Abstract
We present field observations from the Palaeoproterozoic volcano-sedimentary Tampere palaeobasin, where the primary structures have been exceptionally well preserved. We use the observations to construct a new tectonic model for the southeastern margin of the Tampere basin during its inversion and subsequent closure. The observed volcano-sedimentary and structural features suggest a change in the local structural style from thick-skinned inversion to thin-skinned thrusting, to accommodate the crustal shortening during basin closure. Furthermore, it is suggested that there is a genetic relationship between the interpreted palaeothrust and the sulphide-bearing schist horizons in the study area. On a more general note, the results imply that currently subvertical mineralized shear zones may have originally been gently dipping, further suggesting that the mineralized fluids may not necessarily have been sourced from great depths (i.e. from deep within the basement).
Ancient basins archive the early development of plate tectonics, life, and ore-bearing systems. In particular, basins involving volcanic-derived materials contain economically important ore and mineral deposits (e.g. Lentz 1998; Weihed et al. 2005). The formation of these ores is often related to hydrothermal activity and the migration of mineralizing fluids; this can happen in locally extensional environments (notably the formation of the volcanogenic massive sulphide deposits, ‘VMS ores’, which most commonly form in volcanic arc settings; e.g. Ohmoto 1996; Allen et al. 2002) but mineralizing fluids can also move along strike-slip or reverse faults (e.g. Sibson et al. 1988; Teagle et al. 1990; Piessens et al. 2002). Usually the mineralizations formed in compressional settings are related to steep strike-slip or reverse faults where the fluids are derived from great depths; reported mineralizations related to more gently dipping reverse faults are less common (e.g. Nguyen et al. 1998; Piessens et al. 2002).
The Tampere Schist Belt in southern Finland is a c. 1.92–1.88 Ga volcano-sedimentary basin that underwent inversion and closure at c. 1.89–1.88 Ga (Fig. 1; Patchett & Kouvo 1986; Kähkönen et al. 1989; Nironen 1989a; Rutland et al. 2004). The tectonic setting of the belt varied with time from a microcontinental rift to a volcanic arc system (Lahtinen et al. 2009). The closure of the basin resulted in the formation of a large-scale, predominantly tight to isoclinal syncline with steeply dipping fold axial planes (Figs 1b and 2; e.g. Kähkönen 1989; Nironen 1989a). Despite the crustal-scale shortening, the primary sedimentary and volcanic structures are generally remarkably well preserved. The Tampere Schist Belt, therefore, offers an excellent opportunity to examine the evolution of an ancient basin, and the mechanics of and strain distribution during its subsequent closure. In addition, the Tampere Schist Belt is valuable for understanding the evolution of ore-bearing systems within volcano-sedimentary basins as it hosts several mineralized deposits. Probably the best known of these are the Haveri deposit, mined in 1942–1962, and the currently mined Kutemajärvi deposit (Fig. 1b). The Haveri Formation forms the base of the Tampere Schist Belt succession and hosts an orogenic gold deposit (originally a VMS ore that was later remobilized; Nironen 1994; Eilu et al. 2003). The Haveri deposit is hosted by schists and is associated with faults and folds (Nironen 1994). Various sulphide-bearing schist horizons and lenses are also contained higher in the local stratigraphy (e.g. Kähkönen & Nironen 1994). The Kutemajärvi deposit is interpreted to be an epithermal (‘shallow heat’) metamorphic deposit (Eilu et al. 2003; Kinnunen 2008). It, too, is associated with schists and prominent fault zones (Poutiainen & Grönholm 1996).
(a) Simplified geological map of the Fennoscandian shield. The Tampere Schist Belt (indicated with a rectangle) is located in a palaeobasin between the Early Svecofennian orogenic basement to the north and the contemporaneous supracrustal rocks to the south. 1, Phanerozoic sedimentary cover <0.57 Ga; 2–4, younger basement rocks 1.65–0.4 Ga; 5, Late Svecofennian granites and migmatites 1.85–1.77 Ga; 6, Early Svecofennian pre- and synorogenic magmatic rocks 1.95–1.85 Ga; 7, Early Svecofennian supracrustal gneisses and migmatites 2.0–1.85 Ga; 8, 9, pre-1.9 Ga rocks. Modified from Korsman et al. (1997). (b) Generalized geological map of the Tampere Schist Belt and the Tampere area. H, Haveri Cu–Au deposit; K, Kutemajärvi gold deposit; Pu, Pukala intrusion; T, Tohloppi; M, Myllyniemi; Ka, Karppi; P, Pohjala, TSZ, Tampere shear zone. The location of Figure 2 is shown with a rectangle. Modified from Kähkönen (1994, 1999), grid coordinates YKJ.
Overall structure of the eastern part of the central Tampere Schist Belt (see Fig. 1 for location and for key; in addition, S, Sorila; Pi, Pirttiniemi). U–Pb zircon age data are indicated (in billions of years; see text for references). Figure compiled and modified from Nironen (1989a; cross-section; single units not shown) and Kähkönen (1994, 1999; map section). Width of the view is c. 16 km, length c. 28 km. The location of the study area is shown with a rectangle. CFGC, Central Finland Granitoid Complex.
The aim of this study is to investigate the structural development and the architecture of a part of the Tampere Schist Belt in more detail, including the relationships between the volcano-sedimentary sequences, the tectonic structures, and the sulphide-bearing schist horizons. Important insights are gained into understanding the mechanisms of the basin closure and the localization of the sulphide mineralization within the basin. The suggested conceptual model implies that early, gently dipping thrusts acted as both channels and traps for the mineralizing fluids that possibly were sourced either locally or from relatively shallow depths from the base of the basin infill. Continued horizontal compression caused subsequent rotation of the thrusts into their present subvertical position. Although the studied rocks have been somewhat metamorphosed, we have in general dropped the prefix ‘meta’ from rock type names.
