Nickel Laterit (Waheed Ahmad)

NICKEL LATERITES ⎯ A TRAINING MANUAL CHEMISTRY, MINERALOGY & FORMATION OF Ni LATERITES

Waheed Ahmad July 30, 2001

LIST OF CONTENTS

CHEMISTRY Valencies; Metals & oxides; Important oxides; Mineralogical formulas in terms of oxides; Atomic weights; Ionic radii; Ionic replacements; Chemical mobility of elements; Chemical composition in terms of elements; Chemical composition on terms of oxides; Conversion of elemental percentages to oxide percentages. ROCK-FORMING MINERALS Silicon tetrahedron; Primary mafic minerals; Primary feldspars; Primary spinels; Secondary minerals; Nickel hydrosilicates; Garnierites; Pyroxene composition fields; Iron oxides; Chromite; Asbolite OLIVINE GROUP Structure; Chemical composition; Formation of olivines; Nickel in the olivines; Alteration and weathering; SERPENTINISATION OF OLIVINES Conditions for serpentinisation; Serpentinisation by volume change; Serpentinisation at constant volume; Dissociation of serpentine ULTRAMAFIC ROCKS Terminology; Classification; Field occurrence; Nickel in ultramafic rocks MAGMATIC DIFFERENTIATION Continuous-reaction series; Discontinuous-reaction series PHASE DIAGRAMS Ice-water-vapour system; 2-component system with solid solution; 2-component system without solid solution; SiO2-MgO system; Phase chemistry in PT Inco furnaces WEATHERING Processes of change; Chemical weathering; Hydrolysis; Oxidation; Hydration; Solution; Factors that influence chemical weathering; Stability of minerals; pH conditions; Eh conditions; Rate of removal; Role of climate; Role of topography WEATHERING OF ULTRAMAFICS CLASSIFICATION OF SOILS LATERITIC SOILS Definition; Lateritic profile; Limonite zone; Intermediate zone; Saprolite zone; Rate of laterisation

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CHEMISTRY Chemical valencies of elements Valence (or valency) is the capability of an atom to combine in particular proportion with another atom. Cations: Generally metallic elements (Na, K, Cu, Mg, Ca, Fe, Mn, Ni, Co, Pb) Anions: Generally non-metallic elements (O-2, S-2, F-, Cl-, Br-, I-) Valency Other atoms Monovalent 1 Divalent 2 Trivalent 3 Tetravalent 4 Pentavalent 5 Hexavalent 6

Examples H, Li, Na, K, Cu1 Mg, Ca, Co, Cu2, Fe2, Mg, Mn2, Ni, O, Pb2, Zn). Bi, Cr3, Al, Fe3, Sb3 C, Si, Ti, Zr, Mn4, Pb4 P, As, Sb5 Cr6

Many atoms such as Fe, Mn, Cu, Cr, Pb, and Sb have multiple valencies: they are capable of combining with other atoms in different proportions.

Metals and oxides Metals Al (Aluminium) C (Carbon) Ca Co Cr Fe

(Calcium) (Cobalt) (Chromium) (Iron)

H K Mg Mn

(Hydrogen) (Potassium) (Magnesium) (Manganese)

Na Ni P S Si

(Sodium) (Nickel) (Phosphorous) (Sulphur) (Silicon)

Oxides and names Al2O3 (Alumina) CO (Carbon monoxide) CO2 (Carbon dioxide) CaO (Calcium oxide) CoO (Cobalt oxide) Cr2O3 (Chrome) FeO (Ferrous oxide) ⎯ Iron protoxide Fe2O3 (Ferric oxide) ⎯ Iron sesquioxide H2O (Water) K2O (Potassium oxide) MgO (Magnesia) MnO (Mangannous oxide) MnO2 (Mangannic oxide) ⎯ Pyrolusite Na2O (Sodium oxide) NiO (Nickel oxide) P2O5 (Phosphorous pentaoxide) SO2 (Sulphur dioxide) SiO2 (Silica)

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Important oxides in mineralogical studies Univalent metal:

H2O K2O Na2O Cu2O

= Water = Potassium oxide = Sodium oxide = Cuprous oxide

Divalent metal:

MgO MnO NiO CoO CaO CuO FeO ZnO

= Magnesia = Manganous oxide = Nickel oxide = Cobalt oxide = Calcium oxide = Cupric oxide = Iron protoxide (ferrous iron) = Zinc oxide

Trivalent metal:

Cr2O3 Al2O3 Fe2O3

= Chrome = Alumina = Iron sesquioxide (ferric iron)

Tetravalent metal: SiO2 CO2 MnO2 SO2 TiO2

= Silica = Carbon dioxide = Mangannic oxide (Pyrolusite) = Sulphur dioxide = Titanium oxide

Pentavalent metal: P2O5

= Phosphorous oxide

Representation of mineralogical formulae in terms of oxides Most rock-forming minerals can be represented in terms of oxides: Mineral name

Formula

Forsterite olivine Fayalite olivine Enstatite/Bronzite Serpentine Talc Magnetite Hematite Goethite Limonite Chromite

Mg2SiO4 Fe2SiO4 MgSiO3 H4Mg3Si2O9 H2Mg3(SiO3)4 Fe3O4 Fe2O3 H2Fe2O4 H6Fe4O9 FeCr2O4

Representation by oxides = = = = = = = = = =

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2MgO.SiO2 2FeO.SiO2 MgO.SiO2 3MgO.2SiO2.2H2O 3MgO.4SiO2.H2O FeO.Fe2O3 Fe2O3 Fe2O3.H2O 2Fe2O3.3H2O FeO.Cr2O3

Atomic weights Atomic weights of elements are relative weights in comparison to Hydrogen as 1.0 or Oxygen as 16.0. H 1.0

O 16.0

Mg 24.3

Al 26.9

Si 28.0

Ca 40.0

Cr 52.0

Mn 54.9

Fe 55.8

Ni 58.6

Co 58.9

Ionic radii of Common rock-forming elements Ions are measured in terms of their radii. The unit of radius is an angstrom. One angstrom is one ten millionth of a millimetre (1 ºA = 10-7 mm or 10-10 m): Si 0.42

Al 0.51

Cr6 0.52

Mn4 0.60

Co3 0.63

Cr3 0.63

Fe3 0.64

Mg 0.66

Fe2 0.74

Zr 0.79

Mn2 0.80

Na 0.97

Ca 0.99

K 1.33

O 1.40

S-2 1.84

Ti 0.68

Ni 0.69

Co 0.72

Ionic replacements Certain cations can replace each other to form isomorphous series. The replacement requires the following conditions to be met: • • • •

They have about the same ionic radii The minerals they form have the same crystal structure The difference in charge is no more than one valency [if an element of a smaller positive charge comes in, then an additional cations are needed to balance the charge in the lattice] The conditions are suitable (the crystals and the remaining liquid are in contact with each other and the composition of the liquid has an opportunity to alter the composition of the crystals)

Fe replaces Mg in olivines Magnesian and ferrous olivines form a continuous series of solid solutions. Both have orthorhombic crystal structure, both are divalent, and the ionic radii of magnesium and ferrous iron are fairly close (0.66°A and 0.74°A respectively). Mg2SiO4 Fe2SiO4 (Mg,Fe)2SiO4

Magnesian Olivine, Forsterite Iron Olivine, Fayalite Solid solution of Mg and Fe olivines

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A typical olivine in ultramafic rocks is Fo90Fa10. Such an olivine is a solid solution of 90% Forsterite and 10% Fayalite. The proportion of oxides and elements in such an olivine can be written out as follows: Fe = 5.48% FeO = 7.05% Fe2SiO4 = 10% (Fayalite) Atoms of Fe = 0.09821

Mg = 31.10% MgO = 51.57% Mg2SiO4 = 90% (Forsterite) Atoms of Mg = 1.2798

Fe:Mg = 1:5.7 FeO:MgO = 1:7.3 Fa:Fo = 1:9 Fe:Mg = 1:13

Fe replaces Mg in pyroxenes Magnesian and ferrous pyroxenes form a continuous series of solid solutions. Both end members have similar crystal structure, are divalent, and the ionic radii of magnesium and ferrous iron are fairly close (0.66°A and 0.74°A respectively). MgSiO3 FeSiO3 (Mg,Fe)SiO3

Magnesian pyroxene, Enstatite Iron pyroxene, Ferrosilite Solid solution of Mg and Fe pyroxenes, Hypersthene

A pyroxene from Poro, New Caledonia, has the composition: (Mg 0.9 Fe The above composition can be expressed in different ways, as follows: Fe = 5.39% FeO = 6.94% FeSiO3 = 10% (Ferrosilite) Atoms of Fe = 0.09659

Mg = 21.14% MgO = 35.05% MgSiO3 = 90% (Enstatite) Atoms of Mg = 0.86996

0.1)

SiO3.

