Main points from this morning
• Continental crust has been extracted from the mantle by
magmatic processes
• Crust is therefore enriched in incompatible elements
(elements with an affinity for magmas relative to residue)
compared to mantle
• Elements can be incompatible because:
1) They're too big to fit in crystal structure, e.g., Large Ion
Lithophile elements: K, Rb, Cs, Sr
2) They have high charge, e.g., High Field Strength
Elements: Nb, Ta, Hf, Zr
Sr and Nd isotopic evolution of crust and mantle over time
87Rb
 87Sr
147Sm
DRb < DSr
Rb (parent) more incompatible
than Sr (daughter)
mantle
high Rb/Sr
mantle
low Rb/Sr
DSm > DNd
Sm (parent) less incompatible
than Nd (daughter)
BSE
143Nd/144Nd
87Sr/86Sr
continental
crust
 143Nd
high Sm/Nd
continental
crust
BSE
low Sm/Nd
time (Gy)
4.55 Gy
time (Gy)
4.55 Gy
Taken together leads to anti-correlation
between 143Nd/144Nd and 87Sr/86Sr
Isochron method
Example of Sm-Nd system
143Nd/144Nd
t2
t1
plagioclase
bulk rock
garnet
t0
147Sm/144Nd
Evolution of
143Nd/144Nd
over time
normalized to 144Nd, a non-radiogenic isotope.)
(147Sm and
143Nd
are
143Nd/144Nd
= 143Nd/144Ndi + 147Sm/144Nd (eλt - 1)
Isochron
equation
t = age
λ= decay constant of 147Sm = 6.54 x 10-12 years-1
i
= initial isotopic ratio
143Nd/144Nd
143Nd/144Nd
garnet
whole rock
plagioclase
Slope = eλt - 1
The age is
determined from the
slope of the
isochron
Initial isotopic ratio =
143Nd/144Nd
i
Provides information
about the source
147Sm/144Nd
An isochron dates a specific event (e.g. recrystallization of a granite)
0.5128
garnet
Isochron age = 500
My = age of rock
recrystallization (for
example)
0.5124
whole
rock
0.5122
0.5120
0.5118
0.5116
plagioclase
Age = 500 My
0.5114
0.5112
0.5110
0
0.05
0.1
0.15
0.2
0.25
0.3
147Sm/144Nd
0.35
0.514
0.513
Model age (TDM) ~ 2.1
Gy = approximate age of
extraction of rock
precursor from depleted
mantle
143Nd/144Nd
143Nd/144Nd
0.5126
Depleted mantle (DM)
0.512
0.511
0.510
Continental
crust
0.509
BSE
0.508
0.507
0.506
0
1
TDM2
3
time (Gy)
4
5
Isotopic variations in the crust and mantle
Radiogenic isotopes
• Nuclear processes
• Absolute dating
• Powerful source tracers
• Not fractionated by geochemical processes (or minor
fractionation is corrected during analysis)
Stable isotope fractionation
• Physicochemical processes
• Varies in a systematic manner with isotope mass
• Sensitive to geochemical processes and temperature
• Source tracers
"Mass independant" isotope fractionation
• Sensitive to highly specific processes
All isotopes of a
given element have
the same electronic
structure.
Harmonic oscillator model of
H2 molecule
Why do they
fractionate?
ZPE =
zero
point
energy
Energy differences
between isotopes are
small but not negligeable.
Separation between nuclei
White (2009)
Isotopic distribution between two phases
light
v = frequency of oscillation
heavy
h = Plancks' constant
m1*m2
M = reduced mass = m +m
1
2
Albarede (2009)
Heavy isotope enters preferentially into phase with
strongest bonds in order to minimize energy
Factors favoring stable isotope
fractionation
•
•
•
•
Low mass
Large mass difference between isotopes
Tendency to form covalent bonds
Existence of several oxidation states (C, N, S) and/or
forming a large variety of compounds (O)
For these reasons, most commonly used stable
isotopes have been O, H, C, N and S.
These are also elements that can be analyzed in a
traditional gas mass spectrometer.
"Non-conventional"
stable isotope
fractionation
MC-ICP-MS
• Introduction of Multicollector Inductively Coupled Mass Spectrometers ~20 years
ago has permitted stable isotope analysis of a wide range of elements: Fe, Mg,
Ca, Li, Ge, Cu, Zn, Ni, Tl, Mo, Se, Si...
