Download Plate Tectonics: This works as a result of hot mantle asthenosphere

Plate Tectonics: Geological Aspects
Prof. J. TARNEY
Wednesdays at 9.30 (Lecture) & Wednesdays at 11.30 (2 hr Practical or Lecture)
Program Second Semester 1998:
Wed Lecture 9.30 - 10.30
Wed Practical/Lecture 11.30 - 1.30
Week 6 Mantle Topology & Mineralogy (L)
Mid-Ocean Ridge Processes (L)
Week 7 Rifting and Wilson Cycle (L)
Practical: Pacific Ocean
Week 8 Extension & Sedimentary Basins (L)
Subduction and Island Arcs (L)
Practical: Mini practical
Practical: Plate model exercises (1)
Week 9 Active Margins & Accretion (L)
Week 10 Plumes & plateaus (L)
Collision; plate tectonics &
Earth evolution
Practical: Plate model exercises (2)
Useful General References:
PRESS, F. & SIEVER, R. 1998. Understanding Earth (2nd Ed.) p.504–535
+ CD-rom & Internet link: www.whfreeman.com/understandingearth
KEARY, P. & VINE, F.J. 1991. Global Tectonics. Blackwell Scientific Publ., 302pp..
(2nd Edition now out)
Plate Tectonics: This
works as a result of hot
mantle asthenosphere
ascending beneath midocean ridges to form the
plates. These plates then
subside into the mantle
again at subduction
zones, pulled by the
excess density of the
mafic ocean crust, which
transform to eclogite.
Plate Tectonics: GL209
Prof. John Tarney
Lecture 1: Geological Aspects
mafic as a result of extraction of silicic granitic magmas from the
deep crust?
There a number of regions where we think tectonic activity
has brought segments of the lower crust up for inspection. Notable
examples are Kapuskasing and Pikwitonei in Canada (Precambrian
crust), the Lewisian of NW Scotland (also Precambrian), Calabria in
S. Italy and the Ivrea Zone in the Alps (both Phanerozoic).
MANTLE PETROLOGY IN RELATION TO PLATE
TECTONICS
Knowledge of mantle petrology and the constitution of the deeper
mantle is important in trying to understand several aspects of plate
tectonics. For instance, is there whole-mantle convection or twolayer convection? What are mantle plumes? What are superplumes?
Does the subducting slab penetrate into the lower mantle? What
happens to the slab at depth? Is the sub-continental mantle different
from the oceanic mantle? First, some basic facts.
Principal Internal Subdivisions of the Earth
Region
Depth
(km)
Mass
Mass
(1025g) Fraction
Crust
Upper Mantle
Transition Zone
Lower Mantle
Outer Core
Inner Core
0-Moho
Moho-400
400-1000
1000-2900
2900-5154
5154-6371
2.4
62
1000
245
177
12
0.004
0.10
0.17
0.41
0.30
0.02
Fig. 2. Diagram (based on field and geophysical studies) to show
how deep crust can be thrust up to high crustal levels. Kapuskasing
structure, Ontario, Canada).
The Lewisian of NW Scotland can be interpreted similarly, the
high-grade lower crustal granulites being thrust over amphibolites
>2.5 Gyr ago.
Note that the crust makes up quite a small proportion of the
total Earth. The main problems that have occupied geologists over
the years are: What is the nature of the crust-mantle boundary (the
MOHO). What is the nature of the low velocity zone? Is the
lithosphere diferent in composition from the asthenosphere? What
happens in the transition zone? What is the nature of the deep
mantle?
Fig. 3. Estimates as to the extent to which we have sections
thru two segments of Archaean crust and one segment of
‘recent’ Alpine crust. There is uncertainty as to whether
actually we can directly sample the lowest crust. The Alpine
sequence is in a series of thrust slices, but some rocks in
region of the Ivrea zone have been down to greater depths
and have re-bounded again to the surface.
Perhaps the more important question is what causes high-P
rocks to exhume?
The Continental Crust
Though we know quite a lot about the upper crust, there is
still quite a lot of uncertainty about the lower crust. Is there a real
Conrad discontinuity separating the lower from the upper crust? Is
the lower crust made up of dry granulite-facies rocks. Is it more
mafic than the upper crust, perhaps as a result of intrusion of mafic
magmas into the lower crust (called "underplating"). Or is it more
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Plate Tectonics: GL209
Prof. John Tarney
Lecture 1: Geological Aspects
The Moho and the Lower Crust
In the early 1960s there was considerable discussion
between petrologists and geophysicists as to the nature of the
MOHO. The P-wave velocity of most regions of the uppermost
mantle beneath both continents and oceans lies in the range 8.2 ±0.2
km/sec. This in itself would restrict the composition of the mantle
below the Moho to some combination of the following minerals
(which have the appropriate properties):
Olivine Pyroxenes Garnet (minor spinel, hornblende,
phlogopite)
The two principal rock types carrying these minerals are:
PERIDOTITE (olivine + pyroxenes) and ECLOGITE (pyroxene +
garnet), which are of ultramafic and mafic composition respectively.
