Another PGC International Study Tour
Developed & Managed by Porter GeoConsultancy
IOCG 07
Iron Oxide Copper-Gold in South America
4 to 13 June 2007
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CONTENT and DESCRIPTIONS OF ORE DEPOSITS
Image:   The Candelaria open pit, Chile.
Candelaria Pit
Porter GeoConsultancy continued its International Study Tour series of professional development courses by visiting a representative selection of the most important iron oxide copper-gold deposits in South America, in the Andes of Chile and the Carajas region of Brazil.
   The tour commenced in Marabá, Para state, Brazil on the evening of Sunday 3 June and ended in Santiago, Chile on the evening of Wednesday 13 June, 2007.
   Participants were able to take any 4 or more days, up to the full tour, as suited their interests or availability.

   The main components of the itinerary were:

Cristalino ...................... Monday 4 June, 2007.

The Cristalino IOCG deposit is located some 40 km to the east of Sossego in a bifurcation of the major regional Carajás Fault in the Carajás district of Para State, Brazil.

Basement in the area is represented by the Xingu Complex which is >2.86 Ga in age and is composed of a variety of rocks, including the ~3.0 Ga Pium Complex and 2.9 Ga Greenstones.   These are overlain by the 2.76 Ga Grão Para Group of volcanics and sediments, cut by the 2.5 Ga Estrella Granite and subsequently by 1.9 Ga granites but is overlain by the un-metamorphosed 2.7 to 2.6 Ga Águas Claras marine sandstones.

Cristalino is hosted by volcanics of the Grão Para Group composed of orange dacite and green andesite with minor basalt and in association with hydrothermally altered and disrupted banded iron formations within this same sequence.   These iron formations have been upgraded nearby where they constitute part of the Carajás Iron Resources.

Mineralisation is concentrated in a NW-SE trending, sinsitral transpressive zone of shearing over a drilled length of 2200 m and thickness ranging from 10's of metres to 500 m.   The shear zone is several hundreds of metres in width and is a splay of the Carajás Fault.   The ore zone is generally brecciated and is found in the volcanics below the iron formation and in the lower sections of the iron formation itself.   In general the iron formation forms the upper limit to ore and may have acted as a capping.   The hydrothermally altered breccia is composed of 5 to 50% sub-angular to sub-rounded fragments.

Mineralisation is associated with the emplacement of 2.7 Ga diorite to quartz-diorite intrusions into the volcano-sedimentary sequence and iron formation.

There are two styles of mineralisation:  (i). 60% of which is crosscutting stockwork veins and veinlets, and  (ii). 40% breccia ore where the breccia fragments are surrounded by sulphide veins and a sulphide matrix.   Mineralisation is also accompanied by magnetite and associated amphibole alteration.   The principal sulphides are chalcopyrite and pyrite in a 2:1 to 3:1 ratio.   The Copper was introduced after the magnetite and amphibolite alteration, although the highest grades are associated with the amphibole zones.   The iron alteration where it affects the iron formation represents addition, not remobilisation of iron.

Hydrothermal alteration progressed from:  (i). early widespread actinolite-albite; to  (ii). biotite with scapolite and magnetite; to  (iii). amphibole with magnetite as hastingsite, grunerite, actinolite and cummingtonite; to  (iv). chlorite with albite, magnetite and hematite; to  (v). chlorite and carbonate; to  (vi). muscovite and carbonate.

The average 3-5% sulphide mineralisation is associated with the last three overlapping phases of alteration and comprise chalcopyrite, pyrite and lesser arsenopyrite with trace Ni-Co sulphides.   The gold is in the pyrite.

Indications of Cu mineralisation were first noted in the area in the late 60's to early 70's.   Grid geochemistry and geophysics from 1984-87 led to 2 anomalies being drilled in 1988 with some 13 holes in two prospects.   The second phase of work was commenced in 1997-98 with more grid mapping, geochemistry and geophysics, culminating in a drill intersection of 38 m @ 1.4% Cu, 0.25 g/t Au between 76 and 114 m depth.

The resultant approximate resource from the subsequent drilling to 2001 amounted to 500 Mt @ 1.0% Cu, 0.2-0.3 g/t Au. According to CVRD, the reserves amount to 261 Mt @ 0.73% Cu.

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Sossego ...................... Tuesday 5 June, 2007.

The Sossego IOCG deposits are located some 40 km to the south of the Carajás townsite in the state of Para, Brazil. It is approximately 80 km SE of Igarapé Bahia and Alemao, and 40 km west of Cristalino. It is also ~30 km east of the main Carajas Serra Sul iron operation (#Location: 6° 25' 10"S, 50° 3' 55"W).

For details of the setting of the deposits see the Carajas IOCG Province record.

Mineralisation is hosted along a regional WNW-ESE-striking shear zone that defines the contact between the metavolcano-sedimentary rocks of the ~2.76 Ga Itacaiúnas Supergroup and the basement tonalitic to trondhjemitic gneisses and migmatites of the ~2.8 Ga Xingu Complex.

Both the Itacaiúnas Supergroup and Xingu Complex rocks are intruded by granite, granophyric granite, gabbro intrusions and late dacite porphyry dykes in the deposit area. The ages of these rocks are uncertain, although a Pb-Pb zircon age of 2734±4 Ga for a biotite-hornblende granite and a U-Pb zircon age of 2765±39 Ga for a tonalite-trondhjemite (Sardinha et al., 2004) close to the deposit area, are considered the best estimates. In addition, the Palaeoproterozoic Rio Branco granite intrusion crosscuts the hydrothermally altered volcanic and intrusive rocks in the deposit area. This field relationship, together with the lack of hydrothermal alteration and mineralisation of the granite, indicate it was emplaced much later than the formation of the copper-gold ore at Sossego (Xavier et al., 2010).

An ~1.88 Ga A-type alkaline to subalkaline magmatism event extends over an area of ~1000 x 1500 km of the Amazon craton and is represented in the Carajás domain by the Serra dos Carajás Intrusive Suite (Machado et al., 1991; Tallarico et al., 2004).

An extensive, >20 km
2, zone of scapolite alteration representing distal sodic alteration surround the Sossego deposits (Villas et al., 2005; Sousa, 2007). This alteration is characterised by plagioclase replacement and by scapolite in both mafic and felsic protoliths. Scapolite veins, with widths ranging from mms to >1 m, commonly crosscut basement and supracrustal units and have a biotite-scapolite-hastingsite rich halo that grades outward to chlorite.

In the Sossego deposit area, deformation is represented by cm- to metre-wide mylonitic zones, which are regionally crosscut by NE-striking faults. Rocks in the immediate footwalls of these faults are intensely mylonitised with biotite-tourmaline-scapolite and siliceous alteration developed prior to or synchronous with the shearing (Xavier et al., 2012).

The ore is located in two adjacent centres, Sossego Hill (the Sossego-Curral zones) and the larger Sequeirinho (the Pista-Sequeirinho-Baiano zones) which has a length of 1.6 km and thickness of 150 to 200 m in its central section. These two centres are separated by a major high angle fault. The original resource within Sequeirinho and Sossego was distributed ~85 to 15% respectively.

The Sequeirinho-Pista-Baiano orebodies are characterised by deeply emplaced magnetite-(apatite) and albite-actinolite-rich zones, whereas the Sossego Hill (Sossego-Curral) orebodies have predominant potassic and chlorite alteration typical of shallow crustal levels. Dating of hydrothermal monazite (U-Pb LA-MC-ICPMS) and molybdenite (Re-Os NTIMS) of the Sequeirinho-Pista orebodies rendered ages of ~2.71 to 2.68 Ga, while ore-related monazite of the Sossego-Curral orebodies yielded ages of ~1.90 to 1.88 Ga (Moreto et al., 2015). This implies two periods of IOCG style mineralisation at Sossego. The first occurred during coupling of ductile sinistral transpression with NNE-directed oblique shortening and Neoarchean magmatism at ~2.7 Ga. The second, shallowly emplaced Palaeoproterozoic system, formed after progressive exhumation of the deeply emplaced Neoarchaean IOCG deposits, and was coeval with the emplacement of 1.88 Ga A-type granites which may have caused regional circulation of magmatic and externally derived fluids along crustal discontinuities (Moreto et al., 2015). This second Palaeoproterozoic event was also responsible for other deposits in the region including

The Sequeirinho ore zone lies along a NE striking sinistral fault, associated with a positive magnetic anomaly, and comprises an S-shaped, tabular, subvertical body. It has been subjected to regional sodic (albite-hematite) alteration, overprinted by sodic-calcic (actinolite-rich) alteration accompanying with the formation of massive magnetite-(apatite) bodies. Both alteration assemblages exhibit ductile to brittle-ductile fabrics and are cut by spatially restricted zones of potassic (biotite and potassium feldspar) alteration that grades outward to chlorite-rich assemblages (Monteiro, et al., 2007).

The Sossego Hill zone is a subcircular, vertical, pipe-like orebody, with a central breccia surrounded by a stockwork of sulphide veins, faults, and shear zones (Morais and Alkmim, 2005; Carvalho, 2009; Domingos, 2009). The orebodies within the zone display only weakly developed early albitic and very poor subsequent calcic-sodic alteration, although they have well-developed potassic alteration assemblages that were formed during brittle deformation that produced breccia bodies. The matrix of the breccias commonly displays coarse mineral infill suggestive of growth into open space (Monteiro, et al., 2007). The potassic alteration assemblages, which mark the onset of mineralisation, grade outward to a widespread zone of chlorite and late hydrolytic (sericite-hematite-quartz) assemblages crosscut by calcite veins (Carvalho et al., 2005; Monteiro et al., 2008, ).

The orebodies are commonly brecciated. The Sequeirinho breccias contain rounded fragments of hydrothermal magnetite, actinolite and apatite or rocks with low angularity. The Sossego ore breccias comprise fragments of potassic-altered host rocks (e.g., granophyric granite) with high clast angularity, characteristic of breccias that underwent minor transport.

The sulphides of both groups of orebodies were initially accompanied by potassic alteration and a subsequent more important assemblage of calcite-quartz-epidote-chlorite. In the Sequeirinho orebodies, sulphides range from undeformed to deformed, while at the Sossego Hill orebodies they are undeformed. Very late stage, weakly mineralised hydrolytic alteration is present in the Sossego Hill orebodies (Monteiro, et al., 2007).

The dominant sulphides are chalcopyrite with subsidiary siegenite and millerite, and minor pyrrhotite and pyrite in the Sequerinho orebodies, although pyrite is relatively abundant in the Sossego Hill bodies.

Chalcopyrite occurs in the breccia matrix associated with pyrite, gold, siegenite, millerite, Pd melonite, hessite, cassiterite, sphalerite, galena, molybdenite, thorianite, and monazite. The resulting Fe-Cu-Au-Co-Ni-Pd- LREE signature characterises the Sossego deposit.

In early 2001 the total resource was quoted as 355 Mt @ 1.1% Cu, 0.28 g/t Au, encompassing a mineable reserve of 219 Mt @ 1.24% Cu, 0.33 g/t Au at a 0.4% Cu cut-off and stripping ratio of 3.3:1 wate:ore.

At the commencement of mining in 2004, reserves were quoted by CVRD as 250 Mt @ 1.0% Cu. Montiero, et al., (2007) published a reserve of 245 Mt @ 1.1% Cu, 0.28 g/t Au.