• Most useful for studying surface processes, but certain, notably Fe and Mg, can
be significantly fractionated by magmatic processes
• Provide new tracers of crustal material in mantle derived rocks
Describing equilibrium isotopic fractionation
RA and RB= isotopic ratios in phases A and B
α is the "fractionation coefficient"
Example: oxygen isotope fractionation between liquid water and vapor
H218Oliq + H216Ovap <=> H218Oliq + H216Ovap
at equilibrium:
" H 2 18O %
" 18O %
$ 16 '
$ 16 '
# H 2 O & vap # O & vap
= 18
= ( vap )liq
" H 2 18O %
" O%
$ 16 '
$ 16 '
# H 2 O & liq # O & liq
α is a! function of T and P (especially T)
Delta notation
• Since fractionation small, expressed as ‰ variation from a reference (SMOW)
• Facilitates interlaboratory comparison
$ ( 18O/16 O)
'
sample
" 18O = & 18 16
#1) * 1000
& ( O/ O)
)
%
(
SMOW
Fractionation between two phases A and B:
(18O/16 O) A
" OB #" OA = 1000 $ 18 16
(% #1)
( O/ O) SMOW
18
!
!
18
where " B #A
(18O/16 O) A
α is small, so α - 1 ~ ln α; also, 18 16
( O/ O) SMOW ! ~
As a result:
!
" 18OB #" 18OA $ 1000 % ln &
!
1
(18O/16 O) B
= 18 16
( O/ O) A
Equilibrium fractionation
factors α for 18O/16O between
various mineral pairs as a
function of temperature
• Degree of fractionation related
to relative bond strength of the
two phases
• Fractionations small at mantle
temperatures  stable isotope
variations indicate presence of
crustally-derived material
• Oxygen isotopes used as a
geothermometer in metamorphic
processes
White (2009)
Kinetic Isotope Fractionation
Less energy needed to break a bond
with a light isotope than with a heavy
isotope.
α = kB/kA
α is fractionation coefficient
kA and kB are rate constants of isotope
species A and B
Bonds with light isotope are more easily broken
 If reaction does not reach equilibrium reaction
products are enriched in the light isotope
Oxygen isotope compositions in rocks and waters
Albarede (2009)
Mantle has a very uniform composition compared to crust. Allows:
• Tracing crustal contamination of magmas
• Tracing recycling of subducted altered crust in the mantle
Oxygen isotope profile in altered oceanic crust:
Oman ophiolite
Mantle
Low T
High T
Seawater
Gregory and Taylor (1981)
• At low T, strong preference of 18O for solid phase relative to seawater
• At high T, little isotopic fractionation, δ18O closer to seawater value
Tracing recycled
altered oceanic
crust in the
subcontinental
lithospheric mantle
Evidence from
diamond
inclusions
Mantle
Schulze et al. (2003)
Isotopic variations in the crust and mantle
Radiogenic isotopes
• Nuclear processes
• Absolute dating
• Powerful source tracers
• Not fractionated by geochemical processes (or minor
fractionation is corrected during analysis)
Stable isotope fractionation
• Physicochemical processes
• Varies in a systematic manner with isotope mass
• Sensitive to geochemical processes and temperature
• Source tracers
"Mass independant" isotope fractionation
• Sensitive to highly specific processes
Oceanic Crust
Continental Crust
• Thick (~ 40 km on average)
• Thin (6 to 7 km)
• Average composition andesitic
• Average composition basaltic
• Lithologically heterogeneous
• Limited lithologic variation
• No systematic lithologic structure
• Systematic structure
• Very wide age range (0 - 4 Ga)
• Young (<180 Ma)
~ 30% of continental
crust submerged
Rudnick and Gao (2003) after Cogley (1984)
How is average composition of continental crust determined?
Upper crust: getting around heterogeneity problem
1)
Human averaging (weighted average of many sample results
and/or analysis of composite samples). Necessary for major
element and soluble trace element compositions.
2)
Analysis of fine-grained clastic sediments. Does a better job of
averaging, but has problem of elemental fractionation during
weathering and transport. Very useful for insoluble trace element
abundances.
Bulk composition:
• Granodioritic
• Highly enriched in
incompatible elements
such as LREE
Rudnick and Gao (2003)
How is average composition of continental crust determined?
