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Ore Geology Reviews 24 (2004) 315–370

Skarn textures and a case study:

the Ocna de Fier-Dognecea orefield, Banat, Romania

Cristiana Liana Ciobanu*, Nigel John Cook

Geological Survey of Norway, NO-7491 Trondheim, Norway

Received 3 February 2002; accepted 4 April 2003

Abstract

We address the question of the predictability of skarn textures and their role in understanding the evolution of a skarn

system. Recent models of skarn formation show that skarns are ideal for application of self-organisation theory, with self-

patterning the rule in fluid-rock interaction systems rather than the exception. Zonation in skarn deposits, a consequence of

infiltration-driven metasomatism, can also be treated in terms of self-organisation. Other less commonly described features,

such as scalloping, fingering and mineral banding, can be understood by application of reactive infiltration and hydrodynamics

at the skarn front. Devolatilisation may trigger formation of back-flow fluxes that overprint previously formed skarn. The range

of textures formed from such events can be used to discriminate between prograde and retrograde stages. Refractory minerals,

such as garnet, magnetite and pyrite, readily retain overprinting events. Skarns are also composed largely of minerals from solid

solution series (garnet, pyroxene, pyroxenoids, etc.) and therefore skarn mineralogy helps to establish trends of zonation and

evolution. The same minerals can act as ‘chemical oscillators’ and record metasomatic trends.

The Ocna de Fier-Dognecea deposit was formed in a f 10 km deep skarn system. Zonation and evolution trends therefore

represent only the result of interaction between magmatically derived fluids emerging at the source and limestone. From the

same reason, the transition from prograde to retrograde regime is not influenced by interaction with external fluids. Thirdly, the

mineralisation comprises Fe, Cu and Zn-Pb ores, thus facilitating comparison with skarn deposits that commonly are formed in

shallower magmatic-hydrothermal environment. Copper-iron ores (magnetite +Cu-Fe sulphides), hosted by magnesian

(forsterite + diopside) skarn, occur in the deepest and central part of the orefield, at Simon Iuda. Their petrological character

allows interpretation as the core of the skarn system formed from a unique source of fluids emerging from the subjacent

granodiorite. It formed first as a consequence of the local setting, where a limestone indented in the granodiorite permitted

strong reaction at f 650 jC and focussed the up-streaming, buoyant fluids. The first sharp front of reaction is seen at the

boundary between the Cu-Fe core and Fe ores hosted by calcic skarn (Di70-90-And70-90), where Cu-Fe sulphides disappear, and

forsterite gives way to garnet in the presence of diopside (Di90). Following formation of forsterite, devolatilisation and

transient plume collapse is interpreted from a range of piercing clusters and trails. We presume lateral flow to have been

initiated at the source, as the emerging fluids are in excess to the fluids driven into reaction by the plume. Formation of the

other orebodies, up to 5 km laterally downstream in both directions, is interpreted as skarn fingering at the limestone side. The

metasomatic front is perpendicular to the flow along the channel of schists placed between the limestone base and the

granodiorite.

A metal zonation centred onto the source is defined, based on metal distribution: Cu-Fe/Fe/Zn-Pb. The second front of

reaction, at the boundary between the Fe and Zn-Pb zone, has a sulphidation/oxidation character, with diopside giving way to a

0169-1368/$ - see front matter D 2003 Elsevier B.V. All rights reserved.

doi:10.1016/j.oregeorev.2003.04.002

* Corresponding author.

E-mail address: Cristiana.Ciobanu@ngu.no (C.L. Ciobanu).

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370316

Fe-Mn-rich pyroxene, (HedJoh)>60 + pyroxmangiteF bustamite; garnet is minor. Johannsenite-rich pyroxene (Di20-40Hed20-

40Joh40) is found in proximal skarn at the upper part of Simon Iuda, stable with Zn0.95Fe0.05S, at an inferred 570 jC. In distal

skarn from Dognecea and Paulus, Mn-hedenbergite (Di< 10Hed70Joh20-30) formed at f 400 jC is stable with Zn0.84Fe0.16S.

Extensive compositional fields, eutectic decomposition and lamellar intergrowths characterise pyroxene in the Zn-Pb zone,

formed at the magnetite-hematite buffer in the presence of pyrite. Distal skarn has a reducing character, in comparison with the

proximal. A drop in both fS2 and O2, with the zoned system moving closer to the pyrite-pyrrhotite buffer, is induced from the

temperature gradient. Based on pyroxene mineralogy and calculated fS2, the metal zonation is confirmed as being formed

upwards and outwards from the source.

The Fe and Zn-Pb zones both have a patterned side coexisting with the unpatterned one. Patterning is seen at scales from

macroscopic (rhythmic banding, nodular, spotted, orbicular, mossy, mottled textures) to microscopic scales (oscillatory zonation

in garnet and silica-bearing magnetite). Following plume updraft, the path of decarbonation reaction controlled the motion of the

skarn front until, towards the end of the prograde stage, a multiple steady state regime developed and produced rhythmic patterns

on all scales. The activation of powerful patterning operators, represented by Liesegang banding alone, or coupled with

competitive particle growth, show that the skarn front had the characteristics of an unstable coarsening front of reaction.

A second retrograde event, carbofracturing, triggered by erratic decarbonation after cessation of infiltration, can be

interpreted from overprinting textures in the Fe and Zn-Pb zone. A major drop in fO2 is inferred from extensive,

pseudomorphous replacement of hematite by magnetite. Textures show progressive destruction of prograde assemblages, i.e.,

piercing clusters, shock-induced, fluid-pressure assisted brecciation and deformation, followed by healing of the disrupted

assemblages. Release of trace elements accompanies both retrograde events, with a Bi-Te-Au-Ag association common to

both. The importance of shock-induced textures is emphasised in the context of Au enrichment, especially when the

retrograde fluids cross the main buffers in fO2-f S2 space.

The presence of Bi-sulphosalt polysomes in the Fe zone indicates that patterning extends down to the nanoscale. The key

role played by polysomatism in stabilising compositional trends that cannot otherwise be formed at equilibrium is a fertile

ground yet to be adequately explored.

D 2003 Elsevier B.V. All rights reserved.

Keywords: Skarn; Ocna de Fier-Dognecea; Romania; Textures; Zonation; Self-patterning; Liesegang banding

1. Introduction

Skarn deposits, especially those in which refractory

minerals dominate, may have sets of retained textures

that record useful information on primary and over-

printing processes. Prevailing textures may not be

exclusive to skarns, but sequences of preserved tex-

tures may assist interpretation of mineralising process-

es and prograde-retrograde paths within individual

skarn systems. Similarly, patterns observed in hand

specimen and the crystal zonation of various skarn

minerals help to constrain the linear/non-linear/cyclic

evolution of mineralising processes driven by infiltra-

tion mechanisms (e.g., Guy, 1981, 1988; Jamtveit,

1991; Jamtveit et al., 1993, 1995). Deposit and ore-

field-scale zonation patterns are characteristic for

skarn districts and are the result of a range of processes

during skarn formation (Korzhinskii, 1970; Guy, 1984,

1988, 1993). The zonation patterns are themselves

varieties on the theme of textural patterning within

any given deposit, and represent powerful exploration

tools for skarn deposits (e.g., Meinert, 1997).

Skarn deposits typically show evidence of pro-

grade and superimposed retrograde stages (e.g.,

Einaudi et al., 1981). Although prograde skarnifica-

tion results from the action of magmatic fluids, the

retrograde stage commonly, but not necessarily,

includes contributions from hydrothermal or meteoric

waters in near-surface environments (e.g., Meinert et

al., 2003). During retrograde events, mixing of fluid

types, and possible boiling and collapse of the skarn

system strongly influence skarn brecciation (e.g.,

Meinert, 1992). Changes in rock permeability or

fluid/volatile production by devolatilisation (Dipple

and Gerdes, 1998) can lead to an overprinting of

primary textures, including hydrofracturing and brec-

ciation. In skarns, the term carbofracturing is appro-

priate when hydrofracturing and/or brecciation are

caused by a sudden separation of a CO2-rich vapour

(e.g., Bowman, 1998). Carbofracturing can be con-

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 317

sidered directly analogous to hydrofracturing and

brecciation caused by secondary (i.e., resurgent) boil-

ing in porphyry systems (e.g., Bowman, 1998). Rec-

ognition of such textures in skarns facilitates the

interpretation of back-flow paths within an evolving

skarn system (Dipple and Gerdes, 1998).

Numerous skarn deposits occur within the Upper

Cretaceous Banatitic Magmatic and Metallogenetic

Fig. 1. Geological sketch map of the Ocna de Fier-Dognecea orefield. Inset

at the Reichenstein III level (357 m) in the central, proximal part of the oref

groups also have equivalents on the western contact between limestone and

III level. Other orebodies and localities mentioned in the text are indicate

Belt of Southeastern Europe (Ciobanu et al., 2002a).

Typical of these is the Fe-Cu-(Zn-Pb)-skarn orefield at

Ocna de Fier-Dognecea (Fig. 1), in the Banat region of

Southwest Romania. The present contribution focuses

on this orefield, which, together with other deposits in

the region, played a key role in the early development

of skarn theory in the 19th Century (e.g., von Cotta,

1864; Castel, 1869; Marka, 1869; Sjogren 1886).

(bottom right) shows the principal orebodies, including Simon Iuda,

ield projected at the surface. Orebodies in the Reichenstein and Elias

schist, but these are restricted to elevations above the Reichenstein

d.

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370318

Sample material from across the entire orefield, re-

cording a variety of mineralising aspects, forms the

basis for our investigation. Following the central

source model introduced for the deposit by Cook and

Ciobanu (2001), we now address the skarn mineralogy

as well as a broad range of primary and overprinting

textures seen across the entire orefield in order to:

1. present the skarn mineralogy, zonation and evolu-

tion trends;

2. discuss the formation of rhythmic and other related

textures in the prograde stage;

3. document the switch from prograde to retrograde

stage using overprinting textures;

4. use all of the above to constrain a comprehensive

model for skarn formation;

5. assess the predictability of skarn textures.

2. Background to fluid-rock interaction in skarns

In any fluid-rock system, metasomatism represents

broadly the overall chemical changes accompanying

the reactions when a fluid moves through a porous

solid (e.g., Korzhinskii, 1970); it requires initial dis-

equilibria between the fluid and solid. Following

fundamental concepts regarding metasomatism and

metamorphic facies (Goldschmidt, 1911; Eskola,

1921), Korzhinskii (1959) introduced a thermody-

namic basis for mineral paragenesis as a method of

analysis in contact aureoles. This resulted in chro-

matographic modelling and phase equilibrium dia-

grams being used to study skarn assemblages (e.g.,

Newberry, 1991; Guy, 1995). Skarn formation, which

classically occurs at the contact of intrusions with

carbonate protoliths (e.g., Einaudi et al., 1981), is

successfully explained by metasomatic theory. How-

ever, skarn-forming systems lie within the broader

metamorphic domain (e.g., Meinert, 1992) and give

rise to more complex assemblages that reflect (inter

alia) the fluid source and the protolith, e.g., silty

limestone, banded iron formation (BIF), dolomite or

pure limestone.

2.1. Self-patterning in geochemical systems

Field evidence from a broad range of geological

settings (e.g., Joesten, 1991 and references therein)

indicates that fluid-rock interaction within a contact

aureole is more complex than can be predicted by

techniques that assume static equilibrium. In fact,

Joesten (1991) and Kerrick et al. (1991) have used

mineral kinetics in non-equilibrium environments to

interpret patterns frequently seen in contact aureoles.

Examples include mineral coarsening (Ostwald,

1925), metamorphic banding (Turing, 1952), calc-

silicate nodules, rims of calc-silicate on chert nodules

in limestone (Joesten, 1974, 1991) and widespread

oscillatory zonation in garnet (Jamtveit, 1991; Jamt-

veit et al., 1993, 1995).

Nevertheless, much of the basis for present-day

modelling of diffusion-controlled patterning in rocks

relates to a series of experiments known as ‘Liese-

gang phenomena’ (e.g., Krug and Kruhl, 2001).

Using diffusion sources in gels, Liesegang (1913)

obtained bands and rings of precipitates. His experi-

ments proved that rhythmic patterns could sponta-

neously develop in gels, without an inherited

background, i.e., via self-organisation. This was a

landmark in the understanding the intrinsic evolu-

tion of various types of systems (e.g., geological,

biological, chemical, etc.) that can induce self-orga-

nisation as a result of their equilibrium state. It

would seem that many systems under ‘far-from-

equilibrium’ conditions (Glansdorff and Prigogine,

1971) have spontaneously undergone an oscillatory

evolution that concludes with the development of

‘instabilities’ seen as ‘dissipative structures’ (Nicolis

and Prigogine, 1977). Guy (1981) discussed rhyth-

mic textures in skarns in terms of ‘dissipative

structures’. A similar approach was undertaken by

Jamtveit (1991) to interpret chaotic zonation pat-

terns in skarn garnets. In his comprehensive mono-

graph on self-organisation phenomena, Ortoleva

(1994) argues that during fluid-rock interaction there

are many ways in which geochemical systems are

driven out of equilibrium. The potential for pattern-

ing, and implicitly, for development of self-organi-

sation phenomena in geochemical systems is linked

to the existence of several isothermal reaction-mass

transport feedbacks (Ortoleva et al., 1987a). Exam-

ples include supersaturation-nucleation-depletion

cycles, competitive particle growth (CPG), autocat-

alytic crystal growth, and mechanical-chemical cou-

pling and reactive-infiltration instability (Ortoleva et

al., 1987b).

eology Reviews 24 (2004) 315–370 319

2.2. Oscillatory zonation: the record of a ‘chemical

oscillator’

Oscillatory zonation patterns in minerals (Shore

and Fowler, 1996 and references therein) may relate to

a whole range of aspects (variation in the major

components of minerals from solid solution series,

order-disorder phenomena in polysomatic or accre-

tional series, trace elements, adsorption of impurities,

or point defects). Their study therefore represents a

fertile field of investigation, in order to address the

question whether oscillatory zonation in crystals

records the role of a ‘chemical oscillator’ that can

arise spontaneously (e.g., Putnis et al., 1992; Prieto et

al., 1997) or require external control (Yardley et al.,

1991; Holten et al., 1997).

The success of experiments demonstrating the

connection between Liesegang phenomena and oscil-

latory zonation in minerals from solid-solution series

represents an important breakthrough in the applica-

tion of self-organisation patterning theory to crystal

growth (Ortoleva et al., 1994). Major-element oscil-

latory zonation has been obtained in (Ba, Sr)SO4 solid

solution in a Liesegang environment by counter

diffusion of (Ba2 +, Sr2 +) and SO42� ions in a porous

silica-gel transport medium (Putnis et al., 1992). The

experiments proved that an autocatalytic surface at-

tachment reaction (Ortoleva et al., 1987a; Ortoleva,

1990) could take place if threshold supersaturation for

nucleation and growth is strongly dependent on com-

position, as for example in series where end-members

have dissimilar solubility.

2.3. Fluid-rock interaction in skarns

Given their obvious characteristics, such as zonation

at all scales, skarn deposits offer an ideal type of

geochemical system for investigation of fluid-rock

interaction. The following is an attempt to illustrate

this affirmation. Taking both theory and observation

into account, we aim to address the predictability of

skarn textures and link to the causes of their formation.

The model of chromatography applied to infiltra-

tion metasomatic zoning (Korzhinskii, 1970; Guy,

1984, 1988, 1993) is an application of isothermal

reaction-mass transport theory. Oscillatory zonation

models of crystal growth can be tested using zonation

patterns in skarn minerals. Skarn-to-hydrothermal

C.L. Ciobanu, N.J. Cook / Ore G

evolution and/or oscillations raised during fluid-rock

interaction may control such zonation patterns (Jamt-

veit, 1991; Jamtveit et al., 1993, 1995). Widespread,

but less commonly mentioned features of skarns, e.g.,

scalloping, fingering, mineral/isotope banding, brec-

ciation, can be modelled in terms of reactive-infiltra-

tion coupled to hydrodynamics at the skarn front

(Dipple and Gerdes, 1998). This last model has great

importance for understanding skarn textures that are

formed at critical points in fluid evolution attained

during fluid-rock interaction.

2.3.1. Metasomatic zoning

Korzhinskii (1965, 1968, 1970) stressed that fluid

transport occurs by infiltration (controlled by pressure

gradients) and diffusion (dependant upon chemical

potentials), the latter being an order of magnitude

slower than infiltration. He also concluded that a

metasomatic column resulting from infiltration has

sharp reaction fronts, whereas diffusion zoning has

transitional limits. Some of Korzhinskii’s predicted

zonation models have been obtained in a series of

experiments involving granodiorite and limestone

percolated by solutions, the compositions and param-

eters of which have been externally controlled (e.g.,

Zaraisky, 1991).

Guy (1984, 1988, 1993) revised Korzhinskii’s

theory of metasomatic zoning by introducing a math-

ematical framework based on non-linear thermody-

namics. He discussed the assessment of appearance

and stability of compositional discontinuities repre-

senting the zoning, i.e., sharp reaction fronts, on an

isothermal fluid–solid fractionation curve. His ap-

proach shows that zoning might be inevitable in any

fluid-rock system that tends to attain local equilibrium

through infiltration, irrespective of the starting con-

ditions. In other words, the multiple steady state

attained by the system at the sharp front of reaction

is a self-organisational aspect of reaction-transport

formalism when the composition of fluid is a function

of composition of the rock at any time (this excludes

systems driven by dissolution-precipitation reaction).

Zoning can be defined simultaneously by changes in

mineralogy or by adjustments in the composition of

minerals that form solid solution series. Skarn depos-

its are, among all geological applications of this

theory, perhaps the most suited, since their develop-

ment implies reactive fronts with dynamic metasoma-

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370320

tism. Examples of zones defined by changes in the

composition of minerals in solid solution series, either

gradual or sudden, have been documented from skarns

(e.g., Guy, 1988).

2.3.2. Crystal zonation in skarn minerals

In skarns, crystal zonation may correlate with the

mineralising processes in the surrounding environ-

ment. Jamtveit (1991) suggested that non-ideal solid

solution between end-members in the grossular-an-

dradite series could explain the chaotic type of oscil-

latory zonation pattern in garnets he observed in the

Oslo Palæorift, Norway. This was in turn linked to

periodic changes in fluid composition. Zoning of

major and trace elements in garnet from the same

location were later ascribed to the evolution of the

hydrothermal system in P-T-fO2 terms (Jamtveit et al.,

1993). Jamtveit et al. (1995), re-investigated the

grandite solid solution series relative to aqueous

solution equilibrium, in order to establish correlations

between zonation patterns and skarn-forming fluids.

Although zonation patterns were considered as mainly

externally controlled, the supplementary role played

by self-organisation was acknowledged. The external

control was seen in fluctuations in Al/Fe-ratios of the

pore fluid, caused by variable rates of infiltration and

kinetic dispersion in the hydrothermal system, where-

as surface kinetics and local transport processes near

the crystal surface, as part of the spectrum of self-

organisation mechanisms, were invoked to explain

minor variations in garnet composition.

2.3.3. Reactive-infiltration applied to the skarn front

Dipple and Gerdes (1998) show that, at skarn fronts,

reaction-infiltration feedback (RIF) defines two reac-

tion parameters impacting on mineral reaction and

fluid production: over-pressure potential and change

in porosity. Infiltration-driven reactions thus have the

potential to produce transient fluid over-pressure, in-

dependent of the reactive capacity of the host rock, and

can either enhance or slow the flow. Also, large

increases in porosity at the skarn front, coupled with

focused flow parallel to the contact between intrusion

and limestone, assuming reactive-infiltration instabil-

ities provide for this, can potentially produce the

stacking of mineral reactions observed in banded skarn

patterns (Meinert, 1997). It is of note that decarbon-

ation may transiently produce a porosity increase of

more than 30%, whereas volatile-producing devolati-

lisation reactions can retard flow, or increase perme-

ability if volatile production is intense (Dipple and

Gerdes, 1998). Although the latter authors only mod-

elled fluid-producing reactions (i.e., wollastonite for-

mation), they postulated that fluid-consuming

reactions, such as those involved in formation of

volatile-rich phases, would generally enhance flow

by increasing the fluid pressure gradient. As a conse-

quence, sudden spurts of devolatilisation may be

produced, inducing back-flows and skarn-brecciation.

3. Description of the Ocna de Fier-Dognecea

orefield

The orefield (Fig. 1) contains more than 30 irreg-

ularly shaped orebodies situated within a narrow, 10

km long, NNE-SSW striking tract between the vil-

lages of Ocna de Fier and Dognecea. The orebodies

lie within the contact aureole of the Ocna de Fier-

Dognecea granodiorite intrusion and are located along

the boundary between Mesozoic limestone and Pre-

cambrian schists of the Bocs�it�a-Drimoxa Formation.

According to Nicolescu and Cornell (1999), the

skarns formed at a depth of about 10 km under an

estimated pressure of 2.8 kbar and a peak temperature

of 700F 50 jC. The age of mineralisation (76.6F 0.3

Ma, based on the Re-Os age of molybdenite; Ciobanu

et al., 2002a) coincides at the 2r level with the U/Pb

zircon age of the granodiorite (75.5F 1.6 Ma; Nic-

olescu et al., 1999).

Early workers mainly advocated a pyrometaso-

matic origin for the deposit (e.g., von Cotta, 1864;

Castel, 1869; Marka, 1869), although Sjogren (1886)

questioned this and ascribed ore genesis to regional

metamorphic processes. The skarn model was the

focus of numerous authors (e.g., Codarcea, 1930,

1931; Kissling, 1967; Vlad, 1974, 1994; Nicolescu,

1998; Nicolescu and Cornell, 1999; Nicolescu et al.,

1999; Ciobanu, 1999). Despite some agreement, no

single comprehensive model has been introduced,

which satisfactorily accounts for all features observed

across the entire orefield.

