Multistage ore formation at the Ryllshyttan marble and skarn-hosted Zn–Pb–Ag–(Cu) + magnetite deposit, Bergslagen, Sweden

Ore Geology Reviews 69 (2015) 217–242

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Multistage ore formation at the Ryllshyttan marble and skarn-hosted Zn–Pb–Ag–(Cu) + magnetite deposit, Bergslagen, Sweden N.F. Jansson a,⁎, R.L. Allen a,b a b

Division of Geosciences and Environmental Engineering, Luleå University of Technology, SE-971 87 Luleå, Sweden Boliden Mineral, Exploration Department, SE-776 98 Garpenberg, Sweden

a r t i c l e

i n f o

Article history: Received 2 December 2014 Received in revised form 13 February 2015 Accepted 19 February 2015 Available online 21 February 2015 Keywords: Bergslagen Garpenberg Ryllshyttan Skarn Magnetite Sulphide

a b s t r a c t Numerous magnetite skarn deposits and marble- and skarn-hosted base metal sulphide deposits occur in polydeformed and metamorphosed, felsic-dominated metavolcanic inliers in the Palaeoproterozoic Bergslagen region of south-central Sweden, including the Ryllshyttan magnetite and Zn–Pb–Ag–(Cu) sulphide deposit, approximately 2.5 km SW of the large Garpenberg Zn–Pb–Ag–(Cu–Au) deposit. The Ryllshyttan deposit, from which approximately 1 Mt of Zn-rich massive sulphide ore and 0.2 Mt of semi-massive magnetite were extracted, is located near a transition between magnesian skarn and dolomitic marble. The host unit consists of a 10–20 mthick former calcitic limestone of likely stromatolitic origin that is commonly pervasively altered to skarn, locally hosting magnetite skarn deposits. The ore-bearing unit is one of several mineralised marble units within a more than 1 km-thick, felsic-dominated metavolcanic succession that includes a metamorphosed, large caldera-fill pyroclastic deposit, 800 m stratigraphically above the Ryllshyttan host succession. The Garpenberg stratabound Zn– Pb–Ag–(Cu)–(Au) deposit is located higher in the stratigraphy, just below the caldera fill deposits. The metavolcanic succession is bounded to the NW by a large granitoid batholith and intruded by a microgranodiorite pluton less than a 100 m from the Ryllshyttan deposit. Magnetite laminae in bedded skarns and metavolcanic rocks in the hanging wall of Ryllshyttan indicate an early (syngenetic) accumulation of Fe-rich exhalites. In contrast, the sulphide mineralisation consists of stratabound replacement-style ore associated with dolomitisation of the host and with discordant K–Mg–Fe ± Si alteration of volcanic rocks and early porphyritic intrusions in the footwall and hanging wall. The microgranodiorite that intrudes the host succession crosscuts the K–Mg–Fe ± Si alteration envelope and is overprinted by Na–Ca alteration (diopside and plagioclase-bearing mineral associations) that also overprints K–Mg–Fe ± Si-altered rocks. The Na–Ca alteration is interpreted to be associated with the formation of calcic and magnesian iron skarn deposits semi-regionally at a similar stratigraphic position. Despite superimposed amphibolite facies regional metamorphism and substantial syn-D2–D3 remobilisation of sulphides concurrent with retrograde alteration of skarn assemblages, cross-cutting field relationships indicate that the Ryllshyttan magnetite and Zn–Pb–Ag–(Cu) sulphide deposit results from protracted VMS-style hydrothermal activity including early seafloor mineralisation (Fe-rich exhalites), closely followed by sub-seafloor carbonate-replacement-style mineralisation (base metal-bearing massive sulphides). Both mineralisation styles were overprinted by contact metasomatism associated with the formation of abundant magnetite skarn deposits during the emplacement of granitoid intrusions. As for other deposits in the Bergslagen region, the ore-forming system at Ryllshyttan thus has similarities to both metamorphosed VMS deposits and metasomatic Fe and Zn skarn deposits. Our results suggest that the sequence of volcanic, intrusive and hydrothermal events in this region is compatible with prograde heating of a long-lived hydrothermal system, wherein a shift from a convective seawater-dominated system to a contact metamorphic and/or metasomatic environment occurred during the early stage of the 1.9–1.8 Ga Svecokarelian orogeny. This model partly resolves the controversy regarding genesis of the iron oxide and base metal sulphide deposits in Bergslagen, as we recognise that these deposits have a complex history of alteration, metamorphism, deformation and (re)mobilisation, and no unique established genetic model can account for all their features. © 2015 Elsevier B.V. All rights reserved.

1. Introduction

⁎ Corresponding author. E-mail address: [email protected] (N.F. Jansson).

http://dx.doi.org/10.1016/j.oregeorev.2015.02.018 0169-1368/© 2015 Elsevier B.V. All rights reserved.

The Palaeoproterozoic Bergslagen region (Fig. 1) in the Fennoscandian Shield is the birth-place of the term ‘skarn’, which after its introduction by Törnebohm (1875) has developed into a specific class of ore deposits,

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Tungsten oxide deposit Manganese oxide deposit Apatite-bearing, iron oxide deposit Iron oxide deposit in skarn or marble Quartz-rich, iron oxide deposit including BIF

Fig. 1. Regional geological map of the Bergslagen region, modified after Stephens et al. (2009) and Jansson and Allen (2011b). Inset shows area of Fig. 2. Deposits mentioned in the text are indicated; D: Dannemora, F: Falun, G: Garpenberg, P: Persberg, Z: Zinkgruvan. Inset shows the location of Bergslagen in the Fennoscandian Shield.

characterised by an intimate association between ore and diverse calcsilicate rocks. Skarn deposits are regarded as products of a variety of metasomatic processes, driven by fluids of magmatic, metamorphic and/or marine origin, at times of regional or contact metamorphism. Despite variations in commodities, skarn mineralogy, and timing relative to intrusive and metamorphic events at different deposits in the class, a

unifying feature is a strong spatial and temporal relationship between ore minerals and skarn minerals, of which garnet and clinopyroxene are the most characteristic (e.g., Meinert et al., 2005). Over 6000 mineral deposits – more than half of which are iron oxide deposits associated with skarn – are known in the Bergslagen region. The region also contain polymetallic sulphide deposits that can be

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divided into stratiform, sheet-like Zn–Pb–Ag–(Cu) deposits, such as Zinkgruvan, and stratabound, multi-lens carbonate replacement-type Zn–Pb–Ag–(Cu–Au) deposits spatially associated with skarn, such as Garpenberg (Allen et al., 1996). The later deposit class was originally referred to as deposits of ‘Falun-type’ by Geijer (1917), who suggested an intrusion-related, metasomatic origin for the ores and associated skarns. The contact metasomatic model was later largely discounted in favour of models involving VMS-like, syn-volcanic ore formation at or just below the seafloor, prior to ductile deformation and amphibolite facies metamorphism during the 1.9–1.8 Ga Svecokarelian orogeny (e.g., Koark, 1962; Frietsch, 1982; Vivallo, 1985; Allen et al., 1996). It has been demonstrated that the deposits occupy predictable stratigraphic positions relative to cycles of caldera-related felsic volcanism (Allen et al., 1996), although the deposits differ from typical VMS deposits in that sulphide ores replaced stromatolitic limestone units, formed during pauses in volcanic activity (Allen et al., 2003). It has furthermore been shown that zones of metamorphosed, altered volcanic rocks associated with the ores resemble hydrothermally altered zones associated with VMS-type deposits – metamorphosed to amphibolite grade – in terms of mineralogy, zoning and chemical change (e.g., Koark, 1962; Trägårdh, 1991; Vivallo, 1985; Allen et al., 2003). This shift in genetic thinking has led to that the calc-silicate rocks at e.g., Garpenberg have come to be regarded as products of regional metamorphism of pre-existing ores by some authors (e.g., Vivallo, 1985). Nevertheless, the complex textural relationships between skarn minerals and sulphides at Garpenberg – where sulphides occur as inclusions in skarn minerals, intergrown with skarn minerals, or veining skarn minerals – leads to that despite the evidence of early syn-volcanic ore formation below the seafloor (Allen et al., 2003), it is difficult to dismiss an influence of metasomatic, skarn-forming processes at the time of ore formation (e.g., Zetterqvist and Christoffersson, 1996; Allen et al., 2003; Allen et al., 2008; Jansson, 2011). Radiometric dating combined with key cross-cutting relationships have furthermore shown that the c. 1.89 Ga age of the volcanic host stratigraphy and ore formation in the Garpenberg area, overlaps with the ages of subvolcanic, porphyritic intrusions and granitoids adjacent to the deposits (Jansson and Allen, 2011a), of which some record endoskarn alteration (Jansson and Allen, 2013). The spatial and possibly genetic relationship between the stratabound sulphide deposits and iron oxide deposits in Bergslagen also remains partly unresolved. Even though major apatite–iron oxide deposit and banded iron formations occur in western Bergslagen (Fig. 1), most iron oxide deposits are low-P (b0.1 wt.% P) marble- and skarn-hosted magnetite deposits which have traditionally been divided into Mn-poor (b 1 wt.% Mn) and Mn-rich (N1 wt.% Mn) varities (Geijer and Magnusson, 1944). Manganese-rich deposits include the currently producing Dannemora deposit, whereas the Mn-poor class includes the Persberg deposit that represents the type locality of skarn (Törnebohm, 1875). It also includes the 3.7 Mt Smältarmossen deposit in the Garpenberg area, for which a contact metasomatic origin related to a metamorphosed dacite porphyry intrusion at c. 1.89 Ga was recently established (Jansson and Allen, 2013). Jansson and Allen (2013) suggested that the formation of the Smältarmossen Mn-poor magnetite deposit was synchronous with or post-dated the formation of sulphide deposits in the Garpenberg area, even though the genetic relationship between these deposit types was left unresolved as no direct crosscutting relationships were observed. Co-existence of a Mn-poor magnetite skarn deposit and a polymetallic sulphide deposits does however occur at the Ryllshyttan deposit in the same area (Fig. 3), providing an opportunity to study the relationship between these two important mineralisation styles. From the early 16th century until 1944, the Ryllshyttan Mine produced 1 Mt of massive to semi-massive sulphides grading 11 wt.% Zn, plus an unknown amount of Pb and Ag (Allen et al., 1996), as well as 0.2 Mt of semi-massive to massive skarn- and marble-hosted magnetite ore grading of N 25 wt.% Fe (Geijer and Magnusson, 1944; Magnusson

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et al., 1944). An estimated 3 Mt of magnetite-skarn with 20–40 wt.% Fe are known to remain (Allen et al., 1996). In this contribution, we address the relationship between iron oxides, sulphides as well as associated skarn and alteration mineral associations at the Ryllshyttan deposit, assess the relative timing of events and propose a genetic model for each ore type. In detail, we assess to what extent the characteristics of Ryllshyttan conforms to a seafloor hydrothermal system model as opposed to a skarn-style, metasomatic model, and whether the iron oxide and sulphide mineralisation formed as parts of the same hydrothermal system or whether they formed from two distinct, superimposed systems. In terms of exploration perspective, the results provide a basis for identifying other iron oxide deposits in Bergslagen where a potential for polymetallic sulphide ores may exist. As the sulphide and iron oxide deposits of the Garpenberg area share features with both metamorphosed VMS deposits (and associated iron formations) and metasomatic skarn deposits, this contribution will furthermore propose a model which can explain these apparently disparate features. 2. Geological setting The Bergslagen region forms part of the Svecokarelian orogen in the Fennoscandian shield (Fig. 1). The region has been interpreted to record two major tectonic cycles at 1.91–1.86 Ga and 1.86–1.82 Ga, inboards of an active continental margin undergoing migratory tectonic switching (Hermansson et al., 2008; Stephens et al., 2009; Beunk and Kuipers, 2012). The first cycle involved deposition of (the iron oxide- and sulphide-hosting) volcanic and sedimentary successions in a back-arc basin on continental crust, followed by emplacement of voluminous calc-alkaline plutonic rocks at 1.9–1.87 Ga and a phase of ductile strain (D1) involving folding under low pressure/high temperature amphibolite facies conditions (M1) at c. 1.87 Ga. This was followed by emplacement of plutonic rocks of more alkali-calcic character at 1.87–1.84 Ga and intra-orogenic sedimentation (Stephens et al., 2009; Beunk and Kuipers, 2012). The second cycle involved folding and shearing (D2) at 1.84–1.82 Ga under amphibolite facies conditions (M2), emplacement of plutonic rocks of more alkali-calcic character until 1.78 Ga, followed by a transition to shearing and faulting along more discrete structures (Stephens et al., 2009; Beunk and Kuipers, 2012). The mineralised metavolcanic successions occur as isolated, folded inliers that are bounded by plutonic rocks and shear zones. The intensely mineralised NW side of the Garpenberg supracrustal inlier (GSI) comprises a NE-trending, tight and complex F2 syncline, referred to as the Garpenberg Syncline by Allen et al. (2003). F2 Fold axes plunge steeply to the NE or steeply to the SW, and subsidiary tight, doubly-plunging F2 folds, 0.5–1 km in diameter, have been mapped in the Garpenberg mines (Allen et al., 2003). These features have been attributed to D2 sheath folding, associated with intense subvertical stretching (Allen et al., 2003). The F2 designation is based on folding of an older bedding-parallel S1 foliation though F1 folds are rarely recognised in the GSI, possibly due to the strong overprint from D2 (Allen et al., 2003). The Garpenberg syncline is bordered towards the SE by a major reverse D2 shear zone (The Stora Jelken Shear Zone, Fig. 2) and towards the NW by ca. 1895–1890 Ma synvolcanic granitoids (Jansson and Allen, 2011a). The stratigraphy of the GSI is described in detail in Allen et al. (2003). Rhyolitic metavolcanic rocks predominate, however dacitic and basaltic metavolcanic rocks and marble occur in particular stratigraphic intervals. Secondary ion mass spectrometry (SIMS) U–Pb zircon ages suggest that the rocks of the GSI were emplaced during a relatively short and intense period of volcanism and that burial to depths of 2–5 km, granitoid intrusion and mineralisation all occurred between ca. 1895 and 1890 Ma (Jansson and Allen, 2011a). Allen et al. (2003) concluded that the stratigraphic succession records the evolution of a large marine felsic caldera complex and that the 20 to 100 m-thick original limestone