Geological setting
The continental crust of the central Fennoscandian Shield is mainly composed of 1.95–1.80 Ga rocks, forming the so-called Svecofennian domain (Fig. 1a; Gaál & Gorbatschev 1987). The orogenic evolution that produced this orogen was considered as the single Svecofennian orogeny by Gaál & Gorbatschev (1987), whereas Lahtinen et al. (2005) divided it into three separate orogenies: the Fennian arc-accretionary orogeny at c. 2.0–1.85 Ga, the Andean-type Svecobaltic orogeny at c. 1.84–1.80 Ga, and the Nordic orogeny with a continent–continent collision at 1.82–1.79 Ga (Fig. 1a; Lahtinen et al. 2005). Rutland et al. (2004) and Williams et al. (2008) divided the earliest orogeny (‘Fennian’ of Lahtinen et al. 2005) into two orogenic episodes; the Early Svecofennian (or ‘pre-Bothnian) before c. 1.92 Ga, and Middle Svecofennian (or ‘post-Bothnian’) at c. 1.88–1.85 Ga. The two events were separated by rifting–extension of the pre-Bothnian basement and associated volcano-sedimentary deposition. Those researchers also concluded that the pre-Bothnian rocks were deposited in a large marginal basin and that neither the pre-Bothnian event nor the post-Bothnian event was arc-accretionary. In any case, most of the igneous and sedimentary rocks in central and southern Finland formed during the early ‘Fennian’ event(s), before c. 1.85 Ga (e.g. Lahtinen et al. 2005, and references therein).
The east–west-striking Tampere Schist Belt is c. 200 km long and lies in the centre of the Svecofennian domain (Fig. 1), but its best studied central part, from Viljakkala in the west to Orivesi in the east, extends for about 60 km (Fig. 1b). In the north, the Tampere Schist Belt is bounded by the igneous rocks of the c. 1.88 Ga Central Finland Granitoid Complex (e.g. Nironen et al. 2000). To the south of the Tampere Schist Belt lies the Vammala migmatite belt (also known as the Pirkanmaa belt). Near Tampere, the Tampere Schist Belt–Vammala migmatite belt contact is strongly sheared along a curvilinear shear zone, here named as the Tampere shear zone (Figs 1b and 2). In the east, the contact is obscured by the c. 1.88–1.87 Ga Siitama batholith near Orijärvi (Figs 1b and 2; Nironen et al. 2000). The batholith is internally relatively undeformed but clearly sheared along its contact with the Tampere Schist Belt.
The Tampere Schist Belt is characterized by 1.90–1.89 Ga volcanic and related sedimentary rocks with volcanic arc-type geochemistries, and by turbiditic sedimentary rocks originally deposited between c. 1.92 and 1.90 Ga (e.g. Seitsaari 1951; Ojakangas 1986; Kähkönen 1987, 1989; Lahtinen et al. 2009). The rocks are metamorphosed in low-T amphibolite- to transitional greenschist–amphibolite-facies conditions (Kilpeläinen et al. 1994). Overall, the central Tampere Schist Belt defines a main syncline with an east–west-striking, subhorizontal hinge zone and a steeply dipping axial plane (Figs 1b and 2; Seitsaari 1951; Kähkönen 1989; Nironen 1989a). The northern limb is dominated by volcanic rocks, whereas the southern limb is considerably richer in sedimentary rocks, although towards the southernmost margin volcanic rocks again become more abundant (e.g. Seitsaari 1951; Kähkönen 1994, 1999).
Generalized stratigraphy of the Tampere Schist Belt and derived ages of the rock formations are presented in Table 1. The nature of the basement underlying the Tampere Schist Belt volcanic rocks and sediments is unknown, but Lahtinen & Huhma (1997) inferred that evolved crust and associated mantle were in place in the area around the onset of the extension at >1.92 Ga. The lowermost recognized Tampere Schist Belt synrift unit is the volcanic Haveri Formation, with a relatively poorly defined Pb–Pb model age of 1990 ± 25 Ma (Vaasjoki & Huhma 1987), and consisting of (proto)oceanic crust (Fig. 1b; Mäkelä 1980; Kähkönen & Nironen 1994; Kähkönen 2005; Lahtinen et al. 2009). The Haveri Formation is overlain by the sedimentary Osara Formation, which is stratigraphically equivalent to the Myllyniemi Formation near Lake Näsijärvi (Fig. 1; Seitsaari 1951; Kähkönen 1989). Both the Osara and Myllyniemi Formations are stratigraphically correlated with the sedimentary rocks of the Karppi Formation farther east (Fig. 1; Seitsaari 1951; Kähkönen 1989). Sulphide-bearing schist layers are found locally within the lower parts of the Myllyniemi and Karppi Formations and in the upper part of the Haveri Formation. Upwards in the Tampere Schist Belt succession, the Myllyniemi sediments and their lateral counterparts are overlain by rocks of the 1898 ± 8 Ma Pirttiniemi Formation (Table 1; Fig. 2; Kähkönen 1989; Kähkönen et al. 2004). Pirttiniemi Formation volcanic rocks are likely to be equivalent to the ‘Tohloppi–Tesoma–Kiviranta–Sorila volcanic rocks’ that on the map (Fig. 2) appear to lie stratigraphically below the 1.92–1.90 Ga Myllyniemi Formation, but show ages of 1892 ± 2 Ga (Table 1; Kähkönen 1994). The onset of the basin inversion and the timing of the final closure are not well constrained, but it is implied that the crust was already thickened by c. 1.88 Ga: two large plutons, the 1878 ± 3 Ma Värmälä and the 1882 ± 6 Ma Hämeenkyrö plutons, intrude the Tampere Schist Belt (Figs 1b and 2; Table 1). Minor extrusive volcanism and related deposition at this stage is, however, also implied (Kähkönen & Huhma 2012).
