Thursday, November 18, 2010

ORE GEOLOGY4.Ores In Felsic Igneous Rocks

CLASSICAL MAGMATIC HYDL. DEPOSITS
(Porphyry Cu. Mo, Pegmatites, Vein-type ores, Skarns etc)

Origin relates to processes operating at the end-stage of felsic
magmatism

The effectiveness of these end-stage processes depends mainly on
(1)magma comp. (H2O % content, metal, S and Cl)
(2)Geologic environment including depth of magma emplacement
Hence, to understand the genesis of the above ones the first consideration
of end-stage processes of felsic magmatism is a must
Three basic processes, according the geological and geochemical
environmental factors (geologic-geotectonic settings, P, T and phase
assemblages) operate.
These are:
(a) ORTHOMAGMATIC: Melt ⇔Crustal equilibrium
Controlled by viscosity and density contrasts, include
(i) some important processes for generation of hydrous metal-bearing magmas form different sources rocks
(ii) magma emplacement and crystallisation

(b) HYDROTHERMAL equilibrium Crystal ⇔ fluid
The other end of the process continuum- involving aq. fluids and solid
phases
(wall rocks + hydrous minerals)- can be of many types. Broadly
involving fluids of
(i) magmatic origin and
(ii) other extraneous nature

(c) TRANSITIONAL:
Between the two extremes (orthomagmatic & hydrothermal processes)-
Upper limit:
arbitrarily set at the point of separation of a magmatic aqueous phase.
lower limit:
H2O saturated solidus of the magma

ORTHOMAGMATIC PROCESSES
A SOURCE ROCKS AND MAGMA GENERATION
IMORTANT STEPS IN THE CONC. OF ORE FORMING ELEMENTS
Partial melting of
(1) Mafic oceanic crust in the subduction
2) Mafic amphibolites of the lower continental crust (
high average conc. of base and precious metal → the melts get enriched in these elements)(3) Felsic metasediment rocks of the lower continental crust (the melts get
enriched in Sn, W, Be, Ta, Nb
). Prior to melting, H2O in the source rocks is
bound as (OH) in hydrous minerals (
at the great pressure in the lower
continental crust or on the subduction zone) – amphibole or biotite.
But
even in these rocks H2O (tot.) < 2 wt% → places a definite constrain on the
amount of H2O generated at a given depth (P) and T

From
the above the P-T projection the following inferences can be drawn:
(1) At a depth < 70km,
a non-porous mafic amphibolite begins to melt at 940-1040°c and
H2O content of the melt must exceed approx. 2.7%
(30% of the orig. rock can be melted for each % of H20 in the rock.)
(2) At a depth > 75 to 80km -
in the subduction zone -
amphibolite is not stable and melting can begin at a temperature of 660°C – However, at such low temperature for melting to occur ,
the melt must contain 27% H2O-
the amount of initial formed melt <5% for each % H2O in the origin rock
(3) Non-porous mus- bearing meta-sediment rocks begin to melt at a temperature of 670-720°C, at the ambient pressure of
the lower continental crust
→ the first formed melt must contain in excess of 8.4wt% H2O
→Hence only 10 −12 % of the rock can be melted for each % of H2O in the original rock.


Melting Relations: ORE GENETIC SIGNIFICANCE


(1)Partial melting of amphibolites or micaceous metasediment rocks (zero porosity), under high-P produces initial melts, with at least 2−7 wt% H2O, regardless of the amount of H2O originally bound in hyd. minerals

(2)The melts produced, range in composition from granitic to dioritic
(calc-alk.) - Representative of igneous rocks associated with   magmatic/hydrothermal deposits

(3)The amount of initial melt formed is directly proportional to the H2O
content of the initial rock
(geologically reasonable 1% H2O = wt% of the
melt =10−25%)

(a) Min. of 2.7 wt% H2O - Once the magma produced and emplaced in
shallow crustal levels
-will evolve a separated magmatic aq. phase (hydl.
fluid) upon cooling and crystallization

(b)It enhances the solubility of metal sulphide by an order of magnitudes,
compound to the same in anhydrous melts of same composition by the
reaction
2FeS(solid)+2H2O(melt)+SiO2(melt) = Fe2SiO4(melt)+ 2H2S(melt)
a hydrous magma with high S-capacity

(c) Leads to crystallization of Bt and /or Hbl up on cooling and
crystallization at a depth > 2 km
→ providing exploration guides to intrusives associated with hydrothermal ore forming process
Calc-alkaline melt comp. → As the magmatic /hydrothermal ore deposits
are associated with rocks of this compositional range-
the hydrothermal activity is associated with these compositional type of rocks- because hydrous minerals in the source rocks (mica or amph.) play a major role in the generation of hydrous calc. alkaline magmas
Amount of melt vs. %H2O of source rock (dependence)- provides a
mechanism for enrichment of partial melt with certain elements relative to their original conc.

(a) elements (e.g. Cu in FeS in amphibolites, minor
phase) which dissolve completely in the early formed melt: enrichment
factor is inversely proportional to XH2O in the rock
(Cu, en. fact. >5)
(b) On the other hand, elements which are present in major mineral ss., that co-exist with the early formed melt (e.g. Pb in K-feldspar)- enrichment
factors are dependent on the kdi values
amongst the co-existing phases;
Hence for Pb- little enrichment

PARTIAL MELTING → CRITICAL INITIAL MULTIPLICITY PROCESS- e.g. Cu

EMPLACEMENT OF MAGMA

Shallow depth of emplacement (generally <10 km,
commonly 1−2 km)
Geological evidence points towards a rather negligible role played by
gravitational settlingunlike mafic/ultramafic melts
Hydrous magmas of intermediate composition
are generally not in chemical equilibrium with more silicic and potassic wall rocks through which they pass or are emplaced (initial disequilibrium) - tendency to assimilate at marginal part enriching in SiO2, K and perhaps Sn– might be an important factor in porphyry Sn deposits, but not in other
magmatic hydrothermal systems

Critical factors:
• Nature of the source rocks
• Process of hydrous magma generation
Emplacement
• Separation (Evolution) of a separate H2O-rich volatile phase
     -controlled by H2O solubility in Silicate melts
      At P                 H2O Saturation             % crystallization
     2km                 2.7 to 3.0 wt% H2O         33%
     8 km                6.1 to 6.4 wt%H2O          69%
     18km               9 to 10 wt% H2O             80%


TRANSITIONAL PROCESSES
(MELT ⇔ FLUID EQUILIBRIUM. )
The H2O content (Wt %) of an uprising magma at any stage of crystallization (Wt) can be worked out by simple mass balance considerations as follows:
Wt (1- F)+ C(F) = Wto ⇒ Hence,
Wt (in wt%) = [Wto - C(F)] / (1- F)] x 100
where Wto is the wt. fraction of H2O in the completely liquid magma,
C is the wt. fraction of H2O after
a finite wt. fraction of the melt is crystallized (F)

The formation of a separate magmatic aq.(volatile) phase by retrograde
(2nd) boiling
marks the beginning of transitional processes (within the
condensed crystal - melt system = orthomagmatic regime)
Chemically, contrary to xl ⇔ melt equilibrium during the orthomagmatic
processes, transitional processes are dominated by melt⇔fluid (volatile) equilibrium

Physically, orthomagmatic processes are controlled by viscosity and
density contrasts between the melt and xl,
whereas transitional processes are dominated by volume changes accompanying the 2nd boiling reaction:
H2O saturated melt→ minerals + volatile phases

