Geological Society of America
Special Paper 388
2005
From Deccan to Réunion: No trace of a mantle plume
Hetu C. Sheth
Department of Earth Sciences, Indian Institute of Technology (IIT) Bombay, Powai, Mumbai 400 076 India
ABSTRACT
The widely accepted mantle plume model postulates that (1) the currently volcanically active Réunion Island in the Indian Ocean is fed by the narrow “tail” of a
mantle plume that rises from the core-mantle boundary, (2) the Deccan continental
flood basalt province of India originated from the “head” of the same plume during
its early eruptive phase near the end of the Cretaceous, and (3) the LakshadweepChagos Ridge, an important linear volcanic ridge in the Indian Ocean, is a product of
the plume. It is not generally appreciated, however, that this “classic” case of a plume
contradicts the plume model in many ways. For example, there is little petrological
evidence as yet that the Deccan source was “abnormally hot,” and the short (~1.0–
0.5 m.y.) duration claimed by some for the eruption of the Deccan is in conflict with
recent Ar-Ar age data that suggest that the total duration was at least ~8 m.y. The Deccan continental flood basalts (CFB) were associated with the break-off of the Seychelles microcontinent from India. Geological and geophysical data from the Deccan
provide no support for the plume model and arguably undermine it altogether. The
interplay of several intersecting continental rift zones in India is apparently responsible for the roughly circular outcrop of the Deccan. The Lakshadweep-Chagos Ridge
and the islands of Mauritius and Réunion are located along fracture zones, and the
apparent systematic age progression along the ridge may be a result of southward
crack propagation through the oceanic lithosphere. This idea avoids the problem of a
10o paleolatitude discrepancy which the plume model can solve only with the ad hoc
inclusion of mantle roll. Published Ar-Ar age data for the Lakshadweep-Chagos Ridge
basalts have been seriously questioned, and geochemical data suggest that they likely
represent postshield volcanism and so are unsuitable for hotspot-based plate reconstructions. “Enriched” isotopic ratios, such as values of 87Sr/86Sr higher than those for
normal mid-ocean ridge basalts, which have been observed in basalts of the ridge and
the Mascarene Islands, may mark the involvement of delaminated enriched continental mantle instead of a plume. High values of 3He/4He also do not represent a deep
mantle component or plume. The three Mascarene islands (Mauritius, Réunion, and
Rodrigues) are not related to the Deccan but reflect the recent (post-10 Ma) tectonicmagmatic development of the Africa Plate. I relate CFB volcanism to continental rifting, which often (but not always) evolves into full-fledged seafloor spreading. I ascribe
the rifting itself not to mantle plume heads but to large-scale plate dynamics themselves, possibly aided by long-term thermal insulation beneath a supercontinent that
may have surface effects similar to those predicted for “plume incubation” models.
*E-mail: jcsheth@iitb/ac/in.
Sheth, H.C., 2005, From Deccan to Réunion: No trace of a mantle plume, in Foulger, G.R., Natland, J.H., Presnall, D.C., and Anderson, D.L., eds., Plates, plumes,
and paradigms: Geological Society of America Special Paper 388, p. 477–501. For permission to copy, contact [email protected]. © 2005 Geological Society
of America.
477
478
H.C. Sheth
Nonplume plate tectonic models are capable of explaining the Deccan in all its greatness, and there is no trace of a mantle plume in this vast region.
Keywords: PLEASE SUPPLY FIVE OR SIX KEYWORDS.
INTRODUCTION
Many continental flood basalt (CFB) provinces of the world
formed during the rifting and breakup of continents (see e.g.,
Storey et al., 1992 and references therein). An excellent example is the ca. 65–60 Ma Deccan province of India (Fig. 1). It
is one of the larger and better-preserved CFB provinces of the
world, with a present-day areal extent of ~5 × 105 km2 and an
estimated original area of at least 1.5 × 106 km2 (e.g., Wadia,
1975). The Deccan CFB formation was associated with continental rifting and the break-off of the Seychelles microcontinent
from India (e.g., Norton and Sclater, 1979; Mahoney, 1988;
Devey and Stephens, 1991; Fig. 1). Greater India (India plus the
70o E
d
iT
ren
Cam
Bundelkhand
Craton
Nar
Satpura Tren
20o N
Tapi R
ift
DECCAN
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ARABIAN
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Seychelles) was involved in two other continental breakup events
during the Mesozoic prior to the Deccan episode, both of which
were also associated with major flood basalt volcanism. Thus,
Greater India broke off from Madagascar at ca. 88–85 Ma, and
this was associated with the formation of the Indo-Madagascar
flood basalt province. This province, though considerably eroded
today, is represented by extensive lavas and dike swarms in
Madagascar, the submarine Madagascar Plateau south of Madagascar, and some volcanics and dike swarms in southern India
(e.g., Storey et al., 1995, 1997; Anil Kumar et al., 2001; Pande
et al., 2001; Fig. 1). Prior to this event, at ca. 120–116 Ma,
the Rajmahal-Sylhet flood basalt province formed in easternnortheastern India (Baksi, 1995) as part of the early Cretaceous
0o
end
Africa
20o
Granulite
Terrain
10o N
Lakshadweep
(Laccadive) islands
40o
Sri
Lanka
60o
India
Figure 1. Map showing the approximate
boundaries of the Precambrian cratons
making up the Indian basement crust
(e.g., Naqvi and Rogers, 1987; Pandey
and Agrawal, 1999), the granulite terrain (stippled), the Precambrian structural
trends (heavy broken lines), the rift
zones crossing peninsular India (e.g.,
Biswas, 1982), and the present outcrop
areas of the Deccan and Rajmahal flood
basalts (shaded). The Saint Mary’s islands felsic volcanics and some of the
many Karnataka-Kerala dike swarms are
related to the India-Madagascar breakup
(e.g., Anil Kumar et al., 2001; Pande
et al., 2001); there are many Precambrian
dike swarms throughout southern India
as well. The inset shows the breakup of
the Seychelles microcontinent (black)
from India at ca. 65 Ma (after Norton and
Sclater, 1979; Mahoney, 1988).
From Deccan to Réunion: No trace of a mantle plume
The hyphen in
OntongJava was
used in
every
other instance in
the volume, so
I’ve left it
here.
Rajmahal-Kerguelen Province, and its formation succeeded by
a few million years the break-off of India from Australia (e.g.,
Kent et al., 1997).
During the last decade, the mantle plume model for flood
basalt volcanism (Richards et al., 1989; Campbell and Griffiths,
1990) has been widely accepted (e.g., Hooper, 1990; Hill, 1991;
Kent et al., 1992). However, observations anomalous with respect to the predictions of the model are increasingly coming to
light from the flood basalt provinces of the world. For example,
regional prevolcanic lithospheric uplift of 1–4 km is a prerequisite in all thermal models, such as the plume model (e.g., Campbell and Griffiths, 1990; Farnetani and Richards, 1994), but such
uplift was absent in the largest CFB, the Siberian Traps (Czamanske et al., 1998), and absent or far less than expected in the
world’s largest oceanic plateau, the Ontong-Java Plateau (Tejada
et al., 2002, 2004). The very existence of mantle plumes has
been questioned (e.g., Anderson et al., 1992; Smith, 1993; Anderson, 1994, 1999; Sheth, 1999a,b; Smith and Lewis, 1999).
Numerous nonplume mechanisms and explanations have been
forthcoming, including volatile-rich mantle sources (e.g., Bonatti, 1990; Smith, 1993; Janney et al., 2000), fertile (eclogite-rich)
mantle sources (e.g., Foulger et al., 2005 and this volume), crackrelated volcanism (e.g., Smith, 2003 and this volume; Natland and
Winterer, this volume), top-down lithospheric control and edgedriven convection (EDC; e.g., Mutter et al., 1988; King and
Anderson, 1995, 1998; Anderson, 1998a), and asteroid impacts
(e.g., Jones et al., 2002). Besides, with new data, especially
geophysical data, many regions traditionally considered classic
plume areas are now attributed by several workers to nonplume
plate tectonic processes. Two prominent examples of such are
Iceland (e.g., Foulger et al., 2000 and this volume; Lundin and
Doré, this volume) and Yellowstone (e.g., Humphreys et al.,
2000; Christiansen et al., 2002). Iceland is the world’s classic
ridge-centered “hotspot,” and Yellowstone is generally considered the type example of a continental “hotspot.”
The Deccan Traps of India constitute one of the world’s
best-developed and best-known CFB provinces, and many
workers have invoked or supported a plume origin for the Deccan. As first put forth by Morgan (1972, 1981), this model postulates that (1) the Deccan originated from the ancestral Réunion
hotspot, which upwelled beneath India in the late Cretaceous,
and (2) the hotspot, now located on the Africa Plate, is fed by a
deep mantle plume. Fluid dynamic modeling (e.g., Campbell
and Griffiths, 1990) postulated that plumes rise buoyantly from
the core-mantle boundary (CMB) and, by entrainment of surrounding mantle, develop large bulbous “heads” that remain
connected to the source region by narrow “tails.” In the plume
model, the break-off of the Seychelles from India and the associated Deccan volcanism were both consequences of the impingement of the mantle plume on the Indian lithosphere (e.g.,
Hooper, 1990; Courtillot et al., 1999). In the present contribution, I extend my previous arguments (Sheth, 1999a,b, 2000,
2005) that a wealth of geological and geophysical evidence from
the Deccan CFB province, and from the subsequent hypothe-
479
sized products of the Réunion hotspot, is at odds with the predictions of the plume model. I begin by briefly stating what I
mean by a plume.
WHAT IS A PLUME?
Morgan (1972) proposed plumes that were fixed in the
Earth so that they could be used as reference frames to record
plate motions, and because of the perceived fixity of hotspots,
the deep plume explanation was considered superior to other
explanations for island-seamount chains, such as crack-related
volcanism (see Anderson and Natland, this volume). Now it is
thought, however, that plumes sway in the mantle wind and
need not be fixed (Steinberger and O’Connell, 1998). Many
think of “hotspots”—locations of intraplate volcanism—as thermal plumes, but most hotspots are thought to be no hotter than
average (Stein and Stein, 2003; DeLaughter et al., this volume;
Green and Falloon, this volume). Several classic “plumes” show
no evidence for an initial head phase (e.g., Hawaii) or a tail
phase (e.g., Siberia, Ontong-Java), and several plume proponents now suggest that the classic model is too restrictive and
that not all plumes need have heads or tails, and not all plumes
need be deep-sourced (e.g., Cserepes and Yuen, 2000). Tailless
upper-mantle “plumes” of the type conceived by Cserepes and
Yuen (2000) still cannot explain observations such as the absent
or insignificant prevolcanic lithospheric uplift in Siberia or
Ontong-Java. Campbell and Griffiths (1990), originators of the
modern model of plume heads and tails based on fluid dynamic
modeling, postulated that the inflated heads of new plumes contain plume source material into which surrounding cooler midocean ridge basalt (MORB) mantle becomes entrained during
upwelling and that plume tails contain hot plume source material. Van Keken (1997), relying on numerical modeling, came
to the opposite conclusion, however, saying that the heads of
plumes should contain primitive plume source material, and the
tail should entrain surrounding mantle.
