Ch 10 magma generation

Mantle MeltingMantle Melting
Thanks to John WinterThanks to John Winter

Magmatic ProcessesMagmatic Processes
Magmas that reach the Earth’s surface to form lavas are highly
varied, ranging in composition from ultramafic komatiites, through
basalts and andesites, to rhyolites and feldspathoidal felsic rocks.
Although the compositions of these lavas may not represent all
magmas formed in the Earth, they do indicate the enormous
diversity of magmas.
Explaining the origin of this diversity has been the dominant goal of
petrology.
Early in the history of science, most of the different magmas were
thought to have independent origins; some were interpreted as
products of magma mixing and still others the products of magma
splitting.

From the beginning of 20th
century, process of magma differentiation
has been interpreted to give rise to variety of magmas.
N. L. Bowen (1928), championed this interpretation and argued that
basalt was the primary magma from which other magmas were
derived.
Lavas erupted from a single volcano at different times can reasonably
be interpreted to share a common origin.
The same is true of different igneous rocks within a single pluton or
even in separate, but closely related bodies, especially when the rock
types share some striking geochemical signature.
Igneous rocks that are related to a common source are said to be
comagmatic.

On the basis of field evidence, geochemistry or experimental work, a
particular magma might be interpreted as being derivative from
another magma, which can be referred to as parental magma.
That parent may, in turn, have been derived from another magma, and
so on.
As the lineage is traced backward toward the primary magma,
magmas are said to become more primitive.
Conversely, magmas late in the lineage are said to be evolved or
differentiated.
Although easy to define, primary magmas are, in fact, very difficult to
identify, because of the uncertainty in the composition of the source
region and hence in the composition of magma with which it would be

Primary magmas are those derived directly by partial melting of some
source (preexisting rock), and have no characteristics that reflect the
effects of subsequent differentiation.
Those magmas that have experienced some form of chemical
differentiation are referred to as evolved or derivative magmas.
Primitive magmas are unmodified magmas that form through anatexis of
mantle rocks that have not been melted or otherwise changed in
composition since they formed.
The parental magma is the most primitive one found in an area, and thus
the one from which we suppose the others are derived.
Primary magmas have (MgO/MgO+FeO) = 0.66-0.75 (Mg#) and high
contents of Cr (>1000ppm) and Ni (>400-500ppm).

Mantle meltingMantle melting
When we begin talking about magma generation, we have to talk
about partial melting of the mantle.
Common product of this melting is basalt, by far the most
common volcanic rock generated today.
Much of the spectrum of igneous rock types can be derived from
primitive basaltic material by some evolutionary process, such as
fractional crystallization, assimilation etc.
The generation of basaltic magma is thus a critical step in developing
a comprehensive understanding of magma genesis.
Because basalts occur most commonly in the ocean basins, and the
Earth's structure is simplest in the oceanic areas, lacking the physical
and chemical complexities imposed by the continents, lets see what is

2 principal types of basalt in2 principal types of basalt in
the ocean basinsthe ocean basins
Common petrographic differences between tholeiitic and alkaline basalts
Tholeiitic Basalt Alkaline Basalt
Usually fine-grained, intergranular Usually fairly coarse, intergranular to ophitic
Groundmass No olivine Olivine common
Clinopyroxene = augite (plus possibly pigeonite) Titaniferous augite (reddish)
Orthopyroxene (hypersthene) common, may rim ol. Orthopyroxene absent
No alkali feldspar Interstitial alkali feldspar or feldspathoid may occur
Interstitial glass and/or quartz common Interstitial glass rare, and quartz absent
Olivine rare, unzoned, and may be partially resorbed Olivine common and zoned
Phenocrysts or show reaction rims of orthopyroxene
Orthopyroxene uncommon Orthopyroxene absent
Early plagioclase common Plagioclase less common, and later in sequence
Clinopyroxene is pale brown augite Clinopyroxene is titaniferous augite, reddish rims
after Hughes (1982) and McBirney (1993).
Tholeiitic Basalt and Alkaline Basalt

Tholeiites are generated at mid-ocean ridges
Also generated at oceanic islands,
subduction zones
Alkaline basalts generated at ocean islands
Also at subduction zones
Each is chemically distinct