Generalized stratigraphy of the Tampere Schist Belt rock formations
We will not discuss the development and structure of the Vammala migmatite belt in detail here, but the overall relationship between the Vammala migmatite belt and the Tampere Schist Belt (Fig. 1b) is important for understanding of the tectonic setting of the Tampere Schist Belt deposition and the subsequent basin closure. A turbiditic origin is interpreted for the gneisses and migmatites of the Vammala migmatite belt, but, in contrast to the Tampere Schist Belt, volcanogenic rocks are not common and the Vammala migmatite belt is considerably richer in black schists. A relatively deep basin setting for much of the Vammala migmatite belt is implied, but sedimentary rocks more typical for shelf and/or upper parts of submarine fan environments have also been found (Lahtinen et al. 2009). The temporal and spatial relationship between the two belts has been controversial. Nironen (1989a) suggested that the Tampere Schist Belt and Vammala migmatite belt represent a single basin but possibly represent different crustal levels or metamorphic conditions. Kähkönen (1999, 2005) considered the Vammala migmatite belt largely as a subduction-zone complex related to the arc volcanism in the Tampere Schist Belt. However, age determination results imply that the Vammala migmatite belt includes pre-1.91 Ga rocks; Rutland et al. (2004) interpreted this to mean that Vammala migmatite belt rocks formed the basement, which subsequently extended to accommodate the Tampere Schist Belt sediments. On the other hand, Lahtinen et al. (2009) showed that the Myllyniemi-type sediments in the lower part of the Tampere Schist Belt succession and bulk of the Vammala migmatite belt turbidites are identical in detrital zircon age populations, and that a part of the Vammala migmatite belt sedimentary rocks have c. 1.89 Ga detrital zircon probably derived from the Tampere Schist Belt arc volcanic rocks. Lahtinen et al. (2009) also emphasized changes in the sedimentary and tectonic environments of the evolution of the Tampere Schist Belt and Vammala migmatite belt. In their model, the basalts and the sedimentary rocks of the Haveri Formation were formed during rifting of a c. 2.0 Ga microcontinent. As the rifting continued, the sedimentation environment changed into a passive margin at c. 1.92–1.90 Ga, and the Myllyniemi-type sediments were deposited with the equivalent Vammala migmatite belt turbidites farther south. The onset of subduction and Tampere Schist Belt arc-type volcanism at c. 1.90 Ga caused a change from a passive margin–continental slope setting to that of an intra-arc system, in which the 1.90–1.89 Ga Tampere Schist Belt volcanism represents the ‘proper’ volcanic arc and the Vammala migmatite belt acted mainly as the accretionary complex.
The folding that produced the main Tampere syncline is interpreted to be associated with the main basin inversion and closure event (Fig. 2; F1 folding phase of Nironen 1989a). In particular, the southern central Tampere Schist Belt near Tampere seems to contain several tight to isoclinal, upright F1 folds with gently dipping to subhorizontal fold axes or hinge lines, ‘parasitic’ with respect to the main regional syncline (Fig. 2; Nironen 1989a). The southern limb is bounded by the Tampere shear zone; kinematic indicators imply thrusting toward the north (i.e. reverse south-side-up; Figs 1b and 2; e.g. Nironen 1989a; Kähkönen 1999). Nironen (1989a) suggested that the displacement along the thrust was essentially vertical, and that the thrust formed during the first deformation phase D1. The strike-slip kinematics of the Tampere shear zone has not yet been described in the literature; Nironen (1989a) interpreted that near Tampere, asymmetric structures have not been observed in the horizontal plane. He further interpreted that the Hämeenkyrö and Värmälä plutons (Figs 1 and 2) intruded during F1.
The area of this study is located in the eastern part of the central Tampere Schist Belt, SSE of the town of Orivesi (Figs 1 and 2). Kähkönen (1999) identified two units in the southern Orivesi area, just east of the selected study area (Figs 1b and 2): Karppi turbidites (greywackes and mudstones) in the north, and Pohjala volcanic rocks (various volcanic rocks, including plagioclase- and plagioclase–uralite-porphyries, of unknown ages) in the south. The study area was chosen because (1) the northern part, composed of sedimentary rocks presumably equivalent to the Karppi turbidites, represents a low level of the regional stratigraphy but the stratigraphic position of the (Pohjala) volcanic rocks in the southern part is unclear, (2) the detailed tectonic development and present structural geometry of the area is unknown, and (3) the area contains some sulphide-bearing schist horizons or lenses, and geophysical anomalies that are assumed to be associated with the schist horizons. The relationships of the sulphide-bearing schists to the overall volcano-sedimentary architecture and/or the tectonic history are also unknown.
Observations
Like elsewhere within the Tampere Schist Belt, the studied supracrustal rocks show well-preserved primary structures, despite the evidently considerable shortening during basin closure (Fig. 3). In many places, top-of-strata indicators and various other primary structures are clearly identifiable. The strata are usually strongly tilted with dips of >70° (Figs 4 and 5). The first secondary foliation (local S1) is a roughly east–west- to ENE–WSW-striking schistosity (Figs 4 and 5). S1 is almost exclusively subparallel to S0: the most substantial observed differences are shown in Figure 5. A younger schistosity S2 is seen to strike at an angle to S1, usually NE–SW, crenulating S1 (Fig. 3h). S1 and S2 vary in intensity depending on the rock type: the fine-grained, mica-rich mudstones often show intense schistosity, whereas the coarser-grained sedimentary and volcanic rocks may seem almost undeformed (Fig. 3). Kinematic markers (sigma clasts, S–C structures) are observed occasionally; on subhorizontal surfaces, they record both dextral and sinistral sense of shear. However, sinistral kinematic markers become dominant in the south and deformation becomes more intense. Stretching lineations are very rare, but are subvertical where present.
Representative photographs of the various rock types in the subunits described in the text. The volcanic unit subunits VU1–VU6 are shown in (a)–(f), the siliciclastic unit in (g)–(l). (a) VU1, stratified felsic volcanic sandstone; (b) VU2, stratified intermediate volcaniclastic sandstone and conglomerate; erosion surface between the fine and the coarse layers indicates southwards top-of-strata (dotted line); (c) VU3, felsic plagioclase porphyry; (d) VU4, ur–pl porphyry; (e) VU5, sandstone within the ‘southern’ volcaniclastic rocks, subvertical section, view toward west; this exposure is located north of the subunit VU6 and shows southward younging strata; (f) VU6, mafic and felsic tuffs; the tight folds (F1) with moderately (30–40°) eastwards plunging fold axes should be noted; (g) stratified, in part graded sandstones–microconglomerates, top-of-strata to the north (graded bedding); (h) sandstones (in part graded) with mica-rich mudstone layer, top-of-strata to north; noteworthy features are the well-preserved soft-sediment structures in the sandy layer (left of centre of the photograph), and the NE–SW-striking S2 schistosity crenulating S1 in the mudstone (right of the centre); (i) non-sulphidized mudstone; (j) sulphide-bearing mudstone–schist; (k, l) photomicrographs of the sulphide-bearing schist, in plane-polarized light (k), and in cross-polarized light (l). The schistosity is defined by brown mica. The abundant opaque minerals associated with the mica, and the secondary fractures mineralized with red haematite and opaque minerals, should be noted. The opaque minerals are presumably mostly pyrite and chalcopyrite, observable in some hand samples. The non-sulphidized schists do not contain abundant opaque minerals in thin section. Scale bar is 200 μm.
