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PLATE TECTONICS: Lecture 3
THE WILSON CYCLE: RIFTING AND THE DEVELOPMENT OF OCEAN
BASINS
As the concept of sea floor spreading gained acceptance in the late 60's, the
consequences for geology gradually began to dawn. One of the first to recognise
how plate tectonics could be applied to the geological record was J. Tuzo Wilson.
If continents rift apart to form ocean basins, other oceans must close. This may be
repeated throughout Earth history. Example: the IAPETUS ocean between
England & Scotland in the Lower Palaeozoic, closed in the Caledonian; later
opening of the Atlantic, almost in the same place. The cycle is known as the
Wilson Cycle:
(1) Rifting of continents by mantle diapirism
(2) Continental drift, seafloor spreading & formation of ocean basins
(3) Progressive closure of ocean basins by subduction of ocean lithosphere
(4) Continental collision and final closure of ocean basin
The two diagrams below (Figs 1 & 2) illustrate some simple (if old) concepts of
continental rifting (e.g. the Gondwana continent) at the start of the Wilson Cycle.
Uprising plume causes doming of crust with magma chamber developing
underneath. As extension continues, an ocean basin forms, and thick sedimentary
sequences develop at continental margins as rivers dump sediments in deep
water. However in reality may be a bit more complex . . .
CONTINENTAL RIFTING: rrr and RRR triple junctions
Four main stages can be recognised in the tectonic development of a typical rifted
passive margin:
(1) The RIFT VALLEY stage involves early graben formation prior to
continental splitting. This stage may be associated with domal uplift caused
by uprise of hot upper mantle material - but this uplift is not ubiquitous and
may be connected with underlying mantle hotspots. Example: African Rift
Valley.
(2) The YOUTHFUL stage, lasting about 50 my after the onsett of seafloor
spreading, while the thermal effects are still dominant. This stage is
characterised by rapid regional subsidence of the outer shelf and slope, but
some graben formation may persist. Example: Red Sea.
(3) The MATURE stage during which more subdued regional subsidence
may continue. Example: most of the present Atlantic continental margins.
(4) The FRACTURE stage when subduction starts and terminates the
history of the continental margin.
Fig. 3. The continent of
Africa is thought to have
been split by a series of
rift valleys in various
states of development.
Those in East Africa are
still in thick crust. Those
in West Africa are
associated with thick oilbearing sediments. In the
Red Sea area the rifting
has gone so far as to
form a narrow ocean. In
the south-east
Madagascar has been
completely separated
from Africa by rifting.
There are many examples of Stage 1. East African Rift Valley is the classic
example. But also the Midland Valley of Scotland, the Rhine Graben, the Oslo
Graben. These rifts have never got beyond stage 1. Commonly the volcanism
associated with these rifts is highly alkaline and undersaturated in silica.
What initiates rifting? There has been considerable discussion on this over the
years. Some have ascribed rifting to up-doming of the crust over a hot-spot;
certainly parts of the E African rift system are very elevated, compared with other
sectors, suggesting that the doming reflects an underlying hot low-density mantle
plume. In other cases, geophysical models suggest the asthenospheric mantle is
rising to high levels beneath the rift. However it is also apparent that rifting can
take place without extensive uplift; in such cases it may be the convective
processes in the underlying asthenosphere which are causing the extension. To
rift a continent apart it needs the rifts associated with various possible thermal
domes to link together. Morgan (1981, 1983) has suggested that as continents drift
slowly over hotspots the hotspots weaken the plate - like a blowtorch impinging on
the base - and these weakened zones become the sites of continental rifting.
Burke & Whiteman (1973), following the doming hypothesis, suggested that in
these domal regions, three rifts would develop, forming an 'rrr' triple junction.
Although it is possible that all three rifts might develop into an ocean ('RRR'), it is
more likely that two of these rifts would develop into an ocean ('RRr'), leaving the
third rift as a 'failed arm'. They demonstrated / speculated that on many continents
it was possible to recognise these RRr junctions. The 'failed arm' rift would
eventually subside as the thermal anomaly decayed and become the site of a
major depositional basin, or a major river channel and delta. The Benue Trough in
Nigeria is regarded as an example of such a failed arm following the opening of
the S. Atlantic. When oceans eventually close it is possible to recognise these
failed arms as depositional basins oriented perpendicular to the collision mountain
belt (most basins tend to be aligned parallel to mountain belts). These are termed
'aulacogens'.
Fig. 4. A. Doming by a mantle plume associated with volcanicity. B.
Rifting (rrr junction) is initiated. C. Further development results in two
of the rifts developing into an ocean, the third is a failed arm
(aulacogen). D. Less likely is that all three arms develop into oceans.
E. A common situation is that the failed arm develops into a major
river system feeding the continental margin. F. Expansion of oceans
on a finite earth is not possible: there must be plate subduction,
somewhere, sometime. G. Closure of oceans results in island arc
development above the subduction zone. H. Continued closure
results in collision with major fold and thrust belts. But often the
failed arm (aulocogen) is still preserved.
Development of Continental Rifts
Early ideas on the development of rifts are conceptualised in the diagram shown in
Fig. 5. This is based on the African rift system, where there is significant rift
magmatism. There is notable extension, shown by the widening of the diagram
block by at least 50 km. At the same time there is uplift or ascent of the more
ductile mantle, especially the asthenosphere. The crust, and particularly the upper
crust, is assumed to act in a brittle fashion.
Fig. 5a.
Progressive
formation of a
rift valley
through
extension of
the lithosphere
and continental
crust (by about
50 km). Note
that uprise and
decompression
of the
underlying
asthenosphere
results in
magma
formation. The
crust responds
by brittle
fracture. Early
rift sediments
are
downfaulted
into the
developing rift
(graben).
Erosion takes
place on the
sides of the rift
valley.
The first stage assumes that graben-like faults begin to form in the brittle crust.
The second stage shows simultaneous necking of the lithosphere with uprise of an
asthenosphere diapir. The decompression associated with the latter causes
melting of the mantle to give alkaline basaltic magmas. Pre-existing sediments are
downfaulted into the graben.
The third stage is accompanied by significant extension and by more uprise of the
asthenosphere. The latter causes doming of the crust (which is evident along the
E. African rift system, but is variably devloped. New sediments are deposited
within the graben as a result of erosion of the uplifting sides of the graben. So
there are both pre-rift and syn-rift sediments within the developing rift valley, but
sediments on the flanks are progressively erodied away. Note the complex
normal-faulting within the rift valley itself.
The fourth stage (Fig. 5b – below) shows the actually rifting-apart of the continent,
so the asthenosphere rises towards the surface, causing decompression and
extensice melting. New basaltic oceanic crust is formed.
Finally, sea-floor spreading takes over as the ocean basin widens. The rift
sedimentary sequence is buried beneath younger marine sediments.
Note: on this diagram the sediments at the continental margin are shown as not
very thick. This is because the model is based on the East African Rift System,
which does not have a great deal of subsidence associated with rifting. However,
other rifted continental margin sequences are very different, with thick sedimentary
sequences.
Continental Shelf Sediments
The real situation at passive continental margins is shown in Fig. 6 (below). This is
typical of a number of crustal cross-sections across the continental shelf of the
eastern Atlantic seaboard of North America, projected down to 30 km -- based
largely on gravity and magnetic evidence, plus some seismic profiles -- and some
extrapolation from land geology based on deep drill holes.
The critical point is the huge thicknesses of Mesozoic and Tertiary sediments, here
shown as almost 15 km, but in other cross-sections this can be even thicker. Note
that at the bottom of this pile are volcanics and volcanogenic sediments, and
evaporites, which most likely are shallow water. Also, massive carbonate reef
structures, which must also be shallow water, but also must indicate progressive
subsidence .. .. slow enough that shallow water sedimentation can keep pace with
it.
In many sections of the continental shelf off this eastern seaboard of the USA
there is a major coast-parallel magnetic structure, possibly a major intrusion. But
its age is unknown.
Fig. 6. Profile of deep structure of continental shelf off Atlantic coast of
eastern North America -- ?typical of passive continental margins. (Based on
gravity, magnetics and seismic data) Critical points regarding this profile are
(a) the large thickness of post-rift sediments of Mesozoic-Tertiary age, up to
15 km, and (b) that most of these sediments are shallow-water type. Note:
volcanics and evaporites and reef (or carbonate banks)
Rift Terminology
Continental Rift: elongate tectonic depression with which
the entire lithosphere has been modified in extension
Rift System: Tectonically interconnected series of rifts
Modern Rift: A rift that is teconically or magmatically active
Paleorift: A dead or dormant rift
Failed Arm: Branch of a triple junction not developed into an
ocean basin
Aulacogen: Paleorift in ancient platform that has been
reactivated by compressional deformation
Active Rifting: Rifting in response to thermal upwelling of the
asthenosphere
Passive Rifting: Rifting in response to remote stress field
Rifts and Mineralisation
Rifting structures are often good sites for mineralisation. This arises for three
reasons:
(1) They can be the sites of thick clastic sedimentation. These sediments hold vast
amounts of inter-granular salt water (brines). The brines may be in contact with
reducing sediments, such as carbonaceous shales, also a ready supply of
sulphur/sulphate. As the sediments compact, these brines are expelled and can
move laterally for large distances until they move up the rift faults. Having been
buried deep the brines get hot, and can be very corrosive. So en route they can
dissolve considerable amounts of metals. However, when they rise up the rift
faults and cool, these metals will be precipitated out. This can be enhanced
because oxidising meteoric water (groundwater) may also penetrate down these
faults, so metals wil be precipitated out when the two meet.
(2) Rift structures are also thermally anomalous hot zones. This is because they
are frequently underlain by igneous intrusions -- granite (or perhaps in some cases
gabbro) plutons. This magmatic heat drives the hydrothermal systems.
Importantly, these hydrothermal systems can last for many millions of years, so
the hot fluids in these hydrothermal systems can leach away at the rocks within
the rift system and precipitate the leached metals nearer the surface. Because the
rift structures remain topographically low structures for many tens of millions of
years, these metals concentrations can be preserved, without being eroded, for
long periods.
(3) The rift zones may be the sites of diverse rocks, particularly basaltic lavas,
which can release their metals on hydrothermal alteration. However, because the
rift faults can extend very deep (well into the upper mantle in some cases), there
may also be a component of deep fluids and metals in the hydrothermal system.
References
The references below will lead you to some of the discussion on rifting and the
Wilson Cycle:
BAKER, B.H., MOHR, P. & WILLIAMS, L.A.J. 1972. Geology of the eastern rift system of
Africa. Geological Society of America Special Paper 136, 1-67.
BOSWORTH, W. 1985. Geometry of propagating continental rifts. Nature 316, 625-627.
BOSWORTH, W. 1987. Off-axis volcanism in the Gregory rift, East Africa: implications for
models of continental rifting. Geology 15, 397-400.
BOTT, M.H.P 1995. Mechanisms ofrifting: Geodynamic modeling of continental rift
systems. In: K.H. Olsen (ed.) Continental rifts: evolution, structure, tectonics.
Developments in Geotectonics, 25, 27-43. Elsevier, Amsterdam
BRAILE, L.W., KELLER, G.R., WENDLANDT, R.F., MORGAN, P. & KHAN, M.A. 1995.
The East African Rift system. In: K.H. Olsen (ed.) Continental rifts: evolution, structure,
tectonics. Developments in Geotectonics, 25, Elsevier, Amsterdam
BURKE, K. & DEWEY, J.F. 1973. Plume generated triple junctions: key indicators in
applying plate tectonics to old rock. Journal of Geology 81, 406-433.
BURKE, K. & WHITEMAN, A.J. 1973. Uplift, rifting and break-up of Africa. In TARLING,
D.H. & RUNCORN, S.K. (eds) Implications of continental drift to the earth sciences.
Academic Press, London. 735-755.
DEWEY, J.F. & BURKE, K. 1974. Hotspots and continental break-up: implications for
collisional orogeny. Geology 2, 57-60.
DUNCAN, C.C. & TURCOTTE, D.L. 1994. On the breakup and coalescence of continents.
Geology 22, 103-106.
GURNIS, M. 1988. Large-scale mantle convection and the aggregation and dispersal of
continents. Nature 332, 695-699.
MORGAN, W.J. 1981. Hotspot tracks and the opening of the Atlantic and Indian Oceans.
In Emiliani, C. (ed) The Sea. Volume 7, 443-487. Wiley, New York.
MORGAN, W.J. 1983. Hotspot tracks and the early rifting of the Atlantic. Tectonophysics
94, 123-139.
MURPHY, J.B. & NANCE, R.D. 1992. Mountain belts and the supercontinent cycle.
Scientific American 266, 84-91.
OLSEN, K.H. & MORGAN, P. 1995. Introduction: Progress in understanding continental
rifts. In: K.H. Olsen (ed.) Continental rifts: evolution, structure, tectonics. Developments in
Geotectonics, 25, 3-26. Elsevier, Amsterdam
SPOHN, T. & SCHUBERT, G. 1982. Convective thinning of the lithosphere: a mechanism
for the initiation of continental rifting. Journal of Geophysical Research 87, 4669-4681.
WHITE, R.S. & McKENZIE, D.P. 1989. Magmatism at rift zones: the generation of volcanic
continental margins and flood basalts. Journal of Geophysical Research 94, 7685-7730.
WILSON, J.T. 1966. Did the Atlantic close and then re-open? Nature 211, 676-681.
PLATE TECTONICS: Lecture 4
CONTINENTAL MARGIN SUBSIDENCE
Fig. 1.
Simplified
relationship
at a
continenta
margin.
There can
be more
than 10 km
of shallowwater
sediments
at the
margin –
implying
slow
subsidence
How?
Passive continental margins are those associated with continental rifting and the
subsequent formation of ocean basins. They differ from active continental margins
which are associated with subduction. The continental shelves around the Atlantic
are typical passive margins: however there are some quite large differences in the
morphology of continental margins around the Atlantic: the reasons for which are
not fully understood (but see White et al. 1987; White & McKenzie 1989). There is
of course considerable interest in continental margins because of their potential as
major oil reservoirs. Hence much has been learned in the last few years.
One aspect of continental margins that has always been puzzling is the existence
of very thick – but relatively shallow-water – sedimentary sequences. There can be
as much as 15 km of Mesozoic and later sediments at some continental margins
bordering the N. Atlantic. How can these very thick sequences be reconciled with
gradual but progressive subsidence? Over the years various ideas (summarised in
Bott 1979, 1982) have been put forward:
Gravity Loading Hypothesis: This attributes subsidence to sediment load
(effectively replacing seawater with denser sediment), and is based on isostacy.
The amount of subsidence depends on relative densities of seawater (1.03),
sediment (2.15 – 2.55) and the underlying mantle (3.3). If the sea is filled with
sediment then in theory a sediment thickness of over twice the initial depth can
develop. In fact a total thickness of 14 km can form near the base of the initial
slope. If the lithosphere is treated as elastic the downwarping can extend about
150 km beyond the local sediment load. See Fig. 2 below.
Problem: This mechanism is not easily reconciled with substantial sequences of
shallow water sediment. It can only work if the sediments were deposited in deep
water initially. If initial water depth is less than 200 m, then sediment loading effect
is negligible.
Fig. 2.
Gravity
loading
hypothesis.
This
depends on
replacing
low density
water by
higher
density
sediment . .
.
Thermal hypothesis: This assumes that continental lithosphere near the embryo
margin is heated at time of continental rifting - this reduces density of lithosphere
permitting isostatic uplift. Subsequently, as the ocean widens, lithosphere cools
with time-scale of ca. 50 my and will subside to original position. However if
erosion occurred during uplift stage, real subsidence can occur, enhanced by
sediment loading.
Fig. 3.
Thermal
hypothesis of
Sleep. This
was the first
to recognise
that heating
up the mantle
(by a plume or
whatever)
could produce
substantial
crustal uplift
(and erosion),
followed by
thermal
subsidence.
Compare the
models by
McKenzie and
Wernicke later
..
Problem: Even with an extreme initial elevation of about 2 km, the amount of
subsidence, even with sediment loading, is not much more than 2 km. So not able
to explain thick sequences of over 5 km.
A modification of this thermal model assumes that the thermal event transforms
the base of the crust to denser granulite facies mineral assemblages, which may
also be invaded by basic magma. If this causes an increase in density of 0.2, it
can be calculated that the maximum depth of sediment permitted would only be
about 3 or 4 km. Thus insufficient to account for large sediment thicknesses.
Fig. 4.
Modification of
thermal
hypothesis
according to
Falvey (who
argues that
heating will
cause dense
granulite to
form).
