Download The evolving nature of terrestrial crust from the Hadean, through the

Precambrian Research 258 (2015) 48–82
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Precambrian Research
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Review
The evolving nature of terrestrial crust from the Hadean, through the
Archaean, into the Proterozoic
Balz S. Kamber ∗
Department of Geology, School of Natural Sciences, Trinity College Dublin, Ireland
a r t i c l e
i n f o
Article history:
Received 11 September 2014
Received in revised form 1 December 2014
Accepted 18 December 2014
Available online 27 December 2014
Keywords:
Meteorite bombardment
Stagnant lid and mantle overturn
Mafic crust and volcanic resurfacing
Mantle potential temperature
Basalt barrier and transition zone
Mantle refertilisation and crustal growth
models
a b s t r a c t
An outstanding feature of the Archaean Eon is that it was a time of major production and preservation of
continental lithosphere. Here I review the geological, geochemical and basic geophysical data that hold
key information regarding Archaean crust formation and preservation. This insight is then contrasted with
the data for the preceding Hadean and following Palaeoproterozoic, both often portrayed as geological
times of apparently much poorer crust preservation.
It is concluded that two different paths led to formation of preserved Archaean continental crust.
The first was the stabilisation of lithosphere at the Hadean–Archaean boundary, eventually giving rise
to the long-lived, relatively slowly maturing, thick cratonic nuclei found today in several Precambrian shields. The second type of Archaean continental lithosphere formed much more rapidly, over
periods of 150–300 Ma, at several times during the Archaean. Magmas were initially erupted as stacks
of mafic–ultramafic volcanic plateaux that became so thick that they underwent internal differentiation.
No Archaean oceanic lithosphere containing evidence for spreading is preserved, but the existence of
oceanic basins underlain by relatively thin lithosphere is inferred from a number of observations.
An important aspect of studying Archaean lithosphere is to differentiate between the effect of higher
radioactive heat production on the crust as opposed to the mantle. In the crust, the strong distillation of K,
U and Th into the upper layer, via repeated re-melting, caused non-uniformitarian geological phenomena.
Once heat production was concentrated into the top layer, the crust became mechanically strong and its
geology more familiar. The effect of higher heat production on the mantle is much less well understood.
My hypothesis is that the temperature of the Archaean convecting asthenosphere was not uniformly hotter than at present. Rather, very hot (ca. 1750 ◦ C) upwellings from the lower mantle passed through much
cooler (ca. 1400 ◦ C) upper mantle, or the asthenospheric mantle temperature swung rapidly between the
two thermal states.
The defining feature of the Archaean mantle could have been a barrier layer in the transition zone. Hot
Archaean mantle upwellings originated at the base of the barrier, which acted as a thermal boundary
layer. The disappearance of this mechanical barrier in the mantle transition zone could have marked
the end of the Archaean. The ensuing pulse of re-fertilisation of the highly depleted asthenosphere and
the more stable thermal state of the Palaeoproterozoic mantle caused the disappearance of previously
widespread highly magnesian lavas, with wide-reaching consequences for petrology of terrestrial magmas, the structure of the continental crust, and the state of the hydrosphere and atmosphere. If the upper
mantle was refertilised between 2.5 and 2.6 Ga, the geochemical and isotopic evidence for muted continental growth in the Palaeoproterozoic is misleading because almost all crustal growth models assume
a constant mass of the depleted mantle. Until the driver for the Archaean–Proterozoic boundary is better
understood, care should be taken with the interpretation of extent of depletion of Proterozoic magma
sources.
© 2014 Elsevier B.V. All rights reserved.
∗ Correspondence to: Museum Building, Trinity College Dublin, Dublin 2, Ireland. Tel.: +353 873571973.
E-mail address: [email protected]
http://dx.doi.org/10.1016/j.precamres.2014.12.007
0301-9268/© 2014 Elsevier B.V. All rights reserved.
B.S. Kamber / Precambrian Research 258 (2015) 48–82
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Contents
1.
2.
3.
4.
5.
6.
7.
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
The long-lived Archaean continental lithosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.1.
Evidence for longevity of Archaean cratons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.2.
Archaean bi-modal magmatism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.3.
The co-evolution of the crust and the refractory mantle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
The Hadean–Archaean boundary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.1.
The Acasta gneisses . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.2.
The detrital Hadean zircon record . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.3.
Destruction of the Hadean crust and the dawn of the Archaean . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.3.1.
Destruction by bolide bombardment . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.3.2.
Catastrophic lid collapse and mantle overturn . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.3.3.
The dawn of the Archaean . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Archaean oceanic lithosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.1.
The paucity of direct evidence for Archaean oceanic lithosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.2.
Indirect evidence for Archaean oceanic lithosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.2.1.
The arc or dehydration fingerprint . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.2.2.
Lateral accretion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.2.3.
General geophysical and global geochemical constraints . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Rapidly grown juvenile Archaean continental crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.1.
Archaean volcanic plateaux . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.2.
Differences between modern oceanic plateaux and the envisaged formation of Archaean submarine volcanic plateaux . . . . . . . . . . . . . . . . .
5.2.1.
Syn-volcanic granitoids . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.2.2.
Burial of hydrated basalt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.2.3.
Thick cratonic mantle keels . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
The Archaean–Proterozoic boundary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
6.1.
A brief summary of ideas regarding the cause of the A–P boundary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
6.2.
Evidence from the extent of mantle depletion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
6.3.
Implication of A–P boundary for crustal ‘growth models’ . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
6.4.
An episodic or long-lived transition zone barrier? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Conclusions and outlook . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
1. Introduction
The Archaean Eon spans at least 1.25 billion years (Ga) of time
(Bleeker, 2004; Moorbath, 2005; Van Kranendonk et al., 2012). It
has witnessed key geological and biological events that continue
to inspire scientific discovery and endeavour. With the integration
of geological, geochemical and geophysical observations and data,
enormous progress has been made in elucidating the history of
Archaean rocks and extracting information relevant to the ancient
lithosphere, hydrosphere, atmosphere and biosphere. It is not the
ambition of this paper to summarise and detail all these discoveries. Rather, the principal aim is to synthesise much of the newly
available geochemical and isotopic data pertinent to crustal evolution during the Archaean Eon and integrate this synthesis with the
time-honoured geological facts. This necessarily includes a discussion of the nature and significance of the boundaries that book-end
the Archaean: the transition from Hadean and to the Proterozoic.
It stands to reason that unless we can explain its beginning and its
end, our understanding of the Archaean will remain incomplete.
It would be tempting to review the evolution of the early terrestrial crust along a logical timeline, from the oldest events through
to the youngest. But the Hadean geological and isotopic record is
incomparably less complete than that of the Archaean and therefore, the few deductions that planetary geologists can make about
the Hadean pale in the face of the information available for much of
the Archaean. Equally, the expression of the Archaean–Proterozoic
(A–P) boundary in the sedimentary rock record does not exactly
coincide with the compositional evolution (e.g. MgO) recorded in
igneous rocks. Indeed, there is still room for substantial discussion
of how the Hadean–Archaean (H–A) and A–P boundaries should be
defined (see Van Kranendonk et al., 2012 for a historical perspective
and for one currently proposed Precambrian timescale). In view of
the uncertainties regarding the boundaries and due to the paucity
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of geological evidence for the Hadean, an alternative narrative for
this paper was preferred, starting with a discussion of those aspects
of early crustal evolution that we think we know best branching
out into the ever less well constrained. This will start on the firm
ground of long-known long-lived, ancient continental crust of cratons that preserve nearly the full Archaean crustal record. Only once
the processes that have shaped long-lived Archaean continental
lithosphere have been discussed, will we enquire about the significance of the paucity of Hadean rocks and speculate about the nature
of the H–A boundary.
By comparison with the record of long-lived Archaean continental lithosphere, the evolution of oceanic lithosphere is shrouded
in much more uncertainty and controversy. The feature that continues to dominate this debate is whether subduction of oceanic
plates was operational in the Archaean or not, a theme which lends
itself as the next step along the narrative. By addressing how the
Archaean mantle lost heat – through modern style plate tectonics
with fully established spreading and subduction or some form of
heat transfer through longer-lived ‘oceanic’ lithosphere – we find
ourselves also confronted with the question of whether Archaean
plumes existed and how they might have operated.
This then takes us into an under-appreciated aspect of the
Archaean continental crustal record. Namely, as pointed out by
Van Kranendonk (2010) and Kamber (2010), juvenile crustal segments exist, which have formed very rapidly, apparently without
the involvement of much pre-existing, significantly older crustal
nuclei. It will be argued that these juvenile continental remnants
differ markedly from how continental lithosphere was formed in
the Phanerozoic, possibly through the maturation of thick oceanic
plateaux, some of which eventually become the substrate for new
continental lithosphere.
The closing section of this paper will address the A–P boundary, or more specifically, the question of how the Archaean might
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B.S. Kamber / Precambrian Research 258 (2015) 48–82
have come to its relatively abrupt end. Interesting proposals for the
significance of the A–P boundary have been put forward by Condie
et al. (2009), Condie and O’Neill (2010), and Campbell and Griffiths
(2014). The idea I advance in this paper also seeks the answer in
the reorganisation of the mantle heat structure. It too, is intended
as a working hypothesis and as such is deliberately drawing attention to the aspects of the A–P boundary that are in conflict what we
currently know about the evolution of the Precambrian Earth.
In writing this paper, I have attempted to find a balance between
the weight attributed to geological versus geochemical and geophysical data and models. The study of Precambrian rocks has
become analytically very sophisticated, indeed, some of the most
challenging isotopic measurements to date (e.g. O’Neil et al., 2008)
have been performed on Archaean rocks. Notwithstanding the
incredible wealth of information that geochemistry and geophysics
have brought to the study of the ancient Earth, there has also been
a tendency to dismiss the seemingly less sophisticated and more
qualitative geological observations regarding Archaean crust. One
truth that is being lost in the highly technical discussion of isotopic
variations of a few parts per million (ppm) is the fact that the early
geologists who set foot on Archaean cratons instinctively knew that
they were dealing with ‘different’ rocks, even though their mineralogy looked altogether familiar. Having originally been trained as
a structural-metamorphic geologist in the Swiss Alps (e.g. Kamber,
1993), I was deployed into the Archaean of Zimbabwe (e.g. Kamber
et al., 1995) without much preparation and, like so many after him,
I must then have relived the moment of insight by Logan (1857)
when appreciating just how different Archaean rocks are. It is my
hope that this review will do justice to the Archaean and steer away
from the view that our planet has been operating along the familiar
concept of true modern plate tectonics for as long as 4.4 Ga.
2. The long-lived Archaean continental lithosphere
Within a decade of the introduction of modern radiogenic
isotope-based whole-rock methods to geology (then Pb–Pb, Rb–Sr
and K–Ar systems; e.g. Compston et al., 1960; Gast et al., 1964),
Moorbath and co-workers (Black et al., 1971; Moorbath et al.,
1972, 1973) made the spectacular discovery that both the coastal
gneisses of the Godthaabsfjord and the supracrustal rocks in the
Isua area of SW Greenland were of great antiquity, in the range
of 3.60–3.75 Ga old. Earlier attempts at finding very ancient rocks
had unearthed Neoarchaean rocks in a variety of places, for example in the Lewisian of northern Scotland (e.g. Giletti, 1959; Evans,
1965) while Pb-isotope studies (e.g. Kanasewich and Farquhar,
1965) had indirectly hinted at the existence of Mesoarchaean rocks
(note that I use the terminology of Bleeker’s (2004) timescale (see
Fig. 1A) for the subdivision of the Archaean and the timing of the
H–A and P–A boundaries). Although it does retrospectively seem
somewhat serendipitous that such old rocks should have been
found quite early in the era of modern isotope geochronology,
it is important to recall that the discovery of the ancient Eoarchaean rocks in SW Greenland was not simply a stroke of luck but
was guided by field observations. These included the persistent
extent of high-grade metamorphism and deformation, including
recognition of several generations of dyke swarms (some metamorphosed and deformed). This evidence hinted at several phases
of regional deformation (e.g. Bridgwater et al., 1970; McGregor,
1973), which had apparently affected certain lithologies but not
others, suggesting protracted crustal growth. Since the early wholerock studies, reconstruction of the events of the Eoarchaean have
greatly benefitted from high-precision U/Pb zircon geochronology
(e.g. Krogh, 1973; Baadsgaard, 1973) and from the development
of several additional radiogenic isotope systems (Sm–Nd, Lu–Hf
and Re–Os). In SW Greenland, the known rock record has been
pushed back further, to 3.85 Ga, with the discovery of an area
south of the Isua belt that contains many >3.8 Ga gneisses (Nutman
et al., 1999; Kamber et al., 2003; Amelin et al., 2011) with enclaves
of supracrustal and ultramafic rocks of great isotopic antiquity
(e.g. Bennett et al., 2002). The coastal gneisses too have yielded
many zircons older than 3.65 Ga (e.g. Kinny, 1986; Nutman et al.,
2007) clearly indicating that evolved crust formation began there
between 3.80 and 3.84 Ga, with strong evidence that zircon growth
before that time is extremely limited (Horie et al., 2010). Even if a
few more Hadean zircon grains will eventually be found in the area,
the geological record and vast bulk of zircons show an absence of
preserved Hadean rocks. Crust formation and reworking as preserved by the Archaean gneiss complex was not a single stage
process but continued across most of the North Atlantic craton into
the Neoarchaean (Bridgwater et al., 1970), during which time the
much older gneisses were strongly reworked and at least partly
isotopically reset, leaving only the lowest strain gneisses relatively
unaffected.
2.1. Evidence for longevity of Archaean cratons
On account of their wide occurrence, excellent state of preservation and spectacular field exposure, the Itsaq gneisses of SW
Greenland remain the archetype of Eoarchaean terrestrial crust (e.g.
Nutman et al., 1996). The North Atlantic craton as a whole and the
Archaean gneiss complex in particular experienced a complex geological history throughout the Archaean (see e.g. Naeraa et al., 2014,
for a summary) that ended with the intrusion of the 2.55 Ga Qôrqut
granite. This composite suite of intrusions covers more than ca.
100 km2 within the much older Archaean gneiss complex of SW
Greenland and exhibits abundant field and geochemical (Brown
et al., 1981a,b), inherited zircon (Nutman et al., 2010) and isotopic
evidence (Moorbath et al., 1981; Hiess et al., 2011; Naeraa et al.,
2014) for reworking of pre-existing, extremely old crust. The most
impressive and incontrovertible evidence for the involvement of
Eoarchaean crust is the combination of extremely unradiogenic Pb
(Moorbath et al., 1981) and inherited 3.6–3.8 Ga zircon xenocrysts
(Nutman et al., 2010).
Indeed, the coincidence of the occurrence of Eoarchaean zircon
(be it as xenocrysts or as detrital grains in much younger sedimentary rocks) on cratons with extreme Pb-isotope compositions
of Meso- and Neoarchaean granitoids is a very consistent phenomenon. Kamber et al. (2003) originally noticed the association
between cratons bearing Hadean and Eoarchaean zircon and an
evolutionary episode within a crustal precursor with a 238 U/204 Pb
much elevated over coeval mantle (termed ‘high ␮’ by Oversby
(1975)). For readers unfamiliar with the high ␮ terminology, the
concept is explained with a number of examples in Fig. 2. A more
detailed treatment of the high ␮ signature can be found in Kamber
et al. (2003) and Kamber (2007).
The uranogenic Pb-isotope evolution of the Archaean depleted
mantle is well constrained both with data from ancient ores and
feldspars (e.g. Carignan et al., 1993) and with results of forward
models (e.g. Zartman and Haines, 1988; Kramers and Tolstikhin,
1997). By comparison with the Archaean depleted mantle evolution
line (Fig. 2A), many initial Pb-isotope compositions of igneous rocks
from long-lived cratons plot in a position that requires a substantial
episode of evolution in an environment with a substantially higher
238 U/204 Pb than coeval mantle.
A simple example is the Long Lake batholith of the Beartooth
Mountains in the Wyoming craton. The different components
of this batholith (diorite, grey gneiss, granodiorite) define a ca.
2.8 Ga regression line compatible with the 2.79 Ga U/Pb zircon
ages (Wooden and Mueller, 1988). However, the isochron does
not intersect the Archaean depleted mantle evolution line at all
(Fig. 2B) and the initial Pb recorded by leached feldspar (Wooden
B.S. Kamber / Precambrian Research 258 (2015) 48–82
51
Fig. 1. Overview of geological time divisions and global parameters covered in this paper. Panel (A) shows the geological timescale proposed by Bleeker (2004), which is
used throughout this paper when referring to the H–A boundary, the subdivisions within the Archaean and the definition of the end of the Archaean. Panel (B) shows the
alternative timescale proposed by Van Kranendonk et al. (2012) placing the H–A and A–P boundaries significantly earlier and later, respectively. Panel (C) shows the time
series of calculated numbers of bolide impacts on Earth capable of producing basin-sized structures (simplified after Bottke et al., 2012). According to this calculation, 105
giant basin-sized impacts occurred between 3850 and 4100 Ma. Panel (D) shows the calculated radioactive heat production (TW) from the three most important elements
(U, Th, K) individually, as well as the cumulative bulk silicate Earth heat production. The heat production was scaled to yield a present-day value of 16 TW (Korenaga, 2008)
using the bulk silicate Earth Th (81.3 ppb) and U (20.3 ppb) concentrations compiled by Kramers and Tolstikhin (1997) and a K/U ratio of 12,700 (Jochum et al., 1983) within
error of the estimate of Arevalo et al. (2009). Panel (E) shows the reconstructed evolution of the mantle plume potential temperature (◦ C) after Kamber (2010).
and Mueller, 1988; Ulrich et al., 2011) does not plot anywhere near
the composition of the 2.8 Ga depleted mantle. Instead, both plot
above the predicted coeval mantle composition at much greater
207 Pb/204 Pb ratios and somewhat higher 206 Pb/204 Pb ratio. This
phenomenon requires the source of these magmas to have experienced an episode in an environment with a higher U/Pb than the
mantle. The magnitude of the offset between the observed initial
composition and that of the modelled coeval mantle is sensitive to
how much higher the ␮ of this source was than that of the mantle, and how long the source existed. Because 207 Pb is the daughter
product of 235 U, which decayed much faster than 238 U, the observed
offset in combined 207 Pb/204 Pb and 206 Pb/204 Pb can be used to
calculate the minimum age of separation of the high ␮ source from
the depleted mantle (e.g. Kamber et al., 2003). In the case of the
Long Lake batholith, source separation must have occurred no later
than 3.9 Ga (see position of 3.9 Ga high ␮ source in Fig. 2B). It is
important to appreciate that this is a firm constraint. If source separation had occurred later, a higher ␮ would have been needed to
achieve the observed 207 Pb/204 Pb, which would push the modelled
206 Pb/204 Pb to higher values than those permitted by the leached
feldspar. In reality, the increase in source ␮ was likely a multi-stage
process, allowing for the possibility of even older source separation.
It should be noted that the vast majority of Archaean grey
gneisses have present-day whole-rock Pb-isotope compositions
52
B.S. Kamber / Precambrian Research 258 (2015) 48–82
16
3.0 Ga
207
Pb/
204
Pb
15
3.5 Ga
11
Pb
12
13
2.80 Ga
14.0
14
15
13.5
12
14
16
18
20
1 Ga
C
2.55 Ga
0 Ga 14.5
2 Ga
15.0
204
Hadean crust vs.
mantle evolution
15.0
15.5
Pb/
2.80 Ga
15.0
4.3 Ga
10
207
15.5
14.5
4.0 Ga
11
14.0
Wyoming craton
3.5 Ga
4.0 Ga
12
14.5
B
16.0
2.5 Ga
3.0 Ga
14
13
16.5
2.5 Ga
A
14.0
3 Ga
13.5
13.0
13.5
North Atlantic
craton
13.0
Qorqut granite
complex
12.5
D
3.9 Ga
12.5
11
12
13
14
15
16
17
12.0
10
18
20
19
17
E
Pb
207
204
Pb/
16
14
15
16
17
18
19
North China craton
14
Zimbabwe craton
14
12
10
F
13
15
15
13
12
16
18
17
11
3.3 Ga
13
3.35 Ga
15
20
206
Pb/
25
204
Pb
30
12
10
12
14
16
206
Pb/
18
20
204
Pb
Fig. 2. Common Pb-isotope systematics of modelled mantle and crustal reservoirs and selected Archaean rock suites. Panel (A) compares the depleted mantle evolution
(green) with that of the Hadean high ␮ source (blue). Both model curves are shown in 100 Ma intervals. The Hadean source evolution is from Fig. 10 of Kamber et al. (2003)
and separated from the depleted mantle at 4.3 Ga with a constant ␮ of 10.5. The depleted mantle curve is from Kramers and Tolstikhin (1997). Panel (B) shows data for the
2.79 Ga Long Lake batholith suite of the Wyoming craton. Both whole rock (squares) and leached feldspar (circles) data from Wooden and Mueller (1988), in yellow colour,
and Ulrich et al. (2011), in red, are shown. A combined regression line was forced to yield the known zircon U/Pb crystallisation age of 2.79 Ga. The regression line does
not intersect the depleted mantle evolution line (green) at 2.8 Ga. A high ␮ source that separated from depleted mantle at 3.9 Ga with a ␮ of 10.75 evolves, by 2.8 Ga, into
the observed feldspar initial Pb-isotope compositions. If the Long Lake batholith suite does not contain any juvenile component, 3.9 Ga is the age of the pre-existing source.