Mineral abbreviations used in the description be-

low and throughout this paper are in Table 1. The

mineralogy and distribution of skarn and ores within

the Ocna de Fier-Dognecea orefield are summarized

Table 1

Mineral abbreviations used in the text, tables and figure captions of

this paper

Act: actinolite Fa: fayalite Px: pyroxene

Alm: almandine Fe-Act:

ferroactinolite

Pxm: pyroxmangite

And: andradite Fo: forsterite Py: pyrite

Ank: ankerite Gah: gahnite Pyr: pyrope

Ap: apatite Gn: galena Qz: quartz

BD: bismuthinite

derivatives

Gr: grossular Sch: scheelite

Bi-ss: Bi-sulphosalts Grt: garnet Sid: siderite

Bn: bornite Hed: hedenbergite Sil: silicate

Bus: bustamite Hem: hematite Si-Mt: Si-bearing

magnetite

Cal: calcite Her: hercynite Sp: sphalerite

Carb: carbonate Ilv: ilvaite Spl: spinel

Cc: chalcocite Joh: johannsenite Sps: spessartine

Chl: chlorite Lw: ludwigite Srp: serpentine

Cl-Ap: chlorapatite MAS: magnetite-

ankerite selvage

Tlc: talc

Cp: chalcopyrite Mld: maldonite Tr: tremolite

Cpb: cuprobismutite Mt: magnetite Turn: turneaureite

Cz: clinozoisite Opx: orthopyroxene Wo: wollastonite

Di: diopside Ph: phlogopite

Ep: epidote Po: pyrrhotite

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 321

in Table 2, and the spatial position of orebodies and

samples described or discussed in this contribution are

shown schematically on Fig. 2.

Our description is restricted to the exoskarn

assemblages. Nevertheless, it should be noted that

minor bodies of endoskarn, hosted by granodiorite

apophyses and dykes, occur in numerous places

across the area. A single endoskarn occurrence

within the main granodiorite outcrops in the valley

Ogas�ul Vintilii, situated between Reichenstein and

Terezia, along the same strike as the other orebodies

(Fig. 1). Epidote and hematite are the main compo-

nents in the Ogas� ul Vintilii endoskarn. Similarly,

bands of epidote + garnet skarnoid, formed by the

metasomatic alteration of schist and containing spo-

radic ore mineralisation, are widely distributed at the

skarn-schist contact.

Throughout the orefield, the exoskarn is approxi-

mately 50 m thick and yielded 15 million tonnes of

Fe-dominated ore before exploitation ceased in 1993.

This figure includes some 2 million tonnes of Cu-Fe

ore. In the 18th century, some 250,000 tonnes of high-

grade copper ore (5.66% Cu) were won from the

cementation zone at the upper part of Simon Iuda.

Ocna de Fier, the northern and larger part of the

orefield, accounts for f 80% of the ore produced,

and comprises Cu-Fe ores (bornite-chalcopyrite-mag-

netite) hosted by magnesian (forsteriteF diopside)

skarn (Cook and Ciobanu, 2001), and Fe-ores hosted

by calcic (granditeF diopside) skarn (Fig. 2). The Fe

ores are found in each orebody, whereas the Cu-Fe

ores are restricted only to the deepest part of Simon

Iuda, between Ursoanea and � 120 ( + 158 m) levels.

Dognecea, the southernmost segment of the orefield,

comprises dominantly Zn-Pb ores hosted by Mn-

hedenbergite (diopside-hedenbergite-johannsenite se-

ries) skarn or limestone. Smaller bodies of this type of

skarn and ore also occur at Paulus (including Francis-

cus-Ignat�ius and Sofia; Table 2) in the northern part of

the orefield, as well as in the middle of the orefield, in

Grat�ianus and the median part of Simon Iuda orebody.

Monomineralic sulphide bodies are known, e.g., a

galena lens in the upper part of Petru and Pavel

(Castel, 1869), a pyrite body in the Reichenstein group

(Codarcea, 1930), and a Fe-rich sphalerite lens in

Paulus (level + 206 m). A more or less continuous

zone of Zn-Pb ore is seen in the upper parts of each

orebody across the entire orefield.

The orebodies have vertical extents of between 200

and 350 m and reach extinction at the base of

limestone in contact with crystalline schists. Lesser

vertical extents (f 100 m) are known from Ocna

Turceasca and Iuliana. In these cases, neither lime-

stone nor the orebody base was encountered during

exploitation (mining ceased because of difficulties

imposed by intense alteration at depth). The Reich-

enstein and Elias groups are the only ones developed

on both western and eastern sides of the limestone at

the contact to the schist (Fig. 1). Individual orebodies

of these two groups are connected at their deepest part

by branches following the limestone base at contact to

the schist. These orebodies have a concentric internal

structure: massive ores (60% oxidesF sulphides) in

the orebody core, enclosed by an outer zone typically

containing no more than 30% ore minerals. In contrast

to this onion-shell structure seen in each individual

orebody at Ocna de Fier, the mineralising style at

Dognecea is characterised by Pb-Zn-rich chimneys in

limestone at upper levels and low-grade pyroxene-rich

amass at depth. A lateral zonation is also defined for

Dognecea, based upon changes in skarn-ore mineral-

ogy (Vlad, 1974).

Table 2

Distribution of minerals and mineral assemblages within orebodies (from north to south), and zones of the Ocna de Fier-Dognecea deposit

Mineral abbreviations: see Table 1. Other abbreviations: p: prograde, r: retrograde.HVP: high-volatile phases; ETP: exotic trace phases, including Bi–Au–Ag–Co–Se–Te trace minerals in Simon Iuda; GBT: Au and Bi tellurides/selenides.

C.L.Ciobanu,N.J.

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s24(2004)315–370

322

Fig. 2. Schematic diagram showing the spatial arrangement, from south to north, of skarn- and ore-types within the individual orebodies of the

Ocna de Fier-Dognecea orefield. Positions of samples referred to in this publication are also shown. Corresponding mining levels across the

orefield are shown (not to scale).

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 323

Orefield-scale zonation continues to be debated. A

transitional trend, involving Fe-oxides to the north

(Paulus pit), sulphides and Fe-oxides in the middle

(Danila Hill), and sulphides to the south (Dognecea),

was initially reported by Castel (1869). Subsequent

workers (Vlad, 1994; Nicolescu and Cornell, 1999)

attributed this trend to primary, i.e., prograde, skarn

zonation. The latter authors describe the zonation as

‘‘grandite-magnetite (calcic-Fe) skarn at Ocna de Fier

in the north, mixed grandite-hematite Zn-Pb-(Cu)

skarn in the central section, and hedenbergite Zn-

Pb-(Cu) skarn at Dognecea in the south’’. We argue

that rather than representing a primary skarn zona-

tion, this apparent trend is due to several overlapping

factors. The factors include the reducing environment

offered by the change from gneiss to mica schist at

the limestone contact in the south, the extensive

replacement of hematite by magnetite, and not least

to the superimposed supergene alteration in the upper

part of Simon Iuda that contributed to hide the

(HedJoh)-rich character (see Section 4.3) of the

original Zn-Pb zone.

In contrast, Cook and Ciobanu (2001) proposed an

orefield zonation based on metal distribution, i.e., Cu-

Fe/Fe/Zn-Pb (Fig. 2) symmetrically centred onto Si-

mon Iuda. Interpretation of this centric zonation is

based upon recognition of a source of fluids in the

central and deepest part of orefield, subjacent to the

Simon Iuda body. The patterns are consistent with

emplacement outwards and upward from the source of

fluids, producing a proximal central segment and

distal segments in both north and south (Cook and

Ciobanu, 2001; Fig. 2). Zoning is accompanied by a

change in the dominant Fe-oxide from magnetite to

hematite towards the outer shells of the Fe-zone in

individual orebodies in the central part of the orefield

and the presence of hematite rather than magnetite in

distal garnet skarn (Fig. 2). Much of this hematite is

replaced by magnetite but it still recognisable because

of the pseudomorphous character of the replacement.

The deepest part of the Simon Iuda orebody dis-

plays several distinct petrological characteristics.

These include intimate co-genetic (poikilitic) relation-

ships between magnetite, Cu-Fe sulphides and for-

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370324

sterite, an abundance of valleriite, phlogopite, apatite,

ludwigite and other minerals indicative of an abnor-

mally high-volatile environment, and evidence for an

initial high-temperature (600 jC) bornitess that actedas a carrier for a range of exotic trace elements. These

features, described in detail by Cook and Ciobanu

(2001), not only underline the unique, innermost

position of the Cu-Fe ores, but also gives evidence

for a Cu-Fe core to the skarn system. This fundamen-

tal observation has major consequences for the origin

of the skarn system and for the spatial and temporal

trends in the deposit.

Such an interpretation is in sharp contrast with that

of previous authors. Cioflica et al. (1992) considered

that the Cu-ores in the central part of the orefield, i.e.,

Simon Iuda, were formed after the Fe-ores, from a

later surge of fluids that emerged from the granodio-

rite. This interpretation was an extension of the forma-

tional model introduced by Vlad (1974), who

formulated the concept of ‘individual hot centres’.

These centres were configured as clusters of apophy-

ses, each linked to the granodiorite, and which pro-

vided conduits for fluid springs. Vlad (1994) proposed

sequential emplacement for these hot centres, i.e.,

earliest and hottest in the north, latest and coolest in

the south, i.e., a pattern that was in agreement with the

aforementioned north to south zonation trend.

4. Skarn mineralogy

Although the present contribution is primarily

focused on textures, we nevertheless take the oppor-

tunity to present petrological evidence for centric

zonation to supplement that given by Cook and

Ciobanu (2001). Together with the textural data, the

observations in the following paragraphs (and also

Tables 3 and 4; Figs. 3 and 4) underpin the centric

zonation model and form the basis for reconstruction

of the skarn system in both space and time, the

ultimate goal of our investigation.

4.1. The Cu-Fe core

Bornite-chalcopyrite-magnetite ores in the Cu-Fe

core are hosted by forsteriteF diopside skarn. Unlike

in the other zones, the skarn is extremely patchy, with

variably dense pockets of mineralisation distributed

throughout the body. Forsterite (Fo95) is also seen as

droplets in magnetite. The magnetite has character-

istics typical of a ‘magmatic’ affiliation (Zharikov,

1970), including MgO contents up to 6 wt.%, and

fields of exsolved skeletal spinels. We note that, in

some of the larger two-component inclusions (Fig.

5a), the presence of forsterite with as much as 30%

fayalite component is accommodated by higher spi-

nel–hercynite ratios, at constant gahnite components

of f 20 mol%. We recognise a clear distinction

between these relatively simple associations, which

we refer to as prograde, and more complex associa-

tions that are clearly superimposed (i.e., retrograde in

origin).

Alongside widespread serpentinization of forster-

ite, the forsterite-magnetite-sulphide assemblages are

crosscut by ubiquitous clusters of apatite, together

with magnetite, sulphides and even newly formed

forsterite and diopside, which give the overprinted

intergrowths the appearance of a symplectite (Fig. 5b).

Much of the apatite that overprints the magnetite-

sulphide assemblage in the core is the As-bearing

variant, turneaureite, Ca5[(As,P)O4]3Cl. Coexisting

magnetite has a considerable Mn content in the Cu-

Fe core (Fig. 5b; up to 6 wt.% MnO), a feature shared

with the Fe skarn at Langban, Sweden, the type

locality for turneaureite (Dunn et al., 1985). Individual

blebs of apatite in the symplectite clusters are strongly

zoned, with As-poor, Cl-rich cores and As-rich rims

(up to 16 wt.% As; Fig. 5c).

Even though diopside is rare in the core, compared

to forsterite, it is nevertheless observed in both pro-

grade and retrograde assemblages. Compositions (Ta-

ble 4) fall within a limited range (Di93). Diopside can

be formed in equilibrium with apatite (Fig. 5d). We

note that the prograde diopside of the Cu-Fe core has

the highest Al2O3 content (3.13 wt.%; 0.14 Aliv per

formula unit) among all pyroxenes in the orefield.

Pyroxenes from skarns generally have Al2O3 contents

well below 1 wt.% (e.g., Nakano et al., 1994).

Elevated Al2O3 contents in skarn pyroxene (as much

as 24 wt.% Al2O3) from two other occurrences in the

Banatitic Magmatic and Metallogenetic Belt, at

Magureaua Vat�ei and Ciclova, have recently been

reported (Katona et al., 2003). These authors consid-

ered the compositions to be concordant with forma-

tion at very high temperatures (f 800 jC) and high

CO2 activities in the fluid.

Table 3

Mean composition of skarn garnets from Fe zone (columns 1 –16) and Zn–Pb zone (columns 17–21)

Orebody 1 2 3 4 5 6 7 8 10 11 12 13 14 15 16 17 18 19 20 21

sampleMagdalena CGM Ocna Turceasca Li Stefania Elias Mijlociu Reichen Jupiter Paulus Simon Iuda Paulus Dognecea

e1oz and r

(n = 14)

e17

(n = 6)798oz and r

(n= 25)

3101a

(n= 6)

e14oz and r

(n= 12)

St1r

(n = 12)

endosk

3089*

(n = 6)

3089a

(n= 7)

3077r

(n = 8)

3083

(n = 4)

stein

GR*oz

(n= 8)

179

(n= 11)5598oz and r

(n= 20)

82oz and r

(n = 18)

72r

(n= 7)

3913oz

(n= 13)

3364oz and r

(n= 8) endosk

3375oz and r

(n = 9)

40

(n= 7)

3786

(n= 8)

Oxide (wt.%)

SiO2 40.16 35.57 38.76 39.73 40.16 40.67 33.71 39.66 40.94 38.50 32.83 39.79 38.51 39.66 39.96 42.27 40.33 40.36 39.61 38.93

Al2O3 5.15 0.33 3.35 1.56 1.71 4.87 0.25 0.82 5.60 0.12 1.39 0.81 5.50 8.07 7.93 9.16 9.76 7.68 10.26 4.12

MgO 0.06 0.04 0.39 0.15 0.08 0.08 0.30 0.21 1.08 0.25 0 0.08 0.08 0.03 0.02 0 0.06 0.03 0 0.04

FeO 20.17 28.66 23.87 23.37 19.65 26.71 24.26 19.02 24.67 26.07 24.82 21.05 16.62 18.20 15.84 18.22 14.33 21.36

Fe2O3 22.42 31.85 23.70 26.53 25.98 21.84 29.68 26.96 21.14 27.41 28.97 27.59 23.39 18.47 20.23 14.76 17.60 20.25 15.93 23.74

MnO 1.38 1.01 1.04 0.37 1.10 2.08 0.44 1.33 0.60 0.76 1.30 1.34 0.87 2.50 3.37 1.39 0.45 2.49 0.68 2.02

CaO 32.83 34.55 32.71 34.22 33.52 32.47 33.48 33.17 32.32 33.88 32.71 33.30 33.68 32.90 30.28 31.80 32.59 30.91 33.29 33.22

Na2O 0.03 0.02 0 0.05 0.07 0.03 0 0 0.02 0.14 0 0.07 0.07 0.04 0.04 0 0.14 0.05 0.04 0.13

K2O 0 0 0 0 0.03 0 0 0 0 0 0 0.01 0.02 0 0 0 0.05 0 0.02 0.03

TiO2 0 0 0.17 0.15 0.05 0.12 0 0 0.15 0 0 0.02 0.16 0.15 0.16 0.04 0.64 0.08 0 0.07

Total 102.02 103.37 100.13 102.76 102.70 102.16 97.86 102.25 101.84 101.05 97.20 103.00 102.28 101.84 101.99 99.43 101.61 101.85 99.94 102.32

Formula based on O= 6

Si 3.19 2.93 3.17 3.19 3.22 2.97 2.93 3.21 3.22 3.17 2.88 3.20 3.08 3.13 3.15 3.32 3.16 3.18 3.13 3.12

Fetotal 1.34 1.98 1.46 1.60 1.57 1.23 1.94 1.64 1.26 1.70 1.91 1.67 1.41 1.10 1.21 0.88 1.04 1.21 0.95 1.44

Al 0.48 0.03 0.32 0.15 0.16 0.37 0.03 0.08 0.52 0.01 0.14 0.08 0.52 0.75 0.73 0.84 0.89 0.71 0.96 0.38

Fe+ + 1.34 1.97 1.46 1.60 1.57 1.23 1.94 1.64 1.26 1.70 1.91 1.67 1.41 1.10 1.21 0.88 1.04 1.21 0.95 1.44

Total 1.82 2.00 1.78 1.75 1.73 1.60 1.97 1.72 1.77 1.71 2.05 1.75 1.93 1.85 1.94 1.72 1.93 1.91 1.90 1.83

Ca 2.79 3.05 2.87 2.95 2.88 2.54 3.12 2.88 2.73 2.99 3.07 2.87 2.89 2.79 2.56 2.68 2.74 2.61 2.82 2.87

Mg 0.01 0.01 0.05 0.02 0.01 0.01 0.04 0.03 0.13 0.03 – 0.01 0.01 0 0 – 0.01 0 0 0.01

Mn 0.09 0.07 0.07 0.03 0.07 0.13 0.03 0.09 0.04 0.05 0.10 0.09 0.06 0.17 0.23 0.09 0.03 0.17 0.05 0.14

Fe+ 0 0.01 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

Total 2.89 3.13 2.99 2.99 2.96 2.68 3.19 2.99 2.89 3.08 3.17 2.97 2.96 2.96 2.79 2.77 2.78 2.78 2.88 3.01

% And 71.8 95.7 79.0 90.3 88.5 71.7 96.5 92.0 67.0 96.7 90.2 92.4 71.6 55.9 56.7 49.3 52.9 59.3 49.2 75.8

% Gr 24.8 1.6 17.1 8.2 8.6 23.2 1.3 4.1 27.3 0.6 6.8 4.2 26.1 38.3 35.1 47.4 42.6 34.6 48.8 19.4

% Sps +

Alm +Pyr

3.5 2.7 3.9 1.4 2.8 5.1 2.2 3.9 5.7 2.7 3.0 3.4 2.3 5.8 8.1 3.3 4.5 6.1 2.0 4.8

CGM: Composite garnet –magnetite; Li: Liesegang banding; oz: oscillatory zonation; r: retrograde overprint.

*Microprobe analyses; all others by SEM-EDS.

C.L.Ciobanu,N.J.

Cook/Ore

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s24(2004)315–370

325

Fig. 3. Ternary diagram summarising the composition of skarn garnet from Ocna de Fier-Dognecea, in terms of the end-members andradite,

grossular and (spessartine + almandine + pyrope). Mean analyses are plotted for each sample (see also Table 3). Inset shows compositional

variation within one individual sample with oscillatory zonation and retrograde overprint seen in absorption-corrosion boundaries (5598).

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370326

4.2. The Fe zone

In the Fe zone, the main components are garnet,

pyroxene, magnetite and hematite. Pyrite is the single

sulphide that, in some cases, is a major component of

the prograde associations. Relationships between Fe-

oxides and pyrite can, however, be rather complex,

implying local disequilibrium and/or re-equilibration

during the subsequent retrograde stage. Certain

assemblages can also contain other silicates, such as

epidote, tremolite and ferroactinolite, in equilibrium

with garnet or pyroxene, e.g., epidote in endoskarn or

skarnoid rocks.

Compositional variation in garnet (Table 3, Fig. 3)

is constrained by a number of factors, e.g., associa-

tion, oscillatory zonation and retrograde overprinting.

This is also true for pyroxene (Table 4, Fig. 4),

although this mineral lacks oscillatory zonation. Mu-

tual relationships between garnet and pyroxene can be

complicated by the relative proportion and grain size

of the components, especially when the assemblage

undergoes subsequent recrystallisation. ‘Cores’ dis-

playing oscillatory zonation (Fig. 5e) are seen irreg-

ularly within the garnet. Oscillatory zonation is not

only a feature of prograde garnet, but can also be a

retrograde manifestation (e.g., Fig. 5e). Although both

zones of the garnet in this example have the same

compositional range (i.e., And90-70), textural criteria

for discrimination between a prograde core and retro-

grade overgrowths, can be readily applied here, unlike

in other cases.