A glossary of plume definitions and many technical articles
on the subject can be found at www.mantleplumes.org. Clearly,
plume means different things to different workers, but this
should not divert us from the main issue—whether there is evidence for (1) “abnormally hot” mantle that is (2) actively upwelling, and (3) has its origin in the deep mantle or at the CMB.
These three qualities are argued and believed to be the prime
characteristics of plumes that distinguish them from shallower
mechanisms. Nevertheless, Courtillot et al. (2003) have defined
three different types of plumes: their “primary” or “Morganian”
plumes come from the deepest mantle or the CMB, and a second category of plumes originates at the base of the mantle Transition Zone, whereas a third “Andersonian” category includes
plumes that are shallow-sourced. Courtillot et al. (2003) find
that only seven out of the forty-nine hotspots of the world satisfy their criteria for deep origin, and Réunion is one of them.
However, as Anderson (this volume) shows, these criteria are
subjective.
480
H.C. Sheth
In this chapter I take plume to mean an active upwelling that
is narrow (relative to the plates), anomalously hot, and deepsourced (originating in the lowermost mantle or at the CMB).
As the modern plume model is based essentially on fluid dynamical experiments, I also expect that a plume has a bulbous
head connected to the source region by a narrow, pipe-like tail.
Having defined a plume, I now examine field data from the Deccan and India.
GEOLOGY OF THE DECCAN AND ITS BASEMENT
Continents consist of cratons amalgamated by networks of
orogenic belts that contain oceanic, island-arc, and continental
margin rocks besides local fragments of older cratons, and a true
continent must contain long-stable cratons (Rogers, 1996). According to Rogers (1996), a block of continental crust is a craton only if it has been sufficiently stable to provide a basement
for the deposition of shallow-water or subaerial volcano sedimentary suites on platforms, in broad unrifted basins, or in rift
valleys. The Indian subcontinent has a rich rock record from the
Early Archean up to Recent time. At least six Archean to early
Proterozoic cratonic nuclei are recognized. These are the Aravalli, Bundelkhand, Singbhum, Bastar, and Dharwar cratons
and the high-grade granulite terrane in the far south (e.g., Naqvi
and Rogers, 1987; Rogers, 1996; Pandey and Agrawal, 1999;
Fig. 1). These nuclei may have been sutured together in Archean
and Proterozoic time by proto–plate tectonic processes (e.g.,
Naqvi et al., 1974; Valdiya, 1984; Radhakrishna, 1989).
Several major rift zones traverse the subcontinent (Fig. 1).
The Godavari and Mahanadi Rifts lie in the east, the Cambay
Rift in the north-northwest, and the Kachchh Rift in the northwest. The Narmada-Tapi Rift Zone, a major extensional zone
within the Indian peninsula, runs in an ENE-WSW direction
for >1600 km along the central part of India (e.g., Mishra, 1977).
This comprises the Narmada and Tapi grabens, which are separated by an upraised horst block, the Satpura range. The Indian
rifts are known to run along major Precambrian tectonic trends
(e.g., Katz, 1978). The Narmada zone is a prominent, ancient
line of weakness and is considered a Proterozoic protocontinental suture between a northern (Aravalli) and a southern (Dharwar)
protocontinent (Naqvi et al., 1974). The western Indian coast
and the Cambay Rift also developed by faulting parallel to the
NNW-SSE Dharwar orogenic trend of Precambrian age (Raju,
1968; Biswas, 1987). Another major Precambrian orogenic trend,
the NE-SW Aravalli trend, splays into two trends at its southern end: the east-west Delhi trend (along which the Mesozoic
Kachchh Rift has developed) and the main NE-SW Aravalli
trend, which continues right across the Cambay Rift into the
Saurashtra peninsula (Biswas, 1982). Two more rifts, the NWSE-trending Koyna and Kurduvadi Rifts were postulated under
the southern Deccan region (Fig. 1) on the basis of linear gravity lows (Krishna Brahmam and Negi, 1973), but these gravity
lows may instead represent upwarps of the sub-Deccan base-
ment. The extensions of these two hypothesized rifts to the north
and the south are not clear.
Figure 2 shows the main rock formations that constitute
Indian geology: a large portion of the Indian Shield is made up
of Precambrian rocks, and there are many younger sedimentary
basins. The Deccan lava pile, which obscures the basement from
observation over 0.5 million km2, is thickest (~2000 m) along
the Western Ghats region (Sahyadri range) adjacent to the western coast and thins progressively eastward and southeastward,
such that along the eastern fringes of the province the lava pile
is only ~200 m thick. Whereas the Deccan lava pile in the Western Ghats region and in the interior areas of the province is made
up almost completely of tholeiitic basalts (e.g., Beane et al.,
1986), felsic and alkaline magma types are also prominent along
the rift zones and along the west coast. Considerable volumes
of acid and basic tuffs and rhyolite and trachyte lavas exist along
the coast, as at Bombay (Lightfoot et al., 1987; Sheth et al.,
2001a; Sheth and Ray, 2002). The west coast and the rift zones
are also where significant tectonic-structural disturbances such
as listric faulting and monoclinal flexing have affected the lava
pile (e.g., Guha, 1995; Sheth, 1998). The coastal dikes form
regional, dominantly north-south-oriented swarms and are mostly
of basalts, dolerites, and alkaline rocks such as lamprophyres
(e.g., Dessai, 1987, 1994; Murthy, 1987). Significant volumes of
felsic rocks and many alkaline complexes (several of which include carbonatites) are found along the Narmada Rift and in the
northwest of the Deccan proper (the Tavidar felsic volcanics and
the Mer Mundwara and Sarnu-Dandali alkaline complexes; e.g.,
Basu et al., 1993, and Roy, 2003). The Narmada-Satpura-Tapi
zone also contains major linear dike swarms (e.g., Deshmukh and
Sehgal, 1988; Sant and Karanth, 1990; Keshav et al., 1998).
Drilling in the Latur area (Fig. 2), the epicenter of a disastrous earthquake (M = 6.3) in 1993 and situated on the proposed
Kurduvadi Rift, directly encountered Precambrian basement (the
Peninsular Gneiss of the southern Indian Shield) at 338 m depth
(Gupta et al., 1998). Many believe that the pronounced linearity
of the west coast and the continental margin suggests structural
control (e.g., Biswas, 1987). The newly formed continental margin and the rift zones may have constituted major vent areas for
the Deccan lavas, as inferred from abundant mafic dike swarms
and intrusions, high heatflow, and aligned thermal springs (e.g.,
Sheth, 2000). Seismic studies and drilling for oil have shown
that the Deccan basalts continue beyond the west coast and onto
the continental shelf (Biswas, 1982; Chandrasekharam, 1985).
Eruptive centers undoubtedly existed along the present-day submerged shelf. The Cambay Rift and the region offshore of the
west coast are regions of productive oil and gas fields. Much of
the Cambay region is today covered by Tertiary and Quaternary
sediments, and the underlying igneous rocks are not exposed.
Boreholes drilled by the Oil and Natural Gas Commission have
penetrated thick (5 km) Tertiary sediments, and at places the
underlying basalts are known, based on seismic data, to be over
4 km thick (Kailasam and Qureshy, 1964; Mahadevan, 1994).
From Deccan to Réunion: No trace of a mantle plume
481
Sarnu
70o
80o
76o
Mundwara
22o
Kala Dongar
Bhuj
Anjar
Malwa
Plateau
Pavagadh
Barda
Botad
Chogat-Chamardi
Palitana
Rajpipla
Navsari
Girnar
20
Pachmarhi
Mhow
Amba Dongar
Toranmal Satpura Mts.
Dediapada
Chikaldara
Nandurbar
Dhule
o
Bombay
Alibag
dla
L
Nagpur
Nasik
Figure 2. Sketch map of western and
central India showing the main features
of its geology and the outcrop of the
Deccan flood basalts (shaded). Also
shown are localities mentioned in the
text. Based on Wadia (1975).
Igatpuri
Deccan
Plateau
Pune
Mahabaleshwar
18o
Man
Jabalpur
obe
Buldana
Western Ghats
24o
Latur
Kurduvadi
Western Ghats
Koyna
ARABIAN
SEA
16o
Belgaum
Rajahmundry
Goa
0
100
200
km
Archaean-Proterozoic Late Precambrian Paleozoic-Mesozoic
Well-studied
lava sections
Deccan
Felsic lavas, plugs, alkaline-carbonatite
complexes, radial dike swarms
THE PREVOLCANIC UPLIFT ISSUE
It has been argued (e.g., White and McKenzie, 1989) that the
huge volumes erupted in CFB provinces such as the Deccan require a large mass of “abnormally hot” mantle. Such arguments
are based on the assumption that the mantle is normally cold and
subsolidus and has no lateral temperature variations. However,
lateral temperature variations of ~200 oC are evident in the mantle from seismic tomography (e.g., Anderson et al., 1992) and are
a natural consequence of “normal” plate tectonic processes (Anderson, 2000b). Also, the plume model, which is primarily a thermal model, requires substantial (1–4 km) broad-scale prevolcanic
uplift of the lithosphere ~5 m.y. before the onset of flood volcanism (Campbell and Griffiths, 1990; Farnetani and Richards,
1994). If such regional domal uplift was absent, a thermal mechanism would be violated. Such uplift did not predate magmatism
Tertiary-Quaternary
Linear dike
swarms
in many flood basalt provinces (e.g., Menzies, 2000). Regional
domal uplift was absent during the enormous CFB event that
formed the Siberian Traps (ca. 250 Ma) and absent or far from adequate during the formation of the Ontong-Java oceanic plateau,
the world’s largest, which remained completely submerged
during its construction (e.g., Czamanske et al., 1998; Tejada et al.,
2004). In fact, the Siberian CFB are underlain almost everywhere by terrigeneous sediments of the Tungusskaya Series
(320–250 Ma), which include the Tunguska coaliferous basin, the
world’s largest (Czamanske et al., 1998). Regional subsidence,
not regional uplift, is thus seen (Elkins-Tanton and Hager, 2000).
Clues in the Sub-Deccan Rocks
In the case of the Deccan, evidence for regional domal uplift is absent as well. Over most of the Deccan province today,
482
H.C. Sheth
the lava-basement contact is at considerable depths in the subsurface. As noted, there is a huge thickness (~1700 m exposed)
of Deccan basalts in the Western Ghats (Sahayadri) region and
an additional ~500 m in the subsurface (as identified from seismic studies, e.g., Kaila et al., 1981b,c). The Deccan lavas overlie two linear Mesozoic sedimentary basins in the Narmada-Tapi
region (Kaila, 1988; Sridhar and Tewari, 2001). The northern,
the Narmada basin, is 1000 m thick, and the southern, the Tapi
basin, is 1800 m thick. The thick Mesozoic sedimentary pile
under the Deccan (Upper Gondwana Supergroup, Triassic) in
this region is exposed in the Pachmarhi region by tectonic uplift
of the Satpura range (Fig. 2).