The MOHO and the gabbro-eclogite transformation: Basalt
transforms to eclogite at high pressures according to the equation:
The nature of the lower crust is less certain. Exhumed
Precambrian high-grade granulite-facies rocks (as in Figs. 2 & 3),
which have equilibrated at depths of 25-30 km, have an intermediate
(dioritic) bulk composition. But deep crustal xenoliths brought up in
volcanic breccia pipes tend to be more mafic (gabbroic) in
composition. So has the continental crust in part been underplated
subsequently by mafic magma? Also, recent deep seismic
investigations have revealed strong horizontal reflections in the deep
(mainly post-Archaean) crust - what is the cause of these reflections?
Do they represent mafic intercalations, differences in fluid content,
crustal viscosity, or the bottoming out of shear zones (cf. Kusznir &
Matthews, 1988; Meissner & Kusznir, 1987; Warner, 1990; Reston,
1990a). Comparison of crustal reflection profiles across the various
tectonic zones of Europe (Wever et al., 1987; Sadowiak et al., 1991)
has identified several different types of deep crustal structure: (a)
abundant lamellae above the Moho as in BIRPS SWAT 4, (b) bands
of reflectors as in BIRPS WINCH 3, (c) hyperbola-like diffractions
as in BIRPS SWAT 6-9, (d) "crocodile" diverging reflectors
observed in old collision zones, but not so far in the UK, (e) "ramp
and flat" stuctures as in BIRPS SWAT 4 2/3, and (f) "fishbone"
pattern observed across the Brabant massif. These features (cf.
Meissner, 1989) are thought to represent different types of crustal
structure. However it is felt that lower crustal viscosity (Meissner &
Kusznir, 1987) is a more important factor controlling development of
reflectors than is composition, and the current view is that good
lower crust reflectors might characterise mature crust, but that this
pattern could be destroyed by either compression/collision
("crocodiles", etc.), igneous intrusions or significant extension.
Nonetheless it is commonly assumed that the lower crust is
gabbroic in composition, either igneous (gabbro), or metamorphic
amphibolite (wet) or granulite (dry).
Olivine + pyroxene + plagioclase > jadeitic pyroxene + garnet
For this to be capable of explaining the MOHO it must be a
relatively sharp transition, of no more than a few km. Green and
Ringwood (1967) studied this experimentally to 30 kilobars (= ca
100 km) using a quartz tholeiite and an alkali basalt (Fig 5).
The MOHO as a phase transition
Fig. 4 shows various suggestions (made at one time or other) for how
the MOHO beneath oceans and continents could be a phase transition
(change of mineralogy, but not a major change in composition).
Serpentine is the hydrated variety of peridotite (with ca 12% water,
thus lower density). Eclogite is the highpressure form of basalt or gabbro. But are these models realistic? The
serpentine-peridotite model is now discounted.
Fig. 5. Experimental studies of Green & Ringwood (1967)
on quartz tholeiite basalt showed that transformation to
eclogite occurred over a considerable depth interval. Note
that eclogite has a lot more quartz than the equivalent
basalt. But mantle eclogites have no quartz. Where does the
silica go?
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Plate Tectonics: GL209
Prof. John Tarney
With both compositions the transformation was found to be gradual.
The disappearance of the low density phase (plagioclase) and its
replacement by the high density phases (garnet and jadeite) occurred
over a pressure range of ca 10 kb (= ca 25 km). Thus the MOHO
cannot be a phase transformation and must be a compositional
transition.
Lecture 1: Geological Aspects
(4) Modelling "backwards" from erupted basalt compositions.
(1) Olivine rich nodules are quite common in erupted alkali basalts
worldwide, and are almost all of spinel lherzolite composition
(olivine, orthopyroxene, Cr-diopside, spinel). In kimberlite
(diamond) pipes there is a greater diversity in that both garnet
peridotite and eclogite xenoliths occur (the former dominant). Some
nodules contain diopside-rich, phlogopite-rich or amphibole-rich
veins. It is thought that these nodules are representative of the subcontinental mantle. This material seems to be rather refractory (could
not yield much basalt on melting), but at the same time can be quite
enriched in incompatible trace elements such as Sr, Ba, K, Rb and
the light rare- earths.
(2) The peridotitic material in ophiolite complexes (obducted ocean
floor) is mainly HARZBURGITE (ol+opyx) or DUNITE (ol), often
cut by pyroxenite (opyx) veins and sometimes having chromite
segregations (podiform chromites). There is a consensus that this
mantle is the refractory residue left after basalt extraction at midocean ridges. But are the pyroxenite veins the result of silica-rich
fluids coming off the subduction zone? (see in Fig. 5a all the free
quartz present in eclogite in subduction zones) - note that many
ophiolites are thought to be fragments of "back-arc" spreading
centres.
The MOHO as a chemical transition
Possible scenarios for the oceans and continents are shown in Fig. 6.