Sources of mantle material
Ophiolites
Slabs of oceanic crust and upper mantle
Thrust at subduction zones onto edge of continent
Dredge samples from oceanic crust
Nodules and xenoliths in some basalts
Kimberlite xenoliths
Diamond-bearing pipes blasted up from the
mantle carrying numerous xenoliths from depth

Petrology of the mantlePetrology of the mantle
Measured seismic velocities in the upper mantle
are compatible with a rock made of olivine,
pyroxene, and garnet.
These three phases are composed essentially of
the five major chemical components: SiO2,
Al2O3, FeO, MgO, and CaO.
Upper mantle densities are about 3.35g/cm3
,
which is more consistent with peridotite than
with olivine-free, pyroxene-garnet rock
(eclogite, 3.5g/cm3
).

1) Ophiolites:
These are large sheet-like mafic to ultramafic masses, presumed to be
ancient oceanic crust and upper mantle thrust onto the edge of
continents and/or incorporated into mountain belts.
Ophiolites show a considerable range in size, thickness, and degree
of structural integrity.
For now we are primarily interested in the ultramafic rocks in the
lower portions, because these are believed to represent significant
portions of the upper mantle now exposed at the surface of the Earth.

Smaller slivers of presumed ophiolitic ultramafics, now dismembered
and incorporated into deformed mountain belts, are commonly
referred to as alpine peridotites.
The ultramafic portions contain a variety of peridotites, predominantly
harzburgite and dunite, with subordinate wehrlite, lherzolite, and
pyroxenite.
The mineralogy is dominated by Ol, Opx and Cpx, with lesser
amounts of plag and oxide minerals, including mt, il, and Cr-rich
spinel.
Hornblende and serpentine appear to be later hydrous replacement
minerals.
The larger, more intact ophiolites allow us to see the geometric
relationships between various rock types, but our observations are

2) Dredge samples from oceanic fracture zones
Differences in ridge elevation can result in significant scarps at
many ridge-offsetting fracture zones (transform faults).
As with ophiolites, these samples represent only the uppermost
mantle beneath this oceanic crust.
It is also impossible to know the exact location of a dredged sample
or the relationships between any two samples.
Dredged samples are varied but the types are nearly identical to
those exposed in ophiolites, providing strong evidence that ophiolites
are indeed samples of oceanic crust and upper mantle.

3) Nodules in basalts
Ultramafic xenoliths, called nodules, are occasionally carried to the
surface by basalts.
They are usually fist-sized or smaller, and the most common rock types
are gabbro, dunite, harzburgite, spinel lherzolite, wehrlite, garnet
lherzolite, and eclogite.
Some lower crustal xenoliths are also found in many basaltic lavas.
Many of the nodules are autoliths, or cognate xenoliths, meaning
they are genetically related to the magma, and not picked up from the
wall rocks far from the magma source.

Some of these may be cumulates (eg gabbros and pyroxenites) and
others restites, a refractory residuum left behind after partial melts have
been extracted.
Because the basalts have a mantle source, restites are olivine-rich.
In order to carry such dense olivine-rich nodules in suspension to the
surface, transport was apparently rapid.
These nodules are restricted to alkaline basalts and are not found in the
more common tholeiites, implying that the former travel more rapidly,
and have had less time to crystallize or interact with the wall rocks.

This has led several petrologists to conclude that alkali basalts are
more primitive than tholeiites.
The high-pressure garnet-bearing lherzolites occur only in the most
alkaline and silica-deficient basalts, suggesting that perhaps these
basalts have a deeper origin than the less alkalic and more silica-
saturated basalts.
Spinel lherzolites-Alkalic basalts, basanites.
Garnet lherzolites, harzburgites and eclogites-Kimberlites

4) Xenoliths in kimberlite bodies
Kimberlites are unusual igneous phenomena.
Several lines of evidence suggest that kimberlites tap an upper mantle source
as deep as 250 to 350 km, and travel rapidly to the surface, bringing a variety
of mantle and crustal samples to the surface as xenoliths.
All known kimberlites occur in continental areas, so the xenoliths represent
continental crust and subcrustal mantle.
Nonetheless, the mantle at depth, at least, is believed to be similar in both
continental and oceanic areas.
Garnet lherzolite and spinel lherzolite are dominant among the unaltered
deeper kimberlite samples.