PHYSICAL TRANSITIONAL PROCESSES

• A body of hydrothermal magma emplaced in colder wall rocks (whether
shallow seated porphyry magma or deep-seated pegmatite) must lose heat
to the surroundings → crystallization proceed inward from the walls of the
magma chamber
. But due to low diffusivity of dissolved H2O in silicate
melts=H2O saturation starts at the margin
(H2O-satd. rind or carapace in a quasi-static condition.). The interior of magma body is isolated to transfer
matter (except H) =conditions essential for formation of pegmatites,
porphyry Cu/Mo systems and explosive volcanism
•As 2nd boiling operates inside the H2O saturated rind, the magma bodymust either expand or the int. press. must increase due the 2nd boiling
reaction,
which leads to volume expansion at all crustal levels (pressure)
•A first approx. the vol. expansion is directly proportional to the H2O
content of saturation
and inversely proportional to pressure. A body of
H2O saturated pegmatitic melt
at 2 kbar (6.4 wt% H2O)
will expand approx.11%  upon complete crystallization
whereas the same body saturated with H2O
at 5 kbar (10%H2O) will expand only 5%
→ explains why gem-bearing pegmatites (with void space, low-P) vs. non-gem bearing pegmatites (no void space, high-P) The total change in volume accompanying the 2nd boiling reaction (H2O satd. melt→ xls + aq. phase), ΔVr=f(P,T)
Due to
above volume expansion and as at shallow depths the wall rockshave high rigidity,
no plastic deformation is possible to accommodate strain
.

Hence, as cooling and crystallization proceed, pressure inside the
H2O-satd. carapace
must increase and theoretically this excess int. press.can reach several thousands of bars (≈ 3 to 4 kbar), whereas the tensile strength of the strongest wall rock is only few hundred bars
Moreover 2nd boiling leads to production of mechanical Energy
[(PΔVr )Ergs km-3 of magma) x10–23]
because vol. per unit mass of xl + vap >>vol.of an equal mass of H2O saturated melt
Volume expansion and the release of mech. energy leads to brittle
deformation and formation of fractures
– essential for localization of ores
in most magmatic hydl. systems, the fractures types are stock works, large
veins, larger caldera and collapse structure .

CHEMICAL (TRANSITIONAL) PROCESSES

Generation of the magmatic aq. phase is accompanied by partitioning of all elements in the system such that μi or fi of each species is same in the phases in equil.

Cl: The dissolved Cl (in the hyd. magma) gets strongly partitioned into the aq. phase because
(1) chloride minerals are not stable in magmas of calc-alk. comp.
(2) Cl forms Cl– complexes with H, Na, Ka, Ca, Mg and heavy
metals (Cu, Zn, Pb, An) in the aq. soln. at moderate P-T

F: F also forms stable F- complexes in mag. fluids, but high solubility of the element in silicate melts and high thermal stabilities of minerals such as fluorite, topaz, mica→ cause F to be partitioned largely into the condensed crystalline phases

S: S initially dissolved as HS− in the melt is strongly partitioned into the
aq.phase
, provided a sulfide mineral such as Po is not stable

CO2: CO2 is sparingly soluble in the felsic melts and gets strongly
partitioned into the aq. phase
and thus plays a minor chemical role at this
stage
Di values of the above volatile phases (between the aq.fluid and melt)
appear to be relatively T- independent and except for S are also P-
insensitive.
Because S originally occurs as HS– can exist both as H2S and
SO2 (in the aq.phase), DS= f(fH2O,fO2) due to the reaction
H2S +1.5O2 =SO2 +H2O
At a given fO2 - increasing fH2O (hence fH2) increases the H2S/SO2 in the aq.
phase- hence DS ( ΣSv / ΣSm ) decreases. On the contrary, increasing fO2 at
a given pressure increases the SO2 /H2O ratio; hence the DS (ΣSv / ΣSm)
increases → these phenomena, arising out of low solubility of SO2 in the
melt (compared to H2O)
– are of great significance to sulfide ore formation
– by providing mechanisms for generation of high conc. of both S and ore
metals (almost entirely as Cl–) in the magmatic aq.phase.
fO2: The fO2 in the magma prior to the onset of 2nd boiling is largely
determined by the Fe3+/Fe2+ ratio-dependent on the type of source rocks

I -type Magma:
The fO2 in felsic magmas, generated by partial melting of
metamorphosed igneous
rocks is generally higher than the FMQ buffer. Hence, fSO2/fH2S ratio (and mole ratio) of fluids in equilibrium
with these magmas are close to ≥1 (ranges 1 to 10)

S- type Magma:
fO2
in felsic magmas, generated by partial melting of
carbonaceous metasediments
is generally lower than that of FMQ
buffer and the fCO2/fCH4 ratio (and mole ratio) =1. Consequently, the
fSO2/ fH2S ratio << 0.01

Aq. fluids separated from the I-type melts tend to produce S-rich porphyry
Cu-Mo deposits
and
those separated from S- type melts produce S-poor Sn- W deposits.

Chloride
(1) Granitic melt :The major Cl– complexes in equil. with typical granitic
melt are NaCl, KCl ,HCl (NaCl +KCl ≈ 90%)
However, with xllization of phases like mus (containing K+ & OH-), the HCl and KCl contents decrease (HCl drastically decreases) and NaCl content
increases
to maintain the stoichiometry (NaCl / KCl)fl becomes greater than

Na /K ratio in the melt.

(2) In granodioritic melt: Cl– complexes with Ca (CaCl2) and especially Fe
(FeCl2 + FeCl3), MgCl2 is mino
r. The addition of CaCl2 and FeCl2 (+ FeCl3)
complexes does not affect the HCl content, for a total Cl- conc. or the
equality between (NaCl / KCl)aq and (Na/K)m, but ∑ NaCl + KCl is reduced
by 2 or more times the ∑Fe + Ca - Chloride can be complexed with Na, K,
Ca + Fe-more complex situation
(3) In granitic system, pptn. of K-mica causes NaCl / KCl)fl >1.
But in granodioritic system early xllization of Na-baring Hbl- the trend is in reverse direction than Granite irrespective of initial values (eg. NaCl / KCl << 1).
In presence of both Hbl and Bt, NaCl / KCl ≈ 1. - the high (KCl /lNaCl)fl phase assoc. with granodioritic melt (Hbl-bearing) provides an explanation for K-metasomatism in porphyry Cu systems. On the other hand, high conc. of
HCl in the aq. phase prior to appearance of hydrous minerals - occurrence
of topaz in Sn- greisen and pegmatites.
Also very high conc. of Fe in the aq. phase accounts for contact metasomatic skarn ores in carb. rocks and dominance of pyrite in porphyry Cu-Mo ores



ORE METALS
By analogy with Mn and Zn which are partitioned in favor of the aq.phase
by a factor of 2 (∑mCl-)2, it is expected that elements which are more chalcophile than Mn -- will be strongly partitioned into the aq. phase.
Therefore, aq. chloride conc. where fSO2 ≥ fH2S– metals will be strongly
partitioned

HYDROTHERMAL PROCESSES
With falling temp. or decrease in int- fluid press-transitional chem.
processes give way to those hydl. processes which are dominated by
crystal (mineral)- fluid equil.