(3) In the Ivrea Zone and the Lanzo-Seisia Zone of the Italian Alps,
the peridotite slices are overlain by layered gabbros. They are mainly
LHERZOLITE (ol-opyx-cpyx -spinel, or -hornblende or phlogopite). Most carry veins composed of orthopyroxene,
orthopyroxene-spinel, diopside-orthopyroxene, phlogopite-diopside,
or hornblende. Different segments different veins, suggesting a
complex make-up of the sub-continental mantle.
Some of the rock types are similar to those brought up in volcanic
breccia pipes.
(4) Considerable progress has been made in understanding the
compositions of basaltic rocks in recent years, and interpreting them
in terms of melting models. So it is possible to "model back" to the
primary mantle from which the basalt was derived, and estimate its
composition. We now know that there are several distinct types of
mantle, that have been kept separated for many hundreds of millions
of years. They have distinct trace element and isotopic
characteristics. (You may come across references to them as DMM,
HIMU, EM1, EM2 and PREMA), but there is uncertainty as to
where they are located.
In oceanic regions the form of the MOHO is known from dredging at
fracture zones where peridotites (often serpentinised) have been
recovered by dredging and drilling along with cumulate gabbros and
pillow basalts. Actual sections of ocean floor are preserved in
ophiolite complexes.
Upper Mantle Mineralogy
The variation in upper mantle mineral assemblages with
temperature and pressure can be determined experimentally. But
there is uncertainty about composition. Because a lot of observed
mantle material is not primary, but has had a liquid (basalt) fraction
removed by partial melting, Ringwood coined the term ’PYROLITE’
for primitive fertile mantle - in effect dunite with basalt put back in!
The nature of the sub-continental MOHO is less certain, partly due to
lack of knowledge of the lower crust.
Composition of the Upper Mantle
PYROLITE = 3 parts DUNITE + 1 part BASALT
Our petrological knowledge of the upper mantle composition comes
from several sources:
(1) Nodules brought up in volcanic pipes;
(2) Large sections of mantle found in obducted ophiolites;
(3) Slices of mantle thrust up in mountain belts such as the Alps
(Ivrea-Verbano Zone) and the Caledonides (especially Norway); and
Mantle mineralogy varies mainly on account of the nature of the
aluminous phase, which is P-T dependent, i.e.
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Plate Tectonics: GL209
Prof. John Tarney
Olivine (MgFe)2SiO4 + Orthopyroxene (MgFe)2Si2O6 as major
phases, plus:
Plagioclase CaAl2Si2O8
Spinel
(MgFe)Al2O4
Clinopyroxene (NaCa)2(AlSi)2O6
Garnet Mg3Al2Si3O12
Hornblende
Phlogopite
Lecture 1: Geological Aspects
which are already known (isomorphs). For instance we can compare
silicates with germanates because germanates form a series of crystal
structures closely parallel to those of silicates, but the
transformations occur at lower pressures.
Thermodynamics requires that the high pressure
polymorph be denser, which limits possible structures. Once
structure is known the bond lengths between cations and anions
enable densities to be calculated.
The Radius Ratio (Rcation/Ranion) determines type of
crystal structure. At high pressures effective radii contract
differentially, thus altering the radius ratio. Thus a new high pressure
phase may appear when radius ratios exceed certain critical values.
Large ions (e.g. Oxygen 1.40Å) contract more under pressure than
small ions. Oxygen is more polarizable than smaller cations:
or
or
or
or
or
Phase diagram (Fig. 7), though rather complicated, shows that
plagioclase peridotite can exist only at very shallow depths where the
geothermal gradient is high; spinel- and pyroxene peridotites have a
larger stability field in the upper mantle; but garnet peridotite will
occur at deeper levels (hence common as nodules in kimberlite
pipes).
Element
O-Si4+
Mg2+
Polarizability
3.1 × 10-24cm3
0.04 × 10-24cm3
0.12 × 10-24cm3
Radius (Å)
1.4
0.26
0.72
Pressure thus increases covalent component of chemical bond.
Phase transitions in the deeper mantle
Refinement of seismic wave data has shown number of
discontinuities (Fig. 8):
Fig. 7. Summary of phase relations in pyrolite (after Green
& Ringwood) appropriate to upper mantle conditions. Wet
solidus for small amount of hornblende breakdown. Note
that only oceanic geotherm intersects this wet solidus.
Methods for investigating deep mantle mineralogy
Direct sampling of the deeper mantle is obviously impossible.
Observed seismic velocity-depth functions however constrain the
densities of likely mantle phases. Moreover, possible phase
transformations in the transition zone of the mantle are very difficult
to verify experimentally. For instance, in the 400-900 km depth
region pressures are in the range 130-340 kilobars and temperatures
1500-3000°C ... beyond the range of most experimental equipment
until recently. Now with diamond anvil apparatus, laser heating and
on-line X-ray determinations it is possible to reach into this range, at
least momentarily.
Indirect methods have also proved reasonably successful.
Fortunately high pressure phases tend to crystallise in structures
These zones are:
(1) LOW VELOCITY ZONE: from below lithosphere to about
200-250 km. Not always present. Asthenosphere has S-wave
attenuation . . small amount of liquid, perhaps ca 1% melting?
