Original research/Glaciations

The ice ages or glaciations on Earth occurred from the early Proterozoic (Huronian), late proterozoic (Cryogenian), early Paleozoic (Andean-Saharan) during the Ordovician and Silurian periods, late Paleozoic (Karoo Ice Age) during the Carboniferous and early Permian periods, and lately the Quaternary glaciation.

Geologic time is annotated with glacial or ice age periods. Credit: William M. Connolley.
Earth at the last glacial maximum of the current ice age. Credit: Ittiz, based on: "Ice age terrestrial carbon changes revisited" by Thomas J. Crowley (Global Biogeochemical Cycles, Vol. 9, 1995, pp. 377-389.
Recent (black) and maximum (grey) glaciation of the northern hemisphere are during the Quaternary climatic cycles. Credit: Hannes Grobe/AWI.
Recent (black) and maximum (grey) glaciation of the southern hemisphere are during the Quaternary climatic cycles. Credit: Hannes Grobe/AWI.

Although these ice ages are widely separated in geological time, "in most parts of the Earth major climatic and palaeoenvironmental units typically have a duration of the order of half a precession cycle (around 10 ka) rather than half an eccentricity cycle (around 50 ka) so that the level of stratigraphic resolution provided by the Middle Pleistocene [Marine Isotope Stage] MIS (typical duration 50 ka) is not sufficiently fine to constitute a universal stratigraphic template."[1]

Theoretical glaciationsEdit

There are few outdoor pleasures to match negotiating a rough bouldery scree slope thanks to periglaciation processes. Credit: Eric Jones.

Def. the "process of covering with a glacier,[2] or the state of being glaciated;[3] the production of glacial phenomena;[3] an ice age[4]" is called a glaciation.

Def. the geomorphic processes resulting from "the unstable conditions following the retreat of a glacier"[5] is called paraglaciation.

Def. the "geomorphic processes resulting from seasonal thawing of snow in areas of permafrost, the runoff from which refreezes in ice wedges and other structures"[6] are called periglaciation.

Def. the geomorphic processes resulting from being "situated in front of a glacier or ice sheet (and formed from its meltwaters)"[7] is called proglaciation.

The "term "periglacial" is a function of process, "proglacial" is a function of location and "paraglacial" is a function of degree and mode of recovery from the disturbance of continental glaciation. Periglacial and proglacial environments are commonly viewed as being adjusted to contemporary process, though important questions have been raised about relict periglacial landscapes in this regard. Paraglacial environments are explicitly out of adjustment with contemporary process and retain in their configuration a glacial signature. All three concepts are seen to be essential to comprehensive understanding of glaciated environments. It is a nested set of concepts which overlap in the field but none of the terms is redundant."[8]

