Scheelite skarn mineralization associated with the ...

Ν. Jb. Miner. Abh.

2009, Vol. 186/1, p. 37-50, Stuttgart, July 2009, published online 2009 © by E. Schweizerbart’sche Verlagsbuchhandlung 2009

Scheelite skarn mineralization associated with the Tinos pluton, Cyclades Karen St. Seymour, Stylianos Tombros, Nickolaus Mastrakas, Dimitrios Zouzias, Paul G. Spry, George Denes and Prokopis Kranidiotis

With 6 figures and 5 tables

Abstract. Scheelite (CaW 04) mineralization has developed as disseminations, veinlets and mainly pods within garnet-pyroxene skarn and homfelses in two locations within the contact aureole of a syntectonic calc-alkaline granodiorite to leucogranite pluton on the island of Tinos, Hellas. The pluton was emplaced during a transition from compression (granodiorite) to extension (leu­ cogranite). Opening of conduits assisted by carbofracturing allowed fluid circulation during extension, which resulted in a change of character of contact metamorphism from thermal-isochemical to infiltration metasomatism. During infiltration metasomatism ion exchange involved the plutonic and the country rocks and was intensified by carbofracturing resulting from the pyrometasomatic loss of C 0 2 from the marbles. Skarn formation may have been initiated due to thermal effects of the pluton with local mass transfer in the host schists and marbles. The Tinos tungsten skarn is an oxidized skarn that retains relics of an earlier less oxidized stage probably mirroring a change in the redox conditions of the leucogranite. Skarn pyroxene is diopside, however, hedenbergite has been preserved locally. Cores of garnet megacrysts from an early skarn formation are rich in grossular component and exhibit sector-zoning, characteristic of growth under contact metamorphic conditions at log f02 that ranged between -19.8 and -22.9. Gar­ net overgrowths from a later stage are predominantly andradite and display euhedral crystal boundaries, oscillatory zoning, zonal anisotropy and growth as open space fillings, which are features typical of hydrothermal garnet. Scheelite crystals are enclosed within and display equilibrium boundaries to hydrothermal garnet. They have primary fluid inclusions with an average homogeni­ zation temperature Th of 375 °C and salinity of -14 wt% NaCl equivalent. Th of primary fluid inclusions in the coexisting hydrothermal garnet ranges from 310 to 340 °C, of primary inclusions in quartz from quartz-calcite veinlets formed by carbofracturing from - 400 to 375 °C and salinity -14 wt% NaCl equivalent. These quartz-calcite veinlets cut mainly the early garnet cores, rich in the grossular component, but rarely the hydrothermal andraditic garnet overgrowths, indicating that scheelite mineralization is associated with the late leucogranitic phase of the Tinos pluton. Key words: leucogranite, carbofracturing, infiltration metasomatism, diopside-hedenbergite, grossular-andradite, oscillatory zon­ ing.

Introduction The majority of tungsten deposits are spatially related to granitoid plutons. Whether the magma provides the tung­ sten directly via a juvenile fluid or the tungsten is leached from the host rocks at some later time is still debated. However, in either case the tungsten is ultimately sup­ plied from a fluid that has equilibrated with the pluton. Deposition of tungsten may occur in the granitoid and/or silicate and/or carbonate country rocks. In the latter case the tungsten mineralization is typically hosted by skarn. N ewberry (1983) has related variations of skarn mineral chemistry to intrusion depth and suggested they were a DOI: 10.1127/0077-7757/2009/0135

function of intrusion-imposed oxidation state. Compre­ hensive reviews of tungsten skarns have been provided by N ewberry & E inaudi (1981), N ewberry (1983), N ewberry & Swanson (1986), K wak (1987), M einert (1992), R ay (1995) and N ewberry (1998). Scheelite skarn mineralization at Tinos occurs within the contact aureole of a Miocene early granodioritic to late leucogranitic pluton. An exhalative origin has been proposed for this mineralization on the basis that it ap­ pears to be stratabound (Papastavrou & P antoulas 1992). We address here the question of the origin of the tungsten-bearing fluids and of the role of skarn mineral variability and textures in understanding the evolution of 0077-7757/09/0135 $3.50

©2009 E. Schweizerbart’sche Verlagsbuchhandlung, D-70176 Stuttgart

38

K. St. Seymour et al.

a skarn system and the physicochemical conditions and mechanisms that led to scheelite precipitation.