Garnet lacking oscillatory zonation is widespread

in two-component garnet-magnetite or garnet-hema-

tite associations, especially in the inner part of an

Table 4

Mean composition of skarn pyroxenes

Orebody 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20

sampleCu–Fe core Fe zone Buffered inclusions in Fe zone Zn–Pb zone

Simon Iuda P&P Ocna Turceasca Mijlociu Paulus SI OT Stefania Paulus Proximal Distal

Simon Iuda Dognecea Paulus

164

(n= 1)

73*

(n= 5)

PP

(n= 9)

3101*

(n= 9)

e17

(n= 1)

798

(n= 6)

3078

(n= 17)

3077

(n= 9)

3082

(n= 13)

5598

(n= 18)

72

(n= 8)

3762r

(n= 10)

St1r

(n= 11)

82r

(n= 17)

3093

(n= 18)

71a

(n= 14)

126

(n= 16)

3596

(n= 18)

Dogn

(n= 25)

31

(n= 41)

Oxide (wt.%)

SiO2 57.36 50.29 56.01 53.86 54.06 56.09 56.38 52.38 52.57 53.53 52.14 54.28 54.84 53.57 53.16 53.52 52.26 53.62 52.52 51.77

Al2O3 1.45 3.13 0.82 0.16 0.07 0.19 0.55 1.12 1.33 0.76 0.37 0.04 0.12 0.33 0.03 0.04 0.42 0.08 0.10 0.57

MgO 18.00 16.10 16.40 17.93 14.04 15.36 15.68 13.26 10.92 13.02 7.17 8.25 8.40 7.60 2.82 3.78 7.91 6.28 0.57 1.80

FeO 1.9 1.84 1.70 1.41 3.12 3.20 3.13 7.43 8.51 6.77 10.21 11.45 9.23 10.95 10.42 10.95 3.72 12.54 16.83 18.07

MnO 0.49 0.20 0.51 0.40 1.81 1.39 0.40 0.42 0.52 0.65 6.95 2.81 3.87 3.68 11.23 8.65 11.32 4.45 7.69 5.49

CaO 19.44 26.14 24.48 26.66 25.98 23.93 23.91 25.00 25.35 24.89 23.04 22.74 23.46 22.95 22.27 22.29 24.32 22.88 22.11 21.87

Na2O 0.59 0 0.05 0.02 0.00 0 0.05 0.03 0.06 0.12 0.09 0.10 0.10 0.61 0.09 0.10 0.66 0.13 0.13 0.38

K2O 0 0 0.05 0 0.00 0 0.02 0 0.01 0.02 0.00 0 0 0.04 0.03 0 0.06 0.03 0.02 0.00

TiO2 0 0.30 0.05 0 0.04 0.09 0.04 0.02 0.06 0.11 0.06 0 0 0.02 0.04 0.02 0.12 0.05 0.06 0.01

Total 99.23 98.00 100.07 100.44 99.12 100.25 100.15 99.67 99.34 99.87 100.02 99.66 100.06 99.76 100.09 99.34 100.78 100.06 100.03 99.95

Formula based on O= 6

Si 2.06 1.88 2.02 1.96 2.01 2.04 2.04 1.97 1.99 2.00 2.02 2.07 2.07 2.06 2.09 2.10 2.01 2.07 2.09 2.06

Al 0.06 0.14 0.03 0.01 0.00 0.01 0.02 0.05 0.06 0.03 0.02 0.00 0.01 0.01 0.00 0.00 0.02 0.00 0.00 0.03

Mg 0.96 0.90 0.88 0.97 0.78 0.83 0.85 0.74 0.62 0.72 0.41 0.47 0.47 0.43 0.16 0.22 0.45 0.36 0.03 0.11

Fe 0.06 0.06 0.05 0.04 0.10 0.10 0.09 0.23 0.27 0.21 0.33 0.37 0.29 0.35 0.34 0.36 0.12 0.40 0.56 0.60

Mn 0.01 0.01 0.02 0.01 0.06 0.04 0.01 0.01 0.02 0.02 0.23 0.09 0.12 0.12 0.37 0.29 0.37 0.15 0.26 0.19

Ca 0.75 1.05 0.95 1.04 1.04 0.93 0.93 1.00 1.03 1.00 0.96 0.93 0.95 0.94 0.94 0.94 1.00 0.95 0.94 0.93

Ti – 0.01 0.00 – 0.00 0.00 0.00 0.00 0.00 0.00 0.00 – – 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Total 1.78 2.02 1.90 2.07 1.97 1.91 1.88 2.00 1.93 1.95 1.93 1.85 1.84 1.85 1.82 1.80 1.94 1.86 1.80 1.83

Na +K 0.02 – 0.00 0.00 0.00 – 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.02 0.00 0.01 0.01

% diopside 93.0 93.4 93.0 94.6 83.5 85.6 88.8 75.7 68.3 75.6 42.5 50.4 52.7 47.8 18.7 25.3 45.3 39.5 3.9 11.8

% hedenbergite 5.5 6.0 5.4 4.2 10.4 10.0 9.9 23.0 29.9 22.3 34.1 39.7 33.1 39.0 38.8 41.6 17.3 44.5 65.6 67.5

% johansonnite 1.4 0.6 1.6 1.2 6.1 4.4 1.3 1.4 1.8 2.2 23.4 9.9 14.2 13.2 42.5 33.1 37.4 16.0 30.5 20.7

OT: Ocna Turceasca; P&P: Petru and Pavel; SI: Simon Iuda; r: retrograde overprint.

*Microprobe analyses; all others by SEM-EDS.

C.L.Ciobanu,N.J.

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s24(2004)315–370

327

Fig. 4. Ternary diagram summarising the composition of skarn pyroxenes from Ocna de Fier-Dognecea, in terms of the end-members diopside,

hedenbergite and johannsenite. The diagram illustrates the changes in pyroxene composition from Cu-Fe core to Fe and Zn-Pb zone and from

proximal to distal within the Zn-Pb zone. Individual analyses are plotted (see also Table 4). Note the extended compositional fields for the Zn-Pb

zone in comparison with the others.

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370328

orebody (e.g., sample 3083, Mijlociu body), where it

has a characteristic high andradite (And>90%) com-

position (Fig. 3). Monomineralic garnet, which may

display oscillatory zonation, has a rather limited

compositional range (e.g., And95-83, sample GR),

plotting within the same cluster on Fig. 3 as the cases

mentioned above. In contrast, garnet in prograde

associations with pyroxene, with or without coexist-

ing magnetite or hematite, commonly displays oscil-

latory zonation with broader compositional ranges

extending towards And50, although means for entire

samples are in the interval And80-70 (Fig. 3). Such

associations are characteristic for the lower-grade (25

to 30% Fe) margins of high-grade iron ores (samples

3077, 3078, 3082; Mijlociu orebody). Although sub-

sequent reshaping can be seen in zones with absorp-

tion/corrosion boundaries, the overall composition of

garnet appears little affected (Fig. 5e). We also rec-

ognise newly formed garnet, not part of prograde

association, but rather enclosed within calcite that

cements brecciated aggregates of magnetite. Oscilla-

tory zonation in such cases covers a relatively narrow

range of Gr-rich composition (e.g., And63-48 in sample

82; Paulus).

Hematite ore (mostly converted into magnetite)

was prevalent in the upper part of Elias, in Magdalena

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 329

orebody, in the upper part of Iuliana and Ocna

Turceasca, and in the lowest part of orebodies in

Paulus, as well in the lower part of Dognecea skarn

(Fig. 2). Preservation of prograde hematite in garnet-

pyroxene skarn is rare because of its more or less

ubiquitous replacement by magnetite. Nevertheless,

we can recognise the presence of former hematite in

the lamellar contours seen within magnetite (e.g., e1,

5598). In such cases, we consider that only relict high-

andradite garnet (i.e., And90-98) coexisted with hema-

tite. Compositions in the range And70-50 belong to

garnet coexisting with later magnetite (e.g., sample

5598 in Fig. 3, inset). The coexistence of pyrite with

hematite in garnet skarn is also observed (e.g., e14;

Ocna Turceasca).

In endoskarn, garnet coexists with epidote, hema-

tite (3364, 3383) and pyrite (e15, St1). As depicted in

Fig. 3, such garnet plots over an extended range (i.e.,

And80-55. Only in one sample (3364) is garnet the

main component. The average composition is towards

the grossular-rich field even though zones of And95exist in oscillatory-zoned crystals. This large variation

probably relates to the type of intrusive dyke that has

been metasomatised. Taking relict minerals of former

magmatic origin into consideration, we can approxi-

mate that endoskarn in samples e15 and 3364 formed

from dioritic dykes, whereas 3383 and St 1 were

tonalitic to granitic in composition.

To summarise, prograde garnet in exoskarn from

the Fe zone is represented by high-andradite (And>90)

when associated with Fe oxides, but is diluted by

grossular molecule in the range Gr20-30 when associ-

ated with pyroxene (Table 3). The Mn content of

garnet from the Fe zone never exceeds 10% spessar-

tine component.

Similarly to garnet, the composition of pyroxene

associated only with magnetite has a limited compo-

sitional range, Di>90 (e.g., Di95-93 in samples PP,

3101), very close to the pyroxene in the Cu-Fe core

(Fig. 4, Table 4). Although pyroxene associated with

garnet may have a comparable composition, other

individual analyses of pyroxene associated with gar-

net show a wider and continuous compositional

spread with hedenbergite components of Hed10-40 (at

Joh< 10). From three samples in Mijlociu orebody

representative of this association, we see that, in the

presence of abundant magnetite, the Hed component

is lowest (Hed10; 3078), whereas pyroxene enclosed

in a garnet matrix (3077, 3082) has Hed20-30. Simi-

larly, pyroxene in association with garnet in samples

with pseudomorphed hematite has comparable Hed

contents (i.e., Hed20-30; 5598, e1).

Unlike garnet, pyroxene shows a distinct compo-

sitional trend in which the Joh contents in samples

from upper levels of each orebody (e.g., 3762 in Ocna

Turceasca, St1 in Stefania, 72 in Simon Iuda) are

significantly higher (Joh10-20). The diopside compo-

nent is diluted to Di50. However it is difficult to

establish whether this pyroxene is in equilibrium with

the other components in the association since it is seen

only as inclusions in one or the other mineral. More-

over as we show in Fig. 5f, these 10 to 20 Aminclusions, unlike the coarse pyroxene discussed

above, have rather complicated zonation patterns with

absorption boundaries. In Fig. 5f the pyroxene is

enclosed within garnet that itself is enclosed within

pyrite. It shows zonation with absorbed diopside-rich

cores and outer zones, with varying Hed–Joh ratios.

Associated garnet is And-rich (And>90), but includes

grossular-enriched zones (approaching Gr40Sps10).

Such zones have corrosion boundaries against the

And-rich zones. Bundles of hematite lamellae are also

enclosed within the pyrite. In sample 3762, the main

component is hematite (converted to magnetite) with

coarse grains of pyrite that host the pyroxene. Some

indication of prograde silicate chemistry comes from a

sample from Simon Iuda (72). Although primary

hematite is preserved, silicates have been extensively

replaced. Minute skeletal inclusions of both pyroxene

(Di50Hed30Joh20) and garnet (And40Gr55Sps5) are

nevertheless preserved within coarse pyrite. In a

further example, from Paulus (82), relict pyroxene

and garnet, with similar compositions as in the above

example, occur enclosed within a mass of calcite

cementing the magnetite aggregates.

We conclude that pyroxene from the Fe zone is

Mg-rich (Di>90) when associated with magnetite, but

is diluted by Hed (typically Hed10-30) when associ-

ated with garnet. The coexistence of magnetite

appears to diminish the Hed component in pyroxene

more than does the associated garnet. The diopside

content is further diluted at upper levels in the Fe

zone by a significant Joh component. The trend of

increased Joh component as the limit towards the Zn-

Pb zone gets closer and pyroxene composition is

buffered by pyrite (the host for pyroxene inclusions)

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370330

is evident, as is the marked retrograde overprinting in

such associations.

Magnetite in the Fe zone has contents of SiO2,

Al2O3, MnO, CaO, and MgO that can approach 10

wt.% combined. The relative proportion of these

elements can vary but SiO2 is always higher, some-

times as much as 5 to 6 wt.%. Substitution of Si in the

magnetite structure is known from magnetite deposits,

including skarns (e.g., Shiga, 1989; Westendorp et al.,

1991; Shimazaki, 1998).

Epidote shows little compositional variation with

skarn type, is high-clinozoizite in composition, e.g.,

Clz72Ep27 in endoskarn (samples: 3364, 3677, e15,

St1, 3383) and Clz66Ep34 in exoskarn (samples: 82,

3596, 5598). Even though epidote is only weakly

zoned, peculiarly Ce-bearing zones (Ce2O3 < 2 wt.%)

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 331

are noted in epidote from Stefania (St1). Nevertheless,

in exoskarn all these minerals are most frequent in

associations formed during the retrograde stage. Cal-

cite and quartz are ubiquitous components in almost

all associations, in variable proportions. Although

simple one- or two-component associations of the

above minerals in various combinations are quite

frequent, the Fe zone typically contains ternary

assemblages consisting of garnet, pyroxene and one

or the other Fe oxide.

With the exception of valleriite, minerals such as

phlogopite, ludwigite, serpentine and apatite are also

accessories in the Fe zone. Others, such as titanite are

particularly abundant. But, as seen for example in Fig.

5g, apatite is formed at eutectic equilibrium with

calcite. Exsolution of calcite in apatite and their

mutual relationships suggests that both minerals are

formed from a carbonate-rich ‘residuum’, the appear-

ance being somewhat reminiscent of veins in carbo-

natite skarn (Lentz, 1998). Apatite is most abundant in

magnetite where it can be seen as fronts, swarms,

pockets with well-shaped crystals. However, unlike in

the Cu-Fe core, the appearance of apatite, ludwigite,

etc., is not associated with enrichment in the range of

exotic trace elements (Au, Bi, Te, Se, etc., see below).

4.3. The Zn-Pb zone

We define this zone as consisting of mineral

associations in which sphalerite and galena are dom-

inant over Fe-oxides or Cu-Fe sulphides. This zone

therefore extends beyond the skarn/limestone contact

to include Pb-Ag-rich ores that had been historically

exploited from uppermost levels, e.g., at Dognecea

Fig. 5. Back-scattered electron images showing textures from the Cu-Fe co

sample numbers are given in brackets. (a) Two-component inclusions con

crosscutting forsterite and magnetite assemblage. Note the resulted ‘sympl

emplacement. (c) Characteristically zoned apatite, consisting of a chlorap

chalcopyrite. (d) Newly formed diopside in forsterite (partially serpentini

oscillatory zonation in garnet. Note the similarity in composition in the prog

Numbers refer to the And component in garnet. (f) Zoned inclusion of

absorbed Di-rich core in pyroxene (St1). (g) Typical occurrence of chlora

note also the calcite exsolution in apatite (174). (h) Co-existing pyroxene

assemblage (460 m; Simon Iuda). Note the curvilinear equilibrium boun

eutectic decomposition (126). (i) Pyroxene coexisting with pyroxmangit

bustamite is absent. Note the retrograde overprint in pyroxene (71a; Simo

skarn from the north of the deposit. Their composition corresponds to the

same sample as (j). Note the presence of pyroxmangite as inliers within h

zonation) within high-FeS sphalerite from distal skarn in Paulus (40).

(fide Kissling, 1967; Vlad, 1974). In this section, we

discuss the silicate mineralogy of the zone in an

attempt to establish the differences between the Zn-

Pb and Fe zones throughout the deposit. In the central

part of the orefield, the zone is placed at the upper and

to some extent lateral part of each orebody (Fig. 2).

However, because the upper part of each orebody is

more or less affected by supergene alteration, it is not

always easy to establish the original extent and

character of this zone. This is particularly true for

Simon Iuda, where Cu-rich ores exploited from the

upper part of the orebody (from the 470 m level down

to a depth of 60 to 80 m) were formed by secondary

enrichment processes with the corresponding zone of

metal leaching situated underneath the Reichenstein

III gallery (357 m), and above the Ursoanea gallery

(278 m).

The main minerals in Zn-Pb skarn are pyroxene

(Table 4, Fig. 4), sphalerite (Table 5), galena and

pyrite. This zone is highly inhomogeneous, with other

minerals such as pyroxenoids (bustamite, pyroxman-

gite; Table 6), tremolite and garnet (Table 3, Fig. 3), as

well as hematite and minor magnetite, present in

variable amounts.

In the Zn-Pb zone, in comparison to the Fe zone, a

manganese-enriched character is shown by the ubiqui-

tous increase in the Joh component of pyroxene, i.e.,

Joh20-40, (Table 4) as well as by the presence of

pyroxmangite in all associations. Pyroxene shows

extensive compositional variation in the Zn-Pb zone

(Fig. 4), with a discernable difference between proxi-

mal and distal skarn. In proximal skarn from Simon

Iuda, pyroxene has a previously unrecognised Joh

component that is the highest (Joh40) in the entire Zn-

re (a to d: 164), Fe-zone (e to g) and Zn-Pb zone (h to l). Respective

sisting of spinel and olivine within magnetite. (b) Cluster of apatite

ectite’-like appearance in the resulting assemblage due to the apatite

atite core with a turneaureite rim, within a matrix of magnetite and

sed) within one of the apatite clusters. (e) Prograde and retrograde

rade core and retrograde pressure tail in upper right of picture (798).

pyroxene in garnet (itself enclosed within coarse pyrite). Note the

patite in the Fe-zone, showing equilibrium boundaries with calcite;

s from proximal Zn-Pb zone. Bustamite is the main silicate in the

daries within lamellae. The texture is interpreted as the product of

e and sphalerite: pyroxene is the main silicate in the assemblage;

n Iuda, 357 m). (j) Lamellar intergrowth of two pyroxenes in distal

two clusters in Fig. 4 (31). (k) Pyroxene adjacent to hematite in the

ematite (31). (l) Inclusion of Gr50 garnet (showing weak oscillatory

Table 5

Mean composition of sphalerite

Element (wt.%) Dognecea Stefania Simon Iuda Grat�ianus Paulus

Dogn** 3786 3786a 3596** Stef** 3093** 71a** 126** 3910 54 g 40**

(n= 5) (n= 15) (n= 9) (n= 4) (n= 4) (n= 3) (n= 4) (n= 14) (n= 7) (n= 2) (n= 21)

S 33.93 33.30 31.77 36.21 36.16 35.94 35.88 36.26 32.62 37.55 32.72

Mn 0.48 0.46 0.37 0.44 0.60 0.51 0.53 0.92 0.43 0.29 0.37

Fe 8.355 8.79 3.19 4.61 5.81 2.91 2.69 3.47 3.28 2.52 9.28

Co 0.11 0.08 0.01 0 0 0 0 0.00 0.09 0 0.11

Cu 3.29 0.23 0.77 0.99 1.47 0.84 0.55 0.00 0.53 1.35 0.82

Zn 51.16 54.32 60.77 54.90 52.33 56.80 57.01 56.34 59.31 56.73 53.24

Se 0 0.09 0.19 0.23 0 0 0 0.24 0.11 0 0.05

Ag 0 0.05 0.07 0 0 0 0 0.00 0.08 0 0.08

Cd 0.36 0.38 0.50 0.35 0.47 0.14 0.43 0.44 0.51 0.57 0.26

Total 97.69 97.70 97.64 97.73 96.84 97.12 97.07 97.68 96.96 99.00 97.04

% ZnS 78.4 82.6 92.0 88.5 84.9 92.0 92.6 91.2 91.7 91.9 81.0

% MnS 0.9 0.8 0.7 0.8 1.2 1.0 1.0 1.8 0.8 0.6 0.7

% FeS 15.0 15.7 5.6 8.7 11.0 5.5 5.1 6.6 5.9 4.8 16.5

% CdS 0.3 0.3 0.4 0.3 0.4 0.1 0.4 0.4 0.5 0.5 0.2

% (Co,Cu,Ag)S 5.4 0.5 1.3 1.6 2.5 1.4 0.9 0.0 1.1 2.2 1.5

T (jC, estimated) 400 440 470 570 570 570 400

log fS2 � 6.79 � 5.19 � 4.53 � 1.61 � 1.55 � 1.74 � 6.86

log fO2 � 24.14 � 21.74 � 20.74 � 16.36 � 16.27 � 16.56 � 24.24

**Co-existing with pyroxene, all data by SEM-EDS.

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370332

Pb zone, whereas in distal skarn pyroxene has the

highest Hed values, i.e., Hed70 (Table 4). Although

Mn/Fe ratios in pyroxene are generally considered to

reflect the ratio in the original fluid (e.g., Nakano et al.,

1994), the different trends for proximal and distal skarn

can be considered a result of divergent fluid evolution.

Even though compositional fields are characteristic

for all pyroxenes from the Zn-Pb zone (Fig. 4), this

aspect is represented differently in individual assemb-

lages. Proxene stable with bustamite in the uppermost

part of Simon Iuda (126) has a mean composition that

is Di-rich, Hed-poor (Di40Hed20Joh40). This pyrox-

ene, unusual for the Zn-Pb zone, undergoes eutectic

decomposition as shown by the two compositional

types of pyroxene separated by curvilinear boundaries

within individual lamellae (Fig. 5h). In the absence of

bustamite, 100 m vertically below, in the lower part of

Zn-Pb zone in Simon Iuda, we note a different

pyroxene, with higher proportions of Hed and lower

Di, but with the same Joh component as at upper

levels (Di20Hed40Joh40). Decomposition is recognised

in this pyroxene as well, together with retrograde

overprinting (Fig. 5i).

On the contrary, the pyroxene forms coarse and

homogenous lamellae in distal skarn. In Paulus, even

though the dominant pyroxene has a composition

(Di10Hed70Joh20) close to that from Dognecea

(Di< 10Hed60Joh30), a second pyroxene is also present

in the association. This has the composition

Di20Hed70Joh10, (Fig. 4). The two pyroxenes form

lamellar intergrowths (Fig. 5j). We also mention the

presence of pyroxene lamellae that lack any Di

component, i.e., Hed75Joh25. In Dognecea, the pyrox-

ene appears homogenous despite the broad composi-

tional field. However, we observe that the greatest

variance is between the Hed and Di components, like

in proximal skarn, rather than between the Joh and Di

components as in Paulus.

In Dognecea, the Joh component of pyroxene

typical for distal Zn-Pb ore is almost constant at

around 30% and also has constantly the lowest Di

values, typically Di< 10. In comparison to this, pyrox-

ene from Dognecea North, which is stable with

epidote and pyrite as the dominant sulphide, has a

higher Di component and lower Joh (Di40Hed40-Joh20). Such an intermediate zone containing pyrox-

ene richer in Di was first recognised at Dognecea by

Vlad (1974).

As shown in Fig. 5i and k, pyroxene from Zn-Pb

zone coexists with pyroxenoid, sphalerite and hema-

Table 6

Mean composition of bustamite (column 1) and pyroxmangite (columns 2 to 7)

Orebody sample 1 2 3 4 5 6 7

Simon Iuda Dognecea Paulus

126d* (n= 4) 71a (n= 3) 72a (n= 3) 3093 (n= 5) 126 (n= 5) Dogn (n= 7) 31 (n= 4)

Oxide (wt.%)

SiO2 45.37 51.29 48.51 50.92 48.78 50.91 50.11

Al2O3 0 0.04 0.06 0.15 0.74 0.11 0.46

MgO 0.32 1.52 1.56 0.46 1.95 0.47 0.94

FeO 2.68 7.25 6.77 6.31 5.80 10.67 7.25

MnO 30.58 32.50 38.87 31.11 33.56 29.28 32.84

CaO 17.37 5.72 3.76 10.02 9.35 7.45 7.88

Na2O 0 0.07 0.21 0.10 0.63 0.09 0.53

K2O 0 0 0.08 0.08 0 0.05 0

Total 96.32 98.39 99.82 99.16 100.81 99.11 100.00

Formula based on O=18

Si 5.92 6.38 6.14 6.31 6.04 6.34 6.22

Al – 0.01 0.01 0.02 0.11 0.02 0.07

Mg 0.06 0.28 0.29 0.09 0.36 0.20 0.17

Fe 0.29 0.75 0.72 0.65 0.60 0.96 0.75

Mn 3.38 3.42 4.17 3.27 3.52 3.11 3.45

Ca 2.43 0.76 0.51 1.33 1.24 1.01 1.05

Total 6.16 5.22 5.69 5.34 5.72 5.28 5.43

Na – 0.01 0.03 0.01 0.08 0.01 0.06

K – – 0.01 0.01 – 0.00 –

Total – 0.01 0.03 0.02 0.08 0.02 0.06

% MnSiO3 54.8 65.6 73.3 61.2 61.5 58.9 63.6

% MgSiO3 1.0 5.4 5.2 1.6 6.3 3.8 3.2

% FeSiO3 4.7 14.4 12.6 12.3 10.5 18.2 13.9

% CaSiO3 39.4 14.6 9.0 24.9 21.7 19.1 19.3

*Microprobe analyses; all others by SEM-EDS.