Glacial timesEdit

Ice ages Glaciations Stadials Interstadials
Name Timeline Name Timeline Name Timeline Name Timeline
Holarctic-Antarctic, Late Cenozoic 33.9 million years ago Late Weichselian, Devensian, MIS 5d-2 c. 32,000 years ago
Younger Dryas 12,900 a Meiendorf, Flandrian, Holocene, MIS 1 11,660 ± 40 a to present
Altonian c. 13.2 ka Heinrich Event H1 13 ka
Older Dryas 14,000 a Allerød oscillation 13,400 a
Oldest Dryas, Dolní Veštonice cold event 6-7, Altonian 15,700 Bølling oscillation 14,700 a
Jylland c. 22 ka Lascaux 17 ka
Letzteiszeitliches c. 26 ka Laugerie about 23.2 ka
Heinrich Event 3 c. 26.74 ka Dansgaard-Oeschger event 3 29 ka and ends about 26 ka
Klintholm c. 26 ka Møn, Dansgaard-Oeschger event 4 29 ka and ends about 26 ka
Stadial c. 31 ka GIS 5 30.013 to 29.526 ka
Stadial c. 32 ka Ålesund GIS 6 (start) 31.218 [to] GIS 6 (end) 30.849 ka
Middle Weichselian c. 80 kyr BP to c. 32 kyr BP
Stadial c. 33 ka GIS 7 GIS 7 (start) 32.896 [to] GIS 7 (end) 32.15 ka
Heinrich Event 4 33-39.93 ka Huneborg, Denekamp 36.5-38.5 ka
Hasselo 40-38.5 ka Hengelo 38-36.5 ka
Marine Isotope Stage 3 45.76 and 47.16 ka Moershoofd 44-46 ka
Ebersdorf 50 ka Glinde 48-50 ka
Karmøy, Schalkholz 60 ka Oerel 56-59 ka
Marine Isotope Stage 4 about 57-71 ka Odderade 61-72 ka
early Pleniglacial ? ka Ognon III ? ka
Wisconsinian 80 ka Ognon II 61-72 ka
Early Weichselian c. 119 kyr BP to c. 80 kyr BP
Rederstall 96 ka Interstadial 90 ka
Herning 123 ka Brørup 100 ka
Saale, MIS -6 347 to 128 ka
Penultimate 194 to 135 ka Eemian, Ipswichian, Sangamonian, Riss-Würm, MIS 5e 131 to 117 ka
Illinoian, Wolstonian, Saalian, Riss 220-430 ka Holstein, Sangamon, MIS 9? 220 to 80 ka
Pre-Illinoian, Haweran 220-430, 340 ka Marine Isotope Stage 9 (MIS 9) 337,000 to 300,000 years ago ka
Riss, Saalian, Illinoian, Wolstonian, MIS 8-6 374 until 130 ka Riss-Würm
Mindel, MIS -6 424-374 ka
Mindel 424-374 ka Mindel-Riss, Holsteinian, Hoxnian, Yarmouthian, MIS 11
Elster 478-424 ka
Anglian, MIS 12 478-424 ka Yarmouthian 420 to 500 ka
Kansan 500-600 ka
MIS Boundary 13/14 is at 533 ka 500-600 ka Yarmouthian 420 to 500 ka
Anglian, Mindel, Elsterian, MIS Boundary 14/15 is at 563 ka 500-600 ka Yarmouthian 420 to 500 ka
Nebraskan 650-1000 ka
Nebraskan 650-1000 ka Aftonian 600 to 650 ka
Cromerian, MIS 21-13? 866-478 ka interstadial
Beestonian -866 ka interstadial
Danube ~1.8 Ma to 1.0 Ma
Würm, MIS 5d-2 1.806 Ma Riss-Würm, MIS 5e
Pastonian 1.806 Ma Riss-Würm, MIS 5e
Riss, MIS 8-6 1.806 Ma Mindel-Riss, MIS 9
Baventian, Mindel 1.806 Ma Mindel-Riss, MIS 9
Bramertonian, Günz 1.816 and 2.427 Ma Haslach-Mindel
Danube ~1.8 Ma to 1.0 Ma Danube-Günz, Uhlenberg
Biber ~1.8 Ma to 1.0 Ma
Biber ~2.5 Ma to 2.0 Ma Biber-Danube