Data from the above sources lead us to believe that the
mantle is composed of ultramafic rocks.
Most students of the mantle believe that "typical" mantle is
composed of peridotite.
Spinel and garnet peridotites stand out from the array of
mantle samples as prime suspects for pristine mantle
material when partially melted, can yield a basaltic liquid.
Much of the shallowest oceanic mantle, now represented by
dunite and harzburgite in ophiolites and many nodules in
basalts, appears to be related to the lherzolites as refractory
residuum after basalt has been extracted.

More specifically, it is a four-phase lherzolite, composed of Mg-
rich olivine and lesser amounts of pyroxene, usually both Ca-Mg-
rich clinopyroxene and Ca-poor, Mg-rich orthopyroxene; these
three crystalline phases are a stable assemblage to a depth of about
410 km.
Most peridotites contain more Al in the bulk rock than can be held
in solid solution in pyroxenes and olivine, thus stabilizing a
separate minor Al-rich phase, whose nature depends on P and less
on T.
At less than about 8 kbar (30-km depth), the stable Al-rich phase
is plagioclase; from there to about 25 kbar (roughly 75 km,
depending on crustal thickness), it is spinel; and at still higher P it
is garnet.

Garnetiferous peridotites equilibrated at depths greater than about 150
km and occurring as xenoliths in some kimberlitic rock also contain
diamond.
Diamonds commonly contain minute inclusions of minerals stable in
the upper mantle (pyroxene, garnet).
Extremely rare diamonds contain inclusions stable only at ~670 km-
depth in the deep mantle (Mg-Fe-Ca-Al perovskite), testifying to the
depth from which at least some kimberlitic magmas carrying suspended
diamonds are derived.
At shallower depths where diamond is not stable, the stable C-bearing
phase may be graphite.

Phase diagram for aluminous
4-phase lherzolite:
Plagioclase
shallow (< 50 km)
Spinel
50-80 km
Garnet
80-400 km
Si → VI coord.
> 400 km
Al-phase =
Figure 10.2 Phase diagram of aluminous lherzolite with melting interval (gray), sub-solidus reactions,
and geothermal gradient. After Wyllie, P. J. (1981). Geol. Rundsch. 70, 128-153.

This sequence of reactions explains how we can have
compositionally equivalent spinel and garnet lherzolites,
and also tells us that plagioclase lherzolites are a low-
pressure alternative, also with the same chemical
composition.
It further explains why plagioclase, spinel, and garnet
are rarely found together in the same sample, and why
plagioclase lherzolites are found only in shallow mantle
samples (ophiolites and some oceanic basalts), whereas
garnet lherzolites occur more commonly in kimberlites
that tap a deeper mantle source.

Because plagioclase peridotites are limited to depths less than about 30
km, which is less than the thickness of much of the continental crust, we
would expect plagioclase peridotite to be absent in most of the sub-
continental mantle, the top of which is commonly deeper than 30 km.
This explains why it is so rare in kimberlites. The transition from
plagioclase to spinel peridotite, and from spinel to garnet peridotite are
accomplished by the following idealized metamorphic reactions:

15
10
5
0
0.0 0.2 0.4 0.6 0.8
Wt.%Al2O3
Wt.% TiO2
Dunite
Harzburgite
Lherzolite
Tholeiitic basalt
Partial M
elting
Residuum
Lherzolite is probably fertile unaltered mantle
Dunite and harzburgite are refractory residuum after basalt has been
extracted by partial melting
Note that the compositions of
Du, Hz, Lz, and tholeiitic
basalt are collinear, and that
Lz composition lies at an
intermediate position between
the basalt and the other two
rock types.
If the tholeiite is created by
partial melting of the Lz,
extraction of the liquid will
shift the composition of the
remaining material directly
away from the tholeiite toward
harzburgite and dunite.