Toward the western part of the Narmada Rift, the Deccan
lavas overlie sandstones and limestones of the Bagh Formation
deposited during a marine transgression in the Late Cretaceous
(Sheth, 1999b). This is where the center of the plume head and
maximum uplift should have been. The fact is that there were
both local uplifts and subsidences just before volcanism in this
area (Tandon, 2002). Interestingly, Maastrichtian sediments of
the Lameta Formation (which, like the Bagh Formation, underlies the Deccan lavas in this area) include clays derived in part
from the Deccan basalts themselves; clearly, the Deccan lavas
would have erupted and then undergone some uplift to form the
source areas for the Lameta clays (Salil and Shrivastava, 1996,
Salil et al., 1997; Tandon, 2002). Different conditions are found
in the Lametas of the Dongargaraon Basin of the Nagpur area
(Fig. 2). Tandon (2002) has recorded a clear “shallowing up”
trend from shallow lake deposits to a paleosol before the terrain
was buried by the first lava flow. He has related this to prevolcanic surface uplift of the area on the order of meters only,
and it is possibly also related to mock aridity (Harris and Van
Couvering, 1995; Khadkikar et al., 1999). Again, local uplifts and
subsidences (e.g., Jerram and Widdowson, 2005) cannot be used
to support or refute the plume model, because they are easily
related to the filling and emptying of magma chambers, to emplacement of intrusions, to faulting, and to related processes.
Campbell and Griffiths (1990) cited Pachmarhi as the center of a broad, uplifted dome caused by the plume, mentioning
that the lava-basement contact at Pachmarhi is over 1 km above
sea level. They did not consider it important, or were unaware,
that the contact is hundreds of meters in the subsurface over
most of the province. Pachmarhi is an isolated case. Indeed,
even over large parts of the Satpura range (e.g., at Toranmal;
Fig. 2) the lava-basement contact is in the subsurface. Furthermore, the uplift of the Pachmarhi block appears to be a result not
of prevolcanic doming but of postvolcanic, recent uplift because
of the very youthful landscape (kilometer-high escarpments in
the Gondwana sandstones, V-shaped ravines and gorges, and torrential rivers). Choubey (1971) recognized successive planation
surfaces here, and the highest is at ~1300 m above sea level. As
Ollier and Pain (2000) argued, such surfaces must form near the
base level of erosion of a river, which is at sea level in most
cases, and because there is no geomorphic process capable of
creating a planation surface at a high elevation, a planation surface at high elevations above sea level must indicate recent,
rapid uplift. Dixey (1970, reprinted in Subbarao, 1999), based
on field work in this region, gave evidence that the Deccan lavas
were erupted over an old, regional, flat-lying erosion surface
developed over older rocks, and noted that this surface could be
put to good use in deciphering the subsequent tectonic evolution
of the region.
Casshyap and Khan (2000) provided evidence for “preDeccan doming” of the Indian subcontinent, again based on field
studies in the Pachmarhi region. They identified three separate
uplift events, the latest of which resulted in Late Jurassic–
earliest Cretaceous sediments with a source in northwestern
India. Clearly, uplift centered on northwestern India and preceding Deccan CFB volcanism by a long time (~70 m.y.) cannot be
considered evidence for Deccan plume-related prevolcanic uplift.
Broad-scale domal, prevolcanic uplift of the Indian lithosphere
just prior to Deccan volcanism has yet to be demonstrated. However, the old and extensive erosion surface beneath the Deccan
lavas in central India (e.g., Dixey, 1970), suggests that such uplift did not occur.
Clues in the Indian Rivers
Many major Indian rivers originate in the Western Ghats,
not far from the west coast, and yet flow for hundreds of kilometers eastward to eventually meet the Bay of Bengal, which is
a remarkable fact. Cox (1989) speculated that the pronounced
easterly drainage of the Indian peninsula (Fig. 3) was a consequence of plume-caused lithospheric doming. The Narmada and
the Tapi, two of India’s major rivers, flow westward, however,
and Cox ascribed this to their exploiting a rift system in the
dome. He did not address why such a rift system should produce
a westerly drainage (toward the topographically high center of
the uplifted dome). Summerfield (1990) discussed problems
with Cox’s views, with examples from Africa.
Ollier and Powar (1985) observed that the drainage pattern
of the Indian peninsula is dendritic over both the region of the
Deccan lavas and the older basement. This led them to suggest
that the drainage developed subsequent to the eruption of the
Deccan lavas. The uplift of the Western Ghats must have been
still later, noting their very youthful topography and the evidence that formation of several of the major east-flowing Indian
rivers was antecedent. The Western Ghats (Sahyadri range)
constitute a roughly north-south-trending, 1500 km–long, nearly
continuous “Great Escarpment” that reaches heights of >2.5 km
in the Nilgiri and Palni-Kodaikanal massifs of southern India
(Ollier, 1990; Gunnell and Radhakrishna, 2001; Fig. 3). Notice
that the Cauvery River originates on the western flank of the Western Ghats and flows east through the highest part of the plateau.
This is impossible unless the formation of the river was antecedent to the uplift of the plateau (Ollier and Powar, 1985).
Widdowson and Cox (1996) provided similar observations and
From Deccan to Réunion: No trace of a mantle plume
22
o
Narmada R.
R
U
P
T
A
S
W
S.
T
M
A
Tapi R.
Ma
ha
1646 m (Kalsubai)
E
483
na
di
R.
S
Bombay
G
.
iR
ar
av
E
1436 m (Mahabaleshwar)
n Plain
R
G
N
Kr
ish
na
Figure 3. The main elements of the physiography of the Indian peninsula. The
Western Ghats escarpment is indicated
by the heavy broken line. Note the pronounced easterly drainage. After Ollier
and Powar (1985).
R.
N
rR
A
Karnataka
Plateau
.
A
T
BAY OF
BENGAL
R
nne
E
Pe
H
T
o
A
S
14
T
H
G
ARABIAN
SEA
S
od
Konka
18o
Deccan
Plateau
T
E
S
Pon
naiy
Nilgiri
Massif
Cau
Palni-Kodaika
nal Massif
o
0
.
very
Palghat Gap
R.
2695 m (Anaimudi)
10o
72
ar R
500
2670 m (Dodabetta)
76o
km
Va
iga
iR
.
Sri
Lanka
arguments for the area near Mahabaleshwar. The uplift of the
Western Ghats and the associated tilting of the Indian peninsula
from west to east are young, ongoing, and postvolcanic.
In summary, there is no evidence as yet from the Deccan for
regional prevolcanic uplift that can be related to an upwelling
thermal plume. Existing data offer two possibilities: (1) prevolcanic lithospheric uplift occurred, and the presently buried
lava-basement contact over most of the province is due to rapid
decay of the thermal uplift and loading of the crust by the lava
pile (Campbell and Griffiths,1990), or (2) prevolcanic lithospheric uplift never occurred, and a thermal mechanism is invalid. The second scenario is similar to the situation at Siberia
and Ontong-Java and is the explanation I prefer. Campbell and
Griffiths (1990) argued that prevolcanic regional domal uplift
due to a plume head may not be significant due to lateral migration of magma in the crust. Jerram and Widdowson (2004)
argue that while an area may truly be undergoing uplift on the
scale of hundreds of kilometers, areas within it may be undergoing subsidence on the scale of tens of kilometers, and thus
evidence for subsidence cannot be used to disprove uplift. It
seems to me that, given the observations, the plume might as
well not have been there. Fortunately, it is possible to conclude
84
o
that broad-scale prevolcanic uplift did not occur, because (1) there
is no evidence for such uplift, and (2) there is actual evidence
against it (such as the old erosion surface beneath the Deccan
lavas in central India). Apatite fission track data can also help.
Some have recently become available, but remain inconclusive
regarding the timing of Western Ghats uplift (Gunnell et al.,
2003).
THE DECCAN: SOME MYTHS AND FACTS
Having considered the issue of prevolcanic uplift, I find it
is instructive to consider some of the main arguments repeatedly
offered in favor of a plume origin of the Deccan. These arguments have little support in the data themselves.
An Anomalously Hot Mantle Source?
There is no evidence for an “abnormally hot” mantle source
for the Deccan lavas. Some alkaline picritic liquids (identified
based on appropriate whole-rock Mg #s and olivine compositions) are encountered in boreholes in the northwestern Deccan
(around Botad; see Fig. 2) and in parts of the Narmada Rift
484
H.C. Sheth
(Melluso et al., 1995; Krishnamurthy et al., 2000). Campbell
and Griffiths (1990) thought the borehole lavas were hightemperature, high–melt fraction liquids from the plume axis.
Peng and Mahoney (1995), however, found that they are somewhat alkaline and are consistent with high-pressure, low–melt
fraction melting. The thick sequence of the Deccan basalts in the
Western Ghats (Figs. 1 and 2) includes picritic basalts, but these
are enriched in cumulus olivine and clinopyroxene and do not
represent liquid compositions. The parental melts of these picrites are estimated to have contained only ~9%–10% MgO
(Beane and Hooper, 1988; Sheth, 2005). It is possible that the
sub-Indian mantle was warmer than average during Deccan
volcanism; it could have been warm over a very broad region
(thousands of kilometers) as a consequence of long-term thermal insulation under Gondwana.
A Very Short-Lived Catastrophic Eruption?
Very rapid, catastrophic emplacement of the Deccan Traps
has been one of the key arguments for a plume origin, but the
duration of the volcanism remains one of the most debated issues. Recent 40Ar-39Ar data for trachyte and basalt flows from
Bombay (Sheth et al., 2001a,b) suggest that the total duration
was no less than ~8–9 m.y. There may have been a major, rapid,
short-duration eruptive phase in the Western Ghats, estimated
by some to have lasted only 1.0–0.5 m.y. (e.g., Courtillot et al.,
1988; Duncan and Pyle, 1988; Hofmann et al., 2000) and by
others to have been more protracted, 4–5 m.y. (Venkatesan
et al., 1993; Pande, 2002). Allègre et al. (1999) reported an ReOs isochron age of 65.6 ± 0.3 Ma (2s) for basalt lava flows sampled over a large area of the province, arguing for intense, very
short-lived (~0.5 m.y. or less) volcanism right at the CretaceousTertiary boundary. That random, noncomagmatic samples collected over an area 1000 km across and occupying various
topographic-stratigraphic levels should define an isochron is
remarkable, but the goodness-of-fit value (F) for the claimed
isochron, which was not reported, is 22 (Baksi, 2001a). The line
is clearly an “errorchron” (Faure, 1986), and the age uncertainty
is much greater than the 0.3 m.y. claimed by Allègre et al.
(1999).
Sen (2001) calculated an eruptive duration for the thick
Western Ghats lava pile that was ten times shorter than even that
postulated by most “rapid volcanism” proponents. Based on the
estimated formation times of large plagioclase crystals in some
of the Deccan flows (by comparisons to Hawaiian plagioclases)
and on the derivation of a “one-dimensional eruption rate,” he
proposed that the eruption of the entire Western Ghats sequence
took only ~55 k.y. Sheth (2002) analyzed this approach, arguing
that one must use a volumetric eruption rate (presently unknown)
and that the one-dimensional eruption rate was a meaningless
quantity and therefore the calculated duration of 55 k.y. could
be grossly in error.