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Major feldspar-phyric rhyolitic-dacitic pumice±lithic volcaniclastic units Major quartz-feldspar phyric rhyolitic pumice±lithic volcaniclastic units Lower footwall rhyolitic ash-siltstone with minor coarse volcaniclastic units Undifferentiated felsic volcanics Zn-Pb-Ag sulphide deposits Magnetite skarn deposits Bedding, stratigraphic younging 6684000

Major fault or shear / reverse shear zone

Fig. 2. Geological map of the Garpenberg region. Modified after Allen et al. (2003). Ryllshyttan is located in the ‘Lower footwall rhyolitic ash–siltstone with minor coarse volcaniclastic units’ of Allen et al. (2003). Swedish National Grid RT90.

host unit to the Garpenberg deposits formed via microbial stromatolite growth during pauses in volcanic activity. The other thinner marble units of the Garpenberg supracrustal inlier sequence are thought to have originated from similar processes. The uppermost stratigraphic unit in the core of the syncline is interpreted as a metamorphosed, caldera-fill, non-welded pyroclastic flow deposit. The iron oxide deposits of the GSI, including Ryllshyttan, mainly occur near the base of the exposed stratigraphic succession in thin marble units or their skarn-altered equivalents, whereas the polymetallic sulphide deposits such as Garpenberg are mainly associated with a stratigraphically higher marble (Fig. 2). However, there are important exceptions to this pattern: the polymetallic sulphide ore body at the Ryllshyttan Mine is located in the lower part of the stratigraphy; whereas the contact metasomatic Smältarmossen iron skarn deposit is located

higher in the stratigraphy, at a similar stratigraphic position as the Garpenberg polymetallic sulphide deposits. 3. Methodology Detailed geological logging of 34 drill holes was complemented by field mapping on both deposit and semi-regional scale. All rocks in the study area are deformed and were metamorphosed under lower amphibolite facies conditions (Vivallo, 1984; Allen et al., 2003). Primary rock textures are only locally preserved. For brevity and to emphasise the primary features, descriptive sedimentary and volcanic rock names are used in the following sections for the least altered rocks in which primary textures are well preserved. Strongly altered rocks are described using their current metamorphic mineralogy. The term

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‘limestone’ is used when discussing the host in a stratigraphic context, whereas the unit is referred to as ‘marble’ in the description of paragenetic relationships in the ore zone. One hundred and fifteen samples were submitted to Acme Laboratories in Canada for whole-rock lithogeochemical analysis. The lithogeochemical data was obtained to define the composition of variable altered volcanic units and intrusions, support stratigraphic correlations, and document the chemical modifications caused by the hydrothermal alteration. Major elements and certain trace elements were determined by inductively coupled plasma-emission spectroscopy (ICP-ES) whereas refractory trace elements were determined by

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inductively coupled plasma-mass spectrometry (ICP-MS), both following lithium borate fusion and nitric acid digestion. Total sulphur and carbon were determined by the Leco method. Analytical quality was monitored by the use of internal lab standards and our own hidden reference samples submitted with the unknowns. A total of 110 thin sections were studied to characterise the mineralogy of the different rock types as well as paragenetic and textural relationships. Electron microprobe (EMP) analyses were performed on thin-sections at Uppsala University, Sweden (Cameca SX50 WDS) and Copenhagen University (JEOL JXA8200 superprobe), Denmark. The analyses were made using a 15–20 kV acceleration voltage and 10–

Fig. 3. Surface geological map of the Ryllshyttan area. The terms ‘Silver’, ‘Tombak’ and ‘Kompani’ are the historic names of the three main workings of the Ryllshyttan mine. The metamorphosed sedimentary and igneous rocks are presented with interpreted original volcanic and sedimentary facies first, followed by current metamorphic rock type in brackets. Swedish National Grid RT90.

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A third deformation phase (D3) has caused shearing, transposition and possibly refolding of the earlier D2 structures along ENE-striking shear zones, including the Ryllshyttan Shear Zone (RSZ) immediately south of the Ryllshyttan Mine (Figs. 2–5). The youngest structures recognised at Ryllshyttan are brittle faults designated D4 (Fig. 5). These generally strike E–W and have curved surfaces that dip 30–80° north (Figs. 5, 6D). Cross sections and epidote slickenfibres suggest dominantly reverse dip–slip movements with a minor dextral horizontal component.

15 nA beam current. Measured peaks were calibrated against natural and synthetic standards. Where clinopyroxene and garnet compositions are determined, these are reported with end-member compositions. 4. Deformation of the Ryllshyttan deposit and host succession The dominant folds at Ryllshyttan are upright, tight, steeply plunging and disharmonic F2 folds with NE–ENE striking axial surfaces and curvilinear fold axes plunging mainly 70–90° towards ENE but locally WSW (Figs. 3–5). The folds are commonly accompanied by ENEtrending shear zones that commonly coincide with the attenuated, sinistrally sheared short limbs of the folds (Fig. 6A). A strong, steeply SE–SSE dipping S2 foliation comprises a distinct grain-shape fabric, aligned micas and veinlets of quartz, amphibole and/or epidote parallel to F2 axial surfaces. Near the hinge zones of F2 folds, S2 is seen to crenulate an earlier S1 foliation (Fig. 6B, C) that is generally sub-parallel to S0, whereas this foliation is difficult to distinguish from S2 on the F2 fold limbs. S1 is associated with tight-isoclinal mesoscale asymmetric F1 folds. Major D1 structures such as faults, highstrain corridors or map-scale folds have not been identified. The S1–S2 crenulation is strongly annealed with triple junction grain-boundaries indicative of late- to post-D2 static recrystallisation.

5. Stratigraphy Due to substantial textural and compositional modifications resulting from hydrothermal alteration, metamorphism and deformation, traditional stratigraphic analysis had to be supported by wholerock lithogeochemistry for characterisation of primary rock types and correlation of stratigraphic units across the tectonic structures. Element mobility must be assessed before using whole-rock lithogeochemical data to determine primary rock compositions. Barrett and Maclean (1994) showed that if analytical data for samples with different degrees of alteration from an originally chemically homogenous rock plot along a straight line passing through the origin in a binary diagram, then it can

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Fig. 4. Geological map of the 126 m level of Ryllshyttan. Legend as in Fig. 3. Modified after Jansson and Allen (2011a). Swedish National Grid RT90.

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be concluded that the two elements shown on the diagram were immobile during all alteration processes. Spread of sample points along this “alteration line” reflect dilution and concentration of the immobile elements, caused by mass gain and mass loss respectively of the mobile elements during alteration. At Ryllshyttan, Al2O3, Zr, TiO2 (Fig. 7A–C) and Nb display immobile behaviour, whereas K2O, Na2O, CaO, MgO and Fe2O3 were mobile during alteration (Fig. 8). The following stratigraphic description focuses on the least-altered samples, which were defined as samples that have no anomalous Ca, Mg, Mn, Fe or Si-contents or K/Na-ratios relative to common igneous rocks, have a low (b3 wt.%) loss on ignition (LOI), possess a mineralogy akin to primary igneous rock or their metamorphosed equivalents and show good preservation of primary igneous and volcanic textures (Table 1). Even these samples display evidence of minor alteration. Consequently, traditional major oxide classification diagrams are inappropriate for these data. In the Zr/TiO2 vs. Nb/Y diagram of Winchester and Floyd (1977) (Fig. 9), the least-altered samples define a bimodal pattern with one cluster of sub-alkaline rhyolites and another of subalkaline basalts to basaltic andesites. Data for the plutonic rocks at

Ryllshyttan have been included for comparison, as these rocks are coeval with the volcanic rocks (Jansson and Allen, 2011a). The Ryllshyttan limestone is underlain by a ca. 50 m-thick quartzand feldspar-phyric rhyolitic breccia–sandstone (Figs. 10, 11A), whose lower contact is truncated by intrusions (Figs. 3–5). The breccia– sandstone is dark-grey to pink in colour and contains 3–25 vol.% monomict, sub-angular to ragged 1–4 cm clasts in a sandy matrix with 1–3 vol.% 1 mm quartz and microcline phenocrysts. Many discernable clasts have wispy margins, suggesting that they were originally pumice (Fig. 11A). Other clasts are blocky and angular and were probably lithic fragments. The unit is generally massive with an even distribution of clasts and phenocrysts, yet normal-grading is locally observed. The top contact of the rhyolitic breccia–sandstone unit against the Ryllshyttan limestone is sharp and conformable in the least deformed and altered areas. The rhyolite unit is characterised by tightly clustered immobile element ratios, and anomalously high Zr/TiO2 (Fig. 7A–C). The leastaltered sample show a negative Eu anomaly, La/Gdn at 4.14, a nearhorizontal Tb–Lu slope in a chondrite-normalised REE plot (Fig. 12A), and a Zr content below 200 ppm (Table 1) — all typical of calc-alkaline

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Fig. 6. Structural relationships observed at Ryllshyttan. A: S — symmetric F2 fold developed in a bed of interlaminated garnet–epidote–ferroan diopside–magnetite rock in interbedded rhyolitic siltstone/sandstone–aluminous skarn. Note strong attenuation and shearing of the ENE-trending short limb. B: Crenulated S1 and S0 within the hinge zone of the Silver Syncline. Sub-horizontal section viewing down-plunge of the fold–hinge. Sample is a phlogopite–quartz schist from Ryllshyttan's metamorphosed alteration envelope. C: Amygdaloidal amphibolite after a mafic intrusion, with S1 folded around a NE-trending axial surface associated with a S2 foliation. D: E–W trending, north-dipping reverse D4 faults in the Tombak mine. View towards the east.

igneous rocks (Barrett and MacLean, 1999). The monomict clast population, abundant wispy, probable pumice clasts, lack of segregation between crystals and pumice, and the homogeneous immobile element ratios suggests that the poorly-sorted quartz and feldspar phyric rhyolitic breccia–sandstone is a rhyolitic pyroclastic deposit. The massive to normal graded bedforms, poor sorting, association with limestone and the apparent lack of erosion at the upper contact, are most consistent with deposition from clastic mass flows in a subaqueous environment below wave base (cf. Allen et al., 2003). The Ryllshyttan limestone is 10–20 m-thick and comprises massive to coarsely banded, medium to coarsely crystalline (1–2 mm) white dolomitic marble that is heterogeneously replaced by skarn, magnetite and sulphides that form the Ryllshyttan ore bodies. The skarn consists of alumina-poor associations of andradite, diopside–ferroan diopside, tremolite–actinolite, chondrodite, serpentine and carbonates. The skarn mineralogy and assemblages will be further described below. The stratigraphic upper contact of the limestone is sharp and conformable. The low Zr, Ti and Al contents of limestone and skarn in the ore zone indicates that the limestone precursor had a very low detrital volcaniclastic component. The crystal-rich beds at the base of the unit overlying the limestone are interpreted as reworked rhyolitic sediments in which the fine-grained rhyolitic ash has been winnowed out by wave-action, leaving a crystal-rich residue (cf. Allen et al., 1996, 2003). The direct juxtaposition of this facies upon the Ryllshyttan limestone suggests that the limestone accumulated in a shallow marine environment. Consequently, the limestone may have accumulated in the photic zone whereby a stromatolitic origin seems likely. Based on C–O isotope systematics and facies analysis, Allen et al. (2003) concluded that most limestone units in Bergslagen originated as stromatolitic reefs that formed during times of reduced volcaniclastic sedimentation.