Lherzolite: A
type of
peridotite with
Olivine > Opx
+ Cpx
Olivine
ClinopyroxeneOrthopyroxene
Lherzolite
Harzburgite
Wehrlite
Websterite
Orthopyroxenite
Clinopyroxenite
Olivine Websterite
Peridotites
Pyroxenites
90
40
10
10
Dunite
Figure 2.2 C After IUGS
The aluminous lherzolite represents undepleted mantle, also called fertile mantle, with a
composition presumed to be close to that of original mantle

Melting of the mantleMelting of the mantle
Now that we have an idea of the chemical and mineralogical nature of
the mantle, let's return to our original question regarding the feasibility
of mantle melting.
The geotherm does not intersect the solidus for fertile mantle
lherzolites.
There are a number of estimates for the average oceanic geotherm, but
none of them approach the solidus.
So our first problem, because we know that basalts are indeed
generated, is to figure out how the mantle can be melted.
There are three basic ways to accomplish this goal, following the three
principal natural variables.
We can either raise the temperature, lower the pressure, or change the

Melting of the mantleMelting of the mantle

Raising the temperatureRaising the temperature
Heat transfer associated with convective or advective movement of rock or magma is an
important means of raising rock temperatures above the solidus in large rock volumes.
Two global plate tectonic regimes where this takes place include:
1. The descending oceanic lithosphere in subduction zones absorbs heat from the
surrounding hotter mantle. Consequently, partial melting of the less refractory
basaltic crust might occur in young, still hot subducting lithosphere, as contrasted
with subduction of old, cold lithosphere. Shear heating of rapidly subducted
lithosphere can also promote higher temperatures.

2) Already hot, deep continental crust can be heated in excess of its solidus T by
juxtaposition of hotter mantle-derived magma above subducting plates and upwelling
decompressing mantle. Density constraints indicate that ascending basaltic magma
can be arrested at the base of the less-dense feldspathic crust, underplating it, or
perhaps stagnating not far into it above the Moho.

How does the mantle melt??
1) Increase the temperature
Figure 10.3. Melting by raising the temperature.

Perhaps the simplest way to heat the rock is to accumulate enough heat
by the decay of radioactive elements, because this is the only known
source of heat other than that escaping from the primordial
differentiation process.
The prime radioactive elements (K, U, and Th) occur in such low
concentrations in the mantle that they are not capable of producing
enough heat required for melting.
The thermal conductivity of rocks is pretty low, but certainly high
enough to allow this heat to dissipate long before any rocks would even
approach melting.
In fact, this radioactive generation and conduction of heat is exactly the
process responsible for up to half of the heat flow reaching the surface
that creates the geotherm in the first place.

If, in the unlikely event we did somehow accomplish this
concentration and manage to heat the mantle to its solidus, the heat
required for melting must supply the latent heat of fusion of the
minerals, which is about 300 times greater than the specific heat
required to bring the minerals up to the melting point.
The job of producing sufficient melts to be extractable thus becomes
even more difficult.
Additionally, the first melting episode would extract much of the
incompatible K, Th, and U from the source rock, so that subsequent
heat generation would be greatly retarded.
Therefore, radioactive decay alone, in most instances, cannot produce
sufficient perturbation in T to cause magma generation.

Tectonic Thickening: Thrusting and folding thicken continental crust
to >50 km in active orogenic (mountain) belts in subduction zones.
This can lead to supersolidus temperatures in the deep crust.
Several factors are intertwined in complex ways in this heating.
Radioactive heat production in the thickened continental crust is
enhanced, as is the insulating effect on the flux of heat from the
mantle.
The 600°C temperature at the base of a 30-km-thick crust is
subsolidus for most potential source rocks, but, after thickening to,
say, 50 km and readjustment to the same average geotherm of
20°C/km, the T at the base of the crust is 1000°C, which is well above
the solidus of many crustal rocks.

Optimal “incubation” depends on the rate of
thickening and may be curtailed by prompt rapid
uplift and erosion.
Thermal modeling that takes into account these and
other factors suggests optimal heating occurs on
the order of a few tens of millions of years after
orogenic thickening.

The most obvious manifestations of locally high heat flow are the hot
spots, such as Hawaii, which are above narrow pipe-like conduits of
basaltic magma that appear to have a stationary source in the mantle.
The motion of plates over these stationary hot spots results in the
apparent migration of the volcanic activity across the plate through
time, and it has been used to determine "absolute" plate motions.
The origin of these hot spots is not known, but is popularly attributed
to very deep processes present at the base of the mantle, and perhaps
related to heat production or convection in the liquid core.
Of course attributing problems to a deeper source is the simplest way
to avoid them in the geological literature, but it seems to be justified in
this case.