(2) MINOR DISCONTINUITY around 350 km.
(3) MAJOR DISCONTINUITY at 400 km.
(4) MAJOR DISCONTINUITY at 650 km.
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Plate Tectonics: GL209
Prof. John Tarney
(5) Between 900 and 2700 km no major discontinuities, but some
smaller ones. In general increase in seismic velocities and density
explained by self-compression of homogenous material.
Lecture 1: Geological Aspects
Fig. 9. Density changes with depth in the mantle, and
the changes in mineral structure that have been
proposed to explain them
The following explanations have been proposed to explain these
discontinuities (Fig. 9):
A point of interest is whether this sharp increase in density at 650700 km acts as a barrier to mantle convection. If the slab cannot
penetrate this boundary, does it pile up above 700km?
Are there two convecting zones in the mantle: one above, one below
the 700km discontinuity? Does this also coicide with a chemical
boundary? Is there any chemical interchange across the boundary
layer?
350 km. Pyroxene forms a complex solid solution with pre-existing
garnet in which one-quarter of silicon atoms are octahedrally coordinated, leading to 10% increase in density of pyroxene
component:
Mg3(MgSi)Si3O12 & Ca3(CaSi)Si3O12
Fate of the subducted slab: Ringwood 1991 Model
One of the problems of plate tectonics is the fate of the
subducting slab. This can be traced, from seismic evidence, to
descend to about 650 km; but the evidence is somewhat conflicting
regarding the extent to which it penetrates the dense 650 km
discontinuity. (See references by Jordan and Hilst). Because the
phase changes with depth are now known in some detail, both for
ultramafic mantle material and for subducted basaltic ocean crust, it
is possible to calculate their modal compositions with depth. For
instance, the modal composition of pyrolite with depth is shown in
Fig. 10:
400 km. Olivine transforms to beta-Mg2SiO4 which has SPINEL
structure and is 8% denser than olivine.
500-550 km. Calcium silicate CaSiO3 component of garnet
transforms to extremely dense PEROVSKITE structure. Also betaMg2SiO4 transforms to gamma-Mg2SiO4 with 2% increase in
density.
650 km. The spinel structure disproportionates to MgSiO3 with
PEROVSKITE structure and MgO with a ROCK SALT structure,
i.e. Mg2SiO4 > MgSiO3 + MgO. Additionally the MgSiO3.Al2O3
component transforms to an ILMENITE structure and any sodium
present will transform to a high pressure form of NaAlSiO4 having
CALCIUM FERRITE structure.
Lower Mantle. Below 700 km no more major transformations are
possible - the minerals are as close-packed as they can get. There is
thus then a slow progressive increase in density to the mantle-core
boundary.
Fig. 10 Modal composition of pyrolite with depth.
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Plate Tectonics: GL209
Prof. John Tarney
Lecture 1: Geological Aspects
Fig. 13. Density differences between basalt - pyrolite and
harzburgite - pyrolite as the subducted ocean crust sinks.
The basaltic slab becomes less dense than mantle pyrolite
in the depth range 650 - 750 km.
Fig. 11 shows the same calculations for basaltic ocean
crust.
The important point is that the subducted plate sinks
because the basaltic component of the slab (now eclogite) is ca. 0.2 0.1 g/cc more dense than the enclosing host pyrolite to depths of 650
km, and exerts a strong ’slab-pull’ force at subduction zones. The
harzburgite part of the plate may also be slightly more dense initially
because it is cold, but is inherently less dense once it has thermally
equilibrated with the surrounding mantle pyrolite. However, because
of the phase changes in pyrolite at the 670 km discontinuity, the
basaltic crust suddenly becomes 0.2 g/cc less dense than the pyrolite
in the depth range 650-750 km, whereas the harzburgite component
of the slab becomes very slightly more dense. These effects are very
clearly shown in Fig. 13. Ringwood (1991) argues that these changes
then have the effect of trapping subducted basaltic ocean crust at the
670 km discontinuity, as shown in Fig. 14.
Note that the plate which is subducting is not uniform mantle
pyrolite but, because of melting at the ridge axis, it has segregated
into a basaltic ocean crust (ca 5 km thick), residual harzburgite (from
which the basalts were extracted) underlain by ordinary pyrolite.
Knowing the mineral proportions and the densities of the minerals in
each of the main rock types, undepleted pyrolite, depleted
harzburgite, and basaltic ocean crust, it is then possible to calculate
the density changes in each of these rock types with depth. The
thermally equilibrated densities for these three rock types with depth
are shown in Fig. 12.
Fig. 12. Densities (g/cc) of thermally equilibrated basaltic
ocean crust, subducted harzburgite lithosphere compared
with undepleted pyrolite mantle to depths of 800 km. Note
that the ocean crust is mostly more dense and the
harzburgite is less dense than pyrolite down to 650km
depth, but then their positions are reversed.
Fig. 14. The effect of density differences is that basaltic ocean
crust becomes trapped at the 670km discontinuity.
