Available online at www.sciencedirect.com
Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89 – 129
www.elsevier.com/locate/palaeo
Glacio-epochs and the supercontinent cycle after ∼ 3.0 Ga:
Tectonic boundary conditions for glaciation
Nick Eyles
Department of Geology, University of Toronto at Scarborough, 1265 Military Trail, Scarborough ON, Canada M1C 1A4
Received 11 May 2006; received in revised form 5 September 2007; accepted 26 September 2007
Abstract
Tectonic influences on long-term climate change are of considerable current interest and debate. This paper reviews the
relationship between multi-million year periods of glaciation (glacio-epochs) over the last 3 Ga of Earth history and phases of
supercontinent breakup and assembly. A preferred but not exclusive relationship is evident between glacio-epochs and their mostly
glacially influenced marine record, with rifting. The earliest known glaciation (mid Archean ∼ 2.9 Ga) is recorded in the marine
Mozaan Group of South Africa deposited along the passive margin of the Kapvaal Craton then part of the early continent Ur. The
Paleoproterozoic glacio-epoch, exemplified by the Huronian Supergroup of Ontario, Canada (∼ 2.4 Ga) and strata in northern
Europe and the U.S., is associated with rifting of Kenorland. A long Paleo-Mesoproterozoic non-glacial interval (c. 2.3 Ga to
750 Ma?) coincides with continental collisions and high standing Himalayan-scale orogenic belts marking the suturing of
supercontinents Nena-Columbia and Rodinia. A near absence of glacial deposits other than at 1.8 Ga, may reflect lack of
preservation. The extensive and prolonged Neoproterozoic glacio-epoch records either diachronous glaciations or discrete pulses of
cooling between ∼ 750 and ∼ 580 Ma, and is overwhelmingly recorded by substantial thicknesses (1 km+) of glacially influenced
marine strata stored in rift basins. These formed on the mid to low latitude (b 30°) oceanic margins of western (Panthalassa:
Australia, China, Western North America) and eastern (Iapetus: Northwest Europe) margins of a disintegrating Rodinia. The
youngest glacially influenced deposits formed about 580 Ma along the compressional Cadomian Belt exterior to Rodinia (Gaskiers
Formation) possibly correlative with the classic passive margin Marinoan deposits of South Australia.
A short-lived (1 to 15 Ma?) Early Paleozoic ice sheet about 440 Ma grew over highlands on the polar North Africa margin of
Gondwana possibly likely triggered by uplift at high paleolatitudes as large terranes (e.g., Meguma, Avalonia) rifted away from
North Africa. Incised valleys, coarse glacial fills and thick (1 km +) ‘postglacial’ shales suggest a continuing tectonic influence.
Devonian cooling across the Frasnian-Famennian boundary (c. 376 Ma) is recorded by local ice covers in Brazil and Bolivia and
is linked to elevated topography and enhanced erosion of continental crust. The Late Paleozoic glacio-epoch (∼ 350 and 250 Ma)
coincides with a high paleolatitude positioning of Gondwana and the growth of high standing topography when Gondwana
collided with Laurasia to create Pangea. Breakup after 180 Ma moved landmasses into higher northerly latitudes and was the
backdrop to global cooling of the Cenozoic glacio-epoch that commenced after the Paleocene–Eocene Thermal Maximum
(b 55 Ma). Earliest Antarctic ice at ∼ 40 Ma most likely nucleated on the high shoulders of the Transantarctic Rift coeval with
opening of Drake Passage, and coincides with the earliest ice rafting in the Arctic Basin at 43 Ma, followed by another pulse at
34 Ma. Accelerated glacierization in both hemispheres occurred at ∼ 14 Ma during the middle Miocene Transition but
Milankovitch-forced continental-scale ice sheets did not nucleate in the northern hemisphere until after 3.5 Ma on uplifted
borderlands along North Atlantic passive margins.
E-mail address: [email protected].
0031-0182/$ - see front matter © 2007 Elsevier B.V. All rights reserved.
doi:10.1016/j.palaeo.2007.09.021
90
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
A preferred association of the deposits of Proterozoic and Phanerozoic glacio-epochs with rift basins reflects either a causal link
between rift-related uplift and regional cooling, or simply enhanced preservation of glacial sediments. Glacial deposits are poorly
preserved in areas of compressional tectonics.
© 2007 Elsevier B.V. All rights reserved.
Keywords: Earth's glacial record; Archean; Proterozoic and Phanerozoic glacio-epochs; Rifting; Tectonically created topography
Contents
1.
2.
Purpose of this paper: why does ice appear on planet Earth?. . . . . . . . . . . . . . . . . . . . . . .
Glacio-epochs defined. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.1. Structure of this paper . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3. Archean glacio-epochs (c. 4–2.5 Ga) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.1. A sparse record . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.2. Summary. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4. Paleoproterozoic glacio-epoch (c. 2.4 Ga). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.1. Rift related ice covers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.2. Summary. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5. Paleoproterozoic to Mesoproterozoic non-glacial interval (c. 2.3–0.75 Ga). . . . . . . . . . . . . . . .
5.1. Cycles of continental assembly and breakup but no glaciation? . . . . . . . . . . . . . . . . . .
5.2. Summary. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
6. Neoproterozoic glacio-epoch (0.75 Ga to 545 Ma) . . . . . . . . . . . . . . . . . . . . . . . . . . . .
6.1. Stratigraphic record of Neoproterozoic cold climates: the ‘glaciated rift’ debrite–turbidite association
6.2. Neoproterozoic glacioeustasy or tectonics? . . . . . . . . . . . . . . . . . . . . . . . . . . . .
6.3. Diachronous response to regional tectonics or globally synchronous megafreeze events? . . . . .
6.4. An active hydrosphere; climate models and sedimentology . . . . . . . . . . . . . . . . . . . .
6.5. Carbon isotope excursions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
6.6. Summary. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
7. Lower Paleozoic Saharan glacio-epoch (c. 440 Ma) . . . . . . . . . . . . . . . . . . . . . . . . . . .
7.1. Long lasting or short lived? And just how big? . . . . . . . . . . . . . . . . . . . . . . . . . .
7.2. Geomorphic, sedimentologic and stratigraphic evidence of glaciation in North Africa . . . . . .
7.3. Sea level fluctuations; are they exclusively glacioeustatic? . . . . . . . . . . . . . . . . . . . .
7.4. Tectonic influences on glaciation in North Africa; the importance of black shales . . . . . . . .
7.5. Summary. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
8. Late Devonian ice c. 374 Ma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
8.1. Brief cooling triggered by continental collision . . . . . . . . . . . . . . . . . . . . . . . . . .
9. Late Paleozoic Gondwanan glacio-epoch (c. 350–250 Ma). . . . . . . . . . . . . . . . . . . . . . . .
9.1. Long lived cooling triggered by continental collision . . . . . . . . . . . . . . . . . . . . . . .
9.2. Summary. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
10. Mesozoic non-glacial interval (c. 250 to 55 Ma); a role for small ice masses? . . . . . . . . . . . . . .
Ma)
. . . .. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
11. Cenozoic glacio-epoch (b55
(b< 55
Ma)
11.1. Coupled tectono-climate models: the effects of plate collision and dispersal . . . . . . . . . . .
11.2. Glaciation at the tops of the world: Antarctic and Arctic cooling c. 40 Ma. . . . . . . . . . . .
11.3. Circum North Atlantic glaciation: Milankovitch-driven ice sheets . . . . . . . . . . . . . . . .
11.4. Pacific Northwest glaciation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
11.5. Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
12. Discussion. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
12.1. Assembly related glacio-epochs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
12.2. Break up related glacio-epochs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
12.3. Rifting and glaciation: causal link or simply preservational bias? . . . . . . . . . . . . . . . .
13. Concluding remarks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
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N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
1. Purpose of this paper: why does ice appear on
planet Earth?
Ongoing climate change is at the forefront of current
scientific and public debate. Turning to the remote
geologic past, considerable interest is focussed on the
record of cold climates and much-debated Neoproterozoic global freeze events (‘Snowball Earth’; Hoffman
and Schrag, 2002). These putative climatic catastrophes
are regarded by some as triggers for early metazoan life
at around 580 Ma before present (abbreviated to Ma
throughout); similarly, more recent Phanerozoic cold
periods are credited with causing mass extinctions
(Stanley, 1998; Brenchley et al., 2003). Against this
backdrop of widespread interest in the causes and varied
effects of cold climates, this paper contributes to discussion by reviewing the geotectonic setting of Archean, Proterozoic and Phanerozoic glaciations (Fig. 1).
The objective is to see whether there are any broad
generalizations that can be made about the tectonic
boundary conditions required to grow extensive ice
covers.
91
Assessment of the geodynamic setting of ancient
glaciations was last attempted more than a decade ago
(Eyles, 1993) and many data have emerged since to
warrant a new look especially in the light of new
knowledge on the timing of supercontinent assembly
and breakup (e.g., Rogers and Santosh, 2004). Such a
review is timely because longstanding (Le Roex, 1941)
but formerly rejected (Schermerhorn, 1974) notions of
globally extensive glaciations have re-emerged (Hoffman et al., 1998; Williams and Schmidt, 2004) supported by a wealth of geochemical and geophysical data.
According to some, Precambrian glaciations are seen as
being predominantly ‘near equatorial’ with radically
different causes compared to those of the Phanerozoic
which are regarded as exclusively mid to high latitude
phenomena (Hoffman and Schrag, 2000, 2002, p.129;
Evans, 2003a). According to another school of thought
the broad timing of glaciations on planet Earth over the
last 1 Ga can be linked primarily to extraterrestrial
factors such as the periodic variability of the cosmic ray
flux (CRF) and changing solar activity (Shaviv, 2003;
Marcos and Marcos, 2004; Veizer, 2005; Fig. 2). In fact,
Fig. 1. Schematic representation of glacio-epochs in Earth history and their relationship to phases of supercontinent assembly and break up. The
timing of supercontinent growth and rifting is after Condie (2002a), Rogers and Santosh (2004). Earth's glacial record is after Crowell (1999). It
remains unclear whether the passive-margin related Kaapvaal glaciation represents a glacio-epoch or a short-lived event. The timing and number of
glacial events in the Neoproterozoic (3a, b, c) is uncertain (see text). The anomaly of the lack of a glacial record during the Paleo-Mesoproterozoic
growth of Nena-Columbia is clearly evident though Williams (2005) reports evidence of glaciation at 1.8 Ga. The sedimentary record of most glacioepochs occurs in the geodynamic context of intracratonic rifting, crustal extension and the formation of passive margins. Paleoproterozoic (c.2.5 Ga)
and Neoproterozoic glacio-epochs (c. 750–580 Ma) occurred during the breakup of Kenorland and Rodinia respectively. It is also possible that
extension along high latitude continental margins and consequent uplift also played a role in triggering Ordovician glaciation at c. 440 Ma (when
terranes rifted off Gondwana; see text). Most of the Gondwanan glacio-epoch deposits are stored in rift basins even though glaciation was initiated
during the compressional growth phases of Gondwana. Tectonics played a major role in Cenozoic cooling after 55 Ma culminating in continental
scale Northern hemisphere ice sheets only after 3.5 Ma.
92
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
Fig. 2. A: Glacio-epochs of the last billion years and their relationship to supercontinent cycle. Numbering of each glacio-epoch follows that of Fig. 1.
The model of three distinct global pulses of glaciation in the Neoproterozoic glacio-epoch may be simplistic (see Section 6.3). B: Variation in 13C over
last 1.5 Ga years (modified from Hoffman and Schrag, 2000, 2002; Lindsay and Brasier, 2002 with additional data). Note largest excursions occur in
the Neoproterozoic are coincident with breakup of Rodinia. There is no agreement as to their cause. Prominent spikes in 13C during the midProterozoic non-glacial interval are interpreted as the consequence of rapid burial of carbon (Bartley et al., 2001); those of the Neoproterozoic are seen
as driven by global glaciation and suppression of biologic activity (Hoffman and Schrag, 1999) or possible diagenesis in rift basins (Foden et al.,
2001; Knauth 2005) divorced from the global ocean (e.g. Gammon et al., 2005). C: Variation in Cosmic Ray Flux (curve) a timing of Earth's periodic
crossings of spiral arms of the Milky Way (after Shaviv, 2003; Marcos and Marcos, 2004). In this model, long episodes of glaciation occur at times of
high CRF. Charged particles enhance cloud covers thereby increasing the planetary albedo and promoting overall climate cooling on geologic
timescales (Veizer, 2005). D: Estimated global temperature trends (Veizer, 2005). E: Variation in atmospheric carbon dioxide (dashed line after
Rothman, 2002, solid line after Berner, 2003; Veizer, 2005 with additional data from Sheldon, 2006).
the relationship of CRF flux with terrestrial record of
glaciation (and non-glacial periods) is not precise
(Fig. 2) and additional causes such paleogeography,
atmospheric composition and tectonics clearly play a
role (Hay et al., 2002). As stressed by Shaviv (2003) and
Evans (2003a) the key question in paleoclimatology is
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
93
how to identify each factor and its relative importance
through time.
The present paper assembles tectonic, stratigraphic, sedimentologic, paleontologic and other data
appertaining to Earth's record of cold climates over
the last 3 Ga. As we shall show, plate tectonic
processes have played a role in the initiation of some
but not all. During certain times, tectonic activity in
both extensional and compressional settings created
areas of high topography (‘tectonotopography’) on
which ice could nucleate and expand once other
thresholds were crossed (Crowell, 1999; Shaviv and
Veizer, 2003; Marcos and Marcos, 2004; Veizer,
2005; Hay, 1996).
lengthy episodes of glaciation extending (but not
necessarily continuous) over millions of years are
correspondingly termed glacio-epochs. Other terms
that have been used are ‘glacio-eras’ (Chumakov,
1985), ‘glacial eras’ (Hambrey and Harland, 1985;
Chumakov, 1985), ‘icehouses’ (Fischer, 1986; Frakes
et al., 1992), ‘glacioepochs’ or ‘glacioperiods’ (Chumakov, 1978, 1985) and ‘ice age epochs’ (Shaviv,
2003). Glacio-epochs appear to be an intermittent
feature of Earth's climate system as argued by Crowell
(1999; Fig. 1). This conclusion may change in the
future in the light of research focussed on apparently
long non-glacial episodes within the Proterozoic and
Phanerozoic.