Physical volcanology can also offer clues. Some authors
have invoked hot plume heads to explain the extremely high
lava eruption rates of CFB, though such eruption rates have yet
to be demonstrated. A large proportion of the Deccan basalts is
made up of pahoehoe compound lava flows (e.g., Deshmukh,
1988). If new models of protracted, gradual emplacement of
similar lava flows from the Columbia River Province (Self et al.,
1997) are correct, most of the Deccan could have formed at low
eruption rates. In fact, Bondre et al. (2004) find that the scale
of the individual lava lobes in the Deccan pahoehoe flows is the
same as that of modern Hawaiian flows, though the Hawaiian
flows themselves are much smaller than the Columbia River and
the Deccan lava flows. The large volumes of individual Deccan
flows require explanation, however, and may reflect an extrafertile mantle source (relative to peridotite) (Sheth, 2005), a
hydrous or CO2-rich mantle (e.g., Presnall et al., 2002; never
evaluated), a lithospheric regime dominated by extension (Sheth,
2000), and great lengths of the fissure systems (Self et al., 1997).
Dikes 50–60 km in length are common in the Narmada-SatpuraTapi region (e.g., Deshmukh and Sehgal, 1988; Keshav et al.,
1998).
A Systematic Southward Stratigraphic Younging?
Southward stratigraphic younging of the various formations
within the Western Ghats region, apparent in the earlier years of
Deccan geochemical stratigraphy, has also been widely used in
support of the plume model. Stratigraphically younger lavas
have been said to have erupted in progressively more southerly
locations consistent with the passage of the plate over a plume
(Cox, 1983; Mitchell and Widdowson, 1991). The latest works
on geochemical stratigraphy have provided new data from other
parts of the province that contradict this view: thick lava piles
closely resembling (in both elemental and isotopic composition)
some of the youngest formations of the Western Ghats stratigraphy are now known to outcrop in far northern areas of the Deccan (e.g., Mahoney et al., 2000; Sheth et al., 2004), and there is
no evidence to suggest that the source of the lavas or the eruptive centers moved systematically southward with time.
A Cambay Triple Junction?
Originally included by Burke and Dewey (1973) in their
list of plume-generated triple junctions worldwide, the Cambay
triple junction is not real, because the Narmada Rift Zone continues into the Saurashtra peninsula and it, the west coast rift,
and the Cambay Rift form a cross (Sheth, 1999a; Fig. 1). The
Kachchh Rift is another structure that the Cambay triple junction idea does not address. Nevertheless, papers subsequent to
that of Burke and Dewey (1973) have popularized this triple
junction and the Réunion plume model for the Deccan. Another
unfortunate development is the proliferation of model-dependent
interpretations by which every geological and geophysical observation from the Deccan is interpreted as an effect of the
Réunion plume. For example, low–seismic velocity mantle
underlying the Cambay Rift of the Deccan is interpreted as a
From Deccan to Réunion: No trace of a mantle plume
remnant of the plume (Kennett and Widiyantoro, 1999) instead
of as simply passive upper mantle upwelling.
Geophysical data, like geological data, can offer valuable
clues and insights, and to these we now turn.
GEOPHYSICS OF THE DECCAN
AND ITS BASEMENT
Heatflow
The thermal structure of the large stable region of the
southern Deccan is characterized by a heatflow in the range of
40–70 mWm–2, which is the normal low cratonic heatflow
found over the southern Indian basement shield (Gupta and
Gaur, 1984). Heatflow values are 75–93 mWm–2 in the BroachAnkleshwar area, between Cambay and Surat (Fig. 4). Heatflow
is high (average 83 mWm–2) in the northern part of the Cambay
graben, where temperature gradients are >70 oC/km in some
zones (Pandey and Negi, 1995). At 3 km depth, in situ temperatures were estimated, using drill hole data, to be as high as 175
± 25 oC. In the Bombay offshore region, the average heatflow
is ~83 mWm–2 and temperature gradients are 36–78 oC/km. The
expected temperature at a depth of 3 km is ~175 ± 50 oC. In the
Konkan plain between the west coast and the Western Ghats,
some sixty thermal springs are distributed over a linear N-S dis-
41
70
77
24o
51
80
62
77
75
94
64 89
22o
40
70
SAURASHTRA
98
82
55
Surat
67
116
ad
Narm
Cambay
a Riv
122
120
Tapi River
20
o
79
51
67
er
75 49
110
60 96
126
85
61
120
58
41
47
63
Bombay
76
18o
Arabian
Sea
114
128
113
126
41
40
113
16o
70o
72o
74o
76o
tance of 300 km, with temperatures ranging from 34 to 71 oC
(Pandey and Negi, 1995) (Fig. 4). However, along the NarmadaTapi Rift, the Saurashtra peninsula coast, and the west coast, the
heatflow structure is also determined by convective heat transfer, and the measurements may not accurately reflect the crustal
heat production (Roy and Rao, 2000). Biswas (1987) has suggested that the current high thermal regime of the Cambay and
Bombay offshore regions marks a renewed rifting phase. Roy
and Rao (2000) reported high heatflow values from the Cambay
Rift; however, they stated that there was no evidence for thermal transients associated with Deccan volcanism in the Deccan
region proper (south of the Tapi Rift).
A hot plume under the lithosphere is expected to cause thermal erosion of the lithosphere and thereby produce a thinned
lithosphere and high heatflow. Negi et al. (1986) proposed a
drastically thinned and anomalously hot Indian lithosphere, with
estimates of present-day lithospheric thickness as low as 60 km
(and 40 km under Cambay). They suggested that the Indian lithosphere was both greatly thinned and abnormally hot, and this
was the reason for its supermobility (at Cretaceous–Tertiary time
India was moving northward at superfast rates of 15–20 cm/yr;
e.g., Patriat and Achache, 1984). Gupta (1993) questioned this
view, stating that the arguments of Negi et al. (1986) for pronounced thermal erosion of the Indian lithosphere were based
on the scarce data available then. With much more geothermal
data compiled for Precambrian shield areas of India, Africa,
Australia, and Brazil (all of which formed parts of Gondwanaland), Gupta (1993) found no support for the notion that the
Indian Shield was hotter than other shields. He concluded that
the Indian landmass is no hotter than the other Gondwana landmasses and that its supermobility was not a consequence of being hot, and apparently not related to its thermal characteristics.
Gravity Studies and Deep Seismic Sounding:
Normal or Upwarped Moho?
119
86
121
137
97
485
78o
Figure 4. Heatflow values (in mW/m2) and thermal springs over western India (after Singh, 1998). The solid circles denote values observed
conventionally (Singh and Meissner, 1995); the open circles denote
geochemically estimated heatflow values (Ravishanker, 1988). The
stars mark thermal springs (Krishnaswamy and Ravishanker, 1980).
The elliptical areas defined by heavy dotted lines are high-gravity
anomalies (Glennie, 1951; Takin, 1966; Mahadevan, 1994).
The most striking gravity feature, not only of the western
Indian margin but of the entire Deccan province, is perhaps the
60 km–wide high Bouguer anomaly close to the Bombay, Surat,
and Saurashtra coasts (Fig. 4; Mahadevan, 1994). The Bouguer
anomaly values reach extreme lows of –100 mgal some 150 km
east of Bombay and highs of as much as +50 mgal at the Bombay coast. The high gravity anomaly near Bombay was attributed
by Glennie (1951) to a 26 km–wide mafic dike off Bombay and
by Takin (1966) to a differentiated magma chamber (of olivine
gabbro bulk composition and a density contrast of 0.4 g/cm3
with its surroundings). Kaila (1986) suggested that these gravity highs resulted from Moho upwarps and that the crust under
this region was abnormally thin. This suggestion gained further
support from deep seismic sounding (DSS) studies (Kaila et al.,
1981a; Kaila and Krishna, 1992).
There is a broad relative gravity high over the Satpura
Mountains, and the strong gravity high beneath Navsari, located
on the west coast, is particularly noticeable (Fig. 5). The Satpura
486
H.C. Sheth
73o
75o
-40
-10
22o
-20
+40
-50
Pachmarhi
-70 -80
-70
-40
-60
Narm
-20 -30
-10
0
21o
ada
r
-30 -40
A MOUNTAINS
-60
-50
-50
-30
Tapi Riv
Navsari
Rive
SATPUR
er
-30
-40
-40
-40
-50
-30
-60
-50
-60
-60
20o
79o
77o
-40
-60
-70
-60
Figure 5. Bouguer anomaly map of the
Narmada-Satpura-Tapi region (simplified from Singh, 1998). Contours are in
milligals. A thick (16–24 km) layer of
igneous material sits beneath the entire
area of the map.
-50
-70
gravity high has been interpreted variably in the past. Qureshy
(1971) felt that this indicated a horst-type structure for the Satpuras or was related to the migration of upper-mantle material
into the crust (Moho upwarp?). Verma and Banerjee (1992)
postulated high-density mafic intrusive material at midcrustal
levels. Singh and Meissner (1995) carried out 2D density modeling along the four DSS profiles across the Narmada-SatpuraTapi zone (Fig. 5), proposing an upwarped Moho and igneous
accretion at the base of a thinned crust. Later Singh (1998) carried out 3D gravity modeling and considered the Satpura gravity
high anomalous because a topographically high region ought to
have a low Bouguer anomaly due to expected isostatic compensation. Whereas the DSS results indicated a Moho shallower
than 25 km, Singh (1998) found a normal crustal thickness
(Moho at 38 km) along the west coast of India. This is directly
counter to the earlier beliefs that the crust along the west coast
is thin. Singh also proposed an accreted igneous layer (of 15–
20 km thickness and 3.02 g/cm3 density) at the base of the crust
under the Satpura-Tapi region (Fig. 6). This layer is aligned eastwest, and its thickness varies from 8 km beneath the eastern part
to ~16 km beneath the central part of the region. The thickness
of the layer under Navsari is 24 km, and this greater thickness explains well the gravity high over Navsari. Singh (1998) suggested
that what had been imaged by the DSS studies as the Moho may
actually represent the high-density or velocity discontinuity
where normal continental crust transforms into transitional-type
crust, and concluded that the crust in this region is not thinned
but of normal thickness.
Seismic Tomography: Thin or Thick Crust and Lithosphere?
Seismic tomography does not support the notion of plumerelated thermal damage to the Indian lithosphere. For the southern Indian peninsular shield, the lower-crustal shear wave
(S-wave) velocities are higher than those for the Baltic, African,
and Canadian shields, but the upper-mantle velocities are lower.
Also, compared to the southern Indian shield, the Deccan region
exhibits marginally higher S-wave velocity both in the lower
crust and in the upper mantle (Mohan et al., 1997). Compressional wave (P-wave) analyses have given similar results (Iyer
et al., 1989). The seismic signature of the hypothesized causative
plume head is absent under the Deccan proper, and P-wave data
suggest a lithospheric thickness of at least 300 km under the
Deccan region as well as the southern Indian shield (Ramesh
et al., 1993). Crustal thickness in the Indian shield is normal as
well. Ravi Kumar et al. (2001) determined that the dominantly
Archean crust forming the southern Indian shield has a very simple structure without any prominent intra-crustal discontinuities,
has an average Poisson’s ratio close to 0.25, and is 33–39 km
thick. They found that the predominantly Proterozoic crust forming the northern Indian shield is complex, with several seismic
discontinuities, and has a crustal thickness of >40 km.
