Download Lecture 8: Igneous Petrogenesis

Lecture 8: Igneous Petrogenesis
• Igneous
• Phase
relations
• Mantle
• Trace
rock classification
melting
element geochemistry
>70% of Earth’s annual volcanic
budget is erupted in the oceans.
Igneous Rock Classification: Texture
•The first distinction is between volcanic and plutonic rocks.
– Volcanic rocks are erupted at the Earth’s surface and cool very
quickly. There is insufficient time to grow large crystals. This leads
to formation of glass or very fine-grained rocks, or to phenocrysts
(crystals that grew before eruption) in a fine groundmass.
– Plutonic rocks crystallize at some depth, and therefore lose heat
relatively slowly. Crystals have time to grow after nucleation, and
the resulting rocks generally have individual crystals large enough
to see unaided.
– Rocks of exactly the same composition and mineralogy get
different names in their volcanic and plutonic forms, because they
look different.
Basalt
Gabbro
5 mm
5 mm
Aphanitic: mineral grains or
groundmass that are smaller than 1 mm
(need microscope or hand lens to see).
Porphyritic: larger mineral
grains within an aphanitic or
phaneritic matrix.
Phaneritic: mineral grains easily seen
with naked eye >3 mm.
Mineral groups
1) Silicates (SiO4) – make up 96% of minerals, e.g., olivine
2) Carbonates (CO3): e.g, calcite CaCO3
3) Oxides: metal and oxygen (e.g., hematite, magnetite)
4) Sulfides: element + S2 (pyrite – FeS)
5) Sulfates: element + SO4 (gypsum – CaSO4 2H2O)
6) Halides: element + halide (salt - NaCl)
7) Native elements: e.g., Cu, Au, Ag
Key minerals in mafic igneous rocks:
• Olivine: (Mg,Fe)2SiO4 -- Forsterite (Mg), Fayalite (Fe)
Pyroxene: (Mg,Fe,Ca)2Si2O6 -- Mg-, Fe-, Ca-, (Clinopyroxene);
Na-, Al- (Orthopyroxene)
• Feldspar: (K,Na,Ca)AlSi3O8 -- Albite (Na-) ,Anorthite (Ca-),
Plagioclase (Na+Ca), Alkali (K-)
gypsum
•
Igneous Rock Classification: Mineralogy
• The standard classification scheme uses the mineralogy of the
rock (how much quartz, how much plagioclase, etc.)
– There is one important twist…for volcanic rocks you usually
cannot measure the actual minerals present (or it may be a
glass and there are no minerals present).
– In this case, instead of the actual minerals, you classify based
on normative mineralogy
• The norm is a calculation based on the bulk composition of a
volcanic rock, for what minerals would be present if it were
fully crystallized.
• The standard norm calculation is called the CIPW norm, after
Cross, Iddings, Pirsson, and Washington (1902).
Quartz: SiO2
Orthoclase: KAlSi3O8
Plagioclase: NaAlSi3O8
Feldspathoid: feldspar
with Al:Si = 1.
peridotite
pyroxenite
Dunite
Lherzolite
Pyroxenite
Wt.% Al2O3
1. Ophiolites
2. Dredge samples from oceanic fracture
zones
3. Xenoliths in basalts
4. Kimberlites
Tholeiitic basalt
15
l
tia
r
Pa ing
% lt
20 Me
10
5
Lherzolite
0
0.0
Harzburgite Residuum
Dunite
0.2
Dunite
0.4
Wt.% TiO2
0.6
0.8
Lherzolite
Pyroxenite
Other igneous rock classifications
•By silica percentage:
%SiO2
Designation
%Dark Minerals
Designation
Examples
>66
Acid
<40
Felsic
Granite, rhyolite
52-66
Intermediate
40-70
Intermediate
Diorite, andesite
45-52
Basic
70-90
Mafic
Gabbro, basalt
<45
Ultrabasic
>90
Ultramafic
Dunite, komatiite
By alumina saturation (which dark minerals show up):
Chemistry
Designation
%Dark Minerals
Peraluminous
Muscovite, biotite, topaz,
corundum, garnet, tourmaline
Na2O+K2O+CaO>Al2O3 &
Al2O3 > Na2O+K2O
Metaluminous
Melilite, biotite, pyroxene,
hornblende, epidote
Al2O3 ~ Na2O+K2O
Subaluminous
Olivine, pyroxenes
Al2O3 < Na2O + K2O
Peralkaline
Sodic pyroxenes & amphiboles
Al2O3>Na2O+K2O+CaO
Total alkalis + silica (TAS)
classification
Basalt
Rhyolite
Geodynamic setting of igneous rocks
• Igneous rocks are formed today at plate margins or in continental
or oceanic plate interiors (but most of the action is at plate
boundaries).
