THE ONSET OF PLATE TECTONICS IN THE EOARCHEAN

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DOI: 10.1130/abs/2018AM-315821
Conference: GSA Annual Meeting in Indianapolis, Indiana, USA - 2018
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Journal of Earth Science, Vol. 29, No. 6, p. 1291–1303, December 2018 ISSN 1674-487X
Printed in China
https://doi.org/10.1007/s12583-018-0999-6
Kusky, T. M., Windley, B. F., Polat, A., 2018. Geological Evidence for the Operation of Plate Tectonics throughout the Archean:
Records from Archean Paleo-Plate Boundaries. Journal of Earth Science, 29(6): 1291–1303. https://doi.org/10.1007/s12583-018-0999-
6. http://en.earth-science.net
Invited Review
Geological Evidence for the Operation of Plate Tectonics
throughout the Archean: Records from
Archean Paleo-Plate Boundaries
Timothy M. Kusky *1, 2, Brian F. Windley2, 3, Ali Polat4, 2
1. State Key Laboratory for Geological Processes and Mineral Resources, China University of Geosciences, Wuhan 430074, China
2. Center for Global Tectonics, School of Earth Sciences, China University of Geosciences, Wuhan 430074, China
3. Department of Geology, University of Leicester, Leicester LE1 7RH, U.K.
4. Department of Earth and Environmental Sciences, University of Windsor, Ontario, Canada
Timothy M. Kusky: https://orcid.org/0000-0002-4553-620X
ABSTRACT: Plate tectonics describes the horizontal motion of rigid lithospheric plates away from mid-
oceanic ridges and parallel to transforms, towards deep-sea trenches, where the oceanic lithosphere is
subducted into the mantle. This process is the surface expression of modern-day heat loss from Earth.
One of the biggest questions in Geosciences today is “when did plate tectonics begin on Earth” with a
wide range of theories based on an equally diverse set of constraints from geology, geochemistry, numeri-
cal modeling, or pure speculation. In this contribution, we turn the coin over and ask “when was the last
appearance in the geological record for which there is proof that plate tectonics did not operate on the
planet as it does today”. We apply the laws of uniformitarianism to the rock record to ask how far back
in time is the geologic record consistent with presently-operating kinematics of plate motion, before
which some other mechanisms of planetary heat loss may have been in operation. Some have suggested
that evidence shows that there was no plate tectonics before 800 Ma ago, others sometime before 1.8–2.7
Ga, or before 2.7 Ga. Still others recognize evidence for plate tectonics as early as 3.0 Ga, 3.3–3.5 Ga, the
age of the oldest rocks, or in the Hadean before 4.3 Ga. A key undiscussed question is: why is there such a
diversity of opinion about the age at which plate tectonics can be shown to not have operated, and what
criteria are the different research groups using to define plate tectonics, and to recognize evidence of
plate tectonics in very old rocks? Here, we present and evaluate data from the rock record, constrained
by relevant geochemical-isotopic data, and conclude that the evidence shows indubitably that plate tec-
tonics has been operating at least since the formation of the oldest rocks, albeit with some differences in
processes, compositions, and products in earlier times of higher heat generation and mantle temperature,
weaker oceanic lithosphere, hotter subduction zones caused by more slab-melt generation, and under dif-
ferent biological and atmospheric conditions.
KEY WORDS: Archean, tectonics, ophiolite, OPS (oceanic plate stratigraphy), orogeny.
1 GEOLOGIC SIGNATURES OF PLATE TECTONICS
Plate tectonics is recognized through documentation of plate
boundary processes including sea-floor spreading, transform
faulting, and sinking of oceanic crust and lithosphere at deep sea
trenches where oceanic slabs plunge beneath linear chains of
magmatic arcs (Wilson, 1965). These discoveries demonstrate
creation of new lithosphere at mid-ocean ridges, its lateral trans-
port along strike-slip transform faults, and recycling of this juve-
nile lithosphere to the mantle at trenches marking subduction
zones (Fig. 1a). These processes leave distinctive rock records,
*Corresponding author: tkusky@gmail.com
© The Author 2018. This article is published with open access
at Springerlink.com
Manuscript received October 24, 2018.