2. Glacio-epochs defined
2.1. Structure of this paper
The overall timing and number of major glacial
episodes in Earth history is known within broad
bounds (Fig. 1) but there is no consensus regarding
their causes. The words of Broecker (2000, p. 140) that
‘we do not understand how it is that Earth's climate is
capable of achieving its glacial state’ were made in
regard to short-lived (and relatively well known)
Pleistocene glaciations but apply equally to the long
3 billion year pre-Pleistocene record. Despite the sparse
overall sedimentary record of glaciations older than
750 Ma (Figs. 1 and 2) it is difficult to disagree with
Crowell's suggestion (1999, p.2) that glaciers have
always been present somewhere on planet Earth even
at times of marked global warmth. Today, for example,
ice survives in the tropics in many areas of high
geodynamically-produced topography (e.g., Ehlers and
Gibbard, 2004). The sedimentary deposits of many
ancient ice masses were no doubt reworked, disguised
or lost entirely and they remain (so far) invisible to
sedimentologists. As will be shown, this is especially
the case where glaciers occur in collisional tectonic
settings. In contrast, the larger ice bodies that formed
during long lasting episodes of extended glaciation
were able to leave a sedimentary record because their
margins reached sea level. These extended episodes of
glaciation involve long-term climate change in response to processes such as plate tectonically altered
paleogeographies (termed ‘tectonic scale climate
change’; Ruddiman, 2001). These can be contrasted
with the much higher frequency Milankovitch-driven
and/or solar driven climate cycles that are bundled
within them (relatively short lived Ice Ages or glacial/
interglacial stages). The term ‘epoch’ is widely used to
informally designate an unspecified length of geologic
time (Neuendorf et al., 2005) and in this paper such
This paper commences with the tectonic setting of
the poorly known Archean glacial record and then proceeds sequentially to examine Proterozoic and Phanerozoic glacio-epochs ending with the more familiar
Cenozoic (Figs. 1 and 2). This review notes an essential
continuity between all glacio-epochs in that glaciation is
preferentially associated with plate tectonic processes
that act to elevate areas of continental crust. In this light,
the paper discusses the efficiency of extensional vs.,
collisional plate tectonic settings in generating tectonotopography and ice covers, and in preserving a sedimentary record of cold climates.
3. Archean glacio-epochs (c. 4–2.5 Ga)
3.1. A sparse record
Geology is not much help in resolving Archean climates because the sedimentary rock record is so sparse.
Local glaciation is evident at about 2.9 Ga and possibly
again at 2.8 Ga in southern Africa. Both these deposits
may represent ephemeral regional glaciations and may
not strictly qualify as glacio-epochs. The results of
climate modelling are contradictory. On the one hand is
the notion that a ‘hot’ greenhouse' climate prevailed for
much of Archean and Proterozoic time and prohibited
widespread glaciation (e.g., Kasting, 1987; Kramers,
2002). Knauth and Lowe (2003) examined the oxygen
isotope record preserved in Archean cherts and argued
that climates were as much as 40 °C warmer at 3.5 Ga
compared to the long term Phanerozoic norm as a consequence of a CO2 enriched atmosphere with concentrations 10,000 times greater than at present (Kasting
(1993). This model was questioned by Shields and
Veizer (2002) because of a lack of geologic evidence
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N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
and Veizer (2005) later proposed that global warming
was facilitated not by CO2 but by a muted Cosmic Ray
Flux (CRF) creating a reduction in cloud cover and a
lowered planetary albedo.
Others maintain that the Archean was much colder
as a consequence of a ‘faint young sun’ with an output
25–30% lower than that of today. Caldeira and Kasting
(1992) suggested that runaway glaciation resulting
from condensation of CO2 clouds at low temperatures,
was only prevented by greenhouse gases such as NH3
or CH4. Another ‘cold Archean’ model is that of
Zahnle and Sleep (2002) who modeled CO2 fluxes in
and out of the mantle. The rate of outgassing in the
Archean was perhaps twice that of today which would
result in CO2 rich atmospheres that would offset the
effects of a fainter young sun and the tendency for
irreversible glaciation. However, as Zahnle and Sleep
(2002) indicate, rates of Archean subduction (and thus
consumption of carbonate) were also substantially
greater with yet additional CO2 consumed by weathering of the debris thrown out by massive meteorite
impacts. They concluded that very cold, globally
freezing surface temperatures could be expected for
much of the Archean and early Proterozoic until about
2 Ga. Other authors too, have suggested the existence
of ice-covered Archean oceans only occasionally
melted by large bolide impacts; global freeze-thaw
cycles are seen as a key influence on the synthesis of
amino acids (Bada et al., 1994).
Most recently, Kasting and Howard (2006) argue for
a more temperate ‘warm’ early Earth somewhere between the hot and cold models reviewed above, suggesting that oxygen isotope data used to reconstruct
surface temperatures instead reflect changes in the composition of seawater with time.
Whilst the nature of Earth's early climates is
ambiguous, the onset of Earth's glacial geologic
record shortly after 3 Ga is broadly coincident with
the first undisputed evidence of microbial life (Noffke
et al., 2006) and the development of oxygenic
photosynthesis (Ono et al., 2006). By 3 Ga, an early
continent (Ur) is though to have included parts of the
present day Kaapvaal, western Dharwar, Singhbhum
and Pilbara cratons along with smaller crustal blocks
now part of East Antarctica (Rogers and Santosh,
2004). Glaciation is recorded in the late Archean
(∼2.9 Ga) Mozaan Group of South Africa (Young et
al., 1998) on the passive margin of Ur (Fig. 1).
Diamictites (up to 80 m thick) are present in the
Odwaleni Formation near the top of the 5000 m thick
dominantly marine Mozaan Group deposited on the
southeastern passive margin of the Kapvaaal Craton.
Diamictites1 containing striated clasts (von Brunn and
Gold, 1993) are associated with turbidites, pebble
conglomerates, volcanics and iron formations likely
indicating a submarine setting where glaciclastic
sediment was reworked downslope as debrites (von
Brunn and Gold, 1993). Modie (2002) interpreted
Archean diamictites and lonestones in shales within
the 2.782 Ga Nnywane Formation of Botswana as
glacially derived, but this is questioned by Evans
(2003a,b, p. 370).
If a ‘cold Archean Earth’ model is a reality, then
glacially influenced strata should be widespread and
preserved elsewhere other than on the Kaapvaal Craton
and Botswana but this does not appear to be the case. The
question then is where are they most likely to have been
preserved? Glacial strata correlative with the Mozaan
group of South Africa should be found in East Antarctica
given that fragments of the Kaapvaal Craton were rifted
and dispersed during the breakup of Gondwana (e.g.,
Wareham et al., 1998). Further clues are provided by
recent studies of the Archean Slave Province in the
1
It is necessary here to define several terms used throughout this
paper. The term ‘glaciclastic’ refers to sediment produced by erosion
at the base of glaciers and ice sheets. Such sediment is uniquely
fingerprinted by the occurrence of striated and glacially shaped
‘bullet’ clasts and far travelled extrabasinal lithologies. The greatest
yields of glaciclastic sediment occur from temperate wet-based
glaciers able to slide over (and thus abrade) their beds. Glaciers in
areas of permafrost (polar glaciers) are frozen to their beds and are
much less efficient producers unless parts of their mass are lubricated
by wet sediment promoting fast ice flow (‘ice streams’).
‘Glaciomarine’ refers to processes or deposits in a very narrow
zone in direct contact with the margin of a glacier (or ice sheet) that
reaches sea level and which is affected by meltwaters issuing from the
ice front. The term ‘ice contact marine’ or ‘proglacial marine’ are
alternate terms with the same meaning. This environment is recorded
by large, complexly structured ice contact deposits (‘morainal banks’)
dominated by the conglomeratic deposits of energetic subaqueous
meltwater fans. The term ‘glacially influenced’ was introduced by
Eyles et al. (1985) as an umbrella term for deposits accumulating in
more remote parts of a basin beyond the direct reach of glaciers (i.e.,
distal to the glaciomarine realm) but which nonetheless are composed
of glaciclastic sediment reworked by currents or gravity. Areas
affected by glacioisostatic or glacioeustatic sea level changes (or scour
by drifting ice bergs) can also be regarded as glacially influenced.
Within the subaqueous glacial realm, isolated clasts in fine-grained
sediment (lonestones; descriptive term) are dropped by ice (dropstones; genetic term) but discrimination between dropstones left by
glacial icebergs from other types of ice (sea ice, river ice) and the
climatic ramifications therein is problematic. Many lonestones are not
dropstones. Finally, the term diamict(ite) is a non-genetic term
referring to any poorly sorted deposits regardless of depositional
environment. A till(ite) is a diamict(ite) deposited directly from ice
either terrestrially on land or in the glaciomarine realm. A debrite is
the deposit of a debris flow. Benn and Evans (1998) and Eyles and
Januszczak (2004a,b) review the processes and environments that
result in the formation of till and diamict.
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
Northwest Territories in Canada (Polat and Kerrich,
2001; Bleeker, 2002). The Slave contains prominent
basalt and turbidite dominated greenstone belt emplaced
at about 2.6 Ga as autochthonous bodies on incipient
continental blocks during massive rifting events. Bleeker
(2002) argued that voluminous eruptions of basalt on
rifted continental crust created unstable conditions
because basalt has a density of 2.9–3.1 g/cm3 whereas
the underlying gneissic and granitic crust is lighter
(about 2.7 g/cm 3 ). This resulting Rayleigh–Taylor
instability ultimately leads to ballooning of the lighter
gneissic crust and the formation of large high standing
basement domes up to 60 km across. Uplift is coeval with
the appearance of turbidites and conglomerates in nearby
basins deposited by sediment gravity flows (such as the
Burwash Formation of the Slave Province) and are a
marked feature of other Archean cratons (e.g., Zimbabwe, Yilgarn). In Canada, these poorly sorted infills
are known as ‘Temiskaming-type’ deposits after the type
locality (named the Porcupine Group) in the Abitibi
Greenstone Belt of the Superior Province in Northern
Ontario. It can be speculated that a record of Archean
glacials could be preserved in these deposits as a
consequence of ice masses growing on uplifting granite
domes. Still older glacial deposits may be yet identified
recording the earlier assembly of Ur.
Ono et al. (2006) and Kasting and Howard (2006)
suggest a short-lived photosynthetic oxidation event at
2.9 Ga destabilized the methane-rich Archean atmosphere triggering the Mozaan glaciation. As Kasting and
Howard (2006) discuss (p. 1736) the presence of several
small Archean glaciations is more easily reconciled with
moderate Earth surface temperatures rather than the
extremely hot conditions favoured by some workers.
3.2. Summary
There are fundamental uncertainties regarding Archean climates because of the dearth of sedimentary
deposits and climate modelling yields very different,
opposed perspectives. Glaciation is recorded at about
2.9 and 2.8 Ga but is restricted to southern Africa. The
geodynamic setting indicates a passive margin setting. A
systematic search is needed for new deposits in other
basins.
4. Paleoproterozoic glacio-epoch (c. 2.4 Ga)
4.1. Rift related ice covers
The best-known earliest glacio-epoch is recorded at
about 2.4 Ga within the Paleoproterozoic Huronian
95
Supergroup in Ontario, Canada (part of the Southern
Province of Laurentia). Less voluminous and much less
well-exposed strata are preserved within the Baltic Shield
(in the Finnish sector of the Karelia craton), in South
Africa (Kalahari craton) and in North America such as
the Snowy Pass Supergroup of Wyoming (Marmo and
Ojakangas, 1984; Kohonen and Marmo, 1992; see Eyles,
1993, pp. 59–67; Pesonen et al., 2003; Bekker et al.,
2005). Other diamictites of glacial marine origin (the
Makganyene Formation) occur in the ∼2400 Ma Transvaal Supergroup in the Griqualand West Basin of South
Africa (Polteau et al., 2006). A close geographic clustering of these locations certainly appears likely within
what has been called Kenorland (Pesonen et al., 2003;
Fig. 1). Kirschvink et al. (2000) and Melezhik et al.
(2005a,b) argue that the Huronian and other deposits are a
record of one or more ‘global glaciations’ (Polteau et al.,
2006) that occurred in low to middle latitudes (Williams
and Schmidt, 1997; Kirschvink et al., 2000; Evans, 2000).
Geodynamically, Paleoproterozoic glaciation occurred in the context of early, plume-related rifting of
Kenorland (Fig. 1) that eventually resulted in Laurentia
separating from Baltica at about 2.1 Ga (Visser, 1981;
Ojakangas, 1985, 1988; Kohonen and Marmo, 1992;
Heaman, 1997; Ojakangas et al., 2001; Bekker et al.,
2001; Pesonen et al., 2003; Bekker et al., 2006). The
overall plate tectonic context is strikingly similar to
many of the later Neoproterozic glaciations where deposits were preserved for the most part in rift basins
during the breakup of Rodinia (Section 6.1). The
Huronian was deposited after 2.45 and is at least
12 km thick with four distinct tectonostratigraphic
successions (Elliot Lake, Hough Lake, Quirke Lake and
Cobalt). The most well known is the Gowganda
Formation (up to 1.7 km thick) that occurs at the base
of the Cobalt Group. It is little deformed and present
over a wide outcrop area of northeastern Ontario largely
within the structural confines of the so-called Cobalt
Embayment which was a failed rift arm extending into
the southern margin of the Superior Province (‘Superia’;
Stott, 1997; Calvert and Ludden, 1999; Young et al.,
2001; Bleeker, 2002) then part of Kenorland.
Diamictites and associated facies of the Gowganda
Formation have a long history of being interpreted in
classic ‘bed for bed style’ as exclusively climate driven.
‘Tillite’ beds were argued to record ice expansion and
non-tillite facies as interglacials (Young, 1981) but a
more complex glaciomarine setting was recognised by
Young and Nesbitt (1985). Miall (1985) emphasized the
contextual importance of subaqueous mass flow facies
interbedded with diamictites and argued that they had all
been deposited in a deep marine glacially influenced
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N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
basin. Mustard and Donaldson (1987) found additional
evidence of a marine setting. Miall pointed out the
difficulty of recognizing climatically driven glacial/advance retreat cycles in deposits where active tectonics
condition changing water depths, the supply of sediment
to the basin and the timing of mass flow.