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Plate Tectonics: GL209
Prof. John Tarney
Lecture 1: Geological Aspects
Ringwood has suggested that the slab piles up at the base of
the upper mantle, as shown in Fig. 14. By the end of the
Archaean (2500 my ago) he envisages that the mantle structure
around the 650km discontinuity would be as shown in Fig. 15.
This layer is source for diamond-bearing kimberlite magmas
according to Ringwood et al. (1992).
Fig. 16. Models of mantle differentiation involving storage of
ocean crust at the 650 km discontinuity.
This has interesting consequences as a mechanism to generate mantle
plumes and ’hotspots’. Most plumes need to be generated at a
discontinuity, either the 650 km one or at the core-mantle boundary.
The mantle model (Fig. 16) shows that these plumes rise and
penetrate the lithosphere to become the source of hotspot ocean
islands. If these rising diapirs cannot penetrate the lithosphere, they
may just add to the base of the lithsophere, and melts may penetrate
it and metasomatise and chemically alter it. The ocean lithosphere
is young (almost all less than 200 my) whereas the sub-continental
lithosphere is older, cooler, thicker and more complex, as shown in
both Figs. 16 & 17.
Fig. 15. Likely mantle structure at the end of the Archaean
as a result of subducted mafic ocean crust piling up at the
650 km discontinuity (after Ringwood).
Because of the different thermal regimes, and the
influences of plumes, it is likely that there have been differences in
the make-up of the lithosphere during Earth history. Fig. 18 (below)
shows the likely structure of the modern mature Phanerozoic ocean
lithosphere (left), which is regarded as being less depleted with
increasing depth. Ocean plateaus (centre) have a very thick ocean
crust, with (implicitly) a much more depleted harzburgitic mantle
underlying it. In the Archaean (right) one suggestion is that the high
mantle temperatures would have led to very high degrees of melting,
to produce high-Mg komatiitic lavas, and leaving an extremely
depleted pure-olivine dunitic residue. This oceanic structure is much
more like that of modern oceanic plateaus, so was there plate
tectonics in the Archaean or plateau tectonics (see later PlateLect-F)?
Fig. 17. (above) Comparison between oceanic and continental
lithosphere.
Assuming constant spreading rates (present day) it can be
calculated that, throughout Earth history, the amount of ocean crust
which may have accumulated at the 650 km discontinuity would be
at least 100 km thick. However because the harzburgite is inherently
less dense and potentially more buoyant than the surrounding mantle,
then when it heats up it may begin to ascend as blobs or diapirs, as
shown in Fig. 16 (below).
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Plate Tectonics: GL209
Prof. John Tarney
Lecture 1: Geological Aspects
Mantle Convection
Models of mantle differentiation
The diagrams below show some conceptual models of how
the mantle may be convecting, and possible relationships between
the upper and lower mantle (after Allègre et al.). There is still a very
intense debate on whether the lower mantle is involved in
manconvection.
This widely accepted model implies that mantle (or at least
the upper mantle) is continually differentiating to form continental
crust by a two-stage process. The crust formed is permanent and is
not recycled back into the mantle.
(1) Primitive pyrolite mantle rises at mid-ocean ridges, melts to form
basaltic ocean crust overlying refractory harzburgite plate.
(2) Plate sinks back into the mantle at subduction zones. Hydrated
altered ocean crust dehydrates and causes melting of the basaltic
ocean crust and of the overlying mantle wedge to yield andesitic
magmas.
(3) Andesitic magmas fractionate en route to the surface to produce
more siliceous magmas. Hence sialic crust accretes laterally at
continental margins, is of low density and is indestructible.
Fig. 20. Box models for crust-mantle evolution. On the left
continental growth occurs through igneous contributions
from both the upper and lower mantle. On the right the
continental crust has mainly been extracted from the upper
mantle, which is therefore "depleted" relative to lower
mantle.
Fig. 19. Simple "box model" of mantle evolution, showing
how melting at spreading ridges produces ocean crust, which
is then altered by hydrothermal activity and then subducted.
Part of this subducted crust is then melted to form continental
crust, and the residues then subducted to become part of the
reservoir of the depleted (DMM) mantle. Small degree melts
migrate upwards to enrich the sub-continental mantle and
provide the source for alkali basalts. Sediment subduction
may modify the sub-continental lithosphere. (after Tarney et
al. 1980)
Slab Penetration into Lower Mantle?
A consequence of course is that if the continental crust has been
extracted from the convecting mantle, the convecting mantle must
have become progressively depleted in lithophile elements. This is
now known as the ’DM’ mantle reservoir. This is the reservoir that
supplies depleted mid-ocean ridge basalt ("MORB"). The real story
is a little more complicated, as may be deduced form Fig. 19.
Sediments may be subducted and contaminate the lithosphere under
continental margins as well as the material stored at the 650 km
discontinuity. Small degree melts permeate upwards and vein both
the sub-continental and sub-oceanic lithosphere, but because the
former is older, we generally observe more complex effects under
the continents.