Manuscript accepted October 30, 2018.
documenting the operation of plate tectonics in old terrains
(Dewey and Bird, 1970). In the modern tectonic regime, small
fragments of oceanic lithosphere are occasionally scraped off
subducting oceanic plates, along with their overlying sediments
(oceanic plate stratigraphy or OPS; Kusky et al., 2013b), to be
accreted above the trenches at convergent margins, forming rem-
nants of oceanic lithosphere known as ophiolites (Coleman,
2012). Some continental material is eroded and deposited in the
trenches, or scraped off above the Benioff zones, and brought
back to the mantle (von Huene and Scholl, 1993). When these
materials reach depths of 100–120 km volatiles including water
are released, partially melting the overlying mantle wedge, creat-
ing buoyant magmas that rise to create island- or continental-
margin arc magmas, with a distinctive geochemical signature
reflecting this specific history. In the modern plate mosaic, belts
of strongly-deformed OPS with low temperature/high pressure
metamorphism form an accretionary prism between the trench
Timothy M. Kusky, Brian F. Windley and Ali Polat
1292
Figure 1. Active plate tectonic system and ancient craton. (a) Active plate system. Note how the oceanic lithosphere moves away from the oceanic spreading centers,
parallel to transforms, then descends back to the mantle in subduction zones. Water and other volatiles released from the subducting slab at 110 km aid the partial melt-
ing of the mantle wedge, creating arc magmas. (b) Ancient craton. When the ocean basin between two continents is consumed and the continents collide, a diagnostic
suite of orogenic structures is produced. Note the orogenic core with high-grade metamorphic rocks, accreted terranes and ophiolites, grading outward to foreland fold-
thrust belts, foreland basins, then remnant continental platform deposits. Note also on the right side of diagram, the older continent has been rifted, removing an un-
known portion of the craton, and leaving a triple-junction-rift-passive margin sequence. Bottom of figures is approximately the Moho.
and the magmatic arc that has high temperature/low pressure
metamorphism (Brown and Johnson, 2018; Brown, 2006). This
association forms exclusively at island and Andean arcs, and is
recognized as one of the hallmark signatures of plate tectonics
(Fig. 1a).
Lateral motion of plates brings island arcs and continents
into collision, where arc terranes may be added to active conti-
nental margins, or continents collide to form orogenic belts
characterized by internal high-grade metamorphism, succeeded
outwards by accreted terranes, belts of far-travelled nappes,
fold-thrust belts, then low-grade foreland basins (Fig. 1b). This
distinctive tectonic zonation in orogenic belts forms one of the
most-convincing hallmarks of the operation of plate tectonics
in the geologic record (Kusky et al., 2016; Șengör et al., 2014;
Fritz et al., 2013; Hildebrand, 2013; Windley, 1993; Dewey,
1977; Dewey and Bird, 1970; Wilson, 1968, 1965; Collett,
1927).
2 GEOLOGICAL EVIDENCE FOR THE OPERATION
OF PLATE TECTONICS IN THE ARCHEAN
The best way to recognize plate tectonics in very old rocks
is to systematically document from Archean terranes the same
associations of structures, sedimentary and igneous rocks, and
metamorphic patterns that are produced by plate tectonics at
plate boundaries in young Phanerozoic orogens (Fig. 1b). In
this section, we document such plate boundary zone associa-
tions from Archean rocks, producing clear evidence for the
operation of plate tectonics on Earth throughout the Archean.
Archean extensional plate boundaries are preserved in rare
cases as failed rifts in continents, such as the circa 3.1 Ga Pon-
gola structure in South Africa (Gold, 2006; Burke et al., 1985).
Successful extensional plate boundaries form oceanic spreading
centers, where new oceanic lithosphere is created by mantle up-
welling, partial melting, and cooling beneath the worldʼs oceans.
Since no ocean basins older than 200 Ma old are preserved
Geological Evidence for the Operation of Plate Tectonics throughout the Archean
1293
today on Earth, we must seek records of what ancient ocean
crust looked like from remnants of oceans that closed, where
slices of these oceans were offscraped at paleo-convergent
plate boundaries, becoming emplaced as ophiolites on top of
continental margins, or within accretionary prisms (Figs. 2a,
2b). Early work (Coleman, 2012; Casey et al., 1981; Anony-
mous, 1972) suggested that most oceanic lithosphere had a
similar structure, grading down from deep-sea sediments to
pillow lavas, then a sheeted dike complex, into high-level iso-
tropic gabbros into layered gabbros, then into an ultramafic
cumulate section, and finally into depleted mantle rocks includ-
ing harzburgite, and more rarely lherzolite (e.g., Coleman,
2012). Recent studies have revealed much greater diversity in
ophiolites and modern ocean crust on the sea floor (Fig. 2c),
leading to new definitions of how to recognize ancient oceanic
crust (Coleman, 2012; Dilek and Furnes, 2011; Kusky et al.,
2011) from magma-poor types, to magma-rich types, and those
associated with subduction zones (forearc and backarc), and
those with ocean spreading centers (Furnes et al., 2015). These
can be simplified to having a depleted mantle section from
which melts were removed to form the overlying crustal sec-
tion, and should include some plutonic rocks such as gabbros,
and volcanics such as pillow basalts. The lavas have the dis-
tinct chemical signature of MORBʼs (mid-ocean ridge basalts),
a hallmark of oceanic crust with variations dependent on spe-
cific tectonic setting (Fig. 2).