Dilution with juvenile Pb would push the required age of source separation further back in time. Panel (C) Present-day composition of Eoarchaean whole rocks (grey gneisses
and granitoids), as red solid squares from SW Greenland and Labrador (data from Kamber et al., 2003) illustrating that despite many having a high ␮ source origin, they
later experienced U-loss, retarding their further Pb-isotope evolution. This is evident from the position of their Pb-isotope composition compared to the depleted mantle
evolution model (green) of Kramers and Tolstikhin (1997). They all plot at less radiogenic position than modern mantle. Panel (D) shows whole rock and feldspar data for
the 2.55 Ga Qôrqut granite complex reported in Moorbath et al. (1981), yielding a ca. 2.5 Ga regression line. Also shown up to the crystallisation age of 2.55 Ga is the depleted
mantle evolution line. Panel (E) shows the 3.35 Ga regression line and the two least radiogenic data points of the 3.57 Ga Mont d’Or granite (reported in Taylor et al., 1984).
Also shown up to the isotopic rehomogenisation age of 3.35 Ga is the depleted mantle evolution line. Panel (F) shows feldspar Pb-isotope data of the 3.30 Ga Chentaigou
granitoid intrusion in the Anshan region of the North China craton reported by Wan et al. (1997). A regression line with a 3.30 Ga slope was forced to the data, plotting far
above coeval depleted mantle.
that are less radiogenic (and often dramatically so) than modern
mantle or upper continental crust. This is because the gneisses
underwent later partial melting during which most U (and Th and
K) was removed, leaving the gneiss with very low U/Pb ratios (e.g.
Bolhar et al., 2007). This should not be taken as evidence that these
gneisses represent a low ␮ reservoir. Rather, what matters is the
position of the initial Pb-isotope composition of the gneisses relative to the depleted mantle evolution line. The most impressive
B.S. Kamber / Precambrian Research 258 (2015) 48–82
examples of this are the Eoarchaean gneisses of SW Greenland.
They contain some of the least radiogenic Pb (Fig. 2C) preserved
on Earth (e.g. Kamber and Moorbath, 1998) yet several gneisses
nonetheless contain evidence that their original source had a higher
␮ than mantle (e.g. Kamber et al., 2003). The high ␮ source of these
Eoarchaean gneisses must have separated from the mantle in the
Hadean (Fig. 2A).
Another interesting but complex example is the very unradiogenic Pb in the Neoarchaean Qôrqut granite (Fig. 2D), whose
composition plots far below any plausible juvenile source. Instead,
it requires a Hadean/Eoarchaean episode in a high ␮ environment
followed by massive U (and Th) loss leaving the Qôrqut source with
a much retarded Pb-isotope composition by 2.55 Ga (see Naeraa
et al., 2014, for a detailed explanation). A further classic high ␮
source is recorded in the Zimbabwe craton (e.g. Taylor et al., 1991;
Berger and Rollinson, 1997). There are a number of conventional
Pb-isotope studies on a variety of Archaean intrusive suites on this
craton, which show that the high U/Pb reservoir was apparently
tapped repeatedly from the Palaeoarchaean (Fig. 2E) through to
the Neoarchaean (see 10b of Kamber et al., 2003). We proposed
that the North Atlantic, Yilgarn, Wyoming, Zimbabwe and the western Slave cratons belong to this group of long-lived, very ancient
cratons and speculated that the high ␮ source could be greater in
age than 3.8 Ga. With the benefit of many more data, the Anshan
region of the North China craton can now also be added to this
list. Uranium/Pb zircon geochronology has identified three separate
occurrences of in situ Eoarchaean remnants (Song et al., 1996; Wu
et al., 2008; Wan et al., 2012) and a variety of detrital zircons of Eoarchaean age (see Wan et al., 2009, for a summary) likely sourced from
the North China craton. Although only few whole-rock or feldspar
Pb-isotope studies for younger intrusives from this craton exist,
Wan et al.’s (1997) data for the 3.3 Ga Chentaigou granitoid intrusion in the Anshan region clearly show the presence of the ancient
high ␮ substrate (Fig. 2F). The involvement of much older crust
in the source of younger Archaean granitoids across the Anshan
area has also been implied from Nd-isotope signatures (Song et al.,
1996). A final important observation borne out by modern U/Pb
zircon studies is the apparently closely spaced, repeated igneous
activities that followed the crystallisation of the oldest Eoarchaean
protolith (Wu et al., 2008), followed by longer times of igneous
quiescence.
Since 2003, a great number of new zircon dates and isotopic
measurements have become available for the Superior Province.
This very complex super-craton is composed of several different
sub-provinces, which can be identified as chronologically, geologically, geochemically and tectonically distinct and having been
amalgamated into a ‘collage’ largely after the main crust formation
events (e.g. Downey et al., 2009). Many of the Superior Province
sub-provinces are quite juvenile in character and differ markedly
from the high ␮ cratons. However, along the northern and southern
margins, a number of Eoarchaean crustal fragments exist. Building
on work by Bickford et al. (2006), David et al. (2009, p. 150) proposed that “fragments of Palaeoarchaean to Eoarchaean terranes
that rim the Superior Province represent remnants of a larger terrane(s) that was rifted and dismembered prior to the formation
of the 3.0–2.7 Ga Superior Craton”. These dismembered terranes
include the Palaeoarchaean Minnesota River gneisses that require
a long crustal pre-history (Satkoski et al., 2013) and the Assean
Lake gneisses (Bohm et al., 2003), some of which have Nd-isotope
model ages well in excess of 3.8 Ga. Most famously, occurring
along the north-eastern Superior Province in the Minto Block is the
Nuvvuagittuq belt (David et al., 2009). This package of rocks covers a small area (ca. 15 km2 ) and, in my opinion, should be called
an enclave rather than a greenstone belt, on account of its modest size. To all intents and purposes, the assemblage of rock (as
e.g. described in David et al., 2009; O’Neil et al., 2011) resembles
53
typical Archaean granite-greenstone terrains. The main importance of these rocks is that they contain Eoarchaean protoliths and
relatively well-developed 142 Nd isotope deviations from typical
terrestrial rocks (O’Neil et al., 2008). Variations in the 142 Nd/144 Nd
ratio could only have formed within a few hundred Ma of Earth
accretion. In the Nuvvuagittuq samples, 142 Nd/144 Nd deviations
correlate with elemental Sm/Nd ratios and have been interpreted
by some authors to indicate a ca. 4.4 Ga protolith age for the greenstones. By contrast, the conventional 147 Sm–143 Nd systematics do
not support a Hadean age and, in agreement with available zircon
ages (David et al., 2009; Cates et al., 2013; Darling et al., 2013),
are compatible with an Eoarchaean, ca. 3.75–3.82 Ga protolith age.
Indeed, the correlation between 143 Nd/144 Nd, Sm/Nd elemental
ratios and 142 Nd/144 Nd systematics, coupled with the absence of
>4 Ga 143 Nd/144 Nd regression lines suggests very strongly that
the rocks are Eoarchaean in age (Roth et al., 2013) and that the
142 Nd/144 Nd deviations are inherited source features testifying to
very early silicate Earth differentiation. Hafnium isotopes in zircon also hint at inheritance from a Hadean source (O’Neil et al.,
2013).
The oldest preserved rock on Earth, the Acasta gneisses, occur
along the western margin of Slave craton (Canada). This craton is
composed of two halves of very different crustal history, divided
by a very marked N–S running isotopic boundary. The Neo- and
Mesoarchaean rocks to the west of the boundary show incontrovertible evidence for involvement of much older (>3.5 Ga) crustal
precursors (Davis et al., 1996). The Acasta gneiss, which occurs
west of the isotope boundary, is of very small areal extent (ca.
25 km2 ) and the field relationships between the different polyphase
varieties of gneiss are quite complex. The zircon record strongly
suggests that an initial pulse of crust formation occurred between
3.96 and 4.04 Ga (Bowring and Williams, 1999; Bowring et al., 1990;
Iizuka et al., 2006; Mojzsis et al., 2014). Regardless of the impressive new geological, geochemical and isotopic insight, which will
be discussed in Section 3 of this paper, a chief limitation of the
Acasta gneiss complex is that it does not preserve any properly
recognisable mafic protoliths. For example, the most magnesian
Acasta gneisses – termed mafic gneiss (and interpreted as cumulates) – reported by Mojzsis et al. (2014) contain between 5.0
and 5.7 wt% MgO, far below any primary mantle melt. The Acasta
gneisses thus provide an important glimpse of Hadean–Eoarchaean
crust formation that is biased towards the evolved compositional
spectrum and has limitations for generic models of early crust formation.
In summary, there is very strong geochronological evidence for
the existence of >3.5 Ga rocks on several cratons. All of these cratons have complex magmatic histories with multiple and often
closely spaced events of intrusions. For most of these cratons, Pbisotope data are available for Neo- and/or Mesoarchaean granitoids.
A consistent picture is emerging that these cratons have distinctive Pb-isotope compositions requiring re-melting of ancient,
pre-existing crust up until the Neoarchaean. These cratons are here
called the truly long-lived high-␮ continental lithosphere. By contrast, other cratons, most notably the Pilbara and Kaapvaal and
many of the sub-provinces (e.g. Abitibi) of the Superior Province
lack Eoarchaean zircon and their isotopic character is much more
juvenile. Where present, evidence for Eoarchaean precursors is far
more limited (e.g. Kroner et al., 2013). These will be discussed separately. Where present-day cratons display an isotopic dichotomy
(e.g. Slave and Yilgarn) Eoarchaean zircon is restricted to the part
that shows radiogenic isotope evidence for Neoarchaean remelting
of much older crust. Given the exceptionally large geochronological
database for many of these juvenile subprovinces (e.g. Ayer et al.,
2002), which often are rich in Au-deposits (e.g. Mohan et al., 2008),
the lack of Hadean inheritance appears to be real and not an artefact
of insufficient data coverage.
54
B.S. Kamber / Precambrian Research 258 (2015) 48–82
Relative frequency
0.1
A
Archaean volcanic rocks
(n - 12282)
0.08
0.06
0.04
0.02
0
20
25
30
35
40
45
50
55
60
65
70
75
80
SiO2 wt. %
Relative frequency
0.08
B
Archaean plutonic rocks
(n - 02738)
0.06
0.04
0.02
0
20
25
30
35
40
45
50
55
60
65
70
75
80
SiO2 wt. %
Fig. 3. Histograms of SiO2 content (in wt%) of Archaean igneous rocks from
pre-compiled GEOROC datasets. Panel (A) shows volcanic rocks, which have the
expectedly lower abundance of siliceous compositions compared to plutonic rocks
(panel B).
2.2. Archaean bi-modal magmatism
The continental crust is widely believed to be, on average,
andesitic in composition (e.g. Taylor and McLennan, 1995; Rudnick
and Gao, 2003) with SiO2 of ca. 60 wt%, creating an obvious
mass balance issue, considering that mantle melts are overwhelmingly basaltic with much lower SiO2 between 44 and 48 wt%.
The Archaean continental crustal also appears to be, on average,
andesitic but in addition, shows a well-established bi-modal distribution of igneous rocks. This is evident both in volcanic and
plutonic rocks (Fig. 3). The low SiO2 mode, from 12,282 analyses
of Archaean volcanic rocks is at 50 wt% and that of 2748 plutonic
rock analyses is at 49 wt%. These magmas were thus emplaced into
and onto the evolving continental crust with relatively limited fractional crystallisation from an original mantle melt composition.
It is also evident that the proportion of mafic over felsic volcanic
rocks is much higher than with the plutonic equivalents. The high
SiO2 mode is not well defined in the volcanic rock population but
occurs in the region of 68–75 wt%. The very high SiO2 contents of
a minority of analyses (>75 wt%) represent samples that experienced secondary silicification, likely under the influence of Si-rich
seawater (e.g. Thurston et al., 2012). The high SiO2 mode is therefore better estimated by inspection of the plutonic rock histogram
(Fig. 3B), returning ca. 68 and 70 wt%. This mode contains grey
gneisses, the tonalite–trondhjemite–granodiorite (TTG) series, and
typically late- to post-tectonic granites s.s.
The key observation from the frequency distribution is the
paucity of intermediate compositions, with minima in the histograms exactly in the region of the average continental crust
composition around 60 wt% SiO2 . This is unlikely to reflect an
observational bias due to, for example, preferential exposure of
crustal levels with limited andesitic igneous rocks. The compositional spectrum of intermediate and felsic Archaean volcanic rocks
is very well known because these are the major targets in studies
that establish greenstone belt chronologies via U/Pb zircon dating
(e.g. Jayananda et al., 2013). From the published database (Fig. 3A)
it is therefore clear that the erupted felsic lavas are overwhelmingly
rhyolitic, and not andesitic.
Rather, the most plausible explanation for the paucity of
andesite is that the textbook fractional crystallisation mechanism
from basalt to granite was not dominant in the Archaean. Instead,
fractional crystallisation only drove mantle magmas to slightly
evolved compositions between 50 and 55 wt% SiO2 . The highly
siliceous igneous rocks must therefore have a different origin. The
most widely accepted idea for the highly evolved TTGs (and their
highly metamorphosed chemically equivalent grey gneisses) is that
they formed by remelting of a mafic (basaltic) precursor.
The model of Martin (1986) advocated that the most likely geodynamic setting for this to have happened was in a subduction
zone, where relatively hot Archaean oceanic crust was converted
to garnet-bearing amphibolite or eclogite, and melted directly,
akin to end-member adakites that presently form where unusually young (and therefore abnormally hot) oceanic lithosphere is
being subducted. However, closer comparison of the full trace element chemistry of adakites and Archaean TTGs has shown that the
majority of the latter do not share critical features of adakites (e.g.
Kamber et al., 2002). Rather, the Archaean silicic magmas show
strong evidence for fluid-induced melting.
The alternative to melting a mafic protolith in the mantle (i.e.
a subducting slab) is to melt the base of thickened mafic crust
(Rapp and Watson, 1995). In this regard, important new insight
has come from experimental work, modelling and from isotopic
(mainly Lu–Hf) studies of Archaean rocks and zircon. Nagel et al.
(2012) showed that a good fit between observed TTG chemistry
and modelled trace element compositions is obtained by melting
of an (over thickened) stack of Archaean tholeiite at depths between
35 and 50 km. A key finding of the study of Nagel et al. (2012) was
the difference in the phase diagrams calculated for a hypothetical N-MORB vs. an actually observed Eoarchaean metabasalt. The
phase diagram for the Eoarchaean sample (2000–13 of Polat and
Hofmann, 2003), which has an SiO2 content of 50.7 wt% (i.e. exactly
below the mafic mode in Fig. 3A), produces TTG-like magmas at
lower pressures than N-MORB and with less residual plagioclase
and a much better resulting trace element fit than the more aluminous and sodic N-MORB source. The genetic relationship between
TTG and tholeiite is not only supported by phase diagrams but
also evident in the Nd- and Hf-isotope similarity between metatholeiites and TTG reported by Hoffmann et al. (2011a).
A detailed investigation into the high-field-strength-element
systematics of TTG from SW Greenland by Hoffmann et al. (2011b)
revealed that there are interpretable co-variations between the
Nb/Ta ratio and initial Hf-isotopes. Hoffmann et al. (2011b) also
pointed out that any successful model for TTG petrogenesis has
to be able to explain the co-existence (in SW Greenland) of relatively high Nb/Ta with relatively low Nb/Ta TTG. Importantly,
the wide range in TTG and grey gneiss TTG is not restricted to
SW Greenland (see e.g. Mohan et al., 2008; Satkoski et al., 2013).
A likely explanation for this phenomenon is the depth of melting of over thickened tholeiitic crust. At greater depths, where
rutile and ilmenite control Nb/Ta fractionation, partial melting will
generate high Nb/Ta granitoids, whereas at shallower, mid-crustal
levels, partial melting in the presence of amphibole ± titanite will
generate more variable but generally lower Nb/Ta partial melts
(Hoffmann et al., 2011b). The co-occurrence of both high Nb/Ta
and high Nb/Ta TTG thus imply melting over a substantial crustal
depth section.
B.S. Kamber / Precambrian Research 258 (2015) 48–82
160
Arc melt inclusions (n=2,255)
140
120
100
n
The Lu–Hf system is sensitive to the extent of chemical differentiation of a magma source and can therefore be used to characterise
the nature of the melt source. Typical continental crust has a very
low Lu/Hf ratio. If allowed to evolve over a substantial time period,
on the order of 500–1000 Ma, such a source will assume a very
characteristically unradiogenic Hf isotope composition. A silicic
magma that forms by melting such a low Lu/Hf source will crystallise zircon and store the Hf-isotope composition, which can be
reconstructed much more easily than from whole rock. By contrast, a less evolved, mafic crustal precursor source has a much
higher Lu/Hf ratio. Therefore, the initial Hf-isotope composition is
an excellent indicator of the composition of the precursor material.
Several recent Hf-isotope studies of Meso- and even certain Neoarchaean meta-igneous suites have come out in favour of re-melting
of largely basaltic precursor material (e.g. Satkoski et al., 2013;
Huang et al., 2013) and have revised earlier proposals that many
TTG and even granites formed by re-melting of already evolved
material.
For example, on the basis of Pb-isotopes alone, Moorbath et al.
(1981) argued that the 2.55 Ga Qôrqut granite was mainly derived
by re-melting of already highly differentiated tonalitic gneiss. By
contrast, Hf-isotope data of the Qôrqut granite are nowhere nearly
as unradiogenic as could be expected from >1000 Ma evolution of
Eoarchaean granitoids (Naeraa et al., 2014). The relatively homogenous, mildly unradiogenic Hf-isotope composition coupled with
trace element constraints show that the main mass of the Qôrqut
granite formed by re-melting of Eoarchaean mafic lower crust.
In summary, the bimodal distribution of Archaean igneous rocks
represents the emplacement of mantle-derived basaltic melts that
have experienced relatively little fractional crystallisation as well
as partial melts from basaltic precursors. The most recent studies favour re-melting of basaltic precursors within over-thickened
crust but the possibility of direct melting of subducted oceanic crust
remains, at least for those TTG that have high compatible element
contents.
Seismic data across Phanerozoic deep intracratonic sedimentary basins (e.g. the East Barents Sea basin) have identified bodies
of high density rock below the Moho. Models of crustal shortening
(e.g. Gac et al., 2013) show that these bodies likely represent dehydrated dense metamorphosed mafic lower crustal lithologies. This
observation may be important in the context of melting garnetbearing meta-basalt in over thickened Archaean crust. Namely, the
mineralogy of the refractory residue may assume a sufficiently high
density to warrant foundering from the base of the crust (e.g. Huang
et al., 2013), the classic ‘restite-disposal’ mechanism proposed by
Rapp and Watson (1995). The removal of less silicic restite during
Archaean continental crust formation therefore provides a solution to the overall continental crust composition conundrum. The
change from bimodal Archaean to unimodal Phanerozoic igneous
rocks remains a valid observation that still requires an explanation. However, I propose that it is not the Archaean bimodality that
requires explaining, but that the unimodal distribution of Phanerozoic igneous rocks, with its preponderance of andesite, is difficult
to reconcile with the overall mass balance issue of average continental crust, because it apparently does not allow the formation of
dense residues that could be disposed of and recycled back into the
asthenosphere.
In this regard, it is interesting to note the results of analyses of
melt inclusions in phenocrysts from modern arc volcanoes. Reubi
and Blundy (2009) compiled the compositions of 2255 quenched
glassy melt inclusions from phenocrysts of arc igneous rocks. The
distribution in SiO2 in the inclusions is clearly bi-modal (Fig. 4)
with a mafic mode at 53 wt% SiO2 and a silicic mode at 76 wt%,
remarkably similar to the distribution in Archaean plutonic rocks
(Fig. 3B). Based on this observation, notably from globally occurring
arc volcanic rocks, Reubi and Blundy (2009, p. 1269) proposed that
55
80
60
40
20
0
40
45
50
55
60
65
70
75
80
85
SiO (wt%)
2
Fig. 4. Histograms of SiO2 content (in wt%) of quenched phenocryst-hosted melt
inclusions from a global compilation of arc magmas simplified after Fig. 2b of Reubi
and Blundy (2009).