Mutter et al. (1988) and Anderson (1994) have argued that
large-volume basaltic provinces and volcanic rifted margins form
where the transition from thick to thin lithosphere is abrupt, because such an abrupt transition sets up high lateral temperature
gradients. These gradients, in turn, induce small-scale convection and rapid movement of mantle material. An abrupt lateral
change in the lithospheric thickness of a plate focuses both strain
and magma ascent. Anderson et al. (1992) and King and Anderson (1995, 1998) have pointed out that every CFB province is
situated on the margin of a Precambrian craton. In contrast to
India, which is made up of several Precambrian cratons, the
North American continent has only one Precambrian craton (the
Wyoming craton), and the only post-Precambrian CFB province
in North America, the Columbia River flood basalt province, is
located along the edge of the Wyoming craton. Clearly there are
fundamental lithospheric controls on the location of CFB, and
these must be discussed by any realistic geodynamic model for
CFB volcanism. Because of the geological and geophysical
characteristics of the west coast region of India (volcanic eruptive centers, dike swarm concentrations, high heatflow, aligned
thermal springs, etc.), many authors have considered it a region
of crustal thinning (e.g., Chandrasekharam, 1985; Sheth, 1999a,b).
From Deccan to Réunion: No trace of a mantle plume
487
Depth (km)
20
Navsari
30
40
Figure 6. 3D model (simplified) of the
postulated igneous layer at the base of
the crust under the Narmada-Tapi region.
The Moho depths (in km) are indicated
by contours. Modified from Singh (1998).
40
40
39
38
40
39
38 39
78
77
41
40
76
40
N
22
o
75
41
74
N
21
o
o N
20
79
73
o
E
o
E
o
E
o
E
o
E
o
E
o
E
Understandably, thinned lithosphere can permit the underlying
mantle to upwell to shallow depths and decompress, resulting
in extensive melting and a thick lava pile like the one exposed
in the Western Ghats. Therefore, the recent findings of Mohan
and Ravi Kumar (2004), that the crust along the west coast is
thick, not thin, are interesting.
Mohan and Ravi Kumar (2004) performed a receiver function analysis of teleseismic data recorded from 1998 to 2003 by
a ten-station network deployed near Bombay (Fig. 7). The network comprised one broadband station and nine short-period
stations spread over an area 50 km × 100 km across. The receiver
function analysis reveals a continental crustal thickness varying
from 36 to 41 km, which is quite “normal” and shows that this
crust cannot be called thin crust. The observed thick crust is
apparently not due to underplating by basaltic magma, because
such mafic underplated material would be identifiable from its
characteristic Poisson’s ratio (the ratio of lateral strain to longitudinal strain in an elastic body due to uniaxial longitudinal
stress). Mohan and Ravi Kumar (2004) obtained a value of 0.26
(± 0.01) for the Poisson’s ratio for the sub-Deccan crust in this
region and suggested a felsic to intermediate composition similar to that of the Precambrian southern Indian Shield. (The value
of this ratio is at least 0.28 for mafic rocks and 0.30 and higher
for ultramafic rocks.) It is likely, however, that if there were
Deccan-related underplating of a thinned crust by Deccan felsic
magmas, the geophysical methods used would not be able to
distinguish such an underplated crust from true basement crust,
given their closely similar densities and Poisson’s ratios. As
noted, Deccan rhyolite and trachyte flows and dikes are indeed
found along the west coast, particularly at Bombay (Sheth et al.,
2001a; Sheth and Ray, 2002). However, these felsic magmas
probably developed in magma chambers established at uppercrustal levels (shallower than ~10 km) (Sheth and Ray, 2002),
and they cannot have underplated a thinned (say 15–20 km thick)
crust all the way to a depth of 40 km, as is required by the data.
Mohan and Ravi Kumar (2004) note that the crustal thickness,
Poisson’s ratio, and average crustal S-wave velocity (3.7 km/s)
are all similar to the values for the Precambrian Indian Shield,
concluding that the crust under the Deccan basalts in this region
has not been affected by the Deccan volcanism. It is likely, therefore, that the concentrations of eruptive centers and thermal
springs, dikes, and intrusions and the high heatflow along the
west coast are related not to crustal thinning but to fracture control in an otherwise thick crust and lithosphere.
The Cambay Seismic Anomaly
Kennett and Widiyantoro (1999) reported low-P-wavevelocity mantle under the northern part of the Cambay Rift and
interpreted it as a remnant of the Deccan plume. This mantle is
488
H.C. Sheth
Mehmadabad
Cambay
22o
Bharuch
me
21
nt
Surat
o
We
ste
rn
Gh
ats
es c
arp
Navsari
Billimora
(20 km)
Panvel
Flexure
20o
SAK
AGH WAS
ALM
SHN
PHA
CPD
MUL
19o
NER
Bombay
KOP
ste
We
Pune
rn G
Arabian
Sea
hats
Loni
2% slower than the seismically fast Indian Shield, defines a
region of roughly circular cross-section that is ~250 km across,
of depths between 80 and 250 km (Fig. 8), and appears to connect with a larger (600 km across) slow region that extends from
250 km down to 500 km depth. Kennett and Widiyantoro (1999)
did not rule out seismic anisotropy as a cause of the low seismic
velocity, but felt that the velocity anomaly could also reflect high
temperatures (supporting evidence for which comes from the
high heatflow along the Cambay Rift). They suggested that the
anomalous mantle region represents a conduit of the DeccanRéunion plume, because some of the earliest Deccan rocks outcrop in alkaline complexes in this region, and these have been
argued to be plume-derived (Basu et al., 1993) because of their
high 3He/4He ratios (up to 14 times the atmospheric 3He/4He
ratios in pyroxenes).
Anderson (2000a) has presented a lucid discussion of current 3He/4He ratio fallacies and a priori data-filtering practices Italics
detrimental to rigorous statistical averaging. 4He is produced aren’t
needed for
by radioactive decay of U and Th, and Anderson argues that foreign
high 3He/4He ratios do not reflect high 3He abundance (as in the terms that
plume model) but rather low 4He (i.e., revealing a source poor have made
their way
in U + Th). This conclusion is supported by Natland (2003) and into EngMeibom et al. (2003), who argue that during crystallization of a lish dicmagma, olivine crystals trap 3He along with CO2 in fluid inclu- tionaries.
sions, and because olivine is a mineral very poor in U + Th, there
is negligible growth of 4He over time. Olivine crystals therefore
act as “He time capsules.” Thus, mantle-derived mafic or ultramafic rocks with large amounts of olivine are likely to have high
3He/4He ratios. 3He/4He ratios in mantle-derived rocks cannot
30o
ent
Kelsi
(31.5 km)
arpm
esc
18o
Guhagar
(30 km)
To Ch
o
r ochi
+1.5%
Figure 7. Map of western India showing three deep seismic soundings
profiles (straight lines, Kaila et al., 1981a,b,c), which showed an apparent thin crust along the west coast; the thickness of this thinned crust
under Billimora, Kelsi, and Guhagar, derived from deep seismic soundings, is shown in italics. The triangles define the seismic network employed by Mohan and Ravi Kumar (2004). The stations and the crustal
thicknesses found under them are as follows: MUL—Mulgaon, 35.8 km;
PHA—Phalegaon, 36.4 km; SAK—Sakur, 37.4 km; SHN—Shenva,
37.6 km; NER—Nere, 38.2 km; ALM—Alman, 38.6 km; CPD—
Chinbipada, 38.7 km; WAS—Washala, 39.5 km; AGH—Aghai, 39.9 km;
and KOP—Koproli, 41.3 km. MUL is a broadband station, and the
other stations are short-period stations.
25o
Cambay
ati
on
75o
urb
74o
pe
rt
73o
-1.5%
72o
0%
17
o
20o
+1.5%
Bombay
Arabian
Sea
70o
75o
80o
Figure 8. Map view of the compressional wave velocities in the depth
interval around 80 km, showing the Cambay seismic anomaly (modified from Kennett and Widiyantoro, 1999).
From Deccan to Réunion: No trace of a mantle plume
be used to infer mantle source ratios, and definitely do not constitute evidence for plumes.
The age of the Cambay seismic anomaly is unknown. Assuming it dates from the time of Deccan volcanism, it is located
where expected for passive models involving rift convection
(e.g., Sheth, 1999b). Kennett and Widiyantoro (1999) infer that
a 65 Ma anomaly would have to be at least 50 km across originally, with a temperature excess of 300 oC or more (given
thermal diffusion over time). Such hot sources arguably would
produce high–melt fraction subalkaline picritic liquids (as the
plume model postulates). However, there is no petrological evidence from the rocks exposed in this region (or indeed anywhere in the Deccan) for such high mantle source temperatures.
Although picrites are encountered in boreholes near Cambay,
these are somewhat alkaline and are low-degree melts (Peng and
Mahoney, 1995). Even the mafic-ultramafic rocks analyzed by
Basu et al. (1993) are low-degree, high-pressure alkaline melts
(as inferred, for example, from rare earth element patterns suggestive of residual garnet).
There are two suggestions as to possible origins of the Cambay seismic anomaly outside of the plume model: (1) the anomaly is of Deccan age and related to passive rifting and related
upper-mantle convection, or, more likely, (2) the anomaly is of
post-Deccan age and much more recent. Kennett and Widiyantoro’s study ran out of resolution for the area along the west
coast south of 20oS, but they considered a southerly extension
of the anomaly likely. If so, it would support the view that the
age of the anomaly is post-Deccan and that the spectacular postDeccan uplift of the Western Ghats and the development of this
seismic anomaly may be related in some way. The Cambay Rift
has received several km of Tertiary sediments and continues to
be a low-lying area today. This may be because crustal extension
removes the need for uplift. Biswas (1982) has considered the
Cambay Rift a true active rift in India today. If it is one, the Cambay low-velocity anomaly may represent warm, expanded, relatively less dense mantle that has risen to depths of 80–100 km
because of extension in the overlying lithosphere.
We shall now leave the Deccan and India behind as we
enter the Arabian Sea and the Indian Ocean.
THE LAKSHADWEEP-CHAGOS RIDGE
The Lakshadweep-Chagos Ridge is a linear north-south
ridge that runs for 2200 km in the western Indian Ocean (Fig. 9)
and has been postulated to have been produced by the “tail” of
the Réunion plume after the head of the plume was consumed
in the production of the Deccan Traps. The Lakshadweep islands
at the northern end of the ridge are capped by coral reefs. The
ridge takes off where the crudely circular outcrop of the Deccan
lavas ends, and the pair has been considered a classic example
of a plume head and tail. However, new Ar-Ar data have shown
that the systematic southward age progression required in the
plume model does not apply. Sheth et al. (2001a,b) dated three
489
Deccan rocks from Bombay at 19oN by the 40Ar-39Ar technique. Two of them are trachyte lava flows and yielded ages of
60.4 ± 0.6 Ma (2σ) and 61.8 ± 0.6 Ma (2σ), respectively. Their
third sample came from the well-known, thick (>25 m), and
columnar-jointed Gilbert Hill basalt and was dated at 60.5 ± 1.2
Ma (2σ). The very well-developed plateau spectra, isochrons
with good mean square of weighted deviates (MSWD) values,
and atmospheric 40Ar/36Ar intercepts indicate that these dates
are reliable crystallization ages. But these ages are troublesome
for the plume model because, according to the model, the plume
head was consumed at 65–66 Ma and only the narrow plume tail
remained. How can a plume tail 100–200 km wide have produced volcanism simultaneously in Bombay and at Ocean
Drilling Program (ODP) site 715, two locations that are 1000
km apart (Fig. 9)?