Equal angle stereographic projections summarizing the bedding (S0) and first-order schistosity (S1) observations. S1 can be seen to be mostly subparallel to S0.
Geological map of the study area, with the representative observations and interpreted fault or shear zones. The 100 m scale fining-upward cycles (conglomerate–sand–mud) within the siliciclastic unit are gradual, so that the drawn contacts between these units are approximate. Age references are given in the text. The bedding (S0) is normally subparallel to S1 (see also Fig. 5) and hence mostly omitted from the map, except from the few localities where the highest deviations were observed (i.e. 7–17°). SCU, siliciclastic unit; VU, volcanic unit.
At least two folding styles (here called F1A and F1B) are distinguishable in the field. F1A shows tight to isoclinal folds with subvertical fold axial planes and gently to moderately (up to c. 40°) plunging fold axes or hinge lines. F1B is observed in the southern part of the mapping area, where it is seen to refold F1A. F1B displays open to closed folds with steep axial planes and fold axes plunging c. 60–70° (Fig. 5).
The rock types are here divided into two main lithological units: the volcanic-dominated southeastern part (the volcanic unit) and the siliciclastic-dominated northern part (siliciclastic unit) (Figs 3 and 5). There is a sharp transition from one unit to the other, observable in only the eastern part of the mapping area owing to the Siitama batholith intruding the units in the western part (Fig. 5).
The two main units, the siliciclastic unit and the volcanic unit, are very distinct and display different lithological, architectural, and structural characteristics. The volcanic unit is composed mostly of coherent porphyritic rocks, evidently lavas, domes, and/or sills in origin, and of stratified volcanic rocks, but also, for example, pyroclastic deposits and volcanogenic sedimentary deposits. Single unit thicknesses may be up to c. 150 m. Most of the volcanogenic sedimentary deposits are found in the central part of the volcanic unit. S1 is usually (sub)parallel to S0 in the volcanic unit, and is best developed in the finer-grained layers. S1/S0 normally follows an east–west strike throughout the volcanic unit, except at the contact with Siitama batholith. The top-of-strata indicators (graded bedding, erosion surfaces) are common and dominantly young towards the south, except in the southern part of the volcanic unit where the younging direction is towards the north (Fig. 5). The siliciclastic unit rocks on the other hand are dominated by sandstones, but conglomerates (mostly granule conglomerates but some pebble conglomerates) and mudstones (shales and silts) are common as well. Single beds within the siliciclastic unit are of variable thicknesses, from centimetre to metre scale. Top-of-strata indicators systematically, with one possible exception close to the volcanic unit–siliciclastic unit contact, indicate northward younging directions (ripples, graded bedding, erosion surfaces; Fig. 3g and h). As with the volcanic unit, S1/S0 mostly follows an east–west strike except locally in the vicinity of the Siitama batholith. S2 schistosities are commonly observed in mudstones, striking mostly approximately SSW–NNE to WSW–ENE (Figs 3h and 5).
Based on the southwards-younging pattern of the volcanic unit, the stratigraphic succession of subunits of the volcanic unit can be defined (from bottom to top) as follows: VU1, felsic volcaniclastic rocks; VU2, intermediate volcaniclastic rocks; VU3, coherent felsic plagioclase porphyries; VU4, coherent uralite–plagioclase (ur–pl) porphyritic rocks; VU5, volcaniclastic rocks including conglomerates and sandstones (here called ‘southern volcaniclastic rocks’); VU6, mafic and felsic tuffs. The northwards-younging siliciclastic unit rocks are discussed as a single unit, as single beds do not define laterally coherent subunits as within the volcanic unit (Fig. 5). The relative age and stratigraphic position of the siliciclastic unit with respect to the volcanic unit will be discussed below.
VU1
The felsic volcaniclastic subunit is at its base composed of fine volcanic sandstones (‘tuffs’; Fig. 5). These are followed by volcanic conglomerates or breccias (‘lapillistones’) covered by coarse tuffs (Fig. 3a), both of which are interbedded with more fine tuffs. The beds in this unit are relatively well sorted. S1 is strong in the fine tuffs.
VU2
The intermediate volcaniclastic subunit contains two distinguishable types of strata: (a) sandy tuff with few lapilli and gradual variations in grain size leading to diffuse strata boundaries; (b) graded lapillistone interlayered with sandy tuff, including some cobble-sized fragments of porphyritic lavas (Figs 3b and 5). Some of the lapilli are strongly angular. Both types are rich in biotite and show strong S1.
VU3
The felsic plagioclase porphyries contain feldspar phenocrysts of variable sizes (typically 1–3 mm in diameter; Figs 3c and 5). A weak S1 schistosity is observed, and networks of hydrothermal veins are abundant.
VU4
The uralite–plagioclase porphyries envelope the subunits VU5 and VU6 (Figs 3d and 5). They contain abundant c. 1–3 mm phenocrysts of uralite and/or plagioclase in a fine-grained matrix. In the northern part of this unit, uralite phenocrysts are more common, whereas plagioclase phenocrysts gradually become more dominant toward the south. S1 is fairly weak in the northern part of the subunit but becomes strong towards the south. Both dextral and sinistral sigma clasts and S–C structures are observed in the southern part of the unit, but most of the observed indicators show sinistral shear.
VU5
The southern volcaniclastic subunit is exposed within two elongated areas, separated by the subunit VU6 (Figs 3e and 5). VU5 consists mainly of volcanic sandstones, pebbly sandstones, and conglomerates, but some tuff breccias rich in angular clasts are also observed. Weak S1 is present. As is typical for other volcanic unit rocks, ripples and erosion surfaces in this subunit show southward top-of-strata, but only in the exposures that are located north of VU6 (Figs 3e and 5). South of VU6, the VU5 rocks show northward top-of-strata. S1 is weak.
VU6
The mafic and felsic tuffs consist of alternating fine-grained mafic and felsic layers. The felsic material often occurs as lenses within the mafic material (Fig. 3f). F1 tight to isoclinal folds with steep axial planes and c. 40° eastward-plunging fold axes are observed (Fig. 3f). Close to the Siitama batholith, F2 refolds the F1 folds into open to closed folds with fold axes plunging c. 60° toward the SE to ESE (Fig. 5).