The limestone is stratigraphically overlain by a ca 100 m-thick interval of rhyolitic siltstone–sandstone interbedded with aluminous skarn (Fig. 10). The basal meter comprises coarse-grained quartz and feldspar crystal-rich sandstone (as described above) that grades upwards into siltstone with subordinate crystal-rich sandstone interbeds. This in turn grades upwards into finer-grained, thinly planar stratified siltstones that are interbedded with aluminous skarn beds and subordinate beds of massive sandstone with minor 1 mm feldspar and quartz crystals (Fig. 11B). A few thin rhyolitic beds are normal graded, though beds are generally massive. Small-scale cross-bedding is present locally. Polymict breccia–conglomerates with clasts of laminated rhyolitic siltstone, pink fine-grained siliceous clasts and rhyolite porphyry (described below) clasts locally occur (Fig. 3). The skarn interbeds are dominated by andradite–grossular, epidote, ferroan diopside and quartz. They locally contain relics of laminated magnetite iron formation (Jansson and Allen, 2011b). Within each skarn bed, there is an internal, fine interlamination between different skarns minerals, rhyolitic siltstone and locally magnetite. Planar stratification, local normal grading, good sorting and dominantly silty grain-size suggest that the rhyolitic siltstone beds formed by settling of silty volcaniclastic sediment from suspension and dilute turbidity currents in a subaqueous environment below wave base. Only minor evidence of traction reworking was observed. The interbedded aluminous skarn beds are interpreted as reaction skarns formed from metamorphism of rhyolitic siltstone that co-settled with calcareous–ferruginous hydrothermal sediments (Jansson and Allen, 2011b). Laminae in this facies can be followed over tens of metres in the large outcrops, which also indicate deposition from dilute suspension below wave base. The breccia–conglomerate may represent an altered debris flow deposit (Jansson and Allen, 2011b).

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25.00

100

200 Zr ppm vf

300

400

0

100

200 Zr ppm vf

300

400

Granite

C

Microgranodiorite

20.00

Al 2O3 % vf

Mafic intrusion

Basaltic andesite

15.00

Garpenberg - regional metadolerite sill (Fig. 2)

10.00 Coherent rhyolite porphyry and associated peperite (peperite matrix avoided)

5.00

Footwall - poorly-sorted quartz+feldspar phyric rhyolitic breccia-sandstone

0.00 0.0

0.2

0.4

0.6 0.8 1.0 TiO2 % vf

1.2

1.4

1.6

Fig. 7. Binary immobile element plots for Ryllshyttan juvenile volcanic rocks and major intrusions. A large regional dolerite sill near the hinge of the Garpenberg syncline and the granite west of the GSI have been added for reference (Fig. 2). Note that the mafic intrusions include several different intrusive bodies which may not necessarily have shared exactly the same starting composition. ‘Footwall’ here refers to the stratigraphic footwall of the limestone host.

The interbedded rhyolitic siltstone–sandstone–aluminous skarn grades conformably upwards over tens of metres into red to grey, well-sorted, massive to diffusely stratified rhyolitic siltstone and sandstone with subordinate breccia–sandstone intervals (Figs. 10 and 11C). The boundary is expressed by a decreasing content of skarn beds and a slight coarsening. The siltstone–sandstone is commonly both feldspar- and quartz-phyric. 6. Intrusions The stratigraphic succession is intruded by a pinkish-white quartzand microcline-porphyritic rhyolite with ~ 15 vol.% 2–3 mm evenly distributed phenocrysts in a 0.05–0.1 mm groundmass (Fig. 3). Minor (1–3 vol.%) hornblende or biotite is evenly distributed in the groundmass (Table 2). The porphyry intruded into both the lower hanging wall and the footwall of the limestone host. In the hanging wall, the porphyry is a peperitic 20 m-thick sill, comprising blocky to globular bodies of porphyry enclosed by rhyolitic siltstone (Fig. 13A–B). In contrast, the porphyry in the footwall is mainly coherent but the margins are peperitic. The peperitic facies indicates emplacement into wet and unlithified sediment (Kokelaar, 1986; McPhie et al., 1993). The porphyry display clustered Zr/TiO2, which overlaps with the more spread Zr/TiO2 of the rhyolitic siltstones stratigraphically above the limestone host, yet shows no overlap with the distinctly higher Zr/TiO2 of the rhyolitic breccia–sandstone (Fig. 7). This along

with occurrence of rhyolite porphyry clasts in the polymict breccia– conglomerate interbed suggests that the porphyry and rhyolitic siltstones may have formed from similar magmas, whereas the rhyolitic breccia–sandstone formed from a different, more evolved rhyolitic magma. These relationships along with the peperitic margins suggest that the porphyries are high-level subvolcanic intrusions. The leastaltered sample of rhyolite porphyry show negative a negative Eu anomaly in a chondrite-normalised REE plot, La/Gdn at 4.46, a near-horizontal Tb–Lu slope (Fig. 12B) and a Zr content below 200 ppm (Table 1). The porphyry intruded into the rhyolitic breccia–sandstone at Ryllshyttan has yielded a SIMS U–Pb zircon age of 1895 ± 4 Ma (Jansson and Allen, 2011a). Plagioclase–porphyritic mafic intrusions, metamorphosed to amphibolites, occur as dykes, sills and irregular intrusive bodies (Fig. 13C–D, Table 2), and a small intrusion of basaltic andesite composition has been observed south of Ryllshyttan (Fig. 5, Table 2). The presence of S1 and S2 foliations indicates pre- to syn-D1 emplacement, and presence of strained amygdales (Fig. 6C) suggests an emplacement depth shallow enough to allow melt vesiculation (cf. McPhie et al., 1993). The mafic intrusions cross-cut rhyolite porphyry, have sharp and commonly chilled margins and locally contains angular wall-rock fragments, suggesting that emplacement occurred after induration and dewatering of the volcaniclastic rocks. It is possible that these intrusions were feeders to mafic volcanism c. 500 m stratigraphically higher in the GSI (cf. Allen et al., 2003; Fig. 2). The least-altered samples of mafic intrusion

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strong-intense sericite and/or chlorite alteration

90

80 moderate sericite and/or chlorite alteration

Hanging-wall volcaniclastic rocks Footwall volcaniclastic rocks

70

60

50

40

30

Unaltered to weakly altered samples

20 Moderate-strong Na-Ca alteration Rhyolite

Dacite

10

Basalt-andesite

Ishikawa alteration index - 100(MgO+K 2O)/(MgO+K2O+Na2O+CaO)

100

0 0.0

0.1

0.2

0.3

Zr/TiO2 Fig. 8. Ishikawa alteration index (Ishikawa et al., 1976) for Ryllshyttan igneous rocks. Legend as in Fig. 7. Unfilled triangles denote rhyolitic silt- and sandstone from the stratigraphic hanging wall of the limestone host. Dashed vertical line illustrates the different Zr/TiO2 of these volcaniclastic rocks compared with the volcaniclastic rhyolitic rocks of the stratigraphic footwall.

and basaltic andesite display no Eu anomalies and La/Gdn at 2.56 and 4.30 respectively (Fig. 12C–D). The flatter REE pattern for the mafic intrusion compared with the rhyolite suggests a transitional affinity (cf. Barrett and MacLean, 1999), whereas the basaltic andesite has a similar REE pattern to the rhyolite, except for the lack of a negative Eu anomaly. The Ryllshyttan succession is intruded by a synvolcanic batholith with a 500 m thick margin of microgranodiorite (Fig. 3, Table 2) containing xenoliths of mafic and laminated rhyolitic rocks (Fig. 13E). The microgranodiorite is pink-white with a subtle porphyritic texture defined by more than 25–30 vol.% 1–3 mm oligoclase crystals in a medium-grained groundmass. Mafic minerals account for 5–10 vol.% and include diopside, epidote and actinolite, which most likely represent alteration products of former primary magmatic mafic minerals. Titanite, zircon and apatite are accessory minerals. The rock shows variable strain and locally has a mylonitic fabric with a distinct augen texture. Jansson and Allen (2011a) determined a SIMS zircon U–Pb concordia age of 1894 ± 4 Ma for the intrusion. The microgranodiorite defines a distinct cluster on immobile element box plots; a signature similar to that of dacitic rocks spatially associated with the Garpenberg Zn–Pb–Ag–(Ag–Au) deposits (Fig. 14A–B), suggesting that both units are co-magmatic. These dacitic rocks at Garpenberg comprise synvolcanic intrusions and pumice±lithic breccias (Allen et al., 2003), and the intrusive facies are locally related to the formation of iron skarn deposits through contact metasomatism (Jansson and Allen, 2013). Northwest of Ryllshyttan, the microgranodiorite borders a pinkish-red granite with biotite and hornblende as dominant mafic phases (Table 2). The granite yielded a SIMS zircon U–Pb concordia age of 1895 ± 3 Ma (Jansson and Allen, 2011a) and regional observations suggest that the granite intruded into the granodiorite. The granite contains xenoliths of amphibolitic material that locally comprises up to 5 vol.% of the granite. Least-altered samples granite

and microgranodiorite show negative Eu anomalies in chondritenormalised REE plots, La/Gdn at 5.03 and 6.31 respectively, nearhorizontal Tb–Lu slopes (Fig. 12E–F) and Zr contents below 200 ppm (Table 1) — all typical of calc-alkaline igneous rocks (Barrett and MacLean, 1999). 7. Mineralisation In the ore zone, magnetite occurs intimately intergrown with skarn, as veins cutting skarn, as masses veined by skarn (Figs. 15–17) and as porphyroblasts in dolomitic marble together with minor magnesian silicates. The relative proportions of magnetite and skarn vary widely. Massive, fine-grained magnetite with little skarn can grade over a few metres into skarn with minor magnetite. Great lateral variations also exist in the proportions of marble relative to skarn and in the composition of the skarn. These variations are discordant to stratigraphy. Representative mineral chemical data for magnetite, clinopyroxene, garnet and amphibole are provided in Tables 3–6 and whole-rock skarn compositions are presented in Table 7. In the northern part of the Ryllshyttan deposit, most magnetite occurs as blasts in dolomitic marble, accompanied by serpentine, diopside (Di81–89Hd8–16Jhn2–3), tremolite, pyrite, sphalerite, galena, chondrodite and accessory chlorite. Geijer and Magnusson (1944) also report fluoborite (Mg3(BO3) (F,OH)3) and pseudomorphs after ludwigite (Mg2Fe3+BO5), associated with the association chondrodite–magnetite– Fe-poor sphalerite in the marble. Electron microprobe data show that the magnetite blasts are complexly zoned with compositions spanning a broad range within the magnetite–jacobsite–magnesioferrite series, locally with up to 48 mol% and 35 mol% magnesioferrite (MgFe2O3) and jacobsite (MnFe2O3) components, respectively (Table 3). Some blasts occur in strained black patches of intergrown magnetite–serpentine

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227

Table 1 Representative whole-rock lithogeochemical analytical data of least altered samples of Ryllshyttan volcanic and intrusive rocks. Sample No.

20071244

71007

71227

71036

71006

20071212

20071260

71040

Chemical classification

Rhyolite A

Rhyolite B-I

Rhyolite B-II

Rhyolite B-III

Dolerite

Bas. andesite

M.granodiorite

Granite

Affinity

Calc-alkaline

Calc-alkaline

Calc-alkaline

Calc-alkaline

Transitional

Calc-alkaline

Calc-alkaline

Calc-alkaline

Stratigraphic interval

FW

HW

HW

FW–HW

FW–HW

HW

FW–HW

FW

FW/HW

FW/HW

HW

FW/HW

FW

Clastic juv

Fgr clastic

Fgr clastic

Coherent

Coherent

Coherent

Coherent

Coherent

75.13 0.09 11.93 2.72 0.03 2.52 2.46 2.27 2.69 0.007 b0.02 b0.02 0.6 99.85 211 60 17 0.8 0.2 431 73 106 1 4 23 b0.5 b0.1 0.8 0.234 17.72 131.78 3.52 12.44 0.28 52 53 3 b0.1 b0.1 1 18 7 0.01 0.1 b0.1 5 b0.5 4 1.2 0.2 14.7 6.5 b8 0.8 52.11 112.27 14.28 54.83 10.77 1.77 10.5 1.74 10.36 2.17 6.53 1.09 6.54 0.99

80.75 0.09 10.35 0.64 0.03 0.43 0.42 2.9 4.08 0.011 0.09 0.03 0.5 99.72 135 22 12 1.1 0.6 2360 64 35 1 10 15 b0.5 b0.1 0.7 0.15 13.09 114.44 6.1 11.23 0.54 58 44 b0.1 b0.1 b0.1 0 8 5 b0.01 b0.1 b0.1 4 0.5 2 0.9 b0.1 7.9 2.6 b8 b0.5 8.84 33.97 2.4 10.05 2.53 0.39 2.8 0.47 3.98 0.9 2.77 0.45 2.95 0.43

76.51 0.15 12.15 1.71 0.03 0.59 0.83 3.13 4.62 0.008 0.03 0.03 0.9 99.72 257 53 23 1.2 0.4 1878 79 83 1 6 15 b0.5 b0.1 2.4 0.169 21.11 80.27 4.81 11.35 0.42 57 45 3 b0.1 b0.1 0 16 9 b0.01 0.5 b0.1 5 0.6 4 1.6 b0.1 18.5 5.2 8 b0.5 49.85 101.92 13.09 49.95 10.08 1.2 9.54 1.43 9.76 2.16 6.38 1 6.52 0.95