Menapian ~2.6 Ma to 11.70 ka
Beestonian ~2.6 Ma to 11.70 ka Cromerian, Aftonian
Eburonian, Günz, Beestonian, Nebraskan 1.78 Ma to 480 ka Waal
Menapian, Praetiglian, Gelasian ~2.58 Ma to 11.70 ka Tegelen
Piacenzian 3.6 Ma to 2.58 Ma
Piacenzian 104 2.62 Ma to 2.58 Ma Galasian 103 2.58 Ma to 2.57 Ma
G2 2.66 Ma to 2.63 Ma G1 2.63 Ma to 2.62 Ma
G4 2.70 Ma to 2.68 Ma G3 2.68 Ma to 2.66 Ma
Calvario Drift, G6 2.74 Ma to 2.71 Ma G5 2.71
G8 2.78 Ma to 2.77 Ma G7, Unit 8 2.77 Ma to 2.74 Ma
G10, Unit 7 2.83 Ma to 2.79 Ma G9 2.79 Ma to 2.78 Ma
G12, Unit 6 2.86 Ma to 2.84 Ma G11 2.84 Ma to 2.83 Ma
G14 2.90 Ma to 2.88 Ma G13 2.88 Ma to 2.86 Ma
G16 2.974 Ma to 2.972 Ma G15 2.972 Ma to 2.90 Ma
G18 2.98 Ma to 2.975 Ma G17 2.975 Ma to 2.974 Ma
G20 3.025 Ma to 3.00 Ma G19 3.00 Ma to 2.8 Ma
Mid-Piacenzian Warm Period 3.6 Ma to 2.58 Ma
G22, Unit 4 3.05 Ma to 3.045 Ma G21 3.045 Ma to 3.025 Ma
cooling K1 3.08 Ma to 3.07 Ma K1 3.07 Ma to 3.05 Ma
K2 3.1 Ma to 3.09 Ma K1.5 3.09 Ma to 3.08 Ma
KM2, Unit 3c 3.11 Ma to 3.11 Ma K3 3.11 Ma to 3.1 Ma
KM2, Unit 3b 3.15 Ma to 3.127 Ma K5 3.127 Ma to 3.11 Ma
cooling 3.19 Ma to 3.18 Ma K7 3.18 Ma to 3.15 Ma
cooling 3.20 Ma to 3.19 Ma K9 3.19 Ma to 3.19 Ma
stadial 3.211 Ma to 3.210 Ma K11 3.21 Ma to 3.20 Ma
KM6 3.23 Ma to 3.211 Ma K13 3.211 Ma to 3.211 Ma
M2, Unit 3a 3.31 Ma to 3.265 Ma M1 3.265 Ma to 3.23 Ma
MG2, Unit 2 3.34 Ma to 3.34 Ma MG1 3.34 Ma to 3.31 Ma
MG4, Unit 1 3.375 Ma to 3.37 Ma MG3 3.37 Ma to 3.34 Ma
MG6 3.47 Ma to 3.47 Ma MG5 3.47 Ma to 3.375 Ma
MG8 3.53 Ma to 3.52 Ma MG7 3.52 Ma to 3.47 Ma
MG10 3.57 Ma to 3.55 Ma MG9 3.55 Ma to 3.53 Ma
MG12 3.59 Ma to 2.58 Ma MG11 3.58 Ma to 3.57 Ma
Zanclean 5.333 Ma to 3.6 Ma
Zanclean 5.333 Ma to 3.6 Ma interstadial
Messinian 7.246 Ma to 5.333 Ma
Messinian 7.246 Ma to 5.333 Ma interstadial
Tortonian 11.63 Ma to 7.246 Ma
Tortonian 11.63 Ma to 7.246 Ma interstadial
Serravallian 13.82 Ma to 11.63 Ma
Serravallian 13.82 Ma to 11.63 Ma interstadial
Langhian 15.97 Ma to 13.82 Ma
Langhian 15.97 Ma to 13.82 Ma interstadial
Burdigalian 20.44 Ma to 15.97 Ma
Burdigalian 20.44 Ma to 15.97 Ma interstadial
Aquitanian 23.03 Ma to 20.44 Ma
Aquitanian 23.03 Ma to 20.44 Ma interstadial
Chattian 27.82 Ma to 23.03 Ma
Chattian 27.82 Ma to 23.03 Ma interstadial
Rupelian 33.9 Ma to 27.82 Ma
Rupelian 33.9 Ma to 27.82 Ma interstadial
Karoo, Late Paleozoic 360 to 255 Mya
Andean-Saharan 450 to 420 Ma
Varanger 610 to 575 Ma Gaskiers 579.63 ± 0.15 to 579.88 ± 0.44 Ma
Elatina c. 640 and 580 Ma
Ice Brook c. 659 and 651 Ma
Cryogenian 720 to 635 Mya
Vendian, Sturtian c. 740 Ma to 660 Ma
Beiyixi 755 Ma
Makganyene 2.32 Ga to 2.22 Ga
Huronian 2.5 Gya to 2.2 Gya
Pongola 3.49 Gya to 2.78 Gya