Mineralisation (gold) was initially discovered by garimperos (prospectors) in 1984 within CVRD concessions. The area was tendered to Phelps Dodge in 1996 and the first major intersections were in early 1997.

In 2001 the project was controlled by Mineracao Serra do Sossego, a 50:50 joint venture between Phelps Dodge do Brasil and CVRD.   In 2002 CVRD bought Phelps Dodge's share and commenced mining in 2004 with a nominal capacity of 93 000 tya of Cu in concentrates.

Remaining ore reserves at 31 December 2017 were (Vale 20-F form report to the US SEC, 2017):
    Proved Reserves - 110.7 Mt @ 0.68% Cu;
    Probable Reserves - 9.4 Mt @ 0.66% Cu;
    TOTAL Reserves - 120.1 Mt @ 0.68% Cu, with a recovery range of 90 to 95% of contained Cu.

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Igarapé Bahia ...................... Wednesday 6 June, 2007.

The Igarapé Bahia Au-Cu-(REE-U) deposit is located in the Carajás Mineral Province of Para State Brazil, ~60 km west of the Carajás township (or Nucleo Urbano) and ~40 km south of the Salobo deposit.
(#Location: 6° 1' 57"S, 50° 33' 59"W.

Igarapé Bahia is hosted by the Igarapé Bahia Group, considered to be a lower greenschist facies metamorphosed unit of the Archaean (ca. 2.75 Ga) metavolcano-sedimentary Itacaiúnas Supergroup which comprises two lithological and stratigraphic domains: a lower metavolcanic unit composed of metavolcanic rocks and acid to intermediate volcanoclastics; and an upper clastic-chemical metasedimentary unit and volcanoclastic rocks. The Igarapé Bahia orebodies represents a 100 to 150 m thick gossan-laterite zone from which significant amounts of gold (>60 t) were mined until 2003. Where not outcropping, the primary mineralisation is obscured by a 250 m thick unconformable siliciclastic unit referred as the Aguas Claras Formation.

The copper-gold mineralisation at the Igarapé Bahia deposit is hosted by a hydrothermally altered breccia at the contact between the footwall mafic volcanics, with associated BIF and hyaloclastite, and a dominantly coarse to fine-grained metasedimentary sequence in the hanging wall. The breccia unit is exposed at or near the surface as a semicircular annulus, with a form similar to a ring complex with a diameter of approximately 1.5 km. The mineralised breccia unit occurs as a 2 km long by 30 to 250 m thick series of fault dislocated bodies on the southern, northeastern and northwestern sections of this structure, dipping steeply outwards at ~75°, and is nearly concordant with the metavolcanic-sedimentary wallrocks.

The economically extracted ore at Igarapé Bahia is largely developed as a supergene gossan-laterite enrichment within the 150 to 200 m thick oxide profile. Three orebodies have been mined at this contact, forming a semi-circular trace at the surface namely, Acampamento - dipping at around 75° to the north-east, Furo Trinta to the south-east, and Acampamento Norte to the north-west, forming an outward dipping domal structure in three dimensions.

The oxide zone is characterised by supergene enrichment and hematite, goethite, gibbsite and quartz. This is underlain by a transition zone that may be up to 50 m thick with enriched supergene malachite, cuprite, native copper and goethite and minor amounts of digenite and chalcocite responsible for high grade Cu and Au. This zone is in turn underlain by primary Cu-Au mineralisation, represented by hydrothermal breccias containing chalcopyrite, bornite, carbonate, magnetite and minor molybdenite and pyrite.

Strong hydrothermal alteration of the host sequence produced intense chloritisation, Fe-metasomatism, Cu-sulphidation (chalcopyrite and bornite), carbonatisation, silicification, tourmalinisation and biotitisation in the primary zone.

Gold-copper mineralisation is localised at the commonly brecciated contact between the metavolcanics and the meta-volcaniclastics-metasediments and comprises, magnetite/siderite heterolithic breccias and hydrothermally altered metavolcanics. These rocks are enriched in REE (monazite, allanite, xenotime, bastnäsite and parisite), Mo (molybdenite), U (uraninite), F (fluorite), Cl (ferropyrosmalite) and P (apatite).

The deposit was mined by CVRD/Vale between 1991 and 2002 at a production rate of up to 10 t Au per annum for a total production of 97 t of recovered gold. The remaining reserve in 1998 was quoted as 29 Mt @ 2 g/t Au (Tazava and Oleira, 2000). Reserves in 1999 were quoted by Tallarico et al. (2000) as 18.5 Mt @ 1.97 g/t Au

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Alemão ...................... Wednesday 6 June, 2007.

The Alemão IOCG Au-Cu-(REE-U) deposit is part of the Igarape Bahia mineralised system in the Carajás Mineral Province of Para State Brazil (see the Igarapé Bahia record).
(#Location: 6° 1' 35"S, 50° 34' 38"W).

Alemão is hosted by the Igarapé Bahia Group, considered to be a lower greenschist facies metamorphosed unit of the Archaean (ca. 2.75 Ga) metavolcano-sedimentary Itacaiúnas Supergroup which comprises two lithological and stratigraphic domains: a lower metavolcanic unit composed of metavolcanic rocks and acid to intermediate volcanoclastics; and an upper clastic-chemical metasedimentary unit and volcanoclastic rocks. The Alemão ore body underlies the far northwestern margin of the Igarapé Bahia deposit, which represents a 100 to 150 m thick gossan-laterite zone from which significant amounts of gold (>60 t) were mined until 2003. Elsewhere it is obscured by a 250 m thick unconformable siliciclastic unit referred as the Aguas Claras Formation.

The copper-gold mineralisation at the Igarapé Bahia/Alemão deposit is hosted by a hydrothermally altered breccia at the contact between the footwall mafic volcanics, with associated BIF and hyaloclastite, and a dominantly coarse to fine-grained metasedimentary sequence in the hanging wall. The breccia unit is exposed at or near the surface as a semicircular annulus, with a form similar to a ring complex with a diameter of approximately 1.5 km. The mineralised breccia unit occurs as a 2 km long by 30 to 250 m thick series of fault dislocated bodies on the southern, eastern and northern sections of this structure, dipping steeply outwards at ~75°, and is nearly concordant with the metavolcanic-sedimentary wallrocks. The Igarapé Bahia deposit is the thick gossan-laterite zone developed within the top 100 to 150 m of the exposed breccia unit.

The Alemão deposit is located immediately to the northwest of this annular zone, occurring as a particularly magnetite-Cu-Au-enriched down-faulted segment of the Acampamento Norte orebody, the northern most orebody of the Igarapé Bahia deposit. It has dimensions of around 500 m in length, 50 to 200 m thick and has been traced down plunge for at least 800 m below the surface, although the top of the deposit is at a depth of approximately 250 m below Aguas Claras Formation sandstone cover.

The Alemão orebody is hosted by a hydrothermally altered breccia at the contact between the footwall mafic volcanics, with associated BIF and hyaloclastite, and a dominantly coarse to fine-grained metasedimentary sequence in the hanging wall. A set of unmetamorphosed 2.75 to 2.65 Ga quartz diorite and 2579±7 Ma dolerite dykes cut the orebodies, the host metavolcano-sedimentary sequence and the overlying clastic metasedimentary sequence of the Áoguas Claras Formation/Rio Fresco Group.

The breccia has gradational contacts with its wallrocks and is made up of polymitic, usually matrix-supported clasts, composed mainly of coarse, angular to rounded basalt, BIF and chert clasts derived from the footwall unit.

The hydrothermal paragenesis is marked by ferric minerals (magnetite and hematite), sulphides (chalcopyrite-pyrite), chlorite, carbonate (siderite, calcite, ankerite) and biotite with minor quartz, tourmaline, fluorite, apatite, uraninite, gold and silver. Sericite and albite are rare. The mineralisation is represented by hydrothermal breccias and hydrothermally altered rocks. These fall within two groups, namely:

i). massive magnetite-chalcopyrite bands and polymict breccias with a matrix of magnetite, chalcopyrite, siderite, chlorite, biotite and amphiboles; and
ii). brecciated hydrothermally altered volcanics with chalcopyrite, bornite, pyrite, chlorite, siderite and ankerite both in the matrix and disseminated in the altered country rock.

Several generations of late mineralised veins crosscut the ore breccia and are composed of variable concentrations of chalcopyrite, pyrite, quartz, calcite, chlorite, and fluorite. The veins commonly display open space - filling textures (e.g., comb)

The total estimated resource in 2001 was 170 Mt @ 1.5% Cu, 0.8 g/t Au. More recently CVRD has quoted a reserve of 161 Mt @ 1.3% Cu, 0.86 g/t Au.

Alemao Chlorite Breccia

Alemao breccias. Above - Chlorite-carbonate-sulphide breccia. Note the well developed striations are created by the diamond saw blade that cut the surface shown;
Below - magnetite-sulphide breccia. In both images, the smaller scale card graduations are in millimetres.
Images by Mike Porter, 2020, samples collected at Alemao 2001.

Alemao Magnetite-sulphide Breccia


The deposit was discovered by Docegeo, exploration arm of CVRD (now Vale) in 1996.

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Salobo ...................... Thursday 7 June, 2007.

The Salobo 3 Alpha IOCG Deposit is located in the Carajás district of Para State, Brazil, and is some 30 km to the north of Igarapé Bahia and Alemao and ~50 km WNW of the major Carajas N4 and 5 iron deposits of the Serra Norte (#Location: 5° 47' 25"S, 50° 32' 2"W).

The Salobo deposit was discovered in 1977. It lies within the WNW-ESE trending Cinzento shear zone, hosted by a package of rocks that includes the Igarapé Salobo Group which is interpreted to be a supracrustal suite that is part of the Neoarchaean 2.76 to 2.73 Ga Itacaiúnas Supergroup; orthogneisses of the Mesoarchaean basement Xingu Complex (2950 ±25 to 2857 ±6.7 Ma; U-Pb zircon; Melo et al., 2016) and deformed granitoid gneiss of the Neoarchaean Igarapé Gelado suite (2763 ±4.4 Ma; U-Pb zircon; Melo et al., 2016). The Igarapé Salobo Group is composed of paragneiss, amphibolites, quartzites, meta-arkoses and iron-rich schist.

For details of the regional and structural setting see the Carajás IOCG Province record.

The Itacaiúnas Supergroup sequence is in tectonic contact with trondhjemitic gneiss of the basement Xingu Complex which has been partially migmatised. The original stratigraphic relationships and contacts with the basement, as well as within the host sequence, are masked by intense ductile-brittle shearing and over-thrusting/reverse faulting. This includes strong deformation within the broad 2.7 Ga Itacaiúnas Shear Belt that caused imbrication, and tectonic layering of supracrustal rocks alternating with basement gneisses at several scales. It forms a broad, tens of kilometres wide, braided zone of steeply-dipping, WNW-ESE trending ductile shearing and high temperature mylonitic fabrics developed under upper amphibolite facies regional metamorphic conditions (e.g. DOCEGEO 1988; Araújo & Maia 1991). Structural indicators imply a regime of predominantly sinistral transpression with partitioning of deformation that produced linked systems of ductile strike-slip and thrust dominated shear zones (Araújo and Maia, 1990; Costa et al., 1994). One of these was the ~2.5 Ga Cinzento ductile shear zone that hosts the Salobo deposit (Machado et al., 1991; Holdsworth and Pinheiro, 2000). These zones of shearing are crosscut by the undeformed A-type, peralkaline to metaluminous (Lindenmayer, 2003) syn-tectonic Old Salobo Granite dated at 2573 ±2 to 2547 ±5.3 Ma (Machado et al., 1991; Melo et al., 2016). The sequence is also cut by the 1.88 Ga anorogenic, metaluminous, isotropic Young Salobo Granite (Lindenmayer, 1990; 1998) which is also recognised at the Salobo deposit. The development of these shear zones resulted in a widespread and penetrative, sub-vertical, northwest-striking mylonitic foliation in the rocks of the Salobo deposit area, with the exception of the Young Salobo granite and late dolerite dykes (Réquia et al., 2003).