Mantle melting terminology
Geotherm – Vertical
temperature profile in the
earth
Solidus – Temperature at which
a rock will first start to melt
Liquidus – Temperature at
which a rock will be fully
molten.
Adiabat – A packet of the
mantle that moves up/down
without gaining or losing heat.
Another explanation of the adiabat
Imagine the Earth with its present distribution of material but without gravity. The
material is uncompressed and there is no pressure increase with depth. Set the initial
temperature everywhere to the Earth's surface temperature. Now turn gravity back
on. The gravitational pressure causes the material to contract, with material
compressing more at greater depths because of the greater pressure. The
temperature will also increase because of the compression. If this is done such that
no heat is gained or lost by any given piece of the material, the temperature increase
for any parcel of matter will be adiabatic, and the temperature increase with depth
will thus be adiabatic.
Temperature
Depth
No gravity,
isothermal
Gravity turned on,
adiabatic
Another explanation of the adiabat
Another way to achieve an approximate adiabatic temperature distribution is to have
material convect heat from the hotter interior to the cooler exterior. The heat is
carried upwards by the upwards movement or flow of material, while material cooled
near the surface descends. If the temperature gradient (increase of temperature with
depth) is adiabatic, then upwards movement of a parcel of material will not result in a
temperature difference of the parcel with respect to the surrounding material.
However, if the temperature gradient is greater than adiabatic (super-adiabatic), the
temperature of an upwards moving parcel will only decrease by the adiabatic
gradient, and so will be greater than that of its surroundings.
Temp. excess of parcel
at shallower depth
Depth
In situ
geotherm
(super-adiabatic
temperature gradient)
Adiabatic gradient
Upward movement
of material along adiabatic
gradient
Mantle Plumes
Mantle melts between
~1300-1800ºC due to:
• Increase in temperature
• Decrease in pressure
• Addition of volatile
phases
Mid-Ocean Ridges
(and Plumes)
Mantle melts between
~1300-1800ºC due to:
• Increase in temperature
• Decrease in pressure
• Addition of volatile
phases
Subduction Zones
Mantle melts between
~1300-1800ºC due to:
• Increase in temperature
• Decrease in pressure
• Addition of volatile
phases
Percentage of melting (F)
The pressure (or
depth) versus
temperature (P-T)
path of upwelling
mantle beneath a
mid-ocean ridge
leads to a maximum
of ~25% melt.
Phase Diagrams
A phase diagram is common way to represent the system state at specific
pressure (P) and temperature (T) conditions. Lines on the diagram represent
conditions under which a phase change is at equilibrium. At a point on a
line, it is possible for two or more phases to coexist at equilibrium. In other
regions, only one phase exists at equilibrium.
Phase diagram for water
Triple point: where 3 phases
coexist
Phase diagram for Olivine solid solution
Forsterite (Fo)
Fayallite (Fa)
Solidus: the temperature below which the substance is stable in the solid state
Liquidus: the temperature above which the substance is stable in the liquid
state
Lever Rule: to determine quantitatively the relative composition of a mixture in
a two-phase region in a phase diagram
Equilibrium Melting
Equilibrium melting
occurs when the solid
and liquid phases are
kept together as melting
progresses.
Lever Rule
S – solid composition
L – liquid composition
A – system composition
We can write fraction x of solid as
xS + (1-x)L = A
which can also be written as
x (A –S) = (1-x)(L-A)
We can solve the above
equations to get the proportion
of solid
x = (A – L) / (S – L)
Fractional Melting
Fractional melting
occurs if the liquid is
immediately removed
from the solid as the
solid melts.