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 333

tite. The Hed-rich compositions of our associations

are somewhat at odds with characteristically Di-rich

pyroxene previously reported from skarns in which

hematite is stable (Einaudi and Burt, 1982). We

therefore assume that the Joh component has signif-

icantly increased the stability field of the Hed com-

ponent in pyroxene under sulphidation conditions, a

role documented by experiment (Burton et al., 1982).

We note some variability in the Fe/Ca ratio in pyrox-

mangite, which is higher in Dognecea than in Simon

Iuda (Table 6). In all samples pyrite is the stable Fe

sulphide. The FeS content of sphalerite varies between

5 and 16 mol%, positively correlating with the Hed

component of coexisting pyroxene (Table 5). There is

also a direct correlation between the Di( + Joh) content

of pyroxene and mol% FeS content of coexisting

sphalerite (Table 5), which can be used to express

variation in fS2 (Gamble, 1982).

We conclude that pyroxene compositions with

higher Hed and Joh components than in the Fe zone

(within the ranges Hed40-70 and Joh20-40) are charac-

teristic for the Zn-Pb zone (Fig. 4).

Grossular-enriched garnet (Table 3) is present as

inclusions in Fe-rich sphalerite from Paulus (Fig. 5l).

Even higher Gr components are present in garnet from

the uppermost level in Simon Iuda (3913), from

garnet with oscillatory zonation where some of the

zones still have And95, as in the Fe zone. Oscillatory

zonation and retrograde overprinting are seen in

several other samples (3375, 3786, 3786a) whose

average composition lies in the range And60-80.

4.4. Trace mineralogy

A range of exotic trace minerals are known from the

Cu-Fe core (Cook and Ciobanu, 2001), all of which are

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370334

associated with native Au. Whereas cobalt pentlandite,

carrolite, mawsonite and native Au and Ag are

exsolved from high-temperature bornitess, a range of

Bi- and Bi-Ag tellurides and selenides (kawazulite,

volynskite, hessite, etc.) are restricted to an association

with apatite or retrograde forsterite. In the latter, within

bands of Di90 skarn, at the margin of magnetite-

chalcopyrite ore, abundant native Au occurs as minute

grains in association with cervelleite (Ag4TeS) and a

range of Bi-minerals (Cook and Ciobanu, 2003a).

Adding to the list of rare Bi-Te-Se minerals associated

with emplacement of apatite, we now also report the

presence of telluronevskite, Bi3Te2Se. This Bi-Au-Ag-

Te-Se-(Co-Sn) signature seems to be unique to the Cu-

Fe core. To further stress this aspect, we also report

here the presence of blebs with intermediate composi-

tions in the galena-clausthalite solid solution series, as

exsolutions in chalcopyrite. Complementary to the

suite of exotic trace elements in the core, we also draw

attention to the presence of abundant REE phosphates

and uraninite, both intimately associated with apatite

emplacement.

Bismuth sulphosalts are particularly abundant in

the Fe zone and may coexist with associations of the

prograde stage. Their presence has been noted since

the late 19th century. Although phases recognised by

early workers (‘warthaite’ and ‘rezbanyite’; Krenner,

1925; Koch, 1930), were later discredited (Thompson,

1949; Zak et al., 1992), some 20 Bi-sulphosalts can

today be confirmed from the deposit (Cook et al.,

2002). Complex sulphosalt associations that include

galenobismuthite, cosalite, nuffieldite, members of

bismuthinite derivative series and the lillianite, pav-

onite and junoite homologous series are described

from occurrences in Simon Iuda, Magdalena, Ocna

Fig. 6. Macro- and hand-specimen-scale textures characteristic for the patt

contact, Terezia Quarry. Garnet-pyroxene skarn is in shades of grey. (b) B

band of bustamite, carrying sphalerite and galena, (dark patches) is seen in

magnetite and calcite. Note the decrease in the band interval from lower t

magnetite (black), serpentine (bottom), garnet (middle) and marble (whit

middle part of the photo (Gruescu collection). (e) Mossy texture realiz

colouration of the marble (Magnet Quarry; Li-2). (f) Mottled texture fo

magnetite crystals in a calcite matrix (Gruescu collection). (g) Nodular textu

garnet from the upper level of Reichenstein orebody. At the margin of the

texture of garnet (brown) and diopside (shades of green) within marble. Te

2) composed by a core of garnet-pyroxene skarn and an outer shell of g

Orbicular pattern formed by impregnation of fine galena (grey) and sphaler

upper levels (Orb-1). (k) Macroscopic oscillatory-zoned garnet seen as

Sculptured-faced garnet (Gruescu collection).

Turceasca and Paulus mines (Petrulian et al., 1977;

Ciobanu and Cook, 2000; Ciobanu et al., 2002b).

Most recently, members of cuprobismutite series and

related paderaite are also reported (Cook and Ciobanu,

2003b). Most typical of Ocna de Fier occurrences are

various morphological types of fine-intergrowths be-

tween Bi-sulphosalts, hosted either within magnetite

or hematite ore, or in the garnet-pyroxene skarn.

A number of other exotic trace minerals are asso-

ciated with retrograde overprinting in the Fe zone.

Among these, Bi-tellurides, hessite, matildite and gold

are also described from the Cu-Fe core. Gold seems to

be either associated with Bi-minerals or alone, as

minute grains of less than 10 Am diameter. We have

observed examples from Ocna Turceasca, Mijlociu,

Simon Iuda and Paulus. Maldonite, Au2Bi, has been

found only in the in Fe ores from Paulus (Ciobanu and

Cook, 2002; Ciobanu et al., 2003). To the list of

minerals associated with Au, we now add native

indium as < 5 Am inclusions sitting in retrograde

cracks within magnetite from Ocna Turceasca.

In this contribution, we report scheelite for the first

time from Ocna de Fier. The mineral is abundant,

although only microscopic in a number of occurrences

from Ocna Turceasca, Magdalena and Paulus. In each,

hematite is the first-formed Fe oxide and complex

replacement-overgrowth relationships with magnetite

are seen. Scheelite occurs as minute, dusty inclusions

within the Fe oxides, or along fine cracks extending

into the surrounding carbonate-silica matrix. Larger

grains, some Am in diameter, are also seen, especially

in Ocna Turceasca.

The manganese mineralogy at Ocna de Fier-Dog-

necea is complex and poorly constrained at present. In

the present suite of samples we note the presence of

erned skarn. (a) Scallop-shaped reaction front from the marble-skarn

anded skarn from the upper part of Simon Iuda orebody (Bnd-1). A

the centre, between garnet and marble (right). (c) Rhythmic banded

o upper part (Magnet Quarry; Li2a). (d) Rhythmic banding between

e, upper). Gaps and branching of magnetite bands are seen in the

ed by branching of magnetite (dark), surrounded by light brown

rmed by randomly-oriented, needle-shaped aggregates of aligned

re in macro-scale garnet, surrounding a fine-grained massive core of

core, rhythms of magnetite (black) can be seen (Nod-1). (h) Spotted

rezia Quarry (Nod-3). (i) Nodular texture from Terezia Quarry (Nod-

arnet skarn separated by a thin layer of magnetite (dark grey). (j)

ite (light brown) in silica/magnesia-enriched carbonate. Paulus Mine,

various dark and light shades of brown. Reichenstein Quarry. (l)

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 335

pyrophanite, MnTiO3, from Dognecea, in association

with titanite and Mn-bearing amphibole. Abundant

secondary Mn-associations (Ciobanu, 1999) are

formed by alteration of Zn-Pb ores and Mn-bearing

pyroxene. From an occurrence rich in Pb(Zn)-Mn

oxides and carbonates at the upper part of Elias, we

report zincsilite, Zn3Si4O10(OH)2�4(H2O). We point to

the fact that this mineral was first reported (Smol’ya-

ninova et al., 1960) from a comparable secondary

association, as pseudomorphs after diopside from

galena-sphalerite-chalcopyrite skarn at Batystau,

Kazakhstan.

C.L. Ciobanu, N.J. Cook / Ore Geolo336

5. Prograde textures

The Ocna de Fier-Dognecea skarn includes a

variety of macroscopic patterns, each consisting of

minerals found in prograde associations in the Fe

and Zn-Pb zones. Even though any of these minerals

can also be formed, in lesser amounts, during sub-

sequent, i.e., retrograde, stages, their affiliation to a

primary rather than subsequent stage of skarn devel-

opment can still be ascertained in the majority of

cases. As shown in the previous section, composition

alone is rarely sufficiently discriminative for this

purpose. More appropriate is the use of a combina-

tion of textural criteria, e.g., superposition, crystal

shape, zonation, textural relationships with other

minerals in the association.

Taking the above criteria into consideration, the

types of patterns introduced in this section are

considered to have formed during the prograde stage.

These are principally preserved in the outer parts of

the orebodies and at the marble-skarn contacts. The

latter are characterised by scalloped morphologies

(Fig. 6a), and by cm- to dm-scale, banded skarns

comprising distal-type assemblages: garnet-busta-

mite-marble (Fig. 6b), garnet-tremolite-marble and

garnet-hedenbergite-marble. Such banded assemb-

lages are ubiquitous in distal skarns at both Dogne-

cea and Paulus (Fig. 1). In the central part of the

orefield, marble-skarn contacts show a range of

patterns that include rhythmically banded (Fig.

6c,d), mossy (Fig. 6e), mottled (Fig. 6f), nodular

(Fig. 6g), spotted (Fig. 6h), and orbicular textures

(Fig. 6j).

Repetitive patterns also occur at the microscopic

scale, within monomineralic or two-component gar-

net-magnetite associations or individual crystals. Os-

cillatory zonation is widespread in prograde skarn

and may be seen even at the macroscopic scale in

garnet (Fig. 6k). Sculptured-faced garnet is excep-

tionally seen in hand specimen (Fig. 6l), but like

oscillatory zonation in garnet, it is not by itself

indicative of prograde formation, since both textures

also appear in minerals formed during the retrograde

stage. Although each microscopic-scale repetitive

pattern must be considered in its individual context,

they can nevertheless be considered as characteristic

among the broader range of patterns typical for

prograde skarn (Table 7).

5.1. Magnetite in rhythmically banded textures and

their morphological irregularities

Rhythmically banded textures, generically called

‘zebra rocks’, are known from diverse geological

environments and may consist of different mineral

assemblages (Krug et al., 1996). A variety of such

textures, involving magnetite and skarn/marble asso-

ciations, is recognised at Ocna de Fier. These were

described (‘Tiegererz’ for magnetite in a garnet ma-

trix) for the first time by von Cotta (1864) and were

later discussed comprehensively by Kissling (1967),

who interpreted them Liesegang phenomena (see

Section 8.1.1).

Sequences containing rhythmic magnetite bands

(usually in calcite) commonly mark the outer limit

of skarnification in the Fe zone. Exceptionally, rhyth-

mic bands of magnetite in marble can develop over

intervals of as much as 40 m, for example in the

occurrence in the southern wall of Magnet pit.

Within the rhythmic sequences, which commonly

have thickness in the order of cm to dm, the inter-band

distance ranges from mm to cm and correlates posi-

tively with the width of the individual bands. Various

trends in which grain size increases or decreases

across the patterned sequence are seen. In detail, the

magnetite banding has gaps or branching of the bands,

contains interlayered orbicules or small lenses, and

includes speckled and dendritic patterns. Band-bound-

aries can be sharp, ragged or diffuse. More complex

variants include rings with marginal mossy branches

(Fig. 6e), tiling and curving of bands with a tendency

to form 3D ring patterns, and slightly coloured inter-

band layers variously containing a colloidal substrate

in carbonate (Kissling, 1967), or serpentine, garnet or

diopside (Fig. 6d).

The overall complexity of magnetite banding is

illustrated by two samples (Fig. 7a to f): one from

Magnet quarry (Li-1; Fig. 1), and the other from the

Ocna Turceasca orebody (3101, Reichenstein III lev-

el; Fig. 1). Centimetre-scale banding is combined with

mm-scale rows of magnetite grains (Fig. 7a) and

constitutes two parallel rhythms. However, the finer

banding is turned into an oblique ‘alley’ (Fig. 7b) by

‘knotting’ across a band-set on the left of the sample

(Fig. 7a). Garnet clusters mark breaks in the bands and

are nucleated in the oblique alley (Fig. 7b). In a

parallel slice, cut 0.5 cm further into the sample,

gy Reviews 24 (2004) 315–370

Table 7

Summary of prograde textures in the Ocna de Fier-Dognecea deposit (Krug and Brandtstadter, 1999; Pring, 1989; Pring et al., 1999)

Mineral abbreviations: see Table 1.

Other abbreviations: Coll: colloidal-rhythms, RIF: reaction-infiltration feedback; OLC: Ostwald-Liesegang cycle; CPG: competitive particle growth; CGM: composite garnet magnetite crystals.1Example from Magnet quarry.2Example from Ocna Turceasca, Reichenstein III level.

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garnets are not visible. A second knot in Fig. 7b has

adjacent ‘ripples’ that dissipate to form ‘speckles’ in

the enclosing calcite. The thickest band (0.5 cm), two-

thirds of the way from left to right across the sample

in Fig. 7a, is embedded in magnetite ‘strips’ that

‘ripple’ in opposing directions. The parallel slices in

Fig. 7e and f show ring (2D-repetitive) patterns in

mm-sized magnetite bands that alternate regularly

with diopside bands. From bottom to top of the

sample, a ‘wiggle’ pattern gives way to several ring-

C.L. Ciobanu, N.J. Cook / Ore Geolo

centres that are overlain by a coarser wave rhythm.

Above this, the wave rhythms form an apparent triple

junction (Fig. 7e), but this is not seen in the second

slice (Fig. 7f).

Microscale investigation of the transition zone

between the banding and wiggle textures in Fig. 6e

and f shows a regular swinging pattern (oscillations)

between intervals of magnetite and diopside

(Di95Hed4Joh1) (Fig. 8a). As seen in Fig. 8b, coars-

ening of one mineral component (Mt) is seen against a

‘front’ of nucleation that is highly populated with

small-size grains in the other component (Di), at the

mutual boundary between the two minerals. This

boundary may be sharp but, in some cases, comprises

a narrow zone of small diopside crystals and intersti-

tial magnetite. Interfingering of the phases may also

occur (Fig. 8c).

5.2. Other macroscopic patterns

5.2.1. Skeletal-magnetite in mossy and mottled

textures

In addition to rhythmic banding, magnetite can be

a component of more complex types of patterns, such

as mossy (Fig. 6e) and the spectacular mottled texture

shown in Fig. 6f. In the latter, the texture consists of

laths (2 to 3 cm in length) of magnetite with lanceolate

edges randomly oriented in a marble matrix. In detail,

each lath has a structure characterised by skeletal

development of fine grains of magnetite from the

central part to the margin of the lath. Skeletal growth

of magnetite is also a characteristic of the mossy

branches (Fig. 6e). The fronts of dendrite consisting

of magnetite in marble, and described by Kissling

(1967), represent a variation of the mossy texture

characterised by skeletal growth of magnetite along

preferential directions.

Fig. 7. Photographs exhibiting characteristic prograde textures seen in hand

magnetite. (a) Morphological irregularities in a sliced sample from Magnet

magnetite (dark), calcite (light) with minor garnet (medium grey, arrowed).

Section 8.1.1). (b) Detail of part of slice shown in Fig. 5a, showing break

(arrowed). Garnet is nucleated in the knots. (c) Detail of part of slice show

(arrowed) with crystals pointing outwards to a median channel of calcite. (d

sense of ripple patterns. (e) and (f) Two parallel slices of a sample (3101;

rings and their morphological irregularities, representing the pre-nucleation

to ring centres. A triple joint of rings is seen in (e) (upper left, arrowed),

(arrowed).

5.2.2. Nodular, spotted and orbicular textures

Nodular textures, in which cores of massive garnet

are surrounded by rhythms of garnet and/or magnetite,

characterise the patterns in garnet-dominant assemb-

lages. Thus, in Fig. 6g, an 8-cm diameter nodule

consists of clusters of coarse garnet growing radially

outwards from a core of massive fine-grained garnet. A

2-cm wide zone, formed by rhythmic magnetite and

garnet, separates the fine-grained core from the outer

coarse-garnet. Another type of skarn nodule, lacking

the periodic rhythms, is described as ‘spotted’ textures.

Such nodules, several cm to dm in diameter, display

zonation (e.g., garnet core surrounded by pyroxene +

garnet rim, Fig. 6h). They are seen as spots within the

marble. A transitional type of texture between nodular

and orbicular is shown in Fig. 6i. Here, a thin shell of

magnetite separates the diopside core from the outer

garnet. Concentric, periodic shells of varied composi-

tions form orbicular textures. In cross section, they are

ellipsoidal rather than circular, unlike the nodules.

Sulphides in the Zn-Pb zone also form spotted and

orbicular textures. In Fig. 6j, sphalerite and galena is

associated with alternating layers of silica and magne-

sia-enriched carbonate. All these skarn-hosted patterns

also occur where massive skarn abuts lithological

contacts, such as massive magnetite ore, intrusive

rock, or crystalline schist, or enclaves of limestone

preserved in massive skarn (e.g., Terezia pit). The

nodular, spotted and, to some extent, the orbicular

textures are most abundantly seen in Terezia and

Reichenstein pits.

5.3. Repetitive patterns at the microscopic scale

5.3.1. Fronts of ‘crystals’

Fronts of ‘crystals’, each consisting of an alternation

of magnetite and garnet intervals, can be seen in the

gy Reviews 24 (2004) 315–370 339

specimen: morphological variation in precipitate banding involving

Quarry (sample Li-1, see text), showing rhythmic banding between

The sample is representative for the CPG/post-nucleation model (see

s in magnetite bands and an oblique alley formed by knotted bands

n in Fig. 5a, showing paired rows of magnetite in the finer rhythms

) Detail from a further slice parallel to Fig. 5a, showing variation in

Ocna Turceasca Mine, Reichenstein III level), showing precipitate

model (see Section 8.1.1). Awiggle pattern (dashed line) gives way

but not in (f), where a repeat of small ring centres is seen instead

Fig. 8. Back-scattered electron images showing prograde textures. (a) Transition between banding and wiggle textures (3101; Figs. 7a to c),

showing regular swinging between magnetite (light) and diopside (dark). (b) and (c) Illustrate details from (a). (3101). Coarsening of magnetite

(light) towards the boundaries with diopside (dark). In contrast, numerous small crystals of diopside cluster at the boundary. Skeletal magnetite

can be seen between the diopside grains in (b). (d) Deformed composite ‘crystals’ (CGM; e17), consisting of rhythmically intergrown garnet

(dark grey) and magnetite (light grey). Garnet shows brittle cracks, but magnetite, which behaved in a more ductile fashion, is contorted and

develops marginal hooks. (e) Front of magnetite crystals showing oscillatory zoning expressed by alternating bands of dark (Si-Mt) and light

(Si-free) magnetite (PP). (f) A single magnetite crystal with a Si-Mt core, showing oscillatory zoning comparable with that shown in (e) (PP).

Zoning is centred on a grain of galenobismutite (white), surrounded by fine-grained silicate inclusions. (g) Oscillatory-zoned crystal of

magnetite. The grain has a SiO2-free core, surrounded by bands of Si-Mt and Si-free Mt. Note that the zonation is not as clear as in (f), because

of the superimposed retrograde overprinting. (h) Oscillatory zoning patterns in andradite typical of massive garnet skarn from the Fe zone (GR).

(i) Basal section through a prism of pyroxmangite showing slight zonation with Fe-rich rims. In direct contact, at the right-hand corner is

pyroxene typical for distal Zn-Pb skarn in Dognecea (Dogn).

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370340

transition zone between massive magnetite and garnet

skarn. The close partnership of the two minerals in

prograde skarn is further underlined by the fact that

each individual ‘crystal’ is in fact composite, formed by

rhythmic zones of garnet and magnetite. Such hybrids

(CGM; Fig. 8d) closely resemble the type of ‘crystals’

formed by intergrowths between minerals that form

polysomatic series, e.g., biopyriboles (Veblen and

Buseck, 1979). Garnet within the CGM bands lacks

oscillatory zonation and has a limited composition

range (And96; Table 3), comparable with garnet in

massive magnetite ore. Instead, it is the magnetite

within the bands that displays oscillatory zonation.

This is marked by thin, intermittent intervals of Si-

Table 8Summary of retrograde textures in the Ocna de Fier-Dognecea deposit

Mineral abbreviations: see Table 1.

CGM: composite garnet-magnetite crystal fronts.

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bearing magnetite (typically 1 to 3 wt.% SiO2). The

CGM are subsequently deformed (Fig. 8d), in that

garnet is fractured perpendicular to the layers and

magnetite has accommodated this by limited ductile

flow. Such mutual adjustment during deformation is

evidence for the primary co-existence of magnetite and

garnet during ‘crystal’ growth. However, this would

imply crystal-chemical links between the two minerals

that are, as yet, unrecognised.

A second type of ‘front’ consists of crystals of

monomineralic magnetite (Fig. 8e). As in the previous

case, the front displays oscillatory zonation, in this

case expressed by alternating bands of Si-free and Si-

bearing bands.