Problem: such models predict a gap of many m.y. between onset of spreading and
the first marine sedimentation - which is not observed.
Crustal Thinning hypothesis: The continental crust and the lithosphere have an
upper brittle zone, 20 km thick, overlying a much weaker layer which deforms by
ductile flow. Thus crust may thin by progressive creep of middle and lower crustal
material towards the sub-oceanic upper mantle. It is argued that this may give rise
to jerky subsidence.
Fig. 5. After
the initial
rifting the
lower crust
deforms by
plastic flow.
Could the
lower
continental
crust flow
UNDER
oceanic crust
in the manner
shown?
An alternative hypothesis suggests that extreme thinning of the continental crust
can occur in a rift valley setting by plastic necking. Then, as the ocean basin forms
the passive continental margin will gradually subside.
Fig. 6.
Necking of
continental
crust?
Problem: a
typical rift
zone is about 50 km wide, thus transition zone at a continental margin would be
only 25 km wide. Observed continental margin sequences are however much
wider than this.
Normal-fault based mechanisms:
Early hypotheses assumed that graben formation required a wedge of crust about
60 km wide to sink isostatically between inward-dipping normal faults. As the
upper crust forms graben by wedge subsidence the ductile lower crust
compensates by plastic flow.
Fig. 7. Can
normal faulting
lead to
displacement of
ductile mantle by
flow?
Problem: Calculations suggested that a subsidence of ca 5 km could occur for an
initial 20 km wide trough. Not really enough. But getting nearer.
Faulting near continent-ocean contact:
This mechanism permits limited subsidence as normal faulting accompanies
downdrag of the cooling ocean lithosphere. The oceanic lithosphere subsides on a
time scale of about 50 my, so consistent with shallow water sediments. However
note that the zone of subsidence is too narrow.
Fig. 8.
Does
normal
faulting
occur at
continental
margins in
the
manner
shown in
B?
Conclusions
None of the above mechanisms, either alone or together, seem capable of
explaining the observed thick sedimentary sequences at continental margins that
are formed at the start of the Wilson Cycle. New ideas were clearly required.
These began to develop in the late 1970's as we began to understand more about
the thermal behaviour of the lithosphere and about the nature of listric faults.
Continental Lithosphere: The mantle forming the plates is more rigid than the
underlying asthenosphere. But this rigid mechanical boundary layer (MBL) varies
in thickness. It is thin at the ridges, but thickens to 60 or even 100+ km in old
oceanic lithosphere. It may be much thicker under the continents, but it is also
older - in fact the lithosphere under the continents is usually as old as the
continent above. So it may be cool, and may have experienced enrichment by
small degree mantle melts, the components of which may be stored in hydrous
minerals.
REFERENCES:
BOTT, M.H.P. 1979. Subsidence mechanisms at passive continental margins. American
Association of Petroleum Geologists Memoir 29, 8-19?
BOTT, M.H.P. 1982. The mechanism of continental splitting. Tectonophysics 81, 301-309.
KUSZNIR, N.J. & ZIEGLER, P.A. 1992. The mechanics of continental extension and
sedimentary basin formation: a simple-shear/pure-shear flexural cantilever model.
Tectonophysics 215, 117-131.
WHITE, R.S., SPENCE, G.D., FOWLER, S.R., McKENZIE, D.P., WESTBROOK, G.K. &
BOWEN, A.N. 1987. Magmatism at rifted continental margins. Nature 330, 439-444.
WHITE, R.S. & McKENZIE, D.P. 1989. Magmatism at rift zones: the generation of volcanic
continental margins and flood basalts. Journal of Geophysical Research 94, 7685-7729.
CONTINENTAL EXTENSION AND FORMATION OF SEDIMENTARY
BASINS
There is no doubt that when ocean basins open there is considerable subsidence
of the continental shelves over a wide area, and not just over the immediate rifted
margin. This is well exempified by the South Atlantic at ca. 127 Ma, just as the first
oceanic crust formed:
Fig. 9. A
very large
area in the
south
Atlantic
was
submerged
following
break-up
at 127m.y.
Why? Both
Chile and
Argentina
have
modest onland oil
reserves in
Patagonia
to the west
of the
Falkland
Plateau.
DSDP site
330 drilled
oily
sediments
in 1974.
Why did
Argentina
go to war
over the
Falklands?
Drilling at the eastern spur of the submerged Falkland Plateau revealed that it was
continental (granite gneisses) and that there was a dry caliche surface
(Mediterranean climate) just before opening of the Atlantic, but that there had been
at least 2 km subsidence since then. Initial sediments very oily, deposited under
anoxic conditions in a basin with restricted circulation. So the initial rift stage was
the one that favoured oil accumulation. Why? It is important to understand the
mechanism of development of these basins.
Modern Ideas
It became apparent from COCORP-type deep reflection seismic profiling that
many (if not the majority) of steeply dipping normal faults are actually curved
(concave-upward) and become shallow-dipping and sub-horizontal at depth.
These are now known as listric faults. As the lithosphere is stretched during
continental extension, the ductile deeper crust thins by pure shear, while the upper
crust is broken up and pulled apart by listric faults which 'bottom out' in the ductile
layer. At the surface of course these have the appearance of graben. This is the
essence of McKenzie-type and other recent models of basin formation. As the
sub-continental (i.e. mantle) lithosphere is thinned by stretching it is of course
partly replaced by hotter asthenosphere. This will gradually cool on a time scale of
the order of 50 - 100 m.y., and as it cools it becomes denser and the shallow basin
above gradually subsides and is progressively filled with shallow-water sediment.
The amount of subsidence will depend on the initial amount of stretching. This can
usually be estimated and is known as the stretching factor, or "beta factor". The
parameter b is defined quite simpy as b/a where a was the initial width and b is the
stretched width. A b factor of 1.2 will give ca. 3 km subsidence. With complete
rifting (to form ocean crust and an ocean basin) then b approaches infinity.
Note that during the development of sedimentary basins, subsidence occurs in two
stages:
(1) as a result of tectonic stretching – on a short time scale, ca. 10 my, and
(2) as a result of thermal subsidence – long time scale, ca. 50 – 100+ my.
Considerable information is now available on North Sea basins as a result of
drilling operations and syntheses of the large amount of seismic data (see, e.g.
Badley et al. 1988; Gibbs 1984; Sclater & Christie 1980) so their subsidence
history is well known. The northern Viking Graben suffered two episodes of rifting
– in the Permo-Triassic and in the Middle Jurassic – during which the basin was
progressively widened. Stretching factors in the Permo-Triassic were quite small
(b = 1.1 – 1.3), whereas in the Late Jurassic were much larger in the northern N.
Sea (b = >1.6). Each rifting episode was followed by more substantial thermal
subsidence. In the central part of the Viking Graben almost 10 km of sediment has
accumulated since the onset of the first rifting episode. As the second rifting phase
ended 140 my ago at least 90% of the subsidence resulting from thermal
relaxation must have occurred by now. Note that whereas normal faults during the
rifting phase tend to be listric, those accompanying thermal subsidence are planar.
An important secondary factor in such models is that the sediments initially
deposited in such basins will be 'cooked' slightly as a consequence of the
increased heat from the underlying asthenosphere – vital in maturation and
migration of petroleum. But sedimentary basins are not only important as oil
reservoirs: the expulsion of heated fluids from such basins can leach metals too,
thus if suitable host rocks exist valuable mineral deposits can be formed. A
number of important mineral deposits are attributed to this mechanism.
Further development of lithosphere stretching models have been proposed by
Wernicke, by Lister et al., Coward and others (see references below).
The important difference is in the recognition of low-angle detachments
(superficially like thrusts, but with movement sense as in normal fault), first
proposed for the Basin & Range province in the western USA. These may bottom
out in the lower crust or the upper mantle. The main effect is to introduce
asymmetry compared with the pure shear uniform-stretching McKenzie-type
model, so that basins associated with the thermal subsidence phase may be offset
from the thin-skinned basins associated with the initial rifting. Magmatic effects
(melting resulting from the uprising asthenosphere) may be offset from the main
sedimentary basins. Because of the asymmetry, the continental margins on the
two sides of an opening ocean may have very different profiles. Many other
complications may ensue. Consult the references below if you want the full story!
At least 3 types of continental margin have now been recognised:
(1) volcanic,
(2) non-volcanic and
(3) rift-transform.
(1) Volcanic margins tend to be narrow and have a thick igneous crust
between continental and normal ocean crust. A thick zone (3 – 5 km) of
seaward-dipping volcanic reflectors is typical. Suggestions of convective
circulation in uprising asthenosphere to explain volcanism, or that the
underlying asthenosphere was hotter than usual. Examples: Voring
Plateau, western Rockall Bank, East Greenland. See White et al. (1987 &
1988).
White & McKenzie (1989) have developed these models further to
quantitatively relate the volume of volcanics produced at continental
margins to the temperature of the underlying mantle. If the temperature is
100°C above normal the volume of magma will be doubled. Also they have
developed a relationship between the degree of stretching and the
temperature of the mantle to predict whether the rifted margin will rise
above sealevel or subside below it. When rifting occurs above hotspot
plumes there is usually an accompanying large volume of magma.
(2) Lithospheric deformation on non-volcanic margins is dominated by block
faulting and many listric faults. Stretching over a broad zone (100–300 km).
May be sediment starved (Red Sea, Galicia Bank, Goban Spur– Irish Sea)
or heavily sedimented (e.g. eastern USA margin).
(3) Rift-transform margins evolve in environments where there was a
significant component of strike-slip shear as well as extensional strain
deformation during opening (e.g. region between W. Africa and Brazil;
Falklands Plateau; also Gulf of California).
These different types of margin may have very different petroleum potential. Need
to know more about them to aid in locating future supplies. Note that the important
petroleum reservoirs in the North Sea are in 'failed-rifts' – where the North Atlantic
tried (unsuccessfully) to open quite a long time before it eventually succeeded!
There is a rapidly growing literature on models for continental rifting and basin
formation: try to read some of those below, and especially note the diagrams. In
any case they may prove useful to you next year.
Another problem of concern is why do we get basaltic magmatism associated with
some basins and not with others. Latin and White (1990) have tried to argue that
magmatism is more likely with uniform pure shear stretching (McKenzie model)
than the asymmetric simple stretching model of Wernicke. This is because
asthenosphere uprise is more focussed in pure shear model:
Fig. 13.
Comparison of
thermal
conseqences
of McKenzie’s
pure shear
model and
Wernicke’s
pure shear
model of
extensional
sedimentary
basins. It is
argued with
the simple
shear model it
is very difficult
to produce
sufficient
decompression
to allow
magma
formation.
This then has very different thermal consequences:
Fig. 14. With pure shear the temperature of
the uprising asthenosphere an exceed the
Fig. 15. With simple shear the temperature
of the uprising asthenosphere never
solidus of the mantle and allow melting.
reaches the solidus - so no melting occurs.
Basin Inversion
Basins that have formed by rifting and thermal subsidence don't always remain
basins. and may suffer later uplift and erosion. This is known as basin inversion.
This happened to many of the Permo-Triassic basins in Western Europe (see
Ziegler 1982) and is particularly evident in the NW part of the British Isles and
adjacent continental margin. Could this be due to tectonic compression before all
the thermal subsidence took place, with the excess sediment being removed by
erosion? It is apparent that most of NW Britain was blanketed by Mesozoic
sediment that has been removed (viz. Chalk in calderas on Arran) since the early
Tertiary, and deposited in basins to the east. Some offshore basins with b factors
near 2.0 have a short fall in the expected sediment thickness of ca. 4 km. So
something has caused epeirogenic uplift in the early Tertiary over most of NW
Britain.
Unfortunately, there is no evidence of enough tectonic compression (Roberts
1989) to account for this uplift by crustal thickening. So what else? Brodie & White
(1994) have suggested instead that it may result from magmatic underplating by
basalt. They calculate that 5 km of basalt (density 2.8) underplated into the lower
crust above the Moho would initially cause 600m uplift. Additionally, with the
"amplification" effect of erosion this may increase to ca. 2.5 km. Of course in this
general region we know that the Iceland plume was initiated ca. 60 Ma ago (early
Tertiary), and one 'rrr' arm extended down through Western Scotland to Lundy. A
lot of basalt lavas were erupted. But was much more magma underplated? We
know from their geochemistry that many of these basalt magmas have suffered
crustal contamination. Are they just a small representative of much more that was
ponded in the lower crust? See later lecture on plumes.
The interesting point is that many sedimentation features – basin development,
basin inversion, epeirogenic uplift enhancing erosion – may all have their origin in
mantle thermal processes. Hence it is important to understand the mantle!
REFERENCES: Sedimentary Basins & Continental Margins
BADLEY, M.E., PRICE, J.D., RAMBECH DAHL, C. & AGDESTEIN, T. 1988. The structural
evolution of the northern Viking Graben and its bearing upon extensional modes of basin
formation. Journal of the Geological Society, London 145, 455-472.
BARR, D. 1987. Lithospheric stretching, detached normal faulting and footwall uplift. In
COWARD, M.P., DEWEY, J.F. & HANCOCK, P.L. (eds) Continental Extensional
Tectonics. Geological Society of London, Special Publication 28, 75-94.
BARTON, P. & WOOD, R. 1984. Tectonic evolution of the North Sea Basin: crustal
stretching and subsidence. Geophysical Journal of the Royal Astronomical Society 79 9871022.
BRODIE, J. & WHITE, N. 1994. Sedimentary basin inversion caused by igneous
underplating: Northwest European continental shelf. Geology 22, 147-150.
BUCK, W.R. 1991. Mode of continental lithospheric extension. Journal of Geophysical
Research 96, 20161-20178.
BUCK, W.R., MARTINEZ, F. STECKLER, M.S. & COCHRAN, J.R. 1988. Thermal
consequences of lithospheric extension: Pure and simple. Tectonics 7, 213-234.
COCHRAN, J.R. 1983. Effects of finite rifting times on the development of sedimentary
basins. Earth and Planetary Science Letters 66, 289-302.
COOPER, M.A. & WILLIAMS, G.D. 1989. Inversion Tectonics. Geological Society of
London, Special Publication 44, 000pp.
COWARD, M.P. 1986. Heterogeneous stretching, simple shear and basin development.
Earth and Planetary Science Letters 80, 325-336.
GIBBS, A.D. 1984. Structural evolution of extensional basin margins. Journal of the
Geological Society, London 141, 609-620.
HELLINGER, S.J. & SCLATER, J.G. 1983. Some comments on two-layer extensional
models for the evolution of sedimentary basins. Journal of Geophysical Research 88,
8251-8269.
HOUSEMAN, G. & ENGLAND, P. 1986. A dynamical model of lithosphere stretching and
sedimentary basin formation. Journal of Geophysical Research 91, 719-729.
KENT, P., BOTT, M.H.P., MCKENZIE, D.P. & WILLIAMS, C.A. (eds) 1982. Evolution of
sedimentary basins. Philosophical Transactions of the Royal Society, London A305, .
JARVIS, G.T. 1984. An extensional model of graben subsidence - the first stage of basin
evolution. Sedimentary Geology 40, 13-31.
KLEMPERER, S. 1988. Crustal thinning and the nature of extension in the northern North
Sea from deep seismic reflection profiling. Tectonics 7, 803-821.
LATIN, D. & WHITE, N. 1990. Generating melt during lithospheric extension: Pure shear
vs. simple shear. Geology 18, 327-331.
LE PICHON, X., ANGELIER, J. & SIBUET, J.C. 1982. Plate boundaries and extensional
tectonics. Tectonophysics 81, 239-256.
LEEDER, M.R. 1983. Lithospheric stretching and North Sea Jurassic clastic source areas.
Nature 304, 510-514.
LISTER, G.S., ETHERIDGE, M.A. & SYMONDS, P.A. 1986. Detachment faulting and the
evolution of passive continental margins. Geology 14, 246-250.
LISTER, G.S., ETHERIDGE, M.A. & SYMONDS, P.A. 1989. Detachment models for the
formation of passive continental margins. Tectonics 10, 1038-1064.
McKENZIE, D.P. 1978. Some remarks on the development of sedimentary basins. Earth
and Planetary Science Letters 40, 25-32.
MUTTER, J.C., BUCK, W.R. & ZEHNDER, G.M. 1988. Convective partial melting I: a
model for the formation of thick basaltic sequences during the initiation of spreading.
Journal of Geophysical Research 93, 1031-1048.
RESTON, T.J. 1990. Mantle shear zones and the evolution of the northern North Sea
Basin. Geology 18, 272-275.
ROBERTS, D.G. 1989. Basin inversion in and around the British Isles. Geological Society
of London, Special Publication 44, 131-150.
