plate tectonics: geological aspects prof. j. tarney tectonics: geological aspects prof. j. tarney...

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

Week 9 Active Margins & Accretion (L) Practical: Plate model exercises (1)

Week 10 Plumes & plateaus (L) Practical: Plate model exercises (2)Collision; plate tectonics &Earth evolution

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: Thisworks as a result of hotmantle asthenosphereascending beneath mid-ocean ridges to form theplates. These plates thensubside into the mantleagain at subductionzones, pulled by theexcess density of themafic ocean crust, whichtransform to eclogite.

Plate Tectonics: GL209 Prof. John Tarney Lecture 1: Geological Aspects

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Fig. 3. Estimates as to the extent to which we have sectionsthru two segments of Archaean crust and one segment of‘recent’ Alpine crust. There is uncertainty as to whetheractually we can directly sample the lowest crust. The Alpinesequence is in a series of thrust slices, but some rocks inregion of the Ivrea zone have been down to greater depthsand have re-bounded again to the surface.Perhaps the more important question is what causes high-Procks to exhume?

MANTLE PETROLOGY IN RELATION TO PLATETECTONICS

Knowledge of mantle petrology and the constitution of the deepermantle is important in trying to understand several aspects of platetectonics. For instance, is there whole-mantle convection or two-layer convection? What are mantle plumes? What are superplumes?Does the subducting slab penetrate into the lower mantle? Whathappens to the slab at depth? Is the sub-continental mantle differentfrom the oceanic mantle? First, some basic facts.

Principal Internal Subdivisions of the Earth

Region Depth Mass Mass(km) (1025g) Fraction

Crust 0-Moho 2.4 0.004Upper Mantle Moho-400 62 0.10Transition Zone 400-1000 1000 0.17Lower Mantle 1000-2900 245 0.41Outer Core 2900-5154 177 0.30 Inner Core 5154-6371 12 0.02

Note that the crust makes up quite a small proportion of thetotal Earth. The main problems that have occupied geologists overthe years are: What is the nature of the crust-mantle boundary (theMOHO). What is the nature of the low velocity zone? Is thelithosphere diferent in composition from the asthenosphere? Whathappens in the transition zone? What is the nature of the deepmantle?

The Continental Crust

Though we know quite a lot about the upper crust, there isstill quite a lot of uncertainty about the lower crust. Is there a realConrad discontinuity separating the lower from the upper crust? Isthe lower crust made up of dry granulite-facies rocks. Is it moremafic than the upper crust, perhaps as a result of intrusion of maficmagmas into the lower crust (called "underplating"). Or is it more

mafic as a result of extraction of silicic granitic magmas from thedeep crust?

There a number of regions where we think tectonic activityhas brought segments of the lower crust up for inspection. Notableexamples are Kapuskasing and Pikwitonei in Canada (Precambriancrust), the Lewisian of NW Scotland (also Precambrian), Calabria inS. Italy and the Ivrea Zone in the Alps (both Phanerozoic).

Fig. 2. Diagram (based on field and geophysical studies) to showhow deep crust can be thrust up to high crustal levels. Kapuskasingstructure, Ontario, Canada).The Lewisian of NW Scotland can be interpreted similarly, thehigh-grade lower crustal granulites being thrust over amphibolites>2.5 Gyr ago.


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Fig. 5. Experimental studies of Green & Ringwood (1967)on quartz tholeiite basalt showed that transformation toeclogite occurred over a considerable depth interval. Notethat eclogite has a lot more quartz than the equivalentbasalt. But mantle eclogites have no quartz. Where does thesilica go?