Fe:Mg = 1:3.92 FeO:MgO = 1:5.05 Fa:Fo = 1:9 Fe:Mg = 1:9

Al replaces Si in plagioclases Aluminium can replace silicon atoms in plagioclase feldspars since both have nearly the same ionic radii (Si=0.42°A; Al=0.51°A). However, aluminium is trivalent (Al+++) while silicon is tetravalent (Si++++). Thus, if a cation of lower positive charge replaces a cation of higher positive charge, another cation must be introduced into the crystal lattice to balance the deficit. In the case of albite (NaAlSi3O8), every fourth silicon ion (++++) is replaced by aluminium (+++) and sodium ions (+). In the case of anorthite (CaAl2Si2O8), every second silicon ion is replaced by aluminium and a divalent calcium ion. Ni replaces Mg in olivines During the crystallisation of olivines, small quantities of nickel present in the magma can replace Mg in the lattice structure. Such replacement is possible for the following reasons:

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• • •

Ionic radii of Mg and Ni are fairly close (Mg=0.66°A; Ni=0.69°A) Their valencies are the same (Mg++ ; Ni++) Their crystals belong to the same orthorhombic system

The solid solution of Ni and Mg in olivines can be perceived as the solid solution of nickel-olivine in magnesian-olivine. If the proportion of nickel in forsterite is 0.3%Ni, then the relative quantities of nickel and magnesium in the rock can be worked out in terms of elements, oxides or the type of olivine, as follows: Ni = 0.3% NiO = 0.382% Ni2SiO4 = 0.554% (Ni-olivine)

Mg = 34.368% MgO = 56.983% Mg2SiO4 = 99.446% (Fo)

In terms of numbers of atoms:

Ni:Mg = 1:115 NiO:MgO = 1:150 Ni-ol:Mg-ol = 1:180 Ni:Mg = 1:275

Nickel and iron replace Mg in olivines An olivine from Poro, New Caledonia, is: (Mg1.813 Fe0.18 Ni0.007) SiO4. The olivine can be assumed to be a solid solution of three different olivines. The proportion of various elements, oxides, and olivines is given below: Ni-olivine

Fe-olivine

Mg-olivine

0.007/2 = 0.35% Ni = 0.28% NiO = 0.356%

0.18/2 = 9.0% Fe = 6.855% FeO = 8.822%

1.813/2 = 90.65% Mg = 30.07% MgO = 49.856%

Nickel and iron replace Mg in pyroxene An pyroxene from Poro, New Caledonia, is: (Mg0.908 Fe0.09 Ni0.002) SiO3. The pyroxene can be assumed to be a solid solution of three different pyroxenes. The proportion of various elements, oxides, and pyroxenes is given below: Ni-pyroxene

Fe-pyroxene

Mg-pyroxene

0.002 / 1 = 0.20% Ni = 0.1136% NiO = 0.1446%

0.09 / 1 = 9.0% Fe = 4.866% FeO = 6.263%

0.908 / 1 = 90.8% Mg = 21.3795% MgO = 35.447%

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Chemical mobility of elements in ground water Many metallic elements are soluble in ground water although the solubilities are extremely low compared to common salts. Generally speaking, solubilities increase at higher temperatures and in the presence of low-pH (acidic) waters. The common acid in ground water is Humic Acid that is derived from the decaying of ground vegetation. For components of ultramafic rocks, reasonable solubilities exist at temperatures and conditions commonly found in the tropical climate. Mobilities of elements commonly found in the mafic and ultramafic rocks are classified as follows: • • •

Highly soluble and highly mobile Non-soluble and non-mobile Limited solubility and limited mobility

Highly soluble and mobile elements: Ca, Na, Mg, K, Si Easily leached out of the weathering profile. Highly soluble in tropical ground waters (slightly acidic) Non-soluble (residual) elements: Al, Fe+++, Cr, Al, Ti, Mn, Co Insoluble in ground water. These elements make up the bulk of the residual soil. Elements with limited solubility and mobility: Ni, Fe++ Partly soluble in acidic groundwater. Insoluble in the presence of more soluble elements (Si, Mg). Partial solubility of Ni leads to supergene (secondary) enrichment. Relative Mobilities of elements (given below in the order of decreasing values) Ca++ > 3.0 • • • • • • •

Na+ > 2.4

Mg++ > 1.3

K+ > 1.25

SiO2 > 0.20

Fe2O3 > 0.04

Ca++, Mg++, Na+ are highly soluble and readily lost during leaching K+ is readily leached but fixed again as K-bearing clays Fe++ (ferrous iron) is readily leached and can be mobile Si++++ is slowly lost under leaching conditions (crystalline quartz has one-tenth the solubility of amorphous silica). Ti++++ is generally immobile, except if released as Ti(OH)4 Fe+++ (ferric iron) is immobile under oxidising conditions Al+++ is immobile in the normal pH range of 4.5 – 9.5

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Al2O3 0.02

Chemical composition in terms of elements Relative proportions of elements in compounds are expressed in percentages: H2O (water)

Total weight = 1+1+16 = 18 H = 2 / 18 = 11.1% O = 16 / 18 = 88.9%

SiO2 (silica)

Total weight = 28 + 16 + 16 = 60 Si = 28 / 60 = 46.7% O = 32 / 60 = 53.3%

FeO (iron protoxide)

Total weight = 55.8 + 16 = 71.8 Fe = 55.8 / 71.8 = 77.7% O = 16 / 71.8 = 22.3%

Fe2O3 (iron sesquioxide)

Total weight = 55.8 + 55.8 + (3 x 16) = 159.6 Fe = 111.6 / 159.6 = 69.9% O = 48 / 159.6 = 30.1%

Hematite

Fe3O4 (FeO.Fe2O3) Magnetite

2Fe2O3.3H2O Limonite

Total weight = 55.8 + 55.8 + 55.8 + (4 x 16) = 231.4 Fe = 167.4 / 159.6 = 72.3% O = 64 / 231.4 = 27.7% Total weight = (4 x 55.8) + (9 x 16) + (6 x 1) = 373.2 Fe = 223.2 / 373.2 = 59.8% O = 144 / 373.2 = 38.6% H = 6 / 373.2 = 1.6%

Mg2SiO4 (olivine)

Total weight = 24.3 + 24.3 + 28.0 + (4 x 16) = 140.6 Mg = 48.6 / 140.6 = 34.6% Si = 28 / 140.6 = 19.9% O = 64 / 140.6 = 45.5%

MgSiO3 (pyroxene)

Total weight = 24.3 + 28 + (3 x 16) = 100.3 Mg = 24.3 / 100.3 = 24.2% Si = 28 / 100.3 = 27.9% O = 48 / 100.3 = 47.9%

H4Mg3Si2O9

Total weight = 4 + (3 x 24.3) + (2 x 28) + 144 = 276.9 H = 4 / 276.9 = 1.4% Mg = 72.9 / 276.9 = 26.3% Si = 56 / 276.9 = 20.2% O = 144 / 276.9 = 52.0%

Serpentine

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Chemical composition in terms of oxides Atomic weights of important oxides SiO2 MgO FeO Fe2O3 NiO H2O

= 28 + 32 = 24.3 + 16 = 55.8 + 16 = (2 x 55.8) + (3 x 16) = 58.6 + 16 = 1 + 1 + 16

= 60 = 40.3 = 71.8 = 159.6 = 74.6 = 18

Forsterite olivine

Mg2SiO4

=

2MgO.SiO2 = 140.6 MgO = 80.6 / 140.6 = 57.3 SiO2 = 60 / 140.6 = 42.7%

Fayalite olivine

Fe2SiO4

=

2FeO.SiO2 = 203.6 FeO = 143.6 / 203.6 = 70.5% SiO2 = 60 / 203.6 = 29.5%

Enstatite/Bronzite

MgSiO3

=

MgO.SiO2 = 100.3 MgO = 40.3 / 100.3 = 40.2% SiO2 = 60 / 100.3 = 59.8%

Serpentine

H4Mg3Si2O9 =

3MgO.2SiO2.2H2O MgO SiO2 H2O

= 276.9 = 43.7% = 43.3% = 13.0%

Talc

H2Mg3(SiO3)4 =

3MgO.4SiO2.H2O MgO SiO2 H2O

= 378.9 = 31.9% = 63.3% = 4.8%

Hematite

Fe2O3

=

Fe2O3

= 100%

Magnetite

Fe3O4

=

FeO.Fe2O3 FeO Fe2O3

= 231.4 = 31.0% = 69.0%

Goethite

FeO(OH)

=

Fe2O3.H2O Fe2O3 H2O

= 177.6 = 89.9% = 10.1%

Limonite

2Fe2O3.3H2O =

2Fe2O3.3H2O Fe2O3 H2O

= 373.2 = 85.5% = 14.5%

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Conversion of elemental percentages to oxide percentages Ni to NiO:

Atomic weight of Ni = 58.6 Atomic weight of NiO = 74.6 Conversion factor for Ni to NiO = 74.6 / 58.6 = 1.273

Co to CoO:

Atomic weight of Co = 58.9 Atomic weight of CoO = 74.9 Conversion factor for Co to CoO = 74.9 / 58.9 = 1.272

Fe to FeO:

Atomic weight of Fe = 55.8 Atomic weight of FeO = 71.8 Conversion factor for Fe to FeO = 71.8 / 55.8 = 1.287

Fe to Fe2O3:

Atomic weight of Fe = 55.8 Atomic weight of Fe2O3 = 159.6 / 2 = 79.8 Conversion factor of Fe to Fe2O3 = 79.8 / 55.8 = 1.430

Mg to MgO:

Atomic weight of Mg = 24.3 Atomic weight of MgO = 40.3 Conversion factor for Ni to NiO = 40.3 / 24.3 = 1.658

Si to SiO2:

Atomic weight of Si = 28.0 Atomic weight of SiO2 = 60 Conversion factor for Si to SiO2 = 60 / 28 = 2.143

Al to Al2O3:

Atomic weight of Al = 26.9 Atomic weight of Al2O3 = 101.8 / 2 = 50.9 Conversion factor for Al to Al2O3 = 50.9 / 26.9 = 1.892

Cr to Cr2O3:

Atomic weight of Cr = 52.0 Atomic weight of Cr2O3 = 152 / 2 = 76 Conversion factor for Cr to Cr2O3 = 76 / 52 = 1.462

Mn to MnO:

Atomic weight of Mn = 54.9 Atomic weight of MnO = 70.9 Conversion factor for Mn to MnO = 70.9 / 54.9 = 1.291

Ca to CaO:

Atomic weight of Ca = 40.0 Atomic weight of CaO = 56 Conversion factor for Ca to CaO = 56 / 40 = 1.400

Oxide is always heavier than the metallic element since oxide contains oxygen in addition to the metal. Thus: • To convert into oxide, multiply by the factor • T convert into metal, divide by the factor

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ROCK-FORMING MINERALS SiO4

Silicon tetrahedron

The basic structure of all silicate minerals is the silicon tetrahedron where on silicon atom is linked to four oxygen atoms. In the olivines, the silicon tetrahedra exist independently. In the pyroxenes, the silicon tetrahedra are linked together to make a chain. In the amphiboles, two adjacent chains are so linked as to make a ring structure. In the micas, the silicon tetrahedra are arranged in sheets.