ROWLEY, D.B. & SAHAGIAN, D. 1986. Depth-dependent stretching: a different approach.
Geology 14, 32-35.
SAWYER, D.S., SWIFT, B.A., SCLATER, J.G. & TOKSOZ, M.N. 1982. Extensional model
for the subsidence of the northern US Atlantic continental margin. Geology 10, 134-140.
SCLATER, J.G. & CHRISTIE, P.A.F. 1980. Continental stretching and explanation of the
post mid-Cretaceous subsidence of the central North Sea Basin. Journal of Geophysical
Research 85, 3711-3739.
VOORHOEVE, H. & HOUSEMAN, G. 1988. The thermal evolution of lithosphere extending
on a low-angle detachment zone. Basin Research 1, 1-9.
WERNICKE, B. 1981. Low-angle normal faults in the Basin and Range province: nappe
tectonics in an extending orogen. Nature 291, 645-648.
WERNICKE, B. 1985. Uniform-sense normal simple shear of the continental lithosphere.
Canadian Journal of Earth Sciences 22, 108-125.
WERNICKE, B. & BURCHFIEL, B.C. 1982. Modes of extensional tectonics. Journal of
Structural Geology 4, 105-115.
WHITE, N. 1989. Nature of lithospheric extension in he North Sea. Geology 17, 111-114.
WHITE, N. & McKENZIE, D. 1988. Formation of "steer's head" geometry of sedimentary
basins by differential stretching of the crust and mantle. Geology 16, 250-253.
WHITE, R.S. 1987. When continents rift. Nature 327, 191.
WHITE, R.S., SPENCE, G.D., FOWLER, S.R., McKENZIE, D.P., WESTBROOK, G.K. &
BOWEN, A.N. 1987. Magmatism at rifted continental margins. Nature 330, 439-444.
WHITE, R.S. 1988. A hot-spot model for early Tertiary volcanism in the N. Atlantic. In
MORTON, A.C. & PARSON, L.M. (eds) Early Tertiary Volcanism and the Opening of the
NE Atlantic. Geological Society of London, Special Publication 39, 3-13.
WHITE, R.S. & McKENZIE, D.P. 1989. Magmatism at rift zones: the generation of volcanic
continental margins and flood basalts. Journal of Geophysical Research 94, 7685-7729.
WOOD, R. & BARTON, P. 1983. Crustal thinning and subsidence in the North Sea. Nature
304, 561.
ZIEGLER, P.A. 1982. Geological atlas of Western and Central Europe. Shell International,
The Hague, 130pp. [JT has slides of relevant maps]
PLATE TECTONICS: Lecture 6
THERMAL ASPECTS OF SUBDUCTION ZONES
For the last 2 decades, geologists, geophysicists and geochemists have argued
about the physical and chemical conditions which allow melting to occur in
subduction zones. Whereas it is easy to explain magmatism at ocean ridges
where hot mantle is rising, it is not at first easy to explain why magmas appear in
abundance when a cold slab is pushed into the mantle at subduction zones. It
used to be thought that friction between the overriding and under-riding plates was
responsible, but calculations have showed that friction is most unlikely: there is
probably too much hydrous fluid and soft oozy subducted sediment that act as a
lubricant. It is important to try to understand the thermal structure of subduction
zones.
In a classic review paper, Ringwood (1974) suggested that the most primitive
island arc lavas (IAT), which are basaltic, could be related to dehydration of the
hydrated ocean crust (amphibolite) as it transforms to dense eclogite at depths of
ca. 100km. The hydrous fluids rise up into the peridotite mantle wedge, promoting
melting (magmas form at much lower temperatures in the presence of water).
These magmas then rise slowly up to the arc volcanoes above, and crystallise Mgrich olivines and pyroxenes as they ascend, so the magmas become more ironrich. The eruption of basalt (tholeiite) is non-violent. This is shown in cartoon form:
For the calc-alkaline, more silicic andesitic and dacitic magmas or more mature
arcs, Ringwood suggested a slightly different mechanism based on his
experimental work on eclogite. Hydrous melting of eclogite (if Si-poor garnet stays
in the residue produces silica-rich dacitic magmas. These then react with the
mantle wedge and rise up as diapirs and erupt as much more violent hydrous
magmas, of which Mt. St. Helens is a good example.
However, there are a number of problems with these simple models, and it is now
accepted that they only acount for a minor number of features of subduction zone
magmas. For instance, the primitive Mariana arc tholeiites are really a result of
fore-arc diapirism connected with the initiation of a new subduction zone, following
a change in plate motion, as outlined in the last lecture.
How can subduction zones give rise to the following range of magmas? Surely this
must imply a range of P-T conditions that involve both slab and wedge melting?
Boninites (High-Mg andesites): usually formed at early stage of island arcs
Island Arc Tholeiites (IAT): normally restricted to primitive island arcs
Calc-alkaline basalts & andesites: found in mature island arcs and continental
margins
Bajaites (Adakites): High-Mg andesites (but different from boninites) where ridge
subduction occurs or mafic rocks have been underplated.
Shoshonites: Often late-subduction or post-subduction: high-Ba, Sr magmas
Archaean TTG suite: Distinctive, and thought to be derived from subducted ocean
crust (they resemble adakites).
The critical points of issue are:
(a) under what conditions does the slab melt?
(b) what is the difference between subducting old ocean crust and young
ocean crust?
(c) do magmas originate instead in the mantle wedge?
(d) what is the mineralogy of the wedge. Are minerals like hornblende,
phlogopite & K-richterite stable in subduction zones?
(e) how can the difference between primitive island arcs (e.g. Marianas)
that tend to erupt basalts, and mature arcs (e.g. Andean margin) that tend
to erupt andesite, be explained?
Anderson et al. (1978; 1980) were the first to consider the thermal structure of
subduction zones seriously. Wyllie & co-workers, in a series of papers (e.g. Wyllie,
1988), used experimental petrology to try to constrain what will melt under hydrous
conditions, and what the magma compositions would be. He produced some
useful cartoon models, one of which is shown below:
The important points to note are that the ocean crust reaching a subduction zone
will be relatively "cold" and "wet". Just how cold it will be will depend on just how
many hundreds or thousands of km it has travelled from the spreading ridge. It will
be wet as a result of hydrothermal alteration near the ridge axis. As the plate
subducts the basaltic crust will undergo a progressive increase in metamorphic
grade – Greenshist > Amphibolite > Eclogite facies – which is also a series of
dehydration reactions to about 100km depth.
More recently, Peacock (1991) and Bickle & Davies (1991) have produced much
better thermal models. For instance, Peacock (1991) has produced useful thermal
numerical models. He explores the thermal effect of:
(a) Age of the oceanic crust being subducted. Clearly young warm ocean crust will
be more likely to melt if subducted than old cold lithosphere. The diagram below
shows the increase in temperature at 1 my intervals (dots) as ocean crust ranging
in age from 5 my [A] to 200 my [D] is subducted to 200 km. The surprising result is
that only when quite young crust is being subducted is there a possibility of melting
(i.e. temperatures reach >900°C beneath the arc). So as subduction continues and
older and older crust begins to be subducted, it is less easily melted:
The blocks labelled eclogite and blueschist show the P-T conditions found in
exhumed subduction complexes (e.g. the Franciscan of California) which are
consistent with an average age of 50 Ma old for subucted ocean crust.
(b) Amount of previously subducted lithosphere. Clearly the more you stuff cool
oceanic lithosphere into the upper mantle, the more it will cool it (the iced drink
analogy!). With subduction rates of 10 cm/yr it is possible to subduct 100 km of
ocean crust per m.y. [the diagrams are drawn for a much slower subduction rate of
3 cm/yr]
The implication from the diagram below is that the cooling effect of continued
subduction is quite severe, So after less than 600 km (= 6 m.y.) of ocean crust
subduction, temperatures are below those at which the slab melts. But what about
the mantle wedge?
The diagram also shows how the top and base of oceanic crust heats up. The
base is initially hotter, but the top eventually gets hotter because of heat
conducted from the mantle wedge.
(c) Magmas from the mantle wedge? Curves E and F show the temperatures of
the mantle wedge (straddling the depths at which magmas are generated below
arc volcanoes) at 10 m.y. and 20 m.y. after the start of subduction, but without
allowing any convection in the mantle wedge. The cooling effect of the slab is very
important, quickly taking the wedge below temperatures at which magmas would
be generated.
However, Curve G shows the effect of allowing induced convection in the mantle
wedge (a similar curve linked with E would be even higher temperature . . . ). In
this case the temperatures stay above 950 °C as the wedge material is dragged
down, and so hydrous melting would be possible. An important implication from
this diagram is that it is much more likely that arc magmas are derived from the
mantle wedge: conditions for slab melting are very restricted.
(d) Effect of induced convection on slab. Briefly, the modelling shows that induced
convection can enhance the meltability of older slab, but the effect on young
ocean crust is not important:
(e) Temperature or pressure control on magma generation?
The diagram above shows effect of water on the melting behaviour of basaltic
oceanic crust. Under dry conditions melting increases with pressure (red dashed
line). However, under water-saturated conditions (red full line) melting
temperatures plummet by almost 400°C at depths of 50 km.
Importantly, the blue curve shows how hornblende becomes an important mineral
under hydrous conditions; however, note that the curve turns over at ~70-80 km to
become pressure-sensitive – hornblende in mantle breaks down at ~100+ km.
This means that a lot of fluid will be released from hornblende at these depths,
which could promote melting. Is this why most arc volcanoes lie ~100 km above
the Benioff Zone?
(f) Upward and Downward flow in mantle wedge
There is increasng interest in subduction-induced flow in the mantle wedge. At
shallow levels (25-50km) the massive amounts of water entering the subduction
zone may hydrate the mantle wedge to give serpentinite: this rock contains >12%
water and is significantly less dense than normal mantle, and so can rise
diapirically and "intrude" (solid state flow) the fore-arc regions, whether formed of
arc volcanics or accreted sediment.
Further down, cooling of the wedge by the subduction zone itself may make it
negatively buoyant (i.e. denser) and help drag the wedge down, promoting
hornblende breakdown and fluid release. This will enhance induced convection
effects. The amount of coupling between slab and mantle wedge would however
be reduced by soft sediment at the interfact between the two.
Upward flow further back in the mantle wedge would compensate these effects,
particularly if enhanced by low-density fluid and magmas. (to be continued . . . )
REFERENCES
ANDERSON, R.N., DELONG, S.E. & SCHWARTZ, W.M. 1978. Thermal model for
subduction with dehydration in the downgoing slab. Journal of Geology 86, 731-739.
ANDERSON, R.N., DELONG, S.E. & SCHWARTZ, W.M. 1980. Dehydration,
asthenospheric convection and seismicity in subduction zones. Journal of Geology 88,
445-451.
CARLSON, R.L., HILDE, T.W.C. & UYEDA, S. 1983. The driving mechanism of plate
tectonics: relation to age of the lithosphere at trenches. Geophysics Research Letters 10,
297-300.
DAVIES, J.H. & BICKLE, M.J. 1991. A physical model for the volume and composition of
melt produced by hydrous fluxing above subduction zones. Philosophical Transactions of
the Royal Society, London A335, 355-364.
DAVIES, J.H. & STEVENSON, D.J. 1992. Physical model of source region of subduction
zone magmatism. Journal of Geophysical Research 97, 2037-2070.
DEFANT, M.J. & DRUMMOND, M.S. 1990. Derivation of some modern arc magmas by
melting of young subducted lithosphere. Nature 347, 662-665.
DEWEY, J.F. 1981. Episodicity, sequence and style at convergent plate boundaries. In The
Continental Crust and its Mineral Deposits. Geological Association of Canada, Special
Paper 20, 553-572.
GARFUNKEL, Z., ANDERSON, C.A. & SCHUBERT, G. 1986. Mantle circulation and the
lateral migration of subducted slabs. Journal of Geophysical Research 91, 7205-7223.
HARGRAVES, R.B. 1986. Faster spreading or greater ridge length in the Archean?
Geology 14, 750-752.
KINCAID, C. & OLSON, P. 1987. An experimental study of subduction and slab migration.
Journal of Geophysical Research 92, 13832-13840.
MOLNAR, P. & ATWATER, T. 1978. Interarc spreading and cordilleran tectonics as
alternates related to the age of subducted ocean lithosphere. Earth and Planetary Science
Letters 41, 330-340.
PEACOCK, S.M. 1987. Thermal effects of metamorphic fluids in subduction zones.
Geology 15, 1057-1060.
PEACOCK, S.M. 1991. Numerical simulations of subduction zone pressure-temperaturetime paths: constraints on fluid production and arc magmatism. Philosophical Transactions
of the Royal Society, London A335, 341-353.
RINGWOOD, A.E. 1974. The petrological evolution of island arc systems. Journal of the
Geological Society, London 130, 183-204.
STERN, R.J. & BLOOMER, S.H. 1992. Subduction-zone infancy - Examples from the
Eocene Izu-Bonin-Mariana and Jurassic California arcs. Geological Society of America
Bulletin 104, 1621-1636.
SUDO, A. & TATSUMI, Y. 1990. Phlogopite and K-amphibole in the upper mantle:
implications for magma genesis in subduction zones. Geophysics Research Letters 17, 2932.
UYEDA, S. & KANAMORI, H. 1979. Back-arc opening and mode of subduction. Journal of
Geophysical Research 84, 1049-1061.
WYLLIE, P.J. 1988. Magma genesis, plate tectonics and chemical differentiation of the
earth. Reviews of Geophysics 26, 370-404.
THE FATE OF SEDIMENTS AT SUBDUCTION ZONES
The floors of the world's oceans are covered by sediment up to 1 km thick (age
dependent) as a result of slow accumulation of calcareous and siliceous biogenic
oozes capped by fine clays that have been carried in suspension to the middle of
oceans. Additionally, nearer continents there may be much thicker accumulations
of clastic sediments brought in by deltas and turbidity currents, and further redistributed by strong bottom water currents. Sooner or later this sedimenty must
finish up at a subduction zone. What happens to it? Does it get scraped-off, or
does it get dragged down the subduction zone? If the latter, does it just disappear
into the deep mantle, or does it get recycled into island arc magmas? The balance
is shown as follows:
Effectively, subduction at active margins can be likened to a conveyor belt carrying
a lot of loose rubbish moving against a buttress: some material is going to get
scraped-off:
There are many variables in the whole process. So it is important to look at a
number at different tectonic situations.
(1) Primitive Island Arcs: no sediment accretion
At intraoceanic island arcs, such as the Marianas, there is no sediment supply
from the continent (this is trapped by the back-arc basin), and the arc itself
produces only a minor amount of volcanic ash (the eruptions are basaltic and not
violent). Most of the sediment arriving at the subduction zone is abyssal ooze and
clay carried on the subducting plate (on old ocean crust, at least 0.5km thick). It
used to be thought that this abyssal sediment was scraped off to form an
accretionary wedge in the fore-arc. However, dredging and drilling in the Mariana
forearc and trench has shown that there is little on no sediment in the Mariana
trench. Yet during the 40 my since the arc system has been in existence, up to
40 km³ of sedimen / km length of arc should have been scraped off the subducting
plate (which is subducting at 10 cm/year).
The sediment must be subducted - but how? The answer seems to be that, as the
subducting plate bends over to become vertical, the flexure causes horsts and
graben to develop. Sediments are scraped off from the horsts into the graben and
thus encased as the ocean lithosphere deforms (for this reason it was thought this
would be a good place to dispose of nuclear waste!) In fact the ocean crust acts as
a gigantic rasp on the arc too - the forearc is gradually, but slowly, eroded:
However geochemical studies have shown that very little of the sediment is
actually incorporated into the arc volcanics, so most of it must be cycled into the
deeper mantle. Presumably, as the slab at the Marianas is avalanching into the
lower mantle, the sediments may be taken down also.
(2) Northern Chile: no sediment to subduct
Here the sediment supply is also very limited because of the arid climate. Many of
the rivers from the high Andes never quite make it to the ocean, and in any case
there few floods (which produce the turbidity currents that carry the sediment out
into the ocean proper). Also, major faults parallel to the coast tend to obstruct the
rivers, forming saline lakes (were common in N. Chile).
So the situation is similar to that in the Marianas, although the dip of the
subducting slab is not so great. Some geologists have suggested that the rasping
action of the subducting slab has actually eroded back the continental margin of N.
Chile and Peru. Is this why the locus of volcanic activity continually moves
eastwards with time in the N. Andes? And why Palaeozoic batholiths are exposed
right at the coast, close to the trench? (Although difficult to prove it was there when
it has gone!).
Where sediment supply is a little higher, trench gets partly filled with sediment.