5.3.2. Oscillatory crystal zoning and lamellar inter-

growths

Oscillatory zonation in magnetite is frequent, es-

pecially in the transition zone from massive ore to

garnet-pyroxene skarn. A core of silica-bearing mag-

netite, and a contrasting example of a core of Si-free

magnetite surrounded by an outer Si-bearing zone

with oscillatory zonation are shown in Fig. 8f and g,

respectively. Similar oscillatory zonation in magnetite

has been reported from skarn deposits in Japan (e.g.,

Shimazaki, 1998). The zoning is, however, further

complicated by retrograde overprinting, e.g., the

swarms of silicate inclusions and sulphides introduced

at the intersection of radial cracks (Fig. 8f). Even

more complicated textures involving Si-bearing mag-

Fig. 9. Back-scattered electron images (a to f) and transmitted light photom

retrograde stages in Cu-Fe core (a to e; all from sample 164) and the Fe

sphalerite. Note that the cluster cuts boundaries between Sp and Cp (black

are indicated by white arrows. In the dark areas, forsterite and diopside are

grain in the cluster is shown. Note the corrosion on the margins of forsterit

apatite was precipitated in a ‘colloidal’ state. In (c), pyroxene enclosed in

corrodes deeply into the diopside. (d) A needle-like grain of orthopyroxene

Note the pressure trails at the edges of this needle (arrowed). (e) Bleb of ura

cluster. Serpentinisation of diopside appears contemporaneous with uranin

the serpentine (arrowed). (f) Cluster of calcite with quartz emplaced in

(3089a). Associated with such clusters are tiny but numerous grains of go

secondary Bi-minerals (Bialt) within magnetite (Mt). Maldonite is associa

characteristic retrograde trace mineral assemblage in the Fe-zone (SS2). (h

scheelite (5598). (i) Coarser grains of scheelite formed during retrograde re

and chaotic birefringence (white arrow) that overprint an earlier, pre-exis

crossed polars; 5599). (k) Branching filaments consisting of quartz and

Photomicrograph under crossed nicols (3089). (l) ‘Splashed’ clusters of qu

a garnet (photomicrograph under half-crossed nicols (3087).

netite occur in hematite that has been pseudomorphed

by magnetite (see below).

All skarn silicates (i.e., garnet, pyroxene, pyrox-

enoid) at Ocna de Fier-Dognecea meet the require-

ments for oscillatory zonation, i.e., they represent

intermediate members in solid solution series and

each of them displays a broad field of composi-

tional (see above). However, it is only garnet that

shows oscillatory zonation with variable composi-

tional ranges (Fig. 8h). Even though both pyroxene

and pyroxenoid show significant compositional var-

iation (Tables 4 and 6), a tendency towards zona-

tion is only rarely developed, e.g., marginal rims of

pyroxmangite slightly richer in Fe (Fig. 8i).

In pyroxene, the intermediate compositions seen

across lamellar intergrowths in samples from Paulus

(Fig. 5j) and comparable compositions seen across

homogenous, undifferentiated lamellae in samples

from Dognecea (Fig. 8i) could actually be poly-

somatic sequences of intergrowths, i.e., lattice-scale

intergrowths of end-member unit cells. Only trans-

mission electron microscopic (TEM) investigations,

however, could verify this hypothesis and properly

document this type of material. Comparable, proven

sequences in biopyriboles (Veblen and Buseck,

1979) are interpreted as caused by metasomatism.

The relationship between the persistently present

pyroxenoids, themselves showing variable composi-

tion, suggests that they could potentially also be

considered in terms of polysomatism.

icrographs (g to i), showing piercing clusters emplaced during initial

zone (f to l). (a) Apatite cluster in chalcopyrite, with exsolution of

arrow). Tiny blebs of minerals from the Bi-Te-Se-Au-Ag association

enclosed in the apatite (b) and (c) details from (a). In (b), a forsterite

e, as well as the apatite trails across forsterite. The textures show that

the cluster has a marginal rim of serpentine (dark). Note that apatite

(?) protruding into the sulphide matrix from the rim of turneaureite.

ninite at the margin of diopside, precipitated together with the apatite

ite formation. Note the sigmoid termination of the uraninite towards

magnetite. Note the similarity in appearance to the apatite clusters

ld. (g) Trail of blebs of maldonite (Mld) with marginal alteration of

ted with gold, native-Bi, Bi-tellurides and bismuthinite, forming a

) Garnet showing shock-induced brecciation and dusty inclusions of

placement of hematite by magnetite (e10). (j) Development of scales

ting sector zoning and zonal birefringence (photomicrograph under

calcite (light) piercing at the grain boundaries of garnet (dark).

artz-calcite with undulatory extinction (arrowed) piercing the core of

eology Reviews 24 (2004) 315–370 343

5.4. Micron-to-nanoscale intergrowths of bi-sulpho-

salts from polysomatic series

Although not vital to the present discussion, we

would emphasize that patterning in the Ocna de Fier-

Dognecea ore can also be observed on the nanoscale.

Transmission electron microscopic investigation of

cuprobismutite, Cu8.07(Ag0.99Pb0.2Bi12.72)13.91S24,

and paderaite, Cu7.11(Ag0.36Pb1.20)1.56Bi11.28S22.05,

C.L. Ciobanu, N.J. Cook / Ore G

from Paulus has identified microscopic-scale inter-

growths, with inter-layering between the two miner-

als, that extends down to the lattice scale (Cook and

Ciobanu, 2003a,b; Ciobanu et al., 2004, in press). A

series of long- and short-range polysomes are

documented across compositional fields intermediate

between the two minerals. Such polysomes result

from periodic stacking of unit cells of the two

minerals.

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370344

6. Retrograde textures

All textures that indicate superposition of a new

mineral assemblage upon a former one, or modifica-

tion of the pre-existing assemblage, are considered

retrograde (Table 8). Unlike the prograde ones, the

majority of the retrograde textures we present here are

recognisable only at the microscopic level. We distin-

guish several suites of retrograde textures that show

progressive destruction of the prograde assemblages,

i.e., piercing, shock-induced, fluid-pressure assisted

brecciation and deformation textures, and also a suite

of textures that have a healing character, themselves

superimposed on the previously disrupted and

reworked assemblages.

Retrograde textures are particularly marked within

irregular envelopes, varying from dm to m in width, at

the margin of each orebody. Garnet and magnetite in

the prograde assemblages from the Fe zone preserve

evidence of the full range of retrograde textures,

whereas such textures are poorly recorded by near-

massive magnetite in the centre of the orebodies. In

the Zn-Pb zone, we recognise a dominant brecciation,

with few other retrograde textures preserved. The Cu-

Fe core in Simon Iuda displays a distinct range of

minerals in its overprinting textures that have impli-

cations for the reconstruction of skarn development

across the entire orefield, covered in Section 9.

6.1. ‘Piercing’ clusters1

In the Cu-Fe core we observe that magnetite-chal-

copyrite-forsterite-diopside assemblages are invaded

and pierced by clusters of apatite, each with diameters

up to 1-2 mm (Fig. 9a). Fine-grained, second genera-

tion forsterite is seen in comparably shaped clusters

that pierce through magnetite-chalcopyrite assemb-

lages (sample 73). These clusters crosscut the bound-

aries between pre-existing minerals in the association

(e.g., between Fo and Mt in Fig. 5a or between

exsolved Sp and Cp in Fig. 9a,d). The clusters are

characterised by fine intergrowths of minerals, both

pre-existing and newly introduced, and have a ‘sym-

plectitic’ appearance. Patches, not exceeding a few Amin diameter, and consisting of Bi-Te-Se phases, with

1 ‘Piercing’ clusters are symplectite-like and have been super-

imposed onto a pre-existing assemblage through fluid action.

Au (see Section 4.4) are seen within the apatite

clusters. Trails of apatite blebs crosscut the forsterite

(Fig. 9b) and deeply corrode the diopside margins

(Fig. 9c); rims of serpentine are formed at the bound-

ary between apatite and the two silicates. Needles of

orthopyroxene protrude into the sulphide from the

apatite (Fig. 9d). Among the most suggestive indica-

tions that emplacement of these clusters was ‘dynamic’

is the presence of pressure-trails at the margins of the

orthopyroxene. The apatite clusters include abundant

uraninite. The uraninite patches commonly have sim-

ilar sigmoid terminations towards serpentine (Fig. 9e).

Comparable piercing clusters are seen in the Fe

zone. Here, however, they consist of calcite ( + quartz)

and are emplaced in magnetite and garnet (Fig. 9f).

Minute Au grains are associated with the emplacement

of such clusters, in for example, the Elias orebody. In

Paulus, trails of tiny grains of maldonite, and subordi-

nate Bi-Te phases are seen in magnetite (Fig. 9g).

Shock-induced brecciation in garnet (Fig. 9h) is ac-

companied by nucleation of dusty scheelite in distal

skarn (Paulus, 5598). The coarsest scheelite grains

(Fig. 9i) and abundant fine-grained (f 5 Am) gold

are noted along cracks and within silica-filled voids in

those associations that have a complex history of

replacement-overgrowth between hematite and mag-

netite (e.g., at the upper part of Ocna Turceasca ore-

body). In garnet displaying anomalous birefringence

with prograde sector zoning, the appearance of scales

and chaotic birefringence (Fig. 9j) is followed by

emplacement of filaments/patches of calcite and quartz

(Fig. 9k,l). The optical orientations of individual bodies

within the calcite-quartz clusters (Fig. 9l) are not

identical, but can nevertheless be correlated with one

another, in that a progressive extinction across a cluster

is seen with rotation. This, as well as the pressure trails

shown for apatite in the Cu-Fe core, is evidence that

emplacement of these clusters was pierced into the pre-

existing assemblage in a colloidal state. Crystallisation

of the minerals as we see now in the clusters followed

the piercing moment.

6.2. Shock-induced textures

The outer limit of retrograde overprinting in the Fe

zone is marked by magnetite pseudomorphs of lamel-

lar hematite (Fig. 10a). Where bordered by the pseu-

domorphosed hematite, andradite displays overgrowth

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 345

halos and brecciation in the form of tiling2, attribut-

able to shock induction from high fluid pressure (Fig.

10b). This is further underlined by the presence of

pressure-induced corrosion in the band with grossular-

enrichment closest to the margin. The overgrowths

actually consist of fine intergrowths between garnet

and calcite, with domains of preferential orientation

(Fig. 10c). Therefore they represent a kind of pres-

sure-induced halo that is formed in the pressure-

shadow of garnet towards hematite.

Ongoing retrogression may result in magnetite

partially reverting to hematite and the original laths

being rimmed by either small crystals of magnetite

(Fig. 10d) or microscopic rhythms–overgrowths of

hematite and magnetite (Fig. 10e,f). These textures

indicate at least two cycles of stability inversion

between hematite and magnetite. The magnetite rim

at the margin of the hematite overgrowth in Fig. 10e

shows preferential orientation during growth, similar

to intergrowths of garnet and calcite. Replacement of

former hematite in the laths is here characterised by

the appearance of complex domains with oscillatory

zonation between Mt and Si-Mt, themselves inter-

meshed with domains of Mt and relict Hem (Fig. 10f).

The complexity of the relationships between the Fe

oxides indicates that development of replacement-

overgrowths processes was impacted by the same

pressure-induced shocks that affected the garnet.

Similar to garnet, we see also pressure-induced

halos enclosing relict diopside (Fig. 10g). Instead of

tiling, the corresponding shock-induced texture in

diopside is the appearance of shard-like domains with

slight compositional differences (Fig. 10h). Pressure

corrosion affects Si-Mt cores in oscillatory-zoned

aggregates of Mt (Fig. 10i). In this example, the halo

of small blebs surrounding a central sulphide mass is

an internal arrangement typical for impact textures.

The piercing event is contemporaneous with the

pressure corrosion.

Pressure-induced (shock) textures are seen opti-

mally in garnet, even though all components of the Fe

zone record comparable textures. Tiling is seen as a

penetrative front inducing chemical modifications

2 ‘Tiling’ describes the development of a more or less regular

internal chemical rearrangement, resulting in the tile-like appear-

ance of a pre-existing garnet. The texture is interpreted as a

consequence of pressure-induced displacement.

within pre-existing garnet (Fig. 10j). Thin bands of

more Gr-rich garnet are formed in response to the

pressure-induced shocks. We note that modifications

in the periodicity of these bands, as well as the

appearance of stepwise dislocations, are prompted

by dilational cracks. Highly indicative of high-pres-

sure fluxes during crystal growth (Brenan, 1991) is

the appearance of sculptured-faced garnet (Fig. 6l) or

flattened crystal corners (Fig. 10k). Tiling of garnet

marks the latter. Pressure corrosion, resulting in the

appearance of conical compositional shadows (indi-

cated on Fig. 10k), is further superimposed onto the

tiling. The tiling of garnet can be developed as

rippling fronts when superimposed onto prograde

oscillatory zonation (Fig. 10l). The reshaping of

previous zoning in the form of ripples strongly sug-

gests that the tiling was controlled by oscillatory

pressure in the fluid. The shock-induced textures

represent a suite of patterns superimposed onto host

assemblages, without addition of new material. Even

though piercing and shock-induced textures can be

collectively considered as incipient brecciation, they

characteristically lack any direct connection to sys-

tems of dilational cracks or fractures. Instead, shock-

induced patterning is indicated by slight composition-

al modifications or reshaping of previous patterns

within the pre-existing assemblage.

6.3. Brecciation

In the Cu-Fe core, the ductile matrix provided by

Cu-Fe sulphides, assists disruptive brecciation of

refractory minerals such as forsterite (Fig. 11a) or

magnetite (Fig. 11b). Serpentine is formed in situ

from forsterite and carbonate replaces the fractured

grains of magnetite. In the Fe zone, similar disruptive

brecciation occurs in the ‘blown-apart’ fabrics that

characterise the garnet assemblages (Fig. 11c). Here,

the garnet fragments are separated by quartz-carbonate

infill, as well as a series of hydrated minerals, e.g.,

chlorite. Brittle brecciation in oscillatory-zoned garnet

(Fig. 11d) follows shock-induced reshaping, as shown

in the displacements along the fracture in Fig. 11e.

Zones of Gr-rich garnet corrode deeply into former

andradite (Fig. 11f). Similar complex overgrowths and

corrosion-absorption boundaries develop within py-

roxene from the Fe zone (Fig. 11g), or in garnet in the

Zn-Pb zone (Fig. 11h). Such textures are reported in

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370346

metasomatic fronts from metamorphic terranes, where

they are considered indicative for changes in perme-

ability–porosity at the skarn front in response to

infiltration-driven decarbonation (Yardley and Lloyd,

1995).

In the Zn-Pb zone, other features that emphasise

the importance of fluid pressure fluxes are seen in

‘tails’3 (Fig. 11i) and ‘hooks’4 (Fig. 11j). Symplectite

selvages, consisting of magnetite and ankerite occur at

the margins of galena (Fig. 11j–l) in distal skarn from

Paulus. The internal fabric of the symplectites and the

development of the marginal ‘hook’-like structures

suggests that they formed under oriented fluid pres-

sure. Local remobilisation is invoked to explain the

irregular patches of galena that cluster outside the

symplectite-rimmed larger galena bodies (Fig. 11l).

We suggest that cementation is facilitated by ongoing

remobilisation of matrix material.

Evidence of fluidisation, cementation and rotational

movement also exists in garnet-magnetite assemblages

(Fig. 11h). In all zones, rounded garnet and magnetite

is ascribed to the abrasion and milling of brecciated

magnetite or garnet in the presence of more ductile

minerals and fluid fluxes. Such behaviour is directly

analogous to that of rheologically distinct minerals in

metamorphosed and remobilised sulphide ores (e.g.,

Gilligan and Marshall, 1987; Marshall et al., 2000;

Vokes, 2000).

6.4. Other fluid pressure assisted deformation

textures

A range of textures testifies to further retrograde

overprinting of the fractured assemblages following

brecciation. These include sets of en echelon fractures

(Fig. 12a), crossed stress-jointing (Fig. 12b), over-

growths (Fig. 12c), and ‘jigsaw’-shaped margins

(Figs. 11i and 12d), marginal hook distortion (Fig.

12e), pressure shadows (Fig. 12f) and competence

effects (Fig. 12f). In relation to the latter, the magne-

3 We use ‘pressure tail’ as a descriptive term for oriented

overgrowths on grains that display rounding or deformation. The

orientation of the tail is concordant to the sense of deformation

produced by pressure-fluxes, in the same way that the term pressure

shadow is defined in structural petrology.4 ‘Hook’ is used to describe a fine protrusion from a mineral

grain, caused by plastic deformation.

tite bands are contorted and have clearly behaved in a

more ductile fashion than the garnet (Fig. 12g).

Highly relevant for sustained brecciation-deforma-

tion during the retrograde stage is the assemblage of

coarse-grained garnet, pyroxene, magnetite and py-

rite from Ocna Turceasca (sample 798, Fig. 12d to

f,h,i). Thus, a large grain of pyrite (f 1 cm in

diameter) has undergone brecciation such that clasts

are dispersed within adjacent garnet. A ‘jigsaw’

shaped border formed simultaneously with recrystal-

lisation of the magnetite and is evidenced by the

development of regular inliers of garnet within pyrite

(Fig. 12d). Garnet deformation is assisted by devel-

opment of micro-shear fractures with pressure shad-

ow domains between garnet and coarser pyrite, as

well as by contemporaneous re-crystallisation of

magnetite (Fig. 12f). Both equilibrium boundaries

and diffusion-reaction boundaries between garnet

and magnetite are seen as fine intergrowths (Fig.

12h). Distinct bands within the garnet zoning enclose

sets of pyrite clasts that interrupt a second set of

fractures (Fig. 12i).

6.5. Recrystallisation and annealing fabrics, other

welding and sealing textures

In the Cu-Fe core, brecciation and rounding of

magnetite is followed by cementation of the disrupted

grains by carbonate (Fig. 13a). Recrystallisation of

magnetite is indicated by fresh sets of cleavages

crosscutting inclusion-rich, prograde cores (Fig.

13b,c). Formation of typical minerals for the Cu-Fe

core, such as valleriite, is seen within earlier sets of

cleavages in magnetite (Fig. 13d). This indicates that,

during abrasion of magnetite, the fluids induced

remobilisation of matrix chalcopyrite. Reaction be-

tween chalcopyrite and silicates in the magnetite

produced valleriite. Sealing of fractures within mag-

netite is evidenced by arrays of dusty silicate inclu-

sions (Fig. 13e), whereas 120j triple junctions (Fig.

13f) indicate annealing (seen only in massive magne-

tite from the Cu-Fe core).

In the Fe zone, previously deformed, fractured,

rolled and abraded fragments or aggregates of refrac-

tory magnetite are further welded and overprinted by a

complex sequence of sealing and re-cementation and/

or overgrowth cycles (Fig. 13g,h). Further fracturing

occurs following the welding episode. Thin exten-

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 347

sional cracks follow recrystallisation in garnet (Fig.

13i) due to mechanical equilibration of the reworked

ore mass.

7. Skarn formation

On the basis of mineralogical evidence (Cook and

Ciobanu, 2001), temperatures for prograde proximal

skarn in the Cu-Fe core reached 600 to 700 jC. We

consider the lower temperature limit for formation of

the Cu-Fe core to be in the interval 500–600 jC,based on the presence of exsolved spinel in both

forsterite and magnetite. This is the temperature range

for immiscibility gaps in both the (Fe,Mg)Al2O4-

(Fe,Mg)Fe2O4, and the FeAl2O4-Fe3O4 series calcu-

lated for fixed Fe-Mg exchange potential of Fo95-90 on

isothermal sections through the spinel prism (Sack

and Ghioso, 1991). These temperatures are reasonably

close to those obtained for hornfels in the contact

aureole (700–750 jC; Nicolescu and Cornell, 1999).

We therefore assume a minimum temperature interval

of 550–650 jC for the Cu-Fe core. For the Fe zone,

an upper temperature limit slightly in excess of 600

jC can be assumed, based on the application of the

ludwigite geothermometer by Marincea (1999).

In our reconstruction of skarn formation in T-fO2

space (Fig. 14), conditions for the Cu-Fe core are

based on the stability of Cu-Fe sulphides with mag-

netite at a log fS2 value of � 3.25 (with the Mt-Po-Py

invariant point fixed at 550 jC). The upper fO2 limit

is considered to lie close to the Mt/Hem buffer, based

on the presence of forsterite in association with Di>90pyroxene, the corresponding stable pyroxene at this

buffer. Furthermore, we note that Di90 is also the

stable pyroxene in association with magnetite in the

associations that lack garnet from the proximal Fe

zone (e.g., Petru and Pavel).

The appearance of garnet and disappearance of

forsterite marks the boundary between the Cu-Fe core

and the Fe zone. Andradite stability in fO2-fS2 space

is temperature dependent: below 600 jC andradite is

stable with Py, whereas from 500 to 300 jC high-And

garnet is stable with both Py and Mt. Below 300 jC,andradite is stable with both Py and Hem (Gamble,

1982). Therefore, proximal skarns with And>90 gar-

net, associated with Mt or Hem, but without Py (e.g.,

in Magdalena), are consistent with formation at sim-

ilar temperature-fS2 conditions as in the Cu-Fe core,

but instead straddle the Mt-Hem buffer. Only when

co-existing, do garnet and pyroxene show significant

compositional ranges, i.e., Di70-90Hed10-30Joh< 10;

And95-75 (Section 4).

The boundary between the Fe and Zn-Pb zones is

recognised in terms of compositional changes in py-

roxene, accompanied by a dominance of sulphides in

the Zn-Pb zone, i.e., sphalerite, galena and pyrite are

more abundant than Fe oxides. Pyroxene instead of

garnet, along with associated pyroxmangite and/or

bustamite, is the main skarn component. Garnet does

not disappear altogether, but becomes minor. In all

associations, hematite is the stable Fe oxide. Overall,

the pyroxene composition in the Zn-Pb zone, i.e.,

Di10-40Hed20-70Joh20-40, is highly variable when com-

pared to pyroxene from the Fe zone (i.e., Di>70Hed< 30).

It has a conspicuous, Di-depleted, Hed- and Joh-

enriched character, with variation of Di:Hed:Joh ratios

within individual orebodies (e.g., Simon Iuda), and

from proximal to distal settings, i.e., Di20-40Hed20-40Joh40 in proximal, Di< 10Hed60-70Joh20-30 in distal.