78.52 0.09 10.46 1.96 0.06 0.99 0.88 1.32 5.42 0.023 0.05 0.03 1.7 99.72 137 43 12 1.4 0.7 2387 94 100 1 14 37 b0.5 b0.1 2.2 0.151 13.11 115.22 3.18 11.72 0.27 74 71 b0.1 b0.1 b0.1 1 15 4 b0.01 0.3 0.2 4 0.5 3 0.8 b0.1 11.3 3.2 29 0.8 40.26 81.33 9.8 35.82 7.69 1.31 7.53 1.26 7.44 1.5 4.33 0.61 3.87 0.57

48.02 0.69 14.44 11.94 0.31 9.48 10.41 2.92 1.2 0.166 0.06 0.03 1.2 99.58 31 10 2 44.1 36.8 599 33 174 1 8 38 0.5 b0.1 b0.5 0.005 2.15 20.99 3.2 16.16 0.2 45 65 1 b0.1 b0.1 0 13 1 b0.01 0.1 b0.1 45.5 b0.5 b1 b0.1 b0.1 0.5 0.3 270 b0.5 6.28 12.45 1.78 8.1 1.85 0.72 2.04 0.27 1.92 0.41 1.13 0.18 1.13 0.17

56.86 0.99 15.86 11.23 0.17 3.57 6.23 4.17 0.55 0.205 b0.02 0.06 0.6 99.82 90 21 6 22.8 0.7 201 20 191 15 2 23 0.5 b0.1 b0.5 0.009 5.69 16.08 4.2 14.93 0.28 28 33 2 0.1 b0.1 1 19 3 b0.01 2.8 b0.1 34.2 b0.5 2 0.4 b0.1 6.3 2.1 251 0.9 20.72 45.07 5.63 22.94 4.5 1.33 4.02 0.65 3.76 0.76 2.16 0.35 2 0.31

71.2 0.47 14.48 2.18 0.05 0.92 3.52 6.69 0.27 0.139 0.15 b0.02 1.4 99.93 165 31 10 1.8 0.8 102 9 320 1 4 27 0.6 b0.1 b0.5 0.035 11.41 31.04 5.36 16.14 0.33 10 8 2 0.2 b0.1 0 15 4 b0.01 0.3 b0.1 13.2 0.7 3 0.7 b0.1 11 4.2 27 0.71 40.16 71.6 8.14 30.32 5.72 1.31 5.31 0.88 4.98 1.04 3.13 0.5 3.15 0.49

73.61 0.212 13.31 2.52 0.03 0.31 1.53 3.69 3.86 0.046 b0.02 b0.02 0.7 99.82 186 31.1 11.4 1.4 0.9 999 102 202 3.3 8 14 b0.5 b0.1 2 0.088 14 62.8 5.98 16.32 0.37 44 36 2 0.2 b0.1 0.5 13.5 11.4 b0.01 2.8 b0.1 8 b0.5 2 0.9 b0.1 13.4 5.6 b8 0.7 31.8 64.6 7.55 30.2 5.83 1.09 5.28 0.94 5.29 1.05 3.37 0.54 3.51 0.53

Host unit of sills/dykes Facies group SiO2 TiO2 Al2O3 Fe2O3-tot MnO MgO CaO Na2O K2O P2O5 tot C S LOI Raw. Tot Zr Y Nb Co Ni Ba Rb Sr Cu Pb Zn As Ag Au (Zr/TiO2)/10,000 (Zr/Al2O3)/10,000 Al2O3/TiO2 Zr/Y Zr/Nb Nb/Y Ishik. A.I. Modif. A.I. Be Bi Cd Cs Ga Hf Hg Mo Sb Sc Se Sn Ta Tl Th U V W La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppb

ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm

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10.00

Granite Microgranodiorite Rhyolite porphyry Comendite Pantellerite

Zr/TiO 2 * 0.0001

1.00

Phonolite

Garpenberg regional dolerite sill

Rhyolite

Basaltic andesite

0.10

Dacite Rhyodacite

Trachyte

Upper hanging wall - red-grey

Trachy andesite

Lower hanging wall - interbedded rhyolitic siltstone/sandstone and aluminous skarn

Andesite

0.01

Andesite / basalt

Alkali basalt

Basanite nephenite

Footwall - Poorly-sorted Qz+Fspphyric rhyolitic breccia-sandstone

Sub-alkaline basalt

0.00 0.01

0.1

1

10

100

Nb/Y Fig. 9. Least-altered sampled of volcanic rocks from Ryllshyttan plotted in the classification diagram of Winchester and Floyd (1977). Granite and microgranodiorite from the Ryllshyttan area as well as the regional dolerite sill in the GSI has been plotted for comparison. ‘Footwall’ and ‘hanging wall’ here refers to stratigraphic position relative to the limestone host.

(Fig. 15A–B). The intergrowths commonly define round shapes when viewed parallel with the stretching lineation, interpreted as pseudomorphs after chondrodite or olivine. The magnetite-impregnated dolomitic marble of northern Ryllshyttan grades laterally and irregularly to the south into magnesian skarns of variable magnetite-content, dominated by diopside (Di90Hd9– Di92Hd7) and tremolite/actinolite (Mg/(Mg + Fe + Mn) = 0.87–0.89). These skarns become dominant in the western Kompani workings (Figs. 3–4). Most magnetite occurs with tremolite and serpentine whereas diopside-dominant skarns carry less magnetite. In the western Kompani workings, fine-grained massive diopside skarn is veined by irregular intergrowths of coarser magnetite–diopside and magnetite– tremolite (Fig. 16A–B). These veined skarns grade locally into Cadepleted masses of magnetite accompanied by serpentine, chlorite, phlogopite, calcite and accessory fluorite. These masses form some of the richest iron oxide mineralisation. They are commonly transected by veins of sphalerite–galena ± phlogopite (locally chloritised) ± serpentine and occur near shear zones developed in chloritised mafic intrusions (Fig. 16C). Both the dolomitic marble and the magnesian skarns show gradations to 0.1–1 m-thick zones of massive ferroan diopside (Di57–64Hd31–41Jhn5–6) skarn with minor magnetite, allanite, calcite and pyrite at the original stratigraphic contacts of the marble, particularly the upper contact. The ferroan diopside skarns are generally magnetite-poor except where they are cross-cut by veins of more hydrous skarn. The veins locally host chondrodite with an inner selvage of serpentine grading to tremolite (Mg/(Mg + Fe) = 0.97–1) and outer selvage of actinolite and magnesio-hornblende (Mg/ (Mg + Fe) = 0.56–0.57) (Fig. 15D). The magnetite content is high in the tremolitic parts. Minor massive epidote skarn with calcitefilled vugs, lined by adularia and actinolite with accessory fluorite, are locally present (Fig. 15C). Similar magnetite-poor ferroan diopside (Di48–63Hd33–49Jhn3–4) skarn, but with associated andradite (Adr78Gr16Al4Sp2), characterises the eastern Kompani workings (Fig. 17A–C). The eastern Kompani skarn is dominated by ferroan diopside, which occurs in fine-grained massive aggregates showing a granoblastic polygonal texture and as coarse-grained bladed or prismatic crystals (Fig. 17D). Associations of granoblastic polygonal ferroan diopside, andradite, epidote and quartz replace or overgrow the earlier bladed ferroan diopside (Fig. 17C). Magnetite, quartz, calcite, apatite, epidote, sphalerite, and relics of bladed clinopyroxene form inclusions in andradite. Magnetite generally

correlates negatively with garnet in the eastern Kompani workings. Locally, andradite forms 2–5 mm rims rounding quartz patches interpreted as former druses (Fig. 17A). Here, needle-like euhedral actinolite crystals and euhedral growth-zoned epidote crystals have nucleated on andradite and ferroan diopside and grown freely into the former vugs. Magnetite is mainly intergrown with the fine-grained ferroan diopside whereas only minor magnetite occurs as inclusions in the bladed ferroan diopside crystals. The bladed and granoblastic ferroan diopside share the same composition. The sulphide mineralisation partly cross-cuts the folded pattern defined by skarn assemblages and magnetite distribution (Fig. 5). Sulphides are not restricted to the limestone and skarn; shoots of massive sulphide propagate into the footwall and hanging wall, parallel to the dominant tectonic foliations (Fig. 5). The richest concentrations of massive sulphide occur along the sinistrally sheared short limbs of F2 folds, the largest being the lens along the ENE-trending NW limb and hinge of the Kompani anticline (Fig. 4). In contrast, sulphide mineralisation is minor or absent in the NNE-trending long limbs of the folds. Furthermore, sulphides are concentrated in F2 fold hinges in both the Silver and Tombak workings (Fig. 3). For these reasons, it appears that the location of the mined massive sulphide bodies is largely controlled by D2 and D3 structures. A crude deposit-scale metal zonation can be defined from the sphalerite-dominated Kompani section with less galena and minor chalcopyrite (south) to the argentiferous galena–sphalerite dominated Silver section (north; Fig. 3). Concurrent with this zonation is a northward decrease in skarn replacement intensity of the limestone host (Fig. 3). In northern Ryllshyttan, sphalerite and galena form impregnations in dolomitic marble, accompanied by minor tremolite, chondrodite and serpentine. In the skarn of the Kompani workings, vein networks of sphalerite with minor serpentine and chloritised phlogopite overprint massive magnetite in the most Ca-depleted magnesian skarn (Figs. 16C, 18A–C). The serpentine forms pseudomorphs after tremolite, diopside and chondrodite or occurs as fibrous masses. Anomalous F contents (0.75–1.2 wt.%) have been detected by electron microprobe in both types of serpentine. Fluorite is observed along cleavage planes in elongate serpentine grains. Magnusson (1970) reported fluorite and quartz as major gangue minerals in the Ryllshyttan sulphide ore. The margins of chloritised phlogopite crystal bundles against sphalerite are commonly rich in small intergrowths of magnetite and pyrite. Locally, euhedral pyrite crystals are nucleated on the surfaces of chloritised phlogopite (Fig. 18B). The massive sphalerite–galena

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(m) 500

Mud 0.5 2 8 32 mm

Granite (1895 +/- 3 Ma) Microgranodiorite (1894 +/- 4 Ma)

229

form inclusions in the chalcopyrite. The chloritised phlogopite flakes are commonly randomly oriented, yet show slight kinking of individual crystals. Magnetite only rarely occurs in these post-skarn sulphide veins and where present, is always observed as complex intergrowths with pyrite (Fig. 18E).

Mafic intrusion

400

Rhyolite porphyry (1895 +/- 4 Ma)

8. Alteration

Interbedded rhyolitic siltstone/ sandstone-aluminous skarn Red-grey rhyolitic siltstone/ fine sandstone

8.1. Alteration of igneous rocks and effects and relative timing of metamorphism

Dolomitic limestone Poorly sorted Qz+Fsp-phyric rhyolitic breccia-sandstone 300

Skarn Biotite/phlogopite-rich alteration zone Massive sulphide mineralization Magnetite mineralization Garnet porphyroblast Albite alteration

200

100

0 Fig. 10. Schematic graphic log of the Ryllshyttan stratigraphy. From Jansson and Allen (2011b). SIMS U–Pb zircon ages of samples collected in the Ryllshyttan area by Jansson and Allen (2011a) are indicated.

mineralisation commonly contains inclusions of magnetite and pyrite, as well as magnetite–pyrite intergrowths without any associated chloritised phlogopite (Figs. 13E, 18B). Magnetite mineralisation has not been observed to overprint sphalerite- or galena-rich mineralisation in hand-specimen. Grain-boundaries of sphalerite crystals within massive sphalerite commonly show evidence of static recrystallisation with 120° grainboundaries that are commonly lined by small crystals of pyrrhotite. Where galena occurs as a major phase in massive sulphide, it occurs as a matrix to porphyroblastic sphalerite crystals to which it shows incompetent grain-boundaries (Fig. 18A). In the eastern Kompani workings, only minor sphalerite occurs in association with the andradite, quartz and pyrite that overgrow bladed clinopyroxenes. Abundant later chalcopyrite–sphalerite–magnetite– pyrite veinlets transect these clinopyroxene–garnet skarns. In the chalcopyrite-dominated veins, sphalerite inclusions are rimmed by aggregates of chloritised phlogopite, andradite, diopside, allanite and minor apatite and molybdenite. As in the massive sphalerite ore, sphalerite carries abundant sulphide inclusions, yet here they are dominated by chalcopyrite. More rarely, euhedral laths of molybdenite and gypsum