Fields of Mafic minerals

SiO2

PYROXENES

En Kayolinite-Smectite

Fo

CLAYS

Serp.

OLIVINES

MgO Al2O3

Ferrosilite Nontronite

Fa

FeO Fe2O3

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Primary Mafic minerals Olivines Forsterite (Mg) Chrysolite (Mg,Fe) Fayalite (Fe)

Mg2SiO4 (Mg,Fe)2SiO4 Fe2SiO4

Pyroxenes Orthorhombic Enstatite (Mg) Bronzite (Mg,Fe) Hypersthene (Fe,Mg)

MgSiO3 (Mg,Fe)SiO3 (Fe, Mg)SiO3

Monoclinic Clinoenstatite (Mg) Pigionite (Mg,Ca) Diopside (Ca,Mg) Hedenbergite (Ca,Fe) Augite (Ca,Mg/Mg,Fe) Acmite (Na,Fe) Jadeite (Na,Al) Spodumene (Li,Al)

MgSiO3 intermediate between clinoenstatite & diopside CaMg(SiO3)2 CaFe(SiO3)2 CaMg(SiO3)2 with (Mg,Fe)(Al,Fe)2SiO6 NaFe(SiO3)2 NaAl(SiO3)2 LiAl(SiO3)2

Triclinic Rhodonite (Mn) Babingtonite (Ca,Fe,Mn)

MnSiO3 (Ca,Fe,Mn)SiO3.Fe2(SiO3)3

Amphiboles Orthorhombic Anthophyllite (Mg,Fe) Monoclinic Cummingtonite (Mg,Fe) Grunerite (Fe,Mg) Tremolite (Ca,Mg) Actinolite (Ca,Mg,Fe) Hornblende Triclinic Aenigmatite

(Mg,Fe)SiO3

[Si4O11]

(Mg,Fe)SiO3 (Fe,Mg)SiO3 CaMg3(SiO3)4 Ca(Mg,Fe)3(SiO3)4 (K2,Na2,Mg,Ca,Mn)SiO3

[Si4O11] [Si4O11] [Si4O11] [Si4O11] [Si4O11]

(Fe,Na2)(Si,Ti)O3.Na(Al,Fe)(SiO3)2

Micas Biotite, muscovite, phlogopite

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[Si4O11]

Primary Feldspars Plagioclase feldspars Anorthite (Ca) Bytownite Labrodarite Andesine Oligoclase Albite (Na) Potash feldspars Orthoclase Microcline

CaAl2Si2O8 Solid solutions of calcic to sodic plagioclases NaAlSi3O8 KAlSi3O8 KAlSi3O8

An = 100 – 90% An = 90 – 70% An = 70 – 50% An = 50 – 30% An = 30 – 10% An = 10 – 0%

(monoclinic) (triclinic)

Primary spinels (R++O. R+++2O3) Common spinel (Al) Iron spinel (magnetite) Chrome spinel (chromite)

MgAl2O4 Fe3O4 Fe.Cr2O4

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[MgO.Al2O3] [FeO.Fe2O3] [FeO.Cr2O3]

Secondary minerals Hydrous mafic minerals Serpentine (Mg) Chrysotile Serpentine (Fe) Greenalite Tremolite (Mg/Ca) Talc (Steatite or soapstone) (Mg) Kerolite (more hydrous form of talc) Chlorite (Penninite / Clinochlore / Prochlorite) Sepiolite Clay minerals (hydrous silicates) Kayolinite / Nacrite / Dickite (+Al) Hallyosite (endellite) (Al) Illite (K-Al) Smectite (-Al) Montmorillonite / Saponite / Pyrophyllite (Al,Ca,Mg) Nontronite (Fe) Saponite (Mg)

Mg3Si2O5(OH)4 Fe3Si2O5(OH)4 Mg5Ca2Si8O22(OH)2 Mg3Si4O10(OH)2 Mg3Si4O10(OH)2.nH2O Mg5Al2Si3O10(OH)8 Mg4Si6O15(OH)2.4H2O Al2Si2O5(OH)4 Al2Si2O5(OH)4.2H2O KAl3Si3O10(OH)2 Al2Si4O10(OH)2 Fe2Si4O10(OH)2 Mg3Si4O10(OH)2

Non-silicate oxides and hydroxides (also a magnetic polymorph, “Maghemite”) Hematite Fe2O3 Goethite Fe2O3.H2O Limonite 2Fe2O3.3H2O Brucite Mg(OH)2 Boehmite AlO (OH) [Al2O3.H2O] Gibbsite Al(OH)3 [Al2O3.3H2O]

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Nickel hydrosilicates In nickel hydrosilicates, nickel replaces the Mg atoms in such minerals as serpentine, talc, and chlorite. Pure nickel end members do not exist in nature and most nickel hydrosilicates contain (Ni,Mg) in the place of Mg. The term garnierite has been used as a field term to include all hydrous nickelmagnesian silicates. The first member of the garnierite group (chrysoprase, green silica) was defined in the 18th century. Faust (1966) showed that most garnierites are structurally related to talcs and serpentines. Kato (1961) found New Caledonian garnierites to be similar in structure to serpentine, talc and chlorites. Brindley and Pham Thi Hang (1973) determined that garnierites fall into two groups: they are either serpentine-like with 7°A basal spacing, or talc-like with 10°A basal spacing. Springer (1974) proposed the following definition for garnierites: “nickelmagnesian hydrosilicates, with or without alumina contents, having X-ray diffraction patterns typical for serpentine, talc, Sepiolite, chlorite, vermicolite, or mixtures of these”. The field of garnierite minerals has as its end members: kerolite (talc), serpentine, pimelite and nepouite. Compositions of secondary nickel silicates are shown below along with their magnesian end-members: Mg-end member Chrysotile Mg3Si2O5(OH)4 Lizardite Mg3Si2O5(OH)4 Talc Mg3Si4O10(OH)2 Kerolite Mg3Si4O10(OH)2.nH2O Clinochlore Sepiolite

Mg5Al2Si3O10(OH)8 Mg4Si6O15(OH)2.4H2O

Ni-end member (actually Ni,Mg) Pecroaite Ni3Si2O5(OH)4 Nepouite Ni3Si2O5(OH)4 Willemsite Ni3Si4O10(OH)2 Pimellite Ni3Si4O10(OH)2.nH2O Garnierite Ni3Si4O10(OH)2.H2O Nimite Ni5Al2Si3O10(OH)8 Falcondite Mg4Si6O15(OH)2.4H2O

Garnierites are largely of supergene origin being precipitated in the lower parts of the weathered ultramafic profiles, from downward or laterally-moving solutions. Garnierites occur as fillings in open spaces (fractures, joints), or as coatings in joint and fracture surfaces. They range widely in colour from green (light and dark), to yellowing green, to light blue and turquoise blue. The rich green varieties contain more nickel.

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Garnierites

SiO2

10°A basal spacing

Kerolite - Talc

10°A GARNIERITES

Serpentine

Pimelite

7°A

Nepouite 7°A basal spacing

MgO

NiO

Garnierite compositions Sample-1, New Caledonia1 Sample-2, New Caledonia Sample-3, New Caledonia Sample-4, New Caledonia Sample-5, New Caledonia

SiO2 53.0 49.0 53.2 49.8 37.4

MgO 18.1 18.9 15.0 13.5 2.7

FeO 0.08 0.18 0.03 0.21 0.31

NiO 20.9 21.7 24.5 29.2 49.6

Konde, pit-717 Konde, pit-717 Konde, pit-717 Konde, pit-717 Konde, pit-717 Konde, pit-717 Konde, pit-717

43.8 43.7 52.5 48.7 51.1 50.6 37.6

36.8 30.3 25.3 25.2 22.6 16.1 0.7

4.15 5.95 0.20 0.00 0.05 0.04 0.31

1.31 4.30 9.43 12.75 14.67 22.20 47.14

43.7 52.9 47.8 52.3

30.4 18.3 18.6 16.3

5.49 0.23 0.14 n.a.

5.50 16.80 19.6 20.80

(wall)

(centre)

Morro do Cerisco, Brazil Morro do Niquel, Brazil Riddle, Oregon, USA Riddle, Oregon, USA

1

New Caledonian garnierite assays for samples 1-5 are from G. Troly et al., 1979.