The principal lithology in the Salobo deposit area is a biotite-garnet (almandine)-quartz rich rock that has undergone intense iron and potassic hydrothermal alteration at high-temperatures in a ductile regime that has formed a mylonitic rock package containing variable amounts of magnetite, actinolite, grunerite and tourmaline. Lindenmayer (1990) first suggested it was a metagreywacke, but subsequently reinterpreted it to be a hydrothermally altered basaltic-andesite and dacite of the Igarapé Salobo Group. The host package also includes amphibolite to the NE and quartzite to the SW (Xavier et al., 2010). Alternatively, the iron-rich schists have been interpreted to represent sedimentary iron formations within the Igarapé Salobo Group that have been metamorphosed to pyroxene-hornfels facies (e.g., Campo Rodriguez et al., 2019; Xavier et al., 2010; Villas and Santos, 2001; Lindenmayer, 1990). Similar, structurally disrupted 'iron formations' extend intermittently along strike from Salobo over many tens of kilometres throughout the district (Siqueira and Costa,1991). However, the mineralised iron-rich rocks at Salobo differ from the regional iron formations in that they are enriched in Cu, Au, Ag, U, F, Mo, Co and LREE, whereas the banded iron formations (e.g., at Carajás) are depleted in these elements (Réquia and Fontboté, 2000). Melo et al. (2019) interpret the main host unit gneiss, iron-rich schist and structurally overlying amphibolite to be metamorphosed and altered rocks of the Xingu Complex, straddled by the Igarapé Gelado suite, whilst the 2.76 to 2.73 Ga Itacaiúnas Supergroup Igarapé Salobo Group volcanosedimentary sequence is only represented by mylonitic quartzite remnants on the margin of the deposit to the SW (Melo et al., 2016).

The Salobo 3A deposit extends over a NW-SE to WNW-ESE trending strike length of ~4 km and is 100 to 600 m in width. Mineralisation occurs as steeply dipping, irregular, lens-shaped and massive replacement orebodies following the mylonitic foliation, that have been recognised to depths of 750 metres below the surface (Souza and Vieira, 2000).

The deposit occurs as irregularly distributed lenticular shaped ore shoots within the major brittle-ductile WNW-ESE trending Cinzento Shear Zone. The lens-shaped and massive replacement orebodies are parallel to planar S-C structures along the trend of the shear zone, and commonly exhibit plastic flow textures, recrystallisation, mylonitisation and brecciation (Lindenmayer, 1990; Lindenmayer and Teixeira, 1999; Siqueira and Costa, 1991). The host rocks were progressively metamorphosed to pyroxene hornfels facies, at equilibrium temperatures of 750°C and pressures of up to 2 to 3 Kbar (i.e., 7 to 11 km depth), resulting from sinistral transcurrent transpressive shearing accompanied by oblique thrusting/reverse faulting. This metamorphism produced an assemblage with a coarse granoblastic texture, consisting of fayalite, almandine, spessartine, magnetite, hastingsite, chalcopyrite and graphite (Souza and Vieira, 2000).

The structurally-controlled and massive replacement ore bodies are generally associated with a halo of variably magnetite-rich (<10 to >50%) rocks with Mn-almandine, grunerite, Cl-rich hastingsite, fayalite, schorlitic tourmaline, Fe-biotite, allanite and quartz (Réquia et al., 2003; Réquia, and Fontboté, 2000). As described below, the ore occurs within strongly iron-potassic altered rocks in two main zones: i). massive garnet-biotite-fayalite-grunerite rock which generally has >50% magnetite with minor graphite and fluorite, and ii). a foliated, granoblastic, almandine-biotite-grunerite-plagioclase-quartz assemblage with 10 to 50% magnetite, and extends into the adjacent biotite-garnet-quartz schists. There is a direct relationship between copper and iron grades (Viera et al., 1988; Souza and Vieira, 2000).

Further to the generalised description in the previous paragraph, the main mineralised core of the deposit has been variously subdivided by different authors. The NI 43-101 Technical Report prepared by Burns et al. for Wheaton Precious Metals and for Vale S.A., dated 31 December 2019, recognises the following lithotypes:
Biotite Schist - which forms the bulk of the mineralised core of the deposit. It is medium to coarse-grained rock with anastomosed foliation and is characterised by biotite, garnet, quartz, grunerite, magnetite and chlorite. This assemblage is partially replaced by a second generation of biotite and magnetite with chlorite, K feldspar, quartz, hematite and sulphides. Tourmaline, apatite, allanite, graphite and fluorite generally occur throughout the lithotype.
Magnetite Schist, occurring as branching lenses of massive, foliated and banded rocks, predominantly composed of magnetite, fayalite, grunerite, almandine and secondary biotite. Granoblastic textures with polygonal contacts between magnetite and fayalite are common. The southeast portion of the deposit hosts hastingsite, replaced partially by actinolite, grunerite and sulphide minerals. Fluorite, apatite, graphite and uranium oxides are associated with this assemblage. Within more massive magnetite mineralisation there are small veins and irregular masses of secondary biotite and garnet is completely replaced by magnetite, forming pseudomorphs. Away from the massive magnetite, the magnetite content gradually diminishes, giving way to biotite-garnet schist and/or garnet–grunerite schist. The copper content of magnetite-schist is typically >0.8%.
Garnet-Grunerite Schist, found as bands and lenses in the central to northeastern sections of the deposit. It has an isotropic texture, with weak dispersed schistosity, and a granoblastic texture. The principal mineral assemblage consists of almandine and cummingtonite-grunerite, with magnetite, hematite, ilmenite, biotite, quartz, chlorite, tourmaline and subordinate allanite. Fluorite and uraninite generally occur in veinlets related to stilpnomelane, calcite and grunerite.
Feldspar-Chlorite Mylonite, which forms the northwestern margin of the mineralised core. It is principally composed of feldspar, chlorite and quartz with a mylonitic foliation, produced by the orientation of rims of chloritised deformed biotite, hastingsite, elongated quartz and saussuritised plagioclase (K feldspar, epidote and muscovite alteration). Porphyroblastic garnet is partially or totally replaced by chlorite and epidote. Allanite and apatite generally occur throughout this lithology.
Quartz Mylonite, which forms the southwestern margin of the mineralised core of the deposit. It is grey or white in colour, passing through green to red. Where present, Fe-oxides are medium to fine grained, foliated and composed predominantly of quartz, muscovite, sericite, sillimanite and chlorite. Accessories, such as biotite, feldspar, magnetite, almandine, tourmaline, zircon and allanite are common. The following variations have been differentiated: i). red quartzo-feldspathic rocks composed of K feldspar and quartz and which may be a product of shearing between the gneissic basement and the supracrustal rocks; and ii). chlorite schists, mainly composed of chlorite and quartz, that represent intense hydrothermal alteration. This variant is found near the southwestern border of the deposits, close to important brittle shear zones, which may be interpreted as conduits for hydrothermal fluids.

Viera et al. (1988) and Souza and Vieira (2000) instead, have described the distribution of alteration and mineralisation in terms of five different 'schist types' which together defined the host 'Tres Alpha Formation' of the Igarapé Salobo Group. These occur as compositional lenses characterised by particular mineralogies. They are not stratigraphic units and in general do not display any consistent succession or zoning. They were the main lithotypes distinguished in mapping and drill core logging when the deposit was visited in 1992 (Porter, 1992), and are as follows after (Viera et al., 1988):
Schist X1 - occurring as discontinuous lenticular zones in the core of the deposit. It is massive and generally coarsely crystalline, with >50% magnetite. Shear banding is sometimes observable. The subordinate minerals are garnet, biotite, fayalite, grunerite, graphite (1 to 1.5%) and fluorite. X1 may contain up to 5% Cu, but generally has >1 to 1.2% Cu.
Schist X2 - coarse grained, porphyroblastic and foliated, but displaying little compositional banding. It is relatively rare compared to the other schists, and is mainly composed of garnet (2 to 4 mm) and grunerite with <10% magnetite and subordinate biotite and quartz. When very rich in fayalite, X2 may approach the Fe content of X1, but is invariably lower. The main difference to X1 is the magnetite and Cu content. It has the lowest Cu grades of the five schists, usually <0.5% Cu.
Schist X3 - is both foliated and banded. It has a grano- to lepido-blastic texture and is generally similar in appearance to X1. However the main component minerals are biotite and garnet, with a magnetite content of between 10 and 50%. Subordinate minerals are fayalite, grunerite, quartz and plagioclase. The Cu content is usually 0.5 to 1.1% Cu. This lithotype contains the bulk of the ore in the deposit.
Schist X4 - is similar in appearance to X3, being medium grained, foliated and banded, with a porphyroblastic texture and <10% magnetite. In addition, the principal component minerals are biotite, garnet and quartz, with variable amounts of grunerite, olivine and plagioclase. The Cu content is usually <0.5% Cu.
Schist X5 - is foliated, well banded and fine grained and is composed of plagioclase, biotite, quartz and amphiboles, but has no magnetite. Common accessories are garnet and chlorite. This unit generally forms the outer margins of the Tres Alpha Formation, having gradational boundaries with the overlying quartzo-feldspathic rocks of the Cinzento Formation and the underlying Cascata Gneiss. X5 represents zones that underwent greater ductile strain, as a result of both shearing and hydrothermal activity (Souza and Vieira, 2000). This unit always has <0.5% Cu.

Less deformed exposures of these schists yield textures and structures that have been interpreted to resemble volcanic rocks. There is no consistent relationship between each of the schist types, except that X1 is often a core to developments of X3, and the combination of X1 and the more extensive X3 which constitute the orebody, form the core of the overall schist zone. The X1 lenses are generally from a few mm's up to 10 m thick, with a maximum of 30 m on one section, with lateral dimensions of 20 to 500 m, while the combined X3 and X1 may be up to 100 m thick and extend for up to a kilometre. In detail, lenses of X3 also occur within X1 and thin and pinch out along foliation, while others occur in an en echelon pattern. The contact between X1 and the other facies is commonly, but not exclusively, sharp, being generally 1 to 2 mm wide and always parallel to foliation. The transition from X1 to X4 for instance, is marked by a change from mainly magnetite with infrequent grunerite-fayalite to banded fayalite-grunerite with less common 1 to 2 mm magnetite bands. All other contacts are transitional over widths of a few cm's to tens of metres. The spatial discontinuity and lensoid nature observed between the different schist types is interpreted to be the result of intense tectonic dislocation, involving both imbrication and hydrothermal alteration (Souza and Vieira, 2000).