Lowering the PressureLowering the Pressure
The simplest idea, local zones of lower pressure (the pressure
equivalent of hot spots?) is untenable in ductile material such as the
mantle, because high-pressure material would quickly flow to the low-
pressure areas until lithostatic equilibrium were attained.
A more plausible way to lower pressure is to raise mantle rocks to
shallower levels while maintaining their stored heat content.
When material moves upward, the pressure is reduced and the volume
increases slightly, resulting in a slight temperature reduction (10-
20°C/GPa or 0.3-0.6°C/km for mantle rocks.
The upwelling mass would also move into cooler areas, and lose heat
by conduction to the surroundings, so it would then simply follow the
geotherm, never approaching its melting point.

If, on the other hand, the rise were sufficiently rapid to
minimize heat loss to the surroundings, the only
temperature difference would be due to expansion.

If conductive heat loss were zero, the process is referred to as
adiabatic, and any rising rock material would follow a curve
with the ~12°C/GPa slope, called the adiabat.

Of course, a purely adiabatic process is not likely, but upwelling of
mantle material at geologically acceptable rates would be close,
following a curve with a steeper slope than the melting curve
(~130°C/GPa), and could initiate melting.

Once melting begins, the latent heat of fusion will absorb heat from
the rising mass, and cause the adiabatic path to follow a shallower
temperature/pressure slope closer to the solidus curve, thus
traversing the melting interval more obliquely.

As a result, upwelling mantle material will diverge from
the solidus slowly, producing limited quantities of melt.
The process is called decompression melting.

Upwelling of mantle material occurs at divergent plate boundaries,
where two plates are pulling apart, and mantle material must flow
upward to fill in.
Langmuir et al. (1992) calculated that about 10 to 20% melting will
occur per GPa (0.3%/km) of continued pressure release.
Thus, if the upwelling material began at its solidus temperature, it
would have to rise about 65 to 100 km to attain the 20 to 30%
melting estimated to produce mid-ocean ridge basalt.
Of course this would have to be added to the rise necessary to bring
it to its solidus temperature in the first place, which would be on
the order of 150 km.

Addition of VolatilesAddition of Volatiles
Among the possible changes in chemical composition of a solid rock system
that could induce melting on a global scale, increase in water concentration, is
the most significant.
Very locally, an increase in CO2 or other volatiles may induce melting in
already hot rock.
Even small increases in water concentration can significantly depress the
solidus.
In some mantle xenoliths we find phlogopite or an amphibole.
They are minor phases, but they do attest to the presence of some H2O in the
mantle.
Wyllie (1975) estimated that the amount of H2O in normal mantle material is
unlikely to exceed 0.1 wt.%, and suggested that it is not uniformly distributed
(the mantle is probably much more hydrated in subduction zones, however).

Addition of VolatilesAddition of Volatiles
Microscopic examination reveals fluid
inclusions, up to 5 μm in diameter, most of
which contain H2O, but many are filled with
dense liquid CO2.
This, and the occurrence of carbonate inclusions
in some mantle minerals and in the matrix of
kimberlites, suggests that CO2 is also present in
the mantle.

3) Add volatiles (especially H2O)
Dry peridotite solidus compared to several experiments on H2O-saturated peridotites.

Fraction melted is
limited by the
availability of water
15% 20% 50% 100%
Figure 7.22. Pressure-temperature projection of the
melting relationships in the system albite-H2
O. From
Burnham and Davis (1974). A J Sci., 274, 902-940.

Heating of amphibole-bearing peridotite
1) Ocean geotherm
2) Shield geotherm
Figure 10.6 Phase diagram (partly schematic)
for a hydrous mantle system, including the
H2O-saturated lherzolite solidus of Kushiro et
al. (1968), the dehydration breakdown curves
for amphibole (Millhollen et al., 1974) and
phlogopite (Modreski and Boettcher, 1973),
plus the ocean and shield geotherms of Clark
and Ringwood (1964) and Ringwood (1966).
After Wyllie (1979). In H. S. Yoder (ed.), The
Evolution of the Igneous Rocks. Fiftieth
Anniversary Perspectives. Princeton University
Press, Princeton, N. J, pp. 483-520.