Fig. 22. Cartoon showing how subducting slabs may
either lay themselves out along the 650 km discontinuity
(a), or penetrate the discontinuity to enter the lower
mantle as in (b). The latter gives active back-arc
spreading (see later lecture).
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Plate Tectonics: GL209
Prof. John Tarney
Lecture 1: Geological Aspects
Fig. 23 (after Kerr et al. 1995) shows 2-layer convection of
the mantle, as subducting plates lodge at the 670km discontinuity, or
get convected back into the upper mantle; but with periodic
interchange between the two as cold plates avalanche down into the
lower mantle, and deep mantle plumes are displaced and rise to form
major oceanic plateaus and continental large igneous provinces
("LIPs"). There may be composition differences between upper
mantle and lower mantle as a result of such processes through Earth
history.
The most recent analysis of the fate of the oceanic crust as it
subducts into the mantle beneath the West Pacific island arcs (van
der Hilst & Seno, 1993), suggests that whereas that subducting
beneath the Izo-Bonin arc and Shikoku Basin, south of Japan may be
deflected and "laid-out" along the 650 km discontinuity in the
transition zone (Fig. 22(a)), that further south beneath the Mariana
Arc may penetrate into the lower mantle (Fig. 22(b)):
UPDATES (1994/6)
Larson and Kincaid (1996) then go on to argue that the breakup of
major continents, as occurred with the Gondwana supercontinent in
the Mesozoic (ca. 130Ma), leads to more rapid subduction of old
cold ocean crust. These cold slabs then avalanche down and
penetrate the 670km thermal boundary layer into the lower mantle.
One effect is to raise the 670km TBL; another is to displace material
from the deep lower mantle (D") which appears as major mantle
plumes during the mid-Cretaceous magnetic superchron (120 Ma 80 Ma). See later notes on mantle plumes.
Ringwood’s Megalith Model:
The essence of the Ringwood megalith model is that while ocean
lithosphere subduction is initially thermally driven because the
downgoing slab is cold [the basaltic (eclogite) layer is 5% more
dense than the surrounding mantle, and compensates for the depleted
harzburgite which is 2% lighter – see Fig. 13], the compositional
buoyancy difference between the two becomes significant at 700km
depth, resulting in the mechanical separation of basaltic from
harzburgite/dunite components (see Fig. 15). However, recent
modelling by Gaherty & Hager (1994), using a range of viscosity
contrasts for eclogite vs harzburgite, shows that the two are unlikely
to separate. The slab buckles and folds as it reaches the 700 km
discontinuity, but there is no obvious separation of eclogitic and
harzburgitic components. The compositional buoyancy differences
are subordinate to the overall thermal buoyancy.
Nature of the Lower Mantle:
While it is generally known that the convecting Upper Mantle (above
the 670km discontinuity) is chemically depleted in lithophile
elements because of the progressive growth and extraction of the
continental crust from it, it has been commonly thought that the
Lower Mantle is largely undepleted. However, Kerr et al. (1995)
have proposed that the Lower Mantle is also depeleted, in part
because of the return of subducting slab material right through the
670km discontinuity into the lower mantle: see also van der Hilst &
Seno (1993). This implies that there is much more interchange
between Upper and Lower mantle than was first thought. The Lower
Mantle feeds into the upper mantle in the form of large hot plumes
(see later lecture). Figure 23 below shows how cool subducted
material may go right into lower mantle, or get stuck termporarily at
the 670km discontinuity and then 'drip' into the lower mantle:
9
Plate Tectonics: GL209
Prof. John Tarney
Lecture 1: Geological Aspects
RINGWOOD, A.E. 1985. Mantle dynamics and basalt petrogenesis.
Tectonophysics 112, 17-34.
RINGWOOD, A.E. 1986. Dynamics of subducted lithosphere and
implications for basalt petrogenesis. Terra Cognita 6, 67-77.
RINGWOOD, A.E. 1991. Phase transitions and their bearing on the
constitution and dynamics of the mantle. Geochimica et
Cosmochimica Acta 55, 2083-2110.
RINGWOOD, A.E. & IRIFUNE, T. 1988. Nature of the 650-km
seismic discontinuity: implications for mantle dynamics and
differentiation. Nature 331, 131- 136.
RINGWOOD, A.E., KESSON, S.E., HIBBERSON, W. & WARE,
N. 1992. Origin of kimberlite and related magmas. Earth and
Planetary Science Letters 113, 521-538.
SADOWIAK, P., WEVER, T. and MEISSNER, R. 1991. Deep
seismic reflectivity patterns in specific tectonic units of Western
and Central Europe. Geophysics Journal International 105, 4554.
TARNEY, J., WOOD, D.A., SAUNDERS, A.D., CANN, J.R. &
VARET, J. 1980. Nature of mantle heterogeneity in the North
Atlantic: evidence from deep sea drilling. Phil. Trans. Roy. Soc.
London A297, 179-202.
van der HILST, R. & SENO, T. 1993. Effects of relative plate
motion on the deep structure and penetration depth of slabs
below the Izu-Bonin and Mariana island arcs. Earth and
Planetary Science Letters 120, 395-407.