Some of this sediment may get scraped off. But drilling in the Middle America
Trench suggests that the abyssal ocean floor sediments are still subducted (soft
oozes act as a lubricant)
(3) S. Chile and Alaska: high sediment input
Here the climate is temperate and wet. Abundant rivers, some deriving from
glaciers. Floods common. High rate of sediment supply to the ocean. Sediment
supply was even higher during the Pleistocene (and there has not been time yet to
subduct them).
Result is that large amount of sediment is carried into the trench. Trench quickly
gets filled, and sediment then carried out onto subducting plate. As this continues
the weight of sediment actually depresses the plate as it approaches the trench so
that angle of dip is smaller (dip increases under the continental margin proper).
With a shallower dip, no horsts & graben form, and sediment is scraped off. This
can readily be seen from reflection profiles. Layering of sediments disappears as
continent is approached. Low angle thrusts appear. Younger sediments are
progressively underplated. If sedimentation rates are high (as they are in high
northern/southern latitudes) this can give rise to lateral growth of continents. The
process is called subduction-accretion and the structures are called Accretionary
Prisms. The general features are shown below:
(4) Characteristics of Accretionary Wedges/Prisms
Lateral continental growth by subduction-accretion is dependent on (a) the supply
of material from the ocean, and (b) the sediment supply from the continent. These
two might vary over a large range.
(a) Material accreted from oceans
The ocean floor is not smooth. Study of the Pacific map shows that the preTertiary ocean floor is considerably rougher than that generated in the Tertiary.
There are more oceanic plateaus, aseismic ridges, ocean island chains and arcs –
in large part this results from the spate of mantle plumes which punched through
the Pacific ocean plate in the late Cretaceous (120 - 80 Ma).. Many of these
upstanding structures are capped by carbonate banks, because they stayed above
the carbonate compensation depth (CCD) much longer than normal ocean floor.
Ocean floor that is rough and upstanding is more likely to be scraped off when it
reaches subduction zones at active margins. So this sceaped-off material will be a
mixture of mafic rocks (metamorphosed to amphibolite) associated with thick
limestone (marble) sequences, as well as sileceous and carbonate oozes (=
"cypoline schists") and lithified cherts. Large oceanic structures such as plateaus
and arcs may "choke" the subduction zone, causing back-stepping of the
subduction zone, the arcs being left as an ophiolite (e.g. the classic Troodos
complex on Cyprus). However, normal ocean floor, which is smooth and cold, may
not be scraped off at all (it is th is that converts to eclogite to provide the slab-pull
force), so the soft carbonate-siliceous oozes and cloay may not be scraped off
quite so readily.
(b) Material supplied from the continents
This is largely material supplied by river systems feeding active continental
margins. Of course at the present day there are not many rivers feeding active
continental margins -- they are mostly still feeding the passive margins of the
Atlantic, the Indian ocean and around Antarctica/Australia.
It is important to note that in the Upper Palaeozoic and early Mesozoic, the
southern continents formed part of Gondwanaland - a very large continental
landmass. Moreover much of Gondwanaland was rimmed by active margins. The
margin had low relief (the present high Andes is not typical, and results from
Miocene deformation and uplift). So it is probable that very large rivers were
dumping sediment onto the subducting plate, and the sediment was then accreted
back on to the continental margin . . . now exhumed and exposed, particularly in
southern Chile, where they are of late Palaeozoic age (before the Andean
magmatic cycle), and South Island, New Zealand. But they can also be seen in
Alaska, and of course occur in older mountain belts (commonly termed Flysch).
Compared with the partly-lithified material scraped off from the oceanic plate, the
material coming from the continent is unlithified clastic sediment. The two get
tectonically intermixed and intensely deformed (the subduction interface allows
thousands of km of relative movement in just a few tens of Ma ™ far more than
with continental collision), so most rocks from this environment have strong
penetrative foliations and linear fabrics (see New Harbour Group on Anglesey) and
finish up as teconic melanges -- lenses of oceanic rocks in deformed soft
sediment.
As soft wet sediment (greywacke-shale) is continually underplated beneath the
accretionary wedge, it heats up slowly. Water is progressively driven off. Hot water
dissolves silica from sandy beds, and deposits it at higher levels as abundant
cross-cutting quartz veins. However, because underplating is continuous process,
sediments and quartz veins become progressively and very strongly deformed.
Can be almost mylonite-like fabric. No bedding remains. Cross-cutting quartz
veins are stretched out to become sub-parallel to foliation. Very characteristic rock
type. Many tens or even hundreds of km of 'new' crust can accrete laterally onto
continental margins in this way.
Erosion of upper part of accretionary wedge may occur, and younger sediments
deposited on top in fore-arc basins. These may also become deformed, but less
so (could the South Stack Series on Anglesey may represent such fore-arc basin
rocks?).
References
DAVIES, J.H. & von BLANCKENBURG, F. 1995. Slab breakoff: A model of lithosphere
detachment and its test in the magmatism and deformation of collisional orogens. Earth
and Planetary Science Letters 129, 85-102.
von HUENE, R. & SCHOLL, D.W. 1993. The return of sialic material to the mantle
indicated by terrigenous material subducted at convergent margins. Tectonophysics 219,
163-175.
TERRANES
"Terrane" concepts are now quite widely used in interpreting geological
relationships in many parts of the world, and in rocks of many ages. Basically plate
tectonics can move segments of continental crust or oceanic crust (e.g. ocean
plateaus) many thousands of km in just a few tens of m.y., and as plates can
change their direction of motion (c.f. kinik in Hawaiian chain), this can lead to the
juxtaposition of segments of crust that have a completely different geological
histories. So it is not just collision of major continents (e.g. India and Asia to form
Himalayas) but also on a much smaller scale. In particular, major transform faults
can transport differnt crustal segments laterally for many 1000s of km (e.g. San
Andreas Fault). Of course terranes are usually fault- or thrust-bounded.
Terrane Terminology (Jargon)
"A fault-bounded package of strata that has a geological history distinct from the
adjoining geologic units"
Howell (1989) divided terranes as follows:
Stratigraphic (1) representing fragments of continents
(2) fragments ofcontinental margin
(3) fragments of volcanic arc
(4) fragments of ocean basins
Disruptive
Metamorphic
However, a genetic terminology is also prevalent:
Exotic, Suspect, Displaced or Accreted terranes: this implies that the terrane has
been transported some distance to its current position.
Pericratonic: Contains cratonal detritus and formed on attenuated continental
crust.
Terranes are sometimes described in terms of tectonic assemblages, which are
rock-stratigraphic units formed in actualistic tectonic settings, such as island arcs
or ocean floors. A terrane may consist of one or more tectonic assemblages
Domain: A volume of rock, bounded by compositionalor structural discontinuities,
within which there is structural homogeneity; these may contain minor stratigraphic
distinctions as well andcan be viewed as subterranes.
Superterranes: A composite terrane, consisting of two or more compound
terranes, that were amalgamated prior to subsequent orogenesis.
REFERENCES (General)
CONEY, P., JONES, D.L. & MONGER, J.W. 1980. Cordilleran suspect terranes. Nature
288, 329-333.
BEN-AVRAHAM, Z., NUR, A., JONES, D. & COX, A. 1981. Continental accretion: from
oceanic plateaus to allochthonous terranes. Science 213, 47-54.
HOWELL, D.G. 1989. Tectonics of Suspect Terranes. Chapman & Hall, NewYork, 232pp.
BEBOUT, G.E. & BARTON, M.D. 1989. Fluid flow and metasomatism in a subduction zone
hydrothermal system: Catalina Schist terrane, California. Geology 17, 976-980.
References On Alaskan Terranes
VROLIJK, P., MYERS, G. & MOORE, J.C. 1987. Warm fluid migration along tectonic
melanges in the Kodiak accretionary complex, Alaska. Journal of Geophysical Research
93, 10313-10324.
BARKER, F., JONES, D.L., BUDAHN, J.R. & CONEY, P.J. 1988. Oceanic plateauseamount origin of basaltic rocks, Angayuchan Terrane, Central Alaska. Journal of
Geology 96, 368-374.
References On Caledonian Terranes
DEWEY, J.F. & SHACKLETON, R.M. 1984. A model for the evolution of the Grampian
tract in the Caledonides and Appalachians. Nature 312, 115-121
MURPHY, F.C. & HUTTON, D.H.W. 1986. Is the Southern Uplands of Scotland really an
accretionary prism? Geology 14, 54-57.
HUTTON, D.H.W. 1987, Strike-slip terranes and a model for the evolution of the British and
Irish Caledonides. Geological Magazine 124, 405-425.
BENTLEY, M.R., MALTMAN, A.J. & FITCHES, W.R. 1988. Colonsay and Islay: a suspect
terrane within the Scottish Caledonides. Geology 16, 26-28.
HAUGHTON, P.D.W. 1988. A cryptic Caledonian flysch terrane in Scotland. Journal of the
Geological Society, London 145, 685-703.
MARCANTONIO, F., DICKIN, A.P., McNUTT, R.H., & HEAMAN, L.M. 1988. A 1800 million
year old Proterozoic gneiss terrane in Islay with implications for the crustal structure
evolution of Britain. Nature 335, 62-64.
SOPER, N.J., GIBBONS, W. & McKERROW, W.S. 1989. Displaced terranes in Britain and
Ireland. Journal of the Geological Society, London 146, 365-367.
THIRLWALL, M.F. 1989. Movement on proposed terrane boundaries in northern Britain:
constraints from Ordovician-Devonian igneous rocks. Journal of the Geological Society,
London 146, 373-376.
BLUCK, B.J. & DEMPSTER, T.J. 1991. Exotic metamorphic terranes in the Caledonides:
Tectonic history of the Dalradian block, Scotland. Geology 19, 1133-1136.
RYAN, P.D. & DEWEY, J.F. 1991. A geological and tectonic cross-section of the
Caledonides of western Ireland. Journal of the Geological Society, London 148, 173-180.
MURPHY, F.C., ANDERSON, T.B., DALY, J.S. & 16 others, 1991 An appraisal of
Caledonian suspect terrains in Ireland. Irish Journal of Earth Sciences 11, 11-41.
SOPER, N.J., ENGLAND, R.W., SNYDER, D.B. & RYAN, P.D. 1992. The Iapetus suture
zone in England, Scotland and eastern Ireland: a reconciliation of geological and deep
seismic data. Journal of the Geological Society, London 149, 697-700.
BROWN, C. & WHELAN, J.P. 1995. Terrane boundaries in Ireland inferred from the Irish
Magnetotelluric Profile and other geophysical data. Journal of the Geological Society,
London 152, 523-534.
References On Appalachian Terranes
WILLIAMS, H. & HATCHER, R.D. 1982. Suspect terranes and accretionary history of the
Appalachian region. Geology 10, 530-536.
References On Andean Terranes
ASPDEN, J.A. & McCOURT, W.J. 1986. Mesozoic oceanic terrane in the central Andes of
Colombia. Geology 14, 415-418.
Baltic Shield Proterozoic Terranes
PARK, A.F. 1991. Continental growth by accretion: a tectonostratigraphic terrane analysis
of the evolution of the western and central Baltic Shield, 2.50 to 1.75 Ga. Bulletin of the
Geological Society of America 103, 522-537.
References On Archaean Terranes
(to be continued)
PLATE TECTONICS: Lecture 2
OCEAN RIDGE MAGMATISM
Magma production at the Earth's mid-ocean ridge system far exceeds that in any
other tectonic environment, and this has been so since the early Precambrian. It is
the dominant way in which internal heat is dissipated. The structure of a midocean ridge is shown below:
Note how the lithosphere thickens as it moves away from the ridge. Because the
Earth's magnetic field oscillates between north and south at intervals of a few
hundred thousand (or the odd million) years the basalts erupted then take on the
current magnetisation, and so give rise to the seafloor magnetic lineations
(patterns shown above) that can be used to date the ocean floor. Melting of
pyrolite mantle extracts basaltic liquids to form the ocean crust, leaving a residue
of harzburgite (ol+opyx) forming the underlying lithosphere.
The ocean lithosphere suffers extensive hydrothermal alteration at the ridge (see
below), but the rocks eventually finish up subducting back into the mantle:
It is because these fluids are released in the Benioff Zone as the slab is subducted
that magmas are able to be generated in the mantle wedge above the subduction
zone. It is fluid, not friction, which is responsible for active margin magmatism. But
it is ridge processes which make it all possible. So we need to look at these.
Why does melting occur? Melting temperatures of most silicate minerals increase
with increasing pressures. So temperatures of solid mantle material at depth may
be higher than the melting point of mantle near the earth's surface. As hot deep
mantle rises beneath spreading ridges it will, as pressure falls, rise above its
solidus, and begin melting.
The simplified situation is as follows:
As the uprising mantle crosses the geotherm it begins to melt, and as the solidus
temperature of mantle falls with decreasing pressure, the temperature of the melt
increases relative to this solidus, thus effectively giving higher degress of melting
with decompression, as shown. The amount of melt generated will be limited by
the latent heat of fusion (which is high for silicates), and as the melting range of
mantle peridotite lies between ca. 1100°C and ca. 1700°C, it is likely that most
ridge basalts are partial (rather than complete) melts of mantle. The magma may
enter a chamber in the ocean crust and begin crystallising, giving the following P-T
path:
There is the possibility of superheat (i.e. temperature above the liquidus) if the
magma can rise quickly, but it is apparent that most magmas are erupted or
emplaced without superheat (a possible exception are ultramafic lavas called
komatiites).
Because we haven't yet been able to drill very far down into oceanic crust, the only
way we can begin to understand what happens to the basaltic magma as it rises
up at the ridge is to look at ophiolite complexes. There are many of these in the
Alpine belt, although we are not always sure that these mafic slivers represent true
ocean basin crust or whether some (or all) may represent marginal basin crust or
the roots of island arcs.
Nonetheless, by putting together information from a number of ophiolite
complexes, particularly Troodos on Cyprus, we come up with the following
idealised section:
Not every ophiolite has all these components complete, and it is not always for
tectonic reasons. Often the gabbro is missing, or the sheeted dykes, and in some
cases the dykes may intrude the harzburgite. Of course sheeted dykes can only
be formed if there is a continuously extending magma chamber (try doing it
without!). So if sheeted dykes are missing it may mean that there has not been
such a magma chamber. In fact there is a lot of debate on this issue. Some
geophysical studies indicate a possible continuous magma chamber beneath the
East Pacific Rise. However, the EPR is a smooth fast-spreading ridge, and maybe
there is enough thermal input to keep a continuous magma chamber going. On the
other hand in the slow-spreading Atlantic with its central rift valley and irregular
topography, there is no direct evidence for a continuous magma chamber. Some
workers, including those at Leicester, suggest that with slow-spreading ridges,
each eruption may be a distinct event, and that any magma chamber is only shortlived. Some sections of the Atlantic ridge, like the FAMOUS area (south of the
Azores) have numerous small volcanic cones, and this is now being recognised all
over the Atlantic.
A consequence is that that there may be a variety of magma chamber profiles,
with those from fast-spreading ridges having fat "onion" shapes, those from rather
slower-spreading ridges having "leek" shapes. Very slow spreading ridges (e.g.
SW Indian Ridge) may just have dykes feeding lavas which directly overly
peridotite. There are ophiolites with this profile, where the dykes cut harzburgite
tectonite and gabbro is only locally developed. Even with the type Troodos
ophiolite, which has a moderatley thick gabbro section, geochemical studies have
shown that the gabbros are in fact a compound of a number of small bodies.
Transform Fault Effects
It has long been known that the ocean crust is much thinner in the vicinity of
oceanic transform faults. Also that a greater variety of rock types can be drilled or
dredged in the vicinity of transforms, and that there is usually a significant
topographic difference between the two sides of a transform fault (esp. the larger
ones).
The latter effect arises because the ocean crust sinks as much as 3 km over the
first 50 m.y. of its existence. So the greater the age difference of adjacent bits of
ocean crust across a transform, then the greater the height of the transform wall.
Obviously if the wall is 1 km high, then a large amount of rubble will fall down onto
the lower plate, and deeper parts will become exposed. Moreover as the transform
fault moves, the movement can deform the basalts into hornblende schists.
The thinner crust arises from the cold-wall effect, i.e. that the mantle rising up
adjacent to the transform fault are actually in contact with older, and therefore
cooler, oceanic crust on the other side. Cooler conditions give less melt and
therefore thinner crust. Thinner crust also means there is more likelihood of mantle
being exposed in the transform wall, again increasing the variety of rock types.