Melts can be created under realistic
circumstances
Plates separate and mantle rises at mid-ocean ridges
Adibatic rise → decompression melting
Hot spots → localized plumes of melt
Fluid fluxing may give LVL
Also important in subduction zones and other settings

Generation of tholeiitic and
alkaline basalts from a
chemically uniform mantle
Variables (other than X)
Temperature
Pressure
Figure 10.2 Phase diagram of aluminous lherzolite with
melting interval (gray), sub-solidus reactions, and
geothermal gradient. After Wyllie, P. J. (1981). Geol.
Rundsch. 70, 128-153.

Pressure effects:
Figure 10.8 Change in the eutectic (first
melt) composition with increasing
pressure from 1 to 3 GPa projected
onto the base of the basalt tetrahedron.
After Kushiro (1968), J. Geophys. Res.,
73, 619-634.

Liquids and residuum of melted pyrolite
Figure 10.9 After Green and Ringwood (1967). Earth Planet. Sci. Lett. 2, 151-160.

Initial Conclusions:
Tholeiites favored by shallower melting
25% melting at <30 km → tholeiite
25% melting at 60 km → olivine basalt
Tholeiites favored by greater % partial melting (F)
20 % melting at 60 km → alkaline basalt
incompatibles (alkalis) → initial melts
30 % melting at 60 km → tholeiite

Crystal Fractionation of magmas
as they rise
Tholeiite → alkaline
by FX at med to high P
Not at low P
Thermal divide
Al in pyroxenes at Hi P
Low-P FX → hi-Al
shallow magmas
(“hi-Al” basalt)
Figure 10.10 Schematic representation of the fractional
crystallization scheme of Green and Ringwood (1967)
and Green (1969). After Wyllie (1971). The Dynamic
Earth: Textbook in Geosciences. John Wiley & Sons.

Figure 10.11 After Kushiro (2001).
Other, more recent experiments on melting of fertile (initially garnet-
bearing) lherzolite confirm that alkaline basalts are favored by high
P and low F

Primary magmasPrimary magmas
Formed at depth and not subsequently modified
by FX or Assimilation
Criteria
Highest Mg# (100Mg/(Mg+Fe)) really → parental
magma
Experimental results of lherzolite melts
Mg# = 66-75
Cr > 1000 ppm
Ni > 400-500 ppm
Multiply saturated

Multiple saturationMultiple saturation
Low P
Ol then Plag then
Cpx as cool
~70o
C T range
Figure 10.13 Anhydrous P-T phase relationships for
a mid-ocean ridge basalt suspected of being a
primary magma. After Fujii and Kushiro (1977).
Carnegie Inst. Wash. Yearb., 76, 461-465.

Low P
Ol then Plag then Cpx
as cool
70o
C T range
High P
Cpx then Plag then Ol
Figure 10.13 Anhydrous P-T phase relationships for
a mid-ocean ridge basalt suspected of being a
primary magma. After Fujii and Kushiro (1977).
Carnegie Inst. Wash. Yearb., 76, 461-465.

Low P
Ol then Plag then Cpx
as cool
70o
C T range
High P
Cpx then Plag then Ol
25 km get all at once
= Multiple saturation
Suggests that 25 km is
the depth of last eqm
with the mantle

Summary
A chemically homogeneous mantle can
yield a variety of basalt types
Alkaline basalts are favored over tholeiites
by deeper melting and by low % PM
Fractionation at moderate to high depths can
also create alkaline basalts from tholeiites
At low P there is a thermal divide that
separates the two series

0.00
2.00
4.00
6.00
8.00
10.00
atomic number
sample/chondrite
La Ce Nd Sm Eu Tb Er Yb Lu
increasing incompatibility
Review of REEReview of REE

Review of REEReview of REE
Figure 9.4. Rare Earth
concentrations (normalized to
chondrite) for melts produced at
various values of F via melting of a
hypothetical garnet lherzolite using
the batch melting model (equation
9-5). From Winter (2001) An
Introduction to Igneous and
Metamorphic Petrology. Prentice
Hall.