Geological setting Tinos Island belongs to the Attico-Cycladic massif (Fig. 1), which consists of crystalline complexes that connect the Hellenic mainland with the Menderes massif in Asia Minor. The massif consists of an Upper and Lower struc­ tural group of units separated by low-angle faults (DU r r et al. 1978, Schliestedt et al. 1987, O krusch & B rocker 1990). The Upper Group consists of the 'Ophiolite Cover’ and unmetamorphosed sedimentary rocks. The Lower Group is the Cycladic Blueschist Unit which consists of

a Pre-Alpidic Basement of granite, tonalite and gabbro in the islands of Ios, Syros, Paros and Naxos (F ranz et al. 1993, G autier et al. 1993). The basement is overlain by thrust sheets of a metamorphosed volcanosedimentary sequence. The Cycladic Blueschist Unit has experienced two stages of Tertiary metamorphism. The hrst metamorphic episode occurred at eclogite to blueschist facies con­ ditions (T = 450-500 °C, Ρ = 15 ± 3 kbar; O krusch et al. 1978, M atthews & Schliestedt 1984, B rocker et al. 1993). Based on U-Pb zircon chronology it has been suggested that the eclogite - to blueschist facies rocks formed during multiple episodes separated by significant time intervals with U-Pb ages at 78 ± 1 Ma, 61 and 63 Ma (B rocker & E nders 1999) whilst earlier studies based on white mica chronology yielded ages between 50 and

SK

K L ivadas B a y

km

Jr k-

Fuboea

A ttic a

N ^ T in o s

Syros

■37N

24E

Ikarja

Mykonos Naxos Paros

I IPelagoniki 1---unit [ ||g | Sediments Vblcanics I

{Granitoids

^Blueschist ---unit

BIUe|c t K SUni‘

□

Granitoids

B ,U m a ! & nit

E M D U p p e r unit

Basal unit

Fig. 1. Simplified geological map of Tinos Island and location in the Cycladic Belt. S k : Location of studied skarns and scheelite miner­ alization. Panormos Bay (Q): Location of Au-Ag-Te mineralization (TOMBROS et al. 2007). 1. Quartzites and schists of the lower part of the Blueschist Unit, 2. Marbles, and 3. Blueschists, greenschists of the Blueschist Unit. 4. Ophiolites of the Upper Unit, 5. Tinos granitoid pluton (modified after Melidonis 1980).

Scheelite skarn mineralization

40 Ma (A ltherr et al. 1979, W ijbrans & M cD ougall 1986). A second metamorphic event of greenschist to am­ phibolite facies reaching migmatization, occurred at peak pressures of 4-7 kbar (Avigad et al. 1992, B rocker et al. 1993). White mica 40Ar - 39Ar and Rb - Sr dating of greenschist and amphibolite facies rocks yielded ages be­ tween 25 and 18 Ma (A ltherr et al. 1982. W ijbrans & M c D ougall 1988, W ubrans et al. 1990). This second metamorphic episode is considered to result from isother­ mal decompression during uplift for the northern Cyclades or from prograde metamorphism for the southern Cycla­ des. Detailed overviews of the geological evolution of the Attico-Cycladic Massif are given by DOrr et al. (1978), DOrr (1986), Schliestedt et al. (1987) and O krusch & B rocker (1990). Tinos island (Fig. 1) consists of a series of stacked nappes of the Lower and Upper Groups cut by a Miocene granodiorite pluton (M elidonis 1980). In Tinos the Low­ er Group i.e. the Cycladic Blueschist Unit is distinguished into two subunits: the Lower subunit consists of a dolomite-phyllite-quartzite sequence, the Upper subunit con­ sists of retrograde to greenschist facies, blueschists and intercalated marble horizons (i.e., a package of metamor­ phosed tholeiitic mafic to felsic volcanic and sedimentary rocks). The superior nappe is the Upper Unit (i.e., an­ cient oceanic crust metamorphosed to greenschist facies), which was overthrusted at ~18 Ma (B oronkay 1995) or alternatively at ~21 Ma (B rocker & F ranz 1998, Z ef fren et al. 2005).