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Pyroxene Composition fields CaSiO3 Diopside (CaMg)SiO3

Hedenbergite (CaFe)SiO3 Augite

Ferro-augite

Pigionite

I Hypersthene (Mg,Fe)SiO3

Enstatite MgSiO3

Ferrosilite FeSiO3

Pyroxene compositions

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MnO NiO CoO MgO CaO Na2O K2O H2O LOI Totals

Soroako Unserp. ENSTATITE 55.1 0.05 3.23 0.58

Soroako Unserp. CLINOPYROX. 53.2 0.09 3.47 0.86

Poro Harzburgite Orthopyrox. 60.1

Tiebaghi Harzburgite Orthopyrox. 61.8

5.79 0.13 0.076 0.006 33.5 1.86

2.52 0.08 0.05 < 0.006 18.5 21.7

5.8

5.4

0.1

0.06

34.7

32.7

0.40

0.64

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Iron oxides

Fe Iron protoxide FeO

Magnetite Fe3O4 Iron sesquioxide Fe2O3 ( Hematite)

O

Goethite Fe2O3.H2O Limonite 2F2O3.3H2O

H2O

H

Magnetite Magnetite (Fe3O4) occurs in the laterite profile, derived from two sources: • •

From the primary magnetite present in the ultramafics Magnetite derived from the serpentinisation of the olivines

Magnetite eventually breaks down under the process of chemical weathering. The initial product of weathering is maghemite (a magnetic polymorph of hematite with the Fe2O3 formula). Maghemite eventually gets hydrated to goethite in the upper part of the laterite profile.

19

Goethite Ferrous iron released from the weathering of the primary mafic minerals oxidises to ferric iron and is precipitated as an hydroxide. The hydroxide has a poor degree of crystallinity and appears amorphous. It occurs as concretions as well as clay-like earthy mass. The concretions have concentric layers indicating rhythmic precipitation, often around a nucleus or core. Goethite concentrates as a residual mineral due to its insolubility under prevailing pH-Eh conditions in the laterite environment. Although the standard formula of goethite is Fe2O3.H2O, significant quantities of Al2O3 and Cr2O3 may be present. Thus, a plot of Al2O3+Cr2O3 against the Fe2O3 shows an inverse relationship since both alumina and chrome replace the ferric iron in the goethite structure.

Compositions of goethites

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MnO2 NiO CoO MgO CaO Na2O K2O H2O LOI Totals

Soroako Bonsora 3–6m 1.61 0.08 10.24 3.35 71.96

Soroako Bonsora 6–9m 1.33 0.18 11.13 3.37 70.23

Soroako Bonsora 9 – 12 m 2.71 0.09 11.95 3.15 68.79

0.082 0.41

0.04 0.36

0.08 0.13

0.48 0.02 0.00 0.03

0.46 0.01 0.00 0.02

0.47 0.01 0.00 0.03

2

Manganese in the original assays of all Bonsora samples is reported as total manganese or Mn2O3 (MnO + Mn2O3).

20

COMPOSITION OF CHROMITE Soroako Podiform

Soroako Disseminated

FeO Cr2O3

14.93 49.4

19.56 42.6

Fe Cr Cr/Fe ratio

11.6 33.79 2.91

15.2 29.14 1.92

21

COMPOSITION OF ASBOLITES Asbolane or manganese wad is a black, amorphous-looking material commonly found as thin coatings on joints and fractures, and occasionally as nodules and beads. The material is rich in manganese and contains appreciable amounts of Fe2O3, Al2O3, CoO and NiO. Other elements are present only in very small quantities. Significant amount of the water of hydration is also present.

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MnO2 NiO CoO MgO CaO Na2O K2O H2O LOI Totals

Soroako Limonite 3–6m 1.3 0.1 9.0 0.02 18.4

Soroako Limonite 6–9m 0.8 0.1 15.0 0.2 14.3

Soroako Limonite 9 – 12 m 1.2 n.d. 7.0 0.5 36.0

Soroako Saprolite

31.0 1.65 7.12 0.2 0.03 0.1 0.2

33.6 3.44 7.38 0.5 0.01 0.1 0.1

33.0 2.29 4.96 0.5 0.1 n.a. 0.2

32.0 16.17 3.18 2.8 0.1 n.a. 0.01

1.8 n.a. 3.5 n.d. 14.2

New Cal. (Wadsley, 1950)

1.63 19.22 15.95 39.29 6.99

4.723 11.974

Lithiophorite Lithiophorite is a hydrous manganese-oxide with some lithium in it. Frequently, quantities of lithium can be very low. Various formulae have been advanced for this mineral, all containing the main components of Mn, O, and OH and with minor Al and Li as additional cations: (Al,Li)MnO2(OH)2.

3

H2O -120°C

4

H2O +120°C

22

OLIVINE GROUP Forsterite

Chrysolite

Fayalite

Olivine type Formula In terms of oxides

Magnesia Mg2SiO4 2MgO.SiO2

Ferro-magnesian (Mg,Fe)2SiO4 2(Mg,Fe)O.SiO2

Ferro Fe2SiO4 2FeO.SiO2

Density Hardness

3.22 7

4.39 6.5

Melting point

1890°C

1205°C

Structure • •

Olivines belong to the orthorhombic group (three unequal axes, all at right angles) Olivine structure consists of individual silicon-oxygen tetrahedral linked by magnesium atoms, each of which has six nearest oxygen neighbours. The oxygens lie in sheets parallel to the (100) plane and are arranged in approximate hexagonal close packing (see Figure of olivine structure). The silicon-oxygen tetrahedra point alternately either way along both the x and y directions. The magnesian atoms do not occupy one set of equivalent lattice positions: half are located at centres of symmetry and half on reflection planes, the former having as nearest neighbours two oxygens from two adjacent tetrahedra, the latter, two oxygens from one adjacent tetrahedron.

•

In the olivine series, the unit cell dimensions increase with the increasing replacement of the smaller magnesium atom by the larger iron atom. Unit cell size of Forsterite is 4.756 °A while that of Fayalite is 4.817 °A.

23

Chemical composition of olivines •

Complete solid solution of the Mg and Fe olivines to give a complete range of Mg-Fe olivines

•

Chemical composition of olivines is generally indicated by the percentage of Forsterite (Fo) and Fayalite (Fa) present (example: Fo90Fa10)

•

Pure end members of the Mg-Fe olivines are rare in nature. Thus pure forsterite (Fo100) or pure Fayalite (Fa100) is almost never seen in nature.

•

The composition of olivine that forms dunite at the type locality in New Zealand (Dun Mountain) is Fo92.

•

Generally, the first olivines to crystallise from many basic magmas is in the range of Fo88 to Fo82.

•

The average composition of peridotite olivine is Fo88.

•

Olivines with compositions in the range of Fo80 to Fo50 are very common constituents of basic igneous rocks (gabbros, dolerites/diabases, basalts).

•

Olivines in the ferro-gabbros of the strongly differentiated Skaergaard intrusion show a continuous variation from Fo39 to Fo2.

Olivine Compositions

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MnO NiO CoO MgO CaO Na2O K2O H2O LOI Totals

Soroako Unserp. 40.3 0.02 0.41 0.02

Poro Harzburgite 40.8

Tiebaghi Harzburgite 39.2

8.92 0.13 0.37 0.013 50.8 0.07

7.8

9.0

0.5

0.3

49.2

51.4

0.23

24

Formation of olivines •

Members of the olivine group are important rock forming minerals in the basic and ultrabasic igneous rocks.

•

Basic and ultrabasic rocks generally contain magnesian-rich olivines. Ironrich olivines are found only in some alkaline rocks and iron-rich sediments that have been metamorphosed.

•

Olivines are the first mafic mineral to crystallise from a basic magma.

•

Forsterite has a melting temperature of 1890 °C. Fayalite has a melting temperature of 1205 °C.

•

The cation-oxygen bonds are weaker for the larger ferrous cation (ionic radii of 0.74 °A for Fe++ compared to 0.66 °A for Mg++). As more of the larger ferrous cations enter the olivine structure, there is a progressive reduction in the melting temperature of intermediate compositions.

•

Formation of olivines from a pure olivine-composition liquid is shown in the Figure. As temperature drops and the liquidus curve is reached, olivine of composition intersecting the solidus curve is precipitated. The liquid composition then moves along the liquidus curve as the composition of new olivines moves along the solidus curve. This continues until all olivines with the average composition of the starting liquid are crystallised.

•

If the original magma has more silica than can be used by the olivines (>40%)5, then the more siliceous mafic minerals such as pyroxenes will be formed after all olivines have been precipitated (in actual reality some simultaneous or co-precipitation of olivines and pyroxenes may occur).

5

Pure magnesian olivine (forsterite) has maximum silica content of 42.7%. Iron-olivines, however, require less silica to form. Most olivines have some iron in them. Thus, the silica threshold above which pyroxenes begin to form is in fact lower than 42.7% and depends upon the availability of iron in the magma.

25

Nickel in the olivines •

Olivines can take up to 0.41% of NiO (0.322% Ni). Most of the nickel is contained as a replacement of magnesium atoms by nickel atoms that are of the same size. Generally, the Ni:Mg ratio in the olivines is the same as in the basic magmas.

•

Entry of nickel into the olivine structure in excess of the magmatic Ni:Mg ratio is due to nickel replacing some of the iron in the olivines (iron in olivines lowers the thermal stability).