Equilibrium
Solidification
Equilibrium
solidification occurs
when the solid and
liquid phases are
kept together as
solidifications
progresses.
Fractional Solidification
Fractional solidification
occurs if the solid is
immediately removed
from the liquid as it
crystallizes.
Diopside (Clinopyroxene) – Anorthite (Plagioclase)
Diopside (CaMgSi2O6)
Dark mineral
Gabbro (coarse grained
equivalent of basalt)
– Oceanic Crust
Anorthite
(CaAl2Si2O8)
Light mineral
Diopside-Anorthite phase diagram
1600
liquidus
Liquid
Temperature, ˚ C
1500
eutectic
1400
An + liquid
Di + liquid
1300
1200
CaMgSi206
(Diopside)
solidus
Diopside + Anorthite
20
40
60
Anorthite content, mol%
80
CaAl2Si208
(Anorthite)
Eutectic: mixture that has the lowest freezing point (composition/
temperature of the last solids formed when freezing, first melt formed)
Diopside-Anorthite phase diagram
1600
liquidus
Liquid
Temperature, ˚ C
1500
eutectic
1400
An + liquid
Di + liquid
1300
1200
CaMgSi206
(Diopside)
batch melting
solidus
Diopside + Anorthite
20
40
60
Anorthite content, mol%
80
CaAl2Si208
(Anorthite)
Diopside-Anorthite phase diagram
1600
liquidus
Liquid
Temperature, ˚ C
1500
eutectic
1400
An + liquid
Di + liquid
1300
1200
CaMgSi206
(Diopside)
fractional melting
solidus
Diopside + Anorthite
20
40
60
Anorthite content, mol%
80
CaAl2Si208
(Anorthite)
Back to mantle
How melting
does melting
occur in the mantle?
OL
OL
Gar
Cpx
Gar
Opx
OL
OL
Olivine
G
O
Cpx OL
Incr. melt
C
OL
G
O
OL
Opx
OL
OL
O
C
OL
O
OL
Melting of garnet lherzolite begins at
As the extent of melting increa
cpx-cpx-garnet triple junctions in
melt migrates along grain boun
Melting to
of agarnet
lherzolite
begins at opx-cpx-garnet
response
reduction
in pressure.
forming an triple
inter-connected ne
junctions
(eutectic)
to a reduction
pressure.
Olivine
is
Olivine
is not
involvedininresponse
melting at
that in
allows
the melt
to segrega
notstages.
involved in melting at early stages. As the
the unmelted
extent of melting
(F)
early
crystal residue.
increases, melt migrates along grain boundaries forming an interconnected network that allows the melt to segregate from the
Composition
of
melt
depends
on
the
P
and
T
(which
controls
the
extent
of
unmelted crystal residue.
and the phases involved in the melting. Suppose that ~1/3 (opx) + ~1/3 (cpx
Wenlu Zhu
Melting of garnet lherzolite begins at opx-cpx-garnet triple
junctions (eutectic) in response to a reduction in pressure. Olivine is
not involved in melting at early stages. As the extent of melting (F)
increases, melt migrates along grain boundaries forming an interconnected network that allows the melt to segregate from the
unmelted crystal residue.
Trace elements in mantle melting
• Incompatible elements: preferentially partition into the melt phase
(D<1)
• Compatible elements: preferentially partition into the solid phase
(D>1)
• Partition or distribution coefficient (D) = Csolid/Cliquid
Concentrations
normalized to bulk
earth, C1 chondrites,
or primitive mantle
Most incompatible
Less incompatible
Partition coefficients
Partition coefficients are determined for an element between a
unique mineral phase in a unique lattice site and melt, and are
determined by three primary factors…
• Size (ionic radius) is fairly intuitive control, since the substituting
ion needs to fit into a mineral lattice: Too big or too small a
won't be energetically stable.
• Charge (ionic charge) is also intuitive, since charge must be
balance within a lattice and if a charge imbalance is generated
by a substitution, a second substitution must occur to correct
for this.