METAMORPHISM OF OCEANIC CRUST
There has been a very great deal of interest worldwide in the metamorphism and
hydrothermal alteration of oceanic crust. After all there are few geological
situations where you have a large red-hot magma chamber below and a 3 km
column of ocean water on top trying to dowse it. There are a number of important
questions that could be asked:
(a) how extensive is the metamorphism, and how far distant from the ridge
do the metamorphic effects extend?
(b) if there is extensive hydrothermal activity, does this lead to equally
extensive mineral deposits which could be mined?
(c) would the metamorphism affect the magnetic anomaly patterns that are
so useful for dating ocean crust?
(d) is the ocean crust so hydrated that this represents an important source
of fluids at subduction zones?
(e) does the hydrothermal interchange influence the chemical budget of the
oceans?
There is no doubt that ophiolite complexes (obducted bits of ocean crust in
mountain belts) are usually >90% altered, and there was debate during the 60's
and 70's whether this was a result of metamorphism in the mountain belt during
orogensis, or resulted from ocean floor metamorphism. The latter is now the
favoured explanation. Many of the samples of ocean crust recovered from the
ocean floor by drillling or dredging are altered.
Metamorphism
Cann (1979) recognises 5 different mineral assemblage facies in oceanic basalts
recovered by dredging, drilling etc. The rocks characteristically preserve igneous
textures.
(1) Brownstone Facies
Low temperature ocean floor weathering or cool hydrothermal alteration. Products
usually have yellowish brown tint due to oxidising conditions (bluish grey if
reducing). Mineral assemblages not in equilibrium; just replace specific primary
phases.
Olivine replaced by Celadonite (K-rich dioctahedral Fe-illite) under more extreme
alteration. This fills vesicles and replaces glass. Under reducing conditions this is a
Saponite (Mg-rich trioctahedral smectite). Pyrite common. Thus basalt has clay
alteraton products.
Plagioclase remains fresh, though under extreme alteration may be partly replaced
by K-feldspar (Adularia).
Glass: Where basalt glass is common, Palagonite (orange coloured disordered
illite) occurs, often associated with a low temperature zeolite (Phillipsite) and
Calcite.
Fig. 8.
Diagram from
Cann (1979)
tries to
indicate how
the minerals
in a basalt
affected by
hydrothermal
activity
contribute to
the
secondary
phases. At
low
temperatures
it is mainly
the basalt
glass and
olivine which
are unstable
and
contribute to
brownstone
facies
minerals, but
plagioclase
and then
augite and
iron oxide
become
progressively
involved at
higher
grades until
the whole
rock
recrystallises.
(2) Zeolite facies ( Temperature above 50-100°C.)
Here Phillipsite is replaced by higher temperature zeolites - Analcite and Natrolite.
Distinct zones of zeolites occur on Iceland.
Mafic minerals replaced by Saponite or Saponite-Chlorite mixed layer minerals,
coarser grained than in Brownstone Facies. Plagioclase may also be partly
replaced by saponite, but augite stays fresh. Upper limit of facies (250-300°C)
marked by disappearance of zeolites and saponite and incoming of albite and
chlorite.
(3) Greenschist Facies
Albite ± chlorite ± actinolite ± epidote ± sphene. Degree of alteration variable,
primary assemblages may be completely replaced. Augite is commonly a relic,
veins are common, often quartz-bearing. Assemblages may or may not be
equilibrium ones. Upper limit of facies marked by disappearance of albite, chlorite
and actinolite and the appearance of green aluminous hornblende associated with
more calcic plagioclase (An20-30).
(4) Amphibolite Facies
Hornblende+Ca-plagioclase + titanomagnetite±epidote. This assemblage is most
commonly developed in coarser grained rocks - dykes and gabbros - obviously of
deeper origin. Degree of metamorphism variable. Some primary hornblende
occurs in gabbros or diorites, but it is clear that amphibolite facies metamorphic
assemblages are superimposed on this. So metamorphism closely follows
magmatic activity.
The results can be summarised in the following table. Note that it is not just the
mafic ocean crust (basalt or dolerite) that is altered. The mantle itself is often
brought up along faults, transforms and fracture zones, and this is frequently
altered to serpentine (ca. 13% water) at temperatures below 450°C. But other
hydrous minerals such as talc, tremolite and chlorite are possible at higher
temperatures. There are also 3 different varieties of serpentine: antigorite,
chrysotile (the glossy variety) and lizardite.
_________________________________________________________________
_______________
Summary of Mineral Assemblages in Altered Crust
Facies
BASALT
Brownstone Celadonite +
PERIDOTITE
Phillipsite +
Palagonite + Saponite
?
Zeolite
Saponite + mixed layers
?
+ analcite + natrolite
Chlorite + Albite
Greenschist + Actinolite +
Epidote + sphene
Lizardite
Chrysotile
Magnetite
Hornblende +
Amphibolite Plagioclase +
Iron Oxide
Tremolite
+ Olivine
+ Enstatite
Gabbro
Augite +
Plagioclase +
Hypersthene +
Iron Oxide +
Olivine
Enstatite
+ Diopside
Chromite
_________________________________________________________________
___________
General Comments:
Greenschist and amphibolite facies metamorphism of the ocean floor differs from
these facies in normal regional metamorphism in that:
(a) The thermal gradient is very high: can be several-100°C per km
compared with 30 - 50°C/km in regional metamorphism.
(b) No garnet is developed in mafic rocks (pressures are not high enough).
(c) The rocks lack deformation textures (except in samples recovered from
transform fracture zones)
(d) Very variable degree of recrystallization, because lower grade
metamorphic assemblages are superimposed on earlier higher grade ones.
This is because hydrothermal activity continues under cooler conditions as
crust progressively moves away from ridge. (In regional metamorphism it is
more common for the rocks to equilibrate at one set of P-T conditions)
Despite ca. 20 years of drilling, the deepest drill holes in the ocean floor (several
hundred metres) have still only penetrated brownstone- and zeolite-facies rocks.
No greenschist facies or amphibolite facies rocks penetrated. To see what
happens deeper down, we really need to examine ophiolite complexes.
Ocean floor metamorphism - Sarmiento Ophiolite, Chile
The Sarmiento ophiolite (Saunders et al. 1979) is one of a series of discontinuous
mafic lenses that represent the mafic floor of an extensional back-arc basin that
was closed and uplifted in mid-Cretaceous times in the southern Patagonian
Andes. Excellent vertical exposures have enabled the distribution of metamorphic
zones resulting from the hydrothermal metamorphism to be established (Stern &
Elthon, 1982?).
Lithological sequence at Sarmiento consists of:



2 km pillow lavas
300 m sheeted dykes
1 km gabbros with plagiogranite
Four main metamorphic equilibrium mineral facies can be recognised:
(1) Zeolite Facies. Zeolites, palagonitized glass ± smectites ± calcite ±
quartz ± pyrite ± sphene ± albite.
(2) Greenschist facies. Chlorite, epidote, Na-plagioclase, sphene, ± quartz ±
calcite ± biotite ± pyrite.
(3) Lower Actinolite Facies. Low-Al (2-5% Al2O3) fibrous green amphiboles,
Ca-plagioclase, sphene ± biotite ± calcite.
(4) Upper Actinolite Facies. Higher-Al (5-8% Al2O3) brown-green
amphiboles, Ca-plag. (>An50), titanomagnetite, ± ilmenite ± biotite.
These are arranged vertically in the complex, with metamorphic grade increasing
downwards. However, intensity of metamorphism also varies and is at a maximum
within the sheeted dyke unit. Moreover, the lower temperature facies may be
superimposed on the higher grade ones. The histograms show the intensity of
metamorphism in each of the components of the complex for each of the four
metamorphic grades.
Significant Points:
(1) The intensity of metamorphism (recrystallisation) is greatest in the sheeted
dykes. This is because the vertical dyke margins permit easy access of circulating
fluids, coupled with the fact that higher temperatures speed reaction rates. The
high water-rock ratios in the sheeted dyke zone mean that the rocks are strongly
leached. Any chemical elements not required by the newly forming minerals
(hornblendes and chlorite) are removed upwards by the fluids. This means
elements like Rb, U, Th, K, some Sr, Ba and chalcophile elements like Zn, Cu and
Pb are removed from the deep dyke rocks. Some of the former group will be reabsorbed by the zeolites and clays in the uppermost 'Brownstone' part of the
section, others are dissolved in seawater, whereas the chalcophile elements form
valuable mounds of sulphide on the seafloor (black smokers). Overall, the
hydrothermal activity achieves a major amount of vertical chemical redistribution
within the ocean crust. The mobile elements are moved to the top of the ocean
crust. These are the same elements that become "mobile" when the ocean crust is
subducted and arc magmas are generated.
(2) As thermal gradients fall (i.e. crust moves away from ridge axis) circulating
water permits lower grade assemblages to form - superimposed on high grade
ones. But this secondary metamorphism is of lower intensity because circulation
channels become blocked with the growth of secondary minerals. The
sedimentary cap that progressively builds up on top of the basalts will eventually
block circulation. There is a good example of this in the "Megaleg"
Chemical fluxes in oceanic crust - the 'MEGALEG'
The Deep Sea Drilling Project Legs 51-53 drilled two deep 200 m holes in
Cretaceous (110 m.y.) oceanic crust in the western Atlantic near Bermuda. The
holes were only 450m apart, but one hole, Hole 417A, drilled some of the most
altered basalts found on the ocean floor, whereas those in Hole 417D were
relatively fresh. All alteration was at 'Brownstone' facies. Basalts in both holes
petrographically similar. The differences just reflect the relative access by
circulating fluids.
Compare compositions (water free):
417A
hyaloclastite
417A avge 417D avge
SiO2
53.6
49.9
49.4
TiO2
1.13
1.50
1.50
Al2O3
11.4
10.9
10.5
MgO
5.80
5.44
6.14
CaO
3.68
10.3
13.5
Na2O
1.70
2.21
2.40
K2O
4.36
1.79
0.12
Note significantly higher K2O and lower CaO in altered rocks, particularly the most
fragmentary hyaloclastite 417A, which is the most permeable of rock types.
Interpretation: Hole 417A is on basement 'high' which remained uncovered by
sediment for a considerable time thus permitting long-term circulation of warm
(30°C) water. Hole 417D was a located in a topographic depression which became
quickly filled with sediment which blocked extensive water circulation.
The results were first interpreted to suggest that the ocean crust may be a
tremendous 'sink' for K2O and Rb transferred from sea water through hydrothermal
circulation involving major volumes of seawater. However, how much of the Rb,
K2O, etc. is derived through leaching from more altered rocks (e.g. in sheeted dyke
unit) at deeper levels?
To see what other chemical changes occur during alteration of oceanic crust it is
best to look at the "type" ophiolite, the Troodos Complex on Cyprus. This has been
intensively studied through field investigations, mining operations and by scientific
drilling. Troodos formed at 91 Ma, but hydrothermal activity continued for a further
40 Ma after crust formation (Gallahan & Duncan 1994).
Chemical Changes in Oceanic Crust - Troodos Ophiolite
1. Strontium Isotopic Composition
Studies by Spooner et al. (1977) of zeolite- to amphibolite-facies altered basalts on
Cyprus show that 87Sr/86Sr ratios are increased relative to fresh basalts and
gabbros.
87Sr/86Sr
Zeolite
0.80760 ± 3
Altered Basalt
0.7069
Fresh Gabbros
0.70338 ± 10
Cretaceous Seawater
0.7076
Similar results have been obtained on altered ocean basalts. High 87Sr/86Sr in
altered mineral products can be leached away in dilute acid so that unaltered
minerals yield the original magmatic Sr isotopic ratios.
Spooner suggested that interchange of seawater Sr with ocean crust Sr occurs
during hydrothermal circulation and may buffer Sr isotope composition of
seawater:
87Sr/86Sr
Fresh Ocean Crust Av.
0.703
Seawater Average
0.709
Continental Rocks
> 0.712
There is a considerable variation in 87Sr/86Sr in seawater with time that can be
linked to varying plate activity. This is discussed in more detail below.
2. Oxygen and Hydrogen Isotopic Compositions
Spooner et al. (1977) showed that oxygen isotope ratio values in Troodos and
other Mediterranean ophiolites were higher ( d18O = ca. 9) than in fresh 'mantlederived' basalts (d18O = 6) and were consistent with alteration by seawater at high
temperatures of ca. 350°C. The interesting point about this is that because oxygen
is the most abundant element in any rock, it is necessary to exchange almost all
the oxygen in the rock to significantly change the isotopic ratio. In other words,
water-rock ratios are high, or large volumes of seawater interact with ocean crust
at spreading centres. The implication is that if oxygen can be exchanged on this
scale then many other elements can be changed too.
Heaton and Sheppard (1977) showed that the isotopic composition of hydrogen in
water in equilibrium with chlorite and amphibole from altered dykes from Cyprus
was indistinguishable from that of seawater.
Comment: Altered oceanic crust (now with higher 87Sr, 18O and 2H contents)
which is subducted at Benioff Zones may modify the isotopic composition of island
arc magmas from "normal" mantle values. The hydrous fluids driven off as the
subducting slab heats up as it goes down subduction zones will be enriched in the
heavy isotope of these elements. So it is not surprising that island arc magmas
differ in their isotopic ratios from other mantle-derived igneous rocks.
Sulphide Ore Deposits in Oceanic Crust
Sulphide deposits have been found on East Pacific Rise. They occur at positions
of discharge of hydrothermal systems ("black smokers"). On Cyprus ore bodies
are 500 m x 350 m x 50 m, and consist of pyrite and chalcopyrite with accessory
marcasite, sphalerite and galena. Chemical and isotopic data suggest that the
sulphide deposits mostly formed on seafloor:
(a) Fluid inclusions in ore material have composition of seawater.
(b) Ore material has 87Sr/86Sr = 0.7075.
(c) Hydrogen isotope composition same as seawater.
Calculations suggest ore bodies may have formed in 100,000 yrs. Fluid inclusion
studies suggest that the temperature of the plume of rising hydrothermal fluid was
300 - 350°C. Spooner suggests that contained sulphur has isotopic composition of
seawater sulphate. So ocean crust ore sulphide may be largely of reduced
seawater sulphate origin. The following diagrams illustrate some of the processes
of convective seawater circulation and the respective mineral zones in the
formation of hydrothermal mounds on the ocean floor (l.h. side) and other
submarine envirnments (r.h. side of 1st picture):
These diagrams illustrate the importance of fault-control on the location of the
discharge zones of hydrothermal activity, of permeability in allowing the
hydrothermal solutions to circulate, of P-T-pH-Eh in controlling which minerals are
stable and thus which elements are leached and which are deposited.
Comment: Large amounts of sulphide in adddition to chloride and hydroxyl are
added to ocean crust as a result of hydrothermal activity. A lot of pre-concentration
of potential ore metals already occurs in the ocean crust. So what happens when
ocean crust goes down subduction zones?
At subduction zones chlorine- and sulphide-rich fluids are released during
dehydration. Could this give us a possible explanation for porphyry copper
deposits that occur commonly at continental margins like the Andes? Spooner has
stressed that water is needed as transport medium, chloride for metal complexing
and sulphur for fixing the metals as solid phases. All these are present in ocean
crust as it is subducted.
Metal Deposits on Ocean Floor
Sulphide ores common in ophiolites. Could we locate them on ocean floor? Many
hydrothermal discharge zones have now been found at ocean ridges (mostly in the
East Pacific, but also now in the Atlantic) by submersible investigations. These are
potentially important ore reserves in terms of total volume, but individual deposits
are too small to mine economically, even by remote techniques. Discharge areas
may be located by trace element profiles in seawater near the ocean bottom. So
we can find them, but to exploit them is another matter.
REFERENCES: Oceanic Crust
CANN, J.R. 1974. A model for ocean crust structure developed. Geophysical Journal of the
Royal Astronomical Society 39, 169-187.
CANN, J.R. 1979. Metamorphism in the ocean crust. In: TALWANI, M. et al. (eds) Deep
Drilling Results in the Atlantic Ocean: Ocean Crust. Maurice Ewing Series 2, 230-238.
American Geophysical Union
CARLSON, R.L. & JOHNSON, H.P. 1994. On modelling the thermal evolution of the
oceanic upper mantle: An assessment of the cooling plate model. Journal of Geophysical
Research 99, 3201-3214.
ELTHON, D. & STERN, C.R. 1978. Metamorphic petrology of the Sarmiento ophiolite
complex, Chile. Geology 6, 464-468.
FLOWER, M.J. 1991. Magmatic processes in oceanic ridge and intraplate settings. In:
FLOYD, P.A. (ed) Oceanic Basalts. Blackie, Glasgow, pp.116-147.