Significantly, Yeo (1981) and Young and Nesbitt
(1985) suggested that uplift of rift shoulders adjacent to
a source of moisture such as the opening rift system, was
instrumental in promoting Huronian glaciation. Indeed,
it is now agreed that the Huronian stratigraphy likely
records several episodes of pulsatory rifting, uplift-generated cooling and subsidence along the southern continental margin of the Superior Province (Eyles, 1993;
Young et al., 2001). Each succession oversteps the
underlying succession (recording basin subsidence and
expansion) and consists of ‘lower’ synrift submarine
mass flow facies (glacially influenced diamictites,
graded conglomerates), ‘middle’ thick turbidites and a
coarser ‘upper’ stratigraphic cap of shallow marine to
fluvial facies (Fig. 3). The tripartite tectonostratigraphic
cycles of the Huronian are typical of marine rift basins
where there is a close control on deposit type and geometry by repeated reactivation of basin boundary faults
(e.g., Ravnas and Steel, 1998; Gawthorpe and Leeder,
2000; Hinderer and Einsele, 2002). Poorly sorted facies
at the base of each cycle record renewed uplift of rift
flanks and the shedding of coarse clastics from expanding ice covers into a deepening basin. Middle turbidites record maximum subsidence rates and thus
minimum sedimentation rates. As the rate of subsidence
decreases late in the rift cycle, so basin margin facies are
able to prograde and produce a fluvial and shallow
marine cap to each major tectonostratigraphic succession. Simple recognition of any short-lived glacial or
longer-term eustatic signals is rendered unlikely (e.g.,
Miall, 2005; see Section 6.2). The key point here is that
glaciers undoubtedly contributed to the sediment flux
to the rifted basin but that ice contact strata were
not widely preserved (Mustard and Donaldson, 1987).
Kohonen and Marmo (1992) referred to Paleoproterozoic glacials of the Urkkavaara Formation in Finland,
as ‘tectofacies’ because they accumulated in a failed arm
where sediment influxes were controlled by episodes of
faulting. In South Africa, banded iron formations occur
below and well above diamictites of the Makganyene
Formation. Diamictites are interbedded with and
volcanic tuffs marking volcanic activity and the eruption
of flood basalts during rifting (Polteau et al., 2006,
p. 271). Young (2002) drew attention to the
Fig. 3. Schematic glaciated rift basin during Rodinia breakup after 750 Ma (modified from Gawthorpe and Leeder, 2000). Primary glacial sediment is
extensively reworked by mass flow processes and terrestrial glacial facies are seldom preserved. Sedimentation is markedly diachronous as a
consequence of propagating faults and the non-synchronous formation and filling of different sub-basins. The fill of any one sub-basin comprises a
tectonostratigraphic succession recording changing relationship between subsidence and sediment supply. Marked intrabasinal variability in the
timing of rifting and the sedimentation response prohibits correlations of like facies (e.g., diamictites) and also wider extrapolation of age dates on any
one stratigraphic horizon to other basins worldwide.
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
accumulation of banded iron formations (BIFs) with
enclosed or at least semi-restricted rift basins separate
from the global ocean.
The formation and breakup of Kenorland accompanied marked changes in paleoatmospheric composition,
distinct weathering styles, fluctuations in carbon
isotopes and the accumulation of banded iron formations (e.g., the ‘Great Oxidation Event’ of Karhu and
Holland, 1996) (see also Bekker et al., 2001, 2005;
Lindsay and Brasier, 2002; Eriksson et al., 2005). It has
been speculated that climatic cooling may have been
initiated by chemical weathering of large exposures of
continental crust (Knauth, 2005; Melezhik et al., 2005a,
b). Kopp et al. (2005) and Kasting and Howard (2006)
speculate that Archean and Paleoproterozic glaciations
were triggered by rises in atmospheric oxygen concentration, counterbalancing the greenhouse warming of a
methane-rich atmosphere.
4.2. Summary
The well-developed relationship between Paleoproterozic rifting and glacial deposits suggests either a
causal relationship between rift-related uplift and climatic cooling or selective preservation of glaciated rift
deposits. Deposits are dominantly submarine debrites
associated with thick turbidites. They are part of very
thick marine tectono-stratigraphic successions, often
associated with volcanics, recording the changing interplay of subsidence rates and sediment supply as rifting
progresses. An association between glacials and banded
iron formations may reflect deposition in semi-enclosed
basins with incipient spreading centres.
5. Paleoproterozoic to Mesoproterozoic non-glacial
interval (c. 2.3–0.75 Ga)
5.1. Cycles of continental assembly and breakup but no
glaciation?
Earth's known glacial record has a prominent gap
between about 2.3 Ga and 750 Ma spanning the latter
part of the Paleoproterozoic and most of the Mesoproterozoic (Eyles and Young, 1994; Brasier and Lindsay,
1998; Crowell, 1999; Kah et al., 1999; Fig. 1). Veizer
(2005) proposed that the absence of glaciation could
reflect diminished star formation in the Milky Way and a
minimum in the Cosmic Ray Flux thereby producing a
warmed planet (see also Shaviv, 2003; Marcos and
Marcos, 2004). The absence of a glacial record may be
more apparent than real given the recent discovery in the
Kimberley district of Western Australia of 1.8 Ga old
97
glacially cut channels (Williams, 2005). Otherwise, the
lack of any marked glaciation is anomalous because all
other sedimentary environments found today on the
Earth's surface are represented (Eriksson et al., 2005,
p. 33) and geodynamic processes were no different from
those of the Neoproterozoic and Phanerozoic. During
this time, plate tectonics assembled several large continents such as Arctica, Nena and Columbia involving
repeated supercontinent cycles of crustal rifting and
collision (e.g., Starmer, 1996; Condie, 2002a,b; Rogers
and Santosh, 2004). In nearly all continents, long midProterozoic orogenic belts encircle blocks of Archean
crust and testify to protracted collisional events (e.g.,
Trans-Hudson, Capricorn, Trans North China, TransAmazonian, Yavapai/Mazatzal etc) during cratonization. Moreover, large landmasses such as Laurentia and
Australia lay at high polar latitudes by 1.15 Ga (Pesonen
et al., 2003). Throughout much of the Mesoproterozoic
time interval, the necessary preconditions (tectonics,
latitude) were in place for the formation of ice covers on
tectonically elevated topography.
The most conspicuous tectonic event capable of
producing widespread ice covers was the assembly of
Rodinia that commenced after 1.8 Ga. By 1.3 Ga this
had built the largest orogenic belt known to date
(Grenville Orogen) when a Himalayan-scale (glaciated?) topography 600 km wide and more than 10,000 km
in length formed along the suture between early North
America, South America (Amazonia) and Baltica
(Karlstrom et al., 2001). Peneplanation of the mountains
by c. 800 Ma produced large volumes of sediment that
could have contributed to reducing global pCO2. Significantly, Sheldon (2006) reports a significant drop in CO2
sometime between 1.8 and 1.1 Ga. It may well be the
case that glacial sediments survive as reworked deposits
within highly metamorphosed successions requiring
detailed sedimentological scrutiny. Rifting following the
Grenville Orogeny occurs between about 1 Ga and
750 Ma but again, no glacial deposits are reported (e.g.,
Timmons et al., 2001; Brewer et al., 2002; Wingate
et al., 2002).
5.2. Summary
The absence of any extensive glacial record during
the long Paleoproterozoic–Mesoproterozoic interval
between about 2.3 Ga and 750 Ma represents a large
gap in Earth's glacial history. Glaciation should have
been a common phenomenon given the known formation of several large landmasses and the orographic
effects of associated high standing orogens, but this does
not appear to be the case other than briefly and locally in
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N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
Australia at about 1.8 Ga. Possibly the rock record has
not been sufficiently well examined, deposits were extensively reworked and not preserved along active plate
margins, or as yet unknown processes acted to suppress
glaciation.
6. Neoproterozoic glacio-epoch (0.75 Ga to 545 Ma)
The Snowball Earth hypothesis posits that severe
Neoproterozoic glaciation occurred at low latitudes
(Kirschvink, 1992) and currently ranks as the most
controversial and polarized area of debate perhaps in the
whole of the Earth Sciences (see Jenkins et al. 2004;
Allen, 2006, 2007a; Fairchild and Kennedy, 2007, for
succinct summaries of competing ideas). Thick and
well-preserved Neoproterozoic glacial deposits occurs
on all the continents formerly within Rodinia but despite
the wealth of this sedimentary record, wider environmental reconstruction is hampered because of a lack of
age data (Bowring and Condon, 2006; Kendall et al.,
2006) and because of conflicting reconstructions of the
supercontinent's configuration (Fig. 4; Karlstrom et al.,
2001; Powell and Meert, 2001; Wingate et al., 2002;
Pesonen et al., 2003; Meert and Torsvik, 2003; Torsvik,
2003; Meert and Lieberman, 2004; Rogers and Santosh,
2004; Cawood and Pisarevsky, 2006; Trindade and
Macouin, 2007; Johnson et al., 2005). The Neoproterozoic drift of Rodinia from high latitudes to equatorial
latitudes is incontrovertible but the timing of this
movement with regard to glaciation(s) is still unclear.
There is agreement that Neoproterozoic glaciations
occurred against an overall tectonic backdrop of active
crustal extension as Rodinia broke apart. This is clearly
reflected in a glacial stratigraphic record dominated by
the deposits of former rift basins.
6.1. Stratigraphic record of Neoproterozoic cold climates:
the ‘glaciated rift’ debrite–turbidite association
With few exceptions, the sedimentary record of
Sturtian and younger land-based ice covers is largely
(but not exclusively; see below) preserved in subaqueously deposited offshore strata preserved in rifts and
along newly created passive margins (Fig. 3). As a consequence, an understanding of rift basin dynamics is
fundamental to unraveling the nature of the stratigraphic
record. Neoproterozoic ‘glaciated rift’ successions are
dominated by thick, poorly sorted mass flow deposits
(debrites) and associated turbidites found within successions marked by their great volume. Total thicknesses in
Fig. 4. The Neoproterozoic glacio-epoch and the break up of Rodinia. The precise paleolatitudinal position of the supercontinent is not known and
greatly controversial (see text). The purpose of this diagram is to simply show the very well-constrained association of thick glacially influenced
marine deposits and rifting (see text; Fig. 8). A: ‘paleo-Pacific’ rifting after 750 Ma (Sturtian glacio-epoch; Figs. 1 and 2). B: ‘paleo-Atlantic’ phase
after c. 610 Ma (Marinoan-Gaskiers glacio-epoch; Figs. 1 and 2). A: Arabia, Aus: Australia, EAnt; East Antarctica, Gr/Sc; Greenland, Scandinavia,
In; India, NCB; North China Block, NWA; Northwest Africa, SCB; South China Block, T; Tarim.
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
excess of 2 km suggest a dominantly marine not lacustrine origin, and were accommodated in rift complexes
extending through central Australia (Adelaide Rift
Complex; Young and Gostin, 1989, 1991), western
North America (Young, 1995; James et al., 2001;
Clapham and Corsetti, 2005) and South China (Wang
and Li, 2003; Jiang et al., 2003; Dobrzinski and
Bahlburg, 2007) and Oman (Allen, 2007a). These rifts
are regarded as contiguous in most Rodinian reconstructions and share a common tectonic history beginning as
early as 800 Ma (Wang and Li, 2003) possibly related to
mantle plume outbreaks below Rodinia (Li et al., 1999,
2006). New age constraints on the glacially influenced
marine deposits in Australia indicate much younger ages
than hitherto considered (Kendall et al., 2006; see
below). Nonetheless, the sedimentology of these strata
consisting of great thicknesses of debrites and turbidites
(e.g., Young and Gostin, 1991; Eyles et al., 2007)
indicate a rift setting. Thick rift-related glacial marine
successions also occur in Central Africa (Wendorff and
Key, 2006), Oman and Brazil (Young, 1992; Urban et al.,
1992; Leather et al., 2002; Kellerhals and Matter, 2003;
Allen et al., 2004; Tie-bing et al., 2006; Allen, 2007b)
and around the incipient rifted margins of the Iapetus
Ocean now preserved in several basins around the North
Atlantic and Greenland (Crowell, 1999; Bingen et al.,
2005). Many of these latter successions are characterised
by large volumes of cannibalized carbonate debris (e.g.
the Great Breccia of the Port Askaig Formation, Scotland
and similar facies in Norway; Arnaud and Eyles, 2002a,b)
recording instability of platform margins undergoing
extension and collapse (Vernhet et al., 2006; Eyles and
Januszczak, 2007).
Elsewhere, thick glacially influenced marine strata of
eastern Canada and Mongolia accumulated in volcanically influenced back arc basins (Gardiner and Hiscott,
1988; Eyles and Eyles 1989b; Lindsay et al., 1996; Allen
(2007b)). Those in eastern Canada (Gaskiers Formation)
formed on North African basement along the active
northern Gondwana margin (Eyles, 1990). Deposits were
subsequently rifted off as part of Avalonia–Cadomia
during the opening of the early Paleozoic Rheic Ocean to
be subsequently embedded within maritime Canada in
the Silurian (Gutierrez-Alonso et al., 2003). Other related deposits (e.g., Squantum ‘Tillite” and the Granville
Formation of northern France) have been recently cited as
glacial in origin (e.g., Snowballearth.org) in support of a
Marinoan global freeze event at about 580 Ma but this
ignores a long history of previous investigation favouring
a non-glacial origin involving debris flow and mixing of
coarse and fine sediment to form diamictite (see Fairbridge, 1947; Dott, 1961; Eyles, 1990; Crowell, 1999 and
99
refs therein) referred to as ‘mixtites’. The wealth of glacial
abrasional forms, such as striated pavements, ice-contact
and glaciotectonically-deformed deposits typical of
Pleistocene glaciations (e.g., Eyles, 1988; Benn and
Evans, 1998; Boulton et al., 1996; Anderson, 1999; Clark
et al., 2003) is not a characteristic of the Neoproterozoic
record but some may be locally preserved around basin
margins. The bulk of the ‘glacial’ record is biased toward
offshore basin depocentres.
Striated basement surfaces and true terrestrial tillites
deposited below glaciers are rarely preserved (e.g., Wang
and Li, 2003) with the major exception of the North
African craton, which appears to have been tectonically
stable (Proust and Deynoux, 1994) though Shields et al.
(2007) report the widespread association of volcanogenic strata with postglacial deposits across a large
portion of the Taoudeni Basin. In that basin, periglacial
sandstone wedges indicate subaerial contraction and
cracking of permafrozen ground (e.g., Deynoux et al.,
1989) similar to the deposits of the classic Marinoan
passive margin of South Australia (Williams, 1994).