•

Nickel in the nickeliferous magnetite is also originally derived from the nickeliferous olivines.

Alteration and weathering of olivines •

Minerals of the olivine group are highly susceptible to alteration by hydrothermal fluids and weathering processes.

•

Alteration reactions involve: hydration, silicification, oxidation and carbonation.

•

Common alteration products are serpentine, chlorite, amphibole, carbonates, iron oxides, and talc.

Alteration of Forsterite olivine

H2O

+800°C: Fo to En 625° - 800°C: Fo to En to Talc 500° - 625°C: Fo to Talc 200° - 500°C: Fo to Serpentine

Serpentine Talc

MgO

Fo

26

En

SiO2

SERPENTINISATION OF OLIVINES •

One of the most common products of hydrothermal alteration of olivines is serpentine. The serpentine takes on two forms: o Under static conditions, fibrous chrysotile is formed o Under conditions of stress, flaky antigorite is formed o Under certain conditions, structureless serpophite is formed

•

Alteration of olivines generally begins along randomly placed fractures in the crystal. Eventually, the entire crystal may be altered and replaced ⎯ as a pseudomorph ⎯ by the alteration product.

•

Serpentinisation of olivines results in the: o o o o

Addition of water Leaching of magnesia (or addition of silica) Release of iron in the (Mg,Fe) olivine (serpentine has no iron in it) Conversion of released iron from the ferrous to ferric state to form magnetite. Thus, serpentinised rocks are generally more magnetic.

Conditions for serpentinisation In the presence of water and free silica, olivine would alter to serpentine at temperatures ranging from 200 to 500ºC. Above 500ºC, olivine cannot be converted to serpentine and would change to other minerals as follows: @ 200° to 500°C, olivine changes to serpentine @ 500º to 625ºC, olivine changes to talc @ 625º to 800ºC, olivine changes to enstatite and then to talc @ 800º and above, olivine changes to enstatite (pyroxene) Below is a comparison of olivine and serpentine compositions and densities:

Composition: Composition as oxides: MgO SiO2 H2O Density

Olivine Mg2SiO4 2MgO.SiO2

Serpentine H4Mg3Si2O9 3MgO.2SiO2.2H2O

57.3% 42.7%

43.0% 44.1% 12.9%

3.2

2.2 (relative change of 31%) 27

Serpentinisation accompanied by volume and weight increase (No removal of MgO) Since serpentine carries water and more silica than olivine, these two components have to be added to serpentinise the olivines. As the entire water and silica are used up and as serpentine has a lower density, considerable volume increase takes place in such a process of serpentinisation. 3Mg2SiO4 + 4H2O + SiO2 = 2H4Mg3Si2O9 forsterite water silica serpentine 420 gm 131 cc

552 gm 220 cc

(+31% total weight increase) (+68% total volume increase)

1.26 gm Ni 0.3% Ni

1.26 gm Ni 0.228% Ni (relative decrease of 24%)

The problem with the above model is that most serpentinites do not show any signs of significant volume increases. Similarly, pseudomorphs of serpentine after olivine cannot be explained by the above mechanism.

Serpentinisation at constant volume (Removal of MgO and SiO2) To maintain a constant volume during the process of serpentinisation, considerable quantities of magnesia and some silica need to be removed from the system. The total removal amounts to nearly 31% of the weight of the original olivines. This results in practically no volume change from olivine to serpentine.

5Mg2SiO4 + 4H2O = 2H4Mg3Si2O9 + 4MgO + SiO2 forsterite water serpentine 700 gm

72 gm 552 gm

160 gm + 60 gm Removed in solution

220 cc

220 cc

2.1 gm Ni 0.3% Ni

2.1 gm Ni (assuming all nickel stays in the serpentine) 0.38% Ni (relative increase of 27%)

If it is assumed that all original nickel stays in the newly formed serpentine, then the relative proportion of nickel in the ultramafic is increased by about 27%.

28

However, it is very unlikely that all the nickel will stay behind since considerable quantities of magnesia are being removed hydrothermally. The model of serpentinisation involving equal volumes of original and altered rock faces a different problem. To remove 31% of the original material requires extremely large quantities of water since the solubilities of both magnesia and silica are still limited in hydrothermal solution. In nature, no evidence exists of either the availability of such large quantities of water or the nearby deposition of such large quantities of magnesia and silica that are removed from the system.

Dissociation of Serpentine at high temperatures At 500º to 625ºC, serpentine changes to talc. Water of hydration and magnesia are released: 3MgO.2SiO2.2H2O [serpentine]

3MgO.4SiO2.H2O + 3H2O + 3MgO [talc]

At 625º to 800ºC, serpentine changes first to enstatite and then to talc. In both reactions, water of hydration is given out and magnesia is formed. 2(3MgO.2SiO2.2H2O)

4(MgO.SiO2) + 4H2O + 2MgO 3MgO.4SiO2.H2O + 3H2O + 3MgO

At over 800ºC, serpentine changes to enstatite. Water of hydration and magnesia are released: 3MgO.2SiO2.2H2O [serpentine]

2(MgO.SiO2) + 2H2O + MgO [enstatite]

29

[enstatite] [talc]

ULTRAMAFIC ROCKS

Terminology •

Ultramafic rocks are rich in mafic (ferro-magnesian) minerals such as olivines, pyroxenes and amphiboles.

•

Almost all ultramafic rocks contain less than 45% silica (however, pure orthopyroxenite is classified as an ultramafic rock but contains nearly 60% silica).

•

All ultramafic rocks have colour indices of more than 70.

•

Ultramafic rocks generally lack any feldspars.

•

Ultramafic rocks have no exact counterparts among lavas6.

•

The density of ultramafic magma would be too high for it to rise through the sialic portion of the earth.

Classification of ultramafic rocks Dunite Monomineralic ultramafic rock composed entirely of olivine (generally magnesian). Accessory minerals may include: chromite, magnetite, ilmenite and spinel. Pyroxenite Monomineralic ultramafic rock composed entirely of pyroxenes. The pyroxenites are further classified into whether the pyroxene is orthorhombic or monoclinic: • •

Orthopyroxenites: Bronzitites Clinopyroxenites: Diopsidites; diallagites

Hornblendite Monomineralic ultramafic rock composed entirely of hornblende.

6

Such as rhyolites are volcanic counterparts of granites; dacites are volcanic counterparts of granodiorites; andesites are volcanic counterparts of diorites; trachytes are volcanic counterparts of syenites and basalts are volcanic counterparts of gabbros.

30

Serpentinite Monomineralic rock composed entirely of serpentine. However, the rock may have formed by the serpentinisation of dunite, pyroxenite, hornblendite or peridotite. Peridotite Ultramafic rock that contains largely olivine but also other mafic minerals in significant amounts. Depending upon the other mafic minerals, peridotite may be classified as: Pyroxene peridotite Hornblende peridotite Mica peridotite (such as kimberlite) Pyroxene peridotites are one of the most common ultramafic rocks. Depending upon the type of pyroxene, pyroxene peridotites are further classified into: • • •

Harzburgite: olivine + orthopyroxene (enstatite or Bronzite) Wehrlite: olivine + clinopyroxene (diopside or diallage) Lherzolite: olivine + orthopyroxene + clinopyroxene

Field occurrence of ultramafic bodies The field occurrences of the ultramafic rocks can be simplified to essentially three types: 1. Ultramafic rocks associated with layered intrusions. There is clear evidence in these locations for ultramafics to have derived through gravity settling of heavy mafic minerals during crystallisation of a basic magma (Skaergaard intrusion, the Great Dike in Africa) 2. Small bodies composed entirely of ultramafic rocks (lenses, sheets, dikes, stocks, etc. Occasionally, a feeder to the magmatic chamber is clearly present indicating that the ultramafics may have been intruded as solid crystalline masses. 3. Very large ultramafic occurrences that are clearly associated with ophiolites, subduction melange, outer island arcs and orogenic belts (Ural area, Himalayas, New Zealand, New Caledonia, Sulawesi, etc.). Refer to Figures that show the distribution of major ultramafic belts of the world.

31

OLIVINE

ULTRAMAFIC ROCKS

Dunite 90% OL.

Peridotites Harzburgite Lherzolite Wehrlite 40% OL.

Pyroxenites Orthopyroxenite Websterite Clinopyroxenite

Hornblendites

PYROXENE

HORNBLENDE

OLIVINE

ULTRAMAFICS

Dunite

IN TERMS OF OL, OPX, CPX

90% OL

Wehrlite

Harzburgite Lherzolite OL+OPX+CP 40% Olivine

Olivine Orthopyroxenite

OrthoPyroxenite

Olivine Clinopyroxenite

Olivine Websterites 10%

OL+OPX+CP

10%

ClinoPyroxenite

Websterites

ORTHOPYROXENE

CLINOPYROXENE

32

Nickel in the ultramafic rocks •

Nickel in the ultramafic rocks is held primarily in the mafic minerals. The proportion of nickel generally decreases as follows: olivine > opx > cpx.

•

Primary chromite and magnetite may also contain minor amounts of nickel.

•

Within mafic minerals, nickel is held essentially by olivines that crystallise first. Pyroxenes, that crystallise later, contain far less quantities of nickel.

•

Olivines can take up to 0.4%Ni (0.5% NiO). Much of the nickel is held as a replacement of Mg atoms in the olivine structure.