The high magnetite sections of X1 are grey to black, metallic and appear in places to be almost massive magnetite, with well foliated compositional banding, marked in part by variable thin green-yellow fayalite bands. Magnetite occurs as coarse aggregates up to 5 mm across which are usually aligned parallel to foliation. Chalcopyrite is present as up to 2 x 3 mm blobs while more abundant fine 0.25 to 0.5 mm bornite (distinguished from magnetite by its bluish glint) is distributed along foliation planes and to a lesser extent as strings of separated grains filling fractures in a number of directions. In other sections, high grade bornite/chalcocite follows irregular anastomosing fractures. There are also small patches of pale fluorite spread through the schist. The coarse magnetite crystals and aggregates display a texture approaching that of a breccia. In lower magnetite zones of X1 and in X3, greenish yellow fayalite is accompanied by pale greenish-grey grunerite, while garnet is present as oval shaped aggregates up to 1 cm or more long, sometimes as big as 3 x 10 cm. The banding of fayalite and grunerite wraps around the garnets, while in places the larger garnet crystals are commonly shattered with chalcocite and bornite within the cracks. Biotite also develops within the cracks in garnets, while grunerite (an amphibole - Fe
7Si8O22•(OH)2) rims the larger fayalites (an olivine - Fe2SiO4). The garnets appear to be both syn- and post-metamorphism, with many being aligned parallel to the foliation, while others cut it. The main garnet is almandine (Fe3Al2(SiO4)3), although it is commonly rimmed by spessartine (Mn3Al2(SiO4)3). In the more heavily sheared intervals, the cupriferous minerals are coarser. Although comparatively rare, 1 to 30 cm thick carbonate veins, are found every 5 to 10 m in core through the ore zone, comprising dolomite-calcite and less commonly, siderite with coarse bornite and chalcocite. There is a close relationship between fluorite, grunerite and fayalite. Fluorite replaces the silicates with associated chalcocite-bornite and lesser chalcopyrite. In general there is no fluorite in massive fayalite, but where fayalite and magnetite are intergrown fluorite is present and follows the foliation. There is also a relationship between magnetite and Cu sulphides. Where magnetite is interbanded with massive fayalite, the associated Cu sulphides are always within the magnetite. In detail, in massive magnetite bands, sulphides are present as rims around the magnetite grains, or as fine stringers in fractures cutting the magnetite. Magnetite is almost invariably coarser than associated chalcocite-bornite. Au is also closely associated with magnetite. There is an apparent variation in the type of Cu sulphide depending upon the gangue. Massive magnetite is characteristically accompanied by chalcocite-bornite, while in interbanded fayalite and magnetite, chalcopyrite predominates. Overall within the orebody, chalcocite and bornite account for 85% of the Cu sulphides, while chalcopyrite comprises the remaining 15%. There is very little pyrite within the deposit, although small inclusions of pyrite with alteration rims and magnetite are evident within some chalcopyrite accumulations, as are pyrrhotite with pentlandite which occur as exsolutions. The information in this paragraph has been drawn from various sources from the reference list below.

Polished sections of the ore reveal a paragenetic succession in which magnetite was deposited in the early oxide stage, accompanied by small amounts of hematite, the silicate minerals fayalite (Fe olivine), biotite, garnet (Fe almandine), some fluorite (which has a close association with fayalite and magnetite), plagioclase and chlorite. Parts of the deposit were relatively reducing during this phase, as is indicated by the presence of graphite. The sulphide stage followed the oxide phase, with Mo deposited early, exhibiting a close association with graphite, which comprises 1 to 1.5% of X1. Re-Os molybdenite dating of the Salobo ore yielded an age of 2576 ±8 Ma (Réquia et al., 2003). The sulphide stage is characterised by the formation of tetragonal chalcopyrite, followed progressively by bornite and finally chalcocite. While these sulphides are the main ore minerals, significant Co, Ni, As, Ag, Au, Mo, F, rare earth elements (REEs), and U are characteristic of the Salobo ore represented in part by the presence of subordinate covellite, molybdenite, cobaltite, safflorite [(Co,Ni,Fe)As
2], native gold and silver (Lindenmayer, 1990; RĂ©quia et al., 1995). There are two generations of chalcopyrite, the earlier, pre- bornite development and a late, post chalcocite variety found in veins. In sections of the drill core, chalcopyrite is coarse and spectacular without much apparent accompanying bornite-chalcocite. At the end of the sulphide stage, native gold precipitation occurred in spatial association with cobaltite and safflorite. There is a close relationship between magnetite, Cu and both Au and Ag. In the transition from X1 to X3, the amphibole content does not change substantially. There is however, an antipathetic relationship between magnetite and almandine garnet. Petrographic evidence, such as magnetite cutting rotated garnet and chalcopyrite interstitial to fayalite grains or filling its fractures, indicates that at least some of the mineralisation is post the of peak metamorphism (Réquia and Fontboté, 2000).

Most rocks within the deposit area have been strongly altered, with those least affected, possibly closest to the protoliths but still significantly altered, composed of Ca amphibole ±plagioclase ±quartz ±sericite ±epidote ±chlorite, with or without tourmaline, biotite and K feldspar. The composition of the unaltered rocks is taken to be tholeiitic basalts (Réquia and Fontboté, 2000), based on chemical affinities with the overlying amphibolite interpreted to be of that composition (Lyndenmayer, 1990). These are overprinted by a partially preserved high-temperature calcic-sodic hydrothermal assemblage that includes the amphiboles hastingsite [NaCa
2Fe2+4Fe3+(Al2Si6O22)(OH)2] and actinolite [Ca2Mg4.5-2.5Fe0.5-2.5(Si8O22)(OH)2], as well as Ca and Na plagioclase. This phase is marked by rocks with high Na2O contents of up to 4.5 wt.%. Ca is inferred to have preceded Na alteration, as indicated by narrow rims of Na-plagioclase commonly surrounding crystals of Ca-plagioclase. Plagioclase composition ranges from bytownite to sodic oligoclase (Réquia and Fontboté, 2000). Subsequently, silicification, iron-enrichment (almandine-grunerite-magnetite) and tourmaline formation took place. The dominant alteration associated with sulphide mineralisation is potassic, overprinting the Ca-Na phase, characterised by >3.5, up to 4.6 wt.% K2O). It comprises an assemblage that includes K feldspar-quartz ±Ca-amphibole ±cummingtonite ±plagioclase ±sericite ±epidote ±chlorite, with or without biotite, calcite, tourmaline, titanite and kaolinite. The latter assemblage is observed in the central part of the deposit which is also the richest ore zone. Fe-Mg amphibole, represented by cummingtonite, commonly replaces Ca-amphiboles. The local replacement of Mg-hornblende by actinolite is accompanied by epidote, chlorite and quartz formation. Plagioclase crystals, mainly of labradoritic [(Na,Ca)1-2Si3-2O8] composition, are extensively replaced by K feldspar (orthoclase). Biotite dominates in rocks with only minor or no K feldspar, in association with titanite and quartz (Réquia and Fontboté, 2000). This alteration assemblage developed under intense ductile deformation at temperatures between 650 to 550°C, based on the associated mineral assemblages (Lindenmayer, 1990).

The tectonic evolution of the Salobo deposit was complex, including sinistral transpressional ductile shearing with associated thrusts, followed by sinistral, transtensional, brittle shearing (Souza and Vieira, 2000), as follows:
Sinistral transpressional ductile shearing produced a widespread, NW-SE orientated, sub-vertical, mylonitic foliation and imbrication of the lithological units, and the tectonic layering of strips and lenses of supracrustal rocks alternating with gneisses (Siqueira,1996). The peak of this deformation lies somewhere between 2851 ±4 and 2761±3 Ma (U-Pb, zircon; Machado et al., 1991), but must have been more long lived, or was rejuvenated, as it also apparently affected the 2573 ±2 Ma Old Salobo Granite dated at (Machado et al., 1991). This deformation coincided with the anhydrous metamorphic event that produced the coarse granoblastic textured iron silicate and oxide rich assemblage of ferrosilite [FeSiO
3], fayalite, almandine, spessartine and magnetite with hastingsite, chalcopyrite and graphite, characterised by high temperature, low pressure thermal pyroxene hornfels facies metamorphism (750°C; 2 to 3 Kbar; Souza and Vieira, 2000). This high grade metamorphism was followed by the initial calcic-sodic alteration event resulting in the assemblage described previously, which was largely obliterated (Réquia and Fontboté, 2000). Subsequent high temperature potassic alteration led to fluid penetration and hydration of dehydrated minerals, characterised by partial destruction of fayalite, hastingsite and chalcopyrite to produce grunerite, almandine, magnetite, biotite, bornite and chalcocite, as well as the addition of further bornite and chalcocite. This alteration assemblage developed during or in the waning stages of intense ductile deformation at temperatures of between 650 and 550°C (Lindenmayer, 1990). Potassic alteration accompanied the main mineralisation stage with early Mo sulphide mineralisation dated at 2576 ±8 Ma (Re-Os molybdenite; Réquia et al., 2003) and ore samples dated at ~2452 ±14 Ma (U-Pb monazite; Melo et al., 2016). This also temporally corresponds to reactivation of the Cinzento Shear Zone at ~2.5 Ga (Tassinari et al., 2003; Melo et al., 2019), and intrusion of the Old Salobo Granite (2573 ±2 Ma; Réquia et al., 2003). Evidences of this potassic alteration includes growth of grunerite along fayalite cleavage planes; the substitution of fayalite by grunerite plus magnetite; formation of almandine containing inclusions of grunerite; the substitution of chalcopyrite by bornite and chalcocite (Lindenmayer, 1990). The hydrothermal fluid is interpreted to been acidic, weakly oxidising, rich in SiO2 and K+, and also highly saline, given that it introduced Si and K and removed Ca, Mg and Na. The result of the potassic alteration was the enrichment of the host rocks in Fe2+, K, Ce, Th, U and REE (Lindenmayer, 1990).
Sinistral transtensional brittle faulting was the result of further renewal of displacement on the Cinzento Shear Zone and was particularly evident along the contact of quartzites and gneisses in the SW section of the deposit. It overprinted the earlier deformation with a sub-parallel fabric, dated by Mellito et al. (1998), from magnetite in brecciated iron rocks at 2172±23 Ma (Pb-Pb) and from chloritised gneisses at 2135±21 Ma (Rb-Sr, whole rock). This was accompanied by another hydrothermal event at temperatures of <370°C characterised by the infiltration of Ca-bearing fluids accompanied by intense chloritisation of almandine, biotite and hastingsite within the iron-rich rocks and intense chloritisation in wall rocks. Mineralisation associated with this phase represents the late Riedel shear controlled veining and includes quartz, stilpnomelane, fluorite, allanite, chalcopyrite, molybdenite, cobaltite and gold (Souza and Vieira, 2000) with greenalite-fluorite and uraninite fringes encapsulating fayalite and grunerite, accompanied by partial substitution of bornite by chalcocite (Lindenmayer and Teixeira, 1999; Lindenmayer, 2003). These late veins contain the second generation of chalcopyrite described previously. The fluid introduced during this stage was probably acidic, weakly saline and more oxidising than the high-T fluid in the previous hydrothermal stage (Lindenmayer, 1990).