Some workers have suggested that the 60–61 Ma volcanic
activity in Bombay was of minor volume and hence not problematic (e.g., Mahoney et al., 2002; Courtillot and Renne, 2003).
Courtillot and Renne (2003) noted the young, 60–61 Ma, ages
of Sheth et al. (2001a) but stated that Sheth et al. (2001a) had
dated trachyte dikes and that such late-stage dike activity was
expected and should not render the plume model questionable.
Jerram and Widdowson (2005) also imply that the dated trachytes were intrusions. Sheth et al. (2001a) clearly stated that
they were dating trachyte flows, and the Ar-Ar ages on these dipping lava flows provided age constraints on the formation of the
Panvel flexure (Sheth, 1998), of which they are a part. In any
case, we do not know that this volcanic activity was minor in
volume: large volumes of Deccan lava exist in the subsurface
along the west coast, and there is a scarcity of geochronological
data; the three Ar-Ar dates of Sheth et al., (2001a,b) are the only
existing ones for this region of the Deccan. In comparison, the
Western Ghats section has been heavily sampled and dated. Finally, whatever its magnitude, the late-persisting volcanism
must be still explained without ad hoc auxiliary hypotheses. But
it is not. Suggestions such as northward dragging of the plume
tail by the plate (made by several colleagues in personal commun.) are ad hoc, and such drag and tilting would make any systematic age progression impossible in the first place.
Interestingly, there was Deccan-age magmatism in the ca.
116 Ma Rajmahal Traps of eastern India (Figs. 1 and 9). Kent et
al. (2002) dated a ferro-tholeiite dike in the Rajmahal Traps at
65.4 ± 0.3 Ma (2s). If Deccan volcanism was caused by a plume
under western India, how did it generate magmatism 1500 km
away? Long-distance mantle flow or magma flow may be invoked, but this also significantly reduces the value of a fixed
plume model. Besides, there are mafic dikes in Kerala, southwestern India (Fig. 9), dated at ca. 69 Ma (Radhakrishna et al.,
1994), and if they are a part of the Deccan event, which seems
likely, they do not support a systematic southerly age progression from the Deccan. They also complicate the spherical plume
head–narrow plume tail picture by virtue of their location (note
that the Lakshadweep-Chagos Ridge departs from the western
490
H.C. Sheth
50
o
60
o
70
South Tethyan
Suture Zone
80
HIM
Z
72-73
o
90
ARABIA
20o
Arabian
Sea
OFZ
Camba
Br
68.5 M
Kachchh Rift
D
65
60-61
AY
AS
da Rift
Narma
Ma
Rif hana
t
di
Tapi Rift
La
xm
iR
Go
Rif dava
t
ri
64
id
ge
61 G
SMI 85.5
CA
SB
RI
ER
AF
G
RI
DG
E
0o
IAN RID
CENTR
40
E
T RIDG
Rodrigues Ridge
8-10
0-2 Reunion
N RIDG
AN
RI
AN
DI
IN
T
HW
ES
BROKE
DI
E
E
DG
RI
UT
IN
SO
T
AS
HE
UT
o
Y EAS
0.17-8 Mauritius
DG
E
SCA
MAD
AGA
0
ODP Site 713
(49)
MFZ
30
20
10
VFZ
SO
20
Ma
s
Pla care
tea ne
u
ODP Site 706
(33)
R
10
AL IND
ODP Site 707
(64)
Site SM-1
(45)
o
60
ODP Site 715
(57)
50
GE
63
Seychelles
KK
Figure 9. Prominent structural-tectonic
features of southern Asia and the Indian
Ocean Basin (based on Mahoney et al.,
2002). Abbreviations for localities are
as follows: Q—Quetta; Z—Zhob; B—
Barmer; M—Mundwara; D—Dhandhuka; B—Bombay; R—Rajahmundry.
WG indicates the Western Ghats region
(ages from Venkatesan et al., 1993 and
others). Ages (in Ma) are in boldface
numbers. The ca. 64 Ma age for the
Rajahmundry basalts is from Baksi
(2001b). G indicates the ca. 61 Ma Goa
dikes (Widdowson et al., 2000). KK indicates the ca. 90–69 Ma KarnatakaKerala dikes (e.g., Radhakrishna et al.,
1994; Anil Kumar et al., 2001). SMI indicates the 85.5 Ma Saint Mary’s Islands
volcanics (Pande et al., 2001), part of
the Indo-Madagascar continental flood
basalts, which in India are otherwise represented by the KK dikes. The 72–73 Ma
ages for the Quetta and Zhob rocks and
the 65 Ma age for the Dhandhuka-Botad
lavas are from Mahoney et al. (2002), as
is the modeled hotspot track showing
expected ages in Ma (italics). Note the
rift zones underlying the Deccan and the
absence of any triple junction. OFZ—
Owen fracture zone; MFZ—Mauritius
fracture zone; VFZ—Vishnu fracture
zone.
NINET
CA
RL
Bay of
Bengal
R
90-69
LAKSHADWEEP-CHAGOS RIDGE
10
o
RAJMAHAL
TRAPS
~ 116
DECCAN
TRAPS
67-62
WG
B
o
Indus-Tsangpo
Suture Zone
AL
y Rift
Q
o
Indian coast far north of the Kerala dike occurrences). These
dikes support “passive,” rifting-related volcanism.
A model for the origin of the Lakshadweep-Chagos Ridge
must address the following:
1. Visually, the Lakshadweep-Chagos Ridge, Mauritius, and
Réunion together do not form a picture similar to that of the
Hawaiian chain.
2. The Ridge runs along the Vishnu fracture zone (Fig. 9).
3. Baksi (1999 and this volume) has critically evaluated the
Ar-Ar ages of the basalts forming the Ridge (and those of
other “hotspot tracks” in the Atlantic and Indian Oceans)
and has stated that out of about thirty-four published (and
routinely used) ages, only three satisfy the statistical criteria for acceptable ages. If he is correct, we do not know
if the perceived age progression along the LakshadweepChagos Ridge is even real.
4. In the fixed plume model, the 33 Ma date for ODP site 706
From Deccan to Réunion: No trace of a mantle plume
A
-10°
(33 Ma)
(57 Ma)
715 (64 Ma)
706
Palaeolatitude
707
-20°
Reunion
-30°
Deccan
-40°
0
10
20
30
40
50
Age (Ma)
60
B
70
80
CRACK
Time T3
0°
10° S
Volcano 1
(extinct)
CRACK
Time T2
20° S
Volcano 2
(extinct)
Time T1
40° S
Volcano 1
(extinct)
Volcano 2
(active)
Tip
60° S
Volcano 1
(active)
Volcano 3
(active)
PLATE
50° S
PLATE
30° S
CRACK
on the “track” on the Africa Plate, across the Southwest
Indian Ridge, has been considered evidence that a northerly
jump of the Central Indian Ridge transferred the Réunion
hotspot under the Africa Plate. The alternative explanation
is that the Central Indian Ridge jumped over a southwardpropagating crack (track) at 30 Ma and split the track.
5. The 450 km–long Rodrigues Ridge that lies east of Mauritius is oriented at right angles to the predicted hotspot track.
The Rodrigues Ridge was hypothesized to have been produced by lateral flow of Réunion plume material as the
Central Indian Ridge moved away from it, and it has no
age progression itself (the whole ridge is dated at 8–10 Ma;
Duncan, 1990).
6. The geochemistry of the basalts along the track indicates
mixing of “plume” and “MORB” mantle (Fisk et al., 1988;
White et al., 1990), or, alternatively, these basalts are not
shield-stage rocks but postshield basalts (Sheth et al., 2003,
and later discussion). In ocean islands, the postshield and
posterosional (rejunvenated) stages of volcanism occur 1 m.y.
to a few m.y. after the main shield-building stage (e.g.,
Clague, 1987). If these basalts are not shield-stage, they
cannot be used to locate the plume and their ages cannot be
used for plate reconstructions based on hotspots.
7. The paleolatitudes of the basalts along the chain are variable and do not themselves support the plume model without the introduction of further ad hoc parameters (see later
discussion).
491
THE PALEOLATITUDE VARIATION:
MANTLE ROLL OR CRACK PROPAGATION?
The Kerala mafic dikes (Radhakrishna et al., 1994) complicate the simple picture of a plume head (Deccan Province)
and a plume tail (Lakshadweep-Chagos Ridge). The roughly
circular outcrop of the Deccan does not require or indicate a
spherical plume head beneath, but may be a consequence of the
intersecting rift zones (the western Indian margin, the Cambay
Rift, and the Narmada-Tapi Rift). Sheth (1999a) argued that the
age progression along the Lakshadweep-Chagos Ridge and up
to Réunion Island may be explained by southward crack propagation through the oceanic lithosphere and that the narrow
“hotspot track” may represent localized melting and magma
focusing from a wider area (the “transform-fault effect”; Langmuir and Bender, 1984). In support of this is the fact that the
Lakshadweep-Chagos Ridge lies along the Vishnu fracture zone
(Fig. 9). It is not necessary to explain how a crack would have
propagated across two separate plates, because it is quite possible that the current volcanism at Réunion Island is unrelated
to the Deccan and to the Lakshadweep-Chagos Ridge geodynamically, though it taps delaminated Indian continental mantle brought beneath the Africa Plate by a ridge jump at ca. 30 Ma
(see also Burke, 1996, for a similar interpretation).
In my view, relating the Deccan and Réunion to each other,
and both to a mantle plume, has created more problems than it
80° S
PLATE
70° S
Figure 10. (A) Palaeolatitude variation from the Deccan to Réunion Island through the Ocean Drilling Program leg 115 sites (Vandamme and
Courtillot, 1990). (B) Schematic cartoon showing the development of
volcanism resulting from a crack’s propagating more slowly southward
than the plate moves northward in the southern hemisphere.
has solved. The Deccan lavas erupted at a latitude of ~30°S
(Clegg et al., 1956), but Réunion Island is at 21° S today (Fig. 9
and 10A). This large discrepancy has promoted further ad hoc
speculation. Vandamme and Courtillot (1990) proposed true
polar wander (TPW) of the Earth’s mantle. In their view, subsequent to the Deccan eruptions the Réunion plume remained
fixed in the mantle while the mantle itself “rolled” as a ball inside the lithospheric shell, in a northerly direction. According
to these workers, this has brought the Réunion plume northward
from 30oS to 21oS (a distance of ~1000 km) over 65 m.y. Burke
(1996) questioned this postulated TPW and suggested that the
Deccan plume died out at 30 Ma and that the Réunion plume is
new and unrelated to the Deccan plume.