Holarctic-Antarctic Ice AgeEdit

Global average temperature estimates are for the last 540 My. Credit: Glen Fergus.{{free media}}
Map is of the Northern Hemisphere ice during the last glacial maximum. Credit: Hannes Grobe/AWI.{{free media}}

"This late Cenozoic ice age began at least 30 million years ago in Antarctica; it expanded to Arctic regions of southern Alaska, Greenland, Iceland, and Svalbard between 10 and 3 million years ago. Glaciers and ice sheets in these areas have been relatively stable, more-or-less permanent features during the past few million years."[9]

The Late Cenozoic Ice Age,[10][11] or Antarctic Glaciation[12][13] began 33.9 million years ago at the Eocene–Oligocene extinction event and is ongoing.[10] It is Earth's current ice age or Greenhouse and icehouse Earth (icehouse period). Its beginning is marked by the formation of the Antarctic ice sheets.[14] The Late Cenozoic Ice Age gets its name due to the fact that it covers roughly the last half of Cenozoic era so far.

Six million years after the start of the Late Cenozoic Ice Age, the East Antarctic Ice Sheet had formed, and 14 million years ago it had reached its current extent. It has persisted to the current time.[15]

In the last three million years, glaciations have spread to the northern hemisphere. It commenced with Greenland becoming increasingly covered by an ice sheet in late Pliocene (2.9-2.58 Ma ago)[16] During the Pleistocene (starting 2.58 Ma ago), the Quaternary glaciation developed with decreasing mean temperatures and increasing amplitudes between glacials and interglacials. During the glacial periods of the Pleistocene, large areas of northern North America and northern Eurasia have been covered by ice sheets.

The concept that the earth is currently in an ice age that began around 30 million years ago can be dated back to at least 1966.[17]

As a geologic time period, the Late Cenozoic Ice Age was used at least as early as 1973.[18]

The hottest part of this torrid age was the Paleocene-Eocene Thermal Maximum, 55.5 million years ago. Average global temperatures were around 30 °C (86 °F).[19]

During the early Eocene, Australia[20] and South America[21] were connected to Antarctica.

Australia drifted away from Antarctica forming the Tasmanian Passage, and South America drifted away from Antarctica forming the Drake Passage. This caused the formation of the Antarctic Circumpolar Current, a current of cold water surrounding Antarctica.[15] This current still exists today, and is a major reason for why Antarctica has such an exceptionally cold climate.[20]

The Eocene-Oligocene Boundary 33.9 million years ago was the transition from the last greenhouse period to the present icehouse climate.[22][15] At this point CO2 levels had dropped to 750 ppm.[23] This was the beginning of the Late Cenozoic Ice Age. This was when the ice sheets reached the ocean,[24] the defining point.[25]

At 29.2 million years ago, there were three ice caps in the high elevations of Antarctica.[15] One ice cap formed in the Dronning Maud Land.[15] Another ice cap formed in the Gamburtsev Mountain Range.[15] Another ice cap formed in the Transantarctic Mountains.[15] At this point, the ice caps weren't very big yet.[15] Most of Antarctica wasn't covered by ice.[15]

By 28.7 million years ago, the Gamburtsev ice cap was now much larger due to the colder climate.[15]

CO2 continued to fall and the climate continued to get colder.[15] At 28.1 million years ago, the Gamburtsev and Transantarctic ice caps merged into a main central ice cap.[15] At this point, ice was now covering a majority of the continent.[15]

The Dronning Maud ice cap merged with the main ice cap 27.9 million years ago.[15] This was the formation of the East Antarctic Ice Sheet.[15]

Global refrigeration set in 22 million years ago.[14]

By 14 million years ago, the Antarctic ice sheets were similar in size and volume to present times.[10] Glaciers were starting to form in the mountains of the Northern Hemisphere.[10]

Between 3.6 and 3.4 million years ago, there was a sudden but brief warming period.[10]