“true liquids of intermediate composition (59 to 66 wt% SiO2 ) are
far less common in the sub-volcanic reservoirs of arc volcanoes than
is suggested by the abundance of erupted magma within this compositional range” and further that “effective mingling within upper
crustal magmatic reservoirs obscures a compositional bimodality
of melts ascending from the lower crust”. In a case study of the
classic andesite-erupting volcano Mount Hood, Kent et al. (2010)
showed that the distribution of SiO2 in melt inclusions is also
bimodal. Despite the fact that this volcano reliably erupts andesite
of a narrow compositional band (SiO2 varying from 55 to 65 wt%) at
surface, the melt inclusions demonstrate that the magmatic plumbing system is dominated by a bi-modal distribution of magmas.
Careful study of internal compositions of phenocrysts showed that
recharging of a magma reservoir of relatively evolved composition
with fresh basaltic magma promotes magma (and crystal-magma)
mixing, and triggers the eruption of hybrid, andesitic melts.
Thus, the difference between Archaean and Phanerozoic igneous
rocks may be apparent and may only relate to the mechanism of
magmatic emplacement but not a fundamental change in magma
genesis, other than komatiites which will be discussed later.
Modern andesite-dominated volcanoes apparently require longlived complex magmatic plumbing systems and reliably recurrent
introduction of new basaltic magmas, which may not have been
supported in the mechanically weaker Archaean crust.
2.3. The co-evolution of the crust and the refractory mantle
Although this synthesis concerns itself with the evolution of
Archaean crust, it is necessary to at least briefly review the role
of the sub-continental mantle lithosphere (SCLM) in preservation
of Archaean continental crust. This information includes seismic
and heat flux data on the cratonic scale and more punctuated geochemical and isotopic data from SCLM samples or magmas that
were significantly contaminated during ascent through the SCLM.
It has long been known that the heat flux over stable cratonic
areas, where thermal disequilibria induced by orogenesis have fully
relaxed, is significantly smaller than that in orogenic areas (e.g.
Jordan, 1978; Michaut et al., 2009). This reflects a combination of
three main factors: firstly, thermal disequilibria induced by ancient
56
B.S. Kamber / Precambrian Research 258 (2015) 48–82
orogenesis have long since relaxed; secondly, the heat flux across
the cratonic Moho is relatively low; and finally, there is low radiogenic heat production within much of the Archaean crust itself (e.g.
Mareschal and Jaupart, 2013). Thus, it can be inferred that the SCLM
is refractory and depleted in radioactive heat producing elements.
There is very strong 2D and 3D evidence from analysis of seismic
waves that Archaean cratonic nuclei are underlain by significantly
deeper SCLM than Proterozoic mobile belts and Phanerozoic crust.
This is possibly most compellingly shown in the lithospheric architecture of Africa reconstructed by Begg et al. (2009), who found
that (p. 23) “the larger cratons are underlain by geochemically
depleted, rigid, and mechanically robust SCLM; these cratonic roots
have steep sides, extending in some cases to ≥300-km depth”.
Smaller scale studies have revealed variations in SCLM thickness
that are also mirrored in the crustal architecture. For example, the
North China Craton is transected by a north-south running central
zone, which shows up as an area of lesser SCLM depth besides the
Archaean nuclei to the east and west (Tian et al., 2009).
Ever more sophisticated analysis of S receiver functions and
surface-wave anisotropy is revealing frozen-in anisotropy within
the SCLM as well as very deep compositional changes that may be
related to post-Archaean metasomatic/magmatic infiltration (e.g.
Savage and Silver, 2008; Sodoudi et al., 2013). A detailed study
of the Kalahari craton (Adam and Lebedev, 2012) confirmed earlier suggestions that the ancient lithosphere contained anisotropy
(Silver et al., 2004). Adam and Lebedev (2012) were able to resolve
the depth distribution of anisotropy from analysis of surface-wave
measurements into elements of modern anisotropy in the asthenosphere (fast-propagation directions parallel to the current plate
motion) and ancient (fossil) anisotropy in the SCLM of the Palaeoproterozic Limpopo belt with fast directions parallel to the suture,
suggesting preservation of SCLM anisotropy over at least 2 Ga since
the juxtaposition of the Kaapvaal and Zimbabwe cratons into the
Kalahari craton (e.g. Kamber et al., 1995).
On account of its abundance of diamondiferous kimberlite, the
Kalahari craton is a rich source of samples from the underlying
SCLM that can inform us about its age and nature. Smith et al. (2009)
studied mantle xenoliths, kimberlite mineral concentrates and diamond inclusions from two kimberlite pipes that erupted through
the southern Zimbabwe craton. The xenoliths are very refractory
with dunite (olivine Mg# between 0.92 and 0.95) more abundant
than harzburgite and only rare lherzolite and almost no eclogite. In
agreement with the depleted nature of the xenoliths and their constituent minerals, geothermobarometry on kimberlite concentrate
minerals from two pipes defined pressures of 4–7 GPa and temperatures of 1000–1200 ◦ C corresponding to a heat flux of 40 mW m−2 ,
exemplifying the isolation of the SCLM from the higher asthenospheric temperature. In terms of dating refractory SCLM, much
progress has been made using the Re–Os isotope systematics of
sulphides and chromites. Because Re is far more lithophile than Os
during mantle melting, refractory SCLM, representing the mantle
residue after (multiple?) melt extraction, ends up having a very
low Re/Os ratio. The Os-isotope composition of SCLM sulphides is
therefore frozen in and by comparison with a chondritic reference
line, it is possible to calculate Re-depletion ages, which are equated
with the main SCLM melting event. In the study of Smith et al.
(2009), two diamond inclusion sulphides yielded 2.6 and 3.4 Ga Redepletion ages. They also confirmed, on account of their very low
Pd/Ir ratios, an origin from very depleted SCLM (Fig. 5A). Whole rock
xenoliths have more variable Pd/Ir ratios, indicating possible later
re-enrichment, and therefore their 2.9 Ga Re-depletion ages should
be considered minimum ages for melt-extraction. Three 2.7 Ga and
one 2.0 Ga Re-depletion ages were also found for samples with
no Pd/Ir information available. The 3.4–2.0 Ga Re-depletion ages
match known magmatic events in the crust of the Zimbabwe craton (e.g. Taylor et al., 1991; Mkweli et al., 1995; Dodson et al., 2001;
Jelsma et al., 2004; Rollinson and Whitehouse, 2011; Prendergast
and Wingate, 2013), including the ca. 3.56 Ga oldest known gneisses
(Horstwood et al., 1999) and suggest that the growth history of the
crust is matched by the depletion history in the SCLM.
Using a complementary approach, Nagler et al. (1997) had come
to very similar conclusions regarding growth of the Zimbabwe craton lithosphere. These authors analysed Os-isotopes in low Re/Os
chromites from ultramafic complexes ranging in age between 2.7
and 3.5–3.6 Ga. The chromites are all characterised by very unradiogenic Os and define an evolutionary trend of Re-depletion (Fig. 5B).
The oldest of these samples yielded the most unradiogenic Os
analysed in a terrestrial sample to date, even lower than those
of 3.8 Ga old peridotites from southern West Greenland (Bennett
et al., 2002). Nagler et al. (1997) interpreted the Os-isotope evolution trend to reflect the protracted growth of the SCLM of the
Zimbabwe craton that started long before the age of the oldest preserved gneisses, as far back as 3.9–4.0 Ga. Detrital zircon U/Pb ages
of up to 3.95 Ga (Zeh et al., 2014) are in support of the start of continental growth and SCLM depletion of the Zimbabwe craton around
the H–A boundary.
Episodic growth of Archaean cratons and concomitant depletion of the SCLM has also been proposed for many other cratons
(e.g. Griffin et al., 2003). The available Re–Os and Sm–Nd age constraints for diamond inclusions – the most pristine samples of SCLM
– was reviewed by Shirey and Richardson (2011) who reinforced
the view that peridotitic inclusions have remained isolated from
the asthenosphere for 2–3 Ga.
The clear evidence for very deep Archaean SCLM, in many cases
apparently with quite steep topographic boundaries against adjacent, shallower roots corresponding to Proterozoic ‘mobile belts’ or
younger orogenic structures (e.g. O’Reilly et al., 2009; Sodoudi et al.,
2013), raises the question of how these refractory keels could have
survived for as long as they have. The longevity of the Archaean
SCLM suggests a state of near-neutral buoyancy, which may be an
inherent and ephemeral feature of the Archaean SCLM (e.g. Eaton
and Perry, 2013) on account of its greater extent of depletion and
thus more magnesian (olivine-rich composition) and inherently
lower density than younger SCLM (e.g. O’Reilly et al., 2009). A large
remaining issue is the poorly constrained thermal evolution of the
asthenosphere, which affects the calculated long-term stability of
the SCLM (e.g. Michaut et al., 2009). An additional factor that influences the mechanical stiffness of the SCLM is the water content of
its constituent minerals (Lenardic and Moresi, 1999; Arndt et al.,
2009) because anhydrous minerals increase viscosity. Peslier et al.
(2010) discovered that kimberlite hosted olivine from the SCLM
from >5 GPa has a lower water content than shallower olivine. Furthermore, SCLM-derived olivine may, in general, have a lower water
content than that predicted from olivine in the asthenospheric
MORB-source (Peslier et al., 2010). This would promote a larger
viscosity contrast between SCLM and asthenosphere than temperature and mineralogy alone and may have helped the survival of
Archaean continental lithosphere (Lenardic and Moresi, 1999).
High-resolution geophysical data have also been used to study
the structure of the crustal part of the Precambrian lithosphere.
For example, within the western North American plate, the “Deep
Probe” study was capable of resolving the crustal structure of the
Medicine Hat block from the adjacent Hearne and Wyoming cratonic fragments (Gorman et al., 2002). The “Deep Probe” results also
identified a possibly Proterozoic structure at the base of the crust
of both the Medicine Hat block and the Wyoming craton, which
Gorman et al. (2002) interpreted as a magmatic underplate, added
after Archaean amalgamation of the two units. Crustal underplating has also been inferred from Neoarchaean zircon xenocrysts
retrieved from kimberlites in the North China Craton (Zheng et al.,
2009), showing that addition of new material to the lithosphere
continued both at surface and in the lower crust.
B.S. Kamber / Precambrian Research 258 (2015) 48–82
57
Fig. 5. Data relevant to the formation of the SCLM of the Zimbabwe craton. Panel (A) shows a plot of the rhenium depletion age at the time of eruption (determined from Osisotope composition) versus the normalised Pd/Ir ratio (log scale) from Smith et al. (2009). Two diamond inclusion sulphides (open white symbols) have very low Pd/Ir ratios
indicating their derivation from a very depleted SCLM (that experienced greater than 24% melt extraction, green diagram background). One of these inclusions apparently
experienced Re-depletion at 3.4 Ga. Two whole rock xenoliths (red solid squares) have consistent Re-depletion ages of 2.9 Ga. One has a (very depleted) low Pd/Ir ratio
whereas the two others experienced later Pd-enrichment, shifting their compositions into the field of no apparent melt extraction (yellow background). Horizontal broken
lines denote minimum Re-depletion ages of additional samples for which no Pd/Ir values are known. Panel (B) shows evolutionary trend in Os-isotope space for chromitite
concentrates analysed by Nagler et al. (1997). Panel (C) shows a cartoon of the interpreted lithospheric structure across the southern Zimbabwe craton into the Limpopo
belt (simplified from Smith et al., 2009). Vertical broken lines show positions of studied kimberlites transposed into the schematic S–N transect. The important observations
are the far southern extent of the Palaeoarchaean SCLM of the Zimbabwe craton (green colour) as well as the presence of the ancient Tokwe gneisses (purple) as far as the
southern edge of the craton. The Neoarchaean rocks of the Northern Marginal Zone of the Limpopo belt (orange) overly Palaeoarchaean rocks but have themselves isotopic
memories of contamination with precursors of much greater antiquity (Berger and Rollinson, 1997). The Kaapvaal craton and the associated Southern Marginal Zone of the
Limpopo belt have distinct, more juvenile crust (yellow) and SCLM (blue).
In summary, geophysical and geochemical data are in strong
agreement that the crust-SCLM composite of long-lived Archaean
continental plates grew in unison and that the deep Archaean SCLM
keels were critical for survival of the Archaean continental lithosphere.
3. The Hadean–Archaean boundary
3.1. The Acasta gneisses
In the original sense of definition by Preston Cloud, the
beginning of the Archaean coincides with the oldest preserved confidently dated terrestrial rock. The key reason why Van Kranendonk
et al. (2012) have proposed the H–A boundary to be 4.03 Ga (Fig. 1B)
is the preservation of the Acasta gneissic remnants, components
of which date back to 3.96–4.04 Ga. I prefer the more conventional
H–A boundary definition of 3.85 Ga (e.g. Bleeker, 2004) for two reasons. Firstly, because only components (i.e. small portions of zircon)
of some Acasta gneiss can confidently be dated to 4.03 Ga. Secondly,
because there is strong U/Pb zircon evidence that preservation
probability increased sharply at that time, with Eoarchaean zircon
and rocks present on several cratons. According to the conventional time scale (Fig. 1A), the geological record at Acasta dates back
into the Hadean and the gneisses may therefore inform us about
aspects of the geology of the latest Hadean and the H–A transition
period. It is important to recall that it is physically impossible for
a small enclave of Hadean or Eoarchaean granitoid crust to survive
on its own. Because the enveloping early lithosphere of the Acasta
gneisses is apparently not preserved, the geological information is
biased to the insight from the granitoid gneisses and their isotopic
memory.
Roth et al. (2014) measured Nd-isotope systematics that suggest
that the Acasta gneiss magmas were derived from two sources;
a mafic long-lived Hadean crustal source (as old as 4.3 Ga) and a
younger felsic source with chondritic isotope composition. Reimink
et al. (2014) discovered a 4.02 Ga ferroan variety of Acasta gneiss,
that shares certain geochemical similarities with evolved rocks
of modern Iceland, which themselves are known to be sourced
from re-melted hydrated basalt. Based on this similarity, Reimink
et al. (2014) speculated that the early continents may have formed
58
B.S. Kamber / Precambrian Research 258 (2015) 48–82
in similar tectonic settings as modern Iceland. Whereas I agree
that ferroan grey gneisses may hold important clues to the nature
of the evolved Hadean crust, not the least because they may be
present elsewhere (e.g. the >3.85 Ga Fe-rich Nanok gneisses of
Hebron, Labrador; Collerson, 1983) the analogy with Iceland as a
likely origin for early continental crust is poorly supported. Namely,
Icelandic ryholites and dacites today form due to shallow melting of buried hydrated oceanic crust in response to the unusally
high geothermal gradient in a very peculiar tectonic setting (e.g.
Martin and Sigmarsson, 2010). At Hadean radioactive heat production rates (e.g. Fig. 1D), the coincidence of a mantle plume and
oceanic rift are not required to melt hydrated basalt (see e.g. Fig. 5
of Kamber et al., 2005a).
Regardless of tectonic setting, further geochemical studies of
the Acasta gneisses (e.g. Mojzsis et al., 2014) are supporting the
impression from the field appearance that the oldest variety of
gneiss was migmatised, possibly multiply, within a relatively short
time of formation (less than 100 Ma) of the not yet identified oldest protolith. Very detailed sub-sampling studies of slabbed Acasta
gneisses (following the principle for Rb–Sr studied originally outlined by Collerson et al., 1989) confirm in general the conclusion
of an earlier whole-rock study (Moorbath et al., 1997) that they
do not preserve an original 147 Sm–143 Nd memory (Mojzsis et al.,
2014; Roth et al., 2014). By contrast, sophisticated comparison
of zircon and whole rock Lu–Hf isotope systematics with zircon
U/Pb information, whole rock trace element geochemistry, and zircon geochemistry has revealed two clear isotopic populations of
Acasta gneisses (Iizuka et al., 2009; Guitreau et al., 2014). The older
Hadean, ca. 3.96 Ga group shows a chondritic Hf-isotope initial,
whereas the younger, 3.6–3.7 Ga one shows a sub-chrondritic initial, implying evolution in a (crustal?) reservoir with a somewhat
lower Lu/Hf ratio than coeval mantle.
3.2. The detrital Hadean zircon record
The results from the Acasta gneisses fit well with the available
Hadean detrital zircon record. A set of five consistent observations
has emerged.
(1) Nearly all Hadean zircon have very complex U/Pb age structures (e.g. Nemchin et al., 2006). This was identified as an important
phenomenon by Nelson et al. (2000) and interpreted to reflect a
Hadean crust through which heat was transported in episodic, relatively closely spaced thermal events. In the Acasta gneisses, it
is possible to reconstruct a series of closely spaced events during which the original crust was reworked (e.g. Iizuka et al., 2007;
Mojzsis et al., 2014). But it is also possible to ascertain that the
thermal events were associated with addition of new material in
the form of intrusions, with the most likely crust-forming pulses
having occurred at 4.03–3.94, 3.74–3.72 and 3.66–3.59 Ga (Iizuka
et al., 2007).
In addition to age complexity that reflects directly the time(s)
of Pb-loss, closer study of the oldest Hadean zircon has unearthed
apparent age complexities of limited geological significance. Since
the report of 4404 ± 8 Ma zircon (Wilde et al., 2001), Parrish and
Noble (2003) noted a correlation between U-content of zircon and
apparent ages, suggesting the possibility of elevated 207 Pb/206 Pb
ratios due to complexities in the 235 U decay chain. Their analysis
of the data reported by Wilde et al. (2001) suggests that the oldest
zircon could be closer to 4350 Ma old. Nemchin et al. (2006) carefully studied a number of >4.2 Ga zircon and also concluded that the
oldest bona fide igneous zircon was 4363 ± 20 Ma old. Valley et al.
(2014) re-analysed the purported 4404 Ma grain with a sophisticated atom probe and discovered 6–10 nm diameter-sized clusters
enriched in elements such as Y and Pb, but not U and Th. These
clusters have implausibly high 207 Pb/206 Pb ratios, which means
that radiogenic Pb was mobilised within the lattice in the deep
geologic past, most likely during the growth of the 3.4 Ga secondary
zircon rim. Thus, the oldest confidently dated terrestrial zircon
is 4363 ± 20 Ma but the evidence for repeated high-temperature
reworking of the vast majority of Hadean zircon grains remains
valid.
(2) There is a clear association of Hadean material with cratons that show the existence of crust dating back to 3.75–3.85 Ga
and, in some cases, the formation of highly depleted harzburgitic
SCLM. Thus, a critical pre-requisite for the preservation of Hadean
material evidently was the formation of buoyant and viscous Eoarchaean lithosphere possibly superimposed on pre-existing Hadean
lithosphere (Kamber et al., 2003, 2005a).
(3) Hf-isotopes require the Hadean crust to have been relatively
long-lived (e.g. Kemp et al., 2010), otherwise, the consistent Hfisotope arrays emanating from chondritic initial values would not
have survived. Thus, the Hadean crust was likely being internally
reworked but crustal reservoirs persisted for as long as 500 Ma.
(4) The inferred Lu/Hf ratios of this crustal material are most
compatible with an originally mafic composition of the Hadean
crust. This is apparently also consistent with the Nd-isotope data
reported by Roth et al. (2014). The mafic Hadean protocrust may
also be the ultimate source of the high ␮ signature in the longlived Archaean cratons (Kamber et al., 2003). The notion of a mafic
(basaltic, mantle-derived) Hadean proto-crust was initially at odds
with O-isotope data for Hadean zircon. The original studies by
Wilde et al. (2001) and Mojzsis et al. (2001) found isotopically heavy
O in Eoarchaean and Hadean zircon, prompting the notion of a cool,
potentially habitable Hadean Earth with a liquid hydrosphere. In
the ‘cool early Earth’ model, the magmas from which Hadean zircon
crystallised had interacted with low temperature hydrous fluids.
However, the more detailed follow-up studies (e.g. Nemchin et al.,
2006; Valley et al., 2014) have not been able to confirm any zircon >4.2 Ga in age that contain isotopically heavy O. Whereas this
should not be used as evidence against a Hadean Earth with a liquid hydrosphere (e.g. Kramers, 2003), the >4.2 Ga zircon apparently
crystallised from quite ordinary mantle-derived melts.
(5) Finally, as originally pointed out by Nutman (2001) and
Kamber et al. (2005a) Hadean detrital zircon is absent in Eoand even many Palaeoarchaean sedimentary rocks. Instead, most
Hadean zircon is found in Neoarchaean and even Proterozoic sedimentary rocks (e.g. Wyche et al., 2004). This means that during
the Eo- and Palaeoarchaean, Hadean remnants must have largely
been isolated (buried?) from the sedimentary cycle and were only
exposed after cratonisation. This is also evident at Acasta, where
the final metamorphism and associated unroofing may be as young
as Palaeoproterozic (e.g. Moorbath et al., 1997; Roth et al., 2014).