Siliciclastic unit
This unit (Fig. 3g–l) shows laterally varying layer thicknesses in an east–west direction so that the granule conglomerates, the sandstones and the mudstones define interfingered lensoid shapes on the geological map (Fig. 5). The grain-size variations define rhythmic (‘cyclic’) patterns at the map scale, with overall fining-upward sequences (younging towards the north). The only folds that can be observed are very local, tight, asymmetric (dextral) folds in the southern part of the siliciclastic unit (Fig. 5). These folds show west–east- to NW–SE-striking, steeply dipping axial planes and roughly 40° east- to SE-plunging fold axes.
Sulphide-bearing schists appear in three exposures within the siliciclastic unit (Figs 3j and 5). Two of the exposures occur in the central part of the mapping area, close to the contact with the Siitama batholith and the volcanic unit, and one exposure is a thin horizon in the NW. In addition, sulphide schists are also found within the Siitama batholith (Fig. 5). The sulphide schists are often extensively weathered and eroded, and have rusty outcrop surfaces. The bedding or other structures are not easily observable, but a few coarser-grained sandstone lenses are distinguishable in places. S1 is subparallel to S0 (where visible), and fractures are observable in some outcrops. The sulphide mineralization consists mainly of small amounts of disseminated pyrite and chalcopyrite visible to the naked eye, often located in the visible fractures, and of secondary mineralization (mainly haematite) identifiable in the fractures in thin section (Fig. 3k and l). In thin section, S1 is strongly developed, defined by elongate micas and opaque minerals (Fig. 3k and l). It is notable that the opaque minerals, presumably pyrite–chalcopyrite, and haematite, are present both within S1 and within the fractures cross-cutting it. The economic potential of the sulphide schists within the study area has not been assessed, but the grade seems to be low.
Interpretation
Structure
The observed S1 schistosities strike relatively constantly east–west throughout the study area and are (sub)parallel to S0 (Fig. 4). S1 is interpreted to represent the main deformation phase D1 that caused the inversion and closure of the basin and, ultimately, the large-scale folding. The intensity of the S1 schistosity is directly dependent on the lithotype: it is best formed in the mudstones but only weak in the sandstones and many of the porphyritic volcanic rocks. In the absence of clear signs of significant lateral movement (shearing) in most parts of the mapping area, the dominantly east–west strikes of S1 would imply approximately north–south shortening during D1, in accordance with previous studies (e.g. Kähkönen 1989; Nironen 1989a). The shearing becomes more intense towards the south, and sinistral kinematic markers (affecting also the northernmost edge of the Siitama batholith) and F2 folds (see below) can be observed.
S2 is locally observed, and it constantly strikes at an angle of c. 40–50° counterclockwise from S1. S2 is usually, but not always, observed in the mudstones, where it is seen to crenulate S1. S2 is interpreted to represent a D2 deformation phase separate from and younger than S1. As for most of S1, S2 is not seen to be associated with pervasive lateral shear in the field, so that the dominantly NE–SW strike of S2 suggests an approximately NW–SE-oriented compression during D2, consistent with the interpretation of Nironen (1989a).
The polarity observations imply a previously unknown F1A syncline, ‘the Pohjala syncline’, in the volcanic unit (Figs 5 and 6). The tightly folded rocks of subunit VU6 form the core of the syncline. The relatively steep c. 40° eastward plunge of the F1A fold axes in VU6 is interpreted to be representative for the entire syncline. The Pohjala syncline is best defined by the subunit VU5, which shows opposite bedding polarities on the north and south sides of VU6 (Figs 5 and 6). The rest of the subunits (VU1–VU4) form the northern limb of the Pohjala syncline (a narrow zone of subunit VU4 is also present in the southern limb). The c. 40° eastward plunge of the Pohjala syncline is interpreted to at least partly result from the intrusion of the Siitama batholith. In the hinge zone of the Pohjala syncline (in VU6), the F1B folding event is seen to refold the isoclinal F1A folds. The F1B folds are open to closed folds with hinge lines plunging c. 60° toward the SE to ESE. They are interpreted to be related to the sinistral shearing along the southern margin of the study area. Therefore, the shearing event occurred later than or at the late stages of the main basin closure, but probably before the formation of the S2 schistosity because the NW–SE-oriented D2 compression cannot be associated with the sinistral shearing along the margin.
A 3D block diagram presenting the interpreted structural geometry of the study area. The legend is as in Figure 4. The volcanic unit forms the ‘Pohjala syncline’, defined by the opposite younging directions of the volcanic unit units. The volcanic unit is separated from the siliciclastic unit by a fault, interpreted as an early (D1) thrust fault (see text and Fig. 7). The second, curved, right-lateral strike slip fault offsetting the siliciclastic subunits is interpreted to have formed during a later deformation phase (D2).
In the siliciclastic unit, north of the volcanic unit, both the lithology and the polarity indicators abruptly change (Fig. 5): igneous rocks are present only as relatively thin sills or dykes and, except at one location close to the volcanic unit–siliciclastic unit transition, the siliciclastic unit rocks consistently show northward top-of-strata. The siliciclastic unit rocks also show mostly northward-dipping S0. The age of the dykes or sills is unknown, but they show S1 schistosities. The tight, asymmetric folds observed in the southern part of the siliciclastic unit are interpreted to be F1A folds. There is a possibility that the intrusion of the Siitama batholith has caused the local folding. However, owing to the similarity between these folds and the F1A folds within the Pohjala syncline (similar tight to isoclinal folding with moderately eastward-plunging hinge lines), we interpret both folds to have formed during the main deformation (folding) phase. It is unclear whether the dextral asymmetry of the F1A folds at the bottom of the siliciclastic unit reflects a relative north-side-up movement, or whether the folds are parasitic folds of a larger fold within the siliciclastic unit. Because no sinistral F1A folds have been found (corresponding to the other hypothetical limb of a larger fold), and because the folds are only found close to an inferred fault zone (see below) we tentatively prefer the interpretation of these F1A folds as shear folds indicating north-side-up reverse movement.
The data reveal a ‘double-syncline’ (Figs 5 and 6): the Pohjala syncline within the volcanic unit in the south, and the larger Tampere Schist Belt main syncline system in the north (only the siliciclastic-dominated southern limb of which extends into the mapping area). However, no anticline has been observed between the two synclines, and the transition from volcanic unit to siliciclastic unit (and from one bed polarity trend to another) is sharp. The single southward top-of-strata observation within the siliciclastic unit close to the volcanic unit–siliciclastic unit contact may represent remnants of an anticline, but this is very local, as elsewhere along the contact the siliciclastic unit top-of-strata indicators are systematically northward. The abrupt change in structure, lithotypes, and bedding polarities together indicate that faulting must be involved and is responsible for the juxtaposition and the present configuration of the units. We interpret the volcanic unit–siliciclastic unit contact to represent a significant palaeothrust that we here name the Pohjala thrust.