Ocean plate stratigraphy (OPS) is the “sequence of sedi-
mentary and volcanic rocks deposited on oceanic crust substra-
tum from the time it forms at a spreading center, to the time it
is incorporated into an accretionary prism at a convergent mar-
gin (Kusky et al., 2013b; Kusky and Bradley, 1999; Wakita,
1997; Bradley and Kusky, 1992)”. Figure 2a shows the tempo-
ral development of typical OPS on the oceanic substratum, as it
is transported across the ocean basin in Fig. 2b, to be incopo-
rated into the accretionary wedge where it is imbricated and
strongly deformed during accretion. Typical OPS grades up-
wards from the pillow lava section of the oceanic crust, to
deep-sea chert or limestone, banded iron formation, to shales
and mudstones, into distal then proximal turbidites (sandstone
and shale), then into conglomerates or olistostromes (Fig. 2a).
The uppermost turbidites to conglomerates are derived from the
erosion of nearby or far-distant continental blocks and trans-
ported by axial currents along the trench. When the full se-
quence enters a trench, it is strongly deformed and repeated
Figure 2. Cross section of an ocean basin, showing spreading center, and accretionary orogen above a subduction zone. Panel a shows the development of
ocean plate stratigraphy, and its structural disruption as it enters the trench. Panel b shows the oceanic spreading center where oceanic lithosphere is produced,
and moves laterally away towards trenches in a ridge-centered reference frame. On the left, the young oceanic lithosphere is shown intruded by off-axis OIB or
plume type magmas. On the right, the oceanic plate moves slowly towards the trench, slowing accumulating its OPS, which becomes imbricated, offscraped,
and underplated by thrust faults and strongly deformed as it enters the trench. Panel c below the main figure shows different types of oceanic lithosphere that
develop depending on the balance between spreading rate and magma supply.
Timothy M. Kusky, Brian F. Windley and Ali Polat
1294
by numerous thrust faults, or deformed so strongly that it forms
a tectonic mélange. OPS has now been recognized in orogenic
belts of all ages (Fig. 3), ranging from sequences being off-
scraped at convergent plate boundaries today, throughout the
Phanerozoic, Proterozoic, and through the Archean all the way
back to the worldʼs oldest preserved rocks in the circa 4.0 Ga
Nulliak greenstone belt in the Saglek Block of Labrador (Ko-
miya et al., 2015). In the Nulliak Belt (Fig. 3), MORB-basalts
are overlain by a sequence of meta-carbonates, chert, pelites,
and clastic rocks (Komiya et al., 2015). This is positive evi-
dence for the lateral motion of oceanic plates away from ridges,
accumulating the oceanic sedimentary sequence, and being
offscraped and added to the overriding arc or continental plate
at paleo-convergent plate boundaries, and of the progressive
accumulation of pelagic oceanic sediments, and of the final
deposition of clastic, continental- or arc-derived clastic sedi-
ments. The whole package is partly offscraped by thrusting in
the trench and added to an overriding arc or continental plate at
a paleo-convergent plate boundary. Well-documented examples
have now been recorded in accretionary orogenic belts of all
ages, and thus OPS provides the first-order evidence of the
operation of plate tectonics throughout Earth history.
Many Archean greenstone belts have long been known to
contain fragments of arcs and ophiolites, but this has been de-
bated (Kusky, 2004; de Wit and Ashwal, 1997), based inter-
alia on the false preconception that ophiolites must follow the
1973 Penrose definition. Nevertheless, understanding of these
rocks has improved considerably in the last two decades so that
now many different types of ophiolites re recognized ranging
from magma-poor to magma-rich types (Fig. 2c), and from
those generated in mid-ocean ridges to those created in supra-
subduction settings (Furnes et al., 2014; Kusky et al., 2011).
Based on new models of ophiolite structure, some of the best-
documented Archean ophiolites include portions of the 2.5 Ga
Shangyin-Zunhua Belt of North China (Kusky and Li, 2010;
Kusky et al., 2001), the 3.3 Ga Barberton Belt (de Wit et al.,
2018), and the 3.8 Ga Isua Belt (Fig. 4).