The boundary between these two regimes is arbitrarily chosen as the H2O std. solidus of magma may be relatively sharp in some system ( sudden brittle fracturing in porphyry Cu dep) and gradational in some other . Eg. in pegmatites both regimes can co-exist in different parts of the system and communicate with each other through volatile phases → coexistence of two inter-communicating regimes is apparently essential for development of mineralogical zoning in pegmatites Rapid passage of magmatic aq. phase into hydl. regime in development of fracture system in porphyry Cu-Mo dep -→leads to conditions of gross disequilibrium between the hydl. soln. and cooler wall rocks. The extent of disequil. depends on the initial condition of equil. in magmatic system as
well as on the nature of the wall rocks and to the extent of P-T decrease

 Aq. chl. solutions from a high temperature magmatic source tend to be enriched in HCl and react with the feldspathic
wall rocks to produce Al-silicate (mainly andalusite and topaz) alteration
with or without Bt at high temperature and
muscovite (+sericitic+ phyllic ) altn. at low temperature.

Fluid that equilibrated initially with Hbl- bearing
magmas - KCl/NaCl ratio is more
-- when interact with non-carb. wall rocks
– fixation of K in Fel and Bt (K – alteration) by exchange of Na+ & Ca++
lowers the KCl/HCl ratio→enter into the stability field of mus

FLUIDS DERIVED FROM I- TYPE MAGMAS (high fO2)
 If during cooling the fluid interacts with Fe2+ -minerals then
→(SO2 +6FeO +H2O = H2S +3Fe2O3) → aH2S increases at the expense of aSO2
→resultant increase in aH2S causes pptn. of metal sulfide from metal -
chloride complexes of mostly Fe which in turn produces HCl due to the
reaction 4FeCl2 + 7H2S + H2SO4= 4FeS2 + 4H2O+ 8HCl

Production of HCl is further enhanced by pptn. of anhydrite:
CaCl2 +H2SO4= CaSO4 + 2HCl
(CaCl2 is either produced by earlier K⇔Ca reaction with wall rock plag or direct reaction with H2SO4 of Ca- bearing wall rock minerals)-- since the amount of HCl produced in the above reactions is
related (mSfl); the low temperature acid (HCl) alteration may be much
greater than that produced by the HCl in the original magmatic fluid

FLUIDS DERIVED FROM S-TYPE MAGMAS (low fO2)
May contain same amount of H2S as that in high fO2 fFLUIDS DERIVED FROM S-TYPE MAGMAS (low fO2)
May contain same amount of H2S as that in high fO2 fluids - but due to low
fO2 they contain less SO2; hence less ΣS.
Consequently they ppt. less sulfide (mainly Po) but largely oxides (cassiterite) upon cooling.
mCO2≈mCH4≈mH2S>> mSO2, as a result they are of very high mCH4/mSO2 ratio
and mCO2/mCH4 ratio remains const. during cooling and the fO2-T paths tend to lie near CO2/CH4=1 line
Corroborated by rich conc. of C- species (CO2+CH4) in fluid inclusion from
granite-related Sn deposits.
In feldspathic wall rocks the amount of HCl produced is controlled by
hydrolysis reactions involving silicate minerals – commonly yielding mus
or other aluminous minerals.
In carb. rocks, HCl content of the fluid is fixed at lower levels
as: CaCO3 +2HCl = CaCl2 +H2O+ CO2, consumption of HCl in turn leads to pptn. of sulfide minerals
(carb. replacement) by reactions such as: ZnCl2 + H2S=ZnS + 2HCl (skarns)
The above reactions ore commonly accompanied /preceded by other carb.
replacement reactions involving pptn. of Fe- bearing silicates (Gt, Px)
oxides ( Mt) -typical Skarn forming reactions with evolution of CO2 and
further inhibiting calc-silicate formation; such as CaMgSi2O6 + 2CO2
=CaMg(CO3)2 +2SiO2 → resulting in complete replacement of calcite by
dolomite and silica














 




 




ORE GEOLOGY3.Fluid Inclusions Study

Fluid Inclusions : Microgeological Systems
 types:
1)primary:
trapped during xenolith growth
2)Secondary:
formed after the xenolith growth, along healed cracks
3)pseudo-secondary:
occur along healed cracks and fractures that terminate at the grain boundaries → cracks formed during xenolith growth. PRACTICALLY FURNISH THE SAME DATA as primary inclusions.
Mechanisms of trapping:
Any process, which interferes with the growth of a perfect xenolith, may cause trapping of primary inclusions.
(1) A period of rapid crystal growth forming a porous dendritic layer
succeeded by a slower growth controlled xllization, thus covering solid
impervious layers and trapping many inclusions.
(2) If the material (nutrients) is supplied to the growing xenolith faces by a mass flow of fluid, large inclusion may be trapped as a result of temporary
starvation of centre of faces relative to faster growing edges
(having
easier access to the fluid) →the excess fluid is trapped.
(3) If some xenolith-growing blocks grow faster than others
→the surface becomes rough with many angular re-entrants,
which are filled in by the fluids
followed by covering during later growth periods yielding negative xenolith cavities.

ANY INCLUSION WHICH IS TRAPPED ALONG THE PRIMARY GROWTH
ZONE IN A MINERAL GRAIN IS PRIMARY. ISOLATED INCLUSIONS
FORMING A RANDOM 3-D NETWORK IN UNDEFORMED HOST MINERAL
GRAIN IS ALSO CONSIDERED AS PRIMARY

Trail
terminology:
main distinction is
made between
trans-granular (secondary),
inter-granular and
intra-granular
(pseudo- secondary)



Assumptions
(i)Homogeneous entrapment - proof: inclusions showing apparently same phase ratios
(ii) No change in cavity volume after entrapment
(iii) Inclusions behaved as thermodynamically closed systems
(iv) The origin of the inclusion is known

PHASE CHANGES SINCE TRAPPING

(A) SHRINKAGE:
leads to formation of the vapour bubble.
As the volume co-efficient of thermal expansion for most minerals is one to three order of magnitude lower than that for H2O - during cooling the cavity
(mineral) shrinks much less than the fluid inside.

When pressure in the inclusion drops below the total vapour pressure (P < Pv) of the multicomponent fluid (Hence Vfl <Vincl)
→ a bubble nucleates and grows in size. Hence, by reversal of this natural process in the lab. (heating), temperature (Th) can be determined.
Types of: L + V → L, L + V → V,
CRITICAL HOMOGENIZATION: by fading of the L- V miniscus
(B) IMMISCIBILITY:
There is no dividing line between the separation of a vapour bubble by fluid shrinkage and separation by immiscibility.
dH2O(pure H2O) =.0006 at 100°c,
but if the temperature is around 370°C then dH2O ≈ 0.32 grm.cm-3 (500 times greater) most gases present (CO2, CH4etc.) will preferentially enter the vapour phase dV > dL and the gas bubble will sink in the liquid.
If Xco2 is high – then CO2 phase will split into LCO2 and VCO2 and they will homogenize at any temperature below
31.1°c (Tcrit,CO2) down –16°c ( depending on dCO2).
(C) CRYSTALLIZATION ON THE WALLS:
during cooling the host mineral, if supersaturated should xllize.
Then theoretically speaking, it should redisolve on heating-- But seldom seen in inclusions from ore deposits.
(D) DAUGHTER MINERALS:
Fluid could become saturated wrt some
dissolved electrolytes
(salts) and they nucleate as daughter minerals
(NaCl, KCl, CaSO4 etc). These daughter minerals dissolve (Ts,NaCl etc.)
during heating – contrary to captive phases.
(E) METASTABILITY:
Inclusions are very small system even from the atomic viewpoint.
An inclusion of 10μ size containing 30% solution of NaCl may have only 1011-1012 molecules of NaCl.
inclusion containing Similarly, another 10 ppm PbS of the same size would contain around 25x106 molecules of PbS –→ results in metastability i.e., failure to nucleate new but stable phases.
Noticeable metastability during freezing the inclusion (failure to nucleate ice, NaCl .2H2O, CO2 5.75 H2O
etc).