Other wedge-like forms (e.g., Spencer 1971; Nystuen,
1976) occur in strata that have marine characteristics and
may record soft sediment deformation during rift-related
earthquakes (Eyles and Clark, 1985). By and large, the
global record of cold, arid conditions as proposed by the
Snowball model is very restricted compared to the
wealth of the offshore glacially influenced record.
6.2. Neoproterozoic glacioeustasy or tectonics?
Glacioeustatic sea level fluctuations of at least
1.25 km (an order of magnitude greater than those of
the Pleistocene) are invoked for a Marinoan glaciation at
c. 635 Ma based on assumptions of ice thickness on land
and on the oceans (2 km and 0.4 km respectively;
Hoffman et al., 2006). A corollary is that fine-grained
marine strata up to 1 km thick resting on Neoproterozoic
diamictites reflect postglacial glacioeustatic sea level
recovery (see review by Eyles and Januszczak, 2004a)
and are globally correlative (e.g., Christie-Blick et al.,
1988, 1995; Young, 1992; Pyle et al., 2004; de Alvarenga
et al., 2004; Pyle et al., 2004). In contrast, the thickness of
these successions points to an underlying tectonic
control involving thermal subsidence of newly formed
passive margins (see Eyles and Januszczak, 2004a,
pp. 31–2, 56–58). At times of maximum thermal
subsidence, blankets of fine-grained sediment accumulate
above coarser grained mass flow facies (‘rift climax deep
water systems’; Leppard and Gawthorpe, 2006) and
probably accounts for the thick, turbidite successions
found above many Neoproterozoic glacials (see above
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N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
Section 6.1) and hitherto interpreted as postglacial
‘glacioeustatic’ strata. In contrast, postglacial sea level
recovery returns water depths to their preglacial position
and cannot create new additional accommodation space
(see also Section 7.4 below). In general, intrabasinal
tectonic processes (and accompanying relative changes in
water depths) are more important than eustatic changes
(Miall 1997; Miall and Miall, 2001; Miall 2005) resulting
in marked variation in strata from one sub basin to another
(see Martin et al., 1985; Thomas 1991; Gawthorpe and
Leeder, 2000; Lund et al., 2003).
6.3. Diachronous response to regional tectonics or
globally synchronous megafreeze events?
Can Neoproterozoic glacial deposits be considered as
the ‘global benchmarks’ of Calver et al. (2004) (Kirschvink, 1992; Hoffman et al., 1998) or are they tectonically
conditioned and diachronous? On the one hand is the
model of tripartite glaciations named from oldest to
youngest Sturtian (c. 750 Ma), Marinoan (c. 635 Ma) and
Gaskiers; c. 580 Ma; Chen et al., 2004; Halverson et al.,
2005; Trindade and Macouin, 2007; Figs. 1 and 2).
Recently however, Calver et al. (2004) correlated the
Gaskiers with the Marinoan at c. 580 Ma and concluded
that there are only two ‘truly global ice ages' (Sturtian and
Marinoan; see Knoll et al., 2006, p. 15). A model of either
two or three globally-correlated glacial intervals may be
simplistic in the light of new ages on classic “Sturtian”
black shales of 643.0 ± 2.4 Ma in Australia much younger
than considered previously (Kendall et al., 2006) thereby
raising the possibility of diachronous regional ice covers
(Fanning and Link, 2004, 2006) throughout the Neoproterozoic. The limitation on testing the diachronous vs.
discrete models still remains the lack of high precision,
well-calibrated dates (Kennedy et al., 1998; Bowring
et al., 2003; Bowring and Condon, 2006).
The dominantly rift-related tectonic setting of much
of the Neoproterozoic glacial record is central to the
question of the synchroneity or diachroneity of individual glaciations but tends to be ignored in favour
of simple lithostratigraphic correlations. For example,
McCay et al. (2006) pigeonholed the glacial deposits of
the Scottish–Irish Caledonides into a trinity of globally
correlative glaciations without a single specific age date
on any horizon. Yet, this succession is in part, tectonically controlled (Arnaud and Eyles, 2006) made up
of distinct tectonostratigraphic successions (Dempster
et al., 2002; Hutton and Alsop, 2004) recording intrabasinal tectonic controls on the initiation and preservation of glacial deposits. Adherents of the Snowball Earth
model only briefly acknowledge the existence of ‘paleo-
geographic boundary conditions to global climate’ (i.e.,
tectonics) (Halverson et al., 2004, 2005, p. 1201).
The lack of a dating control looms large at a time of
very rapid plate motions. Meert and Torsvik (2003)
categorized existing reconstructions of Rodinia between
1110 and 500 Ma as ‘extremely fluid and controversial’
(p. 282) a finding echoed by Tohver et al. (2006) who
argued that West Gondwana cratons were at middle to
high latitudes between 700 and 575 Ma and did not
reach an equatorial position until the Cambrian well
after the end of Neoproterozoic glaciation. Consequently, Fig. 4 simply depicts the known association of rifted
margins with glacial deposits rather than showing a
specific paleogeography. Evans (2000, 2003a,b), Maloof et al. (2002), Li et al. (2004), Meert and Torsvik
(2004), Kilner et al. (2005), Tohver et al. (2006) and
Trindade and Macouin (2007) review a substantial
paleomagnetic database. Of this, the Elatina Formation
of South Australia is currently regarded as ‘the most
thorough paleomagnetic determination of any Precambrian rock unit’ indicating glaciation at low paleolatitudes (Sohl et al., 1999; Raub and Evans, 2006). The
presence of cold climate ice wedge casts in associated
strata (Williams and Tonkin, 1985; Williams, 1994;
Young, 2002) has been explained by marked changes in
the angle of obliquity of the Earth's spin axis (Williams
and Schmidt, 2004 but see Bingen et al., 2005). Nonetheless, there still remain major uncertainties about the
precise depositional environment(s) recorded by the
Elatina Formation in regard to the diamictites within it,
the paleogeography of any ice covers at the time,
and their age (Lemon and Gostin, 1990; Eyles and
Januszczak, 2004a,b). That the Elatina is ice contact
glacial in origin is based entirely (and tentatively) on the
interpretation of a single thin (∼ 5 m) diamictite horizon,
in contact with the underlying Trezona Formation, as an
‘ice-pushed tillite’ that is noted as the ‘only example’ in
the entire succession (Lemon and Gostin, 1990, p.152;
Eyles and Gammon, 2007). Despite uncertainty regarding the primary ice-contact origin of this classic deposit,
it is the key global reference deposit for low latitude
glaciation (see Knoll et al., 2006). The database of
Trindade and Macouin (2007) includes deposits where a
glacial origin has never been definitively demonstrated
(e.g., Squantum Member of the Boston Bay Group).
Namibian strata (‘Ghaub Tillite’) held to record
Snowball conditions and the growth of glaciers on low
latitude carbonate reefs (Hoffman and Prave, 1996;
Hoffman et al., 1998; Hoffman and Schrag, 2002;
Halverson et al., 2005; Hoffman, 2005) were the focus
of several earlier studies (e.g., Schermerhorn, 1974;
Porada and Wittig, 1983a,b; Miller, 1983; Martin et al.,
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
1985) that could not recognise a glacial influence on
sedimentation in a complex extensional basin (see Eyles
and Januszczak, 2007). Advances in understanding
paleoclimates is fundamentally dependent on detailed
facies and basinal evaluations of deposits by glacial
sedimentologists conversant with Pleistocene and modern glacial depositional systems. This is clearly evident
in regard to climate models that invoke severe freezing
whereas the rock record tells otherwise.
6.4. An active hydrosphere; climate models and
sedimentology
The Neoproterozoic geologic record provides no
evidence as to planetary albedos or obliquity, ocean
circulation, solar luminosity, or atmospheric carbon
dioxide levels etc., but these are key inputs to climate
models (Crowley and Baum, 1993; Frakes and Jenkins,
1998; Chandler and Sohl, 2000; Hyde et al., 2000;
Baum and Crowley, 2001; Crowley et al., 2001; Poulsen
et al. 2001, 2002; Pierrehumbert, 2002, 2004; Poulsen,
2003a,b; Lewis et al., 2003; Tajika, 2003; Ridgwell et al.
2003; Donnadieu et al., 2003, 2004a,b,c; Poulsen,
2003a,b; Peltier et al., 2004; Romanova et al., 2005;
Sohl, 2006; Nédélec et al., 2007a,b; Godderis et al.,
2007). Given the constraints in input data, model
outputs relate more to the limitations of the models
themselves and explain large discrepancies.
Nearly all climate models stress the essential role of
sea ice in lowering planetary albedos as a necessary
condition for a snowball planet, but sedimentary evidence for its importance as a geologic agent is largely
missing. Globally distributed ice rafted dropstone facies
should be an integral part of any Snowball Earth with
frozen ocean surfaces (with ice thicknesses of up to
400 m; Hoffman et al., 2006) and thick shore ice capable
of freezing onto and exporting substrate materials.
Marine deposits with far travelled ice-rafted debris
should be widespread but this is not the case (e.g.,
Macquaker and Keller, 2005). Similarly, there is (as yet)
no convincing record of the scouring and deformation of
passive marine strata by the keels of sea ice such as
occurs today over enormous areas around the margins of
the Arctic Ocean, and which also occurred in water
depths of as much as 1 km below Pleistocene icebergs
(Polyak et al., 2001; Kristoffersen et al., 2004). The
demonstrable absence of ice-rafted debris in successions
that supposedly span assumed extreme climatic conditions (e.g., Polarisbreen Group of Svalbard) has been
explained as a consequence of a ‘fully frozen ocean’
(Halverson et al., 2004, p. 316) and the inability of ice to
drift. This scenario fails to consider the dispersal of
101
floating ice containing englacial debris at the end of any
severe glaciation. In addition, ice rafted sedimentary
facies similar to late Pleistocene Heinrich layers
produced by ocean going armadas of icebergs travelling
several thousand kilometres from their parent ice sheets
(e.g., Andrews, 1998) are persistent features of
Cenozoic (and possibly Paleozoic) glacio-epochs
(Eyles et al., 1997; Crowell, 1999). Such layers may
record a fundamental inherent instability of marine ice
sheets and should be present in the Neoproterozoic.
Their apparent lack of preservation indicates that the
global reach of floating ice was limited. On the other
hand, glendonite within Neoproterozoic cold water
carbonates of the Windermere Supergroup in northwest
Canada suggested to James et al. (2005) extensive
freezing of ocean water. Glendonite is a calcite
pseudomorph (after the mineral ikaite) forming today
in cold, high latitude organic and carbonate-rich bottom
sediments (see Marland, 1975; Suess et al., 1982;
Buchardt et al., 1997). It occurs profusely within highlatitude, shallow marine shelf successions of the early
Permian Sydney Basin in New South Wales when cooltemperate forest covers grew inland and cold Antarctic
bottom currents flowed north from a late Paleozoic
Antarctic ice cap (Eyles et al., 1997, 1998). These
conditions cannot easily be compared with extensive
freezing of Neoproterozoic ocean waters but ice wedge
casts (Deynoux et al., 1989; Williams and Schmidt,
2004) provide convincing evidence of cold conditions
well below − 6 °C close to sea level.
Sedimentologic evidence for an active hydrosphere
finds support from paleobiologists. The notion of a
possible link between glaciation and biologic evolution
is long standing (Rudwick, 1964) and recent views hold
that extreme global cold had a profound effect on life
(Hoffman and Schrag, 2000, 2002; Liang et al., 2006).
This relationship is however currently categorized as
‘overstated’ (Corsetti et al., 2006a,b, p. 127). Complex
metazoans (e.g., Ediacaran fauna) first appear in the
rock record 4 million years after the so-called Gaskiers
glaciation at around 582 Ma but the morphological
complexity of this faunal group suggests they evolved
preglacially (Vickers, 2006). No clear ‘cold effects’ are
identified in the record of sensitive shallow dwelling
marine organisms in glacially influenced and open marine environments (Grey et al., 2003; Gorjan et al., 2003;
Olcott et al., 2005; Corsetti et al., 2006a, b). In fact, in
that part of the Australian record which spans the classic
Sturtian glacial event, the Acraman–Bunyeroo bolide
impact event just after 580 Ma may be of greater
environmental importance (Williams and Gostin, 2005;
McKirdy et al., 2006). Data also suggest a key stimulus
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on Neoproterozoic biologic evolution and in particular,
the emergence of multicellular Ediacaran organisms was
the breakup of Rodinia involving enhanced nutrient
fluxes from mid-ocean spreading centres and generation
of large newly rifted margins with shallow water.
Neoproterozoic rifting and associated paleogeographic
changes are increasingly recognised as a major stimulus
on increased speciation rates and changing paleobiogeographic patterns (Lieberman, 1997, 2005; Lindsay
and Brasier, 2002; Xiao and Kaufman, 2006). By analogy with the breakup of Pangea, global oceanic circulation (and thus climate) is profoundly altered by small
changes in continental positioning and ocean basin
shape and depth (Poulsen et al., 1998; Goddéris et al.,
2007). Neoproterozoic glaciation may have been promoted by biotic weathering of newly ‘greened’ continental surfaces (Hedges, 2003).
Neoproterozoic rift basins contain very substantial
thicknesses many kilometres thick of glacially influenced
submarine fan deposits dominated by muddy debrite/
turbidite facies. Thicknesses of such strata necessitate
erosion by wet-based glaciers sliding over their beds
(e.g., Hooke and Elverhoi, 1996; Eyles, 2006). Normal
glacial and non-glacial sedimentary processes operated
within a fully functioning hydrological cycle (McMechan, 2000; Leather et al., 2002; Condon et al., 2002;
Allen et al., 2004; Arnaud, 2004; Eyles and Januszczak,
2004a; Arnaud and Eyles, 2006; Sohl, 2006; Allen,
2007b; Eyles et al., 2007; Dobrzinski and Bahlburg,
2007). It is not possible to create such thick sedimentary
records of marine glacial environments without the
ability to produce and flush large volumes of glaciclastic
sediment to marine basins by water. Thick glacial marine
deposits do not agree with a hard Snowball Earth scenario
involving dry based ‘static’ glaciers permafrozen to their
beds, thick sea ice and a non-functioning hydrosphere
(see Gaucher et al., 2003, 2005).