REE data for oceanic basalts
Figure 10.14a. REE diagram for a typical alkaline ocean island basalt (OIB) and tholeiitic mid-
ocean ridge basalt (MORB). From Winter (2001) An Introduction to Igneous and Metamorphic
Petrology. Prentice Hall. Data from Sun and McDonough (1989).

Spider diagram for oceanic basalts
Figure 10.14b. Spider diagram for a typical alkaline ocean island basalt (OIB) and tholeiitic
mid-ocean ridge basalt (MORB). From Winter (2001) An Introduction to Igneous and
Metamorphic Petrology. Prentice Hall. Data from Sun and McDonough (1989).

Suggests different mantle source
types, but isn’t conclusive.
Depleted mantle could → both
MORB and OIB.

REE data
for UM
xenoliths
Figure 10.15 Chondrite-normalized REE diagrams for
spinel (a) and garnet (b) lherzolites. After Basaltic
Volcanism Study Project (1981). Lunar and Planetary
Institute.
LREE enriched
LREE depleted
or unfractionated
LREE depleted
or unfractionated
LREE enriched

Review of Sr isotopes
87
Rb → 87
Sr λ = 1.42 x 10-11
a
Rb (parent) conc. in enriched reservoir (incompatible)
Enriched reservoir
Figure 9.13. After Wilson (1989). Igneous Petrogenesis. Unwin
Hyman/Kluwer.
develops more
87
Sr over time
Depleted reservoir
(less Rb)
develops less
87
Sr over time

Review of Nd isotopes
147
Sm → 143
Nd λ = 6.54 x 10-13
a
Nd (daughter) → enriched reservoir > Sm
Enriched reservoir
develops less
143
Nd over time
Depleted res.
(higher Sm/Nd)
develops higher
143
Nd/144
Nd
over time
Nd Sm
REE diagram
Figure 9.15. After Wilson (1989). Igneous Petrogenesis. Unwin Hyman/Kluwer.

Nd and Sr isotopes of Ocean Basalts
“Mantle Array”
Figure 10.16a. Initial 143
Nd/144
Nd vs. 87
Sr/86
Sr for oceanic basalts. From Wilson (1989). Igneous Petrogenesis.
Unwin Hyman/Kluwer. Data from Zindler et al. (1982) and Menzies (1983).

Nd and Sr isotopes of Kimberlite Xenoliths
Figure 10.16b. Initial 143
Nd/144
Nd vs. 87
Sr/86
Sr for mantle xenoliths. From Wilson (1989). Igneous Petrogenesis.
Unwin Hyman/Kluwer. Data from Zindler et al. (1982) and Menzies (1983).

“Whole Mantle” circulation model
Figure 10-17a After Basaltic Volcanism Study Project (1981). Lunar and Planetary Institute.

Upper depleted mantle = MORB source
Lower undepleted & enriched OIB source
“Two-Layer” circulation model
Figure 10-17b After Basaltic Volcanism Study Project (1981). Lunar and Planetary Institute.

Experiments on melting enriched
vs. depleted mantle samples:
Tholeiite easily created
by 10-30% PM
More silica saturated
at lower P
Grades toward alkalic
at higher P
1. Depleted Mantle
Figure 10-18a. Results of partial melting experiments on depleted
lherzolites. Dashed lines are contours representing percent partial
melt produced. Strongly curved lines are contours of the normative
olivine content of the melt. “Opx out” and “Cpx out” represent the
degree of melting at which these phases are completely consumed
in the melt. After Jaques and Green (1980). Contrib. Mineral.
Petrol., 73, 287-310.

Experiments on melting enriched vs.
depleted mantle samples:
Tholeiites extend to
higher P than for DM
Alkaline basalt field
at higher P yet
And lower % PM
2. Enriched Mantle
Figure 10-18b. Results of partial melting experiments on fertile
lherzolites. Dashed lines are contours representing percent partial
melt produced. Strongly curved lines are contours of the normative
olivine content of the melt. “Opx out” and “Cpx out” represent the
degree of melting at which these phases are completely consumed
in the melt. The shaded area represents the conditions required for
the generation of alkaline basaltic magmas. After Jaques and
Green (1980). Contrib. Mineral. Petrol., 73, 287-310.

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