Tinos plutonic rocks and contact aureole The Miocene Tinos pluton is exposed over an area of about 30 km (Fig. 1). The pluton consists of a central body of Itype biotite-hornblende granodiorite emplaced ca. 17 Ma (K-Ar ages from amphibole and biotite, A ltherr et al. 1982, H enjes -K unst et al. 1988). Genesis of the granodi­ orite magma is attributed to partial melting of mafic lower crustal rocks (M astrakas 2007). An S-type biotite-muscovite-garnet-tourmaline leucogranite, of much smaller present day exposure, was emplaced peripherally to the main plutonic body and also as sills at shallow levels. The age of emplacement of the leucogranite has been deter­ mined as ca. 14 Ma (K-Ar ages from A ltherr et al. 1982, H enjes -K unst et al. 1988, B rocker & F ranz 1998). Pluton contacts display thermal contact effects with the country rocks and mylonitic borders. The faulted contacts of the pluton were tectonically active until at least 8-10 Ma (B oronkay 1995). The joint system, which cuts the granodiorite, hosts not only aplites but also frequently pseudotachylites (B oronkay 1995, M astrakas & St .

39

Seymour 2000). Geothermometric studies indicate that the granodiorite was emplaced at a temperature range of 750° to 800 °C, whereas the leucogranite was emplaced ca. 680 °C (M astrakas 2007). The granodiorite crys­ tallized at a pressure of ~4.7 kbar and the leucogranite at -1.5 to 2 kbar and the latter experienced intense ret­ rograde boiling as shown by numerous miarolitic cavi­ ties and alkali-feldspar porphyritic texture (M astrakas 2007). Based on geochemical evidence M astrakas & S t . S eymour (2000) argued that the Tinos intrusive phas­ es were emplaced in a Miocene tectonic regime evolv­ ing from compression (I-type granodiorite) to extension (S-type leucogranite). Differentiated magma pods within the cogenetic Tinos granodiorite were emplaced during a transition from compression to extension. These findings agree with those of B oronkay (1995), which were based on structural data and those by B rocker & F ranz (1998) who also suggested extensional emplacement of the leu­ cogranite. The Tinos pluton displays a weak syntectonic folia­ tion mostly developed parallel to its contacts and a ther­ mal imprint against the Blueschist and Upper Units. The contact aureole has been described in detail by B rocker & Franz (1994, 2000), Stolz et al. (1997) and M astra ­ kas (2007). The metamorphic zones are generally welldefined with the exception of the eastern border of the intrusion. The width of the inner three zones (scapolite to hornblende homfels) varies from 600 m to a maximum of 1100 m. K-Ar dating of contact metamorphism has shown that biotite in the granodiorite and from the contact aure­ ole “closed” at ca. 14 Ma (B rocker et al. 1993). In the contact area west of the pluton, the K-Ar system closed at

Spess+Pyrope+Alm

Spess+Pyrope+Alm

Fig. 2. Compositional variation of garnet from northern and south­ ern areas of Tinos skarn scheelite mineralization (in mol%). Closed symbols: cores, open symbols: rims.

40

K. St. Seymour et al.

ca. 8-10 Ma (B rocker & F ranz 2000). The older ages are attributed to cooling of granodiorite and crystallization of leucogranite, whereas the younger ages are interpreted to represent deformation or a different cooling history for this part of the aureole (B rocker & F ranz 2000).

Tinos skarn and skarn mineralization Skarn rocks developed within amphibole schists and mar­ bles of the Blueschist Unit. Pyroxene-garnet skarn rocks occur in the western part of the contact aureole of Tinos

Table 1. Representative microprobe analyses of garnet from Tinos skarn mineralization. Sn

W

Mo

Cu

Zn

1.5

4



> J > J >

& Ρ

^

Ο (Ν

Ο X

J

«η (Ν

-

Γ - CN I— ■— 1

Ο CN sO t J·

O O O O O ^ O O

J‘“"S.

>

o

o

^

CL ^

— oVO

o

Γ'- —

>> S’

f-a

c o "C o CL ω

CL

CL

-σ cI CO

CO

H

o-

CN CN OO CN cn

P

3I § a Ν

Ν

Ν

Ν

ο σ σ σ i>->^>1 4^ >ΝL4 0) (U