K., Lundqvist, L., Sjo¨stro¨m, H., 1998. Major shear deformation in the post-orogenic Revsund granite in Ja¨mt-land, central Sweden. In: Wilson, R.J. (Ed.), Abstract volume 23. Nordiske Geologiske Vintermøde, p. 125. Hutton, D.H.W., 1988. Granite emplacement mechanisms and
tectonic control: inferences from deformation studies. Transactions of the Royal Society of Edinburgh. Earth Sci. 79, 245 – 255.
Ka¨rki, A., Laajoki, K., Luukas, J., 1993. Major paleoprotero-zoic shear zones of the Fennoscandian shield. Precamb. Res. 64, 207 – 223.
J.-R., Solheim, D., Chepik, A., Kulinich, A., Zhdanova, L., Vaher, R., All, T., Silvee, H., 1999. Preliminary mag-netic and gravity anomaly maps of the Fennoscandian shield 1:10 000 000. Geol. Surv. Finland 27, 173 – 179 Spe-cial Paper.
Korja, A., Heikkinen, P.J., 1995. Proterozoic extensional tec-tonics of the central Fennoscandian shield: results from the Baltic and Bothnian Echoes from the Lithosphere experi-ment. Tectonics 14, 503 – 517.
Korsman, K., Koistinen, T., Kohonen, J., Wennerstro¨m, M., Ekdahl, E., Honkamo, M., Idman, H., Pekkala, Y., 1997. Bedrock Map of Finland, 1:1 000 000. Geological Survey of Finland, Esbo, Finland.
Lindroos, A., Romer, R.L., Ehlers, C., Alviola, R., 1996. Late-orogenic Svecofennian deformation in SW Finland constrained by pegmatite emplacement ages. Terra Nova 8, 567 – 574.
Ludwig, K.R., 1993. PBDAT: a computing program for pro-cessing Pb – U – Th isotope data. Version 1.24. USGS Open-file report 88 – 524, p. 23.
Ludwig, K.R., 1995. ISOPLOT. A plotting and regression program for radiogenic isotope data. Version 2.82. USGS Open-file report 91 – 445, p. 45.
Ludwig, K.R., 1999. Isoplot/Ex version 2.00. A geochronolog-ical toolkit for Microsoft Excel, Berkeley Geochronologi-cal Center, Special Publication No. 2, p. 46.
Lundega˚rdh, P.H., Gorbatschev, R., Kornfa¨lt, K.-A., Stro¨mberg, A., Karis, L., Zachrisson, E., Sjo¨strand, T., Skoglund, R., 1984. Bedrock geology of Ja¨mtland county, scale 1:200 000. Sver. Geol. Unders., Ser. Ca 53. Lundqvist, L., 1996. Kartbladen 18F Bra¨cke NV, NO och 19F
Ha¨ggena˚s SV, SO (in Swedish). In: Wahlgren, C.-H., (Ed.), Regional berggrundsgeologisk underso¨kning sammanfat-tning av pa˚ga˚ende underso¨kningar 1995. Sver. Geol. Un-ders., Rapporter och meddelanden nr. 84, pp. 62 – 66. Lundqvist, L., Korja, A., 1997. Kartbladen 18F Bra¨cke NV,
NO och 19F Ha¨ggena˚s SV, SO (in Swedish). In: Wahlgren, C.-H. (Ed.), Regional berggrundsgeologisk underso¨kning sammanfattning av pa˚ga˚ende underso¨kningar 1996. Sver. Geol. Unders., Rapporter och meddelanden nr. 89, pp. 27 – 34
Lundqvist, T., 1995. Bedrock of Sweden. 1:1 250 000. Na-tional Atlas of Sweden.
Mansfeld, J., Sturkell, E., Broman, C., 1998. Meteorite impact influence on a proterozoic shear zone mineralization at Gullgruvan, Ja¨mtland, Sweden. In: Wilson, R.J. (Ed.), Abstract, 23. Nordiske Geologiske Vintermøde, p. 195. Nironen, M., 1997. The Svecofennian orogen: a tectonic
model. Precamb. Res. 86, 21 – 44.
Passchier, C.W., Trouw, R.A.J., 1996. Microtectonics. Springer, Berlin, p. 289.
Patchett, P.J., Todt, W., Gorbatschev, R., 1987. Origin of continental crust of 1.9 – 1.7 Ga age: Nd isotopes in the Svecofennian orogenic terrains of Sweden. Precamb. Res. 35, 145 – 160.
Persson, L., 1978. The Revsund – So¨rvik granites in the western parts of the province of A,ngermanland central Sweden. Sver. Geol. Unders. Ser. C 741, 59.
Romer, R.L., Smeds, S.-A., 1994. Implications of U – Pb ages of columbite-tantalites from granitic pegmatites for the paleoproterozoic accretion of 1.90 – 1.85 Ga magmatic arcs to the Baltic shield. Precamb. Res. 67, 141 – 158. Romer, R.L., Smeds, S.-A., 1997. U – Pb columbite chronology
of post-kinematic paleoproterozoic pegmatites in Sweden. Precamb. Res. 82, 85 – 99.
Saint Blanquat, M., Tikoff, B., 1997. Development of mag-matic to solid-state fabrics during syntectonic emplacement of the Mono Creek granite, Sierra Nevada batholith. In: Bouchez, J.L., Hutton, D.H.W., Stephens, W.E. (Eds.), Granite: From Segregation of Melt to Emplacement Fab-rics. Kluwer Academic Publishers, Dordrecht, pp. 231 – 252.