MICROTHERMOMETRY
Most commonly used non-destructive analytical technique involving careful
observation of phase changes (e.g., ice melting, homogenization, solid dissolution etc) as a function of temperature (-195°C to 700°C) in individual inclusions, carried out in microscopic heating- freezing stages.The data obtained can be used with the help of experimental data inpertinent systems to constrain the chemical (salinity, gross chemistry) and physical (density, P, T) parameters of the fluid.
The constraints are in general, semi-quantitative for comparing complex
multi- component natural fluids to simplified experimental systems.

The H2O system
Many inclusions have comp. approaching pure H2O.
Hence, phase equil. of the most important compound on Earth is the logical place to start.
In this one component system, the solid- liquid- vapor triple point (+0.01°C an 0.006 bar) is invariant (change in P or T will reduce the no. of phases). As
dLH2O > dice (at the triple point) and
there are high- P polymorphs of ice →several unique features of low- T phase equil. of H2O → potential confusion in interpreting melting data.

Three possible cases can occur:
P-T diagram for pure H2O,
showing boiling curve (L-V equilibrium.) and isochores of various densities.
The inset shows an expanded view of the area near the triple point and melting
paths of two model inclusions.



i)
Relatively low- density H2O inclusion containing a vapor bubble
will freeze→ ice + V (incl. A) which will follow ice + vapor curve during warming until it melts and vanishes at +0.01°C.
(ii)
In somewhat denser inclusion (B), the vapor bubble eliminated during by
expansion of H2O. On warming, the inclusion will begin to melt at temp. <
0°C → final disappearance of ice at +0.01°C, where the bubble should re-nucleate.
(iii)
Still denser inclusion (not shown above), without a vapor bubble at room
temp. → freeze to a mixture of (ice I + ice III) → begin melting at -21°C
and finish melting at temp. below 0°C.

THE INEVITABLE PRESENCE OF OTHER SPECIES IN NATURAL H2O
COMPLICATE THE ABOVE INTERPRETATION

H2O- NaCl system
Most common salt system. Apart from ice, two new solid phases form.
They are: halite and hydrohalite (NaCl.H2O).
For all compositions < 61.9 % NaCl, at temp. < -21.2°C,
an incl. will contain (ice + HH+ V).
Aq. incl. irrespective of their comp. generally must be frozen to –30° to –40°C →instantaneously frozen to a disequilibrium mixture of ice + HH, which may begin to melt upon heating at the metastable eutectic at –28°C.
Let's assume that incl. A has been frozen to an equilibrium mixture of ice HH+ V. Upon heating to –21.2°C, the first melt forms.
The temp. remain fixed until either all the ice or all the HH melts.
Since the bulk comp. is to the left of the eutectic, HH will disappear.
Correct observation of first melting is critical for determining the pertinent incl.fluid comp. Upon further warming of incl. A, the liquid Increases (lever rule) until the ice disappears at a temp. that fixes its comp. For comp. <23.2 wt % NaCl, the final ice melting temp. (Tm) of the ice can be converted to salinity using eqn. (1)
wt % NaCl = (-1.78Tm)-[0.042(Tm)^2]-[0.00057(Tm)^3]....................(1)
For inclusion A (Tm= -10°C), has a salinity of 13.9 wt % NaCl. Note, however,
the same Tm of –10°C can also be observed for bulk comp. 23.2 and 26.3 wt
% NaCl, where HH rather than ice is the final phase to melt, for which
salinity can be calculated using eqn. (2)
wt % NaCl = 26.271 + (0.181 Tm) + [0.002 (Tm)^2].............(2)

Ice- or HH- melting (?)

Ice has a RI lower than H2O,
HH
is strongly birefringent and has a higher RI
Incl. containing > 26.3 wt % NaCl must have a halite daughter xl stable at
room temp. The bulk salinity of such incl. is calculated from the halite
dissolution temp, Ts,NaCl  using eqn. (3)
Wt % NaCl = 26.242 + (0.4928 Ts) + [1.42 (Ts)2]- [0.223 (Ts)3] + [0.04129  (Ts)4]            +[0.006295 (Ts)5]- [0.001967 (Ts)6]+ [0.00011112 (Ts)7](3)
where Ts = Ts,NaCl/ 100
Homogenization temperature (Th)



Influence of salinity on the boiling curve and slope of the
isochores in the H2O- NaCl system

In the above figure, for a given press. the temp. diff. (Tt- Th) increases as
the density decreases
as the isochores (lines of const. density/molar 10volume) flatten out. (Tt − Th) is referred as pressure correction.
Density can be calculated from the freezing data, but in the absence of correct press. estimate, Th is considered as MIN. TEMP. OF ENTRAPMENT (OR
FORMATION)
, unless boiling condition is established.

Since boiling in hydrothermal systems is essentially a near-isothermal decompression (or near-isobaric cooling) process,
inclusions showing both liquid- as well as vapor-state homogenization is indicative of entrapment from a boiling fluid


A MODEL BOILING FLUID ENTRAPMENT

Estimate of fluid density based on Th is a prerequisite for correct
estimation of Tt.
In real inclusions solutes such as NaCl, KCl, MgCl2 and
CaCl2 further complicate the P-T estimation. In the above figure, inclusion
A and B through trapped at the same press (1kb) and homogenizing at the same temperature. (Th=140°c) require grossly different pressure correction
as evident from the difference in T t(A) & T t(B)
Dissolution of daughter minerals
The daughter minerals (halite, sylvite) dissolve at different rates while
heating the inclusion (KCl faster).
For a saturated NaCl–H2O system, the
final solution temperature of halite (TS,NaCl) is directly proportional to the wt% NaCl in the solution.
When more than one daughter phase is present
(I)Correct identification of the phase and
(2) Knowledge of the solubility of each in the mixed salt system

fig:
Dissolution of halite and sylvite as a function of temperature (Ts)

Freezing
All phase changes occurring below room temperature. Freezing is wholly
complimentary to heating and each inclusion should be frozen and heated to complete the microthermometric runs.
Data obtained on freezing
refer to
gross fluid composition and density.
For multicomponent fluids
the normal sequence during freezing is L+V⇔ S+L+V ⇔ S+V.
The presence of additional component extends the temperature interval over
which solidification (and melting) takes place. The melting temperature is a
direct function of composition–thus fluid composition can be determined
provided appropriate experimental data is available.
Inherent problem–
Reluctance to freeze due to metastability, the extent of which is inversely
proportional to inclusion size.

Freezing point is depressed in different binary H2O- salt systems

Salinity is measured by measuring the final ice–melting temperature (Tm)
on re-heating the completely frozen inclusion.