The Pohjala thrust cannot be directly observed as the volcanic unit–siliciclastic unit contact itself is not exposed. The primary structures in both volcanic unit and siliciclastic unit are often very well preserved, even right next to the inferred contact zone, and the lack of asymmetric kinematic markers at subhorizontal outcrop in most parts of the study area suggests either subvertical shearing or coaxial flattening. The only possible asymmetric kinematic markers near the Pohjala thrust are found in the siliciclastic unit mudstones: the asymmetric F1 folds could be tentatively interpreted as shear folds indicating north-side-up reverse movement, supporting the interpretation of a palaeothrust and indicating the direction of its relative movement. The absence of obvious kinematic indicators in the vicinity of the contact zone might be explained by shear localization. Many examples exist worldwide of simple shear (non-coaxial strain) being strongly partitioned into narrow zones that accommodate large amounts of movement. For example, the Glencoul thrust in NW Scotland forms a part of a thin-skinned thrust system that accommodates up to tens of kilometres of transport along a mylonitic (non-coaxial) thrust zone that is only from some decimetres to some metres in thickness, whereas the strain directly below the shear zone is mostly coaxial (e.g. Butler 1984; Law 1987; Law et al. 2010).
Another discontinuity plane is interpreted to offset the subunits of the siliciclastic unit (Fig. 5). The inferred fault has an apparent dextral kinematics and bends into parallelism to S0 in the southern part of the mapped siliciclastic unit (Figs 5 and 6). The age of the inferred fault is uncertain. It may, therefore, represent a late dextral fault zone, formed as a response to D2 with roughly NW–SE compression, or alternatively it might form a part of an imbricate thrust system formed during D1 (the dextral offset therefore being apparent).
The most extensive sulphide schist occurrences in the area are found in the lowest parts of the siliciclastic unit near the Pohjala thrust (Figs 5 and 6). The sulphide mineralization is interpreted to have occurred in two phases. The first mineralization phase occurs within S1 (Fig. 3k and l). Therefore, the mineralization is interpreted to be genetically linked to the formation of S1 and to the closure of the basin. The sulphide-bearing hydrothermal fluids from which the minerals precipitated were probably derived from deeper within the basin. The subsequent, second mineralization phase resulted in deposition of sulphides and oxides into the fractures within and adjacent to the Pohjala thrust. There is no indication in the thin sections that the S1-related mineralizations were remobilized during this event. The timing of the second phase is unclear, but it occurred during a (locally) brittle deformation event at a late D1 stage or after the closure of the basin (possibly during D2). However, D2 structures are not observed to be associated with hydrothermal activity, so we consider it more likely that the second mineralization is related to late D1 faulting and fracturing.
Architecture
The siliciclastic unit and volcanic unit can be reasonably correlated with the Karppi turbidites and the Pohjala volcanic rocks, respectively, east of the mapping area (Fig. 1a; see also Seitsaari 1951). We also correlate the siliciclastic unit with the turbidites of the Myllyniemi Formation (Fig. 1; Kähkönen 1999), also typically located in the southern limb of the main Tampere Schist Belt syncline (Figs 1b and 2; Kähkönen 1989; Nironen 1989a). In accordance with previous studies (see references in Table 1), we interpret the interfingered siliciclastic rocks of the siliciclastic unit to have been formed in a submarine turbidite fan coevally with the other Myllyniemi-type sedimentary rocks at c. 1.92–1.90 Ga. At least five fining-upward cycles at a hundred-metre scale are observed in the mapped part of this unit, suggesting stepwise basin subsidence and sea-level transgression (Fig. 5). Alternatively, the apparent repetition of the cycles might be caused by a development of an imbricate system within the siliciclastic unit.
The architecture of the volcanic unit is simpler than that of the siliciclastic unit, assuming that the porphyries of VU4 are completely extrusive and not sills. The tuffs, lapillistones, and porphyries (VU1–VU4) are overlain by the relatively thin volcaniclastic VU5. The ripples in the sandstones of VU5 are consistent with erosion and redeposition within a dynamic volcanic environment (Fig. 3e). The youngest rocks within the volcanic unit are the fine VU6 tuffs in the core of the Pohjala syncline (Fig. 3f). Rheological differences between the mafic and the felsic material in VU6 during deformation are interpreted from the observed break-up of the felsic material into fragments, whereas the behaviour of the unbroken mafic material seems to have been more ductile.
The relative age of the volcanic unit with respect to the siliciclastic unit is not immediately obvious from the collected data. The age relationship of the units has consequences for the resulting geological model; this is discussed in the following section.
Discussion
Geological model
The age relationship between the volcanic unit and the siliciclastic unit is obscure: the polarity indicators and the lithologies and, therefore, the relative ages of the units cannot be systematically carried over from one unit to the other. However, if the asymmetry of the F1A folds observed close to the Pohjala thrust reflects north-side-up kinematics, then the siliciclastic unit is older than the volcanic unit. Stratigraphic correlation with dated units elsewhere within the Tampere Schist Belt supports this interpretation: the volcanic unit seems to occupy an approximately similar stratigraphic position to the Tohloppi–Tesoma–Kiviranta–Sorila volcanic units further west, whereas the Myllyniemi turbidites, at present north of the volcanic rocks, can be correlated with the siliciclastic unit (Kähkönen 1994; Figs 1b and 2). The depositional age of the Myllyniemi Formation has been determined as 1.92–1.90 Ga whereas the Tohloppi–Tesoma–Kiviranta–Sorila volcanic rocks are younger. Based on these implications that the volcanic unit is younger than the siliciclastic unit, we propose the following geological model for the studied area (Fig. 7).