The glaciation of the Arctic in the Northern Hemisphere commenced with Greenland becoming increasingly covered by an ice sheet in late Pliocene (2.9-2.58 Ma ago).[16]

Earth is currently tilted at 23.5 degrees. Over a 41,000 year cycle, the tilt oscillates between 22.1 and 24.5 degrees.[26] When the tilt is greater (high obliquity), the seasons are more extreme. During times when the tilt is less (low obliquity), the seasons are less extreme. Less tilt also means that the polar regions receive less light from the sun. This causes a colder global climate as ice sheets start to build up.[26]

The shape of Earth's orbit around the sun affects the Earth's climate. Over a 100,000 year cycle, Earth oscillates between having a circular orbit to having a more elliptical orbit.[26]

From 2.58 million years ago to about 1.73 million ± 50,000 years ago, the degree of axial tilt was the main cause of glacial and interglacial periods.[26]

Around 850,000 ± 50,000 years ago, the degree of orbital eccentricity became the main driver of glacial and interglacial periods rather than the degree of tilt, and this pattern continues to present-day.[26]

According to Blue Marble 3000 (a video by the Zurich University of Applied Sciences), the average global temperature around 19,000 BCE (about 21,000 years ago) was 9.0 °C (48.2 °F).[27] This is about 4.8 °C (8.6 °F) colder than the 1850-1929 average, and 6.0 °C (10.8 °F) colder than the 2011-2020 average.

The figures given by the Intergovernmental Panel On Climate Change (IPCC) estimate a slightly lower global temperature than the figures given by the Zurich University of Applied Sciences. However, these figures are not exact figures and are open more to interpretation. According to the IPCC, average global temperatures increased by 5.5 ± 1.5 °C (9.9 ± 2.7 °F) since the last glacial maximum, and the rate of warming was about 10 times slower than that of the 20th Century.[28] It appears that they are defining the present as the early period of instrumental records when temperatures were less affected by human activity, but they do not specify exact years, or give a temperature for the present.

Berkeley Earth puts out a list of average global temperatures by year. It shows that temperatures were stable between the beginning of records in 1850 all the way through 1929. If you average all of the years from 1850 to 1929, the average temperature comes out to 13.793 °C (56.828 °F).[29] When subtracting 5.5 ± 1.5 °C (9.9 ± 2.7 °F) from the 1850-1929 average, the average temperature for the last glacial maximum comes out to 8.3 ± 1.5 °C (46.9 ± 2.7 °F). This is about 6.7 ± 1.5 °C (12.0 ± 2.7 °F) colder than the 2011-2020 average. This figure is open to interpretation because the IPCC does not specify 1850-1829 as being the present, or give any exact set of years as being the present. It also does not state whether or not they agree with the figures given by Berkeley Earth.

According to the United States Geographical Survey (USGS), permanent summer ice covered about 8% of Earth's surface and 25% of the land area during the last glacial maximum.[30] The USGS also states that sea level was about 125 m (410 ft) lower than in present times (2012).[30] The volume of ice on Earth was around 17,000,000 mi3 (71,000,000 km3),[31] which is about 2.1 times Earth's current volume of ice.

"During the last one million years, large ice sheets developed in North America, Eurasia, the Andes, and elsewhere. These ice masses were unstable, growing and self-destructing in cycles averaging about 100,000 years, which correspond to eccentricity in the Earth's orbit around the Sun (Mangerud et al. 1996). The most recent great ice sheets disappeared only 10,000 years ago, but the Holarctic-Antarctic Ice Age still continues in regions of stable glaciation."[9]


Periglacial calcareous patterned ground is shown. Credit: P.Cikovac.
Calcareous scree (block field) is along the Jastrebica leg. Credit: P.Cikovac.

In the periglacial (sub) arctic areas limestone scree slopes occur in a circumpolar manner in completely flat areas. The freezing and thawing processes lead to material sorting and frost pattern soil formation. In the Alps, on the other hand, floating soils can form over limestone soils.[32] In alpine high-altitude zones, calm limestone scree slopes always form habitats with an emphasis on snow, which in primary succession have limestone-snow-soil profiles.