Campo Rodríguez et al. (2019) recognised three stages of magnetite crystallisation with associated sulphides. Stage I, an inclusion free 'massive cystalline magnetite' with ferrosilite, fayalite, hastingsite and associated bornite as well as chalcopyrite with inclusions of pyrite. This is interpreted to have been emplaced towards the end of the main period of ductile shearing and emplacement of the Igarapé Gelado Suite at 2763 ±4.4 Ma. Stage II is a 'magnetite-bearing breccia', comprising inclusion-rich magnetite surrounded by a chalcopyrite matrix, which in turn, hosts pyrite and pyrrhotite close to Fe-rich magnetite and amphibole mineral grains and fragments. Some pyrite and pyrrhotite esxolutions are replaced by chalcopyrite. Inclusions within the magnetite are micro- to nanometre-scale and randomly distributed, composed by REE, zircon, apatite and Cu-bearing minerals, mainly bornite and chalcopyrite. This stage is also interpreted to have been formed late in the peak metamorphic stage and during emplacement of the Igarapé Gelado Suite. Stage III, 'magnetite schist' is taken to have formed at ~2.5 Ga during reactivation of the Cinzento Shear Zone strike-slip faulting. It has high quantities of inclusion-poor magnetite with an equigranular, granoblastic texture, which follows the schist foliation, and is accompanied by fibrous molybdenite and graphite. δ
34S signatures of pyrite, chalcopyrite and pyrrhotite for sulphides associated with stage I and II magnetite range vary from 1.70 to 5.04 (average 2.72‰) and 0.88 to 1.98 (average 1.56‰). The close relationships imply chalcopyrite has inherited δ34S values, at least in part, from the pyrite and pyrrhotite respectively, reflecting reactions between pyrite and an oxidised Cu-rich fluid, which resulted in a replacement of pyrite by chalcopyrite. In addition with the association between high-temperature minerals (i.e., fayalite and ferrosilite) and magnetite-bornite indicate that primary mineralisation included in the stages I and overlapping stage II was formed from the same evolving magmatic fluid at high temperatures. Moreover, the lack of negative or broad Δ33S values, which are close to zero, are interpreted by Campo Rodríguez et al. (2019) to be compatible with high-temperature oxidised-hydrothermal fluids without a contribution from shallow or surficial fluids. Subsequent metamorphism accompanying emplacement of the Old Salobo Granite and reactivation of the Cinzento Shear Zone at ~2.5 Ga generated the stage III magnetite (Campo Rodríguez et al., 2019).
  Melo et al. (2019) investigated oxygen and sulphur isotopes to draw similar conclusions, as follows. The iron enrichment at Salobo, occurred at 565 ±50°C, accompanied by hydrothermal fluids with magmatic or metamorphic compositions. This is evidenced by grunerite with δ
18OH2O = 7.20 to 8.50‰, δDH2O = -25.33 to -16.01‰;   garnet with δ18OH2O = 7.10 to 9.70‰;   and tourmaline with δ18OH2O = 5.07 to 7.37‰, δDH2O = -32.13 to +11.60‰ (Melo et al., 2019). However, the fluid inclusions at Salobo are hypersaline containing 30.6 to 58.4 wt.% NaCl equiv., favouring a magmatic origin in domains where the bulk of metamorphic devolatilisation is restricted to shear zones (Réquia, 1995).
  The fluids that are associated with potassic alteration at 565 ±50°C, also have a typical magmatic/metamorphic composition, indicated by biotite with δ
18OH2O = 7.23 to 18.03‰, δDH2O = -40.94 to -25.94‰;   and quartz with δ18OH2O = 7.52‰. Similarly, the δ34SV-CDT signatures of chalcopyrite = 0.81 to 1.28‰ and bornite = -0.37 to +1.63‰, which are interpreted to suggest an evolving felsic magmatic sulphur source at Salobo with little or no input from shallower basin or meteoric fluids (Melo et al., 2019). Réquia (1995) showed that fluids at Salobo were rich in H2O-CO2-NaCl-(CaCl2-CH4) with homogenization temperatures up to 485°C. The mineralising hydrothermal fluid is therefore interpreted to have been hypersaline, hot, F-rich, oxidised and carrying Cu, derived from a felsic magma coeval with ductile to brittle deformation at a depth of probably >6 km, presumably precipitating Cu when reduced/neutralised by magnetite.

The uncut geological resource in 2000 was estimated to be:
    1926 Mt @ 0.59% Cu, 0.34 g/t Au, 6.07% Fe
3O4, 0.16% C, 0.27% S, 0.23% F (Souza and Vieira, 2000).
The estimated mineral resource prior to 2000 was: 746 Mt with 0.93% Cu, 0.56 g/t Au, 9.79% Fe
3O4 at a cutoff of 0.6% Cu (Souza and Vieira, 2000).

Following the 2004 feasibility study, CVRD quoted reserves of: 986 Mt @ 0.82% Cu, 0.49 g/t Au at a 0.5% Cu cutoff.

The Salobo I processing plant commenced production in 2012 with a total capacity of 12 Mtpy of ore processed. The open pit mine and mill reached planned capacities of 12 Mtpy of ore processed and 197 000 tpy of copper in concentrates in quarter 4 of 2016 (Vale Annual Report, 2016).

Remaining Ore Reserves at 31 December 2017 were (Vale 20-F form report to the US SEC, 2017):
    Proved Reserves - 644.1 Mt @ 0.64% Cu;
    Probable Reserves - 549.3 Mt @ 0.57% Cu;
    TOTAL Reserves - 1193.4 Mt @ 0.61% Cu, with a recovery range of 80 to 90% of contained Cu.

Remaining Ore Reserves and Mineral Resources at 31 December 2019 were (Wheatstone-Vale NI 43-101 Technical Report, 31 December 2019) at a 0.253% Cu
 equiv. cutoff:
    Proved Reserves - 152.7 Mt @ 0.69% Cu, 0.39 g/t Au;
    Stockpile Proved Reserves - 163.4 Mt @ 0.45% Cu, 0.22 g/t Au;
    Probable Reserves - 832.4 Mt @ 0.62% Cu, 0.32 g/t Au;
    TOTAL Reserves - 1148.4 Mt @ 0.60% Cu, 032 g/t Au.
    Measured + Indicated Resources - 193.5 Mt @ 0.61% Cu, 0.31 g/t Au;
    Inferred Resources - 176.1 Mt @ 0.59% Cu, 0.29 g/t Au.
  NOTE: Reserves are exclusive of Resources.

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Travelling - Carajas, Brazil to Copiapo, Chile (via Marabá, Brasilia, Sao Paulo and Santiago) ...................... Friday 8 & Saturday 9 June, 2007.


Andean Overview Workshop ...................... Saturday 9 June, 2007.

An overview workshop was run late on Saturday 9 June, 2007 in Copiapo, Chile at the beginning of the Andean segments of the tour to provide a context to the tectonic, geological and metallogenic setting of the IOCG deposits of the Andes and descriptions of other deposits in the region not on the itinerary. The workshop was led by Dr Carlos Arevalo of the Chilean Servicio Nacional de Geología y Minería (SERNAGEOMIN).

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Punta de Cobre district - Field Workshop ...................... Sunday 10 June, 2007.

A field workshop was run in the Punta de Cobre district. It provided a context setting to the major Candelaria deposit and an impression of the anatomy of an important IOCG district, the geologic, structural and metallogenic setting, the distribution of alteration and different styles of mineralisation. The workshop was led by Dr Carlos Arevalo of the Chilean Servicio Nacional de Geología y Minería (SERNAGEOMIN), an acknowledged expert on the district, its geology and mineralisation.

A series of iron oxide Cu-Au(-Zn-Ag) deposits occur within the Punta del Cobre belt which hosts the Punta del Cobre mine, sensu strictu and La Candelaria the largest of these deposits.
(#Location: 27° 29' 47"S, 70° 15' 6"W)

Other smaller mines that extend over a 25 km north-south interval, centred on La Candelaria and Punta del Cobre include Manuel Antonio Matta, Hernan Videla Lira, Jilguero, Alcaparrosa, Santos, Biocobre and La Plata.

La Candelaria is described in detail in a separate record, which also includes a geological map of the Punta del Cobre district.

The Punta del Cobre belt lies within an Early Cretaceous continental volcanic arc and marine carbonate back-arc basin terrane whose sequences are intruded by Early Cretaceous granitoid plutons that form part of the Chilean Coastal Batholith. The Punta del Cobre belt deposits are found fringing the eastern margin of the batholith within (eg., La Candelaria) or just outside the contact metamorphic aureole (eg., the Punta del Cobre district). Andesitic volcanic and volcaniclastic host rocks are intensely altered by biotite-quartz-magnetite. This style of alteration extends much further to the east of the intrusive contact than the metamorphic mineral associations in the overlying rocks that are clearly zoned outboard. Local areas of intense calcic amphibole veining that overprints all rock types occur within the contact metamorphic aureole.

Chalcopyrite mineralization which crosscuts and thus post-dates all of the major metamorphic and metasomatic assemblages is paragenetically late. Deposits found close to the contact of the batholith and the deeper parts of the sequence in the Punta del Cobre district are characterised by abundant magnetite with associated biotite-quartz alteration, which is overprinted by fracture-controlled calcic amphibole, and chalcopyrite-pyrite mineralization. Potassium feldspar-chlorite and/or biotite ± quartz plus magnetite ± hematite occur in the intermediate parts of the hydrothermal system. Up-section and, in places, laterally, these assemblages grade into pervasive albite-chlorite-calcite-hematite that are spatially associated with Cu-Au mineralization in the more distal portions of the system.

Mineralization is controlled by tectonic structures and their intersection with massive volcanic rocks and overlying volcaniclastic rocks. Isotopic ages of alteration minerals associated with the metallic mineralisation indicates that the bulk of the iron oxide mineralization formed between 116 and 114 Ma, and the main copper-gold mineralization between 112 and 110 Ma, and that hydrothermal activity was coeval with both the emplacement of the Copiapó Batholith and regional uplift. They also imply burial during the mineralisation was no greater than 2-3 Km.

Deposits other than La Candelaria account for another 120 Mt @ 1.5% Cu, 0.2 to 0.6 g/t Au, 2 to 8 g/t Ag, with the larger deposits producing up to 1.5 Mt of ore per annum.

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Mantoverde .......... Monday 11 June, 2007.

The Mantoverde IOCG deposit cluster is located in the Los Pozos district of the Chilean Coastal Cordillera, some 400 km south of Antofagasta and 100 km north of the town of Copiapo in northern Chile. Individual deposits include Manto Verde, Manto Ruso, Manto Monstruo and Monte Cristo (#Location: 26° 33' 57"S, 70° 18' 45"W).

The Los Pozos district falls within a structural segment bound to the east and west by two branches of the north-south trending Atacama Fault Zone (AFZ), an arc-parallel wrench fault system that extends for more than 1000 km along the Chilean coast. The prominent 340 to 345° trending and 40 to 50°E-dipping brittle Manto Verde fault cuts this segment of the AFZ and controls the main Cu-Au ore zone.
  The district is predominantly composed of Jurassic and/or possibly Early Cretaceous andesitic flows and breccias which are interpreted to correlate with either the Jurassic La Negra Formation or the Early Cretaceous Bandurrias Group. These cataclased andesitic volcanics host rocks are intruded by coeval (?) early Cretaceous granitoids of the Chilean Coastal Batholith. Granodiorites and monzonites of the ~130 Ma Las Tazas Plutonic Complex are found in the western part of the Manto Verde deposit, while diorites, monzodiorites, granodiorites and tonalites of the ~127 Ma Remolino Plutonic Complex lie to the east.
  For detail of the regional setting of the Chilean Coastal Cordillera in the Central Andes, see the Central Andes and Bolivian Orocline record.