The Moho and the Lower Crust

In the early 1960s there was considerable discussionbetween petrologists and geophysicists as to the nature of theMOHO. The P-wave velocity of most regions of the uppermostmantle beneath both continents and oceans lies in the range 8.2 ±0.2km/sec. This in itself would restrict the composition of the mantlebelow 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 nature of the lower crust is less certain. ExhumedPrecambrian 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 involcanic breccia pipes tend to be more mafic (gabbroic) incomposition. So has the continental crust in part been underplatedsubsequently by mafic magma? Also, recent deep seismicinvestigations 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 varioustectonic 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) bandsof reflectors as in BIRPS WINCH 3, (c) hyperbola-like diffractionsas in BIRPS SWAT 6-9, (d) "crocodile" diverging reflectorsobserved in old collision zones, but not so far in the UK, (e) "rampand 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 crustalstructure. However it is felt that lower crustal viscosity (Meissner &Kusznir, 1987) is a more important factor controlling development ofreflectors than is composition, and the current view is that goodlower crust reflectors might characterise mature crust, but that thispattern could be destroyed by either compression/collision("crocodiles", etc.), igneous intrusions or significant extension.

Nonetheless it is commonly assumed that the lower crust isgabbroic in composition, either igneous (gabbro), or metamorphicamphibolite (wet) or granulite (dry).

The MOHO as a phase transition

Fig. 4 shows various suggestions (made at one time or other) for howthe 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 high-pressure form of basalt or gabbro. But are these models realistic? Theserpentine-peridotite model is now discounted.

The MOHO and the gabbro-eclogite transformation: Basalttransforms to eclogite at high pressures according to the equation:

Olivine + pyroxene + plagioclase > jadeitic pyroxene + garnet

For this to be capable of explaining the MOHO it must be arelatively sharp transition, of no more than a few km. Green andRingwood (1967) studied this experimentally to 30 kilobars (= ca100 km) using a quartz tholeiite and an alkali basalt (Fig 5).


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With both compositions the transformation was found to be gradual.The disappearance of the low density phase (plagioclase) and itsreplacement by the high density phases (garnet and jadeite) occurredover a pressure range of ca 10 kb (= ca 25 km). Thus the MOHOcannot be a phase transformation and must be a compositionaltransition.

The MOHO as a chemical transition

Possible scenarios for the oceans and continents are shown in Fig. 6.

In oceanic regions the form of the MOHO is known from dredging atfracture zones where peridotites (often serpentinised) have beenrecovered by dredging and drilling along with cumulate gabbros andpillow basalts. Actual sections of ocean floor are preserved inophiolite complexes.

The nature of the sub-continental MOHO is less certain, partly due tolack of knowledge of the lower crust.

Composition of the Upper Mantle

Our petrological knowledge of the upper mantle composition comesfrom 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

(4) Modelling "backwards" from erupted basalt compositions.

(1) Olivine rich nodules are quite common in erupted alkali basaltsworldwide, 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 garnetperidotite and eclogite xenoliths occur (the former dominant). Somenodules contain diopside-rich, phlogopite-rich or amphibole-richveins. It is thought that these nodules are representative of the sub-continental mantle. This material seems to be rather refractory (couldnot yield much basalt on melting), but at the same time can be quiteenriched in incompatible trace elements such as Sr, Ba, K, Rb andthe light rare- earths.(2) The peridotitic material in ophiolite complexes (obducted oceanfloor) is mainly HARZBURGITE (ol+opyx) or DUNITE (ol), oftencut by pyroxenite (opyx) veins and sometimes having chromitesegregations (podiform chromites). There is a consensus that thismantle is the refractory residue left after basalt extraction at mid-ocean ridges. But are the pyroxenite veins the result of silica-richfluids coming off the subduction zone? (see in Fig. 5a all the freequartz present in eclogite in subduction zones) - note that manyophiolites are thought to be fragments of "back-arc" spreadingcentres.

(3) In the Ivrea Zone and the Lanzo-Seisia Zone of the Italian Alps,the peridotite slices are overlain by layered gabbros. They are mainlyLHERZOLITE (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 acomplex make-up of the sub-continental mantle.Some of the rock types are similar to those brought up in volcanicbreccia pipes.