GALLAHAN, W.E. & DUNCAN, R.A. 1994. Spatial and temporal variability on
crystallisation of celladonites within the Troodos ophiolite, Cyprus: Implications for lowtemperature alteration of the oceanic crust. Journal of Geophysical Research 99, 31473161.
LIN, J. & PARMENTIER, E.M. 1989. Mechanisms of lithosphere extension at mid-ocean
ridges. Geophysical Journal International 96, 1-22.
MARSH, B.D. 1989. Magma chambers. Annual Reviews of Earth and Planetary Sciences
17, 439-474.
NICOLAS, A. 1989. Structures of ophiolites and dynamics of ocean lithosphere. Kluwer,
Amsterdam, 367pp.
NICOLAS, A., FREYDIER, Cl., GODARD, M. & VAUCHEZ, A. 1993. Magma chambers at
oceanic ridges: How large? Geology 21, 53-56.
RONA, P.A. 1985. Hydrothermal mineralization at slow spreading centers: Red Sea,
Atlantic Ocean and Indian Ocean. Marine Mining 5, 117-145.
RONA, P.A. 1986. Mineral deposits from sea-floor hot springs. Scientific American XX, 6674.
SAWKINS, F.J. 1976. Metal deposits related to intracontinental hotspot and rifting
environments. Journal of Geology 84, 427-430.
SAUNDERS, A.D. & TARNEY, J. 1984. Geochemical characteristics of basaltic volcanism
within back-arc basins. In KOKELAAR, B.P. & HOWELLS, M.F. (eds) Marginal Basin
Geology. Special Publication of the Geological Society, London 16, pp.59-76.
SAUNDERS, A.D. & TARNEY, J. 1991. Back-arc basins. In: FLOYD, P.A. (ed) Oceanic
Basalts. Blackie, Glasgow, pp.219-263.
SAUNDERS, A.D., TARNEY, J., STERN, C.R. & DALZIEL, I.W.D. 1979. Geochemistry of
Mesozoic marginal basin floor igneous rocks from southern Chile. Geological Society of
America Bulletin 90, 237-258.
SINTON, J.M. & DETRICK, R.S. 1992. Mid-ocean ridge magma chambers. Journal of
Geophysical Research 97, 197-216.
SPOONER, E.T.C., CHAPMAN, H.J. & SMEWING, J.D. 1977. Strontium isotopic
contamination and oxidation during ocean floor hydrothermal metamorphism of the
ophiolitic rocks of the Troodos Complex, Cyprus. Geochimica et Cosmochimica Acta 41,
873-890.
THOMPSON, G. 1991. Metamorphic and hydrothermal processes: basalt - seawater
interactions. In: FLOYD, P.A. (ed) Oceanic Basalts. Blackie, Glasgow, pp.148-173.
WHITE, R.S. 1991. Structure of ocean crust from geophysical measurements. In: FLOYD,
P.A. (ed) Oceanic Basalts. Blackie, Glasgow, pp.30-48.
VARIATION IN Sr ISOTOPIC COMPOSITION of SEAWATER WITH
TIME: the plate tectonics connexion
You may wonder what the strontium isotopic composition of seawater has to do
with plate tectonics? Surprisingly the variation in the Sr isotopic composition of
ocean seawater with time it is turning out to be an excellent monitor of past plate
tectonic activity. But we are only just beginning to understand why. For instance
the present day seawater Sr isotopic composition (expressed as 87Sr/86Sr) is
0.709, and because ocean water is globally well mixed, all modern shells and
limestones that incorporate seawater Sr have this ratio. By measuring the Sr
isotopic ratios in limestones or shells from older geological formations it is possible
to log the Sr isotopic variations in seawater with time. The early work by Peterman
et al. (1970) and others produced the following curve:
It can be seen that there is a progressive increase from the late Jurassic to the
present day. However there is also a decrease from the early Carboniferous to the
late Jurassic. So there must be some geological process or processes which
produce a decrease in 87Sr/86Sr as well as those which produce an increase. So
what are these processes?
Rubidium-87 (87Rb) is radioactive and decays to 87Sr, so that the ratio 87Sr/86Sr
must increase in the Earth with time (86Sr is unradiogenic so stays constant).
However the rate at which the 87Sr/86Sr ratio will increase depends on the
elemental Rb/Sr ratio of the rock. The Earth's mantle is low in Rb relative to Sr, so
mantle derived rocks, which have a very low Rb/Sr ratio, tend to have low 87Sr/86Sr
(the ratio has increased from only 0.699 to 0.703 over the last 4500 m.y.!).
However, crustal rocks such as granites and shales are rich in Rb and have a high
Rb/Sr ratio so, given time, become enriched in 87Sr and have a high 87Sr/86Sr. So if
seawater interchanges chemically with crustal rocks its ratio of 87Sr/86Sr will
increase, wheareas if it interchanges with mantle rocks this ratio will be pulled
down again. Clearly to change the ratios in a reservoir the size of an ocean must
indicate (at least) two very big process, one pushing the ratio up and the other
dragging it down. What are they?
Significance of Increasing Sr Isotope Ratios
Clearly, enhanced erosion of high 87Sr/86Sr continental material will drive the ratio
up. As feldspars and micas weather and breakdown to clays their radiogenic Sr is
released. This is very soluble in river water and finishes up in the sea. But
enhancement can only come if the proportion of mountain belts is increased. And
this will be brought about by continental collision at the end of the Wilson Cycle.
So does the sharply increasing curve since the late Jurassic then monitor the
development of our recent mountain belts such as the Alps, Himalayas and
Andes? There have been a number of attempts to model this most recent 100 m.y.
growth, since the seawater variation curve for this period is very well known (after
Richter et al. 1992):
The Sr flux into the oceans can be estimated from the large rivers, e.g.,
River
Sr Flux (mol/yr)
87Sr/86Sr
Amazon
2.2 × 109
0.7109
Orinoco
0.2 × 109
0.7183
Himalayan-Rivers total
7.7 × 109
0.7127
Global Total
3.3 × 1010
0.711
The growth is most rapid between 40 m.y. to the present, and particularly in the
period 20 - 15 m.y. Richter et al. (1992) find that they can model this rapid growth
mainly as a consequence of uplift of the Himalayan-Tibetan plateau, following the
collision of India with Asia. See the high values for the Sr flux in the Himalayan
Rivers (Indus, Ganges, Irrawady, Yangtze, Mekong, etc.). So can we detect
mountain belt formation in the geological record by rapid increases in the Sr
isotopic ratio of seawater?
This poses an interesting problem as to why, if mountain belts are repeatedly
formed throughout history, the Sr ratio does not just go on up and up. After all, if
the seawater ratio can rise by 0.002 in the last 40 m.y., why has it risen by only
0.010 in the last 4500 m.y. Some equally powerful process must be bringing it
down again.
The curve for the whole of the Phanerozoic (Cambrian to present) was established
by Burke et al. (1982) by very careful work on limestones of various ages, taking
care to avoid rocks that experienced later diagenetic effects:
It can be seen that we actually have to go back to the Cambrian before the
seawater Sr ratio was as high as it is at the present day. In fact there has been a
general decline for a period of almost 400 m.y. between the Cambrian and the late
Jurassic, and superimposed on this is at least five sharp falls.
Significance of Decreasing Sr Isotope Ratios
The main reason for the decrease must be hydrothermal exchange of seawater
with hot basalt at mid-ocean spreading centres. As we have seen from ophiolites,
the breakdown of feldspars with low "mantle" Sr ratios releases Sr into the
seawater. At the same time, zeolites and clays growing in the low-temperature
altered basalts take Sr out of seawater with a high "continental" component.
Because ocean crust is eventually subducted at trenches, the net effect is to
remove some of this "continental" Sr component located in the hyrothermallyaltered ocean crust deep into the Earth's mantle.
So can we correlate these periods of rapidly falling Sr isotope composition with
periods of faster spreading? Or with the breakup of major continents like
Gondwanaland? Note that if there is enhanced ridge activity, this probably means
that the uprising mantle is hotter and less dense, and that most likely it will
displace ocean water volume, like Iceland, and flood the continental shelves. Of
course the more the continents are flooded, the less active will be the rivers, and
the lower their content of "continental" Sr. But, in contrast, once there is a major
continental collision, like that following the closure of Tethys and the formation of
the Alps and the Himalayas, then ridge activity must stop, or slow, for a while until
new plate configurations are established. So is this why the curve oscillates
around. It is advisable to try to model it. This is what Richter et al. (1992) have
tried to do.
Modelling the Phanerozoic Curve
Richter et al. (1992) used the following numbers:
1.25 × 1017 (mol)
0.7092
River flux:
3.3 × 1010 (mol/yr)
0.711
Hydrothermal flux:
0.82 × 1010(mol/yr)
0.7030
Diagenetic flux:
0.3 × 1010 (mol/yr)
0.7084
Total Sr in Oceans:
The latter flux is that due to carbonate diagenesis, but is not an important
controlling factor.
They then tried to model the Sr isotopic variation throughout the Phanerozoic by
making various assumptions about plate motions, collisions, spreading rates, etc.
Obviously things get more uncertain the further back in time, and because there
are parts of the Earth where the geology is not very well known. The result is
shown below:
The model accounts reasonably well for the large-scale structure of the sewater Sr
isotopic curve, but fails to reproduce several of the local maxima and minima,
especially in the period 300 - 100 m.y. However the "highs" in the Cambrian,
Devonian and present day do correlate with extensive mountain building.
The interesting point about such graphs is that they tell us that there may be
something missing from our current plate tectonic models. For instance, Richter et
al. assume that all mountain building is due to continent collision. But there was no
major continent collision as such involved in the uprise of the Andes, which is also
a major contributor to seawater Sr. The Andean uplift may have resulted in part
from the docking and attempted subduction of major ocean plateaus. This would
add a new dimension to the story because the formation of hot ocean plateaus
would substantially enhance the ridge hydrothermal Sr flux, to be followed later by
an enhanced "continental" flux as mountain belts were formed as these hot thick
plateaus tried to subduct.
For further discussion of long-term changes in geological features, and possible
implications, see Moores (1993).
References
BURKE, W.H., DENISON, R.E., HETHERINGTON, E.A., KOEPNICK, R.B., NELSON, H.F.
& OTTO, J.B. 1982. Variation of seawater 87Sr/86Sr throughout Phanerozoic time. Geology
10, 516-519.
HESS, J., BENDER, J.L. & SCHILLING, J.-G. 1986. Evolution of the ratio of strontium 87
to strontium 86 in seawater from the Cretaceous to the Present. Science 231, 979-984.
MOORES, E.M. 1993. Neoproterozoic oceanic crustal thinning, emergence of continents,
and origin of the Phanerozoic ecosystem: A model. Geology 21, 5-8.
PALMER, M.R. & EDMOND, J.M. 1989. The strontium isotope budget of the modern
ocean. Earth and Planetary Science Letters 92, 11-26.
PETERMAN, Z.E., HEDGE, C.E. & TOURTELOT, H.A. 1970. isotopic composition of
strontium in seawater throughout Phanerozoic time. Geochimica et Cosmochimica Acta 34,
104-120.
RICHTER, F.M., ROWLEY, D.B. & DePAOLO, D.J. 1992. Sr isotope evolution of seawater:
the role of tectonics. Earth and Planetary Science Letters 109, 11-23.
VEIZER, J. 1985. Strontium isotopes in seawater through time. Annual Reviews of Earth
and Planetary Sciences 17, 141-156.
PLATE TECTONICS: Lecture 5
SUBDUCTION ZONES and ISLAND ARCS
Subduction Zones are where cool lithospheric plates sink back into the mantle. It
takes about 50 my for the ocean lithosphere that formed in the hot (>1000°C)
environment at mid-ocean ridges to cool to an equilibrium state and sink to its
maximum depth below sea-level. Although there is no universal agreement on the
balance of forces that drives plate tectonics, the "slab-pull" force is thought to be
an important one. For instance the Pacific Plate is the fastest moving plate (ca. 10
cm/yr), and this is the plate that supplies most of the Earth's subducting
lithosphere, and thus where the overall slab-pull force will be the larger. The
normal argument is that the cool ocean crust will more easily convert to dense
eclogite which, as we have seen in Lecture 1, is much more dense than pyrolite.
What is most surprising is the great variation in geological features associated with
subduction. There is a huge difference between the East Pacific and the West
Pacific. Not only that, but there are differences along the Andean margin, and also
quite major differences as we go back in time. But it is important to understand
subduction because this is where the continental crust grew progressively with
time.
Subduction is where tectonics, structural geology, sedimentation, igneous
petrology, metamorphism, geochemistry, geophysics and applied geology all
interact. Typical "textbook" features of a mature continental margin subduction
zone are shown below. The cartoon shows sediment being scraped off the
downgoing plate to form an accretionary wedge, and that a forearc basin is
forming on top of the wedge as it is dragged down (and is presumably fed by
volcanic debris from the arc). However, the cartoon avoids the issue of how and
where the volcanic magmas come from. To what extent does the basaltic
subducted slab contribute to arc magmas? Is it just the fluids carried down in
altered oceanic crust that migrate into the mantle wedge overlying the subduction
zone and cause melting? Ot what extent do sediments carried down the
subduction zone then contribute to arc magmas? Why are arc volcanoes nearly
always situated about 110 km above the Benioff Zone? What happens to material
taken down the subduction zone?
MARGINAL BASINS & BACK ARC SPREADING
Marginal basins are a common feature of the Western Pacific. Examples (north to
south) are the Sea of Japan, the West Philippine Basin, the Parace Vela &
Shikoku Basins, the Mariana Trough, the Woodlark Basin, the Fiji and Lau Basins.
By contrast marginal basins are rarer in the Eastern Pacific. The two examples in
the Atlantic are the Caribbean and the Scotia Sea.
Marginal basins are small oceanic basins, usually adjacent or "marginal" to a
continent, which are separated from larger oceans by an island arc. Some
marginal basins at continental margins may be imperfectly developed and
represented by thinned crust, often associated with basic volcanism. Karig (1971,
1974) divided marginal basins into:
(1) Active marginal basins with high heat flow.
(2) Inactive marginal basins with high heat flow.
(3) Inactive marginal basins with normal heat flow.
The first two are thought to have formed by back-arc spreading, either still active
(1), or recently active (2). The third may represent basins formed by even older
back-arc spreading, or normal ocean crust that has been "trapped" behind a
recently developed oceanic island arc.
FRAMEWORK OF AN ISLAND ARC SYSTEM
The commonly held model of an arc - back-arc system has the following
components:
(1) Subduction Zone
(2) Fore-arc region with accretionary sedimentary prism
(3) Frontal Arc
(4) Active Arc
(5) Marginal Basin with spreading centre
(6) Remnant Arc
(7) Inactive Marginal Basin
Although the extensive fore-arc region of many island arcs was thought to be
composed of off-scraped sediments, drilling has not substantiated this. It appears
that - at least at intraoceanic arcs - abyssal sediments on the downgoing plate are
largely subducted.
That the back-arc region is a zone of asthenospheric upwelling is supported by
seismic evidence which suggests a low-Q (seismic attenuation) zone behind the
arc, compatible with a small amount of melt in the back-arc region:
Magnetic anomalies in back-arc basins are not so well developed, nor have such
symmetrical linear patterns, as those in the normal ocean basins. There have
been difficulties in identifying the anomalies. It has been suggested by Lawver &
Hawkins (1978) that spreading may be more diffuse and not constrained to one
central well-defined spreading centre. Good dateable magnetic anomaly patterns
were first described from the Scotia Sea back-arc basin (IA Hill). Spreading in
some basins may be asymmetric, with accretion favoured on the active arc side.
Models for Back-arc Spreading (see Karig, 1974)
Active Diapirism: One of the earliest models, based on the Mariana Arc System,
is that of an uprising diapir splitting the arc. The diapir is initiated either as a result
of frictional heating at the subduction zone, or more likely through fluids released
from the dehydrating subducting slab. The rising diapir then splits the arc in two
and the two halves are progressively separated by seafloor spreading:
Passive Diapirism: This results from regional extensional stresses in the the
lithosphere across the arc system. In effect the downgoing slab, although acting
like a conveyor belt, also has a vertical component that causes "roll-back". The arc
and forearc then stays with the subduction zone, as a result of a supposed trench
suction force:
Stepwise Migration: Here it is assumed that the subducting slab is snapped off
near the hinge, presumably because something on the downgoing slab is too light
to go down, and so a new subduction is initiated oceanwards. The arc stays near
the hinge and the asthenosphere wells up behind it:
Convection-driven: This model proposed by Toksoz & Bird (1978), and requires
that subsidiary convection cells are driven by the downward drag of the downgoing
slab. Calculations suggest that spreading would occur about 10 my after the start
of subduction. This might explain why back-arc spreading is more common in
oceanic regions ™ the lithosphere is thinner and thus more easily disrupted than
under continents:
Uprising Harzburgite Diapir: This model (Oxburgh & Parmentier 1978) depends
on the fact that refractory lithosphere (which has lost its basalt component at midocean ridges) is less dense and inherently more buoyant than normal fertile
mantle. Thus it would rise if heated to same temperature as surrounding mantle.