6.5. Carbon isotope excursions
A role for tectonics is indicated by recent work in
regard to carbon isotope excursions. Isotope profiles are
routinely used to for precise chemostratigraphic correlations of Neoproterozoic strata (see Halverson et al.,
2005; Corsetti et al., 2006a,b) based on the assumption
that strong negative shifts in 13C must record global
freeze/thaw events (Hoffman and Schrag, 2000, 2002).
An outstanding issue is that once again, ‘global isotope
curves’ are not constrained by radiometric dating but are
reliant on presumed lithostratigraphic correlations from
one basin to another and from hemisphere to hemisphere
(Halverson et al., 2005; Bowring and Condon, 2006;
Knoll et al., 2006). Many isotope shifts after 1500 Ma
are not associated with glacial deposits (e.g., Alene et al.
2006; Azmy et al. 2006; Kaufman et al., 2006; Polteau
et al., 2006). Other Neoproterozoic records are of low
stratigraphic resolution, show trends that diverge from
the idealized Snowball model of negative 13C shifts, and
which vary from basin to basin (e.g., Walter et al., 2000;
Kennedy et al., 2001a,b; Shields et al., 2002; Hoffman
et al., 2002; Lindsay and Brasier, 2002; de Alvarenga
et al., 2004; Halverson et al., 2005; Le Guerroué et al.,
2006; Corsetti et al., 2006a, and refs above).
Where excursions occur in the Neoproterozoic they
are regarded as being uniquely of climatic significance
(Halverson et al., 2005) despite the widespread deposition of host sediments in enclosed rift basins possibly
separated from the world oceans. Elsewhere in the
Mesoproterozoic, prominent carbon isotope excursions
are explained as a product of enhanced burial of organic
carbon during the assembly of Rodinia (Bartley et al.,
2001). Condie et al. (2001) emphasized the importance
of changes in the rate of organic carbon burial in newly
formed rift basins (e.g., Fig. 3) divorced from the global
ocean during the breakup of Rodinia (see also Gammon
et al., 2005). Isotope fluctuations may thus reflect tectonic rather than purely palaeoatmospheric and palaeoceanographic events (Brasier et al., 2002). Tectonic
activity not only buries carbon but it may also under
certain conditions fundamentally limit the use of carbon
isotopes to infer paleotemperatures by altering the isotope system by circulating basinal fluids (Foden et al.,
2001; Knauth 2005). Melezhik et al. (2005a) attribute
major shifts in 13C in the Paleoproterozoic to intrabasinal
geochemical factors arguing that the magnitude of excursions in the Proterozoic is much smaller than assumed. Other complicating factors identified by
Kaufman et al. (2006) and Alene et al. (2006, p. 95)
including diagenetic effects being very different in
Neoproterozoic sediments where bioturbation and mixing is absent compared to Phanerozoic environments.
There is growing evidence for a non-unique relationship
between Neoproterozoic climate and carbon isotope
excursions (Le Guerroué et al., 2006). In the Neoproterozoic Adelaide Rift Complex, positive excursions of 13C
are not restricted to glacial intervals but are primarily
associated with rifting events and volcanic activity
(Foden et al., 2001). These workers reported significant
disturbance of both C-and Sr-isotopic compositions of
fine-grained clastic sediments, carbonates and basalts at
several stratigraphic levels. There, the resetting of isotopic systems was attributed to uplift of basin margins
and the recharge of basin fluids by low salinity meteoric
waters.
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
6.6. Summary
The bulk of the Neoproterozoic glacial record is
stored within thick marine debrite-turbidite successions
that accumulated within rift basins. Terrestrial ‘tillites’
and associated deposits are poorly represented. Neoproterozoic glaciers were wet based and produced abundant
meltwater and sediment incompatible with catastrophically cold conditions of a hard Snowball Earth. The
breakup of Rodinia took place over a 200 million year
period and by analogy with other episodes of rifting
there was significant along-strike diachroneity in the
timing of rifting, basin formation and glacially influenced sedimentation (Kendall et al., 2006). Large-scale
rearrangement of landmasses and oceanic configurations created by an evolving disintegrating supercontinent may have played a key role in climate change.
There is growing recognition that Neoproterozoic glaciations were initiated as regional ice centres (Halverson
et al., 2005, p. 1198) whose growth was diachronous
(op cit., p. 1198) countering the longstanding use of
glacial deposits as precise global time markers. Earlier
ideas of ‘instant glaciation’ involving notional albedofeedback mechanisms and runaway refrigeration are
now underplayed (see Halverson et al., 2005).
In the light of the substantial gaps in knowledge
identified above, and the emerging theme of diachroneity of Neoproterozoic glaciation, it is profitable to revisit
the conceptual underpinning of current efforts to subdivide Proterozoic time using Global Stratotype Sections and Points (GSSP). Knoll et al. (2006, p. 14)
believe that ‘the great ice ages that wracked the later
Neoproterozoic world… were global in impact, and
because they are associated with carbon isotopic excursions larger than any recorded in Phanerozoic rocks, the
glaciations offer what are undoubtedly our best opportunities for the sub-division of Neoproterozoic time’. It
can be argued in fact that the geologic consensus is
moving away from catastrophic global freeze events and
instantaneous deglaciations. In contrasts to ‘wracking’
the world, the Neoproterozoic rock record informs us
that glaciers were wet-based and may have been part of
diachronous events as tectonotopography evolved during the dispersal of crustal blocks.
7. Lower Paleozoic Saharan glacio-epoch
(c. 440 Ma)
7.1. Long lasting or short lived? And just how big?
After the breakup of Rodinia, the North Africa craton
edged northwards within the south polar circle (Scotese
103
et al., 1999; Pharoah, 1999; Vercoli and Le Herisse,
2004; Fig. 5) but despite the high paleolatitude, only at
the end of the Ordovician is there geologic evidence
for ice over North Africa. This is the basis for the
influential model of a single and short-lived (∼ 1 Ma)
Late Ashgillian glacial event (see Brenchley et al.,
1995). An apparent lack of long lasting ice covers could
simply reflect the erosion or non-preservation of any
older glacial strata. Indeed, some authors propose a
10 Ma long North African glaciation that started much
earlier in the Ordovician (Ghienne, 2003) and lasted
well into the Silurian (Grahn and Caputo, 1992; Caputo,
1998; Pope and Read, 1998; Crowell, 1999; Saltzman
and Young, 2005). The last (and thus best known) Late
Ordovician Saharan ice sheet formed during a time of
high (16 × the modern value) atmospheric CO2 (Torsvik
and Cocks, 2004; Fig. 2E). The ice sheet may have been
comparable in size to the last North American Laurentide Ice Sheet (∼ 36 × 106 km3) and expanded eastward
from North Africa onto the Arabian platform (Deynoux
et al., 1985; Vaslet, 1990; Sutcliffe et al., 2000). There
are however, continuing uncertainties over its true dimensions (and thus volumes). A Southern African ice
mass is regarded by some as an entirely separate outlier
(Young et al., 2004) whereas others suggest it was
part of a continuous ‘pan-African ice sheet’ extending
over more than 60° of latitude (Scotese et al., 1999;
Veevers, 2004; Ruban et al., 2007). Traditional evidence
for the volume (and timing) of ice covers, derived from
so-called ‘glacioeustatic’ sea level fluctuations recorded
on distant continental shelves is problematic (see
Section 7.3).
Crowley and Baum (1995) and Kump et al. (1999)
modeled a notional North African ice sheet and argued
that it responded like its Cenozoic counterparts to orbital
Milankovitch variables (see also Williams, 1991). This
finding has been used as the foundation of a sophisticated glacial depositional model by Sutcliffe et al.
(2000) that sought to find lithostratigraphic correlations
between horizons more than 6000 km distant in Northern and Southern Africa otherwise entirely unconstrained by age or biostratigraphic data. They argued
that the deposits reflected two orbitally driven 100 ka
cycles of ice sheet expansion and decay and could be
correlated by ‘process based interpretation of sedimentation’ (p. 968) i.e., conventional bed-for-bed lithostratigraphy, from north to south.
Elsewhere, outlying ice covers formed in areas of
high topographic relief along the tectonically active
proto-Andean margin of Gondwana in what is now
Peru–Bolivia (Astini, 1999; Diaz-Martinez and Grahn,
2007). Diaz Martini and Grahn (2007, p. 77–78) were
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N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
Fig. 5. Palaeogeography of the Late Ordovician Saharan glacio-epoch at c. 440 Ma after various sources (principally Fortey and Cocks, 2003). The
northern margin of Gondwana was the locus of active extension after 480 Ma and extension-related uplift of the Gondwanan Highlands may have
triggered polar Saharan glaciation after 440 Ma. Outlying ice masses lay on the proto Andes and in southern Africa where they reached sea level
(Cancanari Formation and Pakhuis Formation respectively). The Saharan ice sheet was short lived and disappeared by the Early Silurian but ice
remained over the uplifted active margin of South America into the Devonian of Brazil and Bolivia but not over the pole (see text). When Gondwana
collided with Laurentia to form Pangea beginning in the mid-Carboniferous this remnant ice would expand to form an extensive Gondwanan ice
complex (Fig. 6).
at pains to emphasize that existing global palaeoclimatic reviews (e.g., Raymond and Metz, 2004) fail to
consider local tectonic controls on the existence and
timing of ice covers along the western Gondwana
margin. These glaciers left glacially influenced marine
strata dominated by mass flow deposits. Because of
local tectonic controls, the timing of these ice masses is
thought to have been out of phase with the main North
African glaciation(s) (Diaz-Martinez and Grahn, 2007).
Ice covers extended well into the Silurian when Brazil
lay at the South Pole and this area was to subsequently
experience glaciation again in the Devonian (see
Section 8 below). It can be ventured that ice was
present along topographically elevated parts of the
active proto-Andean margin throughout much of the
Early Paleozoic.
7.2. Geomorphic, sedimentologic and stratigraphic
evidence of glaciation in North Africa
Late Ordovician glacial deposits are relatively thin
(b200 m), mostly coarse grained and rest on a prominent channeled unconformity (e.g., Le Heron, 2007).
Erosional stripping of overlying shale has exposed the
original geomorphology of the glacial depositional
systems below. Beuf et al. (1971), Trompette (1973)
and Vaslet (1990) recognised terrestrial glacial and cold
climate landforms such as eskers, moraines, drumlins,
periglacial polygon structures, pingos, and channels cut
by subglacial and proglacial meltwaters (see Tucker and
Reid, 1973; Hirst et al., 2002; El-ghali, 2005; Armstrong
et al., 2005). The presence of multiple bedding planes
that expose striated surfaces across what was soft sand,
is seen either as the product of repeated shear either
below a grounded glacier (e.g., Sutcliffe et al., 2000,
2005; Deynoux and Ghienne, 2004; Le Heron et al.,
2005) or grounding ice floes (Woodworth-Lynas and
Dowdeswell, 1994). In this regard, Moreau et al. (2005)
have mapped what they argued to be streamlined glacial
lineations (drumlins) recording ice streams that flowed
north from an area of higher topography in the interior
(the ‘Gondwana Highland’ of Luning et al., 2000) north
toward the Gondwanan continental margin (Ghienne
and Deynoux, 1998; Ghienne, 2003). The assumption is
that the modern ground surface across a large area of
North Africa is an ancient exhumed landform and that
the depth of post-Ordovician stripping has been
precisely the same everywhere.
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
7.3. Sea level fluctuations; are they exclusively
glacioeustatic?
The size of an extensive North African Saharan ice
sheet is traditionally constrained on the basis of thirdorder glacioeustatic sea level changes recorded on distant continental shelves and cratonic interiors (Dennison, 1976; Ross and Ross, 1992; Pope and Read, 1998;
Izart et al., 2003). This model is deeply rooted in the
long standing assumption that changes in relative water
depths recorded in marine rocks (together with associated variations in 13C and 18O values; e.g., Brand et al.,
2006) are driven exclusively by ice sheet growth and
decay (see Zhang and Barnes, 2002 and refs therein). A
well-cited paleobiological model attributes Late Ordovician extinctions to glacially driven fluctuations in sea
level, ocean circulation and chemistry (see Webby and
Laurie, 1992; Armstrong, 1995; Brenchley et al., 1995,
2003; Sheehan, 2001; Munnecke et al., 2003; Twitchett,
2006).
For a number of reasons the ‘glacioeustatic’ model
for reconstructing ice sheet volumes and explaining
Lower Paleozoic extinctions is full of uncertainty. Late
Ordovician glaciation had little or no effect on phytoplankton living along the northern glaciated margin of
Gondwana (Vercoli and Le Herisse, 2004). Delicate
organic walled microfossils (acritarch and chitinozoan
assemblages) in Ashgillian strata show no extinction
events, in fact, microphytoplankton became more diverse through the glacial interval, which is associated
with new speciations and morphological innovations.
Indeed, the principal extinction event when many taxa
died out occurs earlier in the Middle Ordovician (late
Llanvirn) (Vercoli and Le Herisse, 2004). This is an
important finding because acritarchs occur at the base of
the marine trophic system and are sensitive indicators of
changes in the total biomass. Acritarchs constitute the
‘canary in the mine’ in regard to changes taking place in
the numbers and diversity of primary producers. If the
North African ice sheet had no negative effects on
marine microorganisms living in close proximity, it is
difficult to see how glaciation per se had any widespread
global influence on other groups (Munnecke et al.,
2003).
Late Ordovician extinctions were highly selective and
also diachronous occurring first in deeper water communities before affecting those of shallow water. This is
the opposite of what might be expected had glacioeustatic sea level lowering been the prime cause as widely
invoked. That a relatively small Saharan ice sheet, accompanied by small ice masses in South America perhaps could set in motion a series of global biological
105
crises begs the question of the role of other causative
biogeochemical processes unrelated to glaciation. There
appears to be scant recognition of a facies/substrate
control on marine fauna during the Late Ordovician,
viewed as responding primarily to changes in ocean
water masses. Many so-called ‘extinctions’ appear to be
times of altered diversity followed by reemergence of
‘Lazarus faunas’ (Munnecke and Servais, 2007). The
possible role of tectonically modulated changes in water
depths and thus on substrates and associated ecosystems,
goes unremarked.