WEVER, T., TRAPPE, H. and MEISSNER, R. 1987. Possible
relations between crustal reflectivity, crustal age, heat flow and
viscosity of the continents. Annales Geophysicae 5B, 255-266.
WARNER, M.R. 1990. Basalts, water or shear zones in the lower
continental crust? Tectonophysics 173, 163-173.
WYLLIE, P.J. 1971. The Dynamic Earth. Wiley, London
REFERENCES: Mantle Mineralogy
(These references are probably more than you require at this stage,
but as some aspects are followed up in more detail later in your
courses, they may be useful to you.)
ALLÈGRE, C.J. 1982. Chemical geodynamics. Tectonophysics 81,
109-132.
ALLÈGRE, C.J. & TURCOTTE, D.L. 1986. Implications of a twocomponent marble-cake mantle. Nature 323, 123-127.
BLUNDELL, D.J. 1990. Seismic images of continental lithosphere.
Journal of the Geological Society, London 147, 895-913.
GAHERTY, J.B. & HAGER, B.H. 1994. Compositional vs. thermal
buoyancy and the evolution of subducted lithosphere.
Geophysics Research Letters 21, 141-144.
IRIFUNE, T. & RINGWOOD, A.E. 1987. Phase transformations in a
harzburgite composition to 26 GPa: implications for dynamical
behaviour of the subducting slab. Earth and Planetary Science
Letters 86, 365-376.
IRIFUNE, T. & RINGWOOD, A.E. 1993. Phase transformations in
subducted ocean crust and buoyancy relationships at depths of
600-800 km in the mantle. Earth and Planetary Science Letters
117, 101-110.
JORDAN, T.H. 1975. The continental tectosphere. Review of
Geophysics and Space Physics 13, 1-12.
JORDAN, T.H. 1978. Composition and development of the
continental tectosphere. Nature 274, 544-548.
JORDAN, T.H. 1981. Continents as a chemical boundary layer.
Philosophical Transactions of the Royal Society, Lond. A301,
359-373.
KEARY, P. & VINE, F.J. 1991. Global Tectonics. Blackwell
Scientific Publ., 302pp.
KERR, A.C., SAUNDERS, A.D., TARNEY, J., BERRY, N.H &
HARDS, V.L. 1995. Depleted mantle-plume geochemical
signatures: no paradox for plume theories. Geology 23, 843-846.
KUSZNIR, N. and MATTHEWS, D.H. 1988. Deep seismic
reflections and the deformational mechanics of the continental
lithosphere. Journal of Petrology Special Lithosphere Issue, pp.
63-87.
LARSON, R.L. & KINCAID, C. 1996. Onset of mid-Cretaceous
volcanism by elevation of the 670 km thermal boundary layer.
Geology 24, 551-554.
MEISSNER, R. 1989. Rupture, creep, lamellae and crocodiles:
happenings in the continental crust. Terra Nova 1, 17-28.
MEISSNER, R. and KUSZNIR, N. 1987. Crustal viscosity and the
reflectivity of the lower crust. Annales Geophysicae 5B, 365373.
MEISSNER, R., MATTHEWS, D.H. and WEVER, T. 1986. The
Moho in and around Britain. Annales Geophysicae 4B, 659-666.
MENZIES, M.A. 1990. Archaean, Proterozoic, and Phanerozoic
lithospheres. In: M.A. Menzies (Editor) Continental Mantle,
Clarendon Press, Oxford, pp.67-86.
RINGWOOD, A.E. 1974. The petrological evolution of island arc
systems. Journal of the Geological Society, London 130, 183204.
RINGWOOD, A.E. 1975. Composition and Petrology of the Earth’s
Mantle. McGraw-Hill, New York.
RINGWOOD. A.E. 1982. Phase transformations and differentiation
in subducted lithosphere: implications for mantle dynamics,
basalt petrogenesis, and crustal evolution. Journal of Geology 90,
611-643.
APPENDIX:
Germanates as high pressure models of silicates
Trying to elucidate the petrological nature of the deep mantle is not
easy because it is difficult to re-create such high-pressure - hightemperature conditions in the laboratory. At least for any length of
time: Laser heating and momentary shock treatment can do it for a
short time, but as silicate reactions usually take a long time to reach
equilibrium condition, this leads to huge uncertainties in P-T
parameters. However, in the early years, mineral chemistry
principles were used to predict high pressure behaviour.
Use of germanates to model high pressure silicates was first
suggested by Goldschmidt in 1931. Si and Ge are tetravalent and in
same group in the Periodic Table.
RADII:
Si4+ 0.26A
Ge4+ 0.40A
Silicates and germanates usually isostructural and there is a wide
range of germanate structures. So, if it is possible to synthesize a
germanate structure at moderate pressures it is likely that an
equivalent silicate structure will exist at higher pressures. If a
germanate displays a phase transformation at a given pressure, the
corresponding silicate often displays the same transformation but at a
much higher pressure. This is because the critical radius ratio
10
Plate Tectonics: GL209
Prof. John Tarney
RGe/ROxygen for transformation to a new phase is attained at much
lower pressures with Ge. Some germanates (e.g. GeO2) can
crystallise at zero pressure while the equivalent silicate needs 100 kb
pressure.