Periglacial patterned groundsEdit

This periglacial patterned ground plateau to the west of Mynydd Pen Cyrn is scattered with gritstone boulders arranged in polygons. Credit: Duncan Hawley.
Frost heaved boulders of the shepard rock formation are in Glacier National Park, Montana, USA. Credit: P. Carrara, National Park Service (NPS).

The image on the right is a relict feature of arctic conditions which prevailed in this location at the end of the last glacial period (about 10,000 years ago), and was caused by frost shattering of the underlying rock and subsequent upheaving under cycles of ground freeze/thaw. The technical term for the arctic conditions is periglacial and the patterned ground is a periglacial feature.

On the left is an image of frost heaved boulders of the shepard rock formation in Glacier National Park, Montana, USA.


This is a higher altitude boulder field in Iceland. Credit: Seongjink.

In high-altitude arctic and subarctic regions, scree slopes and talus deposits are typically adjacent to hills and river valleys and usually originate from late-Pleistocene periglacial processes.[33]

Rock glaciersEdit

Aconcagua mountain, aerial view of Glaciar Horcones Inferior, ARG, shows the rock glacier. Credit: Mariordo, Mario Roberto Duran Ortiz.
The rock glacier is descending from Sourdough Peak, Alaska. Credit: NPS Natural Resources.
This image shows that the Sourdough Peak rock glacier (Wrangell-St. Elias National Park, Alaska) looks much more like a glacier when snow covers its rocky surface, coloring it an icy white in the wintertime. Credit: NPS Photo/B. Petrtyl.

With the exception of ice-cored rock glaciers, rock glaciers are a periglacial process. Periglacial rock glaciers require permafrost instead of glacial ice in order to form and are caused by continuous freezing occurring within a talus lobe.[34]

The rock glacier imaged on the right is most likely a glacial rock glacier. But the one shown in the summer and winter images on the left may be a periglacial rock glacier.

Little Ice AgesEdit

Changes in the 14C record, which are primarily (but not exclusively) caused by changes in solar activity, are graphed over time. Credit: Leland McInnes.{{free media}}

The Little Ice Age (LIA) appears to have lasted from about 1218 (782 b2k) to about 1878 (122 b2k).

A "climate interpretation was supported by very low δ’s in the 1690’es, a period described as extremely cold in the Icelandic annals. In 1695 Iceland was completely surrounded by sea ice, and according to other sources the sea ice reached half way to the Faeroe Islands."[35]

In the image at the top, "before present" is used in the context of radiocarbon dating, where the "present" has been fixed at 1950. The apparent decreases in solar activity are called the "Maunder Minimum", "Spörer Minimum", "Wolf Minimum", and "Oort Minimum".

"Northern Hemisphere summer temperatures over the past 8000 years have been paced by the slow decrease in summer insolation resulting from the precession of the equinoxes."[36]

Precisely "dated records of ice-cap growth from Arctic Canada and Iceland [show] that LIA summer cold and ice growth began abruptly between 1275 and 1300 AD, followed by a substantial intensification 1430-1455 AD. Intervals of sudden ice growth coincide with two of the most volcanically perturbed half centuries of the past millennium. [Explosive] volcanism produces abrupt summer cooling at these times, and that cold summers can be maintained by sea-ice/ocean feedbacks long after volcanic aerosols are removed. [The] onset of the LIA can be linked to an unusual 50-year-long episode with four large sulfur-rich explosive eruptions, each with global sulfate loading >60 Tg. The persistence of cold summers is best explained by consequent sea-ice/ocean feedbacks during a hemispheric summer insolation minimum; large changes in solar irradiance are not required."[36]

Weichselian glaciationsEdit

"Recent stratigraphical achievements and long time established chronologies exist for the Late Weichselian, i.e. 10-25 ka BP. During this period Denmark experienced the complex Main-Weichselian glaciation from 25 to about 14 ka BP (Jylland stade, Houmark-Nielsen 1989) followed by the Late Glacial climatic amelioration including the interstadial Bølling-Allerød oscillation (13-11 ka BP), finally leading to the interglacial conditions that characterize the Holocene (Hansen 1965)."[37]