The temperature ranges estimated for prograde

assemblages in the Fe and Zn-Pb zone in Fig. 14

are derived, using the data of Gamble (1982), from the

Di components in pyroxene and, especially, pyroxene

coexisting with sphalerite (see below).

In proximal skarn, the stabilisation of diopside in

retrograde assemblages from the Cu-Fe core requires

temperatures in excess of 470 jC, based on the

stability of pyroxene in the presence of epidote

(Einaudi et al., 1981). Pseudomorphous replacement

of diopside by tremolite at the outer margins of

proximal skarn, i.e., at Terezia, Iuliana and Ocna

Turceasca, implies temperatures above 420 jC. At 2kbar, this is the lowest temperature at which tremo-

lite is stable (Einaudi et al., 1981). Comparable

temperatures for the retrograde stage are indicated

also by presence of maldonite blebs (stable up to 371

jC; Schunk, 1969) seen in trails of Au-Bi-Te min-

erals within magnetite (Fig. 9g) from distal skarn in

Paulus.

7.1. Solid-solution effects in skarn assemblages

The reaction in the Fe zone most pertinent to our

assemblages is: 9Hed + 2O2 = 3And +Mt + 9Qz (Gus-

tafson, 1974). However, this reaction considers only

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370348

ideal components; in skarns both pyroxene and garnet

are generally intermediate members of the two solid

solution series, diopside-hedenbergite-johannsenite

and andradite-grossular. This affects the fO2 value

for the reaction on the fO2-T diagram. The solid

solution effect can nevertheless be calculated using

the formula Dlog fO2ss = 3/2log aAnd� 9/2log aHed

(Bowman, 1998).

In order to constrain the lower fO2-T limits for the

Fe zone, we consider the compositions of skarn asso-

ciations from proximal marginal orebodies such as

Mijlociu and Ocna Turceasca, and from distal ore-

bodies (Paulus), calculating Dlog fO2 in each case.

Because of chemical modifications induced by retro-

grade overprinting (Sections 4 and 6), we have criti-

cally evaluated the mean compositions in Tables 3 and

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 349

4, accordingly eliminated a number of individual

analyses, and then recalculated the means on a more

restricted compositional range, thus giving more plau-

sibleDlog fO2 values. The highest values (Dlog fO2>4)

correspond to associations in which Py is stable to-

gether with Mt at the margins of proximal skarn (Ocna

Turceasca, Mijlociu). In the case of Mijlociu, a profile

from the massive magnetite core to outer skarn, gives

Dlog fO2 values from 4.5 to 2.5. From distal skarn in

Paulus (deepest level), we obtain an intermediate Dlog

fO2 value of 3.3 for an association in which Hem is the

primary Fe oxide (now retrogressed toMt, without Py).

Gamble (1982) showed that at 2 kbar, Di75Hed25,

coexisting with Py, is stable only below 550 jC. If weignore the Joh component (Joh< 10), this can be

considered the representative composition for garnet-

pyroxene assemblages in the Fe zone. On the other

hand, garnet coexists only with Mt or with Py (below

500 jC; see above). Considering the above, fO2-T

conditions for the garnet-pyroxene associations in the

Fe zone (Fig. 14) can be well constrained at temper-

atures around 500F 50 jC, at fO2 values that range

from close to the Mt–Hem buffer, extending across

the Py/Mt buffer. Within this temperature interval,

however, individual orebodies show different condi-

tions. The association from Paulus can be placed

towards the maximum temperature, at conditions

closest to the Mt/Hem buffer, whereas assemblages

from Ocna Turceasca are lower temperature, posi-

tioned close to the Mt/Py buffer. The aforementioned

Fig. 10. Back-scattered electron images (except (d), reflected light photom

zone. (a) Coarse garnet (And96) with marginal magnetite that has replace

corroded and has a thin overgrowth halo (e14). (b) Detail of garnet margin

zonation. The dark band is Gr-enriched (And53Gr44Sps3) whereas the lig

corrosion. The marginal overgrowth halo consists of garnet with a similar c

showing oriented symplectite-like intergrowths between garnet (And93G

magnetite (dark), followed by inversion to hematite (patches of relict magn

magnetite at the margin of the laths in contact with the matrix carbonate

inversion between hematite and magnetite (e14). (e) Overgrowth of hemat

has already been replaced by magnetite (lower right). Note orientation o

relationships between magnetite, Si-Mt and hematite within laths of former

preferential orientation are also seen (e1). (g) Relict pyroxene in calcite i

pyroxene and calcite (e17). (h) Chaotic shock-induced deformation in d

composition (798). (i) Clusters of piercing galena and silicates (black) wit

Also note pressure corrosion boundaries (white arrow) (PP). (j) Tiling in g

enriched composition (And62Gr35Sps3). Note the stepwise displacement p

798. (k) Crystal with shock-induced flattened-edges shown by developm

compositional shadows (arrowed). (l) Detail showing fine ripples reshapin

under oscillatory pressure in the fluids.

profile across Mijlociu shows a change in conditions

from the inner core to the margin of orebody, crossing

the Mt/Py buffer.

7.2. Compositional fields in pyroxene and their

significance

The pyroxenes in each association from the Zn-Pb

zone show characteristically extensive compositional

fields (Fig. 4). Mean compositions (Table 4), rather

than individual analyses, are nevertheless instructive

for the reconstruction of genetic conditions for the

different associations. This implies that an initial

equilibrium is to be assumed for the pyroxene in each

association. The compositional ranges among individ-

ual pyroxene populations in proximal skarn can be

explained in terms of the products of eutectic decom-

position (Fig. 5h), even in cases displaying retrograde

overprinting (Fig. 5i).

In proximal skarn, the two pyroxenoids, bustamite

and pyroxmangite, appear to have opposite influences

on pyroxene chemistry; the former tends to occur with

Di-rich pyroxene (Fig. 5h), whereas the latter favours

Hed-rich pyroxene (Fig. 5i). Whether the coexisting

pyroxenoid can directly influence pyroxene stability

or these observations are a function of equilibrium

fractionation can only be speculated upon.

In distal skarn, however, the presence of an array of

different pyroxene compositions within single, ho-

mogenous lamellae (Fig. 4) may represent extensive

icrograph), illustrating retrograde shock-induced textures in the Fe

d hematite as seen by the lamellar shapes. The magnetite border is

from (a), with incipient tiling seen as shock-induced compositional

ht zone is andradite-rich (And93Gr4Sps3) and also shows pressure

omposition to the And-rich garnet (e14). (c) Detail of the halo in (a),

r3Sps4) and calcite (e14). (d) Replacement of hematite (light) by

etite are arrowed in black). Also shown is a stepwise overgrowth of

(white arrow). The texture indicates at least two cycles of stability

ite with magnetite at the margin at the border of a hematite lath that

f magnetite at the edge of hematite (e1). (f) Complex replacement

hematite. Overgrowths of hematite with marginal magnetite showing

nliers within magnetite. Note the marginal dusty halo composed of

iopside, indicated by shard-like domains with slight variations in

hin cores of Si-Mt. The pre-existing zoning is incipiently distorted.

arnet shown as development of thicker and thinner bands with Gr-

rompted by the dilational crack in the centre (e14). (k to l) Sample

ent of tiling. Superimposed pressure corrosion is seen as conical

g oscillatory zoning in andradite. This is interpreted again as tiling

C.L. Ciobanu, N.J. Cook / Ore Geolo350

local-scale non-equilibrium. However, we propose

that these compositions, all slightly differing from

one another, could have been stabilised under the

same formational conditions, and actually be poly-

somatic sequences of intergrowths. A critical role for

polysomatism in stabilising the pyroxene in distal

skarn, i.e., within the limits Di< 10Hed60-70Joh20-30,

is yet to be documented or confirmed, but could

readily explain some of the observations.

7.3. Sulphidation-oxidation effects on skarn

assemblages

End member hedenbergite is stable with pyrite

only below 288F 16 jC (Burton et al., 1982). Exper-

imental studies of the fO2-fS2-T stability fields for

hedenbergite-johannsenite (Burton et al., 1982) and

hedenbergite-diopside (Gamble, 1982) solid solutions

at 2 kbar show that 15% Joh component, or 50% Di

gy Reviews 24 (2004) 315–370

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 351

component, increases the stability of Hed, coexisting

with pyrite, to 537 and 395 jC, respectively. Eventhough these experimental data show that the influ-

ence of Joh is significantly greater than that of Di,

neither study considered the presence of all three

components in the pyroxene.

More appropriate for the purpose of our investiga-

tion is the mol% FeS content of sphalerite (FeSSp), that

acts as a buffer to monitor the sulphidation state in the

same assemblage that Gamble (1982) used to obtain the

calibration curves for pyroxene under sulphidation

reaction. From a total of eleven sphalerite-bearing

assemblages (Table 5), seven coexist with pyroxene

unaffected by later stages of retrograde alteration (i.e.,

replacement by carbonate, quartz or other secondary

compounds). Copper contents of sphalerite are very

low, typically < 1wt.%. We therefore use FeSSp to

calculate the sulphidation state of associations in which

pyroxene coexists with sphalerite, following Gamble’s

equation: log f S2 = 1.3596(103/T)2� 18.5134(103/

T) + 16.9806� 2log aFeS(Sp), at estimated temperatures

for each case.

Temperature estimates for the proximal skarn are

based on the above considerations regarding fO2-fS2conditions at a minimum temperature for the Cu-Fe

core (550 jC), with the Mt-Hem-Py triple point fixed

at a log fS2 value of � 1.25. As presented above, in

the Fe zone, assemblages consisting of Mt-Py and

Hem-Py follow simple Mt-bearing assemblages. This

implies an increase in f S2, followed by an accompa-

nying slight increase in fO2. Indeed, in the associa-

tions from the Zn-Pb zone, pyroxene is stable with

Py-Hem, and therefore we should obtain according

Fig. 11. Photomicrographs in reflected light (b), in transmitted light (c

brecciation. (a) ‘Blown apart’ forsterite (light grey) replaced by serpentine (

magnetite, representative of the Cu-Fe core, seen as sets of fractures; magne

zone composed of garnet fragments cemented by a silica-carbonate matr

Displacements of reshaped oscillatory zonation, showing that the brecciati

boundaries in garnet with compositional changes. The texture is indicative f

light And91Gr6Sps4; intermediate light: And65Gr30Sps5; dark: And43Gr5corrosion boundaries in calcite from distal Fe zone in Paulus (82). (h

boundaries in garnet from Zn-Pb zone in Paulus (3375). Composition of

And16Gr84. The most And-rich analysis in this sample is: And89Gr2Sps8.

abraded sphalerite (Sph1). Enclosed galena (Gn) has developed a ‘jigsaw’

(j to l) Characteristic fluidisation-recementation textures in the distal Zn-P

magnetite and ankerite (MAS) at the boundary with galena (Gn). Patches

(Carb). (k) Detail of the selvage of magnetite and ankerite at the boundary

behaviour of MAS relative to galena (Gn), shown by deformed marginal h

galena outside the MAS within the matrix carbonate.

fO2-fS2 values. However, if we assume a temperature

of 550 jC, the fO2 value obtained is too low for Hem

stability. A slightly higher temperature, 570 jC,however, gives us a range of fO2-fS2 values (Table

5) that fit the Hem-Py stability field far better. For

distal skarn, we assume temperatures lower than those

considered for the garnet-pyroxene assemblages in

the Fe zone (see above). The best fit with respect to

Hem-Py stability in fO2-fS2 space is obtained at

temperatures of 400 jC (Table 5; Fig. 14). A further

consideration for the choice of temperatures is the fact

that the FeS content of sphalerite in equilibrium with

pyrite will increase as fS2 decreases at any given

temperature (Gamble, 1982).

The overall results (Table 5; Fig. 14) indicate that

even though there are significant differences in abso-

lute values between proximal and distal skarn,

expressing the decrease in both oxidation and sulphi-

dation state (i.e., f 8 log units for fO2, f 5 log units

for fS2), the stability conditions will nevertheless

remain fixed on the Py/Hem buffer as the temperature

decreases from 570 to 400 jC. This is in full accor-

dance with observation of Py and Hem in all associ-

ations from the Zn-Pb zone. Despite the decrease in

fO2 that accompanies the drop in temperature from the

proximal to distal environment, there is no major

change in the oxidation state of the skarn associations

(both proximal and distal are at the Mt–Hem buffer).

On the contrary, the sulphidation state decreases

significantly from proximal to distal skarn. The log

fS2 value of � 7 is much closer to the Po-Py buffer in

distal skarn (positioned at log fS2 =� 7.5) than in

proximal skarn. This is in agreement with the obser-

) and back-scattered electron images (d to l), showing retrograde

dark) in chalcopyrite, from Cu-Fe core (58). (b) Brittle brecciation in

tite is enclosed within bornite (164). (c) Blown-apart garnet in the Fe

ix (3103). (d) Oscillatory zoned garnet with fracturing (5598). (e)

on follows the shock-induced event (798). (f) Absorption-corrosion

or changes in porosity during decarbonation. Garnet composition as:

1Sps6 (e1). (g) Pyroxene inclusion with zonality and absorption-

) Compositional changes during brecciation, absorption-corrosion

garnet: light: And54Gr35Sps10; medium grey: And42Gr54Sps4; dark:

(i) Pressure tail realised by an overgrowth of sphalerite (Sph2) on

border with the sphalerite contemporaneously with abrasion (165).

b zone (43). (j) Hook structure (arrowed) developed by selvage of

of galena (Gn) outline the selvage border with the matrix carbonate

with galena (Gn), showing a symplectite-like character. (l) Ductile

ooks extending into galena (white arrow). Note irregular patches of

Fig. 12. Back-scattered electron images showing retrograde deformation textures in the Fe zone. (a to c) Magnetite deformation (82). (a)

Characteristic sets of en echelon fractures assisting micro-shear deformation (arrowed). (b) Crossed jointing in magnetite, crosscutting

overgrowth patterns. (c) Overgrowth in pressure shadow (arrowed) in abraded magnetite. Abrasion is indicated by recementation of fragments

(slightly darker). (d to i) Aspects of deformation and mutual adjustment between refractory minerals (all 798 except g). (d) Inliers of andradite

within pyrite. Recrystallisation of magnetite in the matrix between andradite grains appears contemporaneous with development of ‘jigsaw’-like

boundaries between And and Py (arrowed). (e) Ductile deformation at margins of magnetite crystal (arrowed). (f) Arrangement of pyrite clasts

and syn-kinematic recrystallisation of magnetite within a pressure shadow in deformed, coarse andradite. (g) Contortion of magnetite, assisted

by brecciation in garnet within CGM (e17). (h) Equilibrium boundaries (black arrow) and diffusion-reaction boundaries (white arrow) between

andradite and magnetite. (i) Bands of andradite enclosing pyrite clasts that interrupt fractures (black arrow). Note ductile deformation of pyrite at

the shoulder of garnet (white arrow).

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370352

vation by Vlad (1974), that pyrrhotite can coexist with

pyroxene of the same composition as we report here

from Dognecea. Also in agreement is our observation

of very rare exsolved pyrrhotite in sphalerite from

Paulus.

Based on this discussion we can say that the distal

skarn has a reducing character in comparison with the

proximal, at comparable oxidation conditions. We

also note the buffering influence of high FeS sphal-

erite (16.5 mol%) in Paulus that favours formation of

inclusions of Gr-rich garnet (And49Gr49Sps2), a garnet

that requires lower fO2 than an And-rich equivalent

(Einaudi and Burt, 1982).

The experimental runs of (Burton et al., 1982)

failed to stabilise Hed70Joh30 under oxygen-buffered

conditions, within the temperature range 600 to 800

jC. Instead, a clinopyroxene that ‘exhibits a vermic-

ular intergrowth texture that was not present in the

Fig. 13. Photomicrographs in reflected light (a, c to g) and back-scattered electron images (b, h, i) showing retrograde healing and sealing

textures. Figs. (a to f) are from the Cu-Fe core; (g to i) from the Fe zone. (a) Rounding and abrasion of a fragment of magnetite and gangue

silicate in a bornite matrix, which also partially replaces the fragment (68). (b) Sets of new cleavages developed across grains of magnetite;

darker areas are forsterite (66). (c) Fresh sets of cleavages (arrowed) crosscutting inclusion-rich cores in magnetite (68). (d) Abrasion of

magnetite, assisted by chalcopyrite. Note inclusions of valleriite (arrowed) in the magnetite core (164). (e) Dusty silicate inclusions in a sealing

fracture in magnetite (arrowed; 66). (f) Annealing in magnetite, indicated by 120j triple junctions (arrowed; CuSI). (g) Two fragments of

prograde magnetite, in which rotated swarms of inclusions are seen. The boundary between the two (arrowed) is sealed (PP). (h) Welding of

disrupted magnetite assemblages. Former abraded fragments are indicated by overgrowths and zones of re-cementation of pieces resulting from

earlier abrasion (arrowed, 82). (i) Late extensional cracking (arrowed) overprinting garnet, subsequent to recrystallisation (798).

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 353

starting clinopyroxene’ was obtained in the presence

of bustamite. We note that mean pyroxenes in both

distal skarns can be considered close to Hed70Joh30,

despite the presence of additional minor Di (Table 4).

Indeed, values around Hed75Joh25, effectively com-

pletely lacking Di, are obtained from individual lamel-

lae in Paulus (Fig. 4). In our material, pyroxmangite is

present instead of bustamite, and the assemblage is

stable with pyrite in the presence of hematite, and

sphalerite containing 16.5 mol% FeS (Fig. 5l). This

stresses the lower sulphidation character of the distal

skarn relative to proximal, and also the lower forma-

tion temperatures that contribute to the absence of

bustamite.

Fig. 14. Sketch illustrating the prograde mineralising path during skarn formation in T-fO2 space. Phase equilibria are based on the experimental

work of Gustafson (1974), Burton et al. (1982), Gamble (1982), Myers and Eugster (1983), and using data presented by Einaudi et al. (1981)

and Bowman (1998). The position of the calc-silicate equilibrium reaction: Cc +Qz!Wo+CO2 is shown assuming an X(CO2) = 0.1, and

Hm+Cal +Qz =And is at X(CO2) = 0.03, considering a decrease in this parameter from proximal to distal settings (data presented by Bowman,

1998). The position of sulphide-oxide equilibria at 550 jC (black dashed lines), and at 400 jC (light grey dashed lines), are from data presented

for Cu-Fe sulphides by Simon et al. (2000). Mineral abbreviations are given in Table 1.

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370354

Our observations suggest that Hed70Joh30 could be

attainable at the Py–Hem buffer if the sulphidation

state is as low as log fS2f� 7, and at lower temper-

atures, i.e., 400 jC rather than higher. Again, we

stress the potential role of polysomatism in stabilising

such pyroxene in the presence of pyroxmangite.

The highest sulphidation effect in the Fe zone is

observed in pyroxene (Di40-50Hed30-40Joh10-20) from

associations in the upper parts of orebodies in both

proximal and distal settings (Simon Iuda, Ocna Tur-

ceasca, Stefania, Paulus, etc.). Such associations are

formed at the contact between the Fe and Zn-Pb

zones, where Hem and Py are stable, even though,

as seen in the association from Paulus, monomineralic

magnetite can also be present at this contact. These

pyroxenes, occurring as inclusions in other minerals,

form a select group that share a number of common

features. The inclusions (e.g., Figs. 5f and 11g) may

display a compositional zonation from core to margin

that parallels the Di-rich to Hed-Joh-rich trend in

pyroxene between the Fe and Zn-Pb zones. With the

exception of the occurrence from Paulus (sample 82),

all are buffered by pyrite. In the Paulus case, we take

the presence of Gr50 inclusions in calcite (see Section

4) to infer a comparable oxidation–sulphidation state

for the calcite buffer as for the aforementioned sphal-

erite elsewhere in the same orebody (e.g., sample 40).

7.4. Zonation trends shown by skarn assemblages

We note a similarity in mean pyroxene composi-

tion in Stefania and Dognecea N. (3596; Fig. 2; Table

4), for which FeSSp (Table 5) suggests broadly com-

parable sulphidation states. Both these pyroxenes are

part of a broad compositional field indicated on Fig. 4,

in which all pyroxene inclusions from the Fe zone, at

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 355

the contact to Zn-Pb zone, also plot. Because of the

buffering conditions of the pyroxene inclusions (see

Section 7.3 above), we are able to see how the

prograde to retrograde evolution in this contact zone

overlaps, in terms of sequences of sulphidation–

oxidation states monitored by pyroxene zonation, with

the evolutionary trend followed during transition from

the Fe to the Zn-Pb zone. The sulphidation state for

each assemblage that includes the buffer mineral

shows, however, an overlap with formation conditions

of some prograde assemblages. This applies not only

to the association in the Zn-Pb zone at the margin to

proximal skarn mentioned above (sample 3596), but

also Hem-Py associations at the contact between Fe

and Zn-Pb zones in proximal skarn without marked

retrograde overprinting (e.g., sample 72; Simon Iuda).

We conclude therefore that, at the boundary between

the Fe and Zn-Pb zones (Fig. 2), assemblages on both

sides share a sulphidation–oxidation state that has

either common position relative to the Py–Hem buffer

(72 versus the rest of the group), or have directly

comparable fS2-fO2 values, that relate to temperatures

of formation and proximal-to-distal position of the

orebody in question (Fig. 14). We thus consider that

the contact between the Fe and Zn-Pb zone can also

be considered ‘transitional’ in terms of sulphidation

state. Although the latter controls pyroxene composi-

tion, a convergence of local factors, especially the

buffering effects of certain minerals, impact signifi-

cantly on pyroxene stability.

In contrast to the Fe zone, the pyroxene from the

Zn-Pb zone displays a significant compositional trend

between proximal and distal setting (Fig. 15a). A

slight, but nevertheless significant decrease in the

Joh component is seen between proximal (Joh40)

and distal skarn (Joh20-30). Diopside has a similar

trend, attaining lowest values in distal skarn (Di< 10),

whereas the Hed component is highest in distal skarn.