•

Some nickel, however, may also be held as a replacement of the larger Fe++ atoms in the olivines, particularly when the Ni:Mg ratio in the olivine is higher than that in the original magma (iron in the olivines lowers the thermal stability and allows nickel to come in with greater ease).

•

Nickel contents of olivines: %Ni Olivine in harzburgite, Poro, New Caledonia 0.39 Olivine in unserp. peridotite, Bonsora-W, Soroako 0.358 Olivine in unserp. peridotite, Konde pit 717, Soroako 0.313 Olivine in unserpentinised peridotite, Soroako 0.29 Olivine in harzburgite, Tiebaghi, New Caledonia 0.24 Olivine in “early rocks” of the Skaergaard Intrusion 0.20

•

Nickel contents of pyroxenes: %Ni 0.10 0.067 0.0635 0.047 0.039

%NiO 0.127 0.085 0.081 0.06 0.05

%Ni Chromite, unserp. peridotite, Konde pit 717, Soroako 0.0657 Chromite, unserp. peridotite, Bonsora-W, Soroako 0.0620

%NiO 0.084 0.079

Opx in harzburgite, Poro, New Caledonia Opx in unserp. peridotite, Konde pit 717,Soroako Opx in unserp. peridotite, Bonsora-W, Soroako Opx in harzburgite, Tiebaghi, New Cal. Cpx in unserpentinised peridotite, Soroako •

%NiO 0.50 0.456 0.398 0.37 0.30 0.25

Nickel contents of chromites:

33

•

Nickel contents of ultramafics: %Ni Average dunite compositions (Edel’shtein, 1960) 0.26 Unserp. peridotite, Bonsora-W, Soroako 0.22 Average peridotite compositions (Edel’shtein, 1960) 0.16 Average pyroxenite compositions (Edel’shtein, 1960) 0.08

•

%NiO 0.33 0.28 0.20 0.102

Santos-Ynigo and Esguerra (1961) found highest nickel grades in laterite associated with dunite, peridotite and serpentinite. The poorest grades were associated with pyroxenite and conglomerates.

34

MAGMATIC DIFFERENTIATION Magmatic differentiation is the process by which a homogeneous magma crystallizes into unlike fractions and yields rocks of different compositions. Magmatic differentiation is achieved through fractional crystallisation ⎯ a process in which different crystals are formed at different temperatures during the cooling of the magma. There is a tendency for the newly formed crystals to remain in equilibrium with the liquid. Different minerals achieve this equilibrium in two different ways: 1. Continuous-reaction series: The first plagioclases to crystallise from a basic magma are rich in lime (anorthitic composition). As the temperature drops, and as the remaining liquid gets enriched in sodic composition, first-formed crystals become progressively more sodic. In this way, a continuous series of homogeneous solid solutions is produced. Crystallisation of this type is referred to as a continuous-reaction series. 2. Discontinuous-reaction series Certain ferro-magnesian minerals are incapable of solid solutions. As the composition of the liquid changes significantly, the first-formed crystals change into entirely new crystals with a completely different crystal structure. During the cooling of a basic magma, the first crystals to form belong to the olivine group. If the magma has more than 40% silica, the olivines will react with the remaining liquid to form pyroxenes that have a completely different crystal structure. In time, the pyroxenes may react with the liquid to form amphiboles and them biotite. Such crystallisation process that yields crystals of completely different structure is called a discontinuous-reaction series. Bowen was the first petrologist to propose the order of minerals in the Continuous and Discontinuous reactions series that are shown below. During the cooling of a basic magma, both series begin to crystallise at about the same time. Thus gabbros contain olivine, magnesian pyroxene and calcic plagioclase. Similarly, low-temperature minerals go together such as mica, alkalic feldspars and quartz (granites and granodiorites). Compatible minerals: (olivine + anorthite); (olivine + pyroxene); (quartz + orthoclase) Incompatible minerals: (forsterite + quartz); (labradorite + quartz); (orthoclase + bytownite)

35

BOWEN’S REACTION SERIES

Discontinuous series

Continuous series

(Mafic Minerals)

(Plagioclases)

Olivine 42% SiO2

Calcic plagioclase

Mg pyroxene 58% SiO2

increasing Fe/Mg

Calci-sodic plagioclase

Sodic-calcic plagioclase

Mg-Ca pyroxene

Amphibole

Sodic plagioclase

Biotites

K-Feldspar Muscovite

Quartz

Zeolites

Hydrothermal Solutions

36

PHASE DIAGRAMS Phase diagrams are used to describe the equilibrium conditions of different components and phases in a system. Components: Chemically different components of a system Phases: Chemically similar components with different physical characteristics

Ice – Water – Vapour system

P R E S S U R E

WATER

ICE 4.8mm

P VAPOUR

0.008°C

TEMPERATURE

This system has only one component (H2O) with three different phases. All three phases can co-exist at 0.008°C and a pressure of 4.8mm. At higher temperatures, water can be turned directly into vapour just by changing the pressure. At lower pressures, water can be turned into ice just by changing the temperature.

37

Two–component system with solid solutions Solid solutions are true, homogeneous solid solutions of one substance into another. They involve two isomorphous members with the same basic crystallographic structure and forming a series in which physical and chemical properties change continuously from one member to the other.

T E M P E R A T U R E

1

Liquidus D

T1

C F

T2

Liquid E

H

T3

G Solidus K

Solid solutions of A and B A

COMPOSITION

B

Consider a liquid made up of two components, A and B. The composition of this liquid is shown in the Figure by point 1. If the liquid of composition 1 cools, no change occurs until the Liquidus curve is reached at temperature T1. At this temperature, crystals appear with a composition D (where the T1 temperature intersects the Solidus curve). These crystals in fact represent a solid solution of A and B, and are richer in component A than the starting liquid. With further cooling, the composition of liquid changes along the Liquidus curve towards E and the composition of the crystals changes along the Solidus curve towards F. During slow cooling, the crystals change their composition by continuous reaction with the liquid. At temperature T2, a solid solution of composition F is in complete equilibrium with liquid of composition E. At temperature T3, the last drop of the liquid is consumed and this liquid is in equilibrium with crystals of composition H. Crystals of H have the same composition as the starting liquid. If the early formed crystals of composition D or F are removed from the system (say by crystal settling), then the composition of the liquid will move beyond G and towards K. This will give rise to later crystals of composition in which proportion of B is higher than the starting liquid. An example of the above system is Forsterite-Fayalite.

38

Two-component system without solid solution The components A and B in the system shown below do not form solid solutions. T1 and T3 are the melting temperatures of A and B respectively.

1

2 T3

T E M P E R A T U R E

Liquid T1 T2

T4 Crystals of B and Liquid in equilibrium

Liquidus Crystals of A and Liquid in equilibrium

F

Te

E

C

D

G

Te

Crystals of A and B in equilibrium

A

COMPOSITION

B

Consider the liquid of composition 1. As the liquid cools and reaches temperature T2, crystals of A appear. With the crystallisation of A, the composition of the liquid moves along the liquidus towards E. Until the composition of liquid reaches the point E, crystals of A keep on forming. At temperature Te, crystals of A and B form simultaneously until all the liquid is used up. Point E at which A and B crystallise simultaneously is called the eutectic point and Te is the eutectic temperature. Crystals of A and B are in the proportion of FC to CG. Consider a cooling liquid of composition 2. At temperature T4, crystals of B appear and the composition of the liquid moves along the liquidus towards E. At point E, crystals of B and A appear simultaneously. The final proportion of A and B is in the ratio of FD to DG. System Diopside-Anorthite is an example of the above type of system with a eutectic at 1270°C and a composition of Di58An42. Another system of above type is Forsterite-Diopside with a eutectic at 1400°C and a composition of Fo12Di88.

39

Phase chemistry in PT Inco furnaces Smelting at PT Inco is carried out in furnaces with a diameter of 18m and a power requirement of 45 MVA. Initial expectation was to process the higher-grade West Block ores that also have high silica to magnesia ratios of 2.3 to 2.5 in the slag. Initial smelting showed that: •

The liquidus temperature of the slag was relatively low (1430 – 1475°C)

•

S/M of 2.2 to 2.5 was too acidic for the magnesia refractory in the furnace

•

A higher superheat was required to continually dissolve the coarse unaltered peridotite in the feed. This superheat made it difficult to maintain a protective slag layer on the refractory walls

To correct the problem, PT Inco started blending West Block ores with lower S/M East Block ores. A S/M ratio of 1.9 was targeted in the slag (during the 1980s). The blending had the following benefits: •

The amount of coarse unaltered peridotite was reduced

•

The blending also reduced the S/M ratio, making the slag less acidic

•

The smelting temperature (liquidus) climbed by 50°C, which made it unnecessary to use superheat in the furnace

•

The slag skimming temperature stabilised at 1550°C.

•

Matte tapping temperature also climbed to 1360°C that saved fuel consumption in the converters.

40

WEATHERING Processes of change There are four major processes under which rocks change their physical or chemical properties: 1. 2. 3. 4.