(4) Considerable progress has been made in understanding thecompositions of basaltic rocks in recent years, and interpreting themin terms of melting models. So it is possible to "model back" to theprimary mantle from which the basalt was derived, and estimate itscomposition. We now know that there are several distinct types ofmantle, that have been kept separated for many hundreds of millionsof years. They have distinct trace element and isotopiccharacteristics. (You may come across references to them as DMM,HIMU, EM1, EM2 and PREMA), but there is uncertainty as towhere they are located.

Upper Mantle MineralogyThe variation in upper mantle mineral assemblages with

temperature and pressure can be determined experimentally. Butthere is uncertainty about composition. Because a lot of observedmantle material is not primary, but has had a liquid (basalt) fractionremoved by partial melting, Ringwood coined the term ’PYROLITE’for primitive fertile mantle - in effect dunite with basalt put back in!

PYROLITE = 3 parts DUNITE + 1 part BASALT

Mantle mineralogy varies mainly on account of the nature of thealuminous phase, which is P-T dependent, i.e.


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Olivine (MgFe)2SiO4 + Orthopyroxene (MgFe)2Si2O6 as majorphases, plus:

Plagioclase CaAl2Si2O8 or Spinel (MgFe)Al2O4 or Clinopyroxene (NaCa)2(AlSi)2O6 or Garnet Mg3Al2Si3O12 or Hornblende or Phlogopite

Phase diagram (Fig. 7), though rather complicated, shows thatplagioclase peridotite can exist only at very shallow depths where thegeothermal gradient is high; spinel- and pyroxene peridotites have alarger stability field in the upper mantle; but garnet peridotite willoccur at deeper levels (hence common as nodules in kimberlitepipes).

Methods for investigating deep mantle mineralogy

Direct sampling of the deeper mantle is obviously impossible.Observed seismic velocity-depth functions however constrain thedensities of likely mantle phases. Moreover, possible phasetransformations in the transition zone of the mantle are very difficultto verify experimentally. For instance, in the 400-900 km depthregion pressures are in the range 130-340 kilobars and temperatures1500-3000°C ... beyond the range of most experimental equipmentuntil recently. Now with diamond anvil apparatus, laser heating andon-line X-ray determinations it is possible to reach into this range, atleast momentarily.

Indirect methods have also proved reasonably successful.Fortunately high pressure phases tend to crystallise in structures

which are already known (isomorphs). For instance we can comparesilicates with germanates because germanates form a series of crystalstructures closely parallel to those of silicates, but thetransformations occur at lower pressures.

Thermodynamics requires that the high pressurepolymorph be denser, which limits possible structures. Oncestructure is known the bond lengths between cations and anionsenable densities to be calculated.

The Radius Ratio (Rcation/Ranion) determines type ofcrystal structure. At high pressures effective radii contractdifferentially, thus altering the radius ratio. Thus a new high pressurephase may appear when radius ratios exceed certain critical values.Large ions (e.g. Oxygen 1.40Å) contract more under pressure thansmall ions. Oxygen is more polarizable than smaller cations:

Element Polarizability Radius (Å) O-- 3.1 × 10-24cm3 1.4 Si4+ 0.04 × 10-24cm3 0.26 Mg2+ 0.12 × 10-24cm3 0.72

Pressure thus increases covalent component of chemical bond.

Phase transitions in the deeper mantle

Refinement of seismic wave data has shown number ofdiscontinuities (Fig. 8):

These zones are:

(1) LOW VELOCITY ZONE: from below lithosphere to about200-250 km. Not always present. Asthenosphere has S-waveattenuation . . 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.

Fig. 7. Summary of phase relations in pyrolite (after Green& Ringwood) appropriate to upper mantle conditions. Wetsolidus for small amount of hornblende breakdown. Notethat only oceanic geotherm intersects this wet solidus.


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(5) Between 900 and 2700 km no major discontinuities, but somesmaller ones. In general increase in seismic velocities and densityexplained by self-compression of homogenous material.