Several irregular and discordant alteration zones have been mapped (Figs. 3, 4 and 5). Proximal to sulphide mineralisation, volcanic rocks and intrusions (except for the microgranodiorite) are altered to phlogopite/biotite ± sericite ± quartz schists (K–Mg–Fe ± Si alteration) with almandine porphyroblasts (Alm50–73Sps9–28Grs9–11Prp7–13Adr0–3) and local cordierite (Fig. 19A). Locally, the almandine porphyroblasts are amalgamated to large pods, which are F2 folded. The S1 foliation locally occurs as a planar quartz grain-shape fabric in xenoblastic garnet. Apatite, microcline, albite, ilmenite and prehnite are accessory phases. Zones of this strong K–Mg–Fe ± Si alteration propagate high up into the stratigraphic hanging wall along hydrothermally altered mafic intrusions. Immediately adjacent to the dolomitic marble, and in the stratigraphic hanging wall, the biotite/phlogopite schists locally grade into almandine porphyroblastic pargasite rocks. Strong muscovite–quartz (K ± Si) alteration occurs immediately stratigraphically below the Silver workings and weak to moderate muscovite–quartz alteration occurs more distal to the K–Mg–Fe ± Si altered zones. All mica schists and associated pargasite rocks are enriched in similar sulphide minerals as those found in the massive sulphide lenses. The variable skarn mineralogy is also reflected in alteration of the mafic intrusions (Fig. 2). Mafic intrusions that crosscut magnesian skarns are commonly altered to magnetite–porphyroblastic chlorite schists, whereas those that crosscut calcic skarns are strongly epidote-altered. Moderate to strong oligoclase–diopside ± actinolite ± epidote and oligoclase–epidote ± actinolite (Na–Ca) alteration have affected the microgranodiorite and volcanic rocks, especially within 20 m from the microgranodiorite contact with the volcanosedimentary rocks. Despite the presence of diopside and tremolite, mass-balance calculations suggest that no Mg has been added to the rock (Jansson, 2011). This alteration ranges from pervasive to vein style and primary rock textures are locally preserved. The altered rocks are characterised by white to pink rocks carrying quartz, oligoclase, diopside (Di71–73Hd25–27Jhn2), actinolite (Mg/(Mg + Fe) = 0.76–0.80), minor K feldspar (Or92Ab6Cls3) and epidote. Actinolite appears to have grown at the expense of diopside and oligoclase. Locally, such as in the rhyolitic siltstones northeast of the Silver and Tombak workings, 1–10 mm mottles of actinolite–epidote–clinopyroxene in oligoclase-altered matrix are common. The mottles are strained into ellipsoidal shapes along the S2 foliation, and actinolite veinlets are folded by F2. Textural relationships suggest that the Na–Ca alteration overprints the K–Mg–Fe ± Si alteration (Fig. 19B). Concretion-like pods of epidote and/or calcic clinoamphibole are abundant in the red-grey rhyolitic siltstone–sandstone member (Fig. 11C). These may represent metamorphosed iron-rich carbonate concretions (cf. Allen et al., 2003). Micaceous alteration styles (phlogopite, biotite, muscovite) in the volcanic rocks are overprinted by S1. This along with the fact that the altered zones are cross-cut by the microgranodiorite, suggest that alteration was pre-D1 and M1. Phlogopite/biotite schists are interpreted as metamorphosed chlorite–sericite alteration zones (cf. Trägårdh, 1991; Allen et al., 1996, 2003). The pargasite rocks have immobile element ratios distinct from mafic intrusions, but similar to the rhyolitic siltstones. They are interpreted as more calcareous equivalents of the phlogopite/ biotite schists. They formed at the contacts between these schists and dolomitic marble and where K–Mg–Fe ± Si alteration overprinted

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Fig. 11. Main stratigraphic units of Ryllshyttan. A. Poorly-sorted quartz + feldspar phyric rhyolitic breccia–sandstone. One chlorite-altered, former pumice clast with an irregular shape and wispy margins has been outlined. B. Interbedded rhyolitic siltstone/sandstone–aluminous skarn. C. Red-grey rhyolitic siltstone–fine sandstone with elliptical actinolite–epidote alteration spots.

calcareous iron formation that was precursor to the hanging wall skarn beds (Jansson and Allen, 2011b). Regional metamorphic minerals are here regarded as prograde minerals that are clearly syn- to post-S1. The almandine porphyroblasts of the altered volcanic rocks vary in style from euhedral with inclusionrich rims (Type 1) to euhedral crystals enclosing a planar quartz grain-shape fabric (Type 2) to strongly xenoblastic, ragged, anhedral crystals overgrowing a folded quartz–mica foliation (Type 3). Foliations are generally deflected around the first two types. For Type 2 the external S2 foliation is commonly not continuous with the internal foliation. These observations suggest that Type 2 garnets are M1 garnets with an enclosed S1 foliation and a deflected S2 foliation, and that M1 outlasted D1. It moreover suggests that the D2 deformation was not continuous with D1 but separated by a tectonic hiatus in which almandine porphyroblasts overgrew a passive tectonic fabric. The Type 1 garnets are pre-tectonic relative to S2 but their relationship to S1 is unclear. The lack of an internal foliation may indicate that Type 1 garnets are wholly pre-tectonic and that their inclusion-rich outer rims represent later, syn-tectonic M1 overgrowths. The ragged xenoblastic Type 3 garnets are interpreted as syn- to post-D2 in timing as they overgrow folded S1. Along with statical recrystallisation of S1/S2 crenulation as well as the annealed texture and local presence of unaligned micas in massive sphalerite–galena ore (Fig. 18B–C), they suggest that M2 regional metamorphism under amphibolite facies conditions outlasted D2.

At the Stormoss and Smältarmossen deposits in the Garpenberg supracrustal inlier (Fig. 2), sulphide-poor magnetite deposits hosted by andradite ± ferroan diopside skarns grade directly into calcite marble without any intermediate zone of dolomite or magnesium-rich skarn. In contrast, the mineralised diopside- and tremolite-rich skarns of the Garpenberg Norra sulphide deposit grade outwards into dolomitic marble, which in turn grades into calcitic marble. Allen et al. (2003) interpreted the dolomitisation as an early Mg-rich alteration style in a zoned sulphide ore-forming system, prior to the addition of significant silica to the marble. At Ryllshyttan, the early skarns pervasively replace the marble with fine-grained diopside (magnesian skarn) or bladed ferroan diopside (calcic skarn). Based on the regional pattern outlined above, it is inferred that a major control on whether ferroan diopside or diopside skarns developed was the extent of pre-skarn dolomitisation. The earliest magnesian and calcic skarns are thus viewed as the products of Fe–Si alteration overprinting dolomitic and calcitic marble respectively. The magnetite-impregnated dolomitic marble of the Silver workings reflects the earliest stage of Fe–Si alteration of dolomitic marble, forming some of the earliest magnetite mineralisation (Fig. 15A), whereas the skarns of the Kompani workings reflect more advanced alteration, leading to complete replacement by skarn and magnetite.

8.2. Alteration of limestone

Subsequent to their formation, the skarns underwent alteration that was not isochemical with respect to Fe, Mg, Si, S and base metals. The zones of strong retrograde alteration of earlier, anhydrous skarns largely coincide with the zones with the most massive occurrences of massive sulphides and magnetite along e.g., the shear and attenuated short limbs of the F2 folds. In the dolomitic marble, the serpentine–magnetite intergrowths and observations of serpentine replacing earlier chondrodite and diopside suggest that some magnetite formed by retrograde alteration of an earlier Mg–Fe–Mn-silicate (Fig. 15B). The precursor to the magnetite– serpentine intergrowths may have been Fe–Mn-rich chondrodites or possibly olivine belonging to the forsterite–fayalite–tephroite series of Mossman and Pawson (1976).

Diopside, tremolite and magnetite-bearing magnesian skarns predominate in the interpreted most central part of the alteration system at Ryllshyttan (defined as the western/footwall side of Kompani). These magnesian skarns grade eastwards and stratigraphically upwards into ferroan diopside–andradite-bearing calcic skarns. The skarns in the Kompani area have nearly completely replaced the former limestone, whereas northwards, laterally along strike of the ore host towards the Tombak and Silver areas, the skarns grade into magnetiteimpregnated dolomitic marble. At the Tombak and Silver workings, calcic skarns similar to those of eastern Kompani (except for a lack of andradite) have formed at the limestone contacts.

8.3. Alteration of the skarns

N.F. Jansson, R.L. Allen / Ore Geology Reviews 69 (2015) 217–242 1000

1000

Rock/Chondrite

Poorly-sorted Qz+Fsp-phyric rhyolitic breccia-sandstone (n = 10)

A 100

10

10

1 La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

1000

1000

Mafic intrusion (n = 11)

Rock/Chondrite

B

Rhyolite porphyry (n = 5)

100

1

Basaltic andesite (n = 2)

C

100

100

10

10

D

1

1 La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

1000

La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

1000

100

100

10

10

1

F

Granite (n = 1)

E

Microgranodiorite (n = 9)

Rock/Chondrite

231

1 La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

Fig. 12. Chondrite-normalised whole-rock rare earth element data for juvenile volcanic rocks and intrusive rocks at Ryllshyttan. A: Rhyolitic breccia–sandstone, B: Rhyolite porphyry, C: Mafic intrusion, D: Basaltic andesite, E: Microgranodiorite, F: Granite. The black lines denote the pattern of the least-altered sample, whereas the shaded areas denote the range in altered equivalents, except for in ‘D’, where only two, unaltered samples exist, and ‘F’, where only one, unaltered sample exists. The REE data was corrected to a volatile-free basis before normalisation against chondritic values according to Boynton (1984).

The brecciation and veining of the fine-grained diopside skarns of the western Kompani workings record an influx of Fe into previously un-mineralised diopside skarn (Fig. 16A). This occurred prior to the development of at least one tectonic foliation, though the timing of this foliation (i.e., S1 or S2) has not been possible to constrain. The relationship of the extensively veined diopside skarn to the adjacent serpentine-rich massive magnetite and cross-cutting sphalerite mineralisation (Fig. 16C) is obscured by sheared contacts. However, the textural evidence outlined earlier (e.g., Fig. 15A–B, D) suggest that the development of high-grade, serpentine-rich magnetite mineralisation likely resulted from retrograde alteration of mineral associations containing clinopyroxene and chondrodite. Fine-grained and granoblastic ferroan diopside, magnetite, andradite, quartz, epidote, allanite, actinolite and locally sulphides (pyrrhotite, pyrite, and sphalerite) overgrow and replace paragenetically early bladed ferroan diopside in the eastern Kompani workings, which indicates that ore minerals partly developed at the expense of the bladed ferroan diopside skarn. The local vug-filling textures exhibited by

these mineral associations, as well as the virtually undeformed euhedral crystals of actinolite and epidote within the vugs suggest that these mineral associations post-date the main stages of deformation. Alternatively, these minerals might have escaped ductile deformation due to their position within competent ferroan diopside–andradite skarn. 8.4. Late stage veinlets Skarn formation and alteration of the skarns was followed by the development of straight, late-stage veinlets. The veinlets do not vary in style but vary in composition depending on which lithology they cut. The veins that crosscut diopside-rich skarns generally contain tremolite, calcite, quartz, sphalerite, pyrite and magnetite (Fig. 16B), whereas veins that crosscut ferroan diopside ± andradite skarns commonly carry actinolite, quartz, chalcopyrite, pyrrhotite, sphalerite, pyrite, andradite, allanite, phlogopite, magnetite and minor molybdenite (Fig. 17B, D, G). Yellow serpentine ± calcite veinlets crosscut the

232

Table 2 Characteristics of intrusive rocks at Ryllshyttan. Mineralogy1

Granite

~30–35% Qz, ~60% Mc N Pl, Calc-alkaline 6.8 5–10% Bt N Hbl. Accessory Tit, Ap, Zrn

Micro-granodiorite 65–70% Pl (oligoclase), 20–25% Qz, 1–3% Hbl, 1–3% Cpx. Accessory Zrn, Ap and Ttn Dolerite Hbl, Pl, Ep, Grt

Basaltic andesite

Rhyolite porphyry

1 2

Affinity

Ti/Zr

Calc-alkaline 16.9–19.8

Transitional

Texture

Geometry and contact relationships

Alteration2

Interpretation

Equigranular with 1–3 mm crystals. distinct foliation defined by mafic phases. Pinkish-red colour Weakly-moderately Pl porphyritic. Grey-pink colour

Direct contact with supracrustal package or microgranodiorite has not been observed

1: Minor selective Ep-replacement of Ca-rich Pl-cores. 2: Weak sericitisation of Pl. 3: Weak chloritisation of mafic minerals 1: Pl–Di ± Ep ± Cam alteration. 2: Pl–Ep ± Cam alteration

Deeper granitic margin of early Svecokarelian granitoid. Pre-D2 emplacement

1: Fracture bound Phl/Bt ± Qz N sulphides (non-destructive to Qz-phenocrysts, colour shift from pink to streaky grey-white). 2: Fracture bound weak (texturally conservative) to intense (texturally destructive) Prg ≫ Alm–Ser–Phl/Bt alteration of groundmass and phenocrysts. 3: Pl–Di ± Ep ± Cam alteration. 4: Pl–Ep ± Cam alteration

Early (pre-lithification) intrusion in footwall and hanging-wall. Syn-volcanic intrusion. Pre-D1-emplacement

122.6–215.9 Massive, amygdaloidal. Weakly Pl-porphyritic.