Fig. 12. Densities (g/cc) of thermally equilibrated basaltic ocean
crust, subducted harzburgite lithosphere compared with undepleted
pyrolite mantle to depths of 800 km. Note that the ocean crust is
mostly more dense and the harzburgite is less dense than pyrolite
down to 650km depth, but then their positions are reversed.
Fig. 13. Density differences between basalt - pyrolite and harzburgite
- pyrolite as the subducted ocean crust sinks. The basaltic slab
becomes less dense than mantle pyrolite in the depth range 650 - 750
km.
Many transformations in germanates involve change from 4- to 6fold co-ordination with oxygen. Compare:
NaAlSi2O8
Albite
NaAlGe2O8
2CoSiO3
Pyroxene
2CoGeO3
>
>
>
>
>
>
Lecture 1: Geological Aspects
NaAlSi2O6 + SiO2 28kb (silicate)
Jadeiite + Rutile str
NaAlGe2O6 + GeO2 5kb (germanate)
2Co2SiO4 + SiO2 100kb (silicate)
Spinel + Rutile str
2Co2GeO4 + GeO2 10kb (germanate)
Fig. 14. The effect of density differences is that basaltic ocean crust
becomes trapped at the 670km discontinuity.
Fig. 15. Likely mantle structure at the end of the Archaean as a result
of subducted mafic ocean crust piling up at the 650 km discontinuity
(after Ringwood).
These lines of reasoning allowed predictions to be made as to what
types of structure might exist at depth in the Earth.
Figure Captions
Fig. 16 (below). Models of mantle differentiation involving storage
of ocean crust at the 650 km discontinuity.
Fig. 1. The Earth in proportion (and the crust thickness has still been
exaggerated!). Upper convecting mantle is quite a thin layer. Do
superplumes strat at core-mantle boundary?
Fig. 17. (above) Comparison between oceanic and continental
lithosphere.
FIg. 19. Simple "box model" of mantle evolution, showing how
melting at spreading ridges produces ocean crust, which is then
altered by hydrothermal activity and then subducted. Part of this
subducted crust is then melted to form continental crust, and the
residues then subducted to become part of the reservoir of the
depleted (DMM) mantle. Small degree melts migrate upwards to
enrich the sub-continental mantle and provide the source for alkali
basalts. Sediment subduction may modify the sub-continental
lithosphere. (after Tarney et al. 1980)
Fig. 2. Diagram (based on field and geophysical studies) to show
how deep crust can be thrust up to high crustal levels. Kapuskasing
structure, Ontario, Canada)
Fig. 3. Estimates as to the extent to which we have sections thru two
segments of Archaan crust and one segment of recent Alpine crust.
Fig. 4. Various early models interpreting the Moho as a phase change
(no real change in composition).
Fig. 20 (above). Box models for crust-mantle evolution. On the left
continental growth occurs through igneous contributions from both
the upper and lower mantle. On the right the continental crust has
mainly been extracted from the upper mantle, which is therefore
"depleted" relative to lower mantle.
Fig. 5a. Experimental studies of Green & Ringwood (1967) on quartz
tholeiite basalt showed that transformation to eclogite occurred over
a considerable depth interval. Note that eclogite has a lot more
quartz than the equivalent basalt. But mantle eclogites have no
quartz. Where does the silica go?
Fig. 21 (below) implies that the size of the convective cell depends
on the size of the ocean: the large Pacific ocean with whole-mantle
convection, the smaller Indian ocean with only upper mantle
convection. But evidence is slight!
Fig. 5b. Transformation of alkali basalt into eclogite also occurs over
considerable depth range. Not sharp enough to explain Moho. Note
no quartz.
Fig. 6. Moho as compositional change: various models.
FIg. 22 (above). Cartoon showing how subducting slabs may either
lay themselves out along the 650 km discontinuity (a), or penetrate
the discontinuity to enter the lower mantle as in (b). The latter gives
active back-arc spreading (see later lecture).
Fig. 7. Summary of phase relations in pyrolite (after Green &
Ringwood) appropriate to upper mantle conditions. Wet solidus for
small amount of hornblende breakdown. Note that only oceanic
geotherm intersects this wet solidus.
Fig. 23 (after Kerr et al. 1995) shows 2-layer convection of the
mantle, as subducting plates lodge at the 670km discontinuity, or get
convected back into the upper mantle; but with periodic interchange
between the two as cold plates avalanche down into the lower
mantle, and deep mantle plumes are displaced and rise to form major
oceanic plateaus and continental large igneous provinces ("LIPs").
There may be composition differences between upper mantle and
lower mantle as a result of such processes through Earth history.
Fig. 9. Density changes with depth in the mantle, and the changes in
mineral structure that have been proposed to explain them.
Fig. 11. Modal composition of subducted basaltic ocean crust with
depth. Compare with pyrolite in Fig. 10.
11