Kamber et al. (2005a) envisaged that voluminous basaltic outpourings repeatedly resurfaced the earliest Eoarchaean crust. This
process could have led to gradually deeper and deeper burial of
hydrated basalt. The availability of these hydrous fluids at depth
could explain the O-isotopes of <4.2 Ga Hadean zircon (Kemp et al.,
2010), which imply that their source must have interacted, at some
point, with low-T waters. Together with the increasing geotherm
of the thickening Hadean crust the expelled fluids from metamorphic dehydration reactions could have driven amphibolite-facies
melting, yielding Hadean and Eoarchaean TTG magmas (Hoffmann
et al., 2011b).
This set of five consistent observations regarding the Acata
gneisses and Hadean detrital zircon has informed the schematic
model illustrated in Fig. 6 (discussed in detail in Section 3.3.3).
3.3. Destruction of the Hadean crust and the dawn of the
Archaean
If the H–A boundary is defined as the time after which
preservation of terrestrial rocks became more widespread (i.e.
B.S. Kamber / Precambrian Research 258 (2015) 48–82
59
Fig. 6. Schematic block diagram of lithospheric processes operating across the H–A boundary. The synthesis draws together ideas developed by Rapp and Watson (1995),
Kamber et al. (2003, 2005a), Kemp et al. (2010), Hoffmann et al. (2011b, 2014), Abramov et al. (2013), and Johnson et al. (2014). Numbers (e.g. 2b) identify areas described
in Section 3.3. Only a brief description is provided here. The diagram contrasts a slice of stabilised Eoarchaean lithosphere (right-hand side) developed on Hadean substrate
with a transient remnant of unstable Hadean lithosphere (left-hand side). Both types of lithosphere initially formed in the Hadean via a continuum process of drip down into
the mantle (area 2b) and return flow of asthenosphere that melted to create more primary crust (area 2d). Over several 100 Ma, this generated Hadean crust (blue colours)
that internally differentiated (areas 2c and 1g). This crust was underlain by very thin lithospheric mantle root (area 2a). During the late heavy meteorite bombardment, most
of the Hadean lithosphere was destroyed. Only in places where large Eoarchaean mantle upwellings (area 4) generated extensive melt (area 1a) and large basaltic edifices
(area 1c) as well as a deep and depleted SCLM (area 1b), did the lithosphere withstand the forces of giant impacts (area 3). The burial of the Hadean substrate (blue colours) by
Eoarchaean basaltic resurfacing, thickened the crust, perturbed the geotherm and led to further internal differentiation of the Hadean substrate, giving rise to Eoarchaean TTG
(areas 1d and 1e). Importantly, evolved zircon-bearing Hadean rocks were buried at depth and did not shed sediment into the Eo- and Palaeoarchaean surface environment.
at 3.85 Ga), it coincides with the apparent pervasive destruction of the Hadean crust. The near-complete disappearance is
significant because Hf-isotopes of Hadean and Eo- to Palaeoarchaean zircon show that the Hadean crust was relatively long-lived
(e.g. Kemp et al., 2010). Thus, the Hadean crust appears to
have persisted up to ca. 3.85 Ga but then disappeared quite
catastrophically. Although opinions regarding the formation of
Hadean crust diverge into those who favour a largely basaltic
precursor that was internally differentiated and those who prefer a uniformitarian model including subduction (e.g. Mojzsis
et al., 2001; Valley et al., 2002; Harrison et al., 2005; Watson and
Harrison, 2005), there is wider agreement about the fact that the
Hadean crust apparently disappeared quite rapidly between 4.0
and 3.85 Ga. In principle, the destruction of Hadean crust could
have been driven from above, within, or below or a combination
thereof.
B.S. Kamber / Precambrian Research 258 (2015) 48–82
3.3.2. Catastrophic lid collapse and mantle overturn
The alternative explanation for the demise of the Hadean crust
is that it either self-destructed or was consumed in a mantle overturning event (e.g. Griffin et al., 2014). In this regard, it is important
to recall that very early chemical differentiation within the mantle
led to the establishment of isotopic anomalies in daughters of relatively short-lived radioisotopes, in particular the decay of 146 Sm
to 142 Nd. There are now a number of Archaean rocks from which
deviations in the 142 Nd/144 Nd ratio have been reported (e.g. Caro
et al., 2003, 2006; Bennett et al., 2007; O’Neil et al., 2008; Rizo
et al., 2011; Roth et al., 2014) and the majority of these require the
Sm/Nd fractionation event(s) to have taken place between 4.4 and
4.5 Ga (e.g. Boyet and Carlson, 2005; Bennett et al., 2007), within the
time frame envisaged for the magma ocean stage. The formation of
cumulate layers from the crystallising magma ocean is one obvious way of creating chemical stratification (e.g. Caro et al., 2005).
Deep mantle minerals such as perovskites and high-Si garnets have
lithophile element partition coefficients that are able to explain the
diverging behaviours of Nd- and Hf-isotopes in Eoarchaean rocks
(e.g. Hoffmann et al., 2011a; Rizo et al., 2011). There is the additional possibility that very large scale melting in the deep mantle
may have led to olivine floatation and further chemical stratification of the mantle (e.g. Nisbet and Walker, 1982) but since olivine
has very low partition coefficients for Sm and Nd, this process is
not a likely explanation for excess 142 Nd. With regard to the significance of the H–A boundary, I regard the survival of 142 Nd/144 Nd well
into the Archaean as important. If the H–A boundary was caused by
a major catastrophic mantle overturn, it would seem implausible
>3849 Ma Tonalites
3750 Ma<>3849 Ma
Gneisses
<3750 Ma Granitoids
Narryer gneisses
Metasediments
Amphibolites
Pyroxene cumulates
Nd-excess (ppm)
20
I
15
10
II
142
3.3.1. Destruction by bolide bombardment
The most obvious explanation, on account of its apparent temporal coincidence (e.g. Marchi et al., 2014), is that the H–A boundary
was caused and marked by the end of the so-called late heavy
meteorite bombardment (LHMB). The flux of truly immense, basinforming bolides to Earth has recently been revised by Bottke et al.
(2012) and is shown in Fig. 1C. Bottke et al. (2012) made the
point that basin-forming impacts must have still occurred after
3.85 Ga and could have severely influenced the Archaean geological record. As far as the destruction of the Hadean crust is
concerned, however, the important question is how high the highest rate of impacts had been between ca. 4100 and 3850 Ma. In
other words, how narrow was the LHMB spike? According to the
Bottke et al. (2012) calculations, ca. 105 basin-forming impacts
occurred between 4100 and 3850 Ma. This equates to 0.42 impacts
of this magnitude per Ma and is ca. 3 times the rate experienced between 4567 Ma and 4100 Ma. On Mercury, the impact
record for LHMB indicates that bombardment was accompanied
by widespread massive volcanic resurfacing (Marchi et al., 2013).
This suggests a possible link between planetary-scale volcanism
and the LHMB. However, before a LHMB cause for the destruction of the terrestrial Hadean crust is accepted too enthusiastically
(e.g. Marchi et al., 2014), better age constraints for lunar and Martian giant impact basins are required (e.g. Norman and Nemchin,
2014) to test whether the temporal coincidence could simply be
fortuitous. Furthermore, although progress is being made towards
understanding the physical consequences of giant impacts on Earth
(with its higher gravity than the Moon and Mercury and its presence of a hydrosphere; Grieve et al., 2006; Abramov et al., 2013), it
is not yet clear whether basin-forming impacts would have had
the capacity to destroy Hadean crust extensively (e.g. Abramov
et al., 2013), particularly since we know very little about the
make-up of the Hadean lithosphere as a whole. There are a few
detailed U/Pb zircon datasets showing that at least some Hadean
terrestrial zircon have high temperature overgrowths that could
have been caused by reworking during the LHMB (Abbott et al.,
2012).
+
60
5
0
2800
3000
3200
3400
3600
3800
Age (Ma)
Fig. 7. 142 Nd/144 Nd systematics of Archaean rocks with 142 Nd excesses and wellconstrained ages. Modern silicate Earth has a 142 Nd-excess of 0. Granitoid gneisses
are colour-coded into oldest (red: older than 3849 Ma), intermediate (orange:
between 3750 and 3849 Ma) and youngest (yellow tones: younger than 3750 Ma).
Compiled data include those of Caro et al. (2006), Bennett et al. (2007), Rizo et al.
(2011), and Debaille et al. (2013). The data of Debaille et al. (2013) were plotted at
2716 Ma, the best age constraint for the Kidd Munro Assemblage (John Ayer, pers.
comm., 2014). The amphibolite data and values for gneisses 155768 and 155774 of
Caro et al. (2006) were not plotted due to poor age constraints. Of the other data
reported by Caro et al. (2006) ages were chosen as follows: the metasediments were
plotted at 3706 Ma, the combined U/Pb zircon age constraint of two of these rocks
from detrital, volcanogenic zircon (Kamber et al., 2005a); grey gneiss SM/GR/98/26
at 3813 Ma using the U/Pb zircon age reported for this rock by Whitehouse and
Kamber (2002); and potassic granodiorites at 3638 Ma after the U/Pb zircon age
of the related rock 125540 reported by Whitehouse and Kamber (2002). The grey
arrows show tentative evolutions of the maximum preserved terrestrial 142 Nd
excess.
that Archaean magmas could have tapped mantle domains that
experienced Sm/Nd fractionation 4.4–4.5 Ga ago.
Bennett et al. (2007) interpreted a trend of decreasing
142 Nd/144 Nd deviation with decreasing age within the Itsaq
gneisses (Fig. 7; grey arrow labelled I) to possibly record partial
but incomplete remixing of early mantle reservoirs from 3.85 to
3.70 Ga. Caro et al. (2008) analysed ≥3.3 Ga peridotitic garnet inclusions in diamonds from the Kaapvaal craton and these did not yield
any 142 Nd/144 Nd deviations. This hinted at the complete re-mixing
of the high Sm/Nd mantle reservoir by 3.3 Ga. However, Debaille
et al. (2013) discovered a consistent +7 ppm 142 Nd-excess in a high
MgO tholeiite from the Neoarchaean Abitibi greenstone belt (Fig. 7).
This finding may be highly significant because the 142 Nd-excess
was recorded in a juvenile province of the Superior craton and not,
as all the others, from an ancient high ␮ craton. Thus, it is unlikely
that the Nd-isotope signature was inherited from pre-existing crust
(via contamination) but reflects the persistence of a high Sm/Nd
mantle domain to 2.7 Ga (grey arrow labelled II in Fig. 7). Additional
evidence for the persistence of early-formed silicate reservoirs
B.S. Kamber / Precambrian Research 258 (2015) 48–82
comes from the survival of 182 W-isotope anomalies, which resulted
from very early (<35 Ma) evolution in high Hf/W mantle domains.
To date, 182 W excesses have been found in 2.82 Ga old komatiites
from Kostomushka greenstone belt of the Baltic shield (Touboul
et al., 2012) that apparently persisted since the Eoarchaean but
has disappeared in post-Archaean times (Willbold et al., 2011).
Although it is not straightforward to compare the systematics of a
lithophile element (Nd) with that of a siderophile/hydrophile element (W) the isotopic anomalies point to the survival of very early
formed mantle domains well past the H–A boundary. The structure within the Eoarchaean data (Fig. 7) is permissible of a partial
re-mixing of these mantle domains during the H–A transition and
earliest stage of the Eoarchaean (Bennett et al., 2007).
Most numerical models that have attempted to test the likelihood of mantle overturn and concomitant whole-sale destruction
of the Hadean crust have explored the theme of stagnant single
plate versus mobile multiple plate tectonics (e.g. Piper, 2013). The
response of compositional, rheological and thermal properties of
mantle (and lower crustal) material to heat loss on a young planet
is complex. However, modelling results show that stagnant lid tectonics are feasible for Venus (e.g. Reese et al., 1999) and the early
Earth (Piper, 2013). Namely, it is possible to cool the interior of a
planet that is covered by a single plate. Furthermore, evidence from
Mars and Venus shows that very long-lived stagnant lids are effective at allowing heat to escape to the planetary surface (e.g. Choblet
and Sotin, 2001). A feature of stagnant lid tectonics is the possibility
of catastrophic destruction of the single plate after a long period of
heat build-up at depth, possibly during episodes of subduction (e.g.
Turcotte et al., 1999). The H–A boundary could thus be viewed as a
major lid collapse episode (e.g. Griffin et al., 2014) in which most of
the Hadean lithosphere (the lid) was ploughed back into the mantle. Because the anomalies in 142 Nd persisted beyond the Hadean
there could have been repeated intermittent switching between
a largely stagnant lid with short-lived plate tectonic regime (e.g.
Debaille et al., 2013; O’Neill et al., 2013). Future models may need to
also consider strong compositional layering of the Hadean mantle,
inherited from the crystallisation of the magma ocean, which may
have left the mantle in an unstable thermal state (Kramers, 2007).
3.3.3. The dawn of the Archaean
Regardless of the unresolved reason for the Hadean crust’s
demise it is possible to summarise the key observations that have
allowed Hadean material to survive on Earth (see Fig. 6). Because
Hadean material is only found on high ␮ cratons, and because several of those share so many similarities of their post-Eoarchaean
evolution, it is possible and indeed plausible, that Hadean rocks and
minerals have only survived on one single Eoarchaean lithospheric
fragment (Fig. 6, lithospheric fragment labelled 1 and enclosed
by heavy broken line). I propose that this fragment of high ␮
lithosphere was later dispersed into several smaller cratonic terranes. They key phenomenon that allowed its survival past the
H–A boundary was its exceptionally deep and strong lithospheric
mantle root.
By contrast, more typical Hadean crust was underlain by a
very limited mantle root. This is predicted by the sophisticated
numerical model of early Earth crust formation of Johnson et al.
(2014). Their model explores the mineralogical transformation not
only in the mantle but also the lower crust on an Earth with a hotter
mantle. It predicts that asthenospheric convection continuously
eroded the depleted lithosphere below the crust (Fig. 6, area 2a).
This stagnant lid model also predicts the removal of newly formed
crust (light green) as high-density drips (Fig. 6, area 2b) as one
way of losing heat.
Portions of the Hadean crust may have persisted for several
hundred Ma (Fig. 6, lithospheric fragment labelled 2 and enclosed
61
in heavy broken line), and may have undergone internal differentiation (Kamber et al., 2005a; Kemp et al., 2010). This could
have happened via melting of previously hydrated mafic volcanic
rocks (Fig. 6, area 2c) that was rapidly resurfaced during periods of
basaltic outpourings (Fig. 6, area 2d). But the Hadean crust lacking
deep mantle roots apparently vanished during the H–A transition.
Although the potential for destroying this Hadean crust by giant
impacts is yet to be established, the absence of a deep and mechanically strong mantle root (Fig. 6, area 2a) would not have improved
impact survival probability. What is certain is that the Hadean crust
was being reworked by giant impacts, leading to internally differentiated melt sheets (Fig. 6, area 3a), giant ring and radiating dyke
swarms (Fig. 6, area 3b) and impact ejecta (Fig. 6, area 3c), including
condensed rock vapour (Abramov et al., 2013).
By contrast, the high ␮ Hadean–Eoarchaean continent was a
site of rapid new lithosphere formation. Very high-degree melting
(Fig. 6, area 1a) below the high ␮ crust formed the first deep and
mechanically stable SCLM (Fig. 6, area 1b) between 3.9 and 3.8 Ga.
This type of melting could have been the response to a very hot
mantle upwelling (Fig. 6, area 4). Only this lithospheric fragment,
thanks to its greater mechanical strength, survived the end-Hadean
processes that destroyed all other Hadean crust.
The eruption of a thick pile of 3.9–3.8 Ga mafic/ultramafic volcanic rocks (Fig. 6, area 1c) onto the Hadean substrate thickened
the protocrust and perturbed the geotherm, eventually promoting
melting of deeply buried Hadean crust. The more evolved melts
arising from this process contributed to Eoarchaean TTG, at least
some of which contain isotopic evidence for Hadean source material (e.g. Kamber et al., 2003; Iizuka et al., 2009; Guitreau et al.,
2014; Hoffmann et al., 2014).
Because remelting of Hadean substrate and buried Eoarchaean
crust occurred over a depth range (Hoffmann et al., 2014), within
(Fig. 6, area 1d) and beyond (Fig. 6, area 1e) the amphibole stability
field (Hoffmann et al., 2011b), the resulting TTGs had contrasting
HFSE chemistries. Deep crustal re-melting was accompanied by
delamination of eclogite residues (Fig. 6, area 1f). In this model,
the Hadean remnants constituted the lower part of the crust, that
was covered by Eoarchaean middle and upper crust. Although most
of the lower crust would have been mafic in composition, volumetrially minor more evolved Hadean lithologies also survived buried
at depth (Fig. 6, area 1g). Zircon from these rocks were only entering
the sedimentary cycle at much later times, when the high ␮ crust
had attained sufficient mechanical strength to permit exhumation.
This could explain why Hadean zircon is mainly found in sedimentary rocks >1 Ga younger.
4. Archaean oceanic lithosphere?
The loss of heat through the formation of new, buoyant and
hot oceanic lithosphere at divergent ridges and the subduction of
older, colder oceanic plates at convergent boundaries is a hallmark of plate tectonics. It is therefore not surprising that the
identification of oceanic lithosphere has assumed a prominent role
in the long-standing and heated debate about Archaean tectonics. This evidence could come from direct geological observation,
from indirect inferences made about subduction processes or from
numerical models and global thermal and chemical mass balance
considerations.
4.1. The paucity of direct evidence for Archaean oceanic
lithosphere
In view of the preponderance of mafic and ultra-mafic rocks in
the Archaean geological record, and considering how well many
greenstone belts are preserved (e.g. Fig. 8), a non-specialist might
62
B.S. Kamber / Precambrian Research 258 (2015) 48–82
Fig. 8. Field photographs of representative outcrops in the Abitibi greenstone belt. Panel (A) shows a variolitic pillow basalt of the Tisdale assemblage with well-preserved
chilled margin, and ocelli around the pillow rim that can be used as strain ellipses. Panel (B) Micro-banded iron formation of Deloro assemblage, north of Timiskaming, showing
obvious interlayering of silica- and oxide-rich bands. Panel (C) Reworked trachytic surge deposit of the 2687–2675 Ma Timiskaming assemblage showing well-preserved
pumice clasts. The outcrop is locality ‘F’ of Mueller et al. (1994). Panel (D) Well-rounded cobble of K-rich subvolcanic syenite in fluvial conglomerate of Timiskaming assemblage
exposed along ‘Government Road’.
be forgiven for being somewhat surprised at the lack of bona
fide geological evidence in favour of preservation of oceanic lithosphere that has been unearthed over a century of investigation.
Bedard et al. (2013) provided a comprehensive critical review of
the evidence that has been used to argue in favour of Archaean
plate destruction involving oceanic plates. I agree with Bedard
et al. (2013) that there is no compelling evidence in the geological record for Archaean oceanic lithosphere that could have formed
at a spreading ridge. Only a few complementary aspects to Bedard
et al. (2013) are covered in this section.
In a much publicised paper, Furnes et al. (2007) claimed to
have discovered a remnant of an Eoarchaean ophiolite, containing
evidence for ridge spreading in the form of sheeted dykes. However, Nutman and Friend (2007) and several other authors later
showed that there is insufficient evidence for an ophiolite origin of
this small outcrop area and a possibility that the dykes might be
much younger (but see Hoffmann et al., 2011a) and unrelated to
the pillow basalts. Possibly the closest Archaean equivalent to an
oceanic crustal section are the sheeted dykes and pillow lavas of the
2.7 Ga Kam Group in the Yellowknife greenstone belt, Slave craton
(Helmstaedt et al., 1986). However, trace element and Nd-isotope
data show that these tholeiites were contaminaed by variable
assimilation of pre-existing (older) continental crust, and not
emplaced as part of juvenile oceanic lithosphere (Cousens, 2000).