Since the fluid comp. is difficult to establish, Tm is conventionally reported
as wt % NaCl eq. (The amount of NaCl which would produce an equivalent
lowering) with an assumption that in a mixed Ca-Na-K-Mg chloride solution
the error in estimation of salinity using freezing pt. dep. method of NaCl eq.
is <5% when studying salt–H2O systems the temperature of first melting
(Tfm)
can be determined, if proper care is taken to observe the temperature
of first appearance of liquid.
Tfm should correspond to the eutectic
temperature from which, the gross fluid comp. can be determined.

 composition                       Eutectic Temp( degree C)
H2O- NaCl- CaCl2- MgCl2   -57
H2O- NaCl- CaCl2               -52
H2O- CaCl2                         -49.5
H2O- FeCl2                         -35
H2O- MgCl2                       -33.6
H2O- NaCl- KCl                  -23
H2O- NaCl                          -21.2
H2O- KCl                            -10.6


NON-AQUEOUS CO2 (± CH4) SYSTEMS

The CO2 SYSTEM
Carbonic inclusions are present in variety of hydrothermal ores and pegmatites. Commonly they contain CH4 and rarely H2S.
Hence phase equilibria in the pertinent systems are prerequisites for interpretation of microthermometric data of these incl.

Sequence of phase changes in a pure CO2 inclusion during freezing- heating run finally homogenizing as liquid
CO2 (LCO2 + VCO2 →LCO2)
Different Stages
at 20 degree C               L+V (2x size). Initial stage
at  -100 degree C            Solid CO2+ V (.0001x size)
at  -56.6 degree C           L+V (4x size)
at 27 degreeC                 L+V(x size)
at 29 degree C                L CO2            Final stage


Density and homogenization temp. (Th,CO2) relationship in carbonic incl. Density is calculated from the observed CO2 homogenization temperature, depending on the nature of homogenization


 FIG:Phase relations in the CO2- CH4 system at sub- ambient temperature. The triple point of CO2 is lowered with increasing CH4 (Tm,CO2 in equilibria with LCO2 and VCO2 is depressed)







 Fig: Large primary incl. in quartz containing a vapor bubble, large cubic halite
crystal and two unidentified daughter (or captive) phases (a and b)







P-T Estimation by intersection of data obtained from different inclusions study


Coeval and cogenetic aqueous and carbonic fluids



Fluid Evolution



Example I
Greenstone- hosted Au- mineralization at Larafella, Burkino
Faso, West Africa
The lode Au mineralization is associated with low- saline CO2- rich
inclusion


Example II

alteration zone
At Hutti, the mineralized quartz veins are associated with
low saline aqueous inclusions (H2O+ Salt) ,
carbonic inclusions (CO2 ±CH4) and
aqueous- carbonic (H2O+ CO2 ± CH4 + NaCl) inclusions








Wednesday, November 17, 2010

ORE GEOLOGY2.Generalized Classification Scheme For Ores In Ultramafic-Mafic Rocks

Mafic and Ultramafic Rock Association
1) Crystal Fractionation
  
Dimond in kimberlite
   Cr ores (Layered and Alpine types)
   Fe-Ti-Oxides in gabbro- anorthosite assoc.

2) Carbonate Association
  
REE, Nb, P, Sr, Ba Zr and sometimesCu

3) Liquid Fractionation

Oxide liquid Immiscibility. (Mt+apatite- Kiruna type)
Sulfide liquid Immiscibility .- Cu+Ni±Co± PGE ores

Ores of UM-M association include those formed by
crystal and liquid fractionations.
Carbonatite association is included here because generation of carbonatitic melt can be either due to
(i) liq. Immiscibility from phonolitic/ nephelinitic/ kimberlitic melt
(ii) by direct melting fertile mantle peridotite at P > 21 kb

unifying factors for the entire spectrum of ore deposits in UM-M settings
(i) highly selective and specific ore- rock association
(ii) a very exclusive suit of elements form their ore deposits in such rocks (whose average crustal concentration in the barren UM-M rocks are order-of-magnitude higher than in any other barren rock)
(iii) ore bodies are physically confined within the hosts
(iv) again barring exceptions, pervasive wall rock alteration is strikingly absent.

ORIGIN OF ESSENTIALLY MONOMINERALIC CHROMITE LAYERS 
The shape of the ol-chr cotectic curve is such that there are
TWO POSSIBLE GEOLOGIC SITUATIONS :

(A)
Common situation: the intrusion contains liquids that has differentiated to
point ‘f’, where opx is on the liquidus and some new primitive (mafic) liquid (comp. a) is introduced.
(1) Initially the primitive liq. with high liquidus temp. crystallize along the olivene-chr cotectic (a→b), producing peridotite. But with progressive mixing, in stead it will follow the path b→f towards comp. ‘c’. Because comp. ‘c’ extends across the edge of the chr- field, continued crystallization should for some time produce only CHROMITEdriving the melt comp. back to the opx field at ‘d’
⇒ opx starts forming.
The whole sequence is
peridotite → chromite→orthopyroxenite
as seen in the Muskox intrusive, NWT, Canada.
(2) Alternatively, if the volume of the fractionated liquid (‘f’) >>> volume of new addition (‘a’) no peridotite will form
and the sequence will be
orthopyroxenite → chromite as observed in the Bushveld Ign. Complex, SA.
(3) Alternatively, if addition of fresh liq. (‘a’) is quite large ⇒ hybrid liq.
Formation and the bulk comp. will return to the ol- chr cotectic after chromite
crystallization leading to a sequence of
peridotite →chromite → peridotite→orthopyroxenite as seen in the cyclic units of Stillwater.

B)
The differentiated liq. Is still on the ol- chr cotectic at a point just above the chr - control line.
The mutual solution (and mixing) of this liq. with a primitive liquid (‘a’) ⇒ yield liquid compositions in the chromite field leading to
formation of peridotite →chromite as observed in Great dyke, Zimbabwe.

TWO POSSIBLE CRYSTALLIZATION SEQUENCES
(A)
Crystallization from a melt --->
sulfide liq.separates out with continued crystallization ofsilicates.
If some trace amount (say about 400ppm) of Ni was present in the initial melt (pt.1), then that would be greatly enriched in the final crystallization product at Eutectic point,
thus giving a magmatic Ni-sulfide deposit, formed due to liquid immiscibility.

(B)
On the other hand, crystallization of a liq.which has
higher Fe3+/ Fe2+ ratio (because of relatively high fO2, =>(NO LIQUID  IMMISCIBILITY).
Mt and silicate crystallize together along the cotectic and
although crystallization stops at Eutectic and a sulfide deposit of essentially
pyrrhotite- magnetite mineralogy will form-- devoid of Ni,
due its early strong partioning in to crystallizing olivine and other
early formed silicates along the cotectic
.
These Ni- olivines serve as protore which on weathering forms Ni- laterites.

Melting relations in the Fe- S- O system

The ternary eutectic (Fe- Wu- Po) occurs at 915 ± 2°C with a ternary liq.
comp. of 68.2% Fe, 24.3%S and 7.5% O
the solidus rises steeply away from the reaction point towards decreasing Fe content.
A mixture of Po with 62.5% Fe and Mt would start melting at 1010°C.
Then the solidus flattens out as Po becomes less Fe- rich in equilibrium with Mt.

Effect of P
pressure has either no or slight effect on the eutectic temp. but causes it to move to more Fe- rich comp.

Effect of other component
 substitution of 2% Cu on lowers the solidus temp by 15 to 20°C

Variation in fO2, fS2 and aFeO

The comp. of a Fe- sulfide- oxide melt at a given temp. is a function of
fO2 and fS2. .
With known value of the activity/fugacity of Fe and O2 (aFe and fO2),
aFeOmelt has been calculated,
using the Gibbs- Duhem eqn. (∑(Xi).dlog(ai) =0),
for the reaction Fe + 1/2O2 = FeO(melt).

Application of Fe- S- O system to natural ore magmassome Sulfide droplets are formed in ore magmas.
aFeO of the magma is controlled by the Fe3+/ Fe2+ ratio (a function of fO2)
at equilibrium, the same value of aFeO and fO2 must apply to the
sulfide melt.