The ‘dynamic topography’ model demonstrates that
Paleozoic depositional successions (‘sequences’) on
‘stable’ shelves and cratons are not primarily the result
of global sea level changes (as was originally thought)
but variation in relative sea level conditioned by mantle
processes below the overlying plate. These act to elevate
or depress continental surfaces independent of global
sea level (see Burgess et al., 1997; Mitrovica et al.,
2000; Miall, 2005; Eriksson et al., 2005; Vakarelov
et al., 2006). Thus the assumption of an exclusively
glacioeustatic control on changing Late Ordovician and
Early Silurian water depths (Brand et al., 2006) can be
challenged. A clear illustration is provided by Zhang
et al. (2006) who found significant differences in water
depth changes recorded in coeval Late OrdovicianSilurian strata on Appalachian and Arctic coasts of
North America. The rapid ‘glacioeustatic’ sea level rise
during the persculptus Zone (supposedly the main
Ordovician deglaciation event; Brenchley, 2004) could
not be recognised in the south. They referred to this as a
‘paradox’ which indeed it would be if global eustasy
were the only control on water depths. In reality of
course, given the very different tectonic settings of these
two widely separated margins, incongruence of relative
sea level records is to be expected because their subsidence histories were different. Zhang et al. (2006,
p. 266) argue that there is ‘significant evidence against a
universal sea level curve in the latest Ordovician-Earth
Silurian’.
From the above, the timing, volume nor longevity of
any North African ice sheet can be readily derived from
changes in water depths recorded at any one site distant
from centres of glaciation (cf. Nielsen, 2003). This represents a remarkable gap in our understanding of
Earth's glacio-epochs.
7.4. Tectonic influences on glaciation in North Africa;
the importance of black shales
A regional, long-term tectonic influence on sedimentation patterns along the Late Ordovician North
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N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
African continental margin can be suggested. This is
based on consideration of north draining paleovalleys
identified as ‘tunnel valleys’ or ‘fiords’ (Ghienne, 2003;
Armstrong et al., 2005; Turner et al., 2005) and their
overlying thick shale cap. Paleovalleys are incised into
underlying preglacial shelf deposits and are blanketed
by black, graptolitic shales of Early Silurian age. This
succession is classically interpreted as a consequence of
glacial erosion of an shelf exposed during glacioeustatic
fall accompanying ice sheet expansion, in turn, followed
by postglacial sea level rise during the Parakidograptus
acuminatus graptolite zone (Luning et al., 2000). Reported paleovalleys are however, broad shallow features
that lack the overdeepened form of glacially cut channels and fiords typical of Pleistocene forms. It remains
to be demonstrated that they are not simply ‘lowstand’
submarine valleys part of a widespread ravinement surface accompanying broad-scale uplift along the northern
Gondwanan continental margin. The model of Seidler
(2000) for submarine channels in the Tertiary of East
Greenland is an appropriate analog. Overlying Early
Silurian shales (that include lowermost ‘hot shales’ rich
in organic matter) are exceptionally thick (1000 m +)
and extend from Morocco to Saudi Arabia. The thickness alone indicates that a simple ‘postglacial flooding’
model is unrealistic because glacioeustatic recovery
does not create additional accommodation space. Final
ice melt simply returns sea level to its preglacial level. In
certain cases, a short-lived (thousand years?) phase of
increased accommodation occurs where glacioisostatic
rebound lags behind the rapid rise in sea level during
initial deglaciation (witness the widespread marine
transgressions at the end of the last glaciation around
Canada's Hudson Bay lowlands and St.Lawrence
Valley). ‘Marine overlap’ is temporary however, and
marine successions are rapidly raised above sea level
and eroded as a consequence of rapid crustal rebound.
One kilometre thick Silurian shales in North Africa
point to the creation of new and tectonically driven
accommodation space along that margin.
The typical Late Ordovician tripartite stratigraphic
motif of paleochannels, coarse clastic infill and thick
overlying shale likely represents a distinct tectonostratigraphic succession. This succession can be interpreted in
terms of an initial phase of crustal extension and upliftrelated climate cooling and glaciation along the North
Gondwana margin, followed by deglaciation, thermal
subsidence and flooding. This model suggests a lengthy
early Paleozoic glacio-epoch lasting much more than
1 million years triggered by broad scale extension along
the North African margin and the detachment of several
terranes. During end Ordovician glaciation Laurentia
was separated from North Africa by a wide ocean
(Iapetus) (Fig. 5) that eventually closed when several
terranes detached from the North African margin (e.g.,
Avalonia, Armorica including Iberia and Sardinia, Meguma; Mallard and Rogers, 1997; Sanchez-Garcia et al.,
2003; Fortey and Cocks, 2003; Vercoli and Le Herisse,
2004; Torsvik and Cocks, 2004; Ruban et al., 2007)
opening up the Rheic Ocean in their wake. The precise
timing of separation is unclear (Scotese and McKerrow,
1991; Torsvik et al., 1996; Gutierrez-Alonso et al., 2003;
Keppie et al., 2003; Cocks and Torsvik, 2006). Some
terranes now found as the Paleozoic massifs of central
and southern Europe may contain a glacially influenced
marine record (Robardet and Dore, 1988) but deposits
await sedimentological study and reevaluation (see
Long, 1991). A glacial influence on sedimentation in
Late Ordovician strata in terranes of eastern Canada is
evident in central Newfoundland (McCann and Kennedy, 1974) and Nova Scotia (e.g., White Rock
Formation; Schenk, 1972; MacDonald et al., 2002).
The earlier literature had assumed that a North African
ice sheet flowed across into eastern Laurentia at a time
when these continents were thought to be immediately
adjacent within Pangea. In fact, these glacially influenced marine strata accumulated autochthonous to
North America as part of the sedimentary cover on
terranes that rifted off from North Africa when the Rheic
Ocean opened. Those in Newfoundland form part of the
Dunnage terrane and, with other terranes (Gander,
Avalonia), was obducted against the North American
continental margin during the Taconic–Salinic Orogeny
(Late Ordovician–Silurian). Those in Nova Scotia are
part of the Meguma terrane emplaced during the midDevonian Acadian Orogeny (Schenk, 1991). The recent
identification of significant post-Gondwana uplift across
central and southern and central Africa in response to a
mantle ‘superswell’ by Gurnis et al. (2000) provides an
additional mechanism for creating an elevated dynamic
topography on which ice could nucleate as an outlier ice
cap. The development of a mantle superswell could
possibly have been a precursor to uplift along the late
Ordovician North African margin. Smith (1997, p. 176)
suggested that extensional faulting and extensive
footwall uplift was the trigger for widespread snow
accumulation across the interior of Gondwana.
Late Ordovician pCO2 levels were as much as sixteen
times that of today (Crowley and Baum, 1995; Wang
et al., 1997; Herrmann et al., 2004; Kaljo et al., 2004;
Tobin et al., 2005; Fig. 2E) begging the question as to
what processes might have drawn down pCO2. Increased
organic productivity of plankton and elevated rates of
carbon burial (Brenchley et al., 1995), accelerated
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
107
accumulation of carbonate on shelves (Villas et al. (2002)
and chemical weathering of sediment released from
the Taconic Mountains in North America (the Cenozoic
‘Himalayan model’ of Raymo and Ruddiman, 1992;
Gibbs et al., 1997; Kump et al., 1999) have all been
invoked. Such processes (if real; see Kerrick and Caldeira,
1999) could have amplified climatic cooling and snow
accumulation accompanying the generation of rift-related
dynamic topography across a polar North Africa
sufficient to briefly override high global pCO2 pressures.
7.5. Summary
Gondwana occupied a southern polar position throughout most of the Early Paleozoic but glaciation is only
evident in North Africa after 440 Ma; an older record
may have been eroded. The Saharan Ice Sheet has for
many years been the villain for paleobiologists seeking to
explain extinctions in Late Ordovician marine fauna as
the result of glacioeustatic sea level and water mass
changes. The volume of the ice sheet is constrained
traditionally by the amplitude of fluctuating sea level
recorded in distant shelf successions. Such fluctuations
vary however, even in contemporaneous strata suggesting a cause other than glacioeustasy, such as dynamic
topography. Thus the volume of the ice sheet remains
unconstrained. In North Africa, a tripartite stratigraphy of
preglacial channels, coarse glacial fill and postglacial
Silurian shales (up to 1 km thick) suggest long-term
tectonic controls, possibly accompanying rifting of
terranes from the North Gondwanan margin. A tectonic
trigger and associated uplift may explain why glaciation
occurs only briefly during a long non-glacial, interval
when northern Gondwana lay in polar latitudes for
many millions of years. Clearly, a high paleolatitude is
not by itself a sufficient requirement for glacierization
(see Ruddiman, 2001). After the demise of the Late
Ordovician Saharan glacio-epoch, Gondwana remained
at high polar latitudes for another 100 million years but
remained largely ice free until the onset of the Late
Paleozoic Gondwanan glacio-epoch at about 350 Ma
(Fig. 1).
8. Late Devonian ice c. 374 Ma
8.1. Brief cooling triggered by continental collision
Uplift related cooling along the active margin of the
South American plate spawned shortlived Late Devonian ice covers in what is now Bolivia and parts of
Brazil (Caputo, 1998; Isaacson et al., 1999; Fig. 6). The
suggested cooling role of a possible bolide impact at this
Fig. 6. Ice growth phases during the Carboniferous-Permian
Gondwanan glacio-epoch (After Crowell, 1999, Eyles, 1993). The
nucleus was Devonian ice then located over the high topography along
the active South American plate margin (Fig. 5). This expanded into
southern Africa and then eastward across India, Australia and
Antarctica concomitant with the drift of Gondwana across the south
polar latitudes and the shift in the pole from North Africa to Antarctica
and Australia by the Permian. The thickest rock record of Gondwanan
ice covers occurs in marine intracratonic basins in Brazil, Oman and
Southern Africa and in rifted margin basins of western Australia.
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time has been argued by Streel et al. (2001) and evidence from fossil lycopsids suggests lowered levels of
atmospheric CO2 (Beerling, 2002). Kaiser et al. (2006,
p. 157) refer to Late Devonian ice volumes as being
similar to that of Quaternary glaciations. A 3–4 °C
cooling across the Frasnian–Famennian boundary has
been linked to weathering-related consumption of C02
during collisional events leading to the building of
Pangea resulting in ‘one of the greatest crises of the
biosphere recorded during the Phanerozoic’ (Averbuch
et al., 2005, p. 32). Evidence for changes in water depths
on distant shelves is still related by paleobiologists to the
waxing and waning of Devonian glaciers (e.g, Saltzman,
2002; Kaiser et al., 2006) but this fails to consider the
effect of other processes on water depths such as dynamic topography as discussed above in reference to
Late Ordovician glaciation (Sections 7.3, 7.4).
9. Late Paleozoic Gondwanan glacio-epoch
(c. 350–250 Ma)
9.1. Long lived cooling triggered by continental
collision
Pujol et al. (2006) concluded that tectonic processes
drove global paleoclimates of the Late Paleozoic. This is
clearly evident after 350 Ma when large ice sheets
formed across India, South America, Southern Africa,
Australia and Antarctica (Crowell 1999; Veevers 2004).
Veevers and Powell (1987) and Powell and Veevers
(1987) showed that ice growth was a direct response to
extensive uplift at high southerly paleolatitudes during
the mid-Carboniferous Variscan and late Carboniferous
Alleghenian collisions of Gondwana with Laurasia.
These collisions coincided with large-scale thermal
doming and uplift across Pangea (Speed et al., 1997;
Veevers, 2000) in conjunction with lowered atmospheric
CO2 levels (Beerling 2002). In general, the locus of
ice covers progressively moved across Gondwana
from South America to Australia (Crowell, 1999)
tracking Gondwana's transpolar trajectory. The lack of
any earlier ice cover across this area from the Late
Ordovician through to the Carboniferous could be due to
a lack of either continental topography high enough to
promote significant cooling, or insufficient moisture in
the supercontinental interior.
Study of the deposits of the Gondwanan glacioepoch has been greatly facilitated by the presence of oil
and gas in glacially related marine rocks and hence the
availability of much subsurface data such as drill core
and downhole and seismic geophysical data (e.g., Eyles
et al., 1993, 1994; Tankard et al., 1994; Potter et al.,
1995; Stephenson et al., 2005; Berthelin et al., 2007).
Good quality age data allow identification of the timing
of deglaciation sequences in Southern Africa (Bangert
et al., 1999) with corresponding climate, floristic and
relative sea level changes in distant basins (e.g., FalconLang, 2005; Feldman et al., 2005). In general, the
landward terrestrial imprint of Late Paleozoic glaciation,
in the form of subglacially deposited tillites and glacially eroded and striated basement surfaces, is locally
prominent (e.g. Southern Africa) but is a very minor
component of the record across Gondwana as a whole
(e.g. Visser 1991, 1997; de Broekert and Eyles, 2001).
Exceptionally, a thick ice-contact succession of glaciolacustrine, glaciofluvial and eolian facies is preserved in
the intracratonic Cooper Basin of Southern Australia
(e.g., Williams et al., 1987). The bulk of ‘glacial’ deposits accumulated offshore in a wide variety of tectonic
settings (Bonorino and Eyles 1995; Lopez-Gamundi
1997). Past practice of simply drawing inferences
regarding the size of the ice covers from the geographic
extent of glacial strata across a reassembled Gondwana
is therefore, of little value because deposits accumulated beyond the landward margins of ice sheets. In
Oman, valley glaciation was triggered by uplift of
rift shoulders associated with an incipient triple junction (the India/Madagascar/Arabia rift) marking early
opening of Neotethys (Blendinger et al., 1990; Angiolini et al., 2003). In South America, marine glacial (and
some terrestrial) strata were deposited within forearc
basins along the western collisional plate margin (e.g.,
Tarija Basin; Lopez-Gamundi 1997) and within several
small (Lopez-Gamundi et al., 1992) and large intracratonic basins (e.g., Parana Basin; Eyles et al., 1993).
Glacially influenced marine and brackish water strata
accumulated in southern Africa within the retroarc
foreland Karoo Basin (Stollhofen et al., 2002; Catuneanu et al., 2002; Moore and Moore, 2004; Scheffler
et al., 2006) and in interlinked intracratonic rift basins in
central Africa (Visser 1997). Changes in relative sea
level in these successions previously interpreted in
terms of classic ‘ups and downs’ of glacioeustatically
controlled sea levels, are increasingly seen as longer
term tectonically driven changes in relative sea level
(Eyles et al., 1993; Stollhofen et al., 2002).