There is also an inverse correlation between the

(Di + Joh)–Hed ratio of pyroxene (Pxi: 1.3–0.25) and

FeSSp (5–16.5 mol%) from central proximal (Simon

Iuda) to distal skarn (Paulus and Dognecea). The two

distal skarns show a striking similarity in both FeSSpand Pxi. Within Simon Iuda itself, there is a similar

inverse correlation, although over a lesser interval

(Pxi: 1.3–1; FeSSp: 5–6.5), across a vertical extent

of 100 m, from the 357 m upwards to the 460 m level

(Fig. 15b). Because of these comparable trends in

sulphidation state of the association controlling over-

all pyroxene stability, a centric zonation can be

defined upwards and outwards from a subjacent centre

(see Section 9).

7.5. Evolution trends shown by skarn assemblages

The ‘transitional’ sulphidation zone, positioned

between the Fe- and Zn-Pb zone is relevant for the

evolutionary trends indicated by pyroxene composi-

tion (Fig. 15a,c). The positive correlation between

Hed and Joh components, going from the Cu-Fe core

and Fe zone to the Zn-Pb zone is in contrast with the

negative correlations between Di and Hed or Di and

Joh. This is consistent with the observation that the

dominant pyroxene in the Cu-Fe core and throughout

the Fe zone (Di>70) gives way to pyroxene in the Zn-

Pb zone, in which (Hed + Joh) is greater than the Di

component (>60%). This is a response to the shifts in

the sulphidation and oxidation states that are induced

upwards and outwards by the decrease in temperature,

i.e., from 650 to 570 jC in Simon Iuda, and from 650

to 450 jC from proximal to distal (Fig. 14). Pyroxene

from the transitional zone has compositions situated

mid-way between those of the Fe and Zn-Pb zones.

However, this transitional zone too is formed at

different temperatures from Simon Iuda to outermost

distal locations. In Simon Iuda itself, it formed at

temperatures somewhat higher than 570 jC, whereasin marginal to proximal or distal settings, formation

temperatures were only 470 to 440 jC (Fig. 14). We

conclude from these observations and arguments that

the boundary between the Fe and Zn-Pb zones has a

transitional sulphidation-oxidation character.

At the uppermost part of Simon Iuda, Di40Hed20-Joh40, in the presence of bustamite, indicates a point

in the Zn-Pb zone with uniquely high fO2 and fS2conditions. FeSSp here too suggests that the sulphida-

tion state is markedly different from pyrite-dominated

assemblages (Table 5). Based on compilations (Burton

et al., 1982; Nakano et al., 1994), such compositions

are unusual for Zn-Pb skarns elsewhere.

The compositional variation of Hed versus Di

components from proximal to outermost distal in

the Zn-Pb zone shows a similar negative trend as

from Cu-Fe core to Zn-Pb zone in Simon Iuda. In

contrast, Hed versus Joh and Joh versus Di both show

opposing trends to that shown in the transition from

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370356

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 357

Cu-Fe core to the Zn-Pb zone in Simon Iuda (Fig.

15c). The decrease in fO2 and fS2 from proximal to

distal in the Zn-Pb zone can therefore be expressed by

a negative correlation between FeSSp versus Pxi (Fig.

15d). Only in the outermost distal skarn from the Zn-

Pb zone does pyroxene show a Hed-dominant com-

position (Hed>60), although still with significant Joh

component (Joh>20). Such a pyroxene is more typical

for Zn-Pb skarns elsewhere (e.g., Burton et al., 1982).

As discussed above, the skarn system remains at the

same Mt-Hem buffered conditions as in the upper-

most proximal skarn (Simon Iuda), but a gradual

decrease in fO2-fS2 values appears towards the out-

ermost distal setting, as temperatures decrease from

570 to f 400 jC. Formation of hedenbergite (i.e.,

Hed>60) is therefore attained in more reducing condi-

tion in the distal skarn.

A similar reducing character is obtained from the

observed drop in fO2 at the beginning of the retro-

grade stage. Although seen in all assemblages, the

drop is especially marked by replacement of hematite

by magnetite (Section 6) or can be inferred from

garnet compositions in assemblages from the Fe zone,

in which magnetite remained the stable Fe oxide. We

consider that this is the reason why the zonation of

pyroxene within individual buffered inclusions (in the

transitional zone) tends to reproduce the composition-

al trends identified between the Fe and Zn-Pb zone.

The overall highly oxidised state of the skarn

system is indicated by the dominantly high-And

garnet. With few exceptions (Section 4), the compo-

sitional variation in garnet, best expressed in oscilla-

tory zonation, will always include an And-rich

component in that zonation, even when the mean

value indicates a Gr-rich composition (e.g., 3913,

Fig. 4). This is in agreement with the interpretation

of Einaudi (1982) that the oxidation state of a given

skarn system is more or less fixed at its initiation. The

Di-rich character of pyroxene in the Cu-Fe core and

Fe zone, comparative to Di-poor compositions in the

Fig. 15. Composition trends in skarn pyroxene. (a) Zonation trends, expre

pyroxene from the Cu–Fe and Fe, Zn–Pb and ‘transitional’ sulphidation z

deposit (from Dognecea in the south to Paulus in the north). (b) Plot of FeS

trends in the Zn–Pb zone, upwards and outwards from Simon Iuda to dista

shown as hedenbergite versus johannsenite, hedenbergite versus diopside a

evolution trends (arrows) shown by pyroxene composition between the Cu

from proximal to distal. (d) Variation in (Di + Joh) versus FeSsp in coexistin

recognised trends, vertically within the proximal deposit, and from the pr

Zn-Pb zone, indicates that an upwards trend towards

relatively higher oxidation and sulphidation states

characterised formation of Cu-Fe core and Fe zone

in proximal skarns, and also outwards in the Fe zone

of the other orebodies. This trend became inversed

during formation of the Zn-Pb zone. Both trends were

controlled by temperature gradients, with local varia-

tion within individual orebodies induced by the im-

mediate skarn setting (see Section 9).

A further parameter influencing skarn evolution is

X(CO2). We note that the disappearance of forsterite

and appearance of garnet, in the presence of the same

Di-rich pyroxene at the boundary between the Cu-Fe

core and Fe zone, probably indicates an isothermal

increase in X(CO2). The presence of tremolite asso-

ciated with fronts of apatite in retrograded assemb-

lages from this boundary may be an additional

indication of zonation resulting from variation in

X(CO2). Considering the zoning sequences reviewed

by Bowman (1998), Di + Fo +Cal/Di + Tr +Cal would

correspond to a quartz-saturated assemblage rather

than one saturated by dolomite. We note that tremolite

can occasionally be seen instead of diopside at the

upper part of orebodies in the Fe zone. This too may

indicate variation in X(CO2). A dramatic X(CO2)

variation may also have accompanied the onset of

the retrograde stage. Pseudomorphous replacement of

diopside by tremolite, or breakdown of bustamite–

pyroxene to calcite and quartz in the Zn-Pb zone can

both be considered as indications for such a variation.

8. Formation of textures

All textures presented in this contribution can be

attributed to the type of reaction-mass transport feed-

backs introduced by Ortoleva et al. (1987a,b). Fol-

lowing the nomenclature in Section 2.1, the prograde

textures seen at the macroscopic scale (Table 7; Figs.

6 and 7) can be discussed as resulting from supersat-

ssed in the diopside, hedenbergite and johannsenite components in

one between the Fe and Zn–Pb zones, along the 10 km strike of the

sp/Pxi along the same 10 km strike. The diagram shows the zonation

l skarn. (c) Correlations between the three components in pyroxene,

nd johannsenite versus diopside. The diagrams are suggestive for the

–Fe core, Fe zone to the Zn–Pb zone, and along the Zn–Pb zone

g pyroxene–sphalerite pairs from the Zn–Pb zone showing the two

oximal to distal skarn. See Section 7.5 for additional discussion.

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370358

uration-nucleation-depletion cycles (precipitate band-

ing), competitive particle growth (CPG), and reactive-

infiltration feedback (RIF) mechanisms. The retro-

grade textures in Figs. 9–13 (see also Table 8) are

interpreted as resulting from the action of back-flow

fluxes (fluxes redirected back into the skarn), formed

during devolatilisation reactions in the manner Dipple

and Gerdes (1998) discussed propagation of RIF and

hydrodynamics at the skarn front.

8.1. Precipitate banding and ‘chemical oscillators’ in

the Fe and Zn-Pb zones

Two patterning operators are considered in con-

nection with precipitate formation: Liesegang banding

and CPG (see below). Their combined action, as well

as the presence of characteristic textures, such as those

we have described as mottled and mossy, is evidence

that the skarn front during formation of the Fe and Zn-

Pb zone was an unstable coarsening front of reaction

in the sense defined by Ortoleva (1994; see below).

The formation of the two ‘chemical oscillators’ (Fig.

8f to h) and their crystal fronts (Fig. 8d,e) can be

interpreted as an example of autocatalytic surface

attachment in a Liesegang environment, perhaps par-

alleling the experiments with minerals in solid solu-

tion series where end members have dissimilar

solubility (Putnis et al., 1992). The scalloped fronts

(Fig. 6a) and mineral banding (Fig. 6b) seen at the

skarn–marble contacts are part of the range of tex-

tures considered by Ortoleva et al. (1987b) and Dipple

and Gerdes (1998) as manifestation of morphological

instabilities during propagation of RIF at the skarn

front.

As seen in Fig. 8a to c, macroscopic textures do not

overlap with those seen at the microscopic scale (Fig.

7d to h). Magnetite within rhythmic banding, for

example, lacks oscillatory zonation, even though

magnetite is one of the two ‘chemical oscillators’ in

the Fe zone. Whereas the minerals forming the mac-

roscopic patterns are compositionally very close to

end members within solid solution series, i.e., Di,

And, and silica free-Mt (Figs. 3 and 4), the ‘chemical

oscillators’ are found in garnet-pyroxene associations

where extended compositional ranges are characteris-

tic for each mineral (see Section 4).

Although macroscopic textures may also form

patterned ‘islands’ within the macroscopically unpat-

terned skarn, they are nevertheless mainly to be found

at marble–skarn contacts, in contrast to the micro-

scopic textures that are more characteristic for the

envelope to the inner core of an orebody. These

observations imply that formation of the macroscopic

patterns could be considered as having followed

formation of the skarn hosting the ‘chemical oscilla-

tors’. Bearing in mind the compositional characters of

the two classes, macroscopic and microscopic, as

discussed above, we may conclude that a recurrent

compositional trend, from And70-Di70, reverting back

to And>90-Di>90, is recorded in garnet, pyroxene and

magnetite throughout the patterning in the Fe zone.

We could alternatively consider that the microscopic

oscillators and macroscopic patterns were formed

more or less simultaneously, contemporaneous with

the unpatterned side of Fe zone giving way to the Zn-

Pb zone, or the skarn front ceasing its advance into the

marble.

As stated by Ortoleva (1994), there are many

geological environments, for example rocks undergo-

ing changes in stress, temperature or compositional

gradients, in which the system is left with small

precipitate particles, after the cessation of nucleation.

As Ortoleva continued to demonstrate, these are the

types of environment that will promote CPG, causing

nodular, spotted, orbicular, mottled and mossy pat-

terning to develop at an unstable coarsening front of

reaction. Chemical waves, as well as other types of

oscillation, may develop spontaneously, especially

close to thermodynamic equilibrium (Chu and Ross,

1990; Hjemfelt and Ross, 1991). Patterning will only

occur, however, if the fluctuations attain their macro-

scopic amplitude at the same time when the minerals

that express the precipitate banding–oscillatory zona-

tion are stable.

In our case, assuming synchronous pattern forma-

tion, we can say that the combined package of

precipitate banding and crystal growth covers two

different ranges of amplitude, at scales differing by

four orders of magnitude (cm- or dm- and Am-scales,

respectively). In the Fe zone, we note that among the

minerals that form solid solutions series, it is only

garnet that responds to chemical fluctuations in terms

of oscillatory zonation. Pyroxene, and especially the

Di-rich pyroxene stable within the local environment

at that time, does not promote such fluctuations.

Nonetheless, pyroxene, when together with magnetite,

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 359

is involved in macroscopic precipitate banding (Figs.

7e,f and 8a to c). Only in the distal Zn-Pb zone, do we

see a Hed-rich pyroxene that can express oscillations

by means of extended compositional fields and lamel-

lar intergrowths (Figs. 4 and 5j). As discussed above

in Section 5, the development of polysomatism would

be a necessary condition to obtain such pyroxenes in

equilibrium. In such a case, oscillation amplitudes in

the range from nano-to micron-scale would be re-

quired. We have earlier considered (Ciobanu et al.,

2004, in press) that long- and short-range polysomes,

on comparable scales as we propose for pyroxene, can

express the compositional and textural features of Bi-

sulphosalts in polysomatic series, also found in the Fe

zone.

Some of the examples presented in Section 5 are

discussed in more detail below, with reference to the

conceptual models, numerical simulations or experi-

ments that are relevant for their interpretation.

8.1.1. Liesegang banding and morphological

instabilities

Liesegang phenomena (Liesegang, 1913; Ostwald,

1925; Ortoleva, 1994; Krug et al., 1996; Krug and

Kruhl, 2001) can be considered as the spontaneous

formation of banded patterns in linear space, local-

ised rings of precipitate in 2D and screw patterns in

3D, obtained by inter-diffusion of two co-precipi-

tates. Ostwald (1925) showed that the mechanism for

band formation could be produced without precipi-

tation before the initiation of inter-diffusion, as a

result of sequential events involving supersaturation-

nucleation-depletion in the zone where co-precipitate

concentration profiles meet. Today, the Ostwald-Lie-

segang cycle (OLC) is recognised as a powerful

pattern-forming operator, which involves coupling

between particle growth and transport (e.g., Ortoleva

et al., 1987a; Ortoleva, 1994).

Morphological irregularities resembling those from

Magnet Quarry (Fig. 7a to d), i.e., fine structure

within bands, lateral gaps and radial alleys of gaps

within bands, apparent band branching and transition

to speckled patterns, were modelled by Krug et al.

(1996) as morphological instabilities during self-pat-

terning by coupling CPG to Liesegang banding (post-

nucleation model; see below). The appearance of

garnet within breaks in the alleys suggests that pat-

terning thresholds may be locally ‘reversed’ or

delayed at points where nucleation involves more

complex compositions (cf. Krug et al., 1996).

The ‘wiggle-Liesegang’ rhythms we describe (Fig.

7e,f) may be evidence of flow-driven OLC, since

Sultan et al. (1990) showed that a ‘wiggle’ pattern is

obtained characteristically between unsteady deposi-

tion and Liesegang banding. Interference of Liesegang

rings, like those in Fig. 7e,f, has seen modelled by

Krug et al. (1996; pre-nucleation model) varying the

diffusion parameters of OLCs. In the transition zone

between banding and ‘wiggle’ texture (Fig. 8a to c),

supersaturation in component A (Mt) is achieved at

the fastest nucleation rate of component B (Di), the

essential kinetic bottleneck required to obtain Liese-

gang banding in an Ostwald cycle (Ortoleva, 1994),

and thereby inducing an instantaneous switch to

crystallisation of component A. Irrespective of wheth-

er the post- or pre-nucleation model is considered,

band branching in natural samples is consistent with

pattern formation as a result of self-organising pro-

cesses (Krug et al., 1996).

8.1.2. Competitive particle growth

Ortoleva (1994) discussed CPG as another type of

feedback instability that may form macroscopic pat-

terns at reactive fronts after cessation of nucleation. It

involves the dependence of the dissolution equilibri-

um constant on particle radius of curvature. The CPG

could be actually considered as a competitive type of

Ostwald ripening process. The CPG model of post-

nucleation states that the competition can also be

cooperative, such that deviations from the local aver-

age particle size tend to amplify themselves and, as a

result, promote the appearance of what are termed

‘greedy giants’. The CPG supplement Liesegang

banding as a patterning operator, and the two may

be coupled in unstable coarsening fronts of reaction.

The skarn-dominated garnetFdiopsideFmagnetite

assemblages, e.g., the nodular, spotted and/or orbicu-

lar textures in Fig. 6g to j, form 3D patterns of the type

that Jakob et al. (1994) ascribe to CPG in Liesegang

fronts of precipitation. Although they are very differ-

ent to other macroscopic patterns, we also attribute the

skeletal magnetite forming mottled textures within

marble (Fig. 6f), as well as the mossy branches of

magnetite in Fig. 6e, to precipitate patterning at

unstable coarsening fronts. Development of such

patterns was obtained by experiment (Ortoleva,

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370360

1978). During these experiments, the self-organisation

of uniform sols into mottled, halo and spiral patterns

was obtained. All involve skeletal growth. Such

experiments were later on numerical simulated by

Ortoleva (1994) using the CPG model.

8.2. Reaction infiltration feedback coupled to devo-

latilisation at the skarn front

The suite of retrograde textures presented in Section

6 indicates a sequential development in which the early

piercing of clusters, tiling and other shock-induced

textures indicate that the fluids recorded peaks of

transient overpressures. Although these early textures

are comparable across all zones, they are different in the

Cu-Fe core compared to the other two zones, in terms

of mineralogy and also the volume of volatiles that

were involved in such events. For example, although

apatite is abundant in all three zones, it is seen in

piercing clusters in the Cu-Fe core (Figs. 5b and 9a),

whereas in the Fe and Zn-Pb zone, apatite is seen as

resulting from the latest residual fluids, forming equil-

ibrated assemblages with calcite (Fig. 5e). The range of

trace minerals discussed in Section 4.4 is associated

only with the piercing clusters in the core; apatite from

the other zones is not implicated in the enrichment of

Au or other trace elements. We thus consider that

propagation of RIF at the skarn front triggered two

devolatilisation events concluding with back-flow

fluxes, which is why they are interpreted as retrograde

events. The first is in the Cu-Fe core, formation of

volatile-rich assemblages by consuming the volatiles in

the fluids, and the second is in the Fe and Zn-Pb zones.

The latter is seen as a devolatilisation reaction triggered

by erratic decarbonation at the skarn front, which is

interpreted as carbofracturing (see below).

The textures shown in Fig. 13 indicate that, at the

end of the action of back-flow fluxes, the skarn system

is healed rather than being connected to a system of

open fractures reaching the surface. This is further

evidence for the fact that retrograde back-flow fluxes

formed in response to skarn system evolution rather

than reflecting a hydrothermal collapse as advocated in

the case of numerous shallow skarns (Meinert, 1992).

8.2.1. Devolatilisation in the Cu-Fe core

In the Cu-Fe core, the widespread ‘piercing’ (Fig.

7a), and associated serpentinisation of early forsterite,

induced brecciation. The pressure tails seen at the

margins of the pierced areas (Figs. 9a and 11d,e)

indicate that fluids accompanied the subsequent brec-

ciation and rounding of magnetite (Figs. 11b and

13a,d). The latter was assisted by local remobilisation

of ductile sulphides (chalcopyrite and bornite). Apa-

tite is especially abundant in the retrograde stage in

the core, as clusters (Figs. 5b and 9a) in virtually most

areas that underwent subsequent brecciation. Subse-

quent to formation of the main skarn-ore assemblage

in the Cu-Fe core a volatile-rich trace mineral associ-

ation (e.g., phlogopite, chlorapatite, turneaureite, val-

leriite and ludwigite) is seen in overprinting trails and

clusters. The process also appears to have enabled the

extraction and transport of Se, Te, Au, Ag, Bi, Au, Co,

Sn, etc., since they are intimately associated with

volatile-rich minerals in the Cu-Fe core.

Contemporaneous formation of valleriite in this

sulphide-rich environment (Fig. 13d) is an important

key to the characterisation of the fluid, since valleriite

typically develops during serpentinisation in high-

volatile environments (e.g., Genkin, 1971; Nickel

and Hudson, 1976).

At the end of devolatilisation, ongoing healing in the

Cu-Fe core is indicated by formation of new sets of

cleavages (Fig. 13b,c), recrystallisation (annealing;

Fig. 13f) of deformed and fractured magnetite and by

cementation of fractures (Fig. 13e). Annealing of

magnetite aggregates is restricted to this zone, affirm-

ing that the retrograde event was accompanied by

higher temperatures (f 550 jC) here than in the rest

of the orefield. The resultant textures are reminiscent to

those preserved in pyrite-rich massive sulphide ores

that have undergone amphibolite facies metamorphism

(e.g., Craig and Vokes, 1993; Cook et al., 1993).

8.2.2. Carbofracturing in the Fe and Zn-Pb zones

In the outer Fe and Zn-Pb zones we recognise the

initiation of collapse in the skarn system from the

range of pressure-induced textures (Figs. 9f to l and

10) and brecciation (Fig. 11). This can be triggered

when infiltration ceases and decarbonation of wall-

rock becomes erratic, implicitly resulting in transient

peaks of CO2-induced overpressure. Therefore, we

consider the entire retrograde episode in these two

zones as carbofracturing.

The interaction of such CO2-induced peaks of

overpressure with waning fluids that reached the

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 361

outermost limits of the system are highly predictable

to produce immiscibility, thereby resulting in lower-

salinity, less dense fluids and the exsolution of vola-

tiles into the vapour phase. In turn, this will cause

volume increase and wallrock disruption as consid-

ered in the comparable case of vapour-fluid mixing at

the instance of resurgent boiling in porphyry systems

(e.g., Drummond and Ohmoto, 1985; Bowman,

1998). Similarly, carbofracturing in skarns relates to

such intertwined chemical and mechanical changes

that overprint former assemblages at the same site,

i.e., hematite pseudomorphosed by magnetite (Fig.

10d to f) and pressure-induced textures in garnet

(Fig. 10b,c,j to l). Retrograde collapse of skarn

systems is common (Meinert, 1997), but moreover,

at Ocna de Fier-Dognecea we see the dual character of

this collapse, further stressing that it can be charac-

terised in terms of a carbofracturing episode. We note

that pseudomorphosis almost entirely converted he-

matite ores into magnetite, a fact that fortuitously, also

significantly improved the quality of the ores for

exploitation purposes.