The following explanations have been proposed to explain thesediscontinuities (Fig. 9):

350 km. Pyroxene forms a complex solid solution with pre-existinggarnet in which one-quarter of silicon atoms are octahedrally co-ordinated, leading to 10% increase in density of pyroxenecomponent: Mg3(MgSi)Si3O12 & Ca3(CaSi)Si3O12

400 km. Olivine transforms to beta-Mg2SiO4 which has SPINELstructure and is 8% denser than olivine.

500-550 km. Calcium silicate CaSiO3 component of garnettransforms to extremely dense PEROVSKITE structure. Also beta-Mg2SiO4 transforms to gamma-Mg2SiO4 with 2% increase indensity.

650 km. The spinel structure disproportionates to MgSiO3 withPEROVSKITE 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 sodiumpresent will transform to a high pressure form of NaAlSiO4 havingCALCIUM FERRITE structure.

Lower Mantle. Below 700 km no more major transformations arepossible - the minerals are as close-packed as they can get. There isthus then a slow progressive increase in density to the mantle-coreboundary.

A point of interest is whether this sharp increase in density at 650-700 km acts as a barrier to mantle convection. If the slab cannotpenetrate this boundary, does it pile up above 700km?Are there two convecting zones in the mantle: one above, one belowthe 700km discontinuity? Does this also coicide with a chemicalboundary? Is there any chemical interchange across the boundarylayer?

Fate of the subducted slab: Ringwood 1991 ModelOne of the problems of plate tectonics is the fate of the

subducting slab. This can be traced, from seismic evidence, todescend to about 650 km; but the evidence is somewhat conflictingregarding the extent to which it penetrates the dense 650 kmdiscontinuity. (See references by Jordan and Hilst). Because thephase changes with depth are now known in some detail, both forultramafic mantle material and for subducted basaltic ocean crust, itis possible to calculate their modal compositions with depth. Forinstance, the modal composition of pyrolite with depth is shown inFig. 10:

Fig. 9. Density changes with depth in the mantle, andthe changes in mineral structure that have beenproposed to explain them

Fig. 10 Modal composition of pyrolite with depth.


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Note that the plate which is subducting is not uniform mantlepyrolite but, because of melting at the ridge axis, it has segregatedinto a basaltic ocean crust (ca 5 km thick), residual harzburgite (fromwhich the basalts were extracted) underlain by ordinary pyrolite.Knowing the mineral proportions and the densities of the minerals ineach of the main rock types, undepleted pyrolite, depletedharzburgite, and basaltic ocean crust, it is then possible to calculatethe density changes in each of these rock types with depth. Thethermally equilibrated densities for these three rock types with depthare shown in Fig. 12.

The important point is that the subducted plate sinksbecause 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 650km, and exerts a strong ’slab-pull’ force at subduction zones. Theharzburgite part of the plate may also be slightly more dense initiallybecause it is cold, but is inherently less dense once it has thermallyequilibrated with the surrounding mantle pyrolite. However, becauseof the phase changes in pyrolite at the 670 km discontinuity, thebasaltic crust suddenly becomes 0.2 g/cc less dense than the pyrolitein the depth range 650-750 km, whereas the harzburgite componentof the slab becomes very slightly more dense. These effects are veryclearly shown in Fig. 13. Ringwood (1991) argues that these changesthen have the effect of trapping subducted basaltic ocean crust at the670 km discontinuity, as shown in Fig. 14.

Fig. 12. Densities (g/cc) of thermally equilibrated basalticocean crust, subducted harzburgite lithosphere comparedwith undepleted pyrolite mantle to depths of 800 km. Notethat the ocean crust is mostly more dense and theharzburgite is less dense than pyrolite down to 650kmdepth, but then their positions are reversed.

Fig. 13. Density differences between basalt - pyrolite andharzburgite - pyrolite as the subducted ocean crust sinks.The basaltic slab becomes less dense than mantle pyrolitein the depth range 650 - 750 km.

Fig. 14. The effect of density differences is that basaltic oceancrust becomes trapped at the 670km discontinuity.

Fig. 11 shows the same calculations for basaltic oceancrust.