• Electronegativity is harder to visualize, but the disruption to the
mineral lattice of replacing a greedy element with a giving
element or vice versa is too much for a lattice to take.
Partition coefficients
Rock Type
Mineral
Z
Elem
Value
Basalt
Garnet
41
Nb
Basalt
Garnet
41
Nb
0.01
Basalt
Ilmenite
41
Nb
0.8
Basalt
41
Nb
0.003
Basalt
Low Calcium
Pyroxene
Magnetite
41
Nb
Basalt
Olivine
41
Nb
Basalt
Plagioclase
41
Nb
Basalt
Plagioclase
41
Basalt
Plagioclase
Basalt
Kd Type
Reference
Experimental
Jenner et al. 1994
Phenocryst-Matrix,
Experimental
Experimental
Keleman & Dunn 1992
McCallum & Charette 1978
Phenocryst-Matrix,
Experimental
Calculated
Keleman & Dunn 1992
0.01
Calculated
McKenzie & O'Nions 1991
0.01
Calculated
McKenzie & O'Nions 1991
Nb
Experimental
McCallum & Charette 1978
41
Nb
Experimental
Bindeman et al. 1998
Plagioclase
41
Nb
Experimental
Aignertorres et al. 2007
Basalt
Rutile
41
Nb
16
Experimental
McCallum & Charette 1978
Basalt
Rutile
41
Nb
136
Experimental
Foley et al. 2000
http://earthref.org/KDD/
Nielsen 1992
Trace Elements in mantle melting
• Incompatible elements: preferentially partition into the melt phase
(D<1)
• Compatible elements: preferentially partition into the solid phase
(D>1)
• Partition or distribution coefficient (D) = Csolid/Cliquid
Melts
Cmelt/Co
10
Relating trace element concentrations
to melt fraction (F) batch (equilibrium) melting
D=0.01
0.1
0.5
1
1
C-T Lee
5
!!"#$
1
=
!
!
!! + ! 1 − !!
!!"#
Equilibrium Melting
10
0.1
0
100
!!"#
!!
!
0.6
0.8
!0.2 = 0.4
!!"# !! + ! 1 − !!
100
Solids
Melts
Cmelt/Co
1
10
10
D=0.01
Csolid/C
!!"#
o
0.1
!!"#$
1
=
!
!
!!0.5+ ! 1 − !!
!!"#
1
!
!!"#
10
5
!!
!
!! + ! 1 − !!
1
=1
0.5
1
0.1
0.1
5
0.1
0
100
10
0.01
10
0.01
0.2
0.4
0.6
0.8
1
F
0
0.2
F
0.4
0.6
0.8
1
F
FIGURE 7.3. Equilibrium Melting.
Solids
before the melt can separate from the solid residue. The critical melt fraction
10
elt
Melt Fraction
10
Relating trace element concentrations to melt 5fraction (F) Cmelt/Co
0.5
10
fractional melting
0.1
1
1
C-T Lee
D=0.01
!!"#$
1
! ! ! !!
=
1
−
!
!
!
!!
!Fractional
!"#
Melting
100
!
!
/C !"#
o
0.001
0
1
1
=
!
!! + ! 1 − !!
!!"#
0.5 /C
Csolid
o
100
10
0.2
0.6
0.8
1
10
!!
=1
!
!! + ! 1 − !!
5
1
0.5
0.1
0.01
D=0.01
0.001
0
0.4
!
0.1
0.1
0.01
!
!!"#
0.2
! ! ! !!
10
!
!!"#
10
0.1
= 1−!
Solid
1
5
!!"#
100
Instantaneous
Melt Fraction
10
!!"#$
0.1
0.01
0.01
0.4
F
0.6
0.8
Solid
1
0.001
0
0.2
0.6
0.8
F
100
Aggregate Melt
10
5
0.4
0.01
1
10
Melt Fraction
10
10
5
1
Relating
trace
element
concentrations
to
melt
fraction
(F)
5
0.5
Cmelt/Co
id/Co
0.5
10
0.1
fractional
melting
0.1
elt
1
0.1
0.01
!!"#$ 1 − 1 − ! !