The 2.5 Ga Shangyin ophiolite is one of the most complete
Archean ophiolites (Kusky and Zhai, 2012), containing a well-
exposed Moho between mantle harzburgites, a 1–2 km thick
mantle transition zone with interlayered harzburgite, mafic and
ultramafic cumulates, gabbros, and overlying gabbros and basalts,
with local dike complexes (Kusky and Li, 2010). Northern parts
of the sequence are extensively intruded by gabbroic, trondh-
jemitic, and tonalitic sills, and cut by numerous diabase dike
swarms. This igneous stratigraphy (Fig. 4) takes into account the
fact that the ophiolite and surrounding area has been affected by
Mesozoic intrusions (Kusky and Zhai, 2012; Kusky and Li,
2010). The upper section contains silicic fine-grained sedimen-
tary and umber deposits and BIFʼs, and the whole sequence is
dismembered, metamorphosed to amphibolite facies, and in-
truded by several generations of younger magmas (Fig. 4).
In Barberton, seven major thrust sheets each with distinctive
tectonic histories have been delineated (de Wit et al., 2018; de
Wit, 2004), each resembling parts of different types of ophiolites
formed in modern oceanic backarc-like settings. New field, geo-
chemical and geochronological data (Grosch and Slama, 2017)
indicate that the primitive Kromberg massive and pillowed mafic
lavas, gabbros, and ultramafic cumulate sequence and overlying
cherts are a fragment of an Archean ophiolite, and were not
Figure 3. Examples of OPS from modern environments, and back through time to 4.0 billion years ago, showing remarkable similarity, demonstrating the
operation of the lateral movement of oceanic plates for at least 4 billion years. Data for sections from the following sources: Kamchatka (Kersting, 1995); Ma-
riana (Plank et al., 2000); Japan (Wakita, 2012); Nankai (Shipoard Scientific Party, 2000); Ballantrae (Sawaki et al., 2010); Lleyn (Maruyama et al., 2010);
Cleaverville (Kato et al., 1998); 3.5 Ga Pilbara (Kato and Nakamura, 2003); Saglek (Komiya et al., 2017, 2015).
Geological Evidence for the Operation of Plate Tectonics throughout the Archean
1295
Timothy M. Kusky, Brian F. Windley and Ali Polat
1296
erupted through or deposited on older continental crust. Detrital
zircons from the 3.3–3.2 Ga Fig Tree and Moodies groups reveal
no evidence for older continental crust in the Barberton source
area during deposition (Drabon et al., 2017).
In the circa 3.8 Ga Isua Belt of Greenland (Fig. 4), pillow
lavas, possible sheeted dikes, gabbroic and ultramafic rocks,
are interpreted as a small fragment of a Paleoarchean ophiolite
(Furnes et al., 2007). There are hundreds if not thousands of
ophiolitic fragments within Archean greenstone belts (Furnes et
al., 2014; Kusky, 2004), so these three examples spanning the
entire length of the Archean show that processes of sea-floor
extension and magmatism were in operation in the Archean, in
a manner very similar to that of today. Accordingly, there is a
record of extensional oceanic plate boundaries and plate tecton-
ics throughout the duration of Earth history (Furnes et al., 2015;
Kusky et al., 2011; de Wit and Ashwal, 1997).
Large strike-slip faults form at transform plate boundaries,
above zones of oblique subduction, or in association with hori-
zontal motions reflecting lateral motions of plates. Phanerozoic
examples include the Alpine fault of New Zealand (600 km
long), the North Anatolian fault of Turkey (1 500 km long),
and the San Andreas of California (1 200 km long). Archean
cratons also contain abundant >1 000 km strike-slip faults
demonstrating horizontal motion of large rigid crustal blocks in
the Archean. Examples include the >1 100 km long 2.7 Ga
Quetico fault, which transects the entire length of the Superior
Craton (Fig. 5), and is a terrane boundary between the Quetico
meta-sedimentary and Wabigoon meta-igneous provinces
Figure 5. Evidence for plate tectonics in Archean terranes includes craton-scale strike-slip faults, and seismically defined paleo-subduction zones that show offsets of the
Moho, remnant slabs, and zones of metasomatized mantle above the Archean paleosubduction zones (Percival et al., 2012). (a) Map of the Superior Craton (Percival et al.,
2012), showing a paleosubduction zone beneath the Quetico Province (panel c), and upper-plate oblique-slip fault analogous to the Great Sumatra fault (panel f) or the
Septentrional-Oriente fault (SF and OF) between the Caribbean and North American plates (panel e). (b) Map of the Yilgarn Craton, and seismic profile (panel d from
Drummond et al., 2000) showing suture and offset Moho along the suture between the Eastern Goldfields Province, and western part of the craton, composed of the South-
ern Cross (SC), Southwest (SW), Murchison (MV) and Narrier (NR) complexes. Abbreviations in panel c as follows: WW. Wawa; QF. Quetico fault; WG. Wabigoon
subprovince; WR. Winnipeg River subprovince; ER. English River subprovince; UC. Uchi subprovince; NC. North Caribou superterrane.