CRYSTALLIZATION OF SULFIDE ORES
The temp. of beginning of crystallization and final solidification of a
sulfide ore are important when considering how the sulfides will
(a) move away from their host intrusion as an ore magma and
(b) whether they are likely to be mobilized as a liquid during high grade
metamorphism.

most ores start crystallizing in the range of 1160° – 1120°C.
Since ore magmas, like silicate melts, can probably move as a mixture of crystals
+ liquid,
the solidus of the ore magma provides the minimum temp. of its mobilization in partly liq. state.
the solidus temp. varies with the Fe content in Po forming the ore deposit.
Almost all the ores are made up of Po containing 62.5 to 60.5 wt% Fe.
Then one would predict that the solidus temp. would be between 1010 and 1050°C.


Sulfur Solubility in silicate melts

Knowledge of S- solubility in mafic and u. mafic silicate melts is
important in understanding
1) how magmatic sulfide deposits form and
2) evaluating the potential ign. bodies as host for ore deposits of
this type.
At low fO2 (< 10-6 atm.) and at temp. of about 1400- 1500°C,
sulfur dissolves primarily as ‘sulfide’ and
the function can be termed as“S- capacity” of the silicate melt (Cs)
The entry of sulfur into silicate melts is governed by the simple
exchange eqn.
[O]melt + 1/2S2 = [S]melt + 1/2O2
oxygen displaced by the above reaction is very low and constant.
Cs = Sm[fO2/fS2]^1/2
where Sm is the sulfur content of the melt.
Cs increases with
 
increasing temp. 
 increasing FeO and MgO contents
 decreasing SiO2 and Al2O3 contents.
But  Cs is very high for FeO-SiO2 system at any XSiO2 value compared to MgO- SiO2 and CaO- SiO2 system 
=>sulfur can displace the oxygens bonded to Fe2+ at ease.

Cs merely relates to the amount of sulfur that will dissolve in a given melt in response to imposed fS2 and fO2.
Maclean (1969), from his studies in the system Fe- S- O- SiO2
found that the S- content of a silicate melt in equilm. with the S-rich liq.
decreases with increasing O- content.
This S- content is henceforth referred to as the ‘sulfur content at sulfide saturation’ (SCSS).
Maclean:
sulfur dissolves in silicate melt by displacing O-bonded to Fe2+ and that increasing oxygen results in an increase in Fe3+ at the expense of Fe2+ in the melt.
FeO(melt) + 1/2S2 = FeS(melt) + 1/2O2
K = [aFes(fO2)^1/2]/[aFeO(fS2)^1/2]
⇒ aFes
= γFeS NFeS = K.aFeO.[fS2 / fO2]^1/2
logNFeS = 1/2 logfS2 + [logK + log aFeO - 1/2 log fO2 - log γFeS]
1)If for small changes in NFeS, γFeS is assumed to remain constant
2)if the amount of FeS formed by the above reaction is sufficiently small then FeS formation has no appreciable effect on aFes,
then it follows that
NFeS vs logfS2 => st. line with slope =1/2 , at const fO2.

Effect of Temperature:

 At const. fO2 and fS2, the dissolved sulfur content increases by a factor of 8.5 times/ 100°C at 1000°C and
3 times/ 100°C at 1400°C.

Effect of Pressure:
(1) the miscibility gap (between silicate and sulfide liquids) expands with increasing pressure.
(2)  SCSS increases with temp. and decreases with press.
=>when magma rises to the surface, its ability to dissolve sulfur
increases, hence it is unlikely to approach saturation with sulfides.

Melt- Melt Equilibrium - partitioning of chalcophile elements
between sulfide and silicate melts: implication on composition of magmatic Ni- Cu sulfide ores


Nernst partitioning coefficient of a minor metal between the silicate
and sulfide melts
Di(SL/SM)= wt% ‘i’ in SL / wt% ‘i’ in SM
SL: sulfide liquid and SM: silicate melt
Partitioning of Ni between sulfide and silicate melts: theoretical basis
1) Fe, Ni, Co, Cu etc are bonded to oxygen  in Silicate Melt
2)Fe, Ni, Co, Cu etc are bonded to sulfur in Sulfide Liquid

NiO(SM) + 1/2 S2 = NiS(SL) + 1/2 O2
 KA = (aNiS)SL / (aNiO)SM x [fO2 / fS2]^1/2

= (γNiS / γNiO) x (NNiS / NNiO) x [fO2 / fS2]1/2
Di = f (fO2, fS2, T, P and comp. of the two phases).

minor element eqlm
.--combined with FeO- FeS eqlm.

NiO(SM) + FeS(SL) = NiS(SL) + FeO(SM)..................(Z)
 KB = {(aNiS) / (aNiO) }.(aFeO /aNiO)
⇒    KB= (γNiS / γNiO).(γFeO /γFeS ).(NNiS / NNiO).(NFeO / NFeS)
KB ≠ f(fS2, fO2),
(i) liquids of same composition as mss (Fe1-xS - Ni1-xS),
γNiS and γFeS will have similar values (γNiS/γFeS =1), although both the γi terms decrease with increasing fS2
(ii) fO2 can affect NFeO by changing the Fe3+/ Fe2+ ratio in
the magma, but if fO2 ≤ 10-8 atm, the effect in basaltic magma is
small.
(iii) fO2 variation also affects the O- contents of sulfide- oxide liquids- hence the a-X relations within them. But provided that the fO2/fS2 ratio (and hence the O/O+S ratio in the sulfide- oxide liquids) remains approx. constant leading to negligible effect in metal partitioning.

Ni and Ni-Cu sulfide deposits are spatially associated with ultramafic and mafic igneous rocks respectively. Ore composition- rock composition relationship

Cu/(Cu+Ni) ratios in natural sulfide deposits show a general increasing trend with decreasing basicity of host igneous rocks

Ni-Cu ores formed are formed by segregation of sulfide droplets fromthe host silicate melts and their concentration by gravitational settling.

Two factors governing composition of these ores
(i) composition of host magma at the timing of liquid immiscibility
(ii) distribution coefficients of Fe, Ni, Cu and Co between the sulfide and silicate melts.


Much of the compositional changes in silicate liquid is due to crystallization and removal of olivine, causing depletion of Ni in the residual liquid (since Ni has a strong preference to olivine relative to silicate liquid.)
Cu on the other hand tends to get enriched in the residual magma, because it lacks any preference to early formed Fe-Mg silicate ⇒ Ni/Cu ratio decreases rapidly during fractional crystallization of u-mafic and mafic magmas

Ni (460) behaves relatively more sulfophile than Cu (243) in andesitic melt.
In olivine basaltic melt the trend seems to have reversed
Since Fe2+ is the principal cation for the magmatic sulfide deposits,
the equilibrium relation between the SL and SM can be written as the
exchange reaction
(FeS)SL + (MeO)SM = (MeS)SL + (FeO)SM
The equilibrium const. (K) for the above reaction can be expressed in
terms of the KD values as shown in Eqn (z)
Where K is the equilibrium constant of the above melt-melt exchangereaction and is independent on the melt compositions (SL and SM) at
a given temperature.