Before attempting to indentify Archaean ophiolites, one very
important point to establish is whether a greenstone belt developed
autochthonously on the continental crust on which it is now preserved or whether it formed elsewhere (in the oceanic realm) and
was accreted or obducted into its current position. With respect to
the high ␮ cratons described in Section 2, there is now considerable
agreement that many greenstones were emplaced onto preexisting continental crust. In the cratons that were least deformed
after cratonisation (e.g. Zimbabwe), it is possible to develop cratonwide greenstone ‘stratigraphies’ that allow lateral correlation of
volcanic units across different greenstone belts (Wilson et al., 1995;
Blenkinsop et al., 1997). It is worth noting that structural data
have been used to argue that at least certain greenstone belts (e.g.
the Belingwe greenstone belt of the Zimbabwe craton; Kusky and
Kidd, 1992) represent obducted oceanic plateaus, but these proposals have not withstood detailed field (e.g. Blenkinsop et al.,
1993), sedimentological and geochemical (Hunter et al., 1998), isotopic (e.g. Shimizu et al., 2005) and inherited zircon (e.g. Wilson
et al., 1995; Shimizu et al., 2004) arguments. Due to the inherently
low incompatible element content of primary mantle melts, the
magmas that now form greenstones were very susceptible to contamination, if they passed through and erupted onto pre-existing
continental lithosphere. Therefore, trace element systematics and
imprinted crustal radiogenic isotope signatures (e.g. Chauvel et al.,
1985) have identified many greenstone belts on the high ␮ cratons
to be autochthonous and ensialic in origin.
A widely held misconception about the uppermost Archaean
continental crust is that it was dominated by TTG, including grey
gneisses. Whereas this is certainly true for the present level of exposure, the make-up of the uppermost crust during the Archaean
was much less felsic. In a landmark contribution, Condie (1993)
showed that if the sedimentary record is used to estimate the
composition of the uppermost Archaean continental crust, a much
more mafic composition is obtained. Thus, the Archaean clastic sedimentary rocks inform us that the main rock type at the
surface was mafic (basalt and komatiite) with subordinate granite and very little TTG. This has since also been confirmed with
evidence from the high Ni/Fe ratio of Archaean banded iron formations, requiring a preponderance of Ni-rich, high-Mg rocks on
the exposed land surface (Konhauser et al., 2009). Voluminous
pulsed outpourings of basalt and komatiite onto of a less dense,
more Si and Al-rich crust, with higher radiogenic heat producing element content leads to gravitational–mechanical instability,
which is expressed in the large-scale geometric arrangement of
greenstone belts. This dome-and-keel structure (Fig. 9) has been
reconstructed from strain pattern analysis in many different cratons (e.g. Bouhallier et al., 1995; Chardon et al., 1996; Collins et al.,
1998; Benn, 2004) and is successfully modelled numerically (e.g.
Robin and Bailey, 2009; Thebaud and Rey, 2013). It is an expression
of the very weak mechanical strength of the Archaean continental
B.S. Kamber / Precambrian Research 258 (2015) 48–82
63
Fig. 9. Diagrams illustrating Archaean (3.2 Ga) intracrustal reorganisation driven by radioactive heat and density inversion. Panels (A–C) show two-dimensional cartoons of
crustal overturn in rapidly grown crust, starting out as a two-layer arrangement (A). This is initially composed of 8 km of mafic crust (grading from green to yellow) overlying
27 km average (andesitic) continental crust (grading from red to yellow). Build-up of radioactive heat leads to mechanical weakening of middle crust (solid orange colour in
panel (B)), partial melting (green to yellow base in panels (B) and (C) show restitic residue) and eventual overturn of the denser basaltic top (dark green drip accumulating
between 24 and 20 km in panel (C)). This concept is qualitatively based on the numerical models of Robin and Bailey (2009). Panel (D) compares two steady-state geotherms.
For both geotherms, the mantle basal heat flow was taken as 30 mW m−2 , the thermal conductivity as 2.5 W m−1 ◦ C−1 , and the surface temperature as 10 ◦ C. The red geotherm
corresponds to the two-layer crustal arrangement illustrated in (A) composed of 27 km andesitic ‘average’ crust and an 8 km lid of MORB. The U, Th and K concentrations
for the composite crust (both layers together) correspond to average continental crust of Rudnick and Gao (2003). Those of the mafic 8 km lid are the MORB values (density
of 2.9 g cm−3 ) from Arevalo and McDonough (2010). The U, Th and K concentrations of the lower layer were obtained by subtracting the content in the 8 km lid from the
composite and distributing the remaining elements into the 27 km andesitic portion. Radioactive heat production rates and geotherms were calculated using the equations
provided in Kramers et al. (2001). The green geotherm corresponds to a 3-layer crust in which the radioactive heat producing elements have been re-arranged into a refractory
10-km thick lower crust (density of 2.9 g cm−3 ) with 6% of the original K, Th and U content, the 8 km thick middle layer (formerly the top layer in a) and a 17 km thick top
layer (density of 2.65 g cm−3 ) containing 94% of the heat producing elements of the original andesitic layer. The block diagram (E) illustrates the final crustal architecture on
a regional (100 km × 100 km × 35 km) scale redrawn after Fig. 13 of Bouhallier et al. (1995), where green colours depict greenstone belts and warm colours internally zoned
diapirs of intermediate (yellow) to granitic (red) compositions.
crust (e.g. Hamilton, 2007), which is important in the context of the
debate of preserving eclogites, another possible “tectonic tracer” of
subduction of oceanic crust.
The absence of truly high P-low T metamorphic rocks
(blueschists and eclogites) of Archaean age, preserved at surface
(e.g. Brown, 2007) should not be used to argue that eclogite and
blueschist did not form at that time. Rather, for the shallow crustal
emplacement and preservation of these high pressure rocks types,
which are not exactly common in the post-Archaean geological
record either, very rapid exhumation is a requirement, regardless of the type of exhumation style envisaged (e.g. Avigad, 1992;
Ernst and Liou, 1999; Wassmann and Stockhert, 2013). The high
radioactive heat production during the Archaean (Fig. 1D) certainly affected the strength of continental crust. It is evident from
the cratonic deformation patterns and implies a weak mechanical
strength of the crust that did not sustain high mountain chains.
Most Phanerozoic blueschists and eclogites are exhumed when
their burial path is reversed in the subduction zone, requiring the
interaction of plates that are rigid even at shallow depth. The lack of
these rock types preserved in the shallow Archaean crust, including interpreted suture zones, informs us about the thermal and
mechanical state of the crust and not of the mantle. In other words,
the lack of these rocks at the exposed Archaean surface does not in
itself constitute evidence against subduction.
B.S. Kamber / Precambrian Research 258 (2015) 48–82
Eclogites are known as xenoliths from kimberlites that erupted
through Archaean cratons. Despite their mineralogic simplicity,
they cover a wide range of petrologic types (e.g. Jacob, 2004) and
at least some are potentially compatible with having originated
from subducted oceanic crust that had interacted with low temperature hydrous fluids. However, when eclogitic minerals are
present as diamond inclusions, they often indicate that the host
eclogite xenolith has been metasomatically modified by later fluids (Smart et al., 2012). Eclogitic xenoliths are notoriously difficult
to date radiometrically (e.g. Jagoutz, 1988) even with modern isotopic tools (e.g. Smit et al., 2014) but it has proven possible to date
sulphide inclusions of eclogitic affinity encapsulated in diamond
(e.g. Richardson et al., 2001). Using a global compilation of dated
eclogite and peridotite diamond inclusions, Shirey and Richardson
(2011) have discovered that eclogites only appear at 3 Ga and that
this age could date the onset of the modern Wilson cycle.
Although the observation of a lack of >3 Ga eclogitic xenoliths is
very strong within the present database, the alternative explanation of this finding is that the eclogites could represent trapped
continental crustal residues (the restites proposed by Rapp and
Watson, 1995) that delaminated from the lower continental crust
and foundered into the largely harzburgitic SCLM. A point in favour
of this alternative explanation is the bi-mineralic nature of most
cratonic eclogite xenoliths. Eclogite xenoliths with mineralogies
typical of subducted oceanic crust (e.g. including kyanite and rutile)
are found to be younger, Palaeoproterozoic in age (e.g. Smart et al.,
2014). Finally, it is worth noting that Ireland et al. (1994) documented that eclogitic diamond inclusions could have been in
equilibrium with TTG melts. At the time of publication of this finding, the prevailing opinion was that TTGs formed by direct melting
of eclogite in subduction zones. However, it is now more widely
believed that many TTGs formed by remelting of thickened, tholeiitic lower continental crust (e.g. Moyen, 2011; Nagel et al., 2012) as
summarised in Section 2.2. Thus, the chemistry of eclogitic diamond
inclusion is also compatible with them being continental restititic
residues that were arrested in SCLM during drip-off. Finally, most
diamonds that host eclogitic mineral inclusions have mantle-like
C- and N-isotope compositions (e.g. Thomassot et al., 2009), indicating that they did not form from recycled, potentially subducted
C.
In summary, there is no compelling evidence for the preservation of Archaean oceanic lithosphere that could have formed at a
spreading ridge. Archaean eclogites and blueschists are not preserved at surface but this could be the consequence of the weak
mechanical strength of Archaean plates and does not preclude their
formation in putative Archaean subduction zones. Cratonic eclogite
xenoliths currently are the strongest geological evidence for subduction but they are difficult to date radiometrically and could also
represent crustal restites that delaminated as residues after TTG
extraction.
4.2. Indirect evidence for Archaean oceanic lithosphere
The bulk of evidence in favour of Archaean plate tectonics has
come from the interpretation of geochemical data and from the
structural fabric of certain Archaean provinces. The former could be
summarised as the ‘arc’ debate and the latter the ‘lateral accretion’
issue.
4.2.1. The arc or dehydration fingerprint
The geodynamic interpretation of geochemical attributes of
both felsic and mafic Archaean igneous rocks is a very controversial issue and has been revisited in numerous reviews (e.g. van
Hunen and Moyen, 2012; Bedard et al., 2013; Arndt, 2013), the
authors of which have come to opposing conclusions regarding
the operation of supra-subduction zone melting in the Archaean.
103
W
N-MORB normalised
64
Pb
102
Be
102
Th
Nb
Li
101
1
10
100
Th
U
Nb Ta La
100
10-1
Nb
Ta
Ti
10-2
Tl Ba Th Nb La Pb Sr Be Hf SmSn Gd Dy HoTm Lu Co Ni
Cs Rb W U Ta Ce Pr Nd Zr Li Eu Ti Tb Y Er Yb Sc Cr
Fig. 10. Extended trace element diagram comparing the composition of a modern
upper crustal sedimentary composite (green circles; from Kamber et al., 2005b) with
that of an Eoarchaean clastic metasediment average (red circles; from Bolhar et al.,
2005). The elements are arranged in order of most (left) to least (right) incompatible
in MORB-style melting (for order of elements and preferred normalisation values
see Kamber et al. (2005b). The normalisation illustrates the great enrichment of
lithophile elements in upper continental crust compared to depleted mantle melts.
The hydrophile elements (light blue backgrounds) are strongly enriched relative to
their normal lithophile neighbours. The high-field-strength elements Nb, Ta and Ti
are highlighted with light brown background and are depleted. The inset shows the
strong depletion of Nb compared to Th. The position of the expected Nb concentration can be estimated from the intersection of the broken line with the light brown
bar. For more information on this diagram see Shirey et al. (2008).
A comparison (Fig. 10) of an extended trace element diagram for a
Phanerozoic upper crustal sediment composite with an Eoarchaean
clastic metasediment average illustrates the similarity in relative
abundance and behaviour of two distinct groups of elements.
Firstly, upper continental crustal magmas of all ages are
enriched in water and in what could be called ‘hydrophile’
elements (Kamber et al., 2005b). Namely those, usually largeion-lithophile, elements that dissolve into the aqueous fluid
phase, which is expelled in prograde dehydration reactions
(Tatsumi and Eggins, 1995). The hydrophile elements also include
elements that, in the absence of aqueous fluids, assume a chalcophile (e.g. Pb, Sn) or siderophile (e.g. W) character. It is
abundantly clear that both Phanerozoic and Archaean upper continental crust is overly enriched in these elements (Fig. 10).
This finding extends to the elements boron, arsenic and beryllium (e.g. Mohan et al., 2008), an overabundance of which
is firmly associated with the subduction process (e.g. Leeman
et al., 1994). Secondly, both upper crustal compositions also
show the well-established depletion in Ti, Nb and Ta, which
is commonly attributed to the presence of a residual Ti-phase
or a Ti-bearing amphibole in the melt source (e.g. Foley et al.,
2002).
Thus, there is strong evidence for element transfer by aqueous fluids and supply of H2 O into the source of Archaean evolved
crust (Arndt, 2013). The crux of the issue is quite simply whether
this dehydration geochemical signature is unique to the subduction zone environment. In other words, is slab-dehydration and
concomitant fluid-assisted melting of the supra-subduction zone
mantle the only petrogenetic environment in which to produce
this geochemical signature? Or, were there alternative geodynamic
settings in which dehydration could have occurred in a suitable
temperature–pressure window? The numerical model of Johnson
et al. (2014) answers the latter question positively. Their model
for the dynamic interplay between the mantle and the lithosphere
in a hotter Earth showed that basalt that erupted at surface and
interacted with the hydrosphere could drip to depths in which
B.S. Kamber / Precambrian Research 258 (2015) 48–82
dehydration occurs within a few Ma, without a need for convergent
plate motion. The production of TTG with ‘arc’ like geochemistry,
including the deficiency in the high field strength elements, can
also occur by melting of over-thickened mafic crust (e.g. Hoffmann
et al., 2011b) but Nagel et al. (2012) pointed out that the tholeiitic
precursor itself must have had an ‘arc’ type character in the first
place.
In summary, many aspects of the trace element geochemistry
of Archaean igneous rocks resemble the modern arc signature.
But important differences remain, including under-appreciated
discrepancies in the major element compositions (e.g. Mg/Al/Si systematics), whose significance needs to be addressed before an arc
setting is uncritically accepted.
4.2.2. Lateral accretion
The Zimbabwe craton has long been known to contain an
interior ‘nucleus’ (the Tokwe segment) in which Palaeo- and Eoarchaean gneisses are abundant. The arrangement of younger, Mesoand Neoarchaean granite-greenstones around this nucleus (e.g.
Wilson et al., 1995; Kusky, 1998) seemed to conform with the
idea of craton growth by lateral accretion of subprovinces around a
proto-Superior craton (e.g. Wilson, 1957). However, with the benefit of better geochronology and isotopic data, this picture has to
be revised.
One indication that the cratonic nucleus could have extended
beyond its originally defined boundaries came with the discovery
of 3456 Ma old gneisses well to the north of the Tokwe segment
by Horstwood et al. (1999) who concluded (p. 707) that the “protocraton is considered to underlie most of the current Zimbabwe
craton”. Smith et al. (2009) summarised Re–Os results of SCLM minerals from kimberlites erupted through the southern Zimbabwe
craton and into the Limpopo belt (well south of the Tokwe segment)
which indicate that ancient, Palaeoarchaean SCLM extends all the
way to the edge of the craton and beyond (Fig. 5C). This is consistent with the discovery of 3.8–3.95 Ga old zircon (Zeh et al., 2014)
from the Central Zone of the Limpopo Belt, which is isotopically of
similarly ancient character as the Zimbabwe Craton (Kramers and
Zeh, 2011). Therefore, the most self-consistent model is that the
younger greenstone belts found at surface developed ensialically
onto an original continent that has not laterally grown by outboard
accretion between 3.8 and 2.6 Ga (e.g. Horstwood et al., 1999).
The arc and terrane accretion model had previously also been
used to explain the fabric of the western Yilgarn craton. With the
benefit of much better field, geochemical, and geochronological evidence, Van Kranendonk et al. (2013) have revised this view. The
key, as in the example of the Zimbabwe craton, was the realisation
that it is possible to construct a cross-craton volcano-stratigraphy
that indicates a shared history, in the case of the western Yilgarn
for at least three volcanic/plutonic episodes and across the eastern
Yilgarn craton, at least from 2720 Ma onwards.
In cratonic areas where less is known about the SCLM, such
as the Nuuk region in SW Greenland (a well-studied part of the
North Atlantic Craton), an ever increasing number of terranes
have been identified (Friend and Nutman, 2005). The area, roughly
40,000 km2 , is proposed to consist of 6 separate terranes that were
assembled in the Neoarchaean, although Friend and Nutman (2005)
state (p. 147) that at least some of the terranes “might be fragments from a larger Palaeoarchaean complex rifted apart from
c. 3500 Ma onwards”. Prior to the introduction of the concept of
tectonostratigraphic terrane assembly, radiometric isotope studies had identified just one ‘terrane’ that experienced a series of
magmatic events and one major tectono-metamorphic episode followed by the largely post-tectonic intrusion of the 2.55 Ga Qôrqut
granite suite, which is widely considered to be of crustal anatectic
origin (see also Section 2.1). It covers ca. 100 km2 and cuts across
four of the proposed terranes yet it is isotopically surprisingly
65
homogeneous (Naeraa et al., 2014) suggesting that the geological complexity seen at surface may not necessarily be shared at
depth. Thus, the complex terrane geology could also be an upper
crustal phenomenon and care has to be taken when equating terrane boundaries with major sutures that extend all the way to the
SCLM.
One of the most widely quoted examples of crustal growth
by lateral accretion of terranes, including ‘oceanic crust’, is the
Neoarchaean Abitibi greenstone belt of Ontario and Quebec. This
sub-province of the Superior craton has experienced strong structural overprint with development of dominant E–W striking fabric
and upright folds. A series of influential papers, including Ludden
et al. (1986), Chown et al. (1992) and Kimura et al. (1993), have
shaped the view that several sub-provinces of the Superior craton,
but in particular the Abitibi greenstone belt, constituted a sequence
of oceanic terranes that were accreted sequentially from north to
south onto the growing craton. The proposed N–S younging direction and identification of ocean crust appeared to constitute very
strong evidence not only for plate movement but also for collisional
tectonics. These ideas were further reinforced when Lithoprobe
seismic transects revealed south-dipping crustal scale fault zones
across the Abitibi greenstone belt that could be linked with shear
zones at surface and interpreted as thrusts and strike-slip structures (e.g. Clowes et al., 1992) and thrust stacking reminiscent of
that in Phanerozoic mountain belts (Lacroix and Sawyer, 1995).
Since then, the rapidly increasing, very high quality U/Pb zircon geochronology database and new mapping revealed a much
more complex evolution than the initial N–S younging direction
proposed by Kimura et al. (1993). By 2002, with the benefit of >2500
very high-precision U/Pb zircon dates (aiding exploration in this
world-class gold mining camp), it had become clear that there is
no younging direction within the Abitibi greenstone belt. Instead, a
largely autochthonous volcanic-sedimentary stratigraphy has been
established (Ayer et al., 2002; Thurston, 2002) with which 9 distinct
units across the entire greenstone belt can be linked across >700 km
mapped by two indepedent geological surveys. The significance of
this finding is that the Abitibi greenstone belt did not grow laterally, regardless of its strong tectonic fabric (which formed after 5
or 6 of the units had been emplaced). It cannot, therefore, be used
as evidence of oceanic arc accretion. The growth style of the Abitibi
greenstone belt will be revisited in Section 5.
In summary, there is no doubt that compressional, extensional and transpressional tectonics operated in the Archaean
crust at various times. The identification of coherent rock packages as terranes and structures that divide them is very useful
to reconstruct regional geological histories. However, in terms of
indirect evidence in favour of subduction, the claimed examples of
Cordilleran-style accretion are not compelling.
4.2.3. General geophysical and global geochemical constraints
Korenaga (2013) provided a detailed overview of the fundamental principles governing plate tectonics. Accordingly, initiation of
plate tectonics depends on many parameters but thermal evolution of the mantle and viscosity contrast across the plate are two of
the most important. Because viscosity is also dependent on water
content, the interaction of the hydrosphere with both the lithosphere and the mantle are a third major influence on likelihood of
plate tectonics. A hotter dry mantle actually makes plate tectonics more likely because water in the mantle lowers the viscosity
contrast. Thus, from a theoretical perspective and as far as current
simulations can predict, Archaean plate tectonics was possible if
not likely. There remain inherent limitations to modelling tectonics on an evolving Earth but it appears that different philosophies
are converging on the consensus that a hotter mantle does not preclude the operation of plate tectonics, at least on a planet with a
hydrosphere.
B.S. Kamber / Precambrian Research 258 (2015) 48–82
1800
Ocean plateau
LIP/continental
rifts
Picritic basalt
MORB
Plume (upwelling)
1600
1500
1400
Asthenosphere
p
T (°C)
1700
A
mixing?
1300
0
1000
2000
3000
4000
Age (Ma)
30
B
CI normalised
concentration
The debate regarding Archaean subduction has tended to focus
on when plate tectonics started, thereby implicitly assuming that
once under-way, plate tectonics could no longer be stopped. This
view needs to be revised. For example, van Hunen and Moyen
(2012) reviewed the importance of the assumptions regarding the
strength of thicker Archaean oceanic plates and argued that despite
the oceanic SCLM being stronger (on account of being fully dehydrated), the crust might have been thick and weak, leading to
frequent slab break-off. Thus, the Archaean might have been a time
of intermittent subduction. In the absence of on-going subduction,
where plate motion is aided by the considerable force of slab-pull,
subduction initiation may have had to re-occur repeatedly. Hynes
(2014) has shown that subduction initiation might have been easier
in the Archaean than it is today. Specifically, this author pointed out
that for subduction initiation, the key parameter is not solely the
magnitude of slab pull, i.e. the extent of density contrast between
the asthenosphere and an oceanic plate column away from the
ridge. Rather, for initiation of subduction, the critical difference in
density contrast is between newly formed oceanic plate and the
aged plate away from the ridge. Hynes (2014) found that even for
high basal heat flows, this difference is large for hypothesised thick
Archaean oceanic plates, increasing the likelihood of subduction
initiation.