The thickest fills preserved on Gondwana accumulated in several rift basins clustered along the western
extensional margin of the Western Australian plate. These
contain exceptional thicknesses (2–5 km) of cold climate
hydrocarbon-hosting marine strata. ‘Postglacial glacioeustatic’ shales that cap these successions can be shown to
have accumulated diachronously along the Australian
margin in response to increases in accommodation space
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
provided by tectonic subsidence (Eyles et al., 2002, 2003,
2006). These basins eventually widened and evolved into
oceans during the break up of Gondwana (e.g., Harrowfield et al., 2005). As terranes rifted off from Gondwana
they exported deposits of pebbly glaciomarine mudstone
to what is now SE Asia (e.g., Sibumasu: China–Burma,
Malaya and Sumatra; Veevers, 2000).
9.2. Summary
A tectonic control on the onset of glaciation is particularly clear for the Gondwanan glacio-epoch where ice
growth accompanies continental collisions and the growth
of Pangea. Nonetheless, a glacial record was not preserved in many areas until the onset of extensional basin
formation and subsidence. Thus the age of the glacial
record does not necessarily fully reflect the onset of
glaciation. It is now evident after many detailed studies,
that the glacial record is overwhelmingly marine. The
most pressing question regarding the Gondwana glacioepoch is understanding the apparent absence of any ice
covers between the Late Ordovician and the Carboniferous; an almost 100 million year long period where
a significant landmass lay at the pole but remained
(apparently) ice free. This re-emphasizes that a polar
position is clearly not a sufficient condition
for glacierization. Key data are required on paleoelevations changes to assess the role of uplift in climate cooling.
10. Mesozoic non-glacial interval (c. 250 to 55 Ma);
a role for small ice masses?
Final deglaciation at the end of the Gondwanan
glacio-epoch occurred in what is now northeastern
Australia during the Late Permian (Kazanian; c. 256
Ma; Crowell, 1999; Fig. 2). This termination was
followed by a lengthy largely non-glacial episode of
more than 200 million years when no large ice sheets
formed on the planet. Possible middle to late Triassic
glaciation in Australia is contentious (Gore and Taylor,
2003). Jurassic and Cretaceous ice rafted debris is known
(Frakes et al., 1995; Chumakov and Frakes, 1997; Lurio
and Frakes, 1999) and sediments and erosional forms
attributed to terrestrial ice masses are identified in
Australia (Alley and Frakes, 2003). Any such ice masses
were small. High frequency (sub million year) fluctuations in sea level and corresponding parasequences of the
Cretaceous Western Interior Seaway of North America,
were formerly interpreted either as glacio-eustatic (Plint,
1991; Plint et al., 1992) or Milankovitch forced (Gale
et al., 2002) but are now viewed as tectonically driven
(e.g., Vakarelov et al., 2006; see also Miall, 1997). It is
109
difficult to envisage ice masses large enough to create 50
to 60 m variations in early Cretaceous sea level proposed
by Stoll and Schrag (1996, 2000) because they require a
global ice volume equivalent approximately to the
modern day ice cover over Antarctica or during the
Pleistocene in North America. Moriya et al. (2007)
concluded that there was no evidence of glaciation in
high-resolution oxygen isotope records for the midCenomanian. Crowell (1999, p. 7–8) suggested that
small Cretaceous glaciers could have formed on uplifted
rift shoulders during early Atlantic opening. If so, any
stratigraphic record now lies deeply buried at the base of
thick passive margin sequences.
11. Cenozoic glacio-epoch (< 55 Ma)
Earth began to cool after the Paleocene-Eocene
Thermal Maximum at around 55 Ma (Figs. 2 and 7)
and a range of tectonic influences are apparent on the
formation of glacial ice covers. Plate tectonic modification of continental elevation, ocean configuration and
oceanic gateways are recognised as keys to understanding
the transition from a long warm Mesozoic to a cooler
Cenozoic (e.g., Poulsen et al., 1998; Hay et al., 2002;
Prothero et al., 2003; Meijer et al., 2004).
11.1. Coupled tectono-climate models: the effects of
plate collision and dispersal
Cooling after 55 Ma is increasingly being related to
the tectonic and bathymetric evolution of the oceans as
Pangea continued to disintegrate (Scher and Martin,
2006; Via and Thomas, 2006) in combination with
shorter term, orbitally forced ocean circulation changes
(Shevenell et al., 2004). The dispersal of Pangea moved
large continental blocks (Eurasia and North America) to
higher latitudes (Turekian, 1996; Wolfe, 1978) where
climatic effects of Milankovitch variations were amplified by uplift. The northward obduction of India against
Eurasia built the glaciated Himalayas and released large
volumes of easily weathered sediment, changed upper
atmosphere flow patterns, planetary albedos (Ruddiman
et al., 1989; Kutzbach et al., 1989, 1993), monsoon
intensity (and thus mechanical and chemical weathering)
and atmospheric CO2 (Raymo and Ruddiman, 1992
though note Kerrick and Caldeira, 1999). In the
southwest continental USA, crustal extension began in
the Oligocene after 30 Ma resulting in broad scale uplift
of the Basin and Range. This affected an area more than
900 km wide and produced elevations greater than
1.5 km above sea level across the Colorado Plateau
(Parsons, 1995). Some models estimate an Earth that was
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Fig. 7. The Late Cenozoic glacio-epoch after 55 Ma. The breakup of Pangea moved large landmasses into higher latitudes, isolated Antarctica and
changed the configuration and bathymetry of ocean basins. The first ice appears at c. 40 Ma in both northern and southern hemispheres but
Milankovitch-forced Northern hemisphere ice sheets only developed after 3 Ma. After Hay et al. (2002), Scher and Martin (2006), Moran et al.
(2006), Via and Thomas (2006), Eldrett et al. (2007). Milankovitch-forced continental ice sheets in the northern hemisphere were the culmination of
some 50 million years of tectonically-influenced cooling. This may provide a model for the generation of glacio-epochs during earlier episodes of
supercontinent breakup (Fig. 1).
warmer by 1–6 °C in the absence of uplift in Central
Asia and elsewhere (Harris, 2006). Ramstein et al.
(1997) argued that the closure of the epicontinental
Paratethys was also important in determining Eurasian
climates after 30 Ma. In the Pacific southeast, the
climatic effects of Andean uplift have been recognised
(Hartley, 2003) as well as the role of glacial erosion in
promoting additional topographic (Lamb and Davis,
2003) and climatic feedbacks (Montgomery et al., 2001).
Additional cooling effects may have resulted from
widespread passive margin uplift and escarpment formation along newly rifted margins around the southern
Atlantic Ocean (e.g., Brown et al., 2000) such as reflected
in widespread periglacial and glacial deposits in southern
Africa (e.g., Lewis and Illgner, 2001) though this is
debated (Boelhouwers and Meiklejohn 2002).
11.2. Glaciation at the tops of the world: Antarctic and
Arctic cooling c. 40 Ma
It has been thought that glacierization occurred early
in the Antarctic by c. 44 Ma followed by ice growth
around the Arctic at 14 Ma (the unipolar ice sheet model
of Perlmutter and Plotnick, 2003) but recent work
continues to push back the onset of glaciation in circumArctic regions ranging from 45 Ma (Moran et al., 2006)
to c. 38 to 30 Ma (Eldrett et al., 2007).
Cenozoic glaciation in Antarctica possibly originated
along the West Antarctic Rift System which is one of the
largest areas on the planet of high standing, extended
crust comparable in size to the East African rift. Abrupt
uplift of the 3500 km long and 4 km high Transantarctic
Mountains after 95 Ma records rifting above a mantle
hot spot (ten Brink and Stern, 1992; Stump, 1995;
Hamilton et al., 2001; Studinger et al., 2002; Winberry
and Anandakrishnan, 2004). Major extension and uplift
culminated just after 50 Ma and it cannot be just coincidental that the earliest ice sheet capable of reaching
sea level (and thus recorded in the marine record) occurred at about 43 Ma before present (Lear et al., 2000).
The separation of Australia and the opening of the Drake
Passage at c. 40 Ma resulted in the thermal isolation of
Antarctica at high polar latitudes and, in conjunction
with changing ocean currents, set up the conveyor belt
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
of oceanic circulation that dominates the Cenozoic
(Prothero et al., 2003; Scher and Martin, 2006).
The onset of North Atlantic Deep Water production
at about 33 Ma (Fig. 7) was tectonically triggered by
subsidence of the Faroe–Icelandic Ridge (Via and
Thomas, 2006).
According to DeConto and Pollard (2003) Antarctic
glacierization just before 40 Ma was aided by CO2
drawdown resulting from weathering of shallow water
carbonates. Tripati et al. (2005) suggest at that time there
was a global sea level fall of 125 m as a consequence of
ice sheet growth in both hemispheres but the geologic
record does not support this model because it requires
global ice volumes similar to that of the classic Late
Cenozoic ice ages after 3.5 Ma. Hitherto, the widely
accepted model has been that significant landbased ice
did not form in the Northern Hemisphere until about
14 Ma (e.g., Cecil and Edgar, 2003). Further resolution
has until recently, been prevented by the lack of a
detailed sediment record from the Arctic basin but this
data gap has now been filled. Arctic drilling on the
Lomonosov Ridge in 2004 identifies a lonestone in
middle Eocene sediment (∼ 45 Ma) accepted as evidence of the onset of iceberg rafting in the Arctic, and
the near synchronous start of glaciation in both hemispheres (Moran et al., 2006). Interestingly, the Lomonosov Ridge is a very large rifted piece of continental
crust that was detached about 57 million years ago at the
Paleocene–Eocene boundary by seafloor spreading
along the Gakkel Ridge (Jokat et al., 1992). The overall
geodynamic setting of this part of the Arctic during
polar cooling after 55 Ma is directly analogous to that
obtaining during the onset of the Late Ordovician
Saharan glacio-epoch at 440 Ma along the north
Gondwana polar margin (Fig. 5; Section 7.4).
Significant Arctic cooling at about 45 Ma is indicated
by abundant ice rafted debris in late Eocene to Oligocene
sediments that indicates isolated calving glaciers in east
Greenland (Eldrett et al., 2007). This is at a time when
atmospheric CO2 concentrations were at least four times
present. Later evidence for circum Arctic glacier ice
occurs during the Middle Miocene Transition (Thiede
and Myhre, 1996; Helland and Holmes, 1997) (recorded
just after 16 Ma at the top of a prominent condensed
interval in Arctic ocean cores; Moran et al., 2006), which
coincides with a major expansion of the Antarctic Ice
Sheet (Shevenell et al., 2004). By 14 Ma the abundance
of Arctic Ocean ice rafted debris greatly increases,
marking the onset of glaciation in Greenland and continues until about 5 Ma when there was a marked spike in
warmth during the early Pliocene (involving an increase
of as much as 10 °C) (Ballantyne et al., 2006). Renewed
111
cooling at 3 Ma ended this warm episode as incipient ice
sheets began to wax and wane over Northern Europe and
North America as part of classic Milankovitch-driven ice
ages and interglacials.
11.3. Circum North Atlantic glaciation:
Milankovitch-driven ice sheets
The high latitude Atlantic margins are built of
thickened Proterozoic crust underplated magmatically
by the Iceland Plume (e.g., Wood et al., 1989; Eyles,
1996; Chalmers and Cloetingh, 2000; Mathieson et al.,
2000; Faleide et al., 2002; Nielsen et al., 2002; Huuse,
2002; Rohrman et al., 2002; Redfield et al., 2005). In the
cooling world after the Eocene, the increase in regional
elevation, involving wholesale uplift of plateau surrounding the North Atlantic Ocean, was likely a major
factor in promoting the growth and survival of perennial
snowfields. These ultimately would develop into
Milankovitch-modulated ice sheets after 3.5 Ma. In
Scandinavia, episodes of uplift along the Norwegian
continental margin are recorded by raised marine
planation surfaces and are precisely correlated with
offshore unconformities and influxes of glacial sediment
in the marine record (Hendriks and Andriessen, 2002;
Huuse, 2002; Hinderer and Einsele, 2002; Stoker, 2002)
thereby providing a clear and well calibrated example of
the causative climatic role of uplift along a passive
margin.
The North Atlantic borderlands share a common
Mesoproterozoic geological and geomorphological heritage. Basement rocks in eastern North America, Greenland and Scandinavia were peneplaned by about 800 Ma
BP. These shields experienced deep weathering during
the Mesozoic (the ‘etch plain’ model of LidmarBergstrom et al., 2000; Lidmar-Bergstrom and Naslund,
2002). Deep clayey regolith was later stripped by
Cenozoic uplift and reworked from the shields by ice
sheets as till. Glacial erosion created the typical ‘glaciated
shield topography’ of structurally controlled, overdeepened valleys (fiords, lakes) cut into little modified
plateau surfaces on which tropical landforms of deep
weathering (e.g., tors) survive locally (Dyke, 2004).
11.4. Pacific Northwest glaciation
The role of tectonotopography in glacierization so
well displayed along the Scandinavian passive margin is
also evident in western North America but along the very
different, compressional Gulf of Alaska margin. Pangean
rifting after 200 Ma altered the trajectory of the North
American plate creating an active plate margin what is
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now British Columbia and Alaska characterised by
terrane accretion and broad scale uplift. In the Pacific
northeast, collision of the North America plate against
the Yakutat block (an exotic terrane carried by the Pacific
plate; Mazzotti and Hyndman, 2002) promoted rapid
uplift along the Gulf of Alaska continental margin after
5 Ma. This created the highest mountain chain in North
America (Chugach-St.Elias) and triggered regional
glaciation of the far North Pacific Ocean (Haug et al.,
2005) accompanied by the release of enormous volumes
of glacially derived sediment to the Gulf of Alaska basin
where the glacially influenced Yakataga Formation is at
least 5 km thick (Lagoe et al., 1993). In contrast, the
terrestrial record is meager (Eyles and Eyles, 1989a,b).
The presence of this high topographic barrier along the
coast contributed to cooling and permafrost in the
interior of northern North America and subsequently, the
development of sea ice covers in the Arctic Ocean after
3 Ma (White et al., 1997; Westgate 2003). At about the
same time, final closure of the Straits of Panama by the
building of a Central American volcanic arc after 3 Ma
may have diverted warmth to the far northern hemisphere triggering changes in global ocean circulation
(Driscoll and Haug, 1998). Ice sheet growth has also
been linked to warm surface waters in the North Pacific
increasing the availability of moisture over the northern
interior of North America (Haug et al., 2005).