Middle and lower crust: Problem of access
Direct samples
• Exposed high grade
metamorphic terranes - highly
heterogeneous, but mafic
lithologies increase with depth
Lewisian Gneiss, Scotland
• Xenoliths in volcanic rocks) - highly
heterogeneous, but gabbros dominate
in most lower crustal xenolith suites
Kilbourne Hole, New Mexico
How is average composition of continental crust determined?
Middle and lower crust: Problem of access
Geophysical constraints
• Seismic velocities - related to lithology through laboratory
experiments. But dependant on T, P, volatiles, and anistropy as
well as bulk composition
Seismic wave velocities
anticorrelated with SiO2
Seismic wave velocities
increase with depth,
indicating increasingly
mafic compositions.
• Heat flow
• From radioactive decay of K,
U and Th, all highly
incompatible elements
• Implies andesitic average
crustal composition and
increasingly mafic composition
with depth
Heat Production
10-6 W/m3
Crustal Heat
Production
1.5
1
0.5
K (wt%)
• Calculated from measured
surface heat flow minus
mantle heat flow
2
3.5
3
2.5
2
1.5
1
0.5
0
U (ppm)
Crustal heat production:
2.5
0
4
3.5
3
2.5
2
1.5
1
0.5
0
16
Th (ppm)
Variation of crustal
composition with
depth: constraints
from heat flow
3
12
8
4
0
Granite
Granodiorite
Andesite
Gabbro
Peridotite
Inferences concerning crustal major
element composition
• Highly heterogeneous throughout its thickness. For this reason,
estimates of average composition vary between studies.
• In general, crustal composition becomes more mafic with depth,
varying from dominantly granodioritic in upper crust to mainly gabbroic
in lower crust, suggesting an andesitic composition for bulk crust.
• In most Archean cratons, evidence for ancient gabbroic lower crust is
lacking. Possible explanations:
- Archean gabbroic lower crust existed but was removed?
- Mechanism of crust formation was different in Archean?
Mantle melts are basaltic but crust is andesitic. Why?
Most likely explanation:
• Intracrustal differentiation produces gabbroic lower crust.
• As crust thickens, it transforms into eclogite.
• Eclogite is denser than surrounding peridotite and thus sinks into
mantle ("delamination" or "density foundering").
• Greater than 50% of original crustal volume would have to be
removed to create andesitic crust from original basaltic melt.
Other suggested explanations:
• Mixing of silicic melts of subducting slab with basaltic melts of peridotite
• Removal of Mg by weathering, followed by transfer of Mg into altered
oceanic crust which is then subducted
• Removal into mantle of ultramafic cumulates intial crust formation
Various estimates
of lower crustal
trace element
composition
• All parts of crust enriched in
LREE and highly incompatible
elements, but lower crust less
enriched than upper crust.
• Lower crust: positive Eu and Sr
anomalies
• Upper crust: negative Eu and Sr
anomalies
Evidence of intracrustal
differentiation involving
plagioclase accumulation in the
lower crust.
Chondrite normalized
Eu anomalies
Chondrite normalized
Primitive mantle normalized
Trace element
composition of
upper, middle and
lower crust
Upper
Middle
Lower
U/Pbupper crust > U/Pblower crust
Rb/Srupper crust > Rb/Srlower crust
Crustal abundances of Rudnick and Gao (2003)
Trace element composition of bulk continental crust
White (2013)
Increasing incompatibility
Bulk continental crust (CC) is strongly LREE enriched relative to
bulk oceanic crust (OC). Sm/NdCC << Sm/NdOC
Characteristic trace element features of continental crust
Negative Nb
and Ta
anomalies
Positive Pb
anomalies
Increasing incompatibility
Continental crust contains 20 - 70% of incompatible element content of
Bulk Silicate Earth despite representing ~ 0.6% by mass
Continental Crust compared to magmas from main
tectonic settings
Radiogenic isotopes
Sr, Nd and Pb isotopes of continental
crust reflect strong but highly variable
enrichment in incompatible elements
and great age.
Differences between upper
and lower crust reflect
intracrustal fractionation of
Rb/Sr and U/Pb ratios.
Continental Crust - main points
• Bulk composition is andesitic, therefore cannot have
been produced by single step melting of the mantle
• Highly enriched in incompatible elements
• Positive anomalies in Pb, negative anomalies in Nb
and Ta
• Highly variable isotopic compositions, reflecting
incompatible element enrichment integrated over
very long time periods