40% Pl, 20–25% Qz, 20–25% Calc-alkaline 65.6–71.1 Hbl, 5–10% Bt, 1–2% Ttn, 0–5% Py Calc-alkaline 3.8–4.0 Phenocrysts (15%) = 10% Qz, 5% Mc. Groundmass (85%) = Qz N Mc, 1–3% chloritised Bt or Hbl. Accessory Zrn

Represents metamorphic mineralogy observed in least altered samples. Represents currently observed metamorphic mineralogy in altered samples.

Discordant to semi-conformable boundaries. Diffuse to sharp margins towards footwall and hanging-wall supracrustal rocks. Truncates ore host Discordant dikes, larger irregular intrusive bodies and subordinate sills. Generally razor-sharp margins towards rhyolitic volcaniclastic rocks

Fine-grained aphyric.

Similar to dolerite

15% 2–3 mm Qz + Mc phenocrysts in 0.05–0.1 mm aphanitic groundmass

Intrudes footwall as well as hanging-wall. Peperitic facies observed towards intruded volcaniclastic rocks.

Granodioritic margin of early Svecokarelian granitoid. Co-magmatic with higher level dacites in the GSI. Pre-D2-emplacement. Relatively shallow, weakly vesiculated 1: Weak–strong pervasive texturally conservative Ep/Cz of groundmass and intrusions emplaced into indurated amygdales. 2: Weak-intense texturally supracrustal wall-rocks. May be syn-volcanic with stratigraphically destructive Phl/Bt ± Crd. 3: higher basalt. Pre-D1-emplacement. Weak-intense texturally destructive Chl–Mag ± sulphides Similar to dolerite Pre-D1 emplacement

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N.F. Jansson, R.L. Allen / Ore Geology Reviews 69 (2015) 217–242

233

Fig. 13. Intrusive facies of the Ryllshyttan area. A. and B. Peperitic rhyolite porphyry intruded into the lower hanging wall. C. Mafic dyke discordantly crosscutting interbedded rhyolitic ash–siltstone and stratiform aluminous skarn. D. Irregular sharp margin between the dyke and interbedded rhyolite siltstone and stratiform aluminous skarns. E. Laminated xenolith of green clinoamphibole– plagioclase altered rhyolitic siltstone in margin of microgranodiorite.

display little to no strain. The veins are interpreted as late- or postD 2 retrograde regional metamorphic veins (Fig. 17B, G). The sulphide-oxide mineral associations within them are interpreted as D 2–D 3 remobilisation of pyrrhotite–chalcopyrite and pyrrhotite– sphalerite where the pyrrhotite has later been gently oxidised to pyrite + magnetite intergrowths (cf. Ramdohr, 1980).

magnetite-impregnated dolomitic marble, magnetite–serpentine skarns and the sphalerite–galena veins. The late stage veins thus reflect the chemistry of the transected lithologies but carry more hydrous mineral associations, suggesting this alteration stage was retrograde and largely isochemical at deposit scale. The veins cross-cut or are subparallel to tectonic foliations and euhedral crystals within the veins 40

2.0

A

30

B

Zr/Al 2O3

Nb/Al2O3

1.5

20

1.0

10

0.5

0

0.0 0

50

Al 2O3/TiO2

100

Granite Microgranodiorite Rhyolite porphyry

Basaltic andesite Garpenberg regional dolerite sill

150

0.0

0.1

Upper hanging-wall - red-grey rhyolitic

Zr/TiO 2

Lower hanging-wall - interbedded rhyolitic siltstone/sandstone and aluminous skarn

0.2

0.3

Footwall - Poorly-sorted Qz+Fsp-phyric rhyolitic breccia-sandstone Regional dacite porphyry intrusions

Fig. 14. Immobile element box-plots for igneous rocks from Ryllshyttan. A: Zr/Al2O3 vs. Al2O3/TiO2. B: Nb/Al2O3 vs. Zr/TiO2. Small black points denote dacite porphyry intrusions close to the stratigraphically higher Smältarmossen and Garpenberg Norra deposits (Fig. 2). ‘Footwall’ and ‘hanging wall’ here refers to stratigraphic position relative to the limestone host. The large scatter in the silty–sandy rhyolitic rocks from the hanging wall compared with intrusive rocks and footwall is interpreted as, reflecting a higher degree of post-eruptive reworking.

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9. Discussion 9.1. Timing of alteration and mineralisation Lindroth (1921, 1924) interpreted sulphide mineralisation, granitoid emplacement and the magnesian skarns at Ryllshyttan as coeval and suggested that a considerable hiatus existed between these events and an earlier stage of magnetite formation, and that the mafic intrusions were emplaced during this hiatus. This model was challenged by Geijer and Magnusson (1944) who interpreted the magnetite as also related to granitoid emplacement and suggested that no significant hiatus existed between iron oxide and sulphide mineralisation. Geijer and Magnusson (1944) related the formation of calcic skarns to sulphide ore formation whereas Fe ores were related to the magnesian skarns. The mafic intrusions were interpreted to post-date the ores and the granitoid, yet pre-date Mg-alteration. These relationships could not be verified during the current study. Instead, our mapping shows that the mafic intrusions are strongly K–Mg–Fe ± Si-altered adjacent to magnesian skarns and K–Mg–Fe ± Si-altered volcanic rocks, whereas the microgranodiorite is unaffected by this alteration style (Figs. 20, 21) and cuts the alteration envelope (Fig. 3). Consequently, the K–Mg– Fe ± Si alteration must have occurred during or after the emplacement of the mafic intrusions but before the emplacement of the microgranodiorite (Fig. 20). The mafic intrusions described by Geijer and Magnusson (1944) may represent a younger generation of intrusions than those observed at Ryllshyttan. The irregular distribution and the textures of magnetite suggest that iron was introduced epigenetically into the limestone. In contrast, the laminated magnetite rocks interbedded with the volcaniclastic rocks overlying the limestone host suggest that iron-rich sediments formed immediately after deposition of the limestone but prior to the epigenetic Fe mineralisation (Jansson and Allen, 2011b). Consequently, at least two distinct generations of iron oxides are recognised; syngenetic iron oxides deposited simultaneously with the volcaniclastic sediments, and epigenetic magnetite deposited after significant burial of the stratigraphic succession and emplacement of intrusions (Fig. 20). In addition, the paragenetic relationships suggest that epigenetic magnetite formation was not a single event. Instead, magnetite appears to have formed both during early hydrothermal alteration (e.g., magnetiteimpregnated dolomitic marble; Fig. 15A), and as a retrograde phase replacing anhydrous skarn minerals and chondrodite (e.g., magnetite– tremolite and magnetite–serpentine assemblages; Figs. 15A, B, D, 16C). Although sulphides consistently occur intergrown with magnetite in retrograde skarns or crosscut magnetite-bearing skarns, it questionable whether these relationships reflect the original timing of mineralising events or the preferential remobilisation of more ductile sulphides into more competent magnetite during deformation (cf. Marshall and Gilligan, 1987). The latter would yield a false impression of genetically younger, overprinting sulphide mineralisation. The strong regional association of sulphide mineralisation to pre-D1 sub-seafloor K–Mg– Fe ± Si alteration (Fig. 21) in the entire GSI, and the association of Fe skarn deposits to the younger phase of dacitic/granodioritic intrusions (Jansson and Allen, 2013) would rather suggest the opposite. Microtextural, mineralogical and structural observations for the late stage skarn veins and textures in the altered skarns suggest that part of the controversy results from syn-metamorphic (re)mobilisation during deformation. Pre-D1 sulphide-mineralisation was (re)mobilised during D2–D3 resulting with massive sulphide bodies now occupying Fig. 15. Skarns and marble of the Silver segment. A: Magnetite (Mag) and serpentine (Srp) impregnated dolomitic marble. B: Dolomitic marble with magnetite impregnation and blotchy patches of black serpentine, diopside (Di), magnetite and chondrodite. C: Epidote (Ep) and actinolite (Act) skarn with calcite (Cal) filled vug lined by hematite-stained adularia (Adl). D: Ferroan diopside (Cpx) skarn cut by chondrodite (Chn) vein with 2 cm selvage of inner serpentine, intermediate tremolite (Tr) + magnetite and outer actinolite. E: Clinopyroxene skarn partly retrograded to actinolite with later cross-cutting actinolite– calcite veinlet. Note offset of earlier actinolite replacement texture.

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235

Fig. 16. Skarn and ore of western Kompani. A: Fine-grained diopside skarn overprinted by magnetite. Both are cross-cut by late-stage planar veins of tremolite–quartz and tremolite– calcite. Hd denotes hedenbergite component in diopside in mol%. B: Transmitted light thin-section view of picture A, showing diopside–magnetite skarn cross-cut by a late-stage tremolite–calcite veinlet. C: Various stages of sphalerite–galena overprinting massive magnetite–serpentine rock.

the sheared short limbs and hinges of F2 folds. The similar retrograde mineral associations in high-grade magnetite and sulphide ores suggest that sulphide (re)mobilisation was facilitated by fluids flowing along the ENE-trending shear zones that altered anhydrous skarns to magnetite-rich hydrous skarns rich in tremolite and serpentine. The timing of intrusive events relative to the early epigenetic ore formation is similar at Ryllshyttan and the Garpenberg deposits. Consequently, these deposits likely formed at a similar time but at different

stratigraphic levels, during the development of a large, marine felsic caldera complex (cf. Allen et al., 2003; Jansson and Allen, 2011a,b). 9.2. Processes of ore formation The epigenetic and stratabound character of ore formation is well documented (Figs. 3–5), the timing of alteration and mineralisation relative to volcanism, sedimentation, intrusions and tectonic fabrics is

Fig. 17. Skarns and marble of Eastern Kompani. A: Massive bladed ferroan diopside skarn with patches of andradite (Adr), locally lining quartz-filled vugs. Grs denotes grossular component in andradite. B: D3-sheared quartz–calcite vein with schlieren of sphalerite (Sph), pyrrhotite (Po), pyrite (Py) and actinolite. C: Transmitted light thin-section view of massive andradite–ferroan diopside skarn with andradite replacing earlier bladed ferroan diopside (relics of one optically continuous crystal is outlined). D: Bladed ferroan diopside skarn with minor andradite. E: Coarse-grained clinopyroxene–epidote skarn with inter-crystal accumulations of pyrrhotite + pyrite + actinolite and pyrrhotite + pyrite + actinolite corroding and replacing ferroan diopside. F: Highlight of inter-crystal pyrrhotite + pyrite + actinolite in E. G: Transmitted light thin-section view of retrogressive micro-shear hosting actinolite, chalcopyrite (Ccp), pyrite and magnetite cross-cutting and corroding coarse-grained ferroan diopside skarn.

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Table 3 Representative compositions of magnetite from dolomitic marble. Sample no.

160.6@52

160.6@52

160.6@52

Setting

P.blast

P.blast

P.blast

Facies

2a

3

4

Point

p8

p19

p6

SiO2 Al2O3 FeO-tot MgO CaO TiO2 MnO Cr2O3 ZnO Total

0.06 1.43 80.64 3.83 0.06 0.03 8.87 0.03 0.05 94.99

0.04 1.19 81.20 1.25 0.03 0.07 11.14 0.00 0.21 95.12

0.02 0.19 94.85 0.01 0.05 0.01 0.08 0.02 0.04 95.28

0.05 0.00 1.95 0.00 0.57 0.35 0.07 0.01 0.00

0.01 0.00 1.99 0.00 0.99 0.00 0.00 0.00 0.00

Number of ions on the basis of 3 cations Al3+ 0.06 Cr3+ 0.00 Fe3+ 1.94 Mn3+ 0.00 Fe2+ 0.51 Mn2+ 0.27 Mg 0.21 Zn 0.00 Ca 0.00

moderately well constrained and a relationship to the development of a large marine volcanic–magmatic system appears evident. However, the processes by which the ores formed are more difficult to resolve due to the effects of subsequent metamorphism and deformation. Ryllshyttan displays features of both metasomatic intrusion-related skarn deposits and metamorphosed sub-seafloor VMS deposits in which metals were precipitated in a limestone trap. Relevant to both models is the documented intense volcanism/magmatism and the development of a caldera (Allen et al., 2003), which is inferred to have been beneficial for the development of a large, metal-rich hydrothermal system. The chemical changes involved in K–Mg–Fe ± Si and K–Si alteration, the immobility of Zr, Al, Ti and Nb and the metamorphic mineral

associations of the altered volcanic rocks are similar to metamorphosed chlorite–sericite and sericite–quartz alteration zones around felsichosted VMS deposits (e.g., Ishikawa et al., 1976; Barrett and MacLean, 1994, 1999; Schardt et al., 2001). Chlorite–sericite alteration is driven by moderately acid hydrothermal fluids at temperatures of 250– 300 °C (Schardt et al., 2001) and is not typical of metasomatic Zn and Fe skarn deposits (Larry Meinert, pers. comm. 2009). Formation of anhydrous skarn minerals such as pyroxene and andradite from hydrothermal solutions in the amounts observed at Ryllshyttan would necessitate higher temperatures unless fluid/rock ratios were very high and CO 2 was efficiently removed (cf. Pattison and Tracy, 1991; Meinert et al., 2005). A relatively low fluid temperature could account for the low Cu-content of the sulphide ores, although the high Zn–Pb vs. Cu content could also reflect the dominantly felsic composition of the volcanic successions in Bergslagen (cf. Allen et al., 1996). Compared with VMS deposits, the stratigraphic succession at Ryllshyttan reflects an anomalously shallow marine environment. The shallow marine environment was however important in allowing formation of a stromatolitic limestone that became the trap to mineralisation. Much of the observed magnetite and sulphides at Ryllshyttan show textural relationships indicating deposition together with hydrous silicates (amphibole–serpentine) after formation of the earliest anhydrous skarn minerals. The link between strong retrograde alteration and mineralisation is a typical feature of metasomatic skarn deposits (Dilles et al., 2000; Einaudi, 2000; Meinert et al., 2005), albeit at Ryllshyttan, retrograde skarn formation appear to be a later phenomena associated with D2–D3 deformation. Ryllshyttan is also similar to other skarn deposits in that the magnesian skarns occupy a central, proximal part of the mineralisation, and that these grade upwards into calcic skarns of more Fe-rich composition (Fig. 20). The gradation from sphalerite mineralisation associated with skarns to galena–sphalerite mineralisation directly in marble (Fig. 20) is common to many Zn skarn deposits (cf. Meinert, 1987). However, unlike typical Zn skarn deposits, no systematic increase in Fe and Mn content and no decrease of garnet to pyroxene ratio are developed towards the skarn front at Ryllshyttan (cf. Meinert, 1987). The Ryllshyttan deposit also differs from Zn skarns in the less Mn-rich and more Mg-rich composition of the skarns; zinc skarns may contain N50 mol% johannsenite component