Geological Evidence for the Operation of Plate Tectonics throughout the Archean
1297
(Percival et al., 2012). In the Yilgarn Craton of Australia, the
circa 1 100 km long Ida oblique slip fault forms the boundary
between the eastern and western Yilgarn terranes (Zibra et al.,
2017), and the upper plate to the east is sliced by numerous
Archean strike-slip faults, forming a system similar to the dex-
tral Sumatran fault zone (Fig. 5) that extends along the topog-
raphic axis of Sumatra and accommodates the oblique compo-
nent of convergence between the Australia/Indian plate and the
overriding Sunda plate (Fitch, 1972), or the Septentrional-
Oriente fault (Fig. 5) that accommodates the strike-slip compo-
nent of strain partitioning between the Caribbean and North
American plates (Dolan and Mann, 1998). Thus, Late Archean
cratons are marked by strike-slip faults with similar scales and
structures as younger plate-bounding transform faults, and the
geological record demonstrates lateral movement of crustal
blocks by 2.7 Ga (Percival et al., 2012). Preserved parts of
older, Eoarchean cratons, are much smaller than the Neoar-
chean cratons, but also preserve evidence of transform tecton-
ics, on the scale of the preserved cratons. In the Australian
Pilbara the 3.5–3.3 Ga Lalla Rookh and Whim Creek belts
formed in pull-apart basins along major craton-scale strike-slip
or transform faults analogous to the San Andreas and North
Anatolian transforms (Krapez and Barley, 1987).
Orogenic belts form in zones of plate convergence, such as
above subduction zones, or in wide plate boundary zones where
two continents collide. Characteristics of these orogenic belts
include early thrust faults indicating crustal shortening (Fig. 6).
Early horizontal thrust faults and nappes are known from orogens
of all ages (McClay, 2012), including the famous Cenozoic
Austro-Alpine nappes, the Paleozoic Appalachian nappes, and in
the Mesozoic Cordillera of western North America. Similarly,
early horizontal thrust and inverted nappe structures are well-
documented in the Early Archean Pilbara and in the North Atlan-
tic cratons, and in the Mid–Late Archean Zimbabwe, Kaapvaal,
Yilgarn, Slave, Superior, North China, and Brazilian cratons
(Kusky et al., 2016; Cawood et al., 2009; Kusky and Vearncome,
1997). Young orogens such as the Alps and Himalaya exhibit
well-defined tectonic zonations, grading from highly-deformed
and metamorphosed hinterlands, including accreted arcs and
other terranes, through zones of nappes, to foreland fold-thrust
belts, and eventually into relatively undeformed foreland basins.
In old orogens in Archean cratons, deeper crustal levels are typi-
cally exposed, and these orogens have been subjected to further
later tectonic overprinting events. Despite this, clear orogenic
tectonic zonations similar to those of the Alps and Appalachians
have been documented from the 2.5 Ga North China (Kusky et
al., 2016), 2.6 Ga Slave (Bleeker and Hall, 2007; Kusky, 1989),
2.7 Ga Brazilian (Hildebrand, 2005), 2.7 Ga Zimbabwe (Kusky,
1998), 3.5 Ga Pilbara (Hickman, 2012) and Yilgarn (Myers,
1995) cratons.
When oceanic plates are subducted, they return to the mantle
to be recycled, but dehydration reactions in the subducting slabs
Figure 6. Schematic composite model of an orogen, showing the classical tectonic zonation from passive margin sequences, to foreland basins, into fold-thrust
belts, obducted arcs and ophiolites, and the high metamorhic grade core of the orogen characterized by intense nappe-style folding, and zones of domal pluton
emplacement. Orogenic zonations like this are found in orogens of all ages, suggesting that plate tectonics has operated since the formation of the first orogens
at 4.0 Ga. Note that the different locations on the map and section correspond to real places in convergent and collision belts on Earth, including Barberton (BB),
Shurugwe (SK). Pilbara (PB), North China Craton (NCC), South China (Yangtze) Craton, Superior Craton (SUP), Himalaya (HM), Transvaal (TR). Imbricated
slabs of subducted ocean lithosphere for the root of the craton (based on the model of Kusky, 1993), and slab-break-off and roll back induce thinning of the
SCLM and magmatism, after the model and sources cited in Kusky et al. (2014b).