  The Cu-Au ores of the Los Pozos district are hosted mainly in specularite-dominated tabular breccia bodies (Mantoverde), breccia pipes (Manto Ruso and Manto Monstruo) and stockwork bodies (Monte Cristo). All are associated brittle faults belonging to the AFZ, which were emplaced during an extensional phase of sinistral strike-slip and dip-slip tectonism of that complex Jurassic to early Cretaceous major regional fault system. The Mantoverde deposit consists of three breccia units paralleling the 12 km long Mantoverde Fault for at least 1500 m. The Mantoverde Fault which strikes NNW and dips at 40 to 50° E, connects two major branches of the AFZ.
  The hangingwall Manto-Atacama breccia is a ±100 m wide, matrix supported, specularite rich, hydrothermal breccia that follows the Mantoverde Fault along strike and down dip although it thins with depth. The degree of brecciation and mineralisation decreases gradually outwards from the fault into the andesite wall rocks.
  In contrast the footwall Manto Verde Breccia is a ±20 m wide cataclasite with relatively sharp mineralisation boundaries, with sulphides being hosted by andesitic volcanic and deformed volcaniclastic rocks where hypogene grades locally exceed 1% Cu with around 0.25 g/t Au and Elevated REE concentrations.
  Other Cu-Au deposits also occur on more northwest-trending bends of this fault or the eastern branch of the Atacama fault zone, or at the intersection of the Manto Verde fault and related second order structures.

  Unlike at Candelaria, early district-wide sodic-calcic alteration is absent in the Los Pozos district (Rieger et al., 2010), although widespread sporadic, weak selective albitisation of plagioclase is taken to represent sub-ocean floor spilitisation (Benevides et al., 2007).
  Benevides et al. (2007) interprets the following sequence of regional- to district-scale alteration: (1) Early, sporadic, widespread, but poorly exposed potassium-iron metasomatism which converted both granitoid and volcanic rocks to orthoclase (with subordinate biotite), accompanied by the deposition of magnetite (at 130 to 126 Ma; Chen, 2010), with associated fluorapatite, and minor pyrite, but no chalcopyrite. Homogenisation temperatures of 550 to 460°C are estimated for this phase (Benevides et al., 2007). (2) Subsequent regional scapolitisation in the area between the Atacama fault system and the marginal basin, characterised by marialitic scapolite. Later in this second stage, much of the scapolite was replaced by chlorite as part of an episode of chlorite and sericite (hydrolytic) alteration and veining (at 350 to 300°C), with minor development of hematite, pyrite and trace chalcopyrite (Benevides et al., 2007). This style of alteration grades from a slightly chloritised country rock, to chlorite-quartz veinlets, into chlorite-quartz-cemented hydrothermal breccias with mostly silicified, and occasionally K-altered, fragments, the Breccia Verde as detailed below (Rieger et al., 2010).
  The Mantoverde IOCG sensu stricto deposits, were subsequently emplaced along the main and, more commonly, subsidiary segments of the major Atacama fault system, together with chalcopyrite-bearing, but sub-economic metasomatic magnetite, and copper-barren magnetite-fluorapatite-pyrite bodies. All are hosted by Middle to Upper Jurassic andesites of the La Negra Formation, and by diorite and monzodiorite assigned to the Lower Cretaceous (126 to 120 Ma) Sierra Dieciocho plutonic complex (Benevides et al., 2007). In the immediate Mantoverde district, mineralisation is developed within an intensely fractured structural block, delimited by the subvertical central and eastern branches of the northsouth Atacama fault system, connected obliquely by the northnorthwest trending, brittle, Mantoverde fault (MVF). The MVF is a releasing strike-slip duplex, representing a transfer zone between the central and eastern branches of the Atacama fault system. Most of the Mantoverde deposits are closely associated with a 12 km interval of the MVF and sub-parallel minor faults (Benevides et al., 2007; Rieger et al., 2010; Rieger et al., 2010; Rieger et al., 2009).
  A loosely constrained Early Cretaceous age of mineralisation is based on K-Ar-dating of two samples of hydrothermal sericite from Mantoverde Norte, with minimum ages of 117±3 Ma from an andesite and 121±3 Ma from a granite dyke (Vila et al., 1996). A two-point Re-Os isochron age derived from magnetite yielded 116 Ma, consistent with the alteration ages (Mathur et al., 2002).
  Five main ore-bearing geological groupings or mineralisation zones/styles are recognised at Mantoverde:
1) Manto Atacama, a specularite-cemented, hydrothermal breccia, averaging 80 m in thickness (locally up to 200 m), in the hanging wall (east) of the 40 to 50° east-dipping MVF. It is composed of sub-angular to sub-rounded clasts, a few centimetres across, of mainly andesitic or granitoid igneous rocks in a mineralised, calcite-bearing, coarse-grained specularite matrix, altered to varying degrees by pervasive K feldspar, with more or less intense chlorite, sericite, silica and/or carbonate.
2) Transition zone, an adjacent specularite stockwork zone, immediately to the east and above the Manto Atacama, containing supergene copper oxide and hypogene sulphides, with essentially the same alteration as the Manto Atacama zone.
3) Mantoverde Breccia, a tectonic breccia in the footwall (west) of the main fault plane of the MVF, commonly 20 to 40 m thick. It comprises a rock flour matrix cataclastite, enclosing angular andesite and/or diorite clasts that are commonly a few millimetres to about 10 cm in diameter, with silica, chlorite and minor to moderate K feldspar alteration.
4) Breccia Verde, which is developed on both sides of the MVF, and has a gradational boundary with both the Transition zone and Mantoverde breccia. It is largely barren, and is composed of silicified and pervasively K feldspar altered volcanic and dioritic clasts within a matrix of chlorite-quartz, containing calcite and subordinate sericite, grading outwards into a chlorite-quartz stockwork. It is cut by sparse to moderately frequent sets of K feldspar±quartz, calcite±siderite, specularite,±quartz and sericite veinlets.
5) Magnetite Zone, developed between the Breccia Verde, an intrusion of the 120 to 127 Ma Sierra Dieciocho pluton to the west, and the Mantoverde Breccia to the east, and mainly at depth. It comprises magnetitechlorite-sericite-K feldspar-cemented breccias, with clasts of igneous rock that are predominantly altered to magnetite (including mushketovite), K feldspar and quartz, cut by veinlets of K feldspar±quartz, calcite, sericite and late specularite-calcite. Pyrite and chalcopyrite occur as small patches, disseminations, or in discontinuous veinlets. This zone is representative of the other magnetite-sulphide bodies in the south of the district (Rieger et al., 2010).
  The bulk of the mining to 2010 exploited supergene mineralisation which occurs within the zone of oxidation that consistently persists to a depth of 150 to 250 m below the surface. Mineralisation comprises chrysocolla, brochantite, atacamite, almagre, malachite and copper-bearing hematite, jarosite, and goethite, disseminated in the breccia matrix, as filling of veinlets, and as a patina on, or patches within rock fragments (Rieger et al., 2010; Benevides et al., 2007). Below the base of supergene oxidation, in the Manto Atacama breccia, the specularite-rich matrix contains pyritechalcopyrite, with the chalcopyrite being locally replaced by digenite and bornite. The breccia is cut by veinlets of K feldspar±quartz, tourmaline, or sericite, with late stage calcite and specularite veinlets (Rieger et al., 2010). Neither native gold, nor electrum have been reported from the Mantoverde district, although there is a close correlation between gold and copper grades, and gold has been detected in both, chalcopyrite and pyrite (Rieger et al., 2010). In the northern half of the district, the mineralised tectonic breccias (e.g., Mantoverde Breccia), specularitecemented, hydrothermal breccias (e.g., Manto Atacama), and outer halo of mineralogically identical veins (Transition zone) are characterised by primary specularite and largely developed in the hanging wall of the MVF. In contrast, in the southern half of the district, there are hypogene, crudely tabular, massive bodies of magnetite-pyrite and magnetitechalcopyrite (e.g., the Magnetite zone), predominantly in the footwall of the Mantoverde fault, while massive and irregular bodies of magnetite-apatite±pyrite are developed along the eastern branch of the Atacama fault system (Benevides et al., 2007).
  The hypogene minerals of the Magnetite zone and the magnetite rich bodies in the south of the district comprise (1) magnetite stockworks and disseminations; (2) elongate magnetite-chlorite-sericite-K feldspar-cemented breccias, with igneous rock fragments that are mainly altered to an assemblage of magnetite, K feldspar and quartz; and (3) massive magnetite bodies. Both primary magnetite and mushketovite are present. These magnetite-rich rocks may be either barren, or contain chalcopyrite (and/ or pyrite), mainly as small patches, disseminations and/or discontinuous veinlets (Rieger et al., 2010). Rieger et al. (2010), notes a vertical zonation of mineralisation-associated iron oxides, with a downward transition from upper hematite to magnetite at depth. Tilting and differential rotation across the MVF has resulted in the magnetite mineralisation being exposed to the south, while drilling confirms the continuity of similar magnetite dominant mineralisation at depth immediately below the Mantoverde Breccia ores in the footwall of the MVF in the north. A lateral transition between specularite breccias and magnetite is not evident to the north, although the two are juxtaposed across the MVF, again due to the differential rotation accommodated by that structure. However, in the south, the Manto Atacama is only poorly developed and gives way to magnetite-rich rocks in the hanging wall, footwall and at depth.
  Benevides et al. (2007) conclude that on the basis of textural relationships in both breccias and veins, the crystallisation of specularite and chalcopyrite was coincident. However, Rieger et al. (2010) interpret them to have been emplaced at different times, with the main mineralising events being (1) an early high-temperature iron oxide stage, comprising the bulk of the specularite and magnetite in the district; (2) a sulphide stage responsible for the main copper-gold mineralisation; and (3) a late stage, represented mainly by calcite±specularite and specularite veining, and pervasive carbonatisation. They suggest the iron oxide stage comprises early hematite (hm-I), followed by early magnetite (mt-I) which wholly or partially replaced hm-I (to form mushketovite), accompanied by pervasive K feldspar alteration, minor tourmaline and weak scapolite. Subsequent, variably intense silicification, pervasive sericitisation and minor pyrite (py-I) was followed by a second phase of magnetite (mt-II) and chloritisation, with specularite locally developing rather than mt-II. Rieger et al. (2010) see the sulphide phase as occurring in isolation, after the development of iron oxides (including the specularite breccia matrix), K feldspar, sericite, silica and chlorite. The iron oxide stage of Rieger et al. (2010) corresponds to both the potassium-iron metasomatism, scapolitisation and sericite-chlorite phases of Benevides et al. (2007), described above. Both agree on a terminal phase of alteration, which comprises largely barren calcite veining, variably accompanied by quartz. Rieger et al. (2010) suggest the specularite associated with ore in the north of the district is their hm-I, on the periphery of the mineralised system, while the magnetite in the mineralisation at depth and to the south is both mt-I mushketovite after hm-I, and mt-II, which is deeper and hotter, proximal to the centre of mineralisation.
  Rieger et al. (2010), noted a similar zonation in sulphur isotope signatures. Chalcopyrite in the Mantoverde district shows a wide range in sulphur isotope composition, with δ
34SVCDT of between -6.6 and +10.0‰ (Rieger et al., 2010; Benavides et al., 2007). Systematic variation of these data reflect the spatial distribution of the sulphides in the orebodies and their position relative to the MVF. Sulphur isotopic compositions around 0‰ δ34SVCDT, which are compatible with a magmatic-derived sulphur component, are characteristic of chalcopyrite in orebodies with a close spatial relationship with the MVF in the southern and deeper, or proximal part of the Mantoverde district, representing the more internal parts of the hydrothermal system, while higher δ34SVCDT values of around +6‰, are typical of the northern part of the district, or the shallower, levels, suggesting sulphur contribution from non-magmatic sources in the peripheral portions of the hydrothermal system (Rieger et al. (2010) report the δ34S values for pyrite (0.2 to 9.4‰) to similarly indicate a dominantly magmatic source, although also showing systematic variations across the district, interpreted to reflect both the relative distance from inferred fluid conduits and the level of deposition within the hydrothermal system, and possible non-magmatic components on the peripheries of the system.
  Homogenisation temperatures of between 550 and 460°C have been reported from hypersaline fluid inclusions (32 to 56 wt.% NaCl
equiv.) in quartz coexisting with magnetite (Vila et al.,1996; Benavides et al., 2007 and unpublished sources quoted therein), while the main hypogene specularite breccia sulphide mineralisation is calculated to have formed at between 250 and 180°C (Benevides et al., 2007). Homogenisation temperatures of a number of determinations for late calcite-chalcopyrite and calcite veins from two and three phase fluid inclusions with salinities of from 32 to 40 wt.% NaCl equiv. and 1 to 10 wt.% NaCl equiv. returned values of 360 to 160°C and 260 and 112°C respectively (Vila et al., 1996 and unpublished sources quoted therein).
  The majority of initial
87Sr/86Sr values of altered volcanic rocks and hydrothermal calcite from the district (0.7031 and 0.7060) are similar to those of the igneous rocks of the region. Lead isotope ratios of chalcopyrite are consistent with lead (and by inference copper) derived from Early Cretaceous magmatism. The sulphur, strontium and lead isotope data of chalcopyrite, calcite gangue and altered host rocks respectively, are compatible with the cooling of metal and sulphur-bearing magmatic hydrothermal fluids (with deposition at <350°C) that mix with meteoric waters or seawater at relatively shallow crustal levels. They suggest, input of additional exotic sulphur is likely, though not essential, for the deposition of copper mineralisation, occurring mainly in the shallow, and distal parts of the ore system.
  Magnetites from the Mantoverde deposit have Os and Re concentrations of 11 to 17 ppt and 4 to 6 ppb, respectively. An initial
187Os/188Os ratio calculated from these is ~0.20, compared to ~0.36 to 0.33 at Candelaria, and from 0.20 to 0.41 for magmatic magnetite from Early Cretaceous batholithic intrusions in the district. These relatively radiogenic ratios are taken to represent a mixture of mantle and crustal components in the ores and batholitic rocks (Mathur et al., 2002).
  Benevides et al. (2007) interpret isotopic and temperature data they quote from regional and shallow hematitic mineralisation to indicate their potassium-iron metasomatism stage was caused by magmatic fluids, possibly products of the second boiling of granitoid magmas, such as the nearby 126 to 120 Ma Sierra Dieciocho complex. However, they suggest the data indicates a change, probably during the scapolitisation and sericitechlorite phases, possibly due to evaporite derived brines mobilised by marginal basin inversion reflected in the regional sodic alteration event, but even more likely during the deposition of chalcopyrite ore.