Saint Blanquat, M., Tikoff, B., Teyssier, C., Vigneresse, J.L., 1998. Transpressional kinematics in magmatic arcs. In: Holdsworth, R.E., Strachan, R.A., Dewey, J.F. (Eds.), Continental Transpressional and Transtensional Tectonics, vol. 135. Geological Society, London, pp. 327 – 340 Special Publication.
Scott, D.J., St-Onge, M.R., 1995. Constraints on Pb closure temperature in titanite based on rocks from the Ungava orogen, Canada: Implication for U – Pb geochronology and
P–T–tpath determination. Geology 23, 1123 – 1126. Skio¨ld, T., 1988. Implications of new U – Pb zircon chronology
to early proterozoic crustal accretion in northern Sweden. Precamb. Res. 38, 147 – 164.
Sjo¨stro¨m, H., Bergman, S., 1995. Deformation zones in central Sweden — relation to the Svecokarelian orogeny. In: Kohonen, T., Lindberg, B. (Eds.), Abstracts, The Twenty-Second Nordic Geological Winter Meeting, p. 194. Sjo¨stro¨m, H., Bergman, S., 1996. Regional structures in
cen-tral Sweden. Geol. Fo¨ren. Stockholm Fo¨rh. 118 (Jubilee issue), A25.
Sjo¨stro¨m, H., Ho¨gdahl, K., Aaro, S., Bergman, S., 2000. The Hassela shear zone in central Sweden, the western part of a paleoproterozoic tectonic boundary across the Baltic shield? In: Eide, E. (Ed.), Abstracts, The 24 Nordiske Geologiske Vintermøte, p. 153.
Stacey, J., Kramers, J.D., 1975. Approximation of terrestrial lead isotope evolution by a two-stage model. Earth Planet Sci. Lett. 26, 207 – 221.
Sta˚lfors, Ehlers, C., 2000. Emplacement mechanisms of late-orogenic granites in southern Finland. In: Eide, E. (Ed.), Abstracts, The 24 Nordiske Geologiske Vintermøte, p. 164. Steiger, R.H., Ja¨ger, E., 1977. Subcommssion of geochronol-ogy: convention on the use of decay constants in geo- and cosmochronology. Earth Planet Sci. Lett. 36, 359 – 362. Stephens, M.B., Wahlgren, C.-H., 1995. Thermal and
mechan-ical responses to bilateral collision in the Svecokarelian orogen. In: Kohonen, T., Lindberg, B. (Eds.), The Twenty-Second Nordic Geological Winter Meeting, p. 203. Stephens, M.B., Wahlgren, C.-H., 1996. Post-1.85 Ga tectonic
K.Ho¨gdahl,H.Sjo¨stro¨m/Precambrian Research105 (2001) 37 – 56 56
Stephens, M.B., Wahlgren, C.-H., Weihed, P., 1994. Geologi-cal Map of Sweden. Sver. Geol. Unders., Ser. Ba no. 52. Stephens, M.B., Wahlgren, C.-H., Weihed, P., 1997. Sweden.
In: Moores, E.M., Fairbridge, R.W. (Eds.), Encyclopedia of European and Asian Geology. Chapman and Hall, London, pp. 690 – 704.
Sturkell, E.F.F., Ekelund, A., To¨rnberg, R., 1998. Gravity modelling of Lockne, a marine impact structure in Ja¨mt-land, central Sweden. Tectonophysics 296, 421 – 435. Tikoff, B., Teyssier, C., 1994. Strain modelling of
displace-ment-field partitioning in transpressional orogens. J. Struct. Geol. 16, 1575 – 1588.
Vavra, G., 1990. On the kinematics of zircon growth and its petrogenetic significance: a cathodoluminescence study. Contrib. Miner. Petrol. 106, 90 – 99.
Wahlgren, C.-H., Stephens, M.B., 1996. Polyphase tec-tonometamorphic reworking in the Transscandinavian ig-neous belt, east of Lake Va¨nern-regional tectonic implications for southern Sweden. Geol. Fo¨ren. Stockholm Fo¨rh. 118 (Jubilee issue), A27 – A28.
Whitehouse, M.J., Claesson, S., Sunde, T., Vestin, J., 1997. Ion microprobe U – Pb zircon geochronology and correla-tion of Archaean gneisses from the Lewisian Complex of
Guinard Bay, northwestern Scotland. Geochim. Cos-mochim. Acta 61, 4429 – 4438.
Whitehouse, M.J., Kamber, B.S., Moorbath, S., 1999. Age significance of U – Th – Pb zircon data from early archaean rocks of west Greenland-a reassessment based on com-bined ion-microprobe and imaging studies. Chem. Geol. 160, 201 – 224.
Wiedenbeck, M., Alle´, P., Corfu, F., Griffin, W.L., Meier, M., Oberli, F., von Quadt, A., Roddick, J.C., Spiegel, W., 1995. Three natural zircon standards for U – Th – Pb, Lu – Hf trace element and REE analysis. Geostandards Newslett. 19, 1 – 23.
Wijbrans, J.R., Beunk, F.F., Pitka, S., van Lil, R., Boeken, W.B., Wahlgren, C.-H., Stephens, M.B., Verdumen, E.A.T., 1995. Laser 40Ar/39Ar and U/Pb dating of right-lateral ductile deformation in the Svecokarelian orogen, south-central Sweden. In: Terra Abstracts, Abstract Sup-plement No. 1 to Terra Nova 7, EUG 8, p. 44.
Welin, E., Christiansson, K., Ka¨hr, A.-M., 1993. Isotope investigation of metasedimentary and igneous rocks in the paleoproterozoic Bothnian basin, central Sweden. Geol. Fo¨ren. Stockholm Fo¨rh. 115, 285 – 296.