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Ringwood has suggested that the slab piles up at the base ofthe upper mantle, as shown in Fig. 14. By the end of theArchaean (2500 my ago) he envisages that the mantle structurearound the 650km discontinuity would be as shown in Fig. 15.This layer is source for diamond-bearing kimberlite magmasaccording to Ringwood et al. (1992).

Assuming constant spreading rates (present day) it can becalculated that, throughout Earth history, the amount of ocean crustwhich may have accumulated at the 650 km discontinuity would beat least 100 km thick. However because the harzburgite is inherentlyless dense and potentially more buoyant than the surrounding mantle,then when it heats up it may begin to ascend as blobs or diapirs, asshown in Fig. 16 (below).

This has interesting consequences as a mechanism to generate mantleplumes and ’hotspots’. Most plumes need to be generated at adiscontinuity, either the 650 km one or at the core-mantle boundary.The mantle model (Fig. 16) shows that these plumes rise andpenetrate the lithosphere to become the source of hotspot oceanislands. If these rising diapirs cannot penetrate the lithosphere, theymay just add to the base of the lithsophere, and melts may penetrateit and metasomatise and chemically alter it. The ocean lithosphereis young (almost all less than 200 my) whereas the sub-continentallithosphere is older, cooler, thicker and more complex, as shown inboth Figs. 16 & 17.

Because of the different thermal regimes, and theinfluences of plumes, it is likely that there have been differences inthe make-up of the lithosphere during Earth history. Fig. 18 (below)shows the likely structure of the modern mature Phanerozoic oceanlithosphere (left), which is regarded as being less depleted withincreasing depth. Ocean plateaus (centre) have a very thick oceancrust, with (implicitly) a much more depleted harzburgitic mantleunderlying it. In the Archaean (right) one suggestion is that the highmantle temperatures would have led to very high degrees of melting,to produce high-Mg komatiitic lavas, and leaving an extremelydepleted pure-olivine dunitic residue. This oceanic structure is muchmore like that of modern oceanic plateaus, so was there platetectonics in the Archaean or plateau tectonics (see later PlateLect-F)?

Fig. 15. Likely mantle structure at the end of the Archaeanas a result of subducted mafic ocean crust piling up at the650 km discontinuity (after Ringwood).

Fig. 17. (above) Comparison between oceanic and continentallithosphere.

Fig. 16. Models of mantle differentiation involving storage ofocean crust at the 650 km discontinuity.


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Models of mantle differentiation

This widely accepted model implies that mantle (or at leastthe upper mantle) is continually differentiating to form continentalcrust by a two-stage process. The crust formed is permanent and isnot recycled back into the mantle.(1) Primitive pyrolite mantle rises at mid-ocean ridges, melts to formbasaltic ocean crust overlying refractory harzburgite plate.(2) Plate sinks back into the mantle at subduction zones. Hydratedaltered ocean crust dehydrates and causes melting of the basalticocean crust and of the overlying mantle wedge to yield andesiticmagmas.(3) Andesitic magmas fractionate en route to the surface to producemore siliceous magmas. Hence sialic crust accretes laterally atcontinental margins, is of low density and is indestructible.

A consequence of course is that if the continental crust has beenextracted from the convecting mantle, the convecting mantle musthave become progressively depleted in lithophile elements. This isnow known as the ’DM’ mantle reservoir. This is the reservoir thatsupplies depleted mid-ocean ridge basalt ("MORB"). The real storyis a little more complicated, as may be deduced form Fig. 19.Sediments may be subducted and contaminate the lithosphere undercontinental margins as well as the material stored at the 650 kmdiscontinuity. Small degree melts permeate upwards and vein boththe sub-continental and sub-oceanic lithosphere, but because theformer is older, we generally observe more complex effects underthe continents.

Mantle Convection

The diagrams below show some conceptual models of howthe mantle may be convecting, and possible relationships betweenthe upper and lower mantle (after Allègre et al.). There is still a veryintense debate on whether the lower mantle is involved inmanconvection.