Table 4 Representative compositions of clinopyroxene from Ryllshyttan. Sample no.

071060

071056

071061

20071218

109@63

TSRY02

160.6@52

Setting

Main mass

Main mass

Main mass

Main mass

Main mass

Main mass

Main mass

Facies

Inter-laminated Adr–Cpx–Ep N Mag rock

Inter-laminated Mag–Adr–Cpx rock

Massive andradite rock

Na–Ca altered footwall rhyolite

Calcic skarn

Magnesian skarn

Dolomitic marble

Point

p9

p3

p3

p2

p1

p1

p6

SiO2 Al2O3 FeO MnO MgO CaO Na2O Total

50.54 0.62 13.49 2.19 8.97 22.85 0.61 99.28

50.48 0.83 15.51 1.54 7.83 22.45 0.62 99.26

51.85 0.63 11.65 2.01 10.18 23.07 0.79 100.18

53.43 0.23 7.98 0.53 13.14 24.62 0.19 100.12

52.44 0.19 11.05 1.04 10.91 23.93 0.36 99.92

54.76 0.04 2.39 0.39 17.21 25.94 0.02 100.75

52.91 0.25 3.43 1.06 16.54 25.47 0.11 99.77

1.99 0.01 0.01 0.23 0.02 0.73 0.98 0.01 4.00

1.99 0.01 0.04 0.31 0.03 0.62 0.97 0.03 4.00

1.98 0.00 0.04 0.03 0.01 0.93 1.00 0.00 4.00

1.94 0.01 0.11 0.00 0.03 0.90 1.00 0.01 4.00

Number of ions based on the Wood & Banno Norm (stoichiometric, 4 cations, 12 charges) Si 1.95 1.97 1.97 Al 0.03 0.04 0.03 3+ Fe 0.11 0.07 0.10 Fe2+ 0.33 0.43 0.27 Mn 0.07 0.05 0.06 Mg 0.52 0.45 0.58 Ca 0.95 0.94 0.94 Na 0.05 0.05 0.06 Total 4.00 4.00 4.00 071060, 071056 & 071061: Hanging-wall banded skarn beds, from Jansson and Allen, 2011b.

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237

Table 5 Representative compositions of garnet from Ryllshyttan. Sample no.

071060

071056

071061

071061

071030

071030

071034

20071227

160.13

160.13

Setting

Main mass

Main mass

Main mass, BSE bright

Main mass, BSE dark

P.blast rim

P.blast core

P.blast core

P.blast core

P.blast rim

P.blast core

Facies

Inter-laminated Adr–Cpx–Ep N Mag rock

Inter-laminated Mag–Adr–Cpx rock

Massive Adr rock

Massive Adr rock

Qz–Sps rock

Qz–Sps rock

Prg rock

Bt schist

Bt schist

Calcic skarn

Point

p6

p1

p6

p7

TR2-01

TR2-L18

p1

p1

p11

p26

SiO2 Al2O3 FeO-tot MgO CaO TiO2 MnO Cr2O3 Total

35.52 4.98 22.76 nd 30.20 0.31 3.39 nd 97.10

35.54 3.07 25.83 nd 30.78 nd 1.80 nd 97.03

35.57 3.67 24.02 0.09 28.84 0.26 4.77 nd 97.22

35.57 7.41 20.92 nd 26.91 0.18 6.28 nd 97.26

36.38 15.61 11.20 0.21 8.27 0.07 27.68 nd 99.54

34.22 3.18 23.31 0.07 15.95 0.31 20.04 nd 97.10

37.16 20.95 24.49 2.95 4.79 0.09 9.92 nd 100.36

37.37 20.97 32.00 1.99 3.35 0.01 4.70 nd 100.39

36.41 20.72 22.96 2.74 3.90 0.02 12.24 nd 98.97

35.64 7.20 21.16 0.03 32.00 0.68 1.10 0.00 97.85

5.93 0.07 0.00 0.00 6.00 3.86 0.01 0.13 4.00 3.14 1.34 0.70 0.82 6.00 16.00

6.02 0.00 0.00 0.00 6.02 3.98 0.00 0.00 3.98 4.30 0.64 0.48 0.58 6.00 16.00

5.90 0.10 0.00 0.00 6.00 3.86 0.00 0.13 4.00 2.98 1.68 0.66 0.68 6.00 16.00

5.84 0.16 0.00 0.00 6.00 1.23 0.08 2.68 4.00 0.22 0.15 0.01 5.62 6.00 16.00

Number of ions on the basis of 24O and Fe3+/Fe2+ estimated on the basis of 16 cations and garnet end-members normative calculation Si 5.93 5.97 5.97 5.92 6.01 5.96 AlIV 0.07 0.03 0.03 0.08 0.00 0.00 Ti 0.00 0.00 0.00 0.00 0.00 0.04 Fe3+(IV) 0.00 0.00 0.00 0.00 0.00 0.00 [R]IV 6.00 6.00 6.00 6.00 6.01 6.00 AlVI 0.91 0.58 0.70 1.38 3.04 0.65 Ti 0.04 0.00 0.03 0.02 0.01 0.00 3+ Fe (VI) 3.06 3.42 3.27 2.60 0.94 3.35 [R3+]VI 4.00 4.00 4.00 4.00 3.99 4.00 Fe2+ 0.12 0.21 0.11 0.31 0.61 0.05 Mn 0.48 0.26 0.68 0.89 3.87 2.96 Mg 0.00 0.00 0.02 0.00 0.05 0.02 Ca 5.40 5.54 5.19 4.80 1.46 2.98 2+ [R ]VIII 6.00 6.00 6.00 6.00 6.00 6.00 Total 16.00 16.00 16.00 16.00 16.00 16.00 071060, 071056, 071061 & 071030: 071060, 071056 & 071061: Hanging-wall banded skarn beds, from Jansson and Allen, 2011b. 071034: Pargasite ≫ garnet altered rhyolite from the Ryllshyttan Lake Outcrop (Fig. 3, Jansson and Allen, 2011b).

Table 6 Representative compositions of calcic clinoamphibole from Ryllshyttan. Sample no.

071056

071061

071030

071034

[email protected]

TSRY02

06-7414

20071218

Setting

Vein-hosted

Vein-hosted

Main mass

Main mass

P.blast

Vein-hosted

Main mass

Main mass

Facies

Inter-laminated Mag–Adr–Cpx rock

Massive andradite rock

Qz–Sps rock

Prg rock

Bt schist

Magnesian skarn

Calcic skarn

Na–Ca altered footwall rhyolite

Point

p8

p19

p6

p6

p1

p5

p3

p2

SiO2 Al2O3 FeO-tot MgO CaO TiO2 MnO Cr2O3 Na2O K2O Total

48.18 4.43 21.42 8.43 11.96 nd 1.82 nd 0.79 0.35 97.38

54.04 1.99 11.41 15.02 12.48 nd 1.65 nd 0.34 0.11 97.05

52.94 1.65 12.64 14.30 11.62 nd 3.68 nd 0.61 0.28 97.71

42.10 15.23 14.75 10.37 11.51 nd 0.80 nd 2.00 0.25 97.00

43.07 14.61 14.66 10.54 11.47 0.27 0.83 nd 1.81 0.32 97.58

57.55 0.51 5.44 21.10 13.29 nd 0.33 nd 0.15 0.04 98.40

51.80 2.49 14.45 14.53 12.29 0.08 1.17 nd 0.41 0.08 97.30

54.16 1.60 10.60 17.45 12.63 nd 0.42 nd 0.31 0.17 97.33

7.83 0.00 0.34 1.38 0.20 3.25 1.94 0.09 0.02 15.06

7.75 0.00 0.28 1.55 0.46 3.12 1.82 0.17 0.05 15.22

6.29 0.00 2.68 1.84 0.10 2.31 1.84 0.58 0.05 15.69

6.38 0.03 2.55 1.82 0.10 2.33 1.82 0.52 0.06 15.61

7.94 0.00 0.08 0.63 0.04 4.34 1.96 0.04 0.01 15.04

7.61 0.01 0.43 1.78 0.15 3.18 1.94 0.12 0.01 15.23

7.77 0.00 0.27 0.00 1.27 0.05 3.73 1.94 0.08 15.15

Cations on the basis of 23 oxygens SiO2 7.39 TiO2 0.00 Al2O3 0.80 FeO 2.75 MnO 0.24 MgO 1.93 CaO 1.97 Na2O 0.23 K2O 0.07 Total 15.36

071056, 071061 & 071030: Hanging-wall banded skarn beds, from Jansson and Allen, 2011b.

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Table 7 Representative analyses of skarn from Ryllshyttan. Sample no.

71016

20071252

20071247

20071253

20071241

20071220

20071235

Description

Dolomitic marble w. 3–5% 1 mm Mag. Impr.

Tr-rich dolomitic marble w. minor Mag.

Tr-rich dolomitic marble w. 35–40% Mag.

Tr + Mag-veined fine-grained Di-skarn

Massive Tr skarn w. 2–3% mag impr.

Massive Ferroan Di-skarn

Massive Adr–Ferroan Di skarn

1.03 b0.01 0.16 10.19 2.31 16.04 34.88 b0.01 b0.01 b0.001 8.58 0.09 35.6 100.25 2.6 5.5 0.2 4.3 b0.1 b1 0.3 32 602 309.1 781 11 1.4 3.8

4.52 0.01 0.22 9.84 0.84 19.78 25.73 b0.01 0.02 0.017 10.73 0.3 37.7 98.7 4.9 18.9 0.3 6.6 1.1 6 1.5 35 203.8 280.9 3108 5.8 1.6 2 0.049 22.27 22 16.33 b1 14.7 11.1 0.1 2.4 0.1 0.01 0.3 0.2 b1 1.3 5 b0.1 b0.1 0.4 0.4 9 13.8 3.7 8.6 1.21 5.7 1.63 0.73 2.58 0.48 2.48 0.44 1 0.13 0.62 0.09

12.98 0.01 0.61 57.99 0.22 5.82 15.51 0.03 0.02 0.059 2.83 1.93 6 99.29 2.5 5.4 1.1 76.4 4.2 2 1 29.9 641.8 45.4 1060 2.3 0.2 3.5 0.025 4.09 61 2.27 2 0.9 4.8 0.1 7 0.1 b0.01 2.5 0.2 1 2 3 b0.1 0.3 0.3 0.3 42 b0.5 2.2 2.1 0.43 1.7 0.62 0.15 0.99 0.17 0.95 0.18 0.37 0.06 0.39 0.09

52.55 b0.01 0.29 4.55 0.47 17.11 22.8 0.06 b0.02 0.009 0.3 0.02 1.9 99.7 0.5 29.7 0.4 6.5 0.2 2 0.6 7.1 14.9 12.4 45 b0.5 0.2 b0.5

57.06 b0.01 0.27 2.24 0.24 23.3 13.3 0.13 0.06 0.012 0.05 0.64 1.7 98.29 1.9 86.4 1.5 23.5 b0.1 4 1.8 4.2 19.8 2.7 N10,000 1.5 b0.1 b0.5