Melting (takes place at very high temperatures) Metamorphism (high temperature / pressure / addition of chemicals) Hydrothermal alteration (through fluids at high temperature) Weathering (at ordinary temperature and pressure

The term “weathering” applies to those superficial changes in rocks that are brought about by atmospheric agencies and result in a complete destruction of the original structure or composition. Deep-seated processes, such as melting, metamorphism and hydrothermal alteration are excluded. Weathering can take place in two ways: •

Physical weathering: mechanical breakdown of rocks (erosion, thermal expansion and contraction, expansion by freezing water, disintegration by plants and animals)

•

Chemical weathering: breakdown of rocks through chemical processes (contact with water, oxygen, and carbon dioxide)

Chemical weathering “The process in which rocks react to atmospheric, hydrospheric and biologic agencies to produce mineral phases that are more stable.” Chemical weathering works in four ways: 1. Hydrolysis: Oxygen, carbon dioxide, ground water, and dissolved acids attack the minerals in the rock and break down their crystal structure. 2. Oxidation: Elements released by chemical weathering are oxidised. 3. Hydration: Reaction with water adds the hydroxyl ion to many newly formed minerals. 4. Solution: The more soluble products of the break down of minerals are dissolved and carried away in ground water.

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Hydrolysis Hydrolysis is the chemical process by which minerals breakdown into more stable components under the influence of chemical weathering. Jenny (1950) has proposed the following explanation for the breakdown of minerals: •

According to Pauling’s rule, the sum of negative and positive charges must be equal within a crystal

•

However, exposed atoms and ions on the surfaces of crystals possess unsaturated valencies and are thus charge

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Contact with water causes hydration of the surface through the attraction of water molecules to the charged surfaces

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Attractive forces are strong enough to cause polarisation of water and its dissociation into hydrogen and hydroxyl ions

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Hydroxyl ions then bond to exposed cations

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Hydrogen ions then bond to exposed oxygens and other negative ions

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Hydrogen ions may also bond to exposed cations, thereby releasing them

Oxidation •

Ferrous ion exists under reducing conditions. At pH of 8 – 8.5, ferrous ion is slightly soluble.

•

Ferric ion exists under oxidising conditions. It is not soluble until pH is reduced to 2.5.

•

Much of the ferrous ions in the weathering profile are converted to ferric state under highly oxidising conditions. This ferric ion state of iron is quite insoluble under normal pH conditions found in ground water.

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Oxidising conditions exist only above the water table. Below the water table, conditions are generally reducing. However, organic matter is a powerful reducing agent and may create reducing environment above the water table, near the upper soil horizon.

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Hot, well-drained environment favours oxidation through the rapid destruction of organic matter and lowering of water table.

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•

Cool, poorly-drained environment promotes accumulation of organic matter and reducing conditions.

Hydration In the presence of hydroxyl ion (OH), many freshly created oxides from the breakdown of minerals are converted to hydroxides. The more common hydroxides that are found in the lateritic soils include: Hydrated oxides of iron: Hydrated oxides of aluminium: Hydrated oxide of magnesium:

Goethite Limonite Boehmite Gibbsite Brucite

Fe2O3.H2O 2Fe2O3.3H2O Al2O3.H2O Al2O3.3H2O MgO.H2O

Many new mafic minerals are formed due to hydration: Serpentine (Mg) Serpentine (Fe) Greenalite Tremolite Talc (Steatite or soapstone) Kerolite (more hydrous form of talc) Chlorite (Penninite / Clinochlore / Prochlorite) Sepiolite

Mg3Si2O5(OH)4 Fe3Si2O5(OH)4 Mg5Ca2Si8O22(OH)2 Mg3Si4O10(OH)2 Mg3Si4O10(OH)2.nH2O Mg5Al2Si3O10(OH)8 Mg4Si6O15(OH)2.4H2O

Hydration also results in the formation of clay minerals that all contain significant quantities of the hydroxyl ion: Kayolinite / Nacrite / Dickite Hallyosite (endellite) Illite Smectite Montmorillonite / Saponite / Pyrophyllite Nontronite Saponite

Al2Si2O5(OH)4 Al2Si2O5(OH)4.2H2O KAl3Si3O10(OH)2 Al2Si4O10(OH)2 Fe2Si4O10(OH)2 Mg3Si4O10(OH)2

Solution For chemical weathering to continue, it is important that all constituents that are broken down from the primary minerals are removed from the environment through the process of solution. Such a process exposes new mineral surfaces to chemical attack.

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•

Dissolved constituents are removed through percolating ground waters

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Ground waters generally travel from top to bottom in a weathering profile

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The ground waters, with dissolved constituents, eventually drain out to rivers, lakes, and the ocean

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The relative proportions of dissolved constituents in ground waters confirm the relative solubilities of various oxides determined in the laboratory

Factors that influence chemical weathering Several factors influence the speed as well as the direction in which chemical weathering takes place. These factors include: 1. 2. 3. 4. 5. 6.

Stability of minerals (crystal structure, melting points) Acidity/basicity (pH) conditions Redox (reduction/oxidation) potential of the environment Rate of removal of dissolved constituents Climate (temperature, rainfall, fluctuation of water table) Topography

1. Stability of minerals •

Godlich (1938) determined the following sequence of decreasing weathering susceptibilities for the common rock-forming minerals: Olivine

Ca-plagioclase

Augite Hornblende Na-plagioclase Biotite K-feldspar Muscovite Quartz

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The above order coincides exactly with that of Bowen (1928) for the crystallisation sequence in a silicate melt. Some minerals with otherwise highly mobile alkali elements can exhibit remarkable resistance to chemical weathering (muscovite, K- and Na- feldspars). In general, the crystal structure of the mafic silicates dictates the ease with which they break down under chemical weathering: o Olivine, with its independent silicon tetrahedra, is the most unstable mineral and thus most susceptible to chemical weathering o Pyroxenes, with their polymerised chains, are relatively more stable and consequently less susceptible to chemical weathering compared to olivines o Amphiboles, with their ring structures, are still more stable and more resistant to chemical weathering o Clays and micas with their sheet-like structure are the most stable minerals and the least susceptible to chemical weathering Reiche (1943) devised a “weathering potential index” based on the formula: 100 x moles (Na2O + K2O + CaO + MgO – H2O) WPI = 100 x moles (Na2O + K2O + CaO + MgO + SiO2 + Al2O3 + Fe2O3) Calculated Weathering Potential Indices WPI calculated for common minerals using the above formula are given below: Mineral Forsterite Enstatite Anthophyllite Augite Hornblende Talc Biotite Orthoclase Quartz Muscovite Kayolinite Gibbsite

Formula Weathering Potential Index Mg2SiO4 66 MgSiO3 55 Mg7(Si4O11)2(OH)2 40 39 36 Mg3Si4O10(OH)2 29 22 12 0 - 10.7 - 67 - 300

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2. Acidity/basicity (pH) conditions •

pH of the ground water has a strong influence on the solution of different materials

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pH values of natural waters normally lie between 4 and 9

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Most oxides show some solubility in natural waters

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Oxides of calcium, magnesium, sodium and potassium are completely soluble in natural waters

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Oxides of titanium, aluminium and ferric iron are completely insoluble in natural waters

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Solubilities of many oxides are pH dependent and change at higher or lower pH levels (oxides of titanium, calcium and ferrous iron)

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Alumina is not soluble in the normal range of groundwater pH. However, at pH values below 4 and above 10, alumina is soluble.

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Where abundant organic matter is available, pH values may drop below 4. Plant roots carry very low pH values, commonly 4 but down to 2.

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Where abundant basic minerals are being weathered (olivine, pyroxene, nepheline), pH conditions may climb to beyond 9.

3. Role of Redox (Eh) potential •

Redox (or reduction/oxidation) potential of a system is a measure of the ability of that system to bring about reduction or oxidation reactions.

•

Reduction is defined as the decrease in the positive valency of an element (Fe+++ to Fe++) or an increase in the negative valency of an element.

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Oxidation is defined as the increase in the positive valency of an element (Fe++ to Fe+++) or a decrease in the negative valency of an element.

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The neutral value for redox potential is zero. At lower values (-), the redox potential represents reducing conditions. At higher values (+), the redox potential represents oxidising conditions.

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•

Certain oxides are capable of existing in multiple valency states: Ti, Fe, Mn. Laboratory tests have demonstrated that the solubility of such oxides is highly variable under different valency states.

•

Thus, under reducing conditions, iron can be dissolved from the weathered profile in the ferrous state. However, under oxidising conditions, iron is stabilised in the weathering profile in the ferric state.

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There are two factors that control the redox potential in most weathering environments: • •

•

Accessibility of atmospheric oxygen (creates oxidising conditions) Availability of organic matter (creates reducing conditions)

In most systems, Redox potentials are dependent on the pH of the system. This is illustrated in the figure given on the following page for the change from ferrous to ferric state.

4. Rate of removal of dissolved material For chemical weathering to proceed at a fast pace, it is necessary to remove the dissolved constituents from the crystal faces and to expose new surfaces to chemical attack. The rate of removal of dissolved constituents in turn depends on several conditions that include: • • •

The relative solubility of oxides The amount of water moving through the system The presence of crystal fractures, cleavages, and porosity of rock

Opposed to the concept of removal of dissolved constituents is the concept of fixation of some elements. While average igneous rocks contain nearly equal quantities of sodium and potassium (3.13% K2O vs. 3.89% Na2O), the concentration of potassium in seawater is only one-tenth of sodium. This is because much of the potassium leached from igneous rocks gets tied up in clay minerals such as illites and does not really leave the systems.

5. Role of Climate Climate has a great influence on the rate of chemical weathering. Climatic factors include the following:

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•

Rainfall: This controls the supply of moisture for chemical reactions and the supply of water for the removal of dissolved constituents. Gentle and persistent rainfall is more effective compared to sudden heavy rains.