Slab Penetration into Lower Mantle?Fig. 19. Simple "box model" of mantle evolution, showinghow melting at spreading ridges produces ocean crust, whichis then altered by hydrothermal activity and then subducted.Part of this subducted crust is then melted to form continentalcrust, and the residues then subducted to become part of thereservoir of the depleted (DMM) mantle. Small degree meltsmigrate upwards to enrich the sub-continental mantle andprovide the source for alkali basalts. Sediment subductionmay modify the sub-continental lithosphere. (after Tarney etal. 1980)

Fig. 20. Box models for crust-mantle evolution. On the leftcontinental growth occurs through igneous contributionsfrom both the upper and lower mantle. On the right thecontinental crust has mainly been extracted from the uppermantle, which is therefore "depleted" relative to lowermantle.

Fig. 22. Cartoon showing how subducting slabs mayeither lay themselves out along the 650 km discontinuity(a), or penetrate the discontinuity to enter the lowermantle as in (b). The latter gives active back-arcspreading (see later lecture).


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The most recent analysis of the fate of the oceanic crust as itsubducts into the mantle beneath the West Pacific island arcs (vander Hilst & Seno, 1993), suggests that whereas that subductingbeneath the Izo-Bonin arc and Shikoku Basin, south of Japan may bedeflected and "laid-out" along the 650 km discontinuity in thetransition zone (Fig. 22(a)), that further south beneath the MarianaArc may penetrate into the lower mantle (Fig. 22(b)):

UPDATES (1994/6)

Ringwood’s Megalith Model:

The essence of the Ringwood megalith model is that while oceanlithosphere subduction is initially thermally driven because thedowngoing slab is cold [the basaltic (eclogite) layer is 5% moredense than the surrounding mantle, and compensates for the depletedharzburgite which is 2% lighter – see Fig. 13], the compositionalbuoyancy difference between the two becomes significant at 700kmdepth, resulting in the mechanical separation of basaltic fromharzburgite/dunite components (see Fig. 15). However, recentmodelling by Gaherty & Hager (1994), using a range of viscositycontrasts for eclogite vs harzburgite, shows that the two are unlikelyto separate. The slab buckles and folds as it reaches the 700 kmdiscontinuity, but there is no obvious separation of eclogitic andharzburgitic components. The compositional buoyancy differencesare subordinate to the overall thermal buoyancy.

Nature of the Lower Mantle:

While it is generally known that the convecting Upper Mantle (abovethe 670km discontinuity) is chemically depleted in lithophileelements because of the progressive growth and extraction of thecontinental crust from it, it has been commonly thought that theLower Mantle is largely undepleted. However, Kerr et al. (1995)have proposed that the Lower Mantle is also depeleted, in partbecause of the return of subducting slab material right through the670km discontinuity into the lower mantle: see also van der Hilst &Seno (1993). This implies that there is much more interchangebetween Upper and Lower mantle than was first thought. The LowerMantle feeds into the upper mantle in the form of large hot plumes(see later lecture). Figure 23 below shows how cool subductedmaterial may go right into lower mantle, or get stuck termporarily atthe 670km discontinuity and then 'drip' into the lower mantle:

Fig. 23 (after Kerr et al. 1995) shows 2-layer convection ofthe mantle, as subducting plates lodge at the 670km discontinuity, orget convected back into the upper mantle; but with periodicinterchange between the two as cold plates avalanche down into thelower mantle, and deep mantle plumes are displaced and rise to formmajor oceanic plateaus and continental large igneous provinces("LIPs"). There may be composition differences between uppermantle and lower mantle as a result of such processes through Earthhistory.

Larson and Kincaid (1996) then go on to argue that the breakup ofmajor continents, as occurred with the Gondwana supercontinent inthe Mesozoic (ca. 130Ma), leads to more rapid subduction of oldcold ocean crust. These cold slabs then avalanche down andpenetrate the 670km thermal boundary layer into the lower mantle.One effect is to raise the 670km TBL; another is to displace materialfrom the deep lower mantle (D") which appears as major mantleplumes during the mid-Cretaceous magnetic superchron (120 Ma -80 Ma). See later notes on mantle plumes.