The copper resource at Mantoverde to the 800 m level was estimated prior to mining at 120 Mt @ 0.72% Cu (0.2% Cu cutoff) including a mineable reserve of 85 Mt @ 0.82% Cu.

In the mid to late 1990s, exploration identified:
    - a Cu oxide resources of the order of 180 Mt @ 0.5% Cu overlying
    - a hypogene sulphide resource of 440 Mt @ 0.56 % Cu, 0.12 g/t Au at a 0.20 % Cu cut-off.

Remaining mineral resources at 31 December, 2014 at a 0.2% Cu
Total cutoff (Anglo American Annual Report, 2015) were:
  Oxide ore (Heap Leach)
      Measured + indicated resource - 67.4 Mt @ 0.35% Cu
Acid Sol.,
      Inferred resource - 2.6 Mt @ 0.29% Cu
Acid Sol.,
  Oxide ore (Dump Leach)
      Measured + indicated resource - 25.3 Mt @ 0.16% Cu
Acid Sol.,
      Inferred resource - 2.3 Mt @ 0.16% Cu
Acid Sol..

The mine is operated and controlled by Minera Mantos Blancos S.A., which was initially an Anglo American Group company, but was sold to be part of Mantos Copper S.A. in 2015.

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Candelaria .......... Tuesday 12 June, 2007.

The La Candelaria deposit is located some 20 km south of Copiapo. Like Mantoverde and Mantos Blancos it is localised near the Atacama Fault Zone within the Central Andean Coastal Cordillera (or Central Andean Coastal Belt) of northern Chile, and the Chilean Iron Belt (#Location: 27° 30' 55"S, 70° 17' 19"W).

The Central Andean Coastal Belt comprises a Late Jurassic to Early Cretaceous volcano-plutonic belt (Sillitoe, 2003) that is is characterised by voluminous tholeiitic to calc-alkaline volcanic piles and plutonic complexes of primitive mantle origin gabbro to granodiorite. It is associated with an extensional to transtensional event when the underlying crust was attenuated and subjected to high heat flow. All of the intrusive rocks are oxidised and belong to the magnetite series (Charrier et al., 2007; Sillitoe, 2003 and sources cited therin). This belt contains a series of IOCG sensu stricto Cu-Au and other iron oxide-alkali altered mineralisation, such as magnetite-apatite deposits, distributed over a north-south interval of >500 km, with related manto deposits further to the north and south.
  The Candelaria deposit lies within the almost continuously altered, but discontinuously mineralised, Punta del Cobre district, which occupies an area of ~20 Ă— 5 km along the eastern margin of the 120 to 97 Ma Copiapó batholith (Williams et al., 2005; Arévalo, 2006). This batholith, which lies immediately to the east of older Jurassic to Early Cretaceous (>125 Ma) granitoids, comprises the larger ~119 Ma La Brea pluton and ~110 Ma San Gregorio plutonic complex close to Candelaria, one of a group of smaller, similarly aged (111 to 106 Ma) intrusions on the eastern margin of the Copiapó batholith (Arévalo et al., 2006).
  The district is characterised by a ~12 Ă— 5 km envelope of sodic±calcic alteration, superimposed upon volcanic, sedimentary and intrusive rocks. This alteration is manifested as either (1) albite or sodic-plagioclase, and/ or (2) scapolite, with or without calcic amphibole (mainly actinolite, ferro-actinolite, or actinolitic hornblende), pyroxene and/or epidote. Voluminous sodic scapolite-rich assemblages, commonly (but not always) associated with calcic-amphibole and/or pyroxene±epidote±andradite, are usually stratabound, and within the Abundancia and upper Punta del Cobre formations, largely above the ore zone, possibly representing metamorphosed evaporitic beds in these Early Cretaceous units. These rocks also host small magnetite±chalcopyrite-pyrite mantos.
  In contrast, where sodic alteration is predominantly albitisation, it is more commonly discordant and pervasive, occurring in igneous rocks, locally with associated minor pyrite±trace chalcopyrite and/or veinlets and disseminations of hematite (Marschik and Fontboté, 2001). Some of the early albite may be due to spilitisation of volcanics, rather than alteration (Ullrich and Clark, 1999). The overall sodic-calcic zone is enveloped by rocks that were affected by propylitisation, and/or contact thermal metamorphic skarn/hornfels alteration related to the Copiapó batholith (Marschik et al., 2003). The contact between the sodic and thermal metamorphism is gradational (Marschik and Fontboté, 2001).
  The more extensive thermal metamorphic aureole of the batholith, extends over a length of >20 km and width of 2 to 5 km from the contact, producing hornfels and skarns with mineralogies that are dependent upon the host rock, and distance from the contact. Skarn minerals in the thermal aureole include proximal diopsidic-hedenbergite, pyroxenescapolite±andraditic garnet, to distal biotite/quartz/pyroxene±epidote±K feldspar hornfels (Marschik et al., 2000). Early pervasive albitisation accompanied the introduction of widespread specularite, occurring in dilational fractures and open spaces. It preceded, and encloses cores of potassic alteration in the district, and occurs below and peripheral to the 116 to 114 Ma pervasive, brown biotite-quartzmagnetite stage (Mathur et al., 2002), and associated almandine±cordierite alteration that accompanied early barren, more intense magnetite mineralisation. This early magnetite was mainly composed of mushketovite after hematite, indicating a shift to more reducing conditions and/or higher temperatures (Marschik and Fontboté, 2001), and more than one pulse and source of iron oxide alteration (Mathur et al., 2002).
  The deposit lies near the core of the district wide Tierra Amarilla antiform, part of the Paipote fold and thrust system, 20 km to the east of the main Atacama Fault Zone (AFZ). The ore zone is relatively flat lying, broadly concordant with the host sequence of coarse-grained volcaniclastics and massive volcanic flows, breccias and tuffaceous rocks, and a broad, similarly flat-dipping shear zone, with hanging wall evaporite-bearing -limestone, shale and epiclastic rocks.
  The deposit is capped by the base of the 2 km thick Chañarcillo Group (comprising the Abundancia, Nantoco, Totoralilio and PabellĂ³n formations) that was laid down immediately prior to the mineralisation, and hence the depth of formation is assumed to correspond to the thickness of that unit.
  The early pre-ore alteration described thus far, is overprinted by a younger (112 to 110 Ma; Mathur et al., 2002), more areally restricted, ore-related intense potassic±calcic alteration (K feldspar-biotite-amphibole) phase, although Re-Os ages of 115 to 114 Ma for molybdenite (Mathur et al., 2002) and two ~111 Ma dates for amphibole and biotite associated with chalcopyrite (Ullrich and Clark, 1999; Arévalo 1999), could represent the main sulphide phase. Fluids associated with the main stage sulphide mineralisation are hypersaline and CO
2-bearing (Marschik and Fontboté, 2001).
  The ore at Candelaria is associated with a complex, multistage event, and occurs: (1) as bodies that are roughly concordant with stratification and comprise replacements and pore filling (mantos); (2) in the matrix to hydrothermal breccias and pseudobreccias; (3) superimposed on massive magnetite replacement bodies; (4) as discontinuous veinlets and stringers in altered host rock; and (5) as massive veins (Marschik et al., 2000; Marschik and Fontboté 2001). The breccia ores represent intervals of high copper grade.
  They generally occur as irregular zones, and sometimes lens-shaped bodies, 1 to 3 m thick, which are concordant with stratification. They comprise a chalcopyrite-pyrite-magnetite matrix between brecciated, biotite altered, metavolcanic rock clasts. Clasts are very angular to locally rounded, from a few to tens of cm across. The margins of breccia bodies are diffuse and defined by a decrease in the amount of sulphide matrix. Individual clast borders are sharp and form jigsaw patterns, indicating limited transport and rotation after fragmentation, although some may represent replacive pseudo-breccia. The concordant lens-like replacement and pore-infill manto bodies resemble breccias where sulphides fill pore spaces between angular to rounded volcanic and sedimentary breccia clasts. Networks of veins as they are enlarged also form a pseudo-breccia texture (Arévalo et al., 2006).