= 0.4
! 0.2
! 0.6
!!"#
!!
D=0.01
!
0.8
0.001
0
1
100
/ Co
!
!!"#
0.2
0.4
! ! ! !!
0.6
10
1
= 0.1
!
!! + ! 1 − !!
10
!!"#
!
!!"#
Csolid/Co
0.5
!
0.8
1
10
!!
=1
!
!! + ! 1 − !!
5
1
0.5
0.1
1
1
0.01
10
0.2
0.1
0.01
5
0.1
0
= 1−!
Solid
0.01
!
!!"#
!!"#
100
Aggregate Melt
!!"#$
0.1
0.01
0.01
0.001
0
1
1
0.4
F
0.6
F
FIGURE 7.4. Fractional Melting
0.8
1
0.001
0
0.2
0.4
0.6
0.8
F
100
Aggregate Melt
0.01
1
0.001
0
0.2
0.4
0.6
0.8
0.001
0
1
ESCI 430
100
0.2
0.4
0.6
0.8
C-T Lee
100
Solid
Aggregate Melt
10
0.01
10
0.1
Csolid!/C
!"#o
=
Equilibrium Melting
0.1
0.1
0.01
0.5
0.01
1
10
0.1
0
C-T Lee
0.2
0.4
0.6
0.8
0.001
0
1
100
0.1
F
Fractional
1
0.2
5
0.4
0
100
FIGURE
7.4. Fractional Melting
C melt!/!"#Co
0.4
10
5
1
1
0.5
5
0.1
0
5
100
1
0.1
10
0.1
10
0
0.8
Solids
1
0.1
0.6
!0.5!
=
!
!
Csolid
!!"#/Co !1!1 + ! 1 − !!
1
0.1
=
!
!!0.5+ ! 1 − !!
1
Equilibrium
Aggregate Melt
0.2
10
D=0.01
1
0.8
0.01
100
Melts
0.6
10
Equilibrium Melting
10
Melts
0.5
D=0.01
1
5
!
!!"#
5
1
1
!!
!
!
0.1
C!melt
! + ! 1 − !!
!"#/Co !10
1
!"#$
C melt !
/C
o =
0.5
!
!
!! + ! 1 − !!
!!"#
Cmelt/Co
!!"#$
10
100
1
1
0.01
0.2
0.4
F
0.6
0.8
1
0
0.1
10
0.2
0.4
0.6
0.2
F
0.4
0.6
F
0.8
1
0.8
1
0.01
F Melting.
FIGURE 7.3. Equilibrium
h rule was added more recently by Ringwood:
ion with the most similar electronegativity to that of the major
ment being replaced will be favored because it destabilizes the
tal lattice the least.
Trace element partitioning evidence for differentiation of
the Earth
GG325 L36, F2013
Radiogenic,StronOum,Isotope,(87Sr/86Sr),System,
Evolution of 87Sr/86Sr in Earth’s geological reservoirs
DRb (olivine) = ~0.003
element partitioning example
DRb (pyroxene) = ~0.002
0.72
agram shows contours of the clinopyroxene-melt
ution (partition)
coefficient=for~0.0005
various ions as a function
DRb (garnet)
rge and radius (i.e., primarily rules #1 and #2).
Rb
140
K
Ba
120
Sr
100
Oceanic
Crust
Highe
Ca
80
60
40
Pb
La
Th
Y REE
Lu
Mn
U
V
Fe
Hf
Li
Co Sc
Mg
Zr
Ni
0.2
Ga
0.5 >1 Cr
0.1
Ti
0.01
P
Be
3
1
2
4
Ionic Charge
Continetal
Crust
87Sr/86Sr
160
Highe
0
Ta , N b
5
0.70
Low R
Bulk Earth
Depleted Earth’s Mantle
6
Figure 7.10. Ionic radius (picometers) vs. ionic charge contoured for
clinopyroxene/liquid partition coefficients. Cations normally present in
clinopyroxene are Ca2+, Mg2+, and Fe2+, shown by symbols. Elements
whose charge and ionic radius most closely match that of the major
elements have the highest partition coefficients.
modified from White, Geochemistry
4.5 Billion yrs
GG325 L36, F2013
Time
Present
Primitive mantle- normalized
1000
Melts and crust
normalized to
primitive mantle
Cont.