From a geological and petrological perspective, it is worth noting that most of these recent computer simulations have attempted
to achieve a fit to the thermal evolution (Fig. 11A) of the potential
mantle temperature (TP ) by Herzberg et al. (2010). Herzberg et al.
(2010) derived their empirical observational dataset from a compilation of 33 “non-arc” basalts (spanning the Precambrian record),
for which their petrological approach returned successful solutions.
These were classified by the original authors reporting the data as:
oceanic plateau basalts (n = 9); large igneous province basalts and
continental rift basalts (n = 5); unclassified picritic basalts (n = 10);
and MORB (n = 9, of which one is a komatiitic basalt). Whereas there
is no reason to question the resulting TP calculations, it is necessary
to recall the inherent issue regarding the scarcity of obtaining Precambrian samples from which the primary melt composition of an
asthenospheric mantle origin could be obtained. Namely, if subduction was operational (at least episodically) as predicted by the
models, then spreading ridges must have existed and these would
have produced oceanic lithosphere in the modern sense (even if
thicker), including MORB, getting progressively older away from
the ridge.
By way of example, rare earth element (REE) patterns of two of
the 9 rocks listed by Herzberg et al. (2010) as being of MORB affinity (‘MORB-SSZ’, where SSZ probably stands for supra-subduction
zone) are shown in Fig. 11B. These are from the 3075 Ma Ivisaartoq
greenstone belt (southern Greenland) and were reported by Polat
et al. (2008). Sample 485418 returned a TP of 1560 ◦ C and sample 485420 a slightly lower TP of 1532 ◦ C (Herzberg et al., 2010). It
is clear that their REE patterns are not light-REE depleted, a hallmark feature of MORB, and therefore do not represent ordinary
asthenospheric mantle melts. Instead, Polat et al. (2008) proposed
a two-stage evolution for the source of these rocks, originating
in the asthenosphere but being re-fertilised in an oceanic arc.
Thus, the original authors did not propose these rocks to have
formed at a ridge but at an extensional fore-arc above a subduction
zone.
The survey of the extensive literature on the subject of Archaean
oceanic lithosphere (Section 4.2.1) established a very wide agreement that no known Archaean tholeiitic basalt actually represents
the equivalent to N-MORB with unequivocal evidence for the ‘R’, as
in ridge. Logically then, the secular evolution of the non-arc mantle
temperature is near impossible to constrain. One of the oldest, and
best-studied Precambrian suites of MORB-like tholeiites is from the
1.9 Ga Flin Flon belt in northern Manitoba. As shown by Stern et al.
20
Modern N-MORB
1.9 Ga Moen Bay
3.075 Ga '485418'
3.075 Ga '485420'
10
9
8
7
La Ce Pr Nd
SmEuGd Tb Dy Ho Er TmYb Lu
1200
C
1000
800
n
66
600
400
200
0
3
6
9
12
15
18
21
24
27
30
MgO (wt%)
Fig. 11. Panel (A): The non-‘arc’ mantle potential temperature (TP ) reconstruction of Herzberg et al. (2010), plotted according to the geochemical/petrological
sample type of the original authors. Panel (B) shows chondrite-normalised (Sun
and McDonough, 1989) REE patterns for modern N-MORB (average of 25 sampled
reported by Arevalo and McDonough, 2010), 1.9 Ga interpreted MORB tholeiites
(average of 11 Moen Bay tholeiites from Babechuk and Kamber, 2011), and two
3075 Ma pillow basalts (from Polat et al., 2008). The latter were used by Herzberg
et al. (2010) for TP reconstruction but like most samples used show significant light
REE enrichment instead of the required depletion. Panel (C) Histogram of MgO
content of 9119 Archaean metavolcanic rocks (from the pre-compiled Archaean
database in GEOROC) with MgO of more than 3 wt% (to remove silicic mode) and
less than 30 wt% (to remove cumulates).
B.S. Kamber / Precambrian Research 258 (2015) 48–82
(1995), these rocks not only have a depleted mantle Nd-isotope signature but also the required depletion in light REE (Fig. 11B). In my
opinion, only such rocks should be used to attempt a reconstruction
of TP . In addition, modern estimates of TP benefit from the fact that
olivine can be studied, whose Mg-content is an important parameter to work out parental melt composition (e.g. Putirka et al., 2007).
But unlike komatiites, very few Archaean basalts preserve olivine. A
histogram of MgO content of Archaean volcanic rocks shows a wellresolved peak at 7 wt% (see Fig. 11C), similar to modern MORB (e.g.
Putirka et al., 2007). These certainly do not represent parental melt
compositions but it is presently unclear whether they derived from
a mantle that was hotter than at present.
An interesting feature of the data compiled by Herzberg et al.
(2010) is the large spread in apparent TP (Fig. 11A) at any one age.
This might by itself be an indication that hot (upwelling) mantle
co-existed with asthenosphere of much more modest (1400 ◦ C)
temperature. In this interpretation, the basalts used by Herzberg
et al. (2010) are hybrids between magmas from hot upwellings and
melts from ambient mantle with as yet unconstrained but lower TP
(Fig. 11A). Thus, my critical reading of the very few publications
that have attempted to constrain non-‘arc’ TP is that the existing
data are inconclusive. Arndt (2013) proposed that the upper mantle temperature has experienced episodic oscillations, caused by
the arrival of very large pulses of hot plumes from the lower mantle.
This concept also calls into question a uniformly hotter Archaean
asthenosphere. Until better consensus is reached regarding the Precambrian asthenospheric mantle TP it is premature to accept the
predictions of numerical tectonic models, whose most important
input is TP .
This review of the knowledge pertinent to Archaean plate tectonics and subduction concludes with the unsatisfying statement
that the evidence neither argues firmly against nor strongly in
favour. In my opinion, however, the concept of plate tectonics has
been over-stretched to explain phenomena of the Archaean geological record and the importance of arc accretion in particular,
has been over-emphasised. Should we be surprised that there is so
little evidence for Archaean oceanic lithosphere and subduction?
Perhaps not. The Archaean continental lithosphere was in a very different thermal and mechanical state and ill-equipped to preserve
evidence of subduction and oceanic crust formation. Even today,
most low-grade obducted ophiolites are exposed in mountain belts
that will in time be eroded. High-grade metamorphic oceanic rocks
are typically preserved as eclogites and, for reasons discussed, could
not be preserved in the Archaean continental suture zones. Most
greenstone belts on pre-existing older continents inform us about
the conditions of mantle upwellings. There is excellent evidence
that Archaean ‘plumes’, or more generically, ‘upwellings’ were substantially hotter than in the Phanerozoic, but the indication for a
hotter ‘mantle’, which really refers to ambient asthenosphere into
which the plumes were emplaced is, in my opinion, quite weak.
I agree with Arndt (2013) that Archaean subduction is the
most plausible means of delivering aqueous fluids to sufficient
depth to produce granitoid melts and to supply the hydrophile
elements into the source of Archaean igneous rocks that have an
arc signature. A further important point was made by Bickle et al.
(1994), who concluded that the existence of deep-water ocean
basins can be inferred from continental freeboard and from the
fact that many continental Archaean basins were deposited in shallow water. However, Bickle et al. (1994, p. 121) also concluded
that “an unequivocal Archean ophiolite has not yet been recognized”. Twenty years later, in my view, we still have not found a
greenstone belt that truly represents oceanic lithosphere created
at spreading ridges (see also Pearce, 2008). Subduction may have
been intermittent through the Archaean, possibly helped by the
episodic arrival of mantle upwellings in the asthenosphere (Arndt,
2013), thereby heating the asthenosphere and aiding subduction
67
initiation and possibly continental crust preservation and net
growth (Hawkesworth et al., 2010). In this respect, it may be
important that the highly siderophile elements that were delivered during the LHMB were not fully mixed back into the mantle
source of komatiites until 2.6 Ga (Maier et al., 2009). It is difficult to
envisage how mantle convection in a plate tectonic regime could
have been so sluggish as to prevent homogenisation (e.g. O’Neill
et al., 2013). Periods of plate tectonics interspersed with stagnant
lid episodes may therefore best satisfy the collective observations.
5. Rapidly grown juvenile Archaean continental crust
Having described the geological, geochemical and isotopic evidence for very long-lived continental lithosphere and having
summarised the evidence for and against the production of true
oceanic lithosphere, we can now better appreciate the significance
of the second type of Archaean continental lithosphere.
5.1. Archaean volcanic plateaux
The idea of a different type of continental crust was developed in
a series of publications by van Kranendonk and co-workers (including: Van Kranendonk and Pirajno, 2004; Smithies et al., 2005, 2009;
Van Kranendonk, 2010) working on the eastern portion of the Pilbara craton. There, in contrast to the long-lived (in my terminology
high-␮) cratons the role of magmatism with an arc-signature is
only of subordinate importance. Instead, the defining feature of
the East Pilbara terrane is a very thick coherent, normally younging stratigraphy dominated by volcanic rocks of mafic–ultramafic
composition. This stratigraphic succession is known as the Pilbara
Supergroup and has a cumulative preserved thickness of 17 km.
This almost certainly overestimates the true stratigraphic thickness, because not all formations are found across the entire cratonic
expanse. Regardless, they are sufficiently well developed to define
a division into three autochthonous groups. The East Pilbara craton
thus contains very strong evidence for repeated volcanic resurfacing, leading to a volcanic stratigraphy.
All three groups contain evidence for basaltic–komatiitic magmas with a plume affinity, having lithophile element characteristics
distinctive from the arc and MORB signatures (e.g. Smithies et al.,
2005). The volcanic-sedimentary history of the three groups spans
just under 300 Ma, from 3.52 to 3.23 Ga. The combination of vertical
growth, sustained and pulsed volcanism and the plume character of
the volcanic magmas, have led to the comparison with thick oceanic
plateaux (Van Kranendonk, 2010). The lower base of the East Pilbara
Supergroup is not exposed but there is geochemical, isotopic, and
limited zircon evidence that it was, at least in part, built in proximity
to pre-existing continental crust (e.g. Tessalina et al., 2010). However, many uncontaminated lavas are found (e.g. Smithies et al.,
2005) suggesting that the main substrate could have been juvenile
oceanic lithosphere with limited potential to contaminate rising
plume magmas.
Whereas Van Kranendonk (2010) uses the dominance of plume
vs. arc type signature as the main criterion to distinguish between
the two types of Archaean continental crust, I prefer to distinguish
between lithosphere that grew very rapidly versus that which grew
over more than a billion years. Where the two views clearly agree
is the significance of a coherent terrane-wide volcano-sedimentary
stratigraphy with a sense of upward younging – the concept of a
volcanic plateau. In Kamber (2010), I used the secular evolution of
the REE chemistry of marine chemical sediments and their Sr- and
Nd-isotope systematics to infer that apart from the highly evolved
continental landmass, a second type of land, emerged oceanic
plateaux, must have existed and been as voluminous if not more
abundant than typical continental land.
68
B.S. Kamber / Precambrian Research 258 (2015) 48–82
Fig. 12. Comparison of oceanic plateau crustal sections and their potential to store water. Panel (A) shows the schematic lithostratigraphy of the Abitibi greenstone belt after
original Fig. 8 of Ayer et al. (2002). The 9 successive assemblages are labelled on the left-hand side, and depositional age ranges are given in (Ma). Conformable contacts are
marked as straight lines whereas unconformities are indicated by undulating lines. Panel (B) shows Arndt’s (2013) petrologic reconstruction of a modern ocean plateau crustal
section (after original Fig. 5.9). For both sections (A) and (B) the vertical scale is on the order of 15–20 km. Panel (C) shows the relative potential of these sections to hold water.
The fat broken dark blue line shows Arndt’s (2013) interpretation of water impoverished oceanic plateau crust (B) in which cumulates and gabbros are essentially anhydrous.
The blue shaded areas are a schematic illustration of deeper crustal water burial caused by volcanic resurfacing over the 75 Ma course of 9 successive volcanic-sedimentary
episodes recorded in the Abitibi greenstone belt (A).
As already mentioned in Section 4.2.2, the Abitibi greenstone
belt is another example of a terrane with a coherent volcanosedimentary stratigraphy (Fig. 12A) that has grown over a very
short period of time. Mafic–ultramafic lavas were emplaced as subaqueous, often pillowed flows (Fig. 8A) in 7 conformable cycles
between 2751 and 2697 Ma, prior to deposition of two unconformable sediment-dominated assemblages between 2696 and
2675 Ma (Ayer et al., 2002; Thurston, 2002). Similar to the East Pilbara craton, the stratigraphy is best exposed on the edges of synto post-volcanic composite granitoid domes that have punched
through the originally near-horizontal geometry (e.g. Ayer et al.,
2002; Benn, 2004). The domes themselves often show zonation in
terms of composition and age, exposing the least potassic and most
deformed oldest tonalites as grey gneisses along the periphery.
Although intermediate and felsic volcanic rocks are in the minority, they have been pertinent for the reconstruction of a greenstone
belt-wide chronology (e.g. Corfu, 1993; Ayer et al., 2002). Based
on this chronological database, it is evident that the pulses of
magmatism were relatively short-lived, on the order of between
3 and 15 Ma. The rocks deposited in between magmatic pulses
are very thin sedimentary units of chert, wackes and bandediron formation. They contain field and petrographic evidence
for sediment-starvation and suggest that for the most part, the
topography of the growing plateau was relatively minor (Thurston
et al., 2008). The geochemistry of the cherts and banded iron formations is consistent with long-term exposure of volcaniclastic flow
B.S. Kamber / Precambrian Research 258 (2015) 48–82
tops to Si-rich seawater and continued hydrothermal activity in the
oceanic column (Thurston et al., 2012).
Although this stratigraphic framework can be applied across
the entire greenstone belt, it is worth noting that the stratigraphy is not fully developed in the entire belt. Many assemblages
are lens-shaped which likely means that eruption centres were
feeding flows locally. It is therefore not possible to estimate one
single thickness of the composite stratigraphy. The sum of the maximum thicknesses for the 7 lower assemblages is 46 km (Ayer et al.,
2002). This incomplete development of the stratigraphy should
not be taken as evidence against vertical growth by volcanic resurfacing. A very strong piece of information that supports vertical
growth is the age distribution of inherited zircon. Namely, in the
younger units, it is possible to recover inherited zircon that matches
in age exactly the emplacement of the older (by inference underlying) volcanic units (Ayer et al., 2002 and references therein).
Very careful, high-precision single-grain U/Pb zircon studies were
necessary to resolve these details that were not available when
the arc-accretion model with a N–S younging had been proposed
(see Section 4.2.2). Many geochemical studies (whole rock Sm/Nd,
zircon Lu/Hf, feldspar and ore Pb–Pb) conclusively show that the
Abitibi plateau is largely of juvenile character and was sourced
from contemporaneous depleted mantle (e.g. Carignan et al., 1993;
Vervoort et al., 1994; Corfu and Noble, 1992). Only towards the
western edge, do Hf-isotopes in zircon suggest that the Abitibi
plateau grew in lateral proximity to existing, evolved continental crust (the Wawa crust; Ketchum et al., 2008). As with the East
Pilbara terrane, the substrate onto which the oldest mafic volcanic
rocks were deposited is not exposed. In the case of the Abitibi greenstone belt, the only indications for its character are the absence
of inherited zircon older than 2.75 Ga (i.e. older than the oldest
assemblage of the Abitibi greenstone belt itself) and its inability to
contaminate the juvenile basalts, both attributes compatible with
plateau construction onto un-evolved (basaltic?) ocean floor.
5.2. Differences between modern oceanic plateaux and the
envisaged formation of Archaean submarine volcanic plateaux
Use of the term oceanic plateau understandably carries the danger of direct comparison with modern oceanic plateaux, which
differ in many ways from the type of terrane being described here.
Three of these differences are discussed here.
5.2.1. Syn-volcanic granitoids
The most obvious difference are the widespread composite
granitoid domes that are found at the present level of exposure.
These domes are not batholiths in the modern sense. They occur
throughout the entire terrane with no preferred geometric arrangement (e.g. Van Kranendonk, 2010) as can, for example, be found
along the western edge of the modern American plates. The domes
were long-lived and emplaced as composites of low-viscosity but
crystalline rocks, crystal-mush mixtures (migmatites) and melts
(e.g. Collins et al., 1998). Their upward mobility was driven by buoyancy arising from the combined effects of density inversion of the
overlying thickening basaltic plate and radioactive heat build-up
in the middle and lower crust, already discussed in Section 4.1 and
illustrated in Fig. 9. In the East Pilbara craton, a 10–100 Ma period of
“conductive incubation” preceded the formation of the main domeand-keel structure (Sandiford et al., 2004). Of critical importance for
the heat budget was the burial of granitoids of the first magmatic
cycle below the thick Euro basalt.
In the Abitibi greenstone belt, there is an exact correspondence
in age of plutonic rocks with volcanic episodes recorded in the volcanic stratigraphy (e.g. Ayer et al., 2002; Benn, 2004; Ketchum et al.,
2008). This strongly suggests that the heat deposited by rising mantle melts feeding the volcanic centres helped to trigger anatectic
69
differentiation of the lower crust. Benn and Kamber (2009) interpreted granulites of the Kapuskasing Structural Zone to be restitic
remnants of the plateau that suffered repeated thermal pulses in
response to internal crustal differentiation and refinement. The
importance of radioactive heating and the vertical transport of
K, Th, and U is nicely illustrated on air-borne radiometry maps
of the East Pilbara domes and the internal geochronology of the
domes (Fig. 1 of Van Kranendonk and Pirajno, 2004). As a general rule, the younger phases of the dome are found towards the
centre (e.g. Van Kranendonk and Pirajno, 2004) and the granitoids
became progressively more potassic with time (e.g. Bickle et al.,
1993) as illustrated in Fig. 13. The geochemistry and Nd-isotope
systematics of many of the granitoids are compatible with an origin of melting pre-existing basalt followed by re-melting of the
less potassic older granitoids (e.g. Smithies et al., 2009) with relatively limited addition of juvenile material. In this context, the
increasing concentration of radioactive elements that is evident in
the radiometric maps shows that as the plateau was maturing, with
on-going crustal differentiation, the extent of re-melting decreased,
leading to smaller degree melts that became very highly enriched
in heat-producing elements.
5.2.2. Burial of hydrated basalt
The second difference relates to the extent of hydration of the
plateaux. Arndt (2013) dismissed the idea that voluminous granitoids could be formed by remelting of hydrated plateau basalt on
account of a lack of sufficient water, as, for example, envisaged by
Smithies et al. (2009) for the East Pilbara Craton. Arndt (2013) constructed a possible crustal lithostratigraphy for the Ontong-Java
plateau that satisfies the constraint that the parental melt had ca.
18 wt% MgO, yet the erupted basalts only ca. 8 wt% MgO. The resulting section, with a very thick basal cumulate layer, is shown in
Fig. 12B. Due to the large proportion of cumulates and intrusive gabbros, this section has a low potential for hydration and a tendency
of storing most of the water in the top, far too high for any remelting. However, as the comparison with the Abitibi greenstone
belt stratigraphy shows (Fig. 12C), the analogy is over-simplified
and illustrates the problem of calling the ancient terranes oceanic
plateaux. Rather, the vertical stratigraphy in the Abitibi greenstone
belt consists of well-resolved units, deposited as innumerable subaqueous volcanic flows and hyaloclastites (e.g. Ross et al., 2011), all
of which were hydrated. Indeed, extensive low-temperature alteration has affected the basalts so pervasively that none of the original
phenocrysts are preserved to conduct proper parent melt composition estimates. In addition to the hydration of the units in the
known part of the stratigraphy, there is potential to store even more
water in the oceanic crustal substrate onto which the plateau was
emplaced, which could potentially carry an older isotopic signature (e.g. Smithies et al., 2009). Burial of hydrated mafic lithologies
by volcanic resurfacing and sagduction remains a viable option to
produce the granitoid domes.
5.2.3. Thick cratonic mantle keels
The final feature that distinguishes the ancient volcanic plateaux
is that they are underlain by a thick, depleted keel of SCLM (Van
Kranendonk, 2010). The Pilbara craton as a whole has an SCLM of
almost 200 km depth, which contrasts with the even deeper and
cooler lithosphere of the adjacent high ␮ Yilgarn craton. Below
the Abitibi greenstone belt, the SCLM is between 200 and 240 km
thick (Darbyshire et al., 2007). By contrast, the base of the lithosphere below the Cretaceous Ontong Java plateau is at 120 km (e.g.