•

Temperature: It influences the rate at which break down of minerals takes place. According to Van’t Hoff’s rule, each 10°C change in the temperature increases the speed of chemical reaction from 2 to 3. It is estimated that chemical weathering and leaching in temperate climates amounts to only 1/20 to 1/40 of the one in tropical areas.

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Vegetation: Extensive accumulation of organic material in tropical environment leads to low pH values in the percolating waters.

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Equatorial Humid Climate: It results in good and constant flushing of magnesia and silica from the laterite system and does not allow the formation of smectite/nontronite clays.

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Tropical Wet-Dry Climate: It leads to poor flushing of the silica and magnesia from the system. The presence of silica and magnesia in the system leads to the development of smectite/nontronite clays.

6. Role of Topography Topography exerts a powerful influence on the rate of weathering, through the following factors: •

Run-off of rain water vs. absorption of rain water into the profile [on steep slopes, much of the rain water runs off and little penetrates the rock. This promotes physical weathering instead of chemical weathering]

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Rate of sub-surface drainage and removal of dissolved material [higher areas afford better drainage than low-lying and flat areas]

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Rate of erosion of the weathered product [High erosion rates keep exposing new surfaces to chemical weathering. Slopes of < 20° are necessary to retain the laterite ahead of erosion].

Ideal conditions for chemical weathering are attained on rolling to gently sloping lands that are elevated and where surface run-off is not excessive and the subsurface drainage is good.

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WEATHERING OF ULTRAMAFIC ROCKS (Dunite, peridotite and serpentinite)

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Dominant minerals include: Olivine, Pyroxene (orthopyroxene in Harzburgites), Serpentine

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Olivine (being a high-temperature mineral) is highly unstable under atmospheric conditions

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Much of the olivine in the ultramafics is magnesia type which is more unstable than the iron-olivine

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Crystal fractures in the olivine assist the access of moisture and water into the lattice structure

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Silica tetrahedra are bonded together by magnesian ions. Magnesium is highly mobile and its removal causes release of individual tetrahedra units.

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Since the solubility of silica is lower than that of magnesia, the released silica cannot be removed from the weathering profile fast enough. Consequently, some of the silica remains behind (temporarily) to form serpentine, clays or even silica boxwork.

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The ferrous iron (if released above the water table) is oxidised into ferric iron. The ferric iron can form: Hematite (or maghemite), Goethite (iron hydroxide), or Limonite (iron hydroxide)

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With increased weathering, serpentine begins to decompose. Magnesia is preferentially leached out of serpentine, leaving behind a silica-enriched phase or montmorillonite (saponite) and chlorite.

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Ni++ and Fe+++ replace the magnesium being leached out of serpentines. This results in the formation of iron-rich nickeliferous serpentine.

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Since nickel is less soluble and mobile than magnesium, the formation of nickeliferous serpentine (and nickeliferous talc, chlorite and smectite), results in more stable minerals that are resistant to further leaching.

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Eventually, montmorillonite (saponite) and chlorite begin to decompose, releasing remaining magnesia and silica and leaving behind oxides and hydroxides of iron.

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•

Some of the alumina present in the mafic minerals is temporarily fixed in the chlorites. But upon the loss of magnesia and silica from this mineral, alumina is finally fixed in gibbsite.

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Magnetite and ilmenite are common accessory minerals in ultramafic rocks. These minerals oxidise readily during weathering to yield hematite, maghemite, and goethite.

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In the limonite zone, nickel occurs largely in goethite, presumably in solid solution. The nickel grade of goethite is about 0.5 to 1.5% Ni.

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Towards the base of the limonite zone, some 10 – 50% of total nickel present is in the asbolite mineral (manganese wad) which ranges in grade from 2 to 14% Ni. Asbolite also carried significant cobalt grades and any zinc or copper present in the profile.

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Nickel has a much higher solubility in acid waters as compared to alkaline waters. Thus, supergene enrichment of nickel requires its solution in the upper acidic portion of the laterite profile, its movement downward towards a low water table, and its ultimate precipitation in the lower saprolite profile where pH conditions are more alkaline and where more soluble magnesia and silica are abundant. In flat topography with a high water table, nickel is not likely to show supergene enrichment in the saprolite zone.

•

Some of the nickel precipitates as garnierite (hydrated silicate of nickel and magnesium). However, much of the nickel commonly found in the saprolite zone is in the form of a nickel-serpentine of secondary origin. This nickel serpentine can carry appreciable quantities of nickel.

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Laterites that are formed in humid environments with considerable annual rainfall, undergo thorough leaching of the profile. In such laterites, the silica and magnesia is flushed out of the system too quickly to form smectite/nontronite clays.

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Laterites that are formed in tropical wet-dry climate, are poorly flushed and generally give rise to the formation of a distinct smectite/nontronite intermediate zone.

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Alumina and chrome in the laterite profile are largely derived from the pyroxene and spinel in the parent rock. Along with iron, alumina and chrome represent truly residual components and their relative proportions in the laterite profile reflect their original ratios in the ultramafic.

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CLASSIFICATION OF SOILS The most commonly accepted classification of soils is the Great Soil Group. Under this classification, soil groups have a wide distribution and a number of common fundamental characteristics. The Great Soil Group classification is based on temperature and humidity (or rainfall). Two major divisions of the soils are based on humidity (or rainfall). As a general rule, annual rainfall of 25” (635 mm) separates the two divisions: Pedocals ⎯ soils with calcium as carbonate) and aluminium, in arid regions) Pedalfers ⎯ soils with aluminium and iron, in humid regions)

T E M P E R A T U R E

HUMIDITY / RAINFALL Diagrammatic representation of the distribution of the Great Soil Groups according to climate (after Millar, Turk, and Forth, 1958)

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LATERITE SOILS

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The term laterite is derived from the Latin word “later” which means brick.

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Buchanan Hamilton first introduced the term in 1807 for the earthy iron crusts that were being cut into bricks for building purpose by the people of south-central India.

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Currently, the term laterite is used for soils that are rich in sesquioxides of iron and aluminium, formed under the influence of chemical weathering with special ground-water conditions.

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Development of laterites require: Availability of rocks that contain iron and aluminium Relatively high temperature (to aid in chemical attack) High rainfall (to aid chemical weathering) Intense leaching (to remove mobile elements) Strongly oxidising environment (to convert Fe, Al to sesquioxides) Gentle topography (to preserve the laterite soil after development)

Laterite Profile Limonite zone: (also called “laterite” zone or oxide zone) • • • •

• •

The upper zone is rich is goethite. The limonite may be remobilised in near surface acid conditions and recrystallised to ferricrete (iron cap, canga, cuirasse de fer). Extremely insoluble minerals may persists in this zone (spinel, magnetite, maghemite, and primary talc). The base of the limonite zone is enriched in manganese, cobalt and nickel in the form of asbolite or manganese wad. This manganese wad usually occurs as extremely thin surface coatings on joint and fracture planes. The limonite zone represents laterite that has collapsed under its own weight. Thus, dry bulk density in the limonite zone is generally higher than in the transition zone. Due to collapse, the original structure and texture of the rock is completely obliterated.

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Smectite or Nontronite zone: (also called intermediate or transition zone) • This is an intermediate zone between lower limonite and upper saprolite • This zone contains soft smectite clays (usually Nontronite) and hard crystalline quartz. • Relict texture and structure of the original rock are well preserved in this zone that has not completely collapsed as yet. • The development of a distinct intermediate zone depends on climatic peculiarities and its occurrence is limited in the world [in areas of heavy rains throughout the year, the silica and magnesia are completely flushed out of the system to give smectite/nontronite clays. In areas that have a tropical wet-dry climate, smectite/nontronite are quite likely to form and persist]. • When the intermediate zone is developed, the occurrence of manganese wad is more prominent in the upper part of the intermediate zone rather than in the lower part of the limonite zone. Saprolite zone: (Serpentine ore) • This is the zone of altered bedrock where the processes of chemical weathering are proceeding most actively. • Chemical attack and weathering is proceeding along joints and fractures in the rock and cleavages and micro-fractures in the crystals. • Saprolisation along joint surfaces leads to the formation of “boulders” within the saprolite zone. • The boulders can have a significant saprolised crust which can carry good nickel grades. • All original rock textures and structures are well preserved. • Most of the parent rock minerals are also preserved. • The zone consists of bedrock fragments, saprolised rims of boulders, and precipitated quartz and garnierite. • In unserpentinised peridotite, saprolisation is limited to boulder surfaces since the fresh rock is extremely hard for the water solutions to penetrate. • In unserpentinised peridotite, unsaprolised boulders remain free of nickel. • In serpentinised peridotite, saprolisation proceeds through much of the rock mass at the same time since rock is soft enough to permit access to water solutions. • In serpentinised peridotite, unsaprolised boulders may contain significant quantities of supergene nickel.

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Rates of laterisation Based on the solubility of various components in an ultramafic rock, some investigators have arrived at a rate of dissolution of 1mm per 100 years. This translates to: 1 metre / 100,000 years 10 metres / million years 50 metres / 5 million years

According to Golightly (1979), water emerging from well-drained ultramafics typically contains: 25 ppm Mg 10 – 20 ppm Silica pH = 7 – 8 The above numbers correspond to a maximum rock dissolution rate of 1.4mm per 100 years and translate to: 1.4 metre / 100,000 years 14 metres / million years 70 metres / 5 million years On an average, one metre of ultramafic rock produces only 0.4m of laterite. Thus the rate of laterite formation would be about 4 – 6 metres per million years.

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