Conceptual tectonic model for the (southern) Tampere basin prior to and during D1; not to scale. The general crustal-scale tectonic evolution (from continental rift to a continental slope in (a), to a subduction setting in (b), and finally to a collisional setting in (d)) is not shown. The internal architecture of the siliciclastic unit and volcanic unit described in the text is omitted for clarity. The ‘basement structures’ (folds) are for illustration purposes only. (a) Deposition of the siliciclastic unit sediments in a rift marginal basin (turbidite fans on a continental slope). (b) Subduction and associated volcanism initiates, and leads to basin inversion and a creation of an (intra-arc) basin at c. 1.89 Ga. The volcanic activity leads to the deposition of the volcanic unit (note that the entire volcanic package in the Tampere Schist Belt is much thicker than implied in the figure). Some of the extensional faults are reactivated as (blind) reverse faults, leading to open folding of the volcanic and turbiditic units along the southern margin of the basin. (c) A southward-propagating thrust, rooting layer-parallel within the siliciclastic unit, develops to accommodate further shortening of the Tampere Schist Belt. Imbricate faults may form at this point within the siliciclastic unit (omitted from the model). Hydrothermal fluids sourcing from the base of the basin infill (which possibly contains Haveri-type VMS mineralizations) may start channelling into the thrust zone at this stage. It should be noted that the northernmost thrust and the associated synclines (outside the study area) are hypothetical; the thrust is inferred only because of the presence of sulphide schist horizons north of the study area. Within the Vammala basin, another thrust with an opposite thrust direction (south-side-up) develops at the southern margin of the basin (or possibly already during the previous stage). (d) The continued shortening leads to a tightening of the folds and the rotation of the thrust faults into an upright position. At depth, the thrusts may be folded together with the rock units that host them. S1 starts to form, trapping the sulphides of the first mineralization phase. At the end of this phase, further significant reverse movements along the old thrust planes are unlikely, but strike-slip deformation can be expected (the sinistral, possibly slightly oblique, movements along the shear or thrust zones). Fracture zones develop in the palaeothrust planes owing to the strike-slip shear movements, forming a trap for the secondary mineralization within the fractures. (e) Basin closure and S1 is finalized, and the Siitama batholith intrudes along the interpreted continuation of the Tampere shear zone (TSZ).
(1) The siliciclastic unit was deposited, like other parts of the Myllyniemi Formation, mainly on a margin of a rifted microcontinent at c. 1.92–1.90 Ga, with the rift-related Haveri-type volcanic rocks partly at the base (Fig. 7a). The lobate architecture of the siliciclastic unit results from deposition largely on the continental slope during continuing but stepwise basin subsidence (Kähkönen 1999).
(2) The onset of northward subduction and related arc-type volcanism at c. 1.90 Ga caused a change to volcanism-dominated deposition in the Tampere Schist Belt for the period of c. 1.90–1.89 Ga. In our study area, the volcanism resulted first in the deposition of the felsic and intermediate tuffs (VU1 and VU2), followed by the plagioclase- and the uralite–plagioclase-porphyries (VU3 and VU4), the southern volcaniclastic rocks (VU5), and finally the mafic to felsic tuffs (VU6).
(3) During continuing subduction, basin inversion initiates with north–south compression (the local D1 event; Fig. 7b). Some of the extensional faults are reactivated as (blind) reverse faults, leading to open folding of the volcanic and siliciclastic units. This stage of local ‘thick-skinned’ deformation is needed to produce the original Pohjala syncline; a purely thin-skinned thrusting of subhorizontal sedimentary packages on top of each other would probably have resulted only in repetition of units, without significant folding of the units in the footwall of the thrust.
(4) As the basin closure progresses, the normal fault inversion cannot accommodate the increasing amount of shortening, and a southward-propagating ‘thin-skinned’ thrust system develops, including the Pohjala thrust (Fig. 7c). Where the thick-skinned shortening is accommodated at this point is not inferable from our data. The thrust planes are approximately layer-parallel and break towards the surface when they meet the buried basement highs defined by the faulted and partly inverted blocks. The thrust system juxtaposes different geological units, including the observed southward-younging and northward-younging units. The earlier anticlines are mostly destroyed as they are thrust upwards and subsequently eroded. At least one major thrust is active within the Vammala migmatite belt: this fault will later develop into a large (sinistral?) shear zone that is here interpreted to potentially be a continuation of the Tampere shear zone (TSZ in Fig 1; see also Nironen 1989a). The details of the development of the Tampere shear zone are outside the scope of this study, but the Tampere shear zone thrust–proto-shear zone helps to accommodate the overall shortening and tightens the Pohjala syncline.
(5) Continuing shortening leads to the development of S1, tightening of the folds, and the rotation of the bedding and the Pohjala thrust into an upright position (Fig. 7d). In their steeper position, the thrust(s) would become mechanically unfavourable for reverse movements. Lateral (shear) movements along the palaeothrust planes are more likely at this stage, the sense of shear possibly being mostly sinistral based on the field observations (the ‘minor movements’ in Fig. 7d). The shear movements were probably associated with fracturing, and we interpret the second mineralization phase to have occurred at this stage. Despite the lateral shear localizing into the palaeothrust plane(s), bedding-perpendicular coaxial (pure) shear while the beds were in an upright position probably accommodates some shortening. At depth, the layer-parallel thrusts are folded together with the rock units that host them.
(6) At the time of the intrusion of the Siitama batholith c. 1.88–1.87 Ga, the sinistral lateral kinematics in this part of the Tampere shear zone is established and the structures have reached their final subvertical orientations (Fig. 7e).
(7) Lastly, the second deformation phase D2, with roughly SSE–NNW compression crenulates the S1 schistosities, and leads to localized dextral faulting within the siliciclastic unit (the curved fault, Figs 5 and 6).
In the present model the bulk of the shortening is accommodated partly by folding, partly by the thrusts, and partly by the bedding-perpendicular shortening (probably including volume change and compaction). The minimum amount of shortening across the volcanic unit can be calculated, on the basis of the folding of this unit, to a minimum of c. 50%. This shortening estimate cannot be directly extrapolated to the scale of the Tampere Schist Belt, because volcanic unit-type folding is not at present identified in the northern parts of the belt. On the other hand, the interpreted overall geometry of the Tampere Schist Belt (Fig. 2) and the implications of the structural model presented above suggest a minimum shortening of a similar order of magnitude. The estimate does not consider volume change, or account for the unknown total number of thrusts within the Tampere Schist Belt. The possible existence of additional thrusts and/or imbricate–duplex systems within the siliciclastic unit would mean that the total shortening was notably larger than 50%.
