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 1851 9 2 Ma Fig. 7A. Cathodoluminescence images on zircon from
the Forsaa˚n CS-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 1849 9 14
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 overprintinglead 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 CS-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 1858 9 9 Ma U – Pb
SIMS result on zircon from the same undeformed granitoid, and supported by 1854 9 8 and 1859 9
11 Ma obtained from other K-feldspar megacryst bearing granitoids in the vicinity Ho¨gdahl and
Sjo¨stro¨m, 2000.
The large difference in titanite habit between the undeformed reference sample and the CS-my-
lonite shows that the titanites in the latter are newly formedrecrystallised. This difference is an
additional support for the interpretation that the 1816 9 2 Ma result is the age of the shear fabric.
The character of the microstructures shows that they were arrested at medium to low grade condi-
tions after the solidification of the intrusive. Nev- ertheless, in the southern massif, locally existing
magmatic foliations indicating syn-magmatic de- formation have been reported Gorbatschev et al.,
1997. Such structures may represent end mem- bers of continuous deformation from magmatic to
solid-state conditions, as demonstrated in the Mono Creek granite in the Sierra Nevada
batholith Saint Blanquat and Tikoff, 1997. In that granite, a magmatic fabric changes progres-
sively to a solid-state high-temperature fabric, and finally to a low-temperature fabric within a shear
zone. However, in our study, information is lack- ing on the spatial distribution and frequency of
magmatic foliations, and most important, an- isotropy of the magnetic susceptibility data
AMS necessary to decipher magmatic fabrics existing in texturally isotropic rocks, are too
scarce. Therefore, the present data is not conclu- sive for whether the deformation was continuous
or discontinuous. In addition, the transpressive conditions, defined by structural data from the
shear zones in the southern massif, were arrested in a solid-state environment. The 120 – 82 Ma
Sierra Nevada batholith shows that kinematic conditions change during the period of magma-
tism Saint Blanquat and Tikoff, 1997; Saint
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 1849 9 14 Ma age of the protolith in the
Forsaa˚n CS-mylonite and the 1851 9 2 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 km
2
, 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 CS-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 1802 9 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 1809 9 6 and
1798 9 2 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.
Also, kinematically, the Forsaa˚n shear zone fits into the regional pattern defined by the SEDZ
and the HSZ that are both characterised by a dextral, horizontal component but deviate in ver-
tical components. The difference in kinematic conditions recorded along the eastern margin of
the granitoid dextral and southwest-side-up and
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