Ishikawa et al., 2011). Thus, most evidence appears to indicate that
the Archaean plateaux lithospheric residue is much thicker.
In terms of composition, the shallow part of the Ontong Java
mantle lithosphere (at ca. 80 km) is dominated by only mildly
depleted lherzolite typical of Pacific oceanic mantle. Between 80
70
B.S. Kamber / Precambrian Research 258 (2015) 48–82
Fig. 13. Granitoid evolution in the East Pilbara craton (after Smithies et al., 2009). The U/Pb zircon ages (yellow to pink boxes) define 7 granitoid events that produced
increasingly potassic magmas, culminating with monzo- and syenogranites. The main TTG forming episode occurred early in the evolution, between 3.50 and 3.45 Ga and
is marked by the taupe-coloured bar. Formation of the SCLM continued between 3.45 and 3.30 Ga and is indicated by the light green bar. The stippled boxes are Nd-isotope
model ages of granitoids (numbers indicate the number of data points, no number equals one data point).
SCLM can be reconstructed from garnet crystals recovered from
mantle-derived potassic volcanic rocks. The garnets are believed
to originate from disaggregated host mantle rocks. A global survey of mantle garnet chemistry shows a much higher degree of
depletion in Al and Ca during the Archaean compared to the Proterozoic and the Phanerozoic (Fig. 14). This implies a much higher
PRIMITIVE MANTLE
(=ASTHENOSPHERE)
4
3
% CaO
and 120 km depth, the local lithospheric mantle is more highly
depleted but has complex Re-depletion ages. There appears to exist
a much older component unrelated to the formation of the Cretaceous plateau and the plume magmatism may only have left a
relatively thin (10 km?) layer of ultra-refractory lithospheric mantle (Ishikawa et al., 2011). Thus, in reality, the true difference
in depth of the lithospheric roots between modern and ancient
plateaux may be much larger than 100–120 km.
Relatively little is known about the composition of the SCLM
of the East Pilbara craton directly from mantle xenoliths. In the
Abitibi greenstone belt, kimberlite-hosted mantle xenoliths from
the Kirkland Lake region inform us about the relatively high proportion of garnet lherzolites (Meyer et al., 1994). These show
clear signs of re-fertilisation of previously highly depleted harzburgite. Miller et al. (2012) argued that the original Neoarchaean
root of the SCLM of the southern Superior Province, including
the Abitibi greenstone belt, had been destroyed by tectonic erosion, later magmatic events and the influx of asthenospheric
fluids. The geochemistry of the Neoarchaean komatiites of the
Abitibi greenstone belt requires both garnet-bearing harzburgite
and ultra-refractory dunite in the residual source region (Sproule
et al., 2002).
Regardless of these complexities, there are two important
aspects of the ancient plateau’s lithospheric mantle. Firstly, its
highly depleted nature implies that it largely formed via high to
very high degree melting. This results in a very distinctive mineral
chemistry in the residual peridotite. Griffin et al. (2003) provided a
comprehensive review of the chemical evolution of SCLM through
time. The well-established greater extent of depletion of Archaean
PHANEROZOIC
2
PROTEROZOIC
1
ARCHAEAN
0
1
2
3
% Al2O3
4
5
Fig. 14. Extent of estimated SCLM depletion expressed as a cross-plot of the incompatible elements CaO vs. Al2 O3 (simplified from Griffin et al., 2003; their Fig. 1). The
compositions were calculated from garnet concentrates and show a clear evolution
from most highly depleted Archaean to least depleted Phanerozoic SCLM.
B.S. Kamber / Precambrian Research 258 (2015) 48–82
degree of melting. Importantly, modern continental crust formation, dominated by arc magmatism, leaves behind a much less
depleted residue that also differs in many other attributes (e.g.
Mg/Si ratio) from Archaean SCLM (Griffin et al., 2003). Therefore,
the main driving force behind magmatism in the ancient oceanic
plateau were very high temperature plumes. Alternative interpretations that invoke plume–arc interaction have been offered by
Wyman et al. (2002) and Wyman and Kerrich (2002).
The second key feature is the much greater depth of the lithospheric keel. The Abitibi lithosphere consists of 185 km SCLM and
ca. 35 km crust (Darbyshire et al., 2007). Extraction of 30% melt
from originally fertile mantle would have produced a total melt
column of almost 60 km basalt, much more than the preserved
crustal thickness. However, the preserved crust is, on average,
andesitic and mass balance requires the delamination of lower
crustal high-density residues that must have foundered through
the growing lithospheric mantle root. In the East Pilbara craton,
TTG formation occurred relatively early (Fig. 13) and most of the
eclogite delamination may have happened before emplacement of
the plume-depleted harzburgite (Smithies et al., 2009) in which
eclogite could be trapped. This is an important point because it
could explain why, in global datasets, eclogite diamond inclusions and xenoliths apparently only occurred after 3 Ga (Shirey
and Richardson, 2011), at least >500 Ma after the oldest peridotitic
sulphide inclusions in diamond.
Ayer et al. (2002) commented on the irregular thickness of the
various volcanic associations across the Abitibi, yet the preserved
crustal thickness is very homogenous. This observation, in agreement with the evidence for dome and keel structures (e.g. Benn,
2004) implies that a positive feedback mechanism regulated crustal
thickness. This likely was the accumulation of radioactive heat by
burial of K, U and Th-bearing felsic lithologies and the continued
resurfacing with basalt (Mareschal and Jaupart, 2013). The resulting reduction in viscosity of the lower crust weakened the crust
as a whole until the heat producing elements migrated upwards
with granitoid melts. Flament et al. (2011) argued that a higher
Moho temperature may have been an additional factor ensuring
that most Archaean plateaux remained subaqueous, despite considerable thickness of the melt column.
There is excellent evidence, both in the East Pilbara craton
and the Abitibi greenstone belt, that once plume magmatism
waned and the crust was re-organised in terms of heat producing elements, the geology started to assume an altogether
more familiar look to that of the Phanerozoic. In the Abitibi
greenstone belt, the 7 main volcanic pulses were disconformably
overlain by two sediment-dominated groups (Fig. 12), the Porcupine and Timiskaming assemblages deposited into well-defined,
fault-bounded, late-orogenic basins (e.g. Mueller and Corcoran,
1998). The Porcupine assemblage is dominated by sandy turbidite deposits and is unconformably overlain by the mixed
sedimentary-volcanic Timiskaming assemblage (Mueller et al.,
1994). This contains remarkably well-preserved trachytic and porphyritic alkaline lava flows and pyroclastic surge deposits (Fig. 8C).
The sedimentary facies is represented by alluvial–fluvial conglomerates and sandstones. Forming a trend roughly parallel to the late
basins, are ubiquitous stocks of strongly alkaline intrusive rocks,
including syenite and subvolcanic porphyries. Importantly, these
occur as well-rounded pebbles and cobbles in the contemporaneous conglomerates (Fig. 8D). The fact that cobbles of subvolcanic
monzonites can be found in fresh-water conglomerates illustrates
that the crust was by now sufficiently strong to support a topography, which in turn allowed rapid exhumation, akin to modern
molasse basins (Mueller et al., 1994). In the much more highly
deformed East Pilbara craton, Van Kranendonk et al. (2010) found
evidence for a ca. 3.2 Ga period of craton-wide extension, forming basins that were initially filled with coarse-clastic sediments.
71
However, according to Smithies et al. (2009), truly potassic magmatism only occurred at 2.8 Ga (see Fig. 13).
In summary, the higher temperature of Archaean plumes evidently permitted the vertical growth of volcanic plateaux of
significant thickness. Unlike plumes that form continental flood
basalts in a matter of 2–5 Ma, the Archaean plateaux appear to
have witnessed repeated magmatic episodes (e.g. Sproule et al.,
2002), eventually producing edifices that began to differentiate
internally, due also to the higher radioactive heat production rate
(Fig. 1D). From the geological record, it is not possible to estimate how common oceanic plateaux were in the Archaean because
of preservational bias. Only those equipped with a strong SCLM
have survived and they would also represent the most internally differentiated. The REE chemistry and Ni/Fe systematics of
Archaean chemical sediments are compatible with the idea that
plateaux were common (Konhauser et al., 2009; Kamber, 2010).
Their relative paucity in the geologic record therefore suggests that
only few became sufficiently evolved to mature into continental
lithosphere.
6. The Archaean–Proterozoic boundary
The definition of the A–P boundary remains a difficult task (see
e.g. Van Kranendonk et al., 2012) yet this transition is meant to
mark the end of the most prolific crust generation and preservation period experienced by the Earth (e.g. Hawkesworth et al.,
2010). Historically, 19th century field geologists began to develop
a sense that there might have been a significant turning point
between deposition of Archaean and Proterozoic rocks. The geologist most widely credited with the first attempt at defining the A–P
boundary was Logan (1857). Working on the southern Canadian
Shield, he noticed the difference between the structure, deformation style and map patterns of Palaeoproterozic sedimentary rocks
of the Huronian Supergroup and the underlying Archaean rocks
of the southern Superior province (then called the ‘Laurentian’
series). An example of the contrasting map patterns is shown in
Fig. 15, where the gently folded Huronian Supergroup sedimentary
rocks rest unconformably on the steeply dipping, much more complexly deformed Archaean granite-greenstones. Notwithstanding
Logan’s (1857) visionary insights (particularly with respect to
the oxidation state of iron in the Huronian sedimentary rocks),
it was perhaps somewhat fortuitous that the rocks in question
indeed turned out to be of the ‘correct’ absolute age for the A–P
boundary.
We now know that geological phenomena compatible with
modern-style, Wilson-cycle plate tectonics are recorded in the
Archaean rocks record, as far back as 3.2 Ga (e.g. Heubeck and
Lowe, 1994; Van Kranendonk et al., 2010) or possibly even further.
This includes particularly the late rift basins and the compressional deformation patterns established during the final assembly
of the Subprovinces into the giant Superior craton. Some of
these processes operated when other Archaean terranes were
still being shaped by non-uniformitarian processes. For example,
Mesoarchaean sediments (e.g. Moodies Group of the Kaapvaal craton; Heubeck and Lowe, 1994) resembling modern facies were
deposited well before final cratonisation. Equally, certain aspects
of greenstone belt architecture have been compared to nappes that
formed prior to the intrusion of late-orogenic potassic granitoids
(e.g. Dirks and Jelsma, 1998). However, this apparent contradiction is simply a consequence of the fact that the geology of the
Archaean continental crust was strongly governed by the distribution of the heat-producing elements. Namely, the upward transport
of radioactive heat producing nuclides in continental crust was
episodic, leading to periods of very weak rheological behaviour
alternating with times of greater crustal strengths, during which
72
B.S. Kamber / Precambrian Research 258 (2015) 48–82
Fig. 15. Simplified geological map of the area surrounding the Town of Elliot Lake, Ontario, illustrating the difference in map pattern of Archaean vs. Palaeoproterozoic
rocks. The latter are shown as four Groups, starting from the oldest, Elliot Lake to the youngest, Cobalt Group. Simplified from Map 2018, Sault Ste. Marie – Elliot Lake Sheet,
Geological Compilation Series, 1:253,440; Onartio Department of Mines, 1967.
the geology assumed a more familiar picture. Furthermore, the
various cratons completed heat redistribution at different times in
the Neoarchaean. Thus, it would now seem unwise attempting to
define the end of the Archaean Eon with the mechanical behaviour
of the continental crust. Furthermore, the exponentially decaying global radioactive heat production function (Fig. 1D) offers
no threshold that could explain the A–P boundary as a turning
point.
Van Kranendonk et al. (2012) proposed, somewhat unconventionally, that the beginning of the Proterozoic was at 2.42 Ga,
marked by the deposition of sediments associated with the first
‘global’ glaciation. This new concept breaks with the view that
the geological expression of changing crustal rheology (e.g. Logan,
1857) can be used to define the A–P boundary. However, I prefer
Bleeker’s (2004) proposal of a transitional period (Fig. 1A) between
ca. 2600 Ma (the end of the Archaean) and ca. 2300 Ma (the base
of the Proterozoic). The transitional period was likely initiated by
a global (mantle?) event, marking the end of the Archaean Eon.
The base of the Proterozoic marks the beginning of a new style of
terrestrial geology. In between, 200–300 Ma of time had elapsed
B.S. Kamber / Precambrian Research 258 (2015) 48–82
during which the Earth’s geology had adjusted to a new thermal
structure.
6.1. A brief summary of ideas regarding the cause of the A–P
boundary
The majority of attempts at explaining the A–P boundary
seek the driving process in the mantle. Regardless of the lack of
agreement about the temperature of ambient Archaean asthenosphere, there is wide consensus about the maximum TP of plumes.
There are three important additional observations (Campbell and
Griffiths, 2014): firstly, only very few komatiites erupted after
2.5 Ga but komatiite was a widespread, if volumetrically quite
insignificant volcanic rock type throughout the Archaean. Secondly, the TP of Archaean plumes appears to have remained
approximately constant at ca. 1700 ◦ C (note that the earlier idea
of komatiite arising from melting hydrated much cooler mantle is not compatible with the petrological data available now;
e.g. Berry et al., 2008). Finally, the TP of 2.0 Ga Palaeoproterozoic plumes was apparently substantially lower, ca. 1500–1550 ◦ C
(Fig. 1E).
Because plumes originate at thermal boundary layers, most
models that deal with the disappearance of very hot Archaean
plumes have interrogated the evolution of the core-mantle or
lower-upper mantle boundaries. Campbell and Griffiths (2014)
proposed a change in the core–mantle boundary as marking the
transition between the Archaean and Palaeoproterozoic. In their
model, Archaean plume TP was constant because it was buffered
by solidification of the inner core. The drop in plume TP from
the Archaean was gradual and caused by increasing insulation of
the boundary by accumulation of a layer of subducting oceanic
slabs in the lowermost few hundred km above the core (i.e.
called D ).
By contrast, Breuer and Spohn (1995) studied the lower-upper
mantle thermal boundary layer. They argued that throughout the
Archaean, two-layer convection was governing the mantle’s temperature structure and maintaining a large temperature gradient
across the spinel-perovskite (670 km) phase transition. The thermal boundary layer at 670 km would have been the site from which
Archaean plumes originated. They proposed that the initial flush
instability that marked the break-down of two-layer convection
constituted the A–P boundary. Their model predicts that the mass
transfer across the boundary was greatest during the initial flush
instability and that this event dramatically reduced the temperature gradient across the 670-km boundary, thereby ending the
capacity to produce plumes hot enough to generate komatiite.
In his numerical model of the dynamical evolution of the mantle through time, Davies (2008) noticed the episodic (100–150 Ma)
storage of subducted basaltic crust at a depth of 660–750 km (i.e.
the very top of the lower mantle) throughout the Archaean. During
episodes of layering, the top of the lower mantle acted as a barrier
and became a significant thermal boundary layer, with uppermost
lower mantle temperatures increasing by 300 ◦ C. Each breakdown
of this ‘basalt barrier’ was followed by influx of hot lower mantle material into the asthenosphere, inducing giant ‘plume’ melting
episodes. The resulting TP of the Archaean asthenosphere therefore
oscillated between 1400 and 1750 ◦ C on relatively short timescales
of 10 s of millions of years. Eventually, the subducted oceanic crust
became strong enough to penetrate the basalt barrier, effectively
marking the A–P boundary and a more stable thermal evolution of
the upper mantle. Although not matching the periodicity predicted
by Davies’ (2008) model, Rino et al. (2004) identified episodes of
crust formation between 1.9–2.3 and 2.5–2.8 Ga based on modern
alluvial zircon age distributions, which they attributed to mantle overturns due to catastrophic collapse of ‘basalt barriers’ at
670 km.
73
6.2. Evidence from the extent of mantle depletion
Many more ideas have attributed the episodic nature of
ultramafic volcanic episodes to re-arrangement of the mantle temperature structure and/or convection style (e.g. Kump et al., 2001;
Condie et al., 2009; Arndt, 2013) and most models agree that the
A–P transition represents a particularly effective mantle ‘event’. A
predicted consequence of increased transfer of material from the
lower mantle into the asthenosphere at the A–P transition is the
re-fertilisation of the MORB-source mantle and this should be evident in the radiogenic isotope composition of Palaeoproterozoic
juvenile rocks.
Of all the commonly applied radiogenic isotope systems (Rb–Sr,
U–Pb, Sm–Nd, and Lu–Hf) the Pb-isotopes are the most sensitive to
mantle depletion via extraction and storage of continental crust
because Pb is by far the most strongly enriched of the daughter elements (see Fig. 10, where Pb plots at a much higher value
than Nd, Sr and Hf). For this reason, Pb-isotopes are better suited
to constrain the continental crust volume versus age curve than
the other isotope systems (e.g. Nägler and Kramers, 1998). When
comparing the Pb-isotope composition of the 1.9 Ga Flin Flon
ocean floor assemblage tholeiites to modelled depleted mantle Pbisotope evolution curves, Babechuk and Kamber (2011) discovered
a significant mismatch. Specifically, at any given 206 Pb/204 Pb, the
1.9 Ga tholeiites plot at far too high a 207 Pb/204 Pb ratio (Fig. 16A).
This observation is true for all measured Palaeoproterozoic juvenile samples. For example, Taylor et al. (1992) reported Sr-, Ndand Pb-isotope data for 2.1 Ga juvenile granitoids from the Birimian terrane in Ghana (West African shield). These yielded very
unradiogenic (juvenile) initial Sr-isotope ratios, yet in terms of
Pb-isotopes plot in almost the same position as the Flin Flon
ocean floor assemblage (Fig. 16B). Finally, ca. 2.0 Ga ore samples
from the Cape Smith belt (part of the juvenile basalt-dominated
Trans Hudson belt) and very unradiogenic ores from the similarly aged Ashanti belt (West Africa) plot into the same position
again (Fig. 16C). The magnitude of the offset between the modelled
depleted mantle composition and the observed data from juvenile
Palaeoproterozoic data is perhaps best appreciated by comparison with Neoarchaean rocks. For example, the 207 Pb/204 Pb ratio
of the 1.9 Ga Moen Bay tholeiites (Fig. 16A) is higher by a full
0.25 than modelled mantle. By contrast, a 2.72 Ga Pb–Pb regression lines through the data for Neoarchaean high-Mg basalts from
Theo’s Flow from the Abitibi greenstone belt (data from Debaille
et al., 2013) almost intersects the mantle evolution line at the
expected location (Fig. 17), only 0.04 higher in the 207 Pb/204 Pb
ratio.
There are two plausible explanations for this finding. Either, all
juvenile Palaeoproterozoic samples were contaminated with preexisting continental material (the yellow mixing lines in Fig. 16) or
the depleted mantle evolution line (here, Kramers and Tolstikhin,
1997) is not accurate for the Palaeoproterozoic. At least in the
case of the 1.9 Ga Moen Bay tholeiites from Elbow Lake, the possibility of contamination can be excluded with some certainty.
Amongst other characteristics, the strong light REE depletion, akin
to N-MORB (see Fig. 11B), is not permissive of contamination with
continental crust (see also Babechuk and Kamber, 2011). Thus,
the existing data suggest that the depleted mantle evolution line
requires revision towards higher 207 Pb/204 Pb ratios. In principle,
an upswing in the trajectory could be achieved by substantial continental recycling (i.e. net loss of volume of continental crust) or
refertilisation of the depleted mantle with more pristine (lower)
mantle. The latter scenario is essentially the idea of a secular expansion of the volume of the depleted mantle (note that the model of
Kramers and Tolstikhin, 1997 assumed a constant volume of the
depleted (upper) mantle that contained ca. 50% of the lithophile
elements of bulk silicate earth).
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B.S. Kamber / Precambrian Research 258 (2015) 48–82
In view of the multiple suggestions for a mantle-reorganisation
across the A–P boundary (see preceding section), Babechuk and
Kamber (2011) proposed that the depleted mantle Pb-isotope
evolution deviated from the Archaean trajectory at the A–P boundary due to the influx of contemporaneous, much less depleted
Pb from the lower mantle (green mixing lines in Fig. 16). The
refertilisation of the MORB-source mantle in common Pb-isotope
space is shown in Fig. 17. The limited relevant data available for
other juvenile Proterozoic rocks appear to confirm the proposal
(see also Lahtinen and Huhma, 1997) but in detail, the reconstruction of the Proterozoic trajectory will require much more
work.
6.3. Implication of A–P boundary for crustal ‘growth models’
Fig. 16. Common Pb-isotope systematics of Palaeoproterozoic rocks and ores. Panel
(A) shows 1.9 Ga Flin Flon (Elbow Lake) ocean floor assemblage samples (red
squares; Babechuk and Kamber, 2011) and forced 1.9 Ga regression line (red line).