A simpler alternative is illustrated by the schematic cartoon
in Figure 10B. This alternative is that the systematically changing paleolatitudes between the Deccan and ODP site 706 (33 Ma)
492
H.C. Sheth
indicate southward crack propagation in a northward-moving
plate with the condition that the northward plate motion was
faster than the southward crack propagation. The situation shown
in Figure 10B is for a plate in the southern hemisphere. At time
T1, there was an active volcano 1 at the crack tip at latitude
60° S. Between times T1 and T2, the crack tip moved southward
by 10°, but because the plate itself had moved north by 20°, the
new volcano 2 at the crack tip had the latitude of 50° S. A similar progression occurred between times T2 and T3. Therefore,
although the crack tip is propagating southward, the paleolatitudes systematically become more northerly. Thus, there
is not necessarily a contradiction between southward crack
propagation and progressively northerly paleolatitudes, and the
Lakshadweep-Chagos ridge is better interpreted as a crack than
as a track. The ridge may mark the location of a major Gondwanic transform (Reeves and de Wit, 2000; Reeves et al., 2004).
of Réunion developed atop a Paleogene fossil accretionary center. The same was also suggested for Mauritius Island, i.e., Mauritius and Réunion are both fossil accretionary centers that were
active on different sides of a triple junction and were carried
away from each other along the later fracture zone in between.
The seismic data available support the idea that the location of
Réunion is related to the structural heterogeneity of the underlying lithosphere (Charvis et al., 1999; de Voogd and Pontoise,
1999).
The Deccan and Réunion Island are probably unrelated. It
is true that some Réunion-type chemical and isotopic signatures
are found in some lavas in the Deccan (e.g., Peng and Mahoney,
1995) and in ophiolitic rocks predating the Deccan (Mahoney
et al., 2002). However, these signatures may be continental mantle signatures.
GEOCHEMICAL-ISOTOPIC DATA
RÉUNION, MAURITIUS, AND RODRIGUES
Burke (1996) has questioned the Deccan-Réunion link. According to him, Réunion, Mauritius, and Rodrigues are members
of the youthful (<30 Ma) population of hotspots on the Africa
plate. He gave evidence that that the Africa plate came to rest
over the underlying mantle circulation at 30 Ma and has moved
little since. If so, the manifestation of the Deccan hotspot should
have remained in the same place over the last 30 m.y. However,
there has been no volcanism in this area (the Saya de Malha Bank,
Nazareth Bank or Cargados Carajos Bank) since 30 Ma. Burke
therefore argued that the Deccan plume died out at 30 Ma, that
it was a different plume from the Réunion plume, and therefore
that the position of Réunion today at 21oS did not provide evidence of true polar wander.
This suggestion is thought-provoking. The first volcanic
activity in the area after ca. 30 Ma clearly began at 8–10 Ma
(Mauritius Shield and Rodrigues stage). Thus, in the Réunion
plume model, the plume must have been inactive for 20 m.y.
Burke (1996) suggested that each of the three Mascarene islands
(Réunion, Mauritius, and Rodrigues) is underlain by a young
plume. However, the distances between these islands, if they are
underlain by deep plumes, are too small for each of them to have
a CMB origin (Anderson, 1998b). Mauritius and Réunion are
located along the Mauritius fracture zone (Fig. 9), and a small
seamount west of Réunion (Fig. 9) (Duncan, 1990) complicates
the picture. Hirn (2002) has discussed structural studies of
Réunion Island using seismic methods and noted significant departures from the Hawaiian case, to which it is traditionally
compared. There is no moat around Réunion of the type that surrounds a Hawaiian Shield volcano, but crustal doming is seen
instead. Two radial reflection lines to the SSW, close to each
other, detect a difference in depth of the oceanic basement, and
this may be related to the presence of a fracture zone, suggested
from the magnetic anomaly pattern seen from reconstruction of
the seafloor spreading history. The magnetic anomaly pattern
previously has been interpreted to indicate that the western part
In many ocean island volcanoes worldwide (e.g., Hawaii,
Mauritius), shield-stage lavas that make up most of the volcano’s volume (99%) have systematically greater 87Sr/86Sr ratios and lower 143Nd/144Nd ratios than post-shield-stage and
post-erosional-stage (rejuvenated) lavas (e.g., Clague, 1987). It
is the shield-stage lavas that, in the framework of the plume
model, are derived from plumes, and by dating the shield stage
one can infer the location of the plume at that time. Although the
ages of the ODP leg 115 (Indian Ocean) basalts have been used
by plate motion modelers such as Müller et al. (1993), almost all
of these ages were questioned by Baksi (1999). A further problem is that the close Nd-Sr isotopic similarity of some of the
ODP leg 115 basalts to post-shield lavas on Mauritius (the Intermediate and Younger Series) (Fig. 11) suggests that the leg 115
basalts could similarly be post-shield lavas (Sheth et al., 2003).
If so, their ages cannot be used for modeling plate motions based
on the plume model, because the age of the shield stage would
remain unknown (it would be variably older.) We cannot compare these ODP leg 115 basalts to the modern Réunion basalts,
because Réunion Island is currently only in the active shieldbuilding stage. White et al. (1990) argued for progressively more
“enriched” geochemistry southward along the Réunion plume
track, based on Nd and Sr isotopic compositions. Their data
points define a gentle, broad array in each isotopic plot (see their
Fig. 5). But if the analytical error bars are considered, and also
the fact that post-shield volcanism on Réunion has yet to take
place, the array in each plot may be statistically no different
from a flat array, and the time-progressive systematic variation
in isotopic characteristics that White et al. (1990) argued for is
not apparent.
Are mantle plumes the cause of the “enriched” geochemistry? Smith (1993) proposed that ocean island volcanism
is derived from enriched continental mantle, often a part of ancient sutures, delaminated from beneath a breaking continent
and dispersed in the new oceanic mantle. “Enriched” isotopic
ratios, such as higher-than-normal-MORB values of 87Sr/86Sr,
From Deccan to Réunion: No trace of a mantle plume
493
+11
+10
Older Series
Intermediate Series
Younger Series
Site 707 (64 Ma)
Site 715 (57 Ma)
Site 713 (49 Ma)
Site SM-1 (45 Ma)
Site 706 (33 Ma)
Carlsberg Ridge MORB
+9
Ce
n
Rid tral In
ge
d
MO ian
RB
εNd
+8
+7
+6
+5
Reunion
+4
+3
0.7026
0.7028
0.7030
0.7032
0.7034
0.7036
0.7038
0.7040
0.7042
Figure 11. Nd-Sr isotopic ranges of the
Ocean Drilling Program (ODP) leg 115,
Mauritius and Réunion lavas (Sheth et al.,
2003). All data are present-day values
(not age-corrected). The data sources for
Mauritius are Mahoney et al. (1989),
Peng and Mahoney (1995), and Sheth
et al. (2003). The fields for the Central
Indian ridge and Carlsberg ridge midocean ridge basalts (MORB) are based
on Ito et al. (1987) and Mahoney et al.
(1989); the data for the Réunion basalts
are from Fisk et al. (1988); and the data
for the ODP leg 115 and Texaco drill site
SM-1 are from White et al. (1990). The
ages for the leg 115 and SM-1 basalts are
from Duncan and Hargraves (1990).
0.7044
87Sr/86Sr
are usually taken as plume signatures. However, such compositions may instead mark the involvement of shallow-level enriched
continental mantle. High values of 3He/ 4He may be explained
by shallow models (e.g., Anderson, 2000a).
The (enriched) plume model is not required to explain continental intraplate volcanism, given the abundance of enriched
mantle domains within the continental lithosphere itself. The
plume model was extrapolated to continental magmatism from
the ocean basins based on the worldview that the oceanic
mantle was entirely “depleted,” MORB-like, convecting, and
homogeneous. As a consequence, anything enriched or anomalous was assumed to come from plumes (Smith and Lewis,
1999). However, if continental mantle can be introduced into
the oceanic mantle, e.g., by delamination during continental
breakup (e.g., Smith, 1993), enriched plumes are not required
to explain either continental or oceanic intraplate volcanism,
and enriched “plums” of continental mantle dispersed within the
oceanic mantle would constitute a better scenario. The whole
plume argument is unnecessary, then, from a geochemistry point
of view. Rather than accepting that the Deccan formed from a
deep mantle plume now located under Réunion Island, we can
now see that Réunion volcanism may be in part sourced from
delaminated Indian continental mantle. Réunion Island is currently in the active shield-building stage. Modern Réunion basalts
are closely similar isotopically to the Older Series (shield-stage)
basalts of Mauritius (Mahoney et al., 1989; White et al., 1990;
Sheth et al., 2003). Both the plume model and the “plum” or
“blob” model (e.g., Sleep, 1984) can explain this.
Mahoney et al. (2002) reported Réunion-like elemental and
isotopic compositions for mafic ophiolitic rocks in Pakistan dated
at 72–73 Ma (Fig. 9). They proposed that some of these represent pre-Deccan oceanic seamounts. The associated intrusions
were emplaced in continental shelf-and-slope-type marine sediments along the northern margin of India. Mahoney et al. (2002)
considered the continental mantle delamination model, but felt
that it did not explain Réunion-type volcanism occurring on
the updrift side of India at 72–73 Ma. They concluded that the
plume model was the most viable option.
Nevertheless, the intrusions in question are located within
the boundary of the Indian continental mantle, and the purely
oceanic seamounts may not have been far from the northern
margin of India. Delaminated continental mantle could have
migrated northward ahead of India in a radial, outward extrusion and fed the seamounts built on oceanic lithosphere. The
continent followed behind, and when it converged upon Asia it
trapped them along the suture. This is a better explanation for
the observations than the plume-head-impact (not incubation)
model favored by Mahoney et al. (2002), despite the ~8 m.y.
time gap between the Pakistani rocks and the 66–65 Ma voluminous volcanism in the Deccan. Mechanisms such as lateral flow
are required even by the plume model, e.g., for the Rodrigues
Ridge (Fig. 9).
DISCUSSION AND CONCLUSIONS
The traditional mantle plume model for CFB like the Deccan suffers from many contradictions and shortcomings. Many
geological and geophysical features of the Deccan have generally been considered to require a mantle plume origin. Singh
(1999), for example, discussed geophysical data for the Laxmi
Ridge in the northeastern Arabian Sea (Fig. 9). The Laxmi Ridge
is 700 km long and 100 km across, parallels the western Indian
continental margin, and consists of isolated submarine structural
highs. Singh related the Laxmi Ridge to Deccan volcanism and
494
H.C. Sheth
noted that both active and passive models were available that
could explain the geophysical observations. He opted for the
plume model, however, “there being ample independent evidence for a plume origin.” The aim of the present paper is to
show that no evidence requires a plume. Rather, a “passive,”
nonplume model, such as rifting and related mantle convection,
explains the observations equally well, if not better, and is free
of the contradictions and problems of the plume model. A massive amount of decompression melting would have taken place
when the Central Indian Ridge jumped onto the northerly drifting Greater India, splitting the Seychelles (Fig. 1). This can
also explain why the Deccan basalts closely resemble mid-ocean
ridge basalts in many elements, a long-noted fact (e.g., Chandrasekharam and Parthasarathy, 1978; Shrivastava and Pattanayak,
1995). The magma type from the Deccan that is least contaminated by continental material is the Ambenali, and the Ambenali
basalts are evolved ferrobasalts (e.g., Mahoney, 1988). The transitional MORB characteristics (e.g., a small Nb peak in a normalized multielement pattern) of the Ambenali do not require
an “enriched” mantle plume mixing with entrained or ambient
MORB source mantle. They have been proposed to reflect a
slight amount of contamination by phlogopite-bearing peridotite
(Sheth, 1999a), or might represent shallow-residing ancient
eclogite (Sheth, 2005). When a mature, steady-state spreading
ridge suddenly jumps to a new location, strong lateral pressuretemperature gradients must be created in the mantle relatively
instantaneously. A quick burst of voluminous volcanism would
be a natural outcome, because a focused flow of mantle and partial melt would be distributed over a wide region toward the new
ridge. A large plume “head,” invoked in current plume theories,
is not indicated or required.