ORE GEOLOGY1.Phase Equilibria In Ore-Bearing Systems

 aim 
to simulate the natural ore-forming environments in the laboratory.
Problems
plague of metastability’
(a) crystal-chemical complexities
extensive substitution (single, coupled, multilateral)+ dominantly covalent (and metallic) bonding => extremely distorted coordination polyhedra and close metal-metal distances in many structures promote non-stoichiometry, charge imbalance and wide solid solution range

(b) faster reaction kinetics and consequent non-quenchability
Phases such as sphalerite, arsenopyrite, tetrahedrite,
pyrite, hematite, magnetite, wolframite have lower metallic characteristics (lowest degree of metallic bonding expressed as low frequency of short M1-M2 bonds) are most useful
On the contrary, Cu-Fe-sulfides and Cu-sulfides take little time to
reequilibrate  Hence, barring the above refractory minerals, laboratory based
phase equilibrium studies have very little application as far as extrapolation to natural ores is concerned.

(c) dominant control of fS2 on phase assemblages
Standard state of sulfur
The standard state of sulfur is conveniently chosen as ideal diatomic sulfur gas, S2, at one bar fugacity and the temperature of interest. This state is used in spite of the fact that the same is not physically attainable due to condensation of solid or liquid sulfur below 614°C. Therefore fS2 is numerically equal to PS2 (and aS2) in atmosphere. The choice of this standard state is dictated because otherwise the sulfidation curves will otherwise bend at the melting, transition and boiling point of the yellow colored orthorhombic sulfur, stable at the normal P-T condition.
Parameters
P,T,X,fugacity,activity
Representation with the help of different diagrams
P-T,
T-X,
T-f(S2),
T-f(O2),
f(S2)-f(O2).
BASIC IDEAS USED
Application of laboratory data to natural assemblages
for quantitative interpretation in the form of phase diagrams requires recognition of
independent variables. The way to achieve this goal is through the Gibbs phase rule, a
fundamental statement in chemical thermodynamics, relating the number of stable phases
(p) in an equilibrating assemblages, to the number of independent components (C) and
the number of independent degrees of freedom (f).
Derivation: 

 If we represent the composition of all the phases in terms of Xi (X1, X2, X3...
etc), then there will be an Xi term of every component in every phase. Hence for ‘p’
phases there will be ‘cp’ compositional variables. In general, there will be two additional physical variables (P and T). Hence the total number of variables (unknowns) = cp + 2.
Since the sum of Xi in each phase = 1.0, then for ‘p’ phases, there will be p-number of c equations like ∑ X i = 1 .
Transfer of chemical components takes place in the decreasing direction of chemical potential (μi), implying that at equilibrium, μI of each component is the same in every phase where it appears. Hence for component ‘i’
μ iα = μ iβ = ............μ ip
(1)
For p-1 pairs, there will be c(p-1) number of the above equations making the total number
of equations = p+ c(p-1).
Now from the simple principle of linear algebra, the variance of the system of equations
= number of variables – number of equations
Hence f = cp + 2 - p-cp + c
f=c–p+2
It must be noted that the above phase rule is predicted at equilibrium with uniform P, T and μI throughout and that the identity of each variable is lost in the process of account, which can not be recovered by manipulation of the equations.
T
heoretically such diagrams are polybaric, but
because most solids and liquids are relatively incompressible, the P variation has little
effect on phase relation and the phase rule becomes f = c –p +1.
P – T diagram

These diagrams are drawn on the basis of Clausius- Clapeyron’s eqn. Given by
dP/dT= ΔS/ΔV=ΔH/TΔV
derived from dG = ΔVdP – ΔSdT = 0 and ΔG = ΔH – TΔS = 0.
fS2 – T diagrams
2MxSy + S2 = 2MxSy+1
ΔG° = –RTln K = RTln fS2 = 2.303RT log fS2
ΔH° – TΔS° = 2.303RT log fS2
log fS2 =ΔH0/2.303RT - ΔS0/2.303R
differntiating above eqn we can plot the diagram.
fS2–fO2 diagrams

Fe3O4 + 3S2 = 3FeS2 + 2O2
2Fe2O3 + 4S2 = 4FeS2 + 3O2
K
can be calculated from these eqns and used in eqn below
ΔG o = − RT ln K − ( P − 1) ΔVS


Eh-pH diagrams
Eh = E° +0.059(log K)/n



DIFFERENT SYSTEMS:
 


The Fe-S system
This is one of the most studied sulfide system, both in mineralogy and metallurgy. Pyrite and pyrrhotite are the only two sulfide minerals to be called rock-forming minerals.
The Fe-S system is also an important bounding binary in many multi-component systems such as Fe-Zn-S, Fe—Ni-S, Cu-Fe-S, Fe-As-S, Cu-Fe-Zn-S etc.).
The phase relation above 400°C is clear and straight forward The central portion of the system is dominated by the large, high- temperature pyrrhotite solid solution field constitutes extreme solid solution from stoichiometric FeS toward more S- rich compositions. This high-T form with hexagonal NiAs-type structure accommodates solid solution by random vacancies on the Fe sites within the lattice. Hence, the composition of high temperature pyrrhotites, except for FeS, is best given as Fe1-XS. The maximum thermal stability of the pyrrhotite solid solution at pressure = 1 bar is 1192°C, above which it melts incongruently. The eutectic between pyrrhotite and Fe is at 988°C, 56% Fe. Brett and Bell (1969) found that the eutectic temperature is P-insensitive up to 30 kbar but then rises rapidly, reaching 1160°C at 100 kbar.
 S-poor portionunimportant for
terrestrial ore forming processes, do apply directly to lunar and meteoritic samples and iron smelting
On the S-rich portion
the minimum temperature of existence of a sulfide liquid is 1088°C; above this temperature, there is a broad field of
liquid immiscibility comprising of a sulfide-rich and a sulfur-rich liquids.
 

The common occurrence of pyrite + pyrrhotite along with temperature dependence of pyrrhotite composition in the above assemblage, led Arnold (1962) to propose the pyrrhotite geothermometers.
utility of the pyrrhotite thermometry is only restricted
to samples where rapid equilibration between pyrite and pyrrhotites (eg. in basalts) takes place.
Otherwise, during slow cooling, the pyrrhotite composition slides down the solvus to a more Fe-rich composition.
A further complication in most ore deposits is the inversion to monoclinic pyrrhotite either by cooling below 254°C or as a low-temperature oxidation product.
X FeS in Po= n FeS/(n FeS + n S 2)
using the Gibbs-Duhem equation for the FeS-S2
binary, which can be written as
XS2 dlogfS2 + XFeS dlogaFeS = 0from this aFeS can be calculated.



The Fe-Zn-S SystemIn addition to phases encountered in the Fe-S systems, only new phases that appear in this ternary are sphalerite and wurtzite polytypes.
Sphalerite has a cubic closed packed (ccp) structure in which every alternate tetrahedral site is empty. Iron, along with Mn and Cd enters the structure by substituting for Zn and results in an increase in cell volume. Because of its refractory nature and wide variation in Fe content, sphalerite has been tried as a geochemical sensor for ore forming environment.
At 742°C,
in the T - X section of Zn-Fe-S system,
pyrite incongruently breaks down to Po + S- rich liquid
From this Point (I), three invariant curves emanate:
(1) Sp + Po + S
(2) Sp + Py + Po
(3) Sp + Py + S

Sp + Py + HPo solvus remains vertical below 550°C at 20.7 ± 0.6 mole % FeS. Hence, for the major portion of geological temperature range this assemblage can not be used as a geothermometer.

Sp Geobarrometer : sphalerite composition in the assemblage Sph + Py + Hpo furnishes excellent pressure values that compare reasonably well with those estimated from silicate barometers.