Timothy M. Kusky, Brian F. Windley and Ali Polat
1298
and overlying sediments hydrate the overlying mantle wedge on
the way down. At depths of about 110–200 km, sufficient water is
released to the overlying mantle wedge to induce partial melting,
and these melts rise to form a magmatic arc (Fig. 2), with a dis-
tinctive chemical signature as Island Arc Tholeiites (IAT). Rock
suites that form in arcs in the present day plate mosaic have ana-
logs with exactly the same lithologies, chemical signatures, rock
associations, structural relationships, and tectonic zonations in-
cluding forearcs, arcs, and backarcs throughout the Archean
(Polat, 2012). Magmatic arc petrological and geochemical sig-
natures are well-documented from the Proterozoic and Archean
of Australia and Greenland going back to at least 3.1 Ga (Szilas
et al., 2016; Korsch et al., 2011; Windley and Garde, 2009),
and to at least 3.7 Ga in SW Greenland (Nutman et al., 2015).
Thus, there is no doubt that geological processes at paleo-
convergent plate boundaries were the same in the Archean as
they are today, at least to a depth where the slab reaches 110
km and the release of volatiles generates arc magmatism. Al-
though it is possible to produce “arc-like” geochemical signa-
tures using other pressure-temperature-fluid melting conditions
(von Huene and Scholl, 1993), the combination of the struc-
tural geology, sedimentology, and volcanology of the fore-arc,
arc, and back-arc regions, rock types, and geochemistry
strongly argues for a subduction-related origin for arc-like
magmas in Archean terranes. Below the arc-magma generation
depth is what we colloquially name the “zone of speculation”
(Fig. 1) where Archean oceanic lithospheric slabs may have
been subducted to the mantle leaving no trace, except perhaps
geochemical signatures in the depleted mantle, isolated mineral
xenocrysts, or stagnant slabs in the transition zone or along the
core-mantle boundary.
In rare cases, samples of deeply subducted Archean oce-
anic lithosphere have been returned to the surface, where oce-
anic slabs have underplated the overlying continents forming
the sub-continental lithospheric mantle (SCLM) (Kusky, 1993),
and kimberlites (Fig. 6) have entrained samples of the underly-
ing eclogite and peridotite as mantle xenoliths (Shirey and
Richardson, 2011; Richardson et al., 2001). Further evidence
for subduction of oceanic lithosphere in the Archean comes
from geophysical data. One of the best examples is from com-
bined deep seismic reflection, refraction and geological data
across the Superior Province, that shows clearly three Archean
paleo-subduction zones (Fig. 4). The first two are between the
2.7 Ga Quetico and Wawa terranes and the 3.2 Ga Winnipeg
River terrane (Percival et al., 2012). These remnant slabs offset
the Moho and extend to circa 300 km depth as shown by deep
geophysical data (Musacchio et al., 2004; Sol et al., 2002).
These data all show strong S-wave anisotropy in the remnant
slabs (Musacchio et al., 2004), that is typical of oceanic litho-
sphere forming subducted remnants beneath Archean cratons
(Percival et al., 2012; Kusky, 1993). When traced to the surface,
two of these slabs coincide with greenstone belts with oceanic
(MORB) pillow lavas, bordered by the metasedimentary, accre-
tionary prism-like Quetico Domain, a clear indication of shal-
low level convergent margin tectonics to deep subduction in
the Archean (Percival et al., 2012). The third corresponds to the
boundary between the 3.0 Ga North Caribou and 3.5 Ga Hud-
son Bay terranes. Additional seismic surveys across the Abitibi
Province to the east also show dipping reflections extending 30
km into the mantle, and are interpreted (Calvert et al., 1995) as
a remnant 2.69 Ga paleo-subduction zone. Similar integration
of geophysical and geological data has revealed fossil, or pa-
leo-subduction zones in the Archean Yilgarn (Fig. 4d), North
China, and Slave cratons (Kusky et al., 2014a; Kusky, 2011;
Cook et al., 1999).