The model largely agrees with previously published literature in terms of the overall temporal and spatial development of the basin. However, the model gives a more detailed suggestion for the structural development and shortening of the Tampere Schist Belt that might explain some of the observed problems within the Tampere Schist Belt (e.g. Nironen 1989a; Kähkönen 1999): most importantly, the data confirm the presence of early thrust(s) within the Tampere Schist Belt, as tentatively speculated by Kähkönen (1999). In addition, the model expands that of Nironen (1989a) in that we include extension and subsequent thrusting in the basement and suggest that the implied main thrusting direction within the southern part of the Tampere Schist Belt is toward the south. A northward-verging thrust system cannot explain the observed structural geometries and the inferred age relationships; however, if the volcanic unit is older than the siliciclastic unit, a northward-verging thrust system is possible. Northward structural vergences within and at the northern margin of the Vammala migmatite belt are consistent with our model, and the model in fact requires that the Vammala area is uplifted (or a subduction complex or accretionary wedge develops) at an early stage of the basin inversion. The southern central Tampere Schist Belt (the Tampere–Pohjala area) would, effectively, have formed a backstop to the developing Vammala wedge or complex.
Implications for the sulphide mineralization
The sulphide mineralization seems to result from a combination of sedimentary processes and hydrothermal activity (remobilization) along restricted zones within the Tampere Schist Belt. In thin section, the sulphide schist displays iron oxide and sulphide mineralization both within the schistosity and along the second-order fractures (Fig. 3k and l). The most prominent sulphide schist exposures within the study area are located close to the palaeothrust. This spatial association suggests a genetic relationship between the localization of the Pohjala thrust, the occurrence of the schistosity-parallel mineralization, and the localization of the mineralized second-order fractures. The model suggests that the first mineralization phase occurred at model stage (5)) when hydrothermal, mineralized fluids moved within and close to the Pohjala thrust. The ore minerals were trapped within the developing S1 schistosity close to the thrust plane (Fig. 3k and l). A second pulse of ore-bearing fluids is evidenced by the haematite- and sulphide-filled second-order fractures (Fig. 3k and l). There is no evidence of remobilization of first-order mineralization in thin section, although this cannot be ruled out, but the fluids were probably sourced from elsewhere. The age of the fractures is unknown but they localized within the Pohjala thrust after the main schistosity within the thrust had formed, and may reflect the late D1 strike-slip movements. D2 structures in the study area are not observed to be associated with hydrothermal activity, so we consider it unlikely that the second mineralization is related to D2.
The fluids were probably derived from within the lower parts of the basin not exposed in the study area. If the base of the basin infill in our mapping area consists of VMS rocks similar to those that are interpreted as the source for at least the Haveri deposit, they may have served as a source of mineralized fluids. Another possible source for the fluids is magmatic. The younger Siitama batholith can be excluded as a possible source as evidenced by the presence of sulphide schist xenoliths within the intrusion (Fig. 5). In addition, sulphide schist has also been observed north of the study area, kilometres away from the batholith, further suggesting a remobilization of sulphides by another mechanism than magmatic fluids. On the other hand, the 1.896 ± 3 Ga Pukala intrusion at the northern margin of the Tampere Schist Belt has been interpreted as the dynamo for the pipe-like epithermal Kutemajärvi gold deposit (Talikka & Mänttäri 2005; Kinnunen 2008). However, the Pukala intrusion was emplaced within the basin sediments in the northern part of the basin, and the alteration zone of the intrusion extends a maximum of c. 1 km from the intrusion (Talikka & Mänttäri 2005; Kinnunen 2008). Therefore, the Pukala intrusion is unlikely to be the source for the sulphidizing fluids in our study area. A nearby magmatic source would be expected to induce sulphidization and alteration throughout the study area, but most of the schists and all of the other rock types lack sulphidization and alteration. However, an unknown intrusion at depth cannot be excluded as a fluid source. Nevertheless, regardless of the original source (i.e. magmatic v. metamorphic), the results imply that the currently subvertical mineralized fault zones may have originally been more gently dipping and that their minerals did not necessarily originate from great depths (i.e. from deep within the basement). It is also noted that if local fracturing and shearing was indeed involved in the formation of the Kutemajärvi deposit, as suggested by Talikka & Mänttäri (2005) and Kinnunen (2008), it is in accordance with our interpretation of active early faulting prior to the rotation of the stratigraphic units to their present vertical position.
If the sulphide schists within the Tampere Schist Belt are related to early thrusts, the presence of sulphide schists might help in deducing the structural geometries and the basin closure processes elsewhere within the basin, and possibly aid mineral exploration in the area. For example, the Kutemajärvi and Haveri deposits are at least partly fault-controlled, and both deposits are hosted by schists (e.g. Nironen 1994; Eilu et al. 2003; Kinnunen 2008). Furthermore, an extensive sulphide schist horizon crops out just north of the mapped area, possibly indicating the presence of a prominent early thrust with similar mineralization. A minor sulphide schist lens is also found in the northeastern part of the mapping area (Fig. 5), implying that a minor thrust might exist here, possibly related to an unconfirmed imbricate system.
Summary
(1) The detailed structure of a part of the Palaeoproterozoic, well-preserved Tampere basin is described.
(2) The observed architecture and the structures of the volcano-sedimentary rocks within the southeastern Tampere basin imply the presence of hitherto unrecognized palaeothrusts within the basin. These thrusts played an important role in accommodating the overall shortening during basin closure.
(3) We propose a conceptual tectonic model that implies a transition from a normal fault inversion-type crustal shortening to locally ‘thin-skinned’ accommodation of shortening as the basin closure continued.
(4) The palaeothrusts were also critical in the localization of the fracture zones that transported and trapped sulphide-bearing hydrothermal fluids, as the sulphide mineralizations are spatially linked to the thrusts. The proposed tectonic model correlates the structural geometries observed in the field and the occurrence of the sulphide-bearing fractured schists.
(5) This study illustrates the importance of understanding the detailed structural development of a basin for the purposes of ore exploration.
Acknowledgments
This study was funded by a research grant from K. H. Renlunds Stiftelse, which is gratefully acknowledged. Editor I. Alsop and reviewers M. Nironen and T. Kilpeläinen are sincerely thanked for constructive comments.
- © 2014 The Geological Society of London
References
Relationships between basin architecture, basin closure, and occurrence of sulphide-bearing schists: an example from Tampere Schist Belt, Finland