Such diapirs could in theory be derived from subducting lithosphere, although it is
doubtful that subducting lithosphere could be heated within 10 my; more likely it
takes 1000 - 2000 my according to megalith concepts of Ringwood (1982):
Old and Young Lithosphere: Molnar & Atwater (1978) have argued that it
depends on the dip of the subducting slab whether extension occurs in the back
arc region. In the W. Pacific it is old (Jurassic), cold and dense lithosphere that is
subducting - with very steep dip and strong vertical component. Thus extensional
conditions in back-arc region. In the E. Pacific, on the other hand, the lithosphere
subducting beneath the Andes is young (Tertiary), warm and less dense, and
subducts at a shallow angle. Thus convergence is more compressive than
extensional. Uyeda & Kanamori (1979) have characterised these two extreme
types of subduction as Mariana and Chilean type respectively. See also Dewey
(1981)
Other models: Various researchers have since commented on the possible
causes of back-arc spreading, including assessments of dependence on absolute
and relative plate motions. Consult some of references listed below. Experimental
laboratory studies have been carried out by Kincaid & Olsen (1987), observing the
effects of continued subduction where the subducting slab 'hits' the 650 km
discontinuity. The results show that steep subduction does produce a significant
roll-back effect on the hinge, which will generate extensional conditions in the
back-arc region. Note that with subduction rates of about 7 cm/yr it would take
about 10 my before newly subducted ocean lithosphere would 'hit' the 650 km
discontinuity and begin to initiate 'roll-back' of the hinge, and thus extensional
conditions.
EVOLUTION OF MARIANA ARC SYSTEM
The Mariana Arc is perhaps the type intra-oceanic arc system, and the most
extensively studied through marine geophysical studies, dredging and drilling
(particularly Legs 58, 59 and 60 of DSDP in late 1970's). From west to east it
consists of the following features:
(1) West Philippine Basin
(2) Kyushu-Palau Ridge (a remnant arc)
(3) Shikoku & Parece-Vela Basins
(4) West Mariana Ridge (a remnant arc)
(5) Mariana Trough
(6) Active Mariana Arc
(7) Mariana Fore-arc (made of old arc)
(8) Mariana Trench (up to 11 km deep)
(9) The subducting Pacific Plate (Jurassic age)
West Philippine Basin: This may be 'trapped' in origin and not strictly formed by
back-arc spreading. It appears to pre-date the Kyushu-Palau Ridge. Magnetic
anomalies suggest active spreading in the early Tertiary (62-40 Ma) with the NWSE trending Central Basin Fault as the spreading centre. The Oki-Daito Ridge in
the northern West Philippine Sea is aligned parallel to this feature and has been
regarded as an old remnant arc: however drilled samples from the Oki-Daito Ridge
are alkaline basalts, not island arc basalts. Drilled samples from the W. Philippine
Basin are fairly typical MORB.
The Philippine Basin is slowly subducting to the west beneath Taiwan, etc. The
subduction rate is much less than that of the Pacific Plate beneath the Marianas.
Kyushu-Palau Ridge: This is over 2000 km long and rises 2 km above the
adjacent basin floors. Consists of vesicular lava flows, dykes and sills, interbedded
with volcaniclastic breccias lying below Middle Oligocene oozes. Lavas all belong
to Island Arc Tholeiite (IAT) Series, typical of the most primitive island arcs. Now
an inactive Remnant Arc that was active between about 42 and 32 my ago.
Parece-Vela and Shikoku Basins: Magnetic anomaly patterns indicate back-arc
spreading between 30 and 17 my in Parece-Vela and between 26 and 15 my in
the Shikoku Basin in north. Basaltic sills common in sediments near basement,
indicating high rates of sedimentation near near ridge axis. Basalts are vesicular.
Similar to MORB.
West Mariana Ridge: Shallower and younger than the Kyushu-Palau Ridge.
Drilling penetrated about 1000 m of volcaniclastic material composed of basalts,
basaltic andesites, rare andesites and plagioclase phenocrysts. Their character is
calc-alkaline, with much higher contents of Ba and Sr than those of K-P Ridge. Arc
was active 17-8 my ago. So now a Remnant Arc. Arc built up when spreading in PV / Shikoku Basins ceased.
Mariana Trough: This is 1500 km long, 250 km wide. Rough topography, high
heat flow. Magnetic lineations poorly developed, but suggest back arc spreading
from about 6 my ago - i.e. when activity on West Mariana Ridge ceased. Near the
West Mariana Ridge metabasalts, gabbros and anorthositic cumulates were drilled
- deeper part of a rifted-apart arc? Basalts in Mariana Trough are MORB-like, but
have some arc characteristics. Vesicular. Spreading still in progress. Further north,
on Iwo-Jima Ridge, there is an incipient back-arc basin just beginning to form - the
Bonin Trough.
Mariana Active Arc: This consists of numerous small islands and seamounts, on
the eastern edge of the extensive Fore-arc region. Lavas are mainly basalts,
basaltic andesites and andesites.
Mariana Fore-arc: The forearc region shows a history of continual subsidence. The
basement is Eocene in age (similar to Kyushu-Palau Ridge) and consists of two
distinct lava types:
(1) Island Arc Tholeiites (very similar in character to those of KyushuPalau Ridge). These magmas can normally be easily distinguished from
calc-alkaline basalts from more mature arc systems.
(2) Boninites, or high-magnesian andesites. These are unusual lavas,
combining high Si with high Mg, Ni and Cr. They are thought to have
formed by wet-melting of rather refractory lithosphere.
(3) Dacites also occur on Guam.
Drilling and dredging in the trench area of the fore-arc has recovered mainly
volcanic materials. No scraped-off sediments from the oceanic plate - with the
implication that all sediment is being subducted, and that the fore-arc itself is
suffering tectonic erosion as a result of the rasping action of the downgoing slab.
TECTONIC EVOLUTION OF MARIANA ARC SYSTEM
Combining evidence from magnetic anomalies, drilling, dredging and
geochronology, the geologic history of the arc system can be pieced together. In
the period immediately preceding the development of the arc, the plate
configuration in the eastern Indian Ocean and western Pacific was dominated by
the rapid movement of India northward. There were some major N - S oriented
transform faults at this time, so about 60 Ma ago the plate tectonic configuration
probably looked like this:
India was just about to collide into Asia to form the Himalayas, Australia had just
begun to separate from Antarctica, and note the very large ridge offsets on the NS transforms. The critical point at this time was that slab-pull associated with the
rapidly-moving Indian Plate will stop as soon as India collides. Similarly, the
spreading ridge in the NE Pacific is going to push itself under the Aleutians, when
upon the slab-pull will also stop. This leaves the northerly pull forces on the Pacific
plate very weak, and very vulnerable to change in plate motion direction. So about
40 my ago the Pacific Plate changed motion from northwards to westward (c.f.
kink in Hawaiian-Emperor seamount chain). The sequence of events can be
tracked as follows:
(1) The Kyushu-Palau Ridge is thought to mark the position of one of these major
transform faults, with younger, warmer and thinner ocean ocean lithosphere to the
west, and older, cooler and denser lithosphere to east. Drawn to scale, the
position immediately before the change in plate motion probably looked like this:
It can easily be envisaged how the eastern side would easily subduct under the
new young warm lithosphere to the west that had recently formed at a spreading
ridge. After the change in plate motion direction, the map then looked like:
A new volcanic arc forms at the site of the easternmost transform, and many
complications develop in SE Asia (Philippines, etc.) because of transforms turning
into arcs, and various subduction-flips as thick (plateau-type) ocean crust refuses
to subduct. A new subduction zone develops north of Australia.
(2) Rapid build-up of Kyushu-Palau Arc in late Eocene – Oligocene through
voluminous eruption of island arc tholeiites and high-Mg boninites. Activity
continued for ca 10 my. So what happened to bring about such a rapid rate of
magma production. It is possible that the earliest stages of subduction looked as
follows:
Note that the downgoing plate not only has "conveyor-belt" motion, but also a
strong vertical component so that it is sinking into the mantle. At this point hot
asthenosphere mantle rushes in to replace it. So in a rather unique rapidly
extensional tectonic environment, wet altered ocean crust is juxtaposed next to
very hot asthenospheric mantle. With an abundance of heat and water, it is not
surprising that huge amounts of magma are generated. This tectonic situation is
actually even more extensional than at a mid-ocean ridge, so it may be expected
that all the features of a "type" mid-ocean ridge are reproduced: pillow lavas,
sheeted dykes, gabbros, etc. This is shown below:
(to come)
(3) Splitting of K-P Arc in half about 30 my ago with formation of PareceVela & Shikoku Basins by back-arc spreading. Spreading stopped about 16
my ago.
(4) Formation of West Mariana Arc between about 17 and 8 my ago through
eruption of calc-alkaline basalts and basaltic andesites.
(5) Splitting of West Mariana Arc abut 6 my ago to form Mariana Trough by
back-arc spreading, and leaving West Mariana Ridge as remnant arc.
(6) Formation of new Mariana Arc 5 my ago to present. Now erupting lavas
with mixed calc-alkaline - island arc tholeiite characteristics.
Presumably the Mariana Arc will continue migrating eastwards into the Pacific.
Magma Compositions
Arc Magmatism
The magmas erupted at the Mariana Arc show a gradual evolution in composition
with time. Note that the whole arc system has evolved entirely within the oceanic
regime (no continental crust or sub-continental lithosphere involved).
The earliest lavas erupted (now seen on Kyushu-Palau Ridge and Mariana Forearc) are island arc tholeiites (IAT) and boninites. These are characteristic of very
primitive oceanic island arcs, and are not usually erupted on continents or in the
later stages of arc development. IAT have similarities with mid-ocean ridge basalts
(MORB), in having depleted rare-earth element (REE) patterns, but are usually
more Fe-rich and with low Cr and Ni contents, very low Nb and Ta, higher K
contents and high K/Rb ratios. Boninites are high-Mg lavas, but have high silica
contents more typical of andesites; they have high Cr and Ni contents, but have
lower Ti contents and higher K, Rb, Ba and Sr contents than would normally be
expected of high-Mg rocks.
Boninites are thought to result from wet melting of the rather refractory Mg-rich
mantle wedge beneath the developing arc - with the wedge being contaminated
with elements such as K, Rb, Ba, Sr transported from the subduction zone during
dehydration of the hydrous ocean crust.
IAT could be melts of the more fertile asthenosphere, the magmas then
undergoing extensive crystal fractionation en route to the surface. Or they could
represent melts of subducted ocean basalt crust (only possible at the very start of
subduction when the ocean lithosphere is pushed down into hot mantle).
After opening of the Parece Vela basin by back-arc spreading, arc volcanic activity
was transferred 17 my ago to the what is now the West Mariana Ridge, and
continued building up that arc for ca. 9 my. The lavas erupted however were
mainly calc-alkaline basalts (CAB) and basaltic andesites, with higher Al contents,
much higher Sr and Ba contents and light rare-earth enriched rather than depleted
REE patterns. These lavas are more similar to calc-alkaline lavas erupted at
continental margins (though the latter are usually dominated by andesite rather
than basaltic andesites).
These CAB magmas may have been derived from the mantle wedge. But if so
there is an implication that the wedge may have been enriched in Ba, Sr, light
REE, etc., perhaps as a result of continued fluid transport of these elements into
the wedge from the dehydrating subducting slab.
Modern lavas erupted at the active Mariana Arc tend to be mainly andesites and
basaltic andesites having characteristics in between those of IAT and CAB. There
is some evidence that a small component (ca. 0.5%) of subducted abyssal
sediment is involved in their source regions.
Perhaps the most interesting aspect of the Mariana arc is that at least three
distinct magma types appear to have been generated from the one subduction
zone. Yet the whole arc system evolved entirely within the oceanic environment.
Back-arc Basalts
In many respects marginal basin basalts (MBB) are similar to normal mid-ocean
ridge basalts (N-type MORB). However during the early stages of back-arc
spreading, when the uprising mantle diapir splits the volcanic arc, the basalt
magmas are derived from the sub-arc mantle. These basalts tend to have an arclike geochemical signature. Thus their REE patterns may be slightly light REE
enriched, they have higher Ba, Sr, K and Rb, but low Nb and Ta. Moreover they
tend to have higher water contents and be vesicular - a consequence of fluids
distilled from the subducting slab. These features are useful discriminants in trying
to characterise ophiolites as being derived from either obducted ocean floor or
marginal basin crust. See Saunders & Tarney (1984; 1991) for summary.
Addition: Schematic cross-section across the Mariana Arc showing the
components involved in magma generation.
Fluids are released from the sub-ducting slab as "wet" amphibolite recrystallises at
ca. 100km depth to dry dense eclogite. These fluids migrate upwards into the
mantle wedge and induce melting of the sub-arc lithosphere. (The more water, the
more melting, and higher the magma production?). However, this mantle varies in
it's fertility because of previous metasomatic events affecting the deeper
lithosphere.
More active mantle diapirism occurs in the back-arc region, and this results in
much more melting and active spreading. Hydrous fluids are still involved in these
mamgas, but to a lesser extent than in the arc rocks.
WHAT CAUSED THE CHANGE IN PACIFIC PLATE MOTION THAT
PRODUCED THE MARIANA ARC?
If we bear in mind that plate motions are dominantly controlled by 'slab pull', then
anything which reduces the slab-pull force will encourage changes in the direction
and speed of plate motion. It is notable that in the southeastern Pacific the Aluk
Ridge (spreading centre) began to progressively subduct along the Antarctic
Peninsula; at the same time, the northwestern Pacific the Kula Ridge began to
subduct beneath the Aleutians - Kamchatka. A result was a marked reduction in
the N™S slab-pull, because recently formed hot lithosphere is not very dense and
not keen to subduct. In combination with other plate re-configuring events
worldwide, this may have been enough to cause switch in Pacific Plate motion
from N – S to E – W. But see Richards et al. (1996)
REFERENCES: Arcs and Marginal Basins
The references below lead to most aspects of interest to island arcs, even if you
just look at the abstracts & diagrams!
BLOOMER, S.H. 1987. Geochemical characteristics of boninite- and tholeiite-series
volcanic rocks from the Mariana forearc and the role of an incompatible element-enriched
fluid in arc petrogenesis. Geological Society of America, Special Paper 215, 151-164.
CARLSON, R.L., HILDE, T.W.C. & UYEDA, S. 1983. The driving mechanism of plate
tectonics: relation to age of the lithosphere at trenches. Geophysics Research Letters 10,
297-300.
CHASE, C.G. 1978. Extension behind island arcs and motions relative to hot spots.
Journal of Geophysical Research 83, 5385-5387.
CHASE. C.G. 1979. Asthenospheric counterflow: a kinematic model. Geophysical Journal
of the Royal Astronomical Society 56, 1-18.
CRAWFORD, A.J., BECCALUVA, L. & SERRI, G. 1981. Tectono-magmatic evolution of
the West Philippine-Mariana region and the origin of boninites. Earth and Planetary
Science Letters 54, 346-356.
DAVIES, J.H. & STEVENSON, D.J. 1992. Physical model of source region of subduction
zone magmatism. Journal of Geophysical Research 97, 2037-2070.
GARFUNKEL, Z., ANDERSON, C.A. & SCHUBERT, G. 1986. Mantle circulation and the
lateral migration of subducted slabs. Journal of Geophysical Research 91, 7205-7223.
HAMILTON, W.B. 1988. Plate tectonics and island arcs. Geological Society of America
Bulletin 100, 1503-1527.
HASTON, R. & FULLER, M. 1991. Palaeomagnetic data from the Philippine Sea plate and
their significance. Journal of Geophysical Research 96, 6073-6098.
HAWKINS, J.W., BLOOMER, S.H., EVANS, C.A. & MELCHIOR, J.T. 1984. Evolution of
intra-oceanic arc-trench systems. Tectonophysics 102, 174-205.
HICKEY, R.L. & FREY, F.A. 1982. Geochemical characteristics of boninite series
volcanics: implications for their source. Geochimica et Cosmochimica Acta 46, 2099-2115.
HILDE, T.W., UYEDA, S. & KROENKE, L. 1977. Evolution of the western Pacific and its
margin. Tectonophysics 38, 145-167.