49.5 0.01 0.99 15.2 0.9 9.59 22.58 0.33 0.02 0.031 0.22 b0.02 0.6 99.75 1.1 16.1 0.1 14.8 0.2 4 1.2 17.5 10.9 12.6 105 b0.5 b0.1 b0.5 0.011 1.11 99 11 6 0.3 0.1 0.2 3.5 b0.1 b0.01 b0.1 b0.1 1 b0.5 3 b0.1 b0.1 b0.2 1.4 b8 b0.5 8.7 23 3.62 16.5 5.03 3.01 5.17 0.77 3.6 0.44 0.85 0.07 0.47 0.06

43.7 0.02 1.77 23.59 1.07 4.86 25.01 0.14 b0.02 0.053 0.02 0.56 0 100.21 1.5 38.2 0.8 13.1 0.7 b1 1.1 5 997.1 1.9 126 1.1 0.1 b0.5 0.015 0.84 177 1.88 4 0.3 0.7 0.2 7.6 b0.1 0.02 3.3 b0.1 b1 b0.5 20 b0.1 b0.1 b0.2 0.3 13 286.9 2.4 6.9 1.58 10.9 5.22 8.32 6.66 1.15 5.23 0.93 2.07 0.24 1.28 0.13

SiO2 TiO2 Al2O3 Fe2O3-tot MnO (%) MgO (%) CaO (%) Na2O K2O P2O5 Tot C (%) S (%) LOI (%) Raw tot. (%) Zr Y Nb Co Ni Ba Rb Sr Cu Pb Zn As Ag Au (Zr/TiO2)/10,000 (Zr/Al2O3)/10,000 Al2O3/TiO2 Zr/Nb Be Bi Cd Cs Ga Hf Hg Mo Sb Sc Se Sn Ta Tl Th U V W La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppb

16.25

ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm

13 b1 2.2 2.5 b0.1 2.4 b0.1 b0.01 0.7 0.5 b1 b0.5 b1 b0.1 b0.1 b0.2 3.5 b8 1.9 13.5 16.5 1.76 5.3 0.94 1.82 1.15 0.18 0.84 0.12 0.25 0.04 0.14 0.02

in the clinopyroxenes in distal positions (Meinert, 1987) whereas clinopyroxenes at Ryllshyttan are more Mg-rich. In this sense, Ryllshyttan is more similar to magnesian Fe skarn deposits. Our observations indicate that the pervasive replacement of limestone by anhydrous skarns in a large part of the Ryllshyttan deposit

1.72

7.03

1.25 3 0.5 b0.1 b0.1 1.9 b0.1 b0.01 1.1 b0.1 b1 b0.5 5 b0.1 b0.1 b0.2 0.2 8 0.6 6.9 15.3 2.19 8.8 2.88 1.96 4.31 0.82 4.45 0.72 1.64 0.17 0.89 0.11

1.27 b1 b0.1 34.1 0.2 b0.5 b0.1 0.03 14.3 b0.1 b1 1.2 b1 b0.1 b0.1 b0.2 1.1 b8 0.7 1.2 1.6 0.46 3 2.2 0.31 6.44 1.66 10.83 2.1 5.37 0.67 2.98 0.34

cannot be accounted for by regional metamorphism of precursor alteration minerals. We rather suggest that the paragenetically early, anhydrous skarns formed during the emplacement of igneous intrusions, prior to the onset of deformation, regional metamorphism and remobilisation, as discussed below.

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239

Fig. 18. Microscopic ore textures observed in the magnetite and sulphide ores of Ryllshyttan. A. Sphalerite ore with interstitial galena (Gn), chlorite (Chl) and minor pyrite. Note foliated texture and abundant pyrrhotite inclusions in sphalerite grains. B. and C. Massive sphalerite ore with inclusions of magnetite, abundant needles of serpentine and chloritised phlogopite (Phl) and minor fluorite (Fl). Note fine rims of magnetite–pyrite intergrowths around serpentine and chloritised phlogopite and the absence of these in magnetite. D. Sphalerite inclusions rimmed by chloritised phlogopite and minor andradite in post-skarn chalcopyrite vein. E. Complex intergrowths of pyrite–magnetite replacing pyrrhotite occurring as inclusions in massive sphalerite ore. F. Sphalerite in a pressure shadow defined by boudinaged magnetite in a matrix of strongly deformed tremolite skarn.

9.3. Relation to intrusions The microgranodiorite cuts the K–Mg–Fe ± Si and K–Si altered zones, but is itself affected by oligoclase–diopside ± actinolite (Na–Ca) alteration, which also affects adjacent volcanic rocks. The oligoclase– diopside ± actinolite assemblage is reminiscent of the plagioclase–

diopside endoskarn in the deeper parts of the Yerington batholith, Nevada, USA, where primary igneous minerals in a quartz monzodiorite have been replaced by plagioclase, diopside, actinolitic hornblende and titanite in the thermal aureole of a porphyry–skarn system (Dilles et al., 2000; Einaudi, 2000). Where epidote occurs instead of diopside at Ryllshyttan, the alteration lithologies are similar to the Na–Ca

Fig. 19. Metamorphosed hydrothermal alteration facies at Ryllshyttan. A: Quartz + phlogopite/biotite + almandine rock formed through intense K–Mg–Fe ± Si alteration followed by metamorphism of rhyolitic siltstones of the Ryllshyttan hanging wall. The rocks record both S1 and S2 foliations. B: Weakly quartz + phlogopite/biotite-altered rhyolite of the Ryllshyttan hanging wall overprinted by pre-D2 (folded) veinlets of albite–oligoclase.

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Fig. 20. Schematic summary diagram illustrating the interpreted pre-folding geometry of alteration zones, mineral- and lithogeochemical variations in the alteration–mineralisation system and the key cross-cutting relationships between alteration and intrusions. Legend as in Fig. 3. Abbreviations; Adr: andradite, Alm: almandine, Bt: biotite, Di: diopside, Grs: grossular, Hd: hedenbergite, Jhn: johansenite, Ms: muscovite, Phl: phlogopite, Prp: pyrope, Ser: sericite, Sps: spessartine.

alteration zones of Yerington, in which primary felsic igneous assemblages have been replaced by mineral associations of albite–oligoclase (andesine), actinolite, quartz, epidote and locally chlorite or pyrite (Dilles et al., 2000). Sodic–calcic alteration also occurs in the deeper parts of semi-conformable altered zones associated with VMS deposits where seawater was heated to temperatures above 170 °C (Galley, 1993). Sodic alteration of feldspars results from the interaction of rocks with high temperature and high Na/K fluids (Giggenbach, 1984). Naaltered rocks are most commonly preserved in the recharge zone of a hydrothermal system, where descending Na-bearing solutions are being heated. The Na-alteration may be accompanied by Mg- and Careplacement of K-silicate minerals if the hydrothermal fluids are enriched in Mg and Ca. In contrast, zones of ascending hydrothermal fluids are commonly dominated by K–Al silicates, reflecting the cooling of a hydrothermal fluid during its passage through the upflow zone (Giggenbach, 1984; Hitzman et al., 1992). At Ryllshyttan, the Na–Ca alteration appears to overprint and therefore post-date the K–Mg–Fe ± Si alteration, suggesting that the two alteration styles resulted from distinct events. This overprinting relationship can however be explained by a prograde evolution of a single intrusion-driven hydrothermal system. At an early stage, a shallow,

distal hydrothermal alteration, far above deeply emplaced plutons, would have been characterised by focused K–Mg–Fe ± Si alteration in the upflow zones. Evidence for early magmatism coeval with early alteration in the host succession includes the presence of mafic dykes, which are spatially and temporally linked with the K–Fe–Mg ± Si alteration and sulphide zones (Fig. 20), suggest that the mafic dykes, their intrusive contacts and the faults and fractures were important hydrothermal and magma conduits. The K–Mg–Fe ± Si alteration style is in agreement with modified seawater-dominant hydrothermal fluids. Following further basin subsidence, burial of the Ryllshyttan succession and granodioritic–dacitic intrusion(s) emplacement high in the stratigraphy in the deposit vicinity, the earlier, shallow alteration would have been overprinted by alteration formed under higher temperatures, such as the observed Na–Ca alteration. During final emplacement and cooling of the microgranodiorite, Na–Ca alteration would have propagated into the microgranodiorite. The fluids involved in that Na–Ca alteration may have modified the Ryllshyttan deposit, including possible redistribution of ore metals. The Na–Ca alteration was associated with Fe depletion in the microgranodiorite (Jansson, 2011), and consequently may have been one possible source for Fe deposited elsewhere in the system. Based on co-magmatism with the Smältarmossen dacite porphyry, and the

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100 Zn+Pb+Cu enriched strongly K-Mg-Fe+/-Si altered rhyolites and mafic rocks

Ishikawa alteration index [(MgO + K2O) / (MgO + K2O + Na2O + CaO)]

90

s trong-intens e s eric ite and/or c hlorite alteration

80 m oderate s eric ite and/or c hlorite alteration

70 60 50

U naltered to weak ly altered s am ples

40 30 20 Na-Ca alteration

10 0 0

200

400

600

800

1000

Zn+Pb+Cu (ppm) Fig. 21. Ishikawa alteration index (Ishikawa et al., 1976) vs. total Cu + Zn + Pb (ppm) for variably altered igneous rocks in Ryllshyttan. Legend as in Fig. 7.

intimate relationship between endoskarn-like Na–Ca alteration and the magnetite skarns both at Ryllshyttan and Smältarmossen, it is conceivable that the microgranodiorite established a contact metasomatic regime in its vicinity, beneficial for the formation of magnetite skarn deposits.

241

(re)mobilised during D2 and/or D3 folding and shearing. This coincided with retrograde alteration of anhydrous skarns to more magnetite-rich, hydrous skarn, considerably contributing to the total magnetite budget of the deposit. The anhydrous skarns could potentially have formed by regional metamorphism of precursor alteration assemblages in the limestone, although this appears unlikely given the pervasiveness of limestone replacement by skarn and sharp contact between thick zones of massive skarn and marble. Instead, a hybrid model is favoured to account for the apparent paradox that early alteration at Ryllshyttan is partly similar to metamorphosed VMS systems and partly similar to intrusion-related skarn systems. In such a scenario it is conceivable that early intrusions generated a convective seawater-dominated hydrothermal system and K–Mg–Fe ± Si alteration as in VMS systems. Subsequent intrusion(s) to higher crustal levels raised the temperature of the hydrothermal system and generated anhydrous skarn and further mineralisation. At each stage the limestone acted as a reactive trap to hydrothermal solutions. The last stage involved deformation, (re)mobilisation and retrograde alteration of the anhydrous skarns adjacent to D2 and D3 shear zones, forming high-grade massive sulphide as well as magnetite ore bodies, with sulphides consistently overprinting magnetite, reflecting the greater mobility of sulphide minerals compared with magnetite during ductile deformation under medium grade metamorphic conditions. This model partly resolves the controversy regarding genesis of iron oxide and base metal sulphide deposits in Bergslagen, as we recognise that these deposits have a long history of alteration, metamorphism, deformation, and (re)mobilisation, and that a single VMS or intrusionrelated skarn model does not encompass all the features pertinent to formation and modification of these deposits. Acknowledgements

10. Conclusions The stratigraphic succession at Ryllshyttan records the evolution of a marine, rhyolitic volcanic terrain where limestone and exhalative, calcareous iron formations accumulated during pauses in volcanic activity. On a larger scale, the Ryllshyttan succession occurs in the roots of a large, marine, felsic caldera volcano, which contains a major calderafill pyroclastic deposit, the base of which is 800 m stratigraphically above the Ryllshyttan ore-bearing strata. Iron oxide and base metal sulphide mineralisation formed by replacement of limestone following burial of the succession and intrusion by rhyolite porphyries. Hydrothermal fluids were introduced along the margins of mafic dykes that cross-cut the succession. Both the dykes and the rhyolitic rocks were altered to chlorite-, sericiteand quartz-rich mineral associations (K–Fe–Mg ± Si alteration) and the limestone was extensively dolomitised and mineralised with sulphides and some magnetite. Evidence from cross-cutting relationships, lithogeochemistry and microfabrics suggests that the earliest iron oxides and sulphides were introduced prior to the early stages of ductile deformation. After burial of the succession, a microgranodiorite pluton was emplaced directly below and adjacent to Ryllshyttan. The microgranodiorite and adjacent rocks were affected by Na–Ca alteration that overprints the K–Fe–Mg ± Si alteration. From the edge of the Ryllshyttan mineralisation towards the centre, the limestone ore host shows decreasing galena vs. sphalerite and galena vs. chalcopyrite ratios and increasing skarn replacement of the limestone. Within the limestone there is a vertical zonation from stratigraphically lower diopside-rich skarns to stratigraphically higher ferroan diopside-rich skarns that locally carry andradite. The distribution of the most massive, high-grade sulphide and magnetite mineralisation suggests a structural control and association with hydrous, retrograde phases (e.g., serpentine, tremolite–actinolite) that appear to have replaced the anhydrous, prograde skarn minerals (e.g., clinopyroxene). This indicates that sulphide mineralisation was

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