3 RECONCILING ARGUMENTS AGAINST ARCHEAN
PLATE TECTONICS
Despite the abundance of geological data that support the
operation of plate tectonics throughout Earth history, the appar-
ent absence of some features have been used to suggest that plate
tectonics did not start at all until at the time when these so-called
diagnostic signatures of plate tectonics have been first docu-
mented (Condie, 2018; Maruyama et al., 2018; Foley et al., 2014;
Kusky et al., 2013a; Dhuime et al., 2012; Næraa et al., 2012; von
Hunen and Moyen, 2012; Rollinson, 2010; Harrison, 2009; Con-
die and Kröner, 2008; Richardson and Shirey, 2008; Stern, 2008,
2007; Brown, 2007; Smithies et al., 2007; Cawood et al., 2006;
Korenaga, 2006; Moyen et al., 2006). The first is the lack of
blueschist facies metamorphism in the older record (Brown and
Johnson, 2018; Liou et al., 1990; Ernst, 1972). Blueschists record
a cold geotherm, characteristic of modern subduction zones
where the old cold subducting plates refrigerate the overlying
accretionary wedges, leading to high-pressure/low-temperature
metamorphic conditions (Ernst, 1973). Blueschists are extremely
rare in modern orogenic belts, and many Phanerozoic orogens
have none, perhaps explaining their paucity in Archean orogens.
Importantly, with 200–300 ºC higher mantle temperatures in the
Archean mantle (Korenaga, 2013; Abbott and Hoffman, 1984),
and younger average ages of subducting slabs (Abbott and
Hoffman, 1984), subduction geotherms would have been signifi-
cantly warmer in Archean subduction zones, forming green-
schists and amphibolites, instead of blueschists and eclogites.
After collision, these would have been strongly overprinted by
regional medium-pressure/temperature metamorphic conditions.
Most Phanerozoic blueschists are in the circum-Pacific and
Tethyan orogens that have yet to experience their “final” colli-
sions, and subsequent overprinting by regional medium P-T
metamorphism that can obliterate all records of previous HP
conditions.
Archean eclogites have been known for some time as inclu-
sions in young kimberlites (Fig. 6) piercing cratons (Richardson
et al., 2001), interpreted to be entrained from imbricated slabs of
buoyant oceanic lithosphere with intervening trapped wedges of
fertile mantle (Fig. 6), comprising the SCLM (Kusky, 1993). The
reason why these eclogites remain beneath the cratons, with no
known examples being exhumed during Archean continental
collisions could be related to different buoyancy under slightly
higher mantle temperatures, or perhaps they simply have not yet
been recognized) (Ganne et al., 2011). Archean eclogites have
been reported for some time from the Belomoran massif of
Scandanavia (e.g., Dokukina et al., 2014), but these have been
widely disputed mostly based on whether the age of the HP
metamorphism is Archean or Paleoproterozoic.
When most authorities discussed the role of eclogites in
subduction tectonics (e.g., Stern, 2008), they only considered
Geological Evidence for the Operation of Plate Tectonics throughout the Archean
1299
low-T eclogites that are likely to occur, often with low-T
blueschists, in the low-grade, upper crustal parts of orogenic belts,
but they did not consider the possibility that high-T eclogites
may occur in the high-T, high-grade, granulite-gneiss, deep lev-
els of Archean orogenic belts (Fig. 6). In southern India well-
authenticated eclogites and garnet websterites occur as lenses and
layers up to several meters thick within garnet-rich, chromite-
layered, anorthosite-gabbro-ultramafic layered complexes such
as the Sittampundi Complex. Garnets in gabbros contain inclu-
sions of omphacite, and calculated phase equilibria indicate that
the peak metamorphic assemblage was garnet-omphacite-rutile-
melt, which formed at 20 kbar and >1 000 ºC (Sajeev et al.,
2009). The crystallization age of the anorthosite is 2 541±13 Ma,
and the high-grade metamorphic age is 2 461±15 Ma (Mohan et
al., 2013). The data above on ophiolites and HP rocks invalidate
the speculative suggestions of Stern (2008, 2007) that plate tec-
tonics did not “start” until the Neoproterozoic, based on the pre-
sumed lack of eclogites older than that age.
Other arguments for deep subduction in the Archean come
from mineral inclusions in Archean diamonds in kimberlites.
Silicate and sulfide inclusions in kimberlitic diamonds (Shirey
and Richardson, 2011) older than 3.2 Ga have only peridotitic
compositions, but after 3.0 Ga, eclogitic inclusions became
common, interpreted to reflect the time of onset of modern
style subduction. Nitrogen and carbon geochemical fingerprints
of mantle-derived diamonds show that oxidized material has
been subducted to the mantle since at least 3.5 Ga, and proba-
bly since 3.8 Ga (Smart et al., 2016).