Crust
100
10
1
Plume
MORB
0.1
0.01
Cs Rb Ba Th
U
Nb La
Ce Pr
Sr
Nd Zr
Hf Sm Gd Tb Dy Ho
Y
Er Yb Lu
FIGURE 7.7. Trace-element abundances of continental crust, mid-ocean ridge basalt
(MORB) and a plume basalt normalized to primitive mantle.
The second reason for normalization is that one can learn something about
geologic processes that fractionate (i.e. change the relative proportions of trace-elements)
trace-element relative abundances from their original relative abundances. For example,
Primitive mantle- normalized
1000
Melts and crust
normalized to
primitive mantle
Cont.
Crust
100
10
1
Plume
MORB
0.1
0.01
Cs Rb Ba Th
U
Nb La
Ce Pr
Sr
Nd Zr
Hf Sm Gd Tb Dy Ho
Y
Er Yb Lu
FIGURE 7.7. Trace-element abundances of continental crust, mid-ocean ridge basalt
(MORB) and a plume basalt normalized to primitive mantle.
The second reason for normalization is that one can learn something about
geologic processes that fractionate (i.e. change the relative proportions of trace-elements)
trace-element relative abundances from their original relative abundances. For example,
Primitive mantle- normalized
1000
Melts and crust
normalized to
primitive mantle
Cont.
Crust
100
10
1
Plume
MORB
0.1
0.01
Cs Rb Ba Th
U
Nb La
Ce Pr
Sr
Nd Zr
Hf Sm Gd Tb Dy Ho
Y
Er Yb Lu
FIGURE 7.7. Trace-element abundances of continental crust, mid-ocean ridge basalt
(MORB) and a plume basalt normalized to primitive mantle.
The second reason for normalization is that one can learn something about
geologic processes that fractionate (i.e. change the relative proportions of trace-elements)
trace-element relative abundances from their original relative abundances. For example,
Reading for next class
y that other
can similarly
iciently high
We have also
ace of Cu for
ater-gas shift
as a result of
for this type
f the working
Supplementary Text
Figs. S1 to S8
Tables S1 to S3
References (33–41)
16 November 2015; accepted 14 December 2015
10.1126/science.aad8868
OCEANOGRAPHY
Enhanced East Pacific Rise
hydrothermal activity during the last
two glacial terminations
D. C. Lund,1* P. D. Asimow,2 K. A. Farley,2 T. O. Rooney,3 E. Seeley,1
E. W. Jackson,4 Z. M. Durham4
Mid-ocean ridge magmatism is driven by seafloor spreading and decompression melting
of the upper mantle. Melt production is apparently modulated by glacial-interglacial
changes in sea level, raising the possibility that magmatic flux acts as a negative feedback
on ice-sheet size. The timing of melt variability is poorly constrained, however, precluding
a clear link between ridge magmatism and Pleistocene climate transitions. Here we present
well-dated sedimentary records from the East Pacific Rise that show evidence of enhanced
hydrothermal activity during the last two glacial terminations. We suggest that glacial maxima
and lowering of sea level caused anomalous melting in the upper mantle and that the
subsequent magmatic anomalies promoted deglaciation through the release of mantle heat
and carbon at mid-ocean ridges.
ea level–driven pressure variations due to
the growth and decay of ice sheets likely
ridge bathymetry reveal Milankovitch-scale frequencies in abyssal-hill spacing, consistent with
Downloaded from http://science.sciencemag.org/ on Septe
lyzes. Water
ace at room
it dissocia(110) surface
was pumped
hat the Cu
atomic resely because
orbed on the
presence of
overed surng water, as
eak in both
hown in Fig.
in Fig. 3C.
t the O peak
tion of H2O
ppears after
se clustering
O pressures
was also obmixtures, as
) surface, on
O, is inactive
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