The line does not intercept the depleted mantle evolution line (blue circles) of
Kramers and Tolstikhin (1997). Also shown are the primitive mantle evolution line
(green circles) of Kamber and Collerson (1999) and the average sediment composition (yellow circles) of Kramers and Tolstikhin (1997). All evolution lines are shown
in 100 Ma steps, from the past up to 1.9 Ga. Hypothetical mixing between 1.9 Ga
depleted mantle and primitive mantle (green line) and average sediment (yellow
line) is also shown. Panel (B) shows 2.05 Ga regression line (red line) through juvenile granitoid samples (red boxes) reported by Taylor et al. (1992) from West Africa.
Samples included for the regression lines are Upper West Ghana, Kumasi and Cape
Coast granitoids used by the original authors for individual isochron calculations.
The evolution and mixing lines are ornamented as in panel (A) but shown up to
2.1 Ga, the intrusion age of the granitoids. Panel (C) shows compositions of ores
from Ashanti (grey circles) and Cape Smith belt (red squares with 2 Ga regression
line). For data sources see Babechuk and Kamber (2011). The evolution and mixing lines are ornamented as in panel (A) but shown up to 2.0 Ga, the approximate
deposition ages of the ores.
For reasons explained in Section 6.2, Pb-isotopes are the most
sensitive tracers of refertilisation of the depleted mantle. However, similar but more subtle evidence for refertilisation across
the A–P boundary is found in all measures of mantle depletion,
notably Nd- and Hf-isotopes and Nb/Th ratios. With respect to
Nd-isotopes Bennett et al. (1993) were the first to propose that
the mass of the depleted mantle portion may have increased with
time and that transient, more highly depleted mantle portions
may have existed in the Archaean. Whereas Bennett et al. (1993)
focused their analysis on the first half of the Archaean, Abouchami
et al. (1990) discovered that juvenile 2.1 Ga ocean floor tholeiites from Birimian terranes (West Africa) have maximum initial
Nd-isotope ratios of εNd (t) +4.3, comparable but no more extreme
than juvenile basalts and komatiites from the 2.7 Ga Abitibi greenstone belt (e.g. εNd (t) of +2.5 to +5; see Fig. 6 of Ayer et al., 2002
for a compilation). Blichert-Toft et al. (1999) revisited the Birimian tholeiites and reported 2.1 Ga Sm–Nd- and Lu–Hf-isotope
regression lines with initial ratios of εNd (t) +3.0 ± 0.8 and εHf (t) of
+5.9 ± 0.4, respectively. In terms of Hf-isotopes, Corfu and Noble
(1992) had found equally radiogenic Hf (εHf (t) of +4.2 to +6.3) in
2.73–2.70 Ga juvenile rocks from the southern Abitibi greenstone
belt (see Fig. 18A). Ketchum et al. (2008) have since confirmed
that uncontaminated Abitibi greenstone belt magmas have εHf (t)
of +5 to +7. Finally, Babechuk and Kamber (2011) reported εNd (t)
+5.4 to +6.1 for the most light REE depleted Moen Bay tholeiites from the Flin Flon ocean floor assemblage. These authors also
drew attention to the fact that the Nb/Th ratio of the 1.9 Ga tholeiite source was 13.0 ± 0.7, only mildly higher than the average
value (11.11 ± 0.79) for juvenile Neoarchaean basalts (Collerson
and Kamber, 1999).
In summary, it appears that the depleted mantle had not become
much more depleted, or was possibly even somewhat re-enriched,
over the 600–800 Ma time span represented by the best-preserved
juvenile Neoarchaean and Palaeoproterozoic samples. This has
important ramifications for estimating the volume of past continental crust by quantifying the extent of depletion of ‘the mantle’.
Over the last decade, it has become feasible to use combined
U–Pb and Hf-isotope zircon data to revisit the question of the
continental crust preservation history (e.g. Condie et al., 2009;
Hawkesworth et al., 2010; Iizuka et al., 2010; Belousova et al.,
2010; Dhuime et al., 2012). A common feature of all these datasets
is the episodicity of zircon preservation and the observation that
between 2.7 and 2.0 Ga, there is a marked apparent decrease in
the proportion of juvenile crust being added to the global crustal
inventory (most clearly explained in Iizuka et al., 2010). This second phenomenon has been interpreted to imply a ‘shutdown’ of
juvenile magmatism after the A–P boundary (Condie et al., 2009),
and/or as an indication that the continental volume had almost
reached 70–100% of its present level by the latest Neoarchaean
(e.g. Belousova et al., 2010; Hawkesworth et al., 2010). In other
B.S. Kamber / Precambrian Research 258 (2015) 48–82
2.7 Ga Abitibi greenstone belt
2.44 Ga Kivakka intrusion
1.33 Voisey's Bay intrusion
1.07 Ga Coldwell carbonatite
15.0
2.5 Ga
207
Pb/
204
Pb
15.5
75
14.5
12.5
2.5 Ga
13.5
14.5
15.5
206
16.5
17.5
18.5
204
Pb/
Pb
Fig. 17. Common-Pb isotope evolution diagram comparing depleted mantle (blue circles; Kramers and Tolstikhin, 1997) with primitive mantle evolution (green circles;
Kamber and Collerson, 1999) shown at 100 Ma intervals. Juvenile Neoarchaean and Proterozoic rock initial Pb-isotope compositions shown as filled squares. Data sources:
Abitibi greenstone belt leached feldspar (compilation of Kamber, 2010); Kivakka intrusion (3 least radiogenic points from Amelin and Neymark, 1998); Voisey’s Bay intrusion
(3 least radiogenic data from Amelin et al., 2000); Coldwell carbonatite (4 least radiogenic points from Coldwell and Prairie Lake carbonatites compiled by Kwon et al., 1989).
The broken purple curve shows an alternative depleted mantle evolution that experienced ca. 35% refertilisation with primitive mantle at 2.5 Ga. The insert shows a Pb–Pb
regression line for the 2716 Ma Theo’s flow (Debaille et al., 2013) intersecting the depleted mantle curve very close to the expected contemporaneous value.
words, the post-Archaean time was dominated by continental recycling.
Here I draw attention to the fact that the interpretation of Hfisotopes in terms of ‘juvenile’ or ‘recycled’ requires comparison
with a known ‘mantle’ depletion evolution. All the cited studies
assume that the Hf-isotope evolution of the depleted mantle was
linear (see e.g. Fig. 18A) and, implicitly, that the mass of the depleted
mantle has remained constant over at least 3.8 Ga. In view of the
evidence reviewed here, it seems equally possible that the mantle event that may have caused the A–P boundary injected enough
fertile material into the asthenosphere to mask the contribution of
Palaeoproterozoic addition to the volume of preserved continental
crust. In addition, there is strong evidence from the less studied Precambrian shields of Africa (e.g. Abouchami et al., 1990; Iizuka et al.,
2010) and South America (Rino et al., 2004) for juvenile magmatism
in the Palaeoproterozoic (see Fig. 18A). Thus, future attempts at
discriminating between juvenile and recycled continental Palaeoproterozoic crust will have to explore the possibility of a dynamic,
likely expanding depleted mantle reservoir (e.g. Bennett et al.,
1993).
6.4. An episodic or long-lived transition zone barrier?
The concept of a mechanical barrier in the mantle transition
zone and/or in the uppermost lower mantle has many attractive
features that could help explain the seemingly disparate elements
of Archaean geology, some of which are uniformitarian and others very different from the familiar Phanerozoic. One of the first
models invoking a barrier was proposed by Nisbet and Walker
(1982) who envisaged a refractory cap of olivine that had formed
by floatation on a high-degree melt layer in the transition zone.
In their model, the refractory cap is a consequence of very large
degree melting itself. By contrast, models that envisage a barrier
from the magma ocean stage or by transient storage of slabs, view
the barrier as the reason for very high degree melting. Accordingly, the initial episodes of build-up of heat below the barrier
permitted very large scale melting and concomitant formation of
ultra-depleted and mechanically strong SCLM. This could have led
to the formation and preservation of the Eoarchaen high ␮ piece of
continental lithosphere (Fig. 6, area 1) and defines the H–A boundary. The final breakdown of the barrier marked the A–P boundary,
after which recycling of oceanic slabs into the lower mantle became
feasible (Davies, 2008) and plate tectonics was established as a
continuous process. The most important prediction of the existence such a barrier is that there was a strong contrast between the
temperature of the uppermost lower mantle and that of ordinary
asthenosphere (i.e. a pronounced thermal boundary layer). If the
barrier was episodic in nature (Davies, 2008), the asthenospheric
temperature changed between a hot and a cool state (Fig. 18B),
but if the barrier was permanent for the duration of the Archaean
(Breuer and Spohn, 1995), the asthenosphere was relatively cool,
punctured by upwellings of hot plumes emanating from below the
barrier (Fig. 18C). Both scenarios are very different from the hotter
and much more muted (Fig. 18D) temperature evolution proposed
by Herzberg et al. (2010) and adopted in many recent numerical
simulations (e.g. van Hunen and Moyen, 2012; Johnson et al., 2014;
Korenaga, 2013).
The episodic build-up and breakdown of the barrier advocated
by Davies (2008) solves a long-standing problem with the interpretation of the Archaean ‘plume’ signature, first comprehensively
discussed by Campbell and Griffiths (1992, 1993). These authors
pointed out that whereas many Archaean mafic–ultramafic rocks
have elevated incompatible trace element inventories reminiscent
of modern ocean island basalts, their radiogenic isotope signatures
are very depleted. This is unlike modern plume-derived magmas,
which also have elevated incompatible element concentrations, but
radiogenic isotope systematics clearly more enriched than MORB
(hence the terms ‘enriched mantle’ EM 1, EM 2, etc.). For the
komatiites, an incompatible element enriched source is required
in the deep mantle but one that could not have persisted for very
long because otherwise it would have started to evolve towards
more radiogenic Pb and less radiogenic Hf and Nd. The episodic
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B.S. Kamber / Precambrian Research 258 (2015) 48–82
16
A
DM
Hf
(t)
12
8
4
T (°C)
0
1700
B
1600
1500
1400
1300
2300
C
T (°C)
2100
1900
1700
1500
1300
T (°C)
D
1600
1500
1400
1300
0.0
0.5
1.0
1.5
2.0
2.5
3.0
3.5
4.0
Age (Ga)
Fig. 18. Panel (A): linear Hf-isotope evolution line (dark red) for depleted mantle adopted by Hawkesworth et al. (2010) superimposed on the initial Hf-isotope values from
zircon data compiled by Griffin et al. (2014) between 3.0 and 1.8 Ga in age from Africa (pale yellow: n = 280) and the Americas (pale grey = 502). Data with 176 Yb/177 Hf >0.05
not plotted. Also shown as solid green circles are pooled initial Hf-isotope ratios of Abitibi greentone belt zircon (>2700 Ma) from Corfu and Noble (1992). Finally initial ratio
of 2.1 Ga whole rock isochron for Birimian greenstones (Blichert-Toft et al., 1999) shown as solid blue square. Panel (B): upper mantle potential temperature evolution (case
3) modelled by Davies (2008), rapidly oscillating betwwen a hot and a cool state during the Archaean. Panel (C): schematic evolutions of upper (dark red) and lower (green)
mantle potential temperatures modelled by Breuer and Spohn (1995), showing collapse of mechanical boundary at A–P boundary. Panel (D) shows the proposed mantle
potential temperature evolution for a Urey ratio of 0.23 after Korenaga (2008) to fit empirical data presented by Herzberg et al. (2010).
but relatively short-lived pile-up of metamorphosed basalt in the
transition zone barrier can explain this contradiction, as partial
melts from this source would be enriched in incompatible elements
but isotopically not yet evolved. The close association of komatiite and tholeiite in Archaean greenstone belt successions (e.g.
Fig. 12) is also compatible with the rapid swings in thermal state of
the asthenosphere (Fig. 18B), variably generating high-degree and
more modest melts in relatively short succession. The operation
of moving plates and subduction of oceanic crust is implicit in the
model and can account for the delivery of hydrophile elements into
the continental melt source. An outstanding issue with the model is
the preservation of isotopic anomalies in daughters of short-lived
isotopes (see Fig. 7).
The alternative, long-lived barrier may also be feasible from
a theoretical modelling perspective (Breuer and Spohn, 1995). At
first glance, the idea of a long-lived barrier, yielding melts with
a depleted radiogenic isotope character but with ‘enriched’ trace
element signatures, seems improbable. However, the notion that
enriched melts cannot be derived in a single stage from ordinary
mantle (e.g. McKenzie and O’Nions, 1995) does not necessarily apply to mantle material at upper-lower mantle transition
depth, where incompatible trace elements partition very differently from shallower mantle. Calcium Si-perovskite in particular,
has the potential of assuming a critical role for producing lithophile
element enriched melts (e.g. Collerson et al., 2010; Kaminsky,
2012). Indeed, transition zone minerals have the required partition
coefficients to permit the evolution of Nd- and Hf-isotopes away
from chondritic or depleted mantle trajectories (Rizo et al., 2011).
It appears that the Nd- and Hf-isotope divergence is best expressed
in Eoarchaean rocks. This could indicate that the integrity of the
B.S. Kamber / Precambrian Research 258 (2015) 48–82
original transition zone barrier was compromised between 3.6 and
3.75 Ga and that this is additionally reflected in the coeval reduction
in extent of 142 Nd anomalies (Fig. 7).
The presence of majoritic garnet in the transition zone is also
attractive to explain the occurrence of Al-depleted vs. undepleted
komatiite. A nearly pure majorite layer could have formed as part
of a magma ocean cumulate. By contrast, the episodic basalt barrier would have produced areas of majorite-bearing eclogite less
extreme in chemical purity and therefore offers less room for making both Al-depleted vs. undepleted komatiite. A final observation
in support of the long-lived barrier is that, in the geochemical
database for Archaean rocks, there is no evidence that during the
cooler mantle stages predicted by the episodic model, magmas
started to assume the character of Proterozoic chemistry. Rather,
Keller and Schoene (2012) identified an abrupt disruption of the
secular lithospheric chemistry evolution at 2.5 Ga, but not before.
In summary, an improved understanding of the nature of the
A–P boundary requires us to revisit the debate regarding mass
transfer between the upper and lower mantle that arose from
observations of noble gas isotope compositions in modern ocean
basalts (e.g. O’Nions and Tolstikhin, 1996; Allegre and Moreira,
2004). The noble gas discussion centred around the question of
how transparent the upper-lower mantle boundary is today. By
contrast, the Archaean perspective explores the upper–lower mantle boundary area as a major thermal boundary layer and therefore
a plausible origin of ubiquitous Archaean mantle upwellings very
different from the near-stationary cylindrical plumes of today.
7. Conclusions and outlook
Much of the discussion about Archaean continental crust has
evolved around the theme of the subduction process. All to often,
this debate has assumed that once subduction was under way, plate
tectonics proceeded as a unidirectional process much in the same
way as we can observe it today. More recent numerical models
are suggesting that subduction initiation on a planet with higher
radioactive heat production was easier to overcome than had been
imagined but also that subduction may have been intermittently
active. It is now well established that many Archaean mafic and
felsic igneous rocks have trace element characteristics reminiscent
of modern rocks that form in arc settings. The association of a particular rock name, such a boninite or adakite, with a modern tectonic
setting has led the debate about Archaean plate tectonics into a
direction of limited progress. New insight will not necessarily come
from more chemical and isotopic analyses of an even greater number of Archaean rocks. Instead, the results of numerical modelling
of heat loss from the mantle have identified the need to quantify
more basic key input parameters.
First and foremost, all attempts at modelling the thermal evolution of the Earth require better constraints on the temperature
of the asthenosphere throughout the Precambrian. Although a few
proposals have been made for the thermal evolution of the ‘upper
mantle’ as a whole (e.g. Herzberg et al., 2010), these are currently
based on an approach with very high uncertainty. This regards both
the way in which TP is estimated in the absence of preserved olivine
phenocrysts and the selection of suitable tholeiite that could represent a melt that formed at an asthenospheric spreading ridge.
There is great potential to inform numerical models with a careful
reconstruction of the asthenospheric TP , from the Phanerozoic into
the Proterozoic and eventually into the Archaean.
Many Archaean cratons harbour excellent geological evidence
that the elevation of the continental lithosphere has remained
close to sea level (e.g. Galer and Mezger, 1998). In addition to
the continents, thick oceanic plateaux would also have been elevated considerably higher than modern abyssal plains leaving the
77
obvious question about the isostacy of the inferred adjacent deep
ocean basins, itself a function of the thickness of the oceanic lithosphere. Because this, in turn, is a function of asthenospheric TP ,
future numerical models of a cooling planet, whether in favour of
stagnant lid of moving plate tectonics, need to incorporate isostacy
and test the model outcomes against the strong sedimentological
evidence for continental elevation.
Because of the excellent preservation potential of zircon, much
future insight will undoubtedly come from further in situ analyses
of this mineral, including particularly combined U/Pb age determination and initial Hf-isotope reconstruction. Combined datasets
have the potential to aid a more accurate reconstruction of continental volume through time. However, regardless of the amount
of data collected, zircon Hf-isotope data need to be compared to
a reference framework of terrestrial reservoir evolution lines. Of
these, the depleted mantle evolution line is currently the least
constrained. Preferred depletion lines are based on a number of
mutually exclusive models. At one extreme is depletion in response
to a single, severe primordial event that resulted in a linear isotopic
evolution of the depleted reservoir (e.g. Hawkesworth et al., 2010).
These types of models regard continental crust extraction as a secondary influence on the depletion history. At the other extreme is
a highly dynamic evolution of the depleted reservoir, representing crust extraction and preservation (leading to greater depletion)
and expansion of the volume of the depleted mantle (leading to
re-enrichment), as advocated here. Finally, intermediate models
include non-linear curves that reflect episodic extraction and storage of evolved continental crust from a fixed volume of depleted
mantle (e.g. Belousova et al., 2010). Unless these discrepancies are
addressed with a unifying approach to mantle depletion history,
the conclusions drawn from new Hf-isotope datasets will remain a
reflection of a subjective preference for a certain depletion model.
Finally, as explained in detail by Arndt (2013) and originally
pointed out by Campbell and Taylor (1983), the role of water in
the formation of Archaean continents is ubiquitous as is the evidence for release of hydrous fluids into the melt sources of Archaean
magmas. Notwithstanding the proven ability of subducted oceanic
slabs to transport water to sufficient depth, alternative mechanisms for the vertical transport of hydrated lithologies need to be
considered (e.g. Bedard et al., 2013; Johnson et al., 2014) including estimates of how effectively ancient subaqueous mafic igneous
provinces became hydrated (Arndt, 2013).
The greatest remaining uncertainties regarding the evolution
of the Archaean continental crust are associated with the poor
understanding of the evolution of oceanic lithosphere. The major
obstacle to further progress is the extreme paucity of preserved
and uncontaminated Precambrian oceanic crustal remnants. Their
identification will require a combination of careful geological,
petrological and geochemical work. Once identified, the few rock
suites with a demonstrable depleted mantle origin would benefit from central curation and from a policy of open access of this
collection to the wider petrological and geochemical community.
Acknowledgments
This paper arose from my keynote talk at the 2013 V.M.
Goldschmidt conference in the session “Geodynamics and Crust
Formation in the Archean–Palaeoproterozoic”. I am indebted to Elis
Hoffmann for the invitation and Tim Horscroft for prompting me
to write the paper, the contents of which were influenced by many
mentors and colleagues but the shortcomings of which remain
entirely my own. Jan Kramers introduced me to the Archaean
and taught me to approach it with an open mind. Stephen Moorbath, Martin Whitehouse and Ross Taylor sparked my interest in
the Hadean and reminded me not to forget first principles. Ken
78
B.S. Kamber / Precambrian Research 258 (2015) 48–82
Collerson got me interested in trace element geochemistry and the
structure of the mantle. Many ideas developed in this paper arose
when I was working at Laurentian University, learning from the
insight of my colleagues most notably Mike Lesher. Phil Thurston
and John Ayer introduced me to the Abitibi greenstone belt and continue to probe my ideas about the Precambrian. Over the years, my
work on the Precambrian has benefitted from analytical collaborations, none more so than the one with Martin Whitehouse. Mike
Babechuk and Harold Gibson spent many memorable hours with
me in Flin Flon, where the concept of the mantle transition zone failure first occurred to me. My wife Claire kindly drew the diagrams
for Figs. 1, 6 and 12–14. Elaine Cullen produced earlier versions
of Figs. 5, 9 and 15. Stephen Moorbath and David Chew kindly
provided feedback on a draft of the manuscript. Elis Hoffmann,
an anonymous reviewer and editor Randy Parish provided many
helpful suggestions for improvement of the original manuscript.
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