HOLE, M. J., SAUNDERS, A. D., MARRINER, G. F. & TARNEY, J. 1984. Subduction of
pelagic sediment: implications for the origin of Ce-anomalous basalts from the Mariana
Islands. Journal of the Geological Society, London 141, 453-472.
HSUI, A.T., MARSH, B.D. & TOKSOZ, M.N. 1983. On melting of the subducted ocean
crust: effects of subduction induced mantle flow. Tectonophysics 99, 207-220.
IDA, Y. 1983. Convection in the mantle wedge above the slab and tectonic processes in
subduction zones. Journal of Geophysical Research 88, 7449-7456.
JURDY, D.M. 1979. Relative plate motions and the formation of marginal basins. Journal of
Geophysical Research 84, 6796-6802.
JURDY, D.M. & STEFANICK, M. 1983. Flow models for back-arc spreading.
Tectonophysics 99, 191-200.
KARIG, D.E. 1974. Evolution of arc systems in the Western Pacific. Annual Reviews of
Earth and Planetary Sciences 2, 51-78.
KARIG, D.E. 1971. Structural history of the Mariana island arc system. Geological Society
of America Bulletin 82, 323-344.
KARIG, D.E. 1971. Origin and development of marginal basins in the Western Pacific.
Journal of Geophysical Research 76, 2542-2561.
KARIG, D.E. 1982. Initiation of subduction zones - Implications for arc evolution and
ophiolite development. Geological Society of London, Special Publication 10, 563-576.
KINCAID, C. & OLSON, P. 1987. An experimental study of subduction and slab migration.
Journal of Geophysical Research 92, 13832-13840.
KUSHIRO, I. 1990. Partial melting of mantle wedge and evolution of island arc crust.
Journal of Geophysical Research 95, 15929-15939.
LAWVER, L.A. & HAWKINS, J.W. 1978. Diffuse magnetic anomalies in marginal basins:
their possible tectonic and petrologic significance. Tectonophysics 45, 323-339.
MARSH, B.D. 1979. Island arc development: some observations, experiments and
speculations. Journal of Geology 87, 687-713.
MOLNAR, P. & ATWATER, T. 1978. Interarc spreading and cordilleran tectonics as
alternates related to the age of subducted ocean lithosphere. Earth and Planetary Science
Letters 41, 330-340.
MUELLER, S. & PHILLIPS, R.J. 1991. On the initiation of subduction. Journal of
Geophysical Research 96, 651-665.
NATLAND, J.H. & TARNEY, J. 1982. Petrological evolution of the Mariana Arc and Backarc Basin System: a synthesis of drilling results in the South Philippine Sea. Initial Reports
of the Deep Sea Drilling Project 60, 877-908 (Washington: U.S. Government Printing
Office).
PEACOCK, S. M. 1990. Fluid processes in subduction zones. Science 248, 329-337.
RICHARDS, M.A. & LITHGOW-BERTELLONI, C. 1996. Plate motion changes, the
Hawaiian™Emperor bend, and the apparent success and failure of geodynamic models.
Earth and Planetary Science Letters 137, 19-27.
RINGWOOD, A.E. 1974. The petrological evolution of island arc systems. Journal of the
Geological Society, London 130, 183-204.
SAUNDERS, A.D. & TARNEY, J. 1984. Geochemical characteristics of basaltic volcanism
within back-arc basins. In KOKELAAR, B.P. & HOWELLS, M.F. (eds) Marginal Basin
Geology. Geological Society of London, Special Publication 16, 59-76.
SAUNDERS, A.D. & TARNEY, J. 1991. Back-arc basalts. In FLOYD, P.A. (ed) Oceanic
Basalts. Blackie, Glasgow, pp. 219-263.
SHEMENDA, A.I. 1993. Subduction of the lithosphere and back arc dynamics: insights
from physical modeling. Journal of Geophysical Research 98, 16167-16185.
SPENCE, W. 1987. Slab pull and the seismotectonics of subducting lithosphere. Reviews
of Geophysics 25, 55-69.
STERN, R.J. & BLOOMER, S.H. 1992. Subduction-zone infancy - Examples from the
Eocene Izu-Bonin-Mariana and Jurassic California arcs. Geological Society of America
Bulletin 104, 1621-1636.
STERN, R.J., BLOOMER, S.H., LIN, P.-N. & SMOOT, N.C. 1989. Submarine arc
volcanism in the southern Mariana arc as an ophiolite analogue. Tectonophysics 168, 151170.
TARNEY, J., SAUNDERS, A.D. & WEAVER, S.D. 1977. Geochemistry of volcanic rocks
from the island arcs and marginal basins of the Scotia Arc region. In: TALWANI, M. &
PITMAN, W.C. (eds) Island Arcs, Deep Sea Trenches and Back-arc Basins. American
Geophysical Union, Maurice Ewing Series 1, 367-378.
TARNEY, J., SAUNDERS, A. D., MATTEY, D. P., WOOD, D. A. & MARSH, N. G. 1981.
Geochemical aspects of back-arc spreading in the Scotia Sea and Western Pacific.
Philosophical Transactions of the Royal Society of London A300, 263-285.
TARNEY, J., PICKERING, K.T., KNIPE, R.J. & DEWEY, J.F. 1991. Fluids and subduction
zone processes. In TARNEY, J., PICKERING, K.T., KNIPE, R.J. & DEWEY, J.F. (eds)
Behaviour and Influence of Fluids in Subduction Zones. The Royal Society, London. (i-vi)
TATSUMI, Y., MURASAKI, M. & NOHDA, S. 1992. Across-arc variation of lava chemistry
in the Izu-Bonin Arc: identification of subduction components. Journal of Volcanology and
Geothermal Research 49, 179-190.
TAYLOR, B. & KARNER, G.D. 1983. On the evolution of marginal basins. Reviews of
Geophysics 21, 1721-1741.
TOKSOZ, N. & BIRD, P. 1977. Formation and evolution of marginal basins and continental
plateaus.In TALWANI, M. & PITMAN, W.C. (eds) Island Arcs, Deep Sea Trenches and
Back-arc Basins. American Geophysical Union, Maurice Ewing Series 1, 379-393.
UYEDA, S. & KANAMORI, H. 1979. Back-arc opening and mode of subduction. Journal of
Geophysical Research 84, 1049-1061.
WYLLIE, P.J. 1988. Magma genesis, plate tectonics and chemical differentiation of the
earth. Reviews of Geophysics 26, 370-404.
ZHAO, D., HASEGAWA, A. & HORIUCHI, S. 1992. Tomographic imaging of P and S wave
velocity structure beneath northeastern Japan. Journal of Geophysical Research 97,
19909-19928.
TECTONICS OF SUBDUCTION ZONES
Contrasts between West & East Pacific
Uyeda & Kanamori (1979) emphasised that there were two contrasting types of
subduction zone: Mariana Type and Chilean Type - with of course many
intermediate types. The Mariana Type is characterised by a very steeply dipping
slab; the Chilean Type by a shallow-dipping slab. These differences were further
amplified by Dewey (1981).
Mariana Type has:
1. Deep open trench (up to 11 km deep) that subducts old cold Jurassic crust.
2. A very steep Benioff Zone
3. Extensive faulting, subsidence and tectonic erosion of the outer trench wall.
4. Widespread intra-arc extension and back-arc spreading.
5. More earthquakes in the under-riding than in the over-riding plate.
6. A rather thin mafic-intermediate composition volcanic-plutonic crust.
7. Extensive volcanism; mainly basaltic with only minor andesites.
8. Little or no sedimentary accretion at the trench.
9. Subdued morphological expression.
10 Lavas have quiet eruptive style.
11 Volcanoes are mainly submerged cones with fringing reefs.
12 Poorly developed volcaniclastic dispersal fans.
Chilean Type has:
1. Shallower trench (up to 6 km) that subducts younger, warmer, Eocene age
oceanic crust.
2. Thrust faulting common on outer trench wall.
3. Major thrust faulting in the under-riding Nazca Plate up to 200 km west of the
trench.
4. A Benioff Zone with a very shallow dip down to about 200 km, and then a
steeper deeper portion
below a seismic gap.
5. Widespread intra-arc compression and back-arc thrusting over a foreland
trough.
6. More, and higher energy, earthquakes in the over-riding than in the under-riding
plate.
7. Plutonism is dominant over volcanism.
8. Volcanism is dominantly of andesite-dacite-rhyolite type; basalts being much
rarer.
9. Thick (ca 70km) continental crust gradually tapering trenchward to less than 10
km.
10 Because of dominant compression, continental arc has high uplift rates.
11 Violent eruptive style. High viscosity lavas. Extensive volcaniclastic dispersal
fans.
12 Spectacular geomorphological expression.
Difference in seismic characteristics: The steep dip of the Benioff Zone in the
Mariana type means that the contact interface between the subducting slab and
the mantle wedge lithosphere is less than 100 km, hence not much frictional drag.
In any case tectonic conditions are extensional. In Chilean type however, the
shallow slab dip and greater thickness of continental lithosphere means that the
contact interface can be as much as 400 km. Hence considerable resistance and
friction and much greater seismic activity.
Tectonic Erosion and Accretion: In the Mariana Arc there is no accretion of
abyssal sediments at the trench. Yet considerable volumes of sediment are
entering the trench: sediments are 0.5km thick on Pacific Plate entering the trench,
subduction rate 10 cm/yr for ca. 40 m.y. (work out how many cubic km per unit
length of arc!). Instead forearc is undergoing tectonic erosion ("subcretion"). Most
of the sediment is being subducted - only a small proportion of it is re-cycled into
arc volcanics. Along Chilean margin the sediment supply varies: very little in north
where desert conditions, but much more in south where rainfall is high. It has been
suggested that the continental basement may be eroding by subcretion in
Northern Chile, but growing by sediment accretion in Southern Chile. Where
sediment supply is high, sediments may fill the trench and flood over on to the
oceanic plate; thus depressing it so that it approaches subduction zone at a
shallow angle.
Explanation for differences between East and West Pacific Margins
Contrast cannot be explained simply by differences in convergence rate, since
Chilean, Mariana, Japanese and Tonga arcs all have head-on convergence rate of
about 10 cm/yr. Contrast must be related to balance between "roll-back" of hinge
and convergence rate. If roll-back is faster than convergence rate then back-arc
extension results; if slower, then back-arc compression.
Roll-back may be determined by age of subducting lithosphere (Molnar & Atwater
1978). Old cold lithosphere is denser and subducts at steeper angle . . presumably
takes less time to reach 650 km discontinuity. If it cannot penetrate discontinuity
then splays back (see experiments of Kinkaid & Olsen (1987)) and induces rollback of hinge at subduction zone, giving extensional tectonics. However, with
shallower angle subduction of younger warmer lithosphere the slab will take longer
to reach 650km discontinuity, and will warm up more and become less coherent
and less able to induce roll-back effect. So no extension. An additional factor is
that in the Eastern Pacific the American Plate is over-riding the Pacific (Nazca)
Plate due to the opening of the Atlantic . . although the rate is quite small.
Wider implications: If the balance between compression and extension at
convergent plate margins is related to dip of slab (and hence age of lithosphere
subducting), then it may explain why intraoceanic island arcs are essentially a
Phanerozoic phenomenon, and become rare or absent in the middle to early
Precambrian. Higher thermal gradients in Precambrian would mean greater ridge
length and smaller plates (see Hargraves 1986), so subducting plates would be
younger and warmer, and less likely to subduct at steep angle. Hence much less
likely to induce extensional conditions at convergent plate boundaries. Is it only
when there is extension that island arcs are produced?
References
DEWEY, J.F. 1981. Episodicity, sequence and style at convergent plate boundaries. In:
The Continental Crust and its Mineral Deposits. Geological Association of Canada, Special
Paper 20, 553-572.
Drilling at sea: Hydrocarbon Exploration
Drilling at sea poses a few obvious problems. Drilling crews have to operate through the
depth of the sea water column.. Some environments, the North sea for example, have
hostile climates, inhospitable to man and corrosive to machines. To deal with these
specialized problems, drilling contractors have fashioned new tools. This machinery is
employed once suitably promising locations have been pin pointed by cheaper geophysical
methods. The operating company sets up their exploration drilling program, usually using
either a jack-up unit, a semi-submersible rig or a drill ship.
Image used with
permission from UK
Offshore Operators
Association Ltd
A jack-up unit is a
barge with legs that can
be lowered or raised.
The barge is towed to
the drilling location with
its legs in the raised
position. Once in
position, the legs are
lowered. When they
reach the sea-bed, the
barge's body is hoisted
above the water,
creating a stable drilling
platform. The length of
the legs determines the
depth of water in which a
jack-up barge can be
used. They can
generally be used in up
to 100 meters of water.
Jack-up barges are
widely employed in the
relatively shallow waters
of the North Sea's
Southern basin.
A semi-submersible drilling rig
is normally a self-propelled working
platform supported by vertical
columns on submerged pontoons.
By varying the amount of ballast
water in the pontoons, the unit can
be raised or lowered in the water.
The lower the pontoons lie beneath
the surface of the water, the less
they are affected by wave action.
This reduces vertical movement
and allows drilling to continue in
moderately rough seas.
A semi-submersible vessel is
normally held in position by up to
eight very large anchors, or by
dynamic positioning. Dynamic
positioning systems use computer
Image used with permission
from UK Offshore Operators
Association Ltd
controlled directional propellers to
keep the vessel stationary relative
to the sea-bed, compensating for
wind, wave or current.
Semi-submersibles can drill in
water depths to 300 meters or more
all year round.
Drill ships have a
broadly
conventional ship's
hull, but also
feature a large
aperture, known
as a "moon pool",
through which
drilling takes
place.
Either purpose
built, or converted
from some other
use, drill ships can
be moved easily
between locations.
They can carry
large stocks of
Images used with permission from supplies, but are
UK Offshore Operators Association
not as stable as
Ltd
semisubmersibles.
Drill ships use
either anchors or
dynamic
positioning to
maintain station.
The latest drill
ships can operate
in 1,500 meters of
water.
Almost without exception, exploration drilling
units are owned and operated by drilling
contractors who are hired by the oil companies.
The duration of drilling contracts can range from
a few months, in the case of a single well, to as
long as three years for a series of wells.
Once the drilling unit is in position, the well is
started, or "spudded", typically by drilling a 36 inch
hole in the sea-bed. To prevent the hole from caving
in, heavy steel pipe casing is cemented in to form a
solid lining.
The casing becomes progressively narrower as the
hole deepens so that, for example, a 4,000 ft well may
start with 30 inch casing and finish with 7 inch casing
(Imperial measures remain standard for most drilling
activities). Formation pressure can be unpredictable,
and is therefore potentially hazardous. If the bit
penetrates a high pressure zone unexpectedly, oil or
gas, sometimes a mixture of both, may rush up the
bore hole, producing an uncontrollable gusher at the
well head. This is known as a blow-out, and to help
reduce this danger all wells are fitted with an
emergency valve called a blow-out preventer (BOP)
A production test involves the drill crew attaching a
perforating gun to the bottom of the drill string and
lowering it to the target zone. The gun fires off
explosive charges which perforate the casing,
allowing hydrocarbons to flow into the bore hole and
up to the surface
© Copyright
J. Rochester (1998)
Drilling techniques have been constantly improved to
maximize performance and reduce costs. For
example, directional drilling is a widely used
technique that allows the drilling crew to reach
different parts of the reservoir from a single drilling
location. The crew use special fittings to deviate the
drill string to specific predetermined angles.
Horizontal drilling is now being carried out in the
North Sea. This technique is primarily used to
improve recovery from low permeability reservoirs.
Other techniques developed for North Sea exploration drilling include poly-crystalline
diamond drill bits, which are harder and faster than conventional steel bits, and turbo drills
or mud turbines which use the pressure of the mud to drive the bit, eliminating the need to
rotate the whole drill string from the surface.
Summary table of drilling methods available at sea and off land, adapted from
Evans (1995)
Types
Common use
Capacity
Common water depth
limits
Floating rigs
Drill ships
Exploration Drilling
Deep waters
4,000 feet
Exploration
Development
drilling
Semi-submersibles Exploration drilling
Bottom supported
rigs
Exploration Drilling
Jack up rigs
Exploration
Submersibles
Development
drilling
Development
drilling
Platform rigs
Production
operations
Drilling barges
Shallow waters
Swamps
Deep waters
4,000 feet
Shallow waters
350 feet
Swamps
Shallow or Deep
waters
1,000 feet
Drilling Rigs:
From Left to right, Land based oil exploration rig, Platform rig, Jack-up rig,
Semi-submersible,
Drill ship and another deeper water Semi-submersible.
Image used with permission from Offshore Technology
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