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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 yourcourses, 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 two-component 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. thermalbuoyancy and the evolution of subducted lithosphere.Geophysics Research Letters 21, 141-144.

IRIFUNE, T. & RINGWOOD, A.E. 1987. Phase transformations in aharzburgite composition to 26 GPa: implications for dynamicalbehaviour of the subducting slab. Earth and Planetary ScienceLetters 86, 365-376.

IRIFUNE, T. & RINGWOOD, A.E. 1993. Phase transformations insubducted ocean crust and buoyancy relationships at depths of600-800 km in the mantle. Earth and Planetary Science Letters117, 101-110.

JORDAN, T.H. 1975. The continental tectosphere. Review ofGeophysics and Space Physics 13, 1-12.

JORDAN, T.H. 1978. Composition and development of thecontinental 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. BlackwellScientific Publ., 302pp.

KERR, A.C., SAUNDERS, A.D., TARNEY, J., BERRY, N.H &HARDS, V.L. 1995. Depleted mantle-plume geochemicalsignatures: no paradox for plume theories. Geology 23, 843-846.

KUSZNIR, N. and MATTHEWS, D.H. 1988. Deep seismicreflections and the deformational mechanics of the continentallithosphere. Journal of Petrology Special Lithosphere Issue, pp.63-87.

LARSON, R.L. & KINCAID, C. 1996. Onset of mid-Cretaceousvolcanism 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 thereflectivity of the lower crust. Annales Geophysicae 5B, 365-373.

MEISSNER, R., MATTHEWS, D.H. and WEVER, T. 1986. TheMoho in and around Britain. Annales Geophysicae 4B, 659-666.

MENZIES, M.A. 1990. Archaean, Proterozoic, and Phanerozoiclithospheres. In: M.A. Menzies (Editor) Continental Mantle,Clarendon Press, Oxford, pp.67-86.

RINGWOOD, A.E. 1974. The petrological evolution of island arcsystems. Journal of the Geological Society, London 130, 183-204.

RINGWOOD, A.E. 1975. Composition and Petrology of the Earth’sMantle. McGraw-Hill, New York.

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APPENDIX:Germanates as high pressure models of silicates

Trying to elucidate the petrological nature of the deep mantle is noteasy because it is difficult to re-create such high-pressure - high-temperature conditions in the laboratory. At least for any length oftime: Laser heating and momentary shock treatment can do it for ashort time, but as silicate reactions usually take a long time to reachequilibrium condition, this leads to huge uncertainties in P-Tparameters. However, in the early years, mineral chemistryprinciples were used to predict high pressure behaviour.

Use of germanates to model high pressure silicates was firstsuggested by Goldschmidt in 1931. Si and Ge are tetravalent and insame group in the Periodic Table.

RADII: Si4+ 0.26A Ge4+ 0.40A

Silicates and germanates usually isostructural and there is a widerange of germanate structures. So, if it is possible to synthesize agermanate structure at moderate pressures it is likely that anequivalent silicate structure will exist at higher pressures. If agermanate displays a phase transformation at a given pressure, thecorresponding silicate often displays the same transformation but at amuch higher pressure. This is because the critical radius ratio


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RGe/ROxygen for transformation to a new phase is attained at muchlower pressures with Ge. Some germanates (e.g. GeO2) cancrystallise at zero pressure while the equivalent silicate needs 100 kbpressure.

Many transformations in germanates involve change from 4- to 6-fold co-ordination with oxygen. Compare:

NaAlSi2O8 > NaAlSi2O6 + SiO2 28kb (silicate) Albite > Jadeiite + Rutile str NaAlGe2O8 > NaAlGe2O6 + GeO2 5kb (germanate) 2CoSiO3 > 2Co2SiO4 + SiO2 100kb (silicate) Pyroxene > Spinel + Rutile str 2CoGeO3 > 2Co2GeO4 + GeO2 10kb (germanate)

These lines of reasoning allowed predictions to be made as to whattypes of structure might exist at depth in the Earth.Figure Captions

Fig. 1. The Earth in proportion (and the crust thickness has still beenexaggerated!). Upper convecting mantle is quite a thin layer. Dosuperplumes strat at core-mantle boundary?