We interpret the shock-induced textures, especially

tiling (Fig. 10j to l) as resulting from mechanical-

chemical coupling during the action of such back-flow

fluxes. At Ocna Turceasca, (Fig. 12d to f,h,i), sus-

tained overpressure fluctuations could enhance rip-

pling fronts of diffusion as we see in the reshaped

prograde garnet zoning (Figs. 10l and 11e). This and

other evidence of pressure-induced patterns in garnets

from Ocna de Fier-Dognecea (e.g., sculptured crystal

faces in Fig. 6l, flattened crystal edges in Fig. 10k) are

strong evidence of the fact that oscillatory fluid-

pressure was a critical parameter (cf. Brenan, 1991)

at this stage. Such a conclusion is also fully supported

by the deep setting of the skarn (Nicolescu and

Cornell, 1999).

‘Blown-apart’ textures (Fig. 11a,c), overgrowths

(Figs. 10b,c and 12c,g), tails (Fig. 11i), hooks (Figs.

11j and 12e), jigsaw borders (Fig. 12d) and other

evidence of brecciation, cementation and remobilisa-

tion (Fig. 11l) are consistent with sustained transient

fluid fluxes. Widespread absorption-corrosion bound-

aries (Fig. 11f to h) developed during the brecciation

episode represent the kind of textures interpreted by

Yardley et al. (1991) and Yardley and Lloyd (1995) as

typical for development of transient enhanced poros-

ity during quick propagation of decarbonation reac-

tion, as well as evidence for infiltration metasomatism

in metamorphic terranes. We argue that such textures

are highly indicative of the considerable impact

caused by development of decarbonation reaction

during patterning under RIF. Similarly, micro-shear

assisted deformation in refractory assemblages of

garnet-magnetite (F pyrite) are also fully concordant

with the persistent action of fluid fluxes (Fig. 12). All

these features are similar to those ascribed to fluid-

isation and ‘fluid pumping’ in volcanic–subvolcanic

breccia pipe ores (e.g., Burnham, 1985; Sillitoe,

1985), despite the latter environment involving sig-

nificantly shallower depths than the skarn setting. The

near-ductile behaviour of refractive minerals during

micro-shear fluid assisted deformation is again remi-

niscent of deformed massive sulphides in metamor-

phic terranes (Gilligan and Marshall, 1987; Craig and

Vokes, 1993; Marshall et al., 2000). It is clear that

fluid fluxes and temperatures of 400–550 jC charac-

terised the deformation and reworking of prograde

assemblages during the retrograde stage. The intensity

of the carbofracturing episode was amplified in the

context of the buried skarn system, a fact especially

shown in the formation of back-flow fluxes under

high-pressure.

9. Evolution of skarn system in a source centred

model

Given the anvil-shape of the intrusion, centred on

the median part of the orefield (Fig. 13), we believe

that formation of a self-focused source of hydrother-

mal fluids was initiated in the sink beneath Simon

Iuda. Development and focussing of this source into a

plume can be considered in terms of an idealised

model for intrusion with marginal convective flow

generated towards the end of crystallisation (e.g.,

Candela, 1991; Shinohara and Kazahaya, 1995). The

Cu-Fe core has a petrological character that contrasts

with other parts of the entire orefield. Each of these

characteristics (forsterite as main skarn silicate; inti-

mate co-crystallising relationships between forsterite,

Cu-Fe sulphides and magnetite, and the Mg-bearing

character of the latter with exsolved skeletal spinels) is

fully concordant with proximity to the source. We

thus interpret the core formation as a result of this

unique setting, in which the first up-streaming buoy-

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370362

ant fluids were met by limestone indented in grano-

diorite. Well-localised plumes will be up-drafted,

provided that self-focused buoyant fluids can easily

interact with a porous medium at the point of inter-

action with country rock (Ortoleva, 1994; Renard et

al., 1998). Therefore, strong reaction between the up-

streaming fluids and the wedge-shaped limestone

block are ingredients causing rise of the plume at

the initiation of the skarn system. The presence of a

reactive indent inferred by this local setting would

have been critical for sustained self-focussing of the

source, especially when the 10 km depth of formation

estimated for the orefield (Nicolescu and Cornell,

1999) is taken into consideration. The change from

magnesian to calcic skarn defines the boundary be-

tween the Cu-Fe core and Fe zone as a sharp front of

reaction (see above). The change in skarn mineralogy

across this front, in which forsterite disappears and

andradite appears instead, is mirrored by differences

in ore mineralogy, notably by the disappearance of

Cu-Fe sulphides.

The superimposition of an early retrograde stage is

seen in the core by an abundance of volatile-bearing

phases, accompanied by minerals containing a range

of exotic elements (Se, Te, Ag, Bi, Au, Co, Sn, etc.).

These unique associations occur in ‘piercing’ trails

and clusters (Fig. 9a to e). We can assume that

forsterite formation increased porosity at the reaction

front by volume reduction (Rubin and Kyle, 1998).

This would induce an initial decrease in transient

pressure of the fluids leading to destabilisation of

volatile species, including boron, fluorine and phos-

phorus. However, the apparent shock-induced charac-

ter of the suites of textures we have described in these

ores suggests that peaks of transient overpressure

rapidly followed volatile extraction. The sudden devo-

latilisation was manifested with different intensities

throughout the Cu-Fe core.

We note that the early-stage concentration of Cu-Fe

sulphides, together with the exotic trace mineralogy,

in this inner core of magnetite ore is remarkably

similar to the early cores with potassic alteration in

porphyry copper deposits (e.g., Elatsite, Bulgaria;

Tarkian et al., 2003). Although, the presence of

phlogopite in the Simon Iuda core is also reminiscent

of alteration in porphyry environment, the abundance

of apatite, its association with gold and the presence

of uraninite and various REE-phosphates in the skarn

also forces comparison with mineral associations

generally considered typical for Fe-oxide-Cu-Au

deposits (e.g., Hitzman, 2000). These parallels point

to a convergence of processes at certain critical points

in fluid evolution, including the partitioning of vola-

tiles between fluid and rock, i.e., devolatilisation in

skarn and secondary boiling in porphyry systems. We

point to the role played by such mechanisms in

enrichment of exotic trace elements, notably gold, in

all the above types of deposits.

In the Cu-Fe core, the extent of healing textures

and lack of significant overprinting of magnesian

skarn by calcic skarn, a feature often reported in

deposits elsewhere (Einaudi et al., 1981), indicates

to us that the volatile-depleted fluids remaining at the

end of plume collapse were redirected upwards.

Volatile extraction always enhances the flux (e.g.,

Dipple and Gerdes, 1998) and we see that even

though the Cu-Fe core has a considerable vertical

extent (120 m), the Simon Iuda orebody also features

calcic-skarn hosted ore, no less than twice this thick-

ness, situated directly above the magnesian core.

The Cu-Fe core represents only 1.3% of the total

tonnage of ore in the orefield. The remaining ore

comes from orebodies placed symmetrically on both

sides of Simon Iuda along the 10-km strike of the

orefield. We consider that a formational model for all

the other bodies must originate in the presumption that

lateral flow was initiated from the source at the time

when emerging fluids were in excess of fluids driven

into reaction by the plume updraft in Simon Iuda.

Downstream channelling through the package of

schists (f 150 to 200 m thick), placed between the

granodiorite roof and the lower part of limestone,

would also be enhanced by the difference in lithostatic

pressure between the point of emergence and higher

‘topographic’ levels represented by the limestone-

schist contact at the margins of the median part of

the orefield (Fig. 16). The ‘metasomatic front’ is

actually placed parallel to the flow at the limestone

side of this channel, similar to the way Yardley and

Lloyd (1995) considered a metasomatic side parallel

to the flow path at its edge in metamorphic terranes.

Nicolescu and Cornell (1999) proposed a formational

model for Ocna de Fier-Dognecea in which an early

metamorphic dewatering event was followed by later

metasomatism involving interaction between fluids

and limestone. Such a model advocates a ‘metasomat-

Fig. 16. S–N profile across the Ocna de Fier-Dognecea orefield showing the metal distribution and a schematic synthesis of the proposed genetic model. Vertical scale slightly

magnified. The proximal part of the deposit, centred upon the fluid source beneath Simon Iuda and the two distal extremities of the orefield at Dognecea and Paulus are indicated.

Orebody morphologies and geological relationships have been simplified from available maps and exploitation records. The upper contour of the granodiorite is documented in the

central part, but has been estimated for Dognecea and Paulus. The irregular contour at the limestone base is known across the whole orefield from exploitation galleries. The fluid flow

at the limestone base is documented. Insets show other zoning features in proximal (Elias, Reichenstein) and distal (Dognecea) fields. Onion-shapes characterize proximal bodies,

with massive ore concentrated in the innermost shell. In contrast, distal ores have low-grade amass at depth and high-grade chimneys at upper levels.

C.L.Ciobanu,N.J.

Cook/Ore

GeologyReview

s24(2004)315–370

363

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370364

ic front’ that advances parallel to the intrusion contact,

rather than parallel to flow along a channel as con-

sidered in this contribution. The model of Nicolescu

and Cornell (1999), however, cannot explain the

formation of the Cu-Fe core.

We believe that formation of each individual ore-

body resulted from skarn ‘fingering’ at the limestone

metasomatic side of the flow. Skarn ‘fingering’ is

another type of RIF instability at the skarn front with

great impact of metal deposition as Dipple and Gerdes

(1998) discussed in their modelling. The shapes of

some orebodies in the proximal setting are particularly

indicative of such ‘fingering’, especially Elias (Fig.

16, inset). Formation of each orebody is thus more or

less contemporaneous with formation of Fe zone in

Simon Iuda. The characteristic onion-shell structure

defined by the Fe-grades in each of these orebodies is

evidence that infiltration propagated from inside to

outside within each of them (Fig. 16, inset). More-

over, in the distal setting, as the infiltration rates slows

down, the mineralising style is changed from low-

grade at the base to rich-chimneys at the upper parts

(Fig. 16, inset).

In skarns, the prograde stage is the record of all the

sharp fronts of reactions (in the sense of Guy, 1993)

that were achieved at local equilibrium during fluid

advance into the limestone. The resulting sequence of

mineral assemblages forms the basis for orefield–

deposit zonation (Meinert, 1992, 1997). At Ocna de

Fier-Dognecea the prograde stage is developed in

response to intense flow-driven infiltration that con-

cludes with two sharp reaction fronts. Only one of

these represents a clear change in skarn type, from

magnesian to calcic, the other has a sulphidation

character, which we have interpreted from changes

in pyroxene compositions (i.e., >60% Joh +Hed). The

assemblages of the three skarn zones resulting from

the two reaction fronts consist of: FoFDi>90; And90-

70FDi90-70; (HedJoh)>60. The zonation imposed by

the fronts is also reflected by ore distribution in the

corresponding sequence: magnetite + Cu-Fe sul-

phides, Fe-oxides and Zn-Pb sulphides. Whereas the

Fe zone is present in all orebodies, the Cu-Fe zone is

restricted to deepest part of Simon Iuda. The Zn-Pb

zone is seen mainly, though with variable extent, at

the upper part of the each orebody. Therefore the Cu-

Fe zone forms a core in the central part of the orefield,

enveloped by the Fe and Zn-Pb zones.

Since all the prograde textures are part of the two

well-defined skarn zones (Fe and Zn-Pb), we point at

the critical role played by the interface between rates

of infiltration (the mechanisms that assists skarn

zonation; Guy, 1993) and diffusion rates, the mecha-

nism assisting precipitate banding–oscillatory zona-

tion. Such interfaces define the boundaries between

the patterned and non-patterned domains of skarn. In

infiltration-driven metasomatism the fluid flux is

controlled, among other factors, by decarbonation

reaction (Dipple and Gerdes, 1998). In such a model

only a steady state (stationary) flow coupled to

reaction, decarbonation in this case, promotes the

reactive infiltration at the skarn front. However, mul-

tiple steady states regimes can be induced from many

factors among which we consider that the slow-down

of infiltration as the fluids reach the outermost part of

the skarn system can be one. According to the theory

of self-organisation in geochemical systems (Ortoleva,

1994), the multiple steady state situation means that

there is an interval where any of the descriptive

variables in the system, e.g., concentration, diffusion,

infiltration, etc., will have multiple values across the

same interval of the system. For example Guy (1993)

considers that the sharp fronts of reactions that attract

formation of zones, represents one of the multiple

steady state situation in the skarn system.

In our case, the multiple steady state of decarbon-

ation reaction represents one of the ‘far from equilib-

rium’ states in which patterning operators (mass-

transport feedback mechanisms) can be readily acti-

vated. Liesegang banding and/or CPG are the two

patterning operators that could enhance some of the

small fluctuations, perpetually present in the steady

state regime, and amplify them to a macroscopic

amplitude. The resulting patterns, (Fig. 6c to k) are

the expression of these feedback mechanisms activat-

ed at the multiple steady state regime and they are

called ‘dissipative structures’. This is also the way the

system can evolve further, in our case at the end of the

Fe zone will move into the Zn-Pb zone, where, at the

end of this last zone, a new set of macroscopic

patterns (e.g., Fig. 6j) are realised.

Two distinct retrograde events are interpreted from

the range of reworking textures that overprint the

main skarn associations. The first, restricted to the

Cu-Fe core, is driven by sudden extraction of vola-

tiles from fluids (devolatilisation) following forma-

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 365

tion of magnesian skarn. The second has been recon-

structed from series of sequential textures seen at the

marginal parts of all orebodies and described under

‘carbofracturing’ above. This stage results from in-

teraction between limestone and fluids reaching the

edges of the skarn system. Contrary to the core where

a broad range of elements are involved in the

devolatilisation episode, the involvement of volatiles

during carbofracturing is due only to impact of erratic

CO2 release (i.e., decarbonation). Such changes, as

well as the drop in fO2 evidenced by pseudomorpho-

sis of hematite by magnetite, indicates the overall

reducing character of carbofracturing. It is not unre-

alistic to consider that breakdown and replacement of

skarn minerals by calcite and quartz, as well as by a

range of hydrated phases (chlorite, ferroactinolite,

talc, etc.), was associated with the carbofracturing

episode rather than with a later episode of alteration

like that which caused supergene enrichment in

Simon Iuda.

10. Assessment of skarn textures

Even though skarn features are attributes of vir-

tually all contact aureoles, proper skarns form only

under conditions of infiltration-driven metasomatism

(Korzhinskii, 1970; Guy, 1993; Dipple and Gerdes,

1998). As we have illustrated, there is a wide range

of patterns inherent to prograde development in a

skarn system, each of which is predictable from the

underlying theory. Unlike the isolated orbicules and

nodules known from numerous contact aureoles

(Joesten, 1991 and references therein), most of the

textures attributed to reaction-mass transport feed-

back mechanisms at Ocna de Fier-Dognecea (Table

7; Figs. 6 and 7) are located in the Fe and Zn-Pb

zones. This preference is not only because of the

close spatial relationship between the patterned and

non-patterned parts of both zones, but also because

the components involved in these textures are those

that define each zone. Rhythmically banded textures

in the Fe zone always involve the precipitation of

magnetite in a skarn–marble matrix of variable

composition. In one or the other more complex

textures, e.g., mossy, mottled, nodular and orbicular

(Fig. 6e to j), magnetite may, however, be absent. In

the Zn-Pb zone we note that sulphides are involved

in the orbicular (Fig. 6j) or spotted textures. This is

further evidence that the magnetite and sulphides

accompanied the skarn silicates in all expressions of

prograde skarn formation.

Following initial plume updraft, the path of the

decarbonation reaction controlled the motion of the

skarn front until, towards the end of the prograde

stage, a multiple steady state regime developed and

produced rhythmic patterns on all scales. During

formation of the Fe zone, and also the Zn-Pb zones,

the activation of powerful patterning operators repre-

sented by Liesegang banding alone, or coupled with

CPG (as seen from the textures above), show that the

front had the characteristics of an unstable coarsening

front of reaction.

In contrast to the prograde stage, fluids that are

driven back into the prograde skarn (back-flows)

produce retrograde events with an overprinting char-

acter. The back flow fluxes are either associated with

changes in porosity, as had earlier been the case in the

formation of forsterite at the first reaction front, and/

or addition of CO2 to the fluids (decarbonation), as is

the case during carbofracturing. Any of the shock-

induced textures we have presented is a valuable

indicator for peaks of overpressure affecting a skarn

assemblage. The impact of such transient overpres-

sures on fluids is particularly important because it

causes destabilisation of volatile species that have

potential as significant carriers of precious metals. As

we have shown, release of a suite of exotic trace

elements accompanies both retrograde events, of

which the Bi-Te-Au-Ag association is common to

both. The importance of shock-induced textures must

be emphasised in the context of Au enrichment in

skarns of all types, especially when the retrograde

fluids cross the main buffers in fO2– fS2 space.

Maldonite, a component of the trails shown in Fig.

9g, is a clear indication for activation of Bimelts as a

scavenger for Au (cf. Douglas et al., 2000), at temper-

atures above the melting point of Bi (271 jC). Such a

fact is concordant with the pronounced Au-Bi asso-

ciation widely described in Au-skarns elsewhere

(Meinert, 2000).

In contrast to the Cu-Fe core, the retrograde

carbofracturing in the Fe and Zn-Pb zones was the

latest event in the evolution of system and developed

more or less simultaneously throughout the orefield.

Unlike in shallow skarn environments (e.g., Meinert

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370366

et al., 2003), the intensity of the late healing event in

all three zones is consistent with the piercing, tiling

and brecciation failing to reach surface and entrain

meteoric waters. The retrograde features incorporate

elements of hydrothermal brecciation in sub-volcanic

environments (e.g., brecciation and fluidisation) and

recrystallisation in regional metamorphosed terranes

(e.g., deformation and annealing). Development of the

retrograde stage, which was intense and assisted by

sequential back-flow fluxes, was markedly con-

strained by the deep skarn setting. Unlike most

shallow skarns where collapse of the skarn trans-

gresses the contact aureole at upper levels (Meinert,

1992), brecciation processes at Ocna de Fier-Dogne-

cea are arrested in the collapse envelopes that enclose

each orebody.

Even though skarns are generally viewed as an

integral part of the magmatic-hydrothermal range of

deposits (e.g., Sillitoe, 1995), the evolution of fluids

driven by infiltration during metasomatism differs

fundamentally from the evolution of fluids in a

hydrothermal system. Reported skarn zonation (e.g.,

Meinert, 1997) is a direct consequence of metasoma-

tism driven by infiltration and requires that fluids

remain unchanged at source (e.g., Korzhinskii, 1970;

Guy, 1993). Such an assumption contrasts sharply

with a hydrothermal system in which alteration and/

or ore zonation is frequently explained by overlapping

fluxes of fluid or changes in the fluid(s) at source

(e.g., Hedenquist, 1995; Hedenquist et al., 2000). The

positioning of skarn systems in hydrothermal terms

has given the widely accepted impression that ore

mineral formation is an attribute of a post-skarn/

metasomatic stage, associated with the latest, hydro-

thermal and/or retrograde event and involving non-

magmatic fluids (e.g., Meinert et al., 1997, 2003).

Although certainly true for many shallow skarn sys-

tems, and/or those genetically linked with adjacent

porphyry (F epithermal) deposits (Einaudi et al.,

1981), buried skarn systems such as Ocna de Fier-

Dognecea or indeed tungsten skarns in a wide variety

of settings (Newberry, 1998) represent a substantially

different group of deposits.

Precipitation of sulphides, unlike magnetite,

implies mass loss in the fluids, a fact that cannot be

easily incorporated into Guy’s chromatographic ap-

proach. The intimate relationships between sulphides

and silicates (e.g., symplectitic, poikilitic or eutectoid

textures, etc.) hint at the fact that they are more or less

simultaneously formed and therefore have to be in-

corporated into any model addressing mutual changes

suffered by fluids and rocks during prograde skarn

formation. Ignoring them neglects the reality of the

zonation defined using metal distribution or grades in

skarns of all types (e.g., Meinert, 1997). We empha-

sise that comprehensive models for skarn formation

should not neglect such intimate, co-genetic relation-

ships between skarn and ore minerals, as has been

shown throughout this contribution, by Cook and

Ciobanu (2001) and indeed elsewhere (e.g., Mozgova

and Borodaev, 1995). Perhaps among the most sig-

nificant results that have emerged from recent advan-

ces in the understanding of skarn systems is the

recognition that development of porosity–permeabil-

ity during RIF at skarn fronts increases the potential

for ore deposition (e.g., Dipple and Gerdes, 1998).

Both mineralogical and petrographic characterisa-

tions are required to properly diagnose the anatomy of

any mineralising system. We have applied textural

analysis to mineral assemblages representative for the

entire skarn at Ocna de Fier-Dognecea. This is an

outstanding example of a skarn that displays pattern-

ing at all scales (from nanoscale to metres). Moreover,

the patterning phenomena involve most of skarn and

ore minerals present in the system. The deposit also

contains a wide range of superimposed textures that

reveal subsequent stages of skarn evolution. Recon-

struction of all sets of patterns in their geological

context, as presented above, substantiates an interpre-

tation for critical points in fluid evolution with impli-

cation for the sequence of skarn development. Only

the use of petrological data, with complimentary

textural analysis allows a comprehensive reconstruc-

tion of this skarn system in time and space. Although

we concede that our methodologies are not a substi-

tute for quantitative techniques for fluid study (e.g.,

isotope or fluid inclusions), we would however stress

that adequate attention to textures is a necessary

prerequisite in any attempt to reconstruct the evolu-

tion of a given skarn system.

Acknowledgements

We are grateful to the many people in Romania and

abroad, with whom we have enjoyed lively discus-

C.L. Ciobanu, N.J. Cook / Ore Geology Reviews 24 (2004) 315–370 367

sions and exchanged ideas on the geology of the Ocna

de Fier-Dognecea deposit. We are especially grateful

to Constantin Gruescu for access to his collection of

material from Ocna de Fier-Dognecea and for

permission to photograph specimens. The compre-

hensive comments of Khin Zaw, an anonymous

referee, and Brian Marshall, greatly improved the

manuscript. The senior author acknowledges a NATO

post-doctoral fellowship during which time this

manuscript was prepared. The support of the Geo-

logical Survey of Norway is greatly appreciated.

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