The argument that no ophiolites or ophiolitic mélanges are
known in terranes older than 1 Ga is incorrect. 2.5 Ga ophioli-
tic mélanges are well-documented in the North China Craton
(Wang et al., 2016, 2013), the 2.7 Ga Slave (Kusky, 1989), and
Superior provinces of Canada (Kusky and Polat, 1999), and at
deeper levels would only be represented as banded gneisses in
cryptic sutures. Relicts of many greenstone belts are now rec-
ognized as ophiolitic fragments in accretionary orogens extend-
ing back to 3.8 Ga (Furnes et al., 2014), and OPS can be re-
garded as a proxy for sea floor spreading to 4.0 Ga (Kusky et
al., 2013b), demonstrating the lateral motion of oceanic litho-
sphere away from ridges inexorably towards trenches (Fig. 2).
4 SECULAR CHANGES IN PLATE TECTONIC STYLE
THROUGH THE ARCHEAN, BUT STILL PLATE TEC-
TONICS
We present many examples of extensional, transform, and
convergent plate boundary structures and rock associations
throughout Earth history. The styles of deformation are similar,
at all scales, throughout time, and the mineralogical and geo-
chemical components are the same in similar plate boundary
settings (Keller and Schoene, 2018). It is clear from the geologi-
cal record that plate tectonics, in a form similar to that of today,
has operated on planet Earth since at least 4.0 Ga, the age of the
oldest preserved rocks, as documented by abundant geological,
geochemical, isotopic, and theoretical data (Maruyama et al.,
2018; Komiya et al., 2017; Nutman et al., 2015; Korenaga, 2013;
Polat, 2012; Shibuya et al., 2010; Richardson and Shirey, 2008).
Despite this, there are some differences between rocks produced
by plate tectonics on the early Earth, and those produced by simi-
lar processes in the modern world. The first of these relates to
secular cooling of the Earth, which has produced a steady grad-
ual change in the trace element chemistry of magmas in exten-
sional and perhaps other settings (Keller and Schoene, 2018),
differences in the thickness of oceanic crust due to higher de-
grees of partial melting (Foley et al., 2003; Sleep and Windley,
1982), and perhaps a dominance of more-shallow subducting
young slabs with more frequent slab break-off events (Foley et
al., 2003; von Huene and Scholl, 1993), on a planet with more
smaller plates (Abbott and Hoffman, 1984), than in the present
plate mosaic. The second main change between the early and
modern Earth is the change in the biosphere and consequent
chemistry of the oceans and atmosphere (Duncan and Dasgupta,
2017), with resultant changes in tectonic signatures at plate
boundaries. In the Early Archean the atmosphere was much more
reducing than at present—there were no extensive biogenic car-
bonate platforms, so carbonates formed by chemical precipitation
processes. Oceans were saturated in Si, so that hydrothermal
processes at mid-ocean ridges produced silica-rich chimneys and
deposits (cherts, BIFʼs, and the magnetite-quartzite-basalt asso-
ciation common in some greenstone belts) instead of the black
smoker mounds and associated sulfide deposits of the modern
oceans (Shibuya et al., 2010).
In summary, an analysis of the rock record shows that
there is no evidence that plate tectonics did not operate in a
manner similar to modern style tectonics on the early Earth, 4.0
billion years ago, much as it does today. The planet lost heat
then as now, by making and ageing oceanic crust, moving it
laterally away from ridges along transform faults, and returning
crustal material to the mantle at subduction zones to be recy-
cled to the deep mantle. Arcs formed above subduction zones,
forming more highly differentiated rocks that gradually grew in
volume, melted again in later collisions, building the continents
that are extant on the planet today.
ACKNOWLEDGMENTS
This work was supported by the National Natural Science
Foundation of China (Nos. 91755213, 41672212, 41572203),
the MOST Special Fund (No. MSFGPMR02-3) and the Open-
ing Fund (Nos. GPMR201607, 201701) of the State Key Labo-
ratory of Geological Processes and Mineral Resources, China
University of Geosciences (Wuhan). Miss Yating Zhong is
thanked for assistance with final manuscript preparation and
figure drafting. We dedicate this contribution to the memory of
Kevin Burke, for his lifelong contribution to Precambrian tec-
tonics. The final publication is available at Springer via
https://doi.org/10.1007/s12583-018-0999-6.
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