VFeS is large, compared to VZnS
--increase in cell edge of sphalerite with Fe content .
In a system where aFeS is buffered by Py + Hpo, sphalerite should become less Fe-rich with increasing pressure.
this geobarometer is T-independent within much of the
geologically important P-T ranges.
It is important to note that such T- insensitiveness increases with increasing pressure implying that sphalerite barometry is works better in situations when the ore are metamorphosed to relatively higher grades (P ≥5 kbar).
  
for successful application of this barometer certain prerequisites
1) coexistence of sphalerite with both the Fe-sulfides(Po,Py) to
ensure the aFeS- buffered condition
2)Additionally, sphalerite should be analyzed in grains without any chalcopyrite or pyrrhotite blebs in order to make certain preservation of its pristine P-dependent composition.

Assemblage Sp + Py
 fS2 – T relationship can be calculated using
FeS (in Sp) + 0.5 S2 = FeS2
considering a value of 2.3 for γ FeS
log X FeS in Sp = 6.65 –7340/T – 1/2 log fS2

Assemblage Sp + Tr + Fe
applications are only limited to meteorites
Sphalerite composition, buffered by troilite (Tr) and Fe (2Fe + S2 = 2FeS)
fixes aFeS at unity for all temperature. The effect of pressure on Sp + Tr solves is large,
The thermodynamic deduction is given below
X FeS in Sp= a FeS / γ FeS
differentiating this wrt P,
(XFeS)P2−(XFeS)P1 in Sp=−[(VolumeSp−VolumeTr)(XFeS)](P2− P1)/RT


 The Cu-Fe-S system
The phases occurring in the system Cu-Fe-S are found in many geological environments and in lunar and meteoritic materials.
Phase relation is not properly understood due to, large number of phases, extensive solid solution, non-stoichiometry, non-quenchability and metastability.
Phase equilibria at 400°C and above are dominated by three solid solutions (1) chalcocite - digenite - bornite
(2) (Iss)
(3) pyrrhotite solid solution

absence of complete solid solution between Cc-Di-Bn and the Iss, in spite of their similarity of structure and cell size,suggests differences in the type of bonding.
The Iss has a cell edge of 5.4Å, sphalerite-type, fcc unit cell  and
chalcopyrite cell is same excepting doubling in the C-axis direction.

1)On cooling below 557°C,
chalcopyrite appears as an ordered (tetragonal) phase.
Above this temp -transforms to a cubic Iss phase (and small amount of pyrite).
2)The orthorhombic polymorph of cubanite (CuFe2S3) is only stable up to ≈210°C  above which it inverts to cubic Iss.
3)Natural intergrowths of chalcopyrite-bornite, chalcopyrite-cubanite;
and talnakhite (Cu9Fe8S16), mooihoekite (Cu9Fe9S16) and haycockite (Cu4Fe5S8) are likely formed as decomposition products of initially deposited high temperature Iss.


The Fe–Ni–S system
1)Ni-bearing magmatic ore deposits formed by liquid immiscibility. 2)pentlandite,(Ni,Fe)9S8 occurs in some hydrothermal ore associations.


At 1000°C, the Fe–Ni–S system is dominated by
1)a large liquid field
2)a solid solution field extending from pyrrhotite (Fe1-xS) towards nickel monosulfide (Ni1-xS).

At temperatures <992°C, the monosulfide solid solution (mss) becomes complete and dominate the central portion of the system

At 650°C,  the mss coexists
with (Fe,Ni)S2 or (Ni,Fe)S2 on the S-rich side
or
(Ni,Fe)3±xS2 or γ– or α–(Fe,Ni)on the metal rich side.

On further cooling, the mss field narrows down

below 610°C
pentlandite appears on the metal rich side
violarite on the S-rich side.

Pentlandite is therefore entirely a subsolidus phase, formed by exsolution of the mss.

at 400°C
disulfides (Ni7S6 and Ni3S2), that have limited solid solubility with various (Ni,Fe) alloys

below 300°C
The mss begins to unmix

Below 212°± 13°C,
the pentlandite-pyrite tie line is established.

near 200°C
mss is shown breaking into three segments.

at 25°C from the study of natural assemblages.
Although the pentlandite-violarite natural assemblage is observed, it is generally metastable. Hence, the preferred tie lines are shown by joining millerite with pentlandite and pyrite.

At 600°C temperature,
diffusion rates
will be too high but
the driving force for nucleation that depends on the degree of undercooling below the solvus will be small
Slow cooling
Fast diffusion +heterogeneous nucleation will lead to the formation
of granular pentlandite. 
rapid cooling
 a greater degree of undercooling is necessary to cause unmixing. Here,
homogeneous nucleation may lead to oriented flames.

Fe-As-S System
only one ternary phase, arsenopyrite, FeAsS.
Important factors to study :
* compositional variation of arsenopyrite in terms of
 As/S ratio
on the FeS2 (pyrite) –FeAs2 (löllingite) join, and
  its dependence on formation (or equilibration) temperatures.

In the high temperature portion,
arsenopyrite coexists with pyrrhotite, löllingite, As, or an (As,S) liquid in assemblage 1 through 4.

Between 491° and 363°C,
arsenopyrite can coexist with pyrite, pyrrhotite, löllingite, As, or (As,S) liquid in the assemblages 5 through 8.

**
Since arsenopyrite composition is a function of both temperature and fS2, the same can be used a geothermometer, if arsenopyrite occurs in an fS2- buffered assemblage.(On the other hand, ƒS2 – unbuffered assemblages
furnish a range in T and fS2)
The common ƒS2 – buffered univariant assemblages include
Asp + Lo + Po,
Asp + Py + Po,
Asp + Py + As,
in which atomic %As in arsenopyrite uniquely fix both T and fS2


The Fe – Zn – As – S system
Thermobarometric interpretation of sulfide ores is hampered by problems of mineral reequilibration. However, this problem can be minimized by making use of refractory ore minerals that offer best candidates for preserving their composition pertaining to ore forming environments.
The Fe-Zn-As-S system is potentially one of the most useful
systems for thermobarometric interpretations
because it involves three of the most common and refractory sulfide minerals, namely
1)pyrite,
2)sphalerite
3)arsenopyrite

#
From Arsenopyrite composition- both fS2 and T can be evaluated.
#Sp barometry
can be pursued as Sp composition is unaffected by Arsenic.


The Cu – Fe – Zn – S system
more common in hydrothermal ore deposits than those in the arsenide- bearing system but,
in contrast,
commonly undergo post-depositional reequilibration.Not only the Cu- and Cu-Fe-sulfides readjust their composition during cooling, presence of minerals such as chalcopyrite promotes changes in the otherwise refractory sphalerite.
chalcopyrite disease :
(A very common feature in Zn- and Cu-bearing hydrothermal veins, volcanogenic, and metamorphosed massive sulfide deposits is the texture of dispersed blebs and rods of chalcopyrite in sphalerite)

[This texture has been notionally thought to be an exsolution product.
But experimental studies oppose this thought.
1)up to 500°C, solubility of chalcopyrite in sphalerite is negligible.
2)natural sphalerites containing chalcopyrite disease occur in ores such as carbonate- hosted Pb-Zn ores (100°–150°C) and unmetamorphosed VMS ores (200°–300°C)]

Explaining Cp Disease  in Sp:
(i) epitaxial growth during formation of sphalerite or
(ii) selective replacement of Fe- rich zones in sphalerite by a Cu-bearing chloride fluid,

Sphalerite stars or crosses,:
interpreted to have been formed due to exsolution,
because of appreciable solubility of ZnS in CuFeS2.