Candelaria Setting

The geology and superimposed alteration of the Candelaria-Punta del Cobre district of northern Chile (a). All of the known deposits lie within the Paipote fold and thrust system (which includes the steeply dipping Ojancos-Florida Shear Zone separating the Copiapó Batholith from the host Punta del Cobre Formation west of the Candelaria mine; as well as the Paipote, Ladrillos and Cerrillos thrusts and Tierra Amarilla anticline). This structural corridor is part of the regional Chivato fault system, which is ~20 km east of the main Atacama Fault Zone (top-left). The largest deposits, Candelaria and Carola are associated with potassic alteration within the broader sodic-calcic zone. Note the potassic core surrounding the Candelaria deposit is the exposure within the open pit mine, not an original pre-mine outcrop. The sodic-calcic alteration includes both stratabound alteration within the overlying Abundancia and Nantoco formations (possibly related to hydrothermal modification of evaporite-bearing beds within those units), and pervasive, discordant alteration within the volcanic and volcaniclastic rocks of the Punta del Cobre Formation (after Marschik et al., 2000; Marschik and Fontboté, 2001; Arévalo et al., 2006).
East-west cross section through the Candelaria mine (b). The diagram illustrates the distribution of the manto-like early magnetite and the outline of the overprinting +0.4% Cu orebody, both of which are largely confined to the coarse stratified volcaniclastic breccias and interlayered massive andesites of the Lower Andesites, and the flat-lying Candelaria Shear that caps this unit and the ore deposit. Note the concentration of steep faults near the core of the deposit. Distribution of alteration within the section comprises albite/sodic plagioclase-quartz-biotite-magnetite±K feldspar minor Ca amphibole below the orebody; biotite-quartz-magnetite±K feldspar and abundant Ca amphibole (largely actinolite) within the ore zone; biotite-quartz- almandine±cordierite and common Ca amphibole within the upper Candelaria Shear; biotite-amphibole with K or Na feldspar in the tuffs and volcaniclastic sediments and Upper Andesites above the shear; and scapolite±quartz±pyroxene±Ca amphibole within the overlying Abundancia Formation (after Marschik et al., 2000; Marschik and Fontboté, 2001; Arévalo et al., 2006).



  Copper ore is associated with magnetite and/or hematite and is dominantly composed of chalcopyrite and pyrite, with gold occurring as small inclusions in the chalcopyrite, within micro-fractures in pyrite and as a mercury-goldsilver alloy (Williams et al., 2005).
  At the deeper levels in the mineralised system, chalcopyrite has a close spatial association with calcic amphibole (mainly actinolite) in an assemblage that also includes biotite, K feldspar±epidote±sodic plagioclase. Magnetite is ubiquitous, occurring as massive bodies, with or without superimposed sulphide mineralisation, although hematite is rare. Intermediate levels, are characterised by potassic alteration (biotite and/or K feldspar), with or without local developments of calcic amphibole±epidote, sodic plagioclase, and/or local anhydrite. In shallower and distal parts of the system, chlorite is developed at the expense of biotite and amphibole, and albite, chlorite and carbonate alteration increases in intensity. Peripheral to the mineralisation, hematite becomes the dominant iron oxide (Williams et al., 2005). This alteration pattern is complicated by the influence of contact metamorphism related to components of the 120 to 97 Ma Copiapó batholith, as described above. The increasing sodiccalcic alteration in the upper levels of the Candelaria pit area, above the potassic and chloritic zones, may reflect the overlying evaporite-bearing Abundancia Formation (Marschik and Fontboté, 2001).
  Marschik and Fontboté (2001), describe a paragenetic sequence over-printing the early, district scale stage of iron metasomatism and associated pervasive albite alteration, comprising:
  1) a high temperature (600 to 500°C), pre-ore iron oxide stage, comprising pervasive magnetite (mushketovite)-quartz-biotite;
  2) the main sulphide ore stage at 500 to 300°C, represented by chalcopyrite and pyrite; and
  3) the late stage at <250°C, with hematite-calcite and locally minor sulphides.
  Sulphides from Candelaria and some other occurrences in the Punta del Cobre district yielded δ
34SCDT values largely between -3.2 and +3.1‰, with some as high as 7.2‰ for late stage mineralisation, or up to 6.8‰ in the marginal parts of the system (Rabbia et al., 1996; Ullrich and Clark, 1999; Marschik and Fontboté 2001). Marschik and Fontboté (2001) interpret these isotopic data to be consistent with a dominantly magmatic source for sulphur, with a minor contribution during their stage 2, but definite influence in stage 3, from a peripheral evaporite-bearing sedimentary host sequence (e.g., Mathur et al., 2002; Marschik and Fontboté 2001; Ullrich and Clark, 1999).
  Barton et al. (2005) reported that unpublished Sr isotope data for altered and host rocks in Candelaria-Punta del Cobre district, during both early sodic-calcic and late potassic stages of hydrothermal activity imply there are large contributions of non-igneous Sr, implying the ore systems involved influx of fluids from outside the local batholithic granitoids.
  Arévalo et al. (2006) note that the sulphide ages quoted above, corresponds closely to that of the San Gregorio plutonic complex of the Copiapó batholith, supporting a magmatic origin for the main sulphide mineralisation. They interpret the sulphide mineralisation to have been the product of magmatic fluids of a cooling hydrothermal system, emplaced during synplutonic deformation and dilation at the ductile-brittle transition, in the thermal aureole of the San Gregorio plutonic complex.
  An isochron calculated by Re/Os ratios from hydrothermal magnetite and sulphides at Candelaria and the small satellite deposit Bronce, constrains initial
187Os/188Os ratios of 0.36±0.10 and 0.33±0.01 respectively. These values are broadly similar to the calculated initial 187Os/188Os ratio for magmatic magnetite in nearby batholithic rocks that range from 0.20 to 0.41. These relatively radiogenic ratios also represent a mixture of mantle and crustal components in both the ores and batholitic rocks (Mathur et al., 2002).

The pre-mining mineable reserve comprised 470 Mt @ 0.95% Cu, 0.22 g/t Au, 3.1 g/t Ag within a geological resource of 600 Mt at a similar grade.

The mine is operated by the Freeport subsidiary Compania Contractual Minera Candelaria.

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Mantos Blancos .......... Wednesday 13 June, 2007.

The Mantos Blancos copper-silver deposit is located in the Coastal Range of northern Chile, some 45 km NE of the Pacific coastal city of Antofagasta in northern Chile (#Location: 23° 25' 52"S, 70° 3' 22"W).

It lies within the Atacama Fault Zone and is hosted by a Triassic sequence of acid volcanics, mainly rhyolites and dacite which dip at 10 to 45° SE and, cut upper Palaeozoic sediments and metasediments. These are in turn overlain by Jurassic clastics and limestones, Jurassic andesites and Cretaceous andesites and dacites. In the mine area the host volcanics are intruded by a sill like sub-volcanic andesite body, by sheets of dacite and abundant andesite dykes.

The Mantos Blanco mineralisation displays two superimposed hydrothermal events, namely:

i). an older phyllic alteration probably related to felsic magmatic-hydrothermal brecciation at ~155 Ma, and
ii). a younger (141-142 Ma) potassic, propylitic and sodic alteration phase, coeval with dioritic and granodioritic stocks and sills, and dioritic dykes.

The principal ore formation is genetically related to the second hydrothermal event, and comprises hydrothermal breccias, disseminations and stockwork-style mineralisation, associated with sodic alteration.

The ore stage alteration is dominated by albitisation and silicification that is distributed homogeneously through the volcanic sequence. Specular hematite is found in the barren upper levels and red hematite in the mineralised zones. These represent four principal alteration types, namely:

i). Na metasomatism manifested by albitisation of feldspars as well as albite veining and pore filling;
ii). incipient to intense addition (or locally removal) of Mg and Fe, reflected by chloritisation or bleaching;
iii). intense hematisation in the form of disseminated and stringer specularite and by intense pervasive red hematisation to many of the rocks within the deposit; and
iv). silicification, represented by quartz phenocrysts, microcrystalline aggregates in the groundmass and as occasional veinlets and amygdule fillings.

The mineralisation occurs as irregular bodies of oxide and sulphide copper with economically significant associated silver. The oxide minerals atacamite and chrysocolla are common in the upper levels of the sulphide body associated with faulting and intense brecciation.

The hypogene sulphide assemblages have distinctive vertical and lateral zoning, centred on magmatic and hydrothermal breccia bodies, which are interpreted to constitute feeders to the main mineralisation which is largely distributed in irregular, lenticular bodies roughly parallel to stratification. A barren pyrite root zone is overlain by pyrite-chalcopyrite, and followed upwards and laterally by chalcopyrite-digenite or chalcopyrite-bornite. A digenite-supergene chalcocite assemblage characterises the central portions of high-grade mineralisation in the breccia bodies. Silver is found in the lattices of both the oxide and sulphide minerals and correlated with the Cu grade.

Economic grade ore is found over an interval of 3 x 1.5 km and to a depth of 450 m. In 1995 the pre-mining resource was calculated at 170 Mt, of which 91 Mt were oxide ore @ 1.4% Cu and   89 Mt of sulphide ore @ 1.6% Cu and 17 g/t Ag.

Remaining mineral resources at 31 December, 2014 at a 0.2% Cu
Total cutoff (Anglo American Annual Report, 2015) were:
  Sulphide ore (flotation)
      Measured + indicated resource - 90.8 Mt @ 0.64% Cu at a 0.2% Cu
Total,
      Inferred resource - 22 Mt @ 0.56% Cu
Total,
  Oxide ore (Vat and Heap Leach)
      Measured + indicated resource - 18.4 Mt @ 0.43% Cu
Acid Sol.,
      Inferred resource - 16.3 Mt @ 0.29% Cu
Acid Sol.,
  Oxide ore (Dump Leach)
      Measured + indicated resource - 11.0 Mt @ 0.17% Cu
Acid Sol.,
      Inferred resource - 70.7 Mt @ 0.18% Cu
Acid Sol..

The mine is operated by Minera Mantos Blancos S.A., which was initially an Anglo American Group company, but was sold to be part of Mantos Copper S.A. in 2015.

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The summaries above were prepared by T M (Mike) Porter from a wide range of sources, both published and un-published.   Most of these sources are listed on the "Tour Literature Collection" available from the IOCG 07 Tour options page.

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For more information contact:   T M (Mike) Porter, of Porter GeoConsultancy   (mike.porter@portergeo.com.au)

This tour was designed, developed, organised, managed and escorted by
T M (Mike) Porter of Porter GeoConsultancy Pty Ltd.

Porter GeoConsultancy Pty Ltd
6 Beatty Street
LINDEN PARK, 5065
South Australia
Mobile: +61 422 791 776



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