Crustal Extension and Rift Convection
Lithospheric extension was operative before, during, and after Deccan flood volcanism (Sheth, 2000). Geological and geophysical data suggest that intense eruptive and intrusive activity
occurred along the rift zones and the new rifted continental margin. The Deccan lavas were produced under the imminent continental margin, along the boundary between thinned-extended
and thick crust, and flowed on thick cratonic lithosphere immediately adjacent to the margin. More supporting evidence for
this is the relative scarcity of dikes in the Deccan in the more interior parts (eastern and southeastern) of the province, whereas
dikes and intrusions are far too numerous along the west coast
and along the Narmada-Satpura-Tapi zone and the Kachchh
Rift (they are buried beneath Tertiary sediments in the Cambay
Rift). Also, in the eastern and southeastern parts of the Deccan
province the lava pile is very thin (~400 m maximum, including exposed and subsurface portions), compared to the massive
1700 m exposed along the Western Ghats, with an additional
500 m or so in the subsurface (e.g., Tiwari et al., 2001, and references therein). A huge thickness of the lavas along the Ghats
is exactly what is expected in the rift convection model. (Note
that the present position of the Western Ghats escarpment,
~50 km east of the coast, is a result of erosion-induced parallel
scarp retreat; e.g., Ollier and Powar, 1985; Widdowson and
Mitchell, 1999). The role of rift convection in causing Deccan
volcanism was evidently not confined to the west coast alone but
also significant along the Deccan rifts.
Besides, we have noted that, contrary to earlier views (based
on DSS) that the continental crust along the western Indian coast
is greatly thinned, recent gravity modeling and seismic tomographic studies indicate a continental crust of normal thickness
(36–40 km) along the west coast (Mohan and Ravi Kumar,
2004). If, as in the edge-driven convection model of King and
Anderson (1995, 1998), rifted-margin flood basalts like those
of the Deccan largely form and erupt where cratonic lithosphere
suddenly changes thickness, it is important to explain how this
lithospheric asymmetry forms in the first place. The juxtaposition of thin and thick crust along the present west coast of India
is not difficult to explain, because, as already noted, prior to the
Deccan episode, Greater India (India plus the Seychelles) broke
off from Madagascar at ca. 85 Ma along the coast (e.g., Storey
et al., 1995; Anil Kumar et al., 2001; Pande et al., 2001; Fig. 1).
This does not mean, of course, that the crust was thinned. It
is possible that strike-slip motions were involved in IndiaMadagascar separation (e.g., Chand and Subrahmanyam, 2003;
Raval and Veeraswamy, 2003), and it is not known whether
crustal thinning was involved besides crustal extension and fracturing along the then–western continental margin of India. King
and Anderson (1995) argued that CFB events may be the products of lithospheric splitting along preexisting discontinuities
rather than of plume-caused uniform lithospheric thinning or
distributed stretching. The splitting of the lithosphere permits
adiabatic ascent from great depth (>150 km) and extensive melting and yields a high melting column.
The present Indian continental shelf, where many horstgraben complexes with basement ridges and Deccan lava outliers occur, is a region of thinned, extended crust that subsided
below sea level subsequent to Deccan volcanism and has received
several kilometers of Tertiary sediments, and the present west
coast of India is apparently the eastern limit of this stretched
and thinned crust. Thus, lithospheric extension during the IndiaMadagascar breakup event, and also immediately preceding the
Deccan event, created the pronounced lithospheric asymmetry.
The strongly structure-controlled volcanism itself occurred at
the time of the ridge jump that India experienced (Fig. 1).
An Asteroid Impact?
Why should the ridge have jumped? Plume proponents
would propose that the impingement of a plume head at the base
of the lithosphere weakened the lithosphere and provided the
stress necessary for rifting it, which would also explain, for them,
Deccan flood volcanism itself. An alternative active mechanism
From Deccan to Réunion: No trace of a mantle plume
495
Sie
rra
Ma
25o N
dre
ra
Or
r
Sie
ien
tal
en
cid
Oc
m
Ta
o
ay
FZ
Ce
20o N
TZ
Co
R
MG
Mexico
City
ChR
rra
East Pacific
Rise
COCOS
PLATE
Ma
dre
de
lS
MID
AM
ro
zc
o
FZ
DLE
O
110o W
Veracruz
Sie
Rivera FZ
PACIFIC
PLATE
Figure 12. The tectonic setting of the
Mexican volcanic belt (shaded). FZ—
fault zone; MG—Mesozoic granitoids;
TZR—Tepic-Zacoalco Rift; CoR—
Colima Rift; ChR—Chapala Rift. The
boundaries of proposed terranes forming
southern Mexico are also shown as dashed
lines. Based on Sheth et al. (2000).
SLP
Guadalajara
R
RIVERA
PLATE
Gulf of
Mexico
tal
dre
Ma
Altiplano
100o W
ERI
0
ur
CA
200
400
km
TRE
NCH
to a plume is asteroid impact. Hartnady (1986) proposed a Late
Cretaceous asteroid impact in the Indian Ocean Basin, and Chatterjee (1992) proposed that this impact resulted in a large impact
crater (which he coined as the “Shiva Crater”) off western India
and the jump of the Central Indian Ridge to the new location,
which caused the Seychelles to split off from India. The Shiva
Crater, according to Chatterjee, was split in two by the ridge, and
the two halves were carried away from each other by subsequent
seafloor spreading. Large asteroid impacts have been proposed
to have caused the formation of flood basalt provinces, and
quantitative modeling of the process has been performed (Jones
et al., 2002). An impact model has also recently been proposed
for the Ontong-Java Plateau (Ingle and Coffin, 2004), because
several first-order features of the plateau are at variance with
what is expected from a mantle plume-head origin (Tejada et al.,
2004). However, ridge jumps need not be caused by “active” upwellings from below, but may be related to the dynamics of the
plates themselves, including evolving plate boundaries, continental collisions, and crustal thickening.(e.g., Hamilton, 2002).
If so, it is of interest to note that the Indian shield is supposed to
have already contacted Asia by 70 Ma or so (Jaeger et al., 1989),
and the early stages of the collision with Asia may have changed
the long-distance stress field across the plate.
Early Alkaline Magmatism Due to Incipient Ridge Jumps?:
The Deccan and Mexico
Important alkaline mafic magmatism, though comparatively
of smaller volume, took place before, during, and after the main
flood basalt phase (e.g., Basu et al., 1993; Sheth, 1999b; Ray et al.,
2003) and was apparently restricted to the rift zones, which were
in existence well before volcanism and contained thick Mesozoic sediments. On the other side of the globe, the Mexican
volcanic belt provides an interesting comparison (Fig. 12). This
is a linear belt ~1000 km long, and Miocene to Recent in age,
with many monogenetic cone fields and several active andesitic
stratovolcanoes (Verma, 2001, 2002). Ocean island basalt (OIB)–
like alkaline magmatism is relatively low in volume compared
to the widespread andesitic volcanism, but is found throughout
the belt. There is a prominent triple junction in the western part
of the belt (the Guadalajara triple junction), and both the triple
junction and the OIB-type magmatism in the belt have been
ascribed to a mantle plume (e.g., Márquez et al., 1999). Some
others (e.g., Sheth et al., 2000) propose the origin of the OIBtype magmatism in an enriched sub-Mexican mantle lithosphere.
Earlier Luhr et al. (1985) proposed that the recent and ongoing
magmatism in western Mexico is related to abortive attempts of
the East Pacific Rise to relocate onto the continent. They did not
conceive an “active” cause. The Deccan-Indian scenario prior
to the flood basalt episode is strikingly similar—several rifts that
hosted thick sediments and in which low-volume, OIB-type
magmas were emplaced (Sheth, 1999b). Can these rifting events
before the flood volcanism be related , then, to incipient jumps
of the Central Indian Ridge onto the Indian subcontinent, as has
been argued for western Mexico? If so, it was only the last successful event that eventually split the Seychelles from India.
In summary, the Deccan volcanic episode was significantly
controlled by lithospheric structure and was the end product of
long-duration continental rifting and alkaline magmatism, followed by full continental breakup and decompression melting
496
H.C. Sheth
volume). I also thank an anonymous reviewer and Gill Foulger
for helpful comments on the present manuscript. This work was
supported in part by Research Grant 03IR014 from the Industrial Research and Consultancy Centre (IRCC), IIT Bombay.
RIFT
sediments
A
CRATON
intrusions & magma chambers
Enriched Mantle (EM)
Depleted Mantle (DM)
FLOOD BASALT PILE
B
EM
DM
DM
C
EM
DM
D
Mid-Ocean Ridge
oceanic crust
pure DM
Figure 13. Cartoon depicting (A) the formation of a rift and attendant
sedimentation, with (B) attendant mantle metasomatism and alkaline
magmatism with ongoing lithospheric extension, followed by (C) neartotal rifting and flood basalt volcanism, and eventually (D) seafloor
spreading and production of pure mid-ocean ridge basalts (MORB) at
a mid-ocean ridge. Early, pre–continental flood basalt alkaline magmatism derives from the EM; flood basalts are produced in the DM and
contaminated by what remains of the EM (and the continental crust).
Pure MORB erupts only when the EM is totally consumed and seafloor
spreading has become mature and steady-state (after Sheth, 1999b).
(Fig. 13). A “passive” model of rifting-induced convection fits
the observations best.
ACKNOWLEDGMENTS
I dedicate this paper to Don Anderson for his monumental contributions to geodynamics, his efforts for a plume-free planet,
and the inspiration he has provided. Participating in the 2003
Geological Society of America Penrose Conference, Plumes IV:
Beyond the Plume Hypothesis, in Iceland, was invaluable. I am
grateful to the conference conveners, Gillian Foulger, James
Natland, and Don Anderson, for inviting me to it, and to the GSA
and the International Association of Volcanology and Chemistry
of the Earth’s Interior (IAVCEI) for the financial support that
enabled my participation. Interaction with the conference delegates was an enriching experience and a pleasure. I also thank
G. Mohan for discussions and Mike Widdowson, Godfrey Fitton, Gill Foulger, and an anonymous referee for helpful review
comments (on a different manuscript submitted earlier to this
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