11.5. Summary
The Late Cenozoic glacio-epoch highlights the interlinking of tectonically driven changes in ocean basin
configuration and bathymetry, uplift adjacent to sources
of oceanic moisture, and orbitally driven changes in ocean
circulation. The broader geodynamic context was the
breakup and dispersal of Pangea and the movement of
large landmasses into more northern latitudes. Significant
cooling in the Arctic just before 40 Ma is broadly
coincident with the first appearance of ice in Antarctica
following its divorce from South America and Australia
and uplift along the Transantarctic Rift.
12. Discussion
This paper has reviewed geodynamic factors that
contribute to the appearance of the major glacio-epochs
over the last ∼3 Ga. Glacio-epochs appear to be a
response to complex interactions between changes in
continental positioning and elevation, paleolatitude,
ocean circulation and ventilation (e.g., Worsley and
Kidder, 1991; Worsley et al., 1994; Poulsen et al., 1998;
Crowell, 1999; Ruddiman, 2001; Rothman, 2002;
Berner, 2003; Smith and Pickering 2003; Veizer, 2005;
Figs. 1 and 2). Of all possible factors, variation in
atmospheric CO2 appears to be the most poorly
correlated with glacio-epochs, at least for the Phanerozoic (Fig. 2). The fundamental timing of glacio-epochs
at least over the last 3 Ga appears may be related to
celestial drivers such as variation in cosmic ray flux
(CRF; Shaviv, 2003) suggesting that glacio-epochs are
periodic phenomena. Other complicating factors are at
work however insofar as some CRF highs are indeed
matched by glacio-epochs whereas others are not such
as about 900 Ma and after 200 Ma (Fig. 2C). Different
processes combine to create a tipping point where climatic thresholds are crossed and ice can form over
higher, tectonically generated elevations. A major challenge remains the lack of data regarding paleoelevations, as emphasized by Hay et al. (2002), Lamb and
Davis (2003), Shuster et al. (2005), Harris (2006) and
Ghosh et al. (2006).
Being cognisant of these complexities, it is possible to
group Earth's glacio-epochs into two major tectonoclimatic types; those that occur during supercontinent
assembly and those that occur during or shortly after
supercontinent break up. This is necessarily an oversimplification of earth's tectonic history given the diachronous nature of supercontinent assembly and breakup.
Nonetheless, it identifies an important relationship
between breakup and earth's glacial record. The bulk
of Earth's glacial record is preserved within sedimentary
basins formed on extended continental crust.
12.1. Assembly related glacio-epochs
Glacio-epochs associated with plate collisions and
regional uplift associated with supercontinent assembly
appears to be rare, or at least the record of such
environments is sparsely preserved (Fig. 1). Regional
cooling across the Frasnian–Famennian boundary after
380 Ma has been linked to collisional orogenesis
(Averbuch et al., 2005) but the only glacio-epoch
initiated in this fashion is that of the Late Paleozoic
Gondwanan (325 to 250 Ma) triggered by uplift on
active margin mountains as Gondwana collided with
Laurasia (Section 8). Nonetheless, the fact remains that
the stratigraphic record of this glacio-epoch is restricted
with few exceptions, to marine extensional basins.
The general lack of ‘assembly related’ glacio-epochs
is surprising because orogenesis allows maximum exposure of terrestrial sediment to weathering processes
during long overland transport to marine basins, thereby
maximizing the drawdown of pCO2. Major glacioepochs should be associated with the Grenville Orogeny
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
at 1 Ga and others with the protracted orogenic activity
during the Mesoproterozoic but this is not the case other
than a single, short-lived event in Australia (Williams,
2005; Section 5). Averbuch et al. (2005) argued that
cooler but non-glacial phases in Earth history are linked
with enhanced mountain building and enhanced chemical erosion of continental crust such. In contrast, such
effects do not appear at first blush, to have had a major
effect on global cooling over the past 3 Ga (Fig. 1).
Otherwise, glacio-epochs would be more commonly
identified with phases of supercontinent assembly and
mountain building particularly during the Mesoproterozoic and Neoproterozoic. This absence supports
Kerrick and Caldeira's (1999) thesis that the effects of
tectonically driven CO2 drawdown are overstated
because the process is temperature dependent and its
efficacy is greatly reduced during cooling. The apparent
rarity of active margin glacio-epochs may also reflect
the low preservation potential of glacial sediments in
such settings. For example, Cenozoic glaciation of major
mountain ranges created within continental interiors by
compressional plate forces (such as the Himalayas) has
left no primary glacial sediment (even ice-rafted debris)
in the offshore sediment record (e.g., in the Indian
Ocean). This is because glacial sediment is stored in
floodplains or entirely reworked by rivers during long
overland transport to the marine realm. The deposits of
glaciated backarc basins along compressional margins
are similarly underrepresented in the rock record. On the
other hand, the longest and possibly most complete
record of late Cenozoic glaciation anywhere on Earth is
preserved in the 5 k thick Yakataga Formation of the Gulf
of Alaska compressional margin (Lagoe et al., 1993;
Eyles, 1993; Section 11.4). There, vast volumes of
glaciogenic sediment released from rapidly uplifting
coastal mountains and glaciers have been fed directly
into deepwater submarine fans systems. The equivalent
terrestrial glacial record inland is no more than a few
hundred metres thick and only locally preserved (Eyles
and Eyles, 1989a,b). Similarly, dramatic increases in
sedimentation rates that accompanied glaciation at the
same time along the Chilean convergent margin, choked
offshore forearc basins with thick successions of
glacially influenced marine strata (Melnick and Echter,
2006). The long-term fate of these thick successions is to
be incorporated within the continental margin as highly
deformed accretionary slivers.
12.2. Break up related glacio-epochs
Most glacio-epochs have occurred in the geodynamic
context of rifting and crustal extension (Figs. 1 and 2).
113
Paleoproterozoic (c. 2.4 Ga) and Neoproterozoic (c.
750-580 Ma) glacio-epochs are ‘break up related’
formed during the breakup of Kenorland and Rodinia
respectively (Fig. 1). It is also possible that extension
along high latitude continental margins and consequent
uplift also played a role in triggering Ordovician glaciation at c. 440 Ma (when terranes rifted off Gondwana;
Section 7.4) also Cenozoic ice covers after 55 Ma
around the Arctic (rifting of the Lomonosov Ridge?)
and in the Antarctic along the Transantarctic Rift. The
protracted disintegration of Pangea still continues in
Africa to the present day and glaciers occur along the
uplifted but low latitude East African Rift system.
Given the recurring association of glacio-epochs with
continental break up throughout Earth history (Fig. 1) it
appears that there is an essential continuity of tectonic
boundary conditions for Proterozoic and Phanerozoic
glacio-epochs. This begs the question as to the nature of
possible causal links between rifting and climate change,
or simply whether glacial sediments have been selectively preserved in this geodynamic setting.
12.3. Rifting and glaciation: causal link or simply
preservational bias?
Allen (2007a) provides a detailed basin analysis of
Neoproterozoic glacial deposits in Oman (Huqf Supergroup), that supports the central thesis of this paper of a
preferred relationship of glacial successions with rift
basins. Allen (2007a) suggested this was clear evidence
of a preservational bias rather than of causation between
rift topography and cooling. Intracratonic extension is
associated with significant regional uplift and thus
cooling. Rift flank uplift is the consequence of flexural
rebound at the edges of rifted plates combined with
warming and thermal uplift arising from the upward
movement of the lithosphere-asthenosphere boundary
below the rift (e.g., Garfunkel, 1988; Bott, 1995; Van
Der Beek, 1997). The magnitude of flexural arching
varies however and depends among other things, on the
thickness and rigidity of the rifted plate (Braun and
Beaumont, 1989; Weissel and Karner, 1989; King and
Ellis, 1990; Kooi and Beaumont, 1994; Fig. 8). Rifting
can generate a dynamic topography of up to 1 km in
elevation but is a transient condition followed by postrift thermal subsidence compounded by sediment
loading. A lithospheric plate with a high effective elastic thickness (Te) characteristic of old thick
continental crust produces a rift flank uplift that is
high and broad (Fig. 8) sufficient to cross a climatic
threshold allowing snow and ice to accumulate. Extension of cold and thick (e.g., 30–60 km) continental
114
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
Fig. 8. Geometry of continental extension (after Ebinger et al., 2002) as occurred during the Paleoproterozic and Neoproterozoic glacio-epochs
(Fig. 1). The most extensive uplifts are created where crust is old, thick and thus flexurally rigid (creating a high effective elastic plate thickness:
EEPT). The resulting isostatic rebound and uplift is distributed across a large area. This was the case during the break up of Rodinia after 750 Ma that
resulted in the breaking of thick Paleoproterozoic continental crust that has amalgamated mainly after 2.5 Ga (Fig. 1). Additionally, mantle convection
below the broken margin, and magmatic underplating contributes to uplift of passive margins and was instrumental in creating uplifted plateau on
which late Cenozoic ice sheets grew on the margins of the North Atlantic Ocean after 3 Ma (e.g., Laurentide, European ice sheets).
lithosphere such as in East Africa and on the margins of
the Red Sea, has created a dynamic topography 1200 m
above sea level across a zone more than 1000 km wide
(e.g., Khalil and McClay, 2001; Ebinger et al., 2002;
Morley, 2002). Partridge (1997) emphasized the climatic effects of uplift along the rift flanks. The magnitude of uplift in Kenya approaches 2 km (Morley,
2002) with tropical conditions along the rift floor and
glaciated volcanic peaks of mounts Kilimanjaro and
Kenya rising above.
The breakup of Rodinia was likely marked by enhanced dynamic topography where rifts developed in old
thick Archean and Proterozoic crust such as in Australia,
NW Canada and in northern Europe. It is worth emphasizing that there can be a significant delay between
rifting, uplift and any climate response. Illies (1977) and
Garfunkel and Bartov (1977) showed that rift flank uplift
commenced as much as 15 million years after rifting is
initiated. The delay between rifting and climatic cooling
increases to several tens of million years in the case of
Cenozoic glaciations that developed around the North
Atlantic borderlands long after Atlantic Ocean rifting
was initiated (Eyles, 1993; Nielsen et al., 2002). Similarly, the diachronous nature of rifting along propagating
spreading centers (e.g., Huchon et al., 2001; Morley,
2002) means that rift flank uplifts and corresponding
regional changes in climate and a glacial influence on
sedimentation, will be time transgressive.
Other aspects of continental breakup may also favour
glaciation. A rift creates several elevated zones flanking
sources of moisture (an incipient ocean; Fig. 7 and 9),
which is the key requirement for building ice sheets, not
cold temperatures. This aspect was emphasized by Hay et
al. (2002) in an excellent review of the links between
tectonics and climate in the Late Cenozoic. In addition,
rifting reverses plate motions creating active plate margins
and elongate areas of glaciated mountainous topography
on the outboard sides of the newly rifted plates. Above all,
very short transport distances between uplifted source
areas and rapidly subsiding basins (e.g., Fig. 3) minimizes
sediment reworking promoting preservation of glacially
influenced (mostly marine) facies. Uplifted source areas
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
115
Fig. 9. Simplified diagram illustrating principal differences between glacio epochs resulting from uplift resulting from continental collision (A1 and
A2) and resulting from continental extension (B1, B2). In the first case, uplift occurs along a narrow intracratonic mountain belt in the dry interior of
the continent. These glaciers leave only a minor terrestrial record of glaciation that is easily destroyed by large rivers. This contrasts with the much
larger ice covers that grow on broad uplifted rift shoulders close to moisture sources provided by opening oceans. Ice lobes reach sea level and leave a
record of glacially influenced marine strata in rift basins. Rifting is markedly diachronous resulting in a non-synchronous relationship between
glaciation, the formation of fault bounded sub basins and sedimentation along growing rifts (Fig. 3). Additional tectonotopography is created by
terrane accretion along the now active margins of the two new plates.
are separated from basins only by steep slopes along the
basin margin (Withjack et al., 2002). These slopes are
rapidly incised by glacial erosion allowing direct delivery
of sediment to the basin from upland source areas.
Virtually the entire sediment flux in extensional settings is
produced close to basin depocentres and is rapidly buried
with minimal exposure to fluvial reworking and terrestrial
weathering with limited potential to influence atmospheric composition. In this way, Earth's glacial record may be
selectively associated with rifting not just through the
effects of uplift-generated cooling and the availability of
moisture, but also because of a preservational bias (see
Fig. 9). Allen (2007b) has suggested that glacial deposits
are selectively preserved in rifts at times when global
climate was cooled by mountain building in collisional
tectonic settings. As we have identified above, this was
certainly the case in the long Permo-Carboniferous glacioepoch that affected Gondwana after 350 Ma (Section 9.1).
The glacio-epoch was triggered by collision but its
depositional record is found in later rift basin fills. This
model however, does not fit well with the preferred
relationship of glacio-epochs with times of supercontinent
breakup (Fig. 1) where global tectonics were dominated
by extension not collisional processes.
13. Concluding remarks
A review of Earth's glacial record over the last 3 Ga
indicates a close relationship between glacio-epochs and
times of enhanced crustal extension during the Proterozoic and Phanerozoic; most of Earth's glacial record appears to be preserved in extensional basins.
Tectonically generated topography produced by crustal
extension may be an important control on cooling in
conjunction with increased availability of moisture.
Clearly there are times in Earth history of rifting with no
ice, and ice with no rifting but the marked association
between the two for most ancient glacio-epochs cannot
be simply coincidental. This association could simply
reflect better preservation of glacial strata in extensional
basins and the tendency for glacial deposits to be
entirely reworked by large rivers in compressional
settings. Having recognised the importance of tectonic
preconditions under which glacio-epochs develop and
glacial deposits are preserved, detailed consideration
of the role of tectonics in influencing climate and
controlling water depths, sediment supply and the age of
sedimentary successions, is essential in future basin
investigations and climate models.
116
N. Eyles / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 89–129
Acknowledgements
The Natural Sciences and Engineering Research
Council of Canada is thanked for generous research
support. Several colleagues provided informative correspondence and discussions principally Paul Ramaekers, Carolyn Eyles, Jan Veizer, Louise Daurio, Angie
Falcon, Nicole Januszczak, Vic Gostin, Andrew Miall,
Kath Grey, Phil Allen, Paul Gammon, Jeff Lewis, Dick
Peltier, Noel James, John Crowell and Tony Tankard.
Any errors of interpretation or omission remain entirely
mine. Two anonymous journal reviewers and Finn
Surlyk are thanked for their very helpful comments.
This is a contribution to IGCP 512 ʻNeoproterozoic
Glaciations'.
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