Fig. 2. Diagram (based on field and geophysical studies) to showhow deep crust can be thrust up to high crustal levels. Kapuskasingstructure, Ontario, Canada)

Fig. 3. Estimates as to the extent to which we have sections thru twosegments 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. 5a. Experimental studies of Green & Ringwood (1967) on quartztholeiite basalt showed that transformation to eclogite occurred overa considerable depth interval. Note that eclogite has a lot morequartz than the equivalent basalt. But mantle eclogites have noquartz. Where does the silica go?

Fig. 5b. Transformation of alkali basalt into eclogite also occurs overconsiderable depth range. Not sharp enough to explain Moho. Noteno quartz.

Fig. 6. Moho as compositional change: various models.

Fig. 7. Summary of phase relations in pyrolite (after Green &Ringwood) appropriate to upper mantle conditions. Wet solidus forsmall amount of hornblende breakdown. Note that only oceanicgeotherm intersects this wet solidus.

Fig. 9. Density changes with depth in the mantle, and the changes inmineral structure that have been proposed to explain them.

Fig. 11. Modal composition of subducted basaltic ocean crust withdepth. Compare with pyrolite in Fig. 10.

Fig. 12. Densities (g/cc) of thermally equilibrated basaltic oceancrust, subducted harzburgite lithosphere compared with undepletedpyrolite mantle to depths of 800 km. Note that the ocean crust ismostly more dense and the harzburgite is less dense than pyrolitedown 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 slabbecomes less dense than mantle pyrolite in the depth range 650 - 750km.

Fig. 14. The effect of density differences is that basaltic ocean crustbecomes trapped at the 670km discontinuity.

Fig. 15. Likely mantle structure at the end of the Archaean as a resultof subducted mafic ocean crust piling up at the 650 km discontinuity(after Ringwood).

Fig. 16 (below). Models of mantle differentiation involving storageof ocean crust at the 650 km discontinuity.

Fig. 17. (above) Comparison between oceanic and continentallithosphere.

FIg. 19. Simple "box model" of mantle evolution, showing howmelting at spreading ridges produces ocean crust, which is thenaltered by hydrothermal activity and then subducted. Part of thissubducted crust is then melted to form continental crust, and theresidues then subducted to become part of the reservoir of thedepleted (DMM) mantle. Small degree melts migrate upwards toenrich the sub-continental mantle and provide the source for alkalibasalts. Sediment subduction may modify the sub-continentallithosphere. (after Tarney et al. 1980)

Fig. 20 (above). Box models for crust-mantle evolution. On the leftcontinental growth occurs through igneous contributions from boththe upper and lower mantle. On the right the continental crust hasmainly been extracted from the upper mantle, which is therefore"depleted" relative to lower mantle.

Fig. 21 (below) implies that the size of the convective cell dependson the size of the ocean: the large Pacific ocean with whole-mantleconvection, the smaller Indian ocean with only upper mantleconvection. But evidence is slight!

FIg. 22 (above). Cartoon showing how subducting slabs may eitherlay themselves out along the 650 km discontinuity (a), or penetratethe discontinuity to enter the lower mantle as in (b). The latter givesactive back-arc spreading (see later lecture).

Fig. 23 (after Kerr et al. 1995) shows 2-layer convection of themantle, as subducting plates lodge at the 670km discontinuity, or getconvected back into the upper mantle; but with periodic interchangebetween the two as cold plates avalanche down into the lowermantle, and deep mantle plumes are displaced and rise to form majoroceanic plateaus and continental large igneous provinces ("LIPs").There may be composition differences between upper mantle andlower mantle as a result of such processes through Earth history.

plate tectonics: geological aspects prof. j. tarney tectonics: geological aspects prof. j. tarney...

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