"The Weichselian of Europe covers the interval from the end of MIS-5e (c. 119 kyr BP) to the start of the Holocene at c. 11.5 kyr BP and corresponds to the isotope stages 5d to 2. Much of the Weichselian chronology is relative only, determined by stratigraphic relationships of successive glacial and interglacial deposits. Radiocarbon ages for the younger interval and Thermo-Luminescence (TL), Optically Stimulated Luminescence (OSL) and Electron Spin Resonance (ESR) dates for the earlier period are used where available although reliable results remain few. We assume here, therefore, that the Russian􏰀European succession of major glacials and interglacials follows the oscillations of the global sea-level curve of Lambeck & Chappell (2001). The start of stadials is defined by the onset of a sea-level fall and the end is defined by the midpoint between successive lowstands and highstands, in recognition of the lag in icesheet and sea-level response to warming. The Early Weichselian spans the interval from c. 118 kyr BP to c. 80 kyr BP and corresponds to the two stadials MIS-5d and 5b and the two interstadials MIS-5c and 5a. The Middle Weichselian corresponds to the isotope stages MIS-4 and MIS-3 spanning the interval from c. 80 kyr BP to c. 32 kyr BP [...]. As more information becomes available, the interstadials 5c and 5a reveal a more complex structure and each may consist of two or three relative highstands (Potter et al. 2004), implying that ice margins were not constant during these intervals, [...]."[38]

"The early Middle Weichselian is assumed to correspond to the period 80-􏰀62kyr BP and to MIS-4. During this interval, average sea levels reached lower values than during the Early Weichselian and ice extent can be expected to have been substantial. But, as the global sea-level oscillations in this interval are also large, substantial ice-volume fluctuations can be anticipated across northern Eurasia within this stage."[38]

"The last substantial ice movement over arctic Russia is the retreat at the end of MIS-4 back to the Kara Sea and eventually back to the arctic islands such that after c. 55 kyr the major land areas were and remained essentially ice-free. The Scandinavian ice sheet, however, continued to fluctuate throughout Stage 3, with at least two periods of extensive ice-free conditions corresponding to the Bø interstadial (at c. 52 kyr BP) when the ice retreated to northern Sweden, and the Ålesund interstadial (at c. 35 kyr), when much of Scandinavia may have been ice-free. At least one major advance (the Jæren-Klintholm-Skjonghelleren advance at c. 45-40 kyr BP) occurred in between these two interstadials (Olsen 1997; Larsen et al. 2000; Arnold et al. 2003; Houmark-Nielsen & Kjær 2003). The LGM and post-LGM ice model adopted is that previously constrained by rebound data across Scandinavia and northern Europe (Lambeck et al. 1998b; Lambeck & Purcell 2003)."[38]

Baltic GlaciationsEdit

"After c. 22 ka BP during the Jylland stade (Houmark-Nielsen 1989), Late Weichselian glaciers of the Main Weichselain advance overrode Southeast Denmark from the northeast and later the Young Baltic ice invaded from southeasterly directions. Traces of the Northeast-ice are apparently absent in the Klintholm sections, although large scale glaciotectonic structures and till deposits from this advance are found in Hjelm Bugt and Møns Klint (Aber 1979; Berthelsen 1981, 1986). At Klintholm, the younger phase of glaciotectonic deformation from the southeast and south and deposition of the discordant till (unit 9) were most probably associated with recessional phases of the Young Baltic glaciation. In several cliff sections, well preserved Late Glacial (c. 14-10 ka BP) lacustrine sequences are present (Kolstrup 1982, Heiberg 1991)."[37]

Vistulian GlaciationsEdit

"The Lower Vistula formation is a distinct and spatially well defined lithostratigraphical unit [...]. It comprises sediments formed since the decline of the Middle Polish Glaciation up to the Toruń Glaciation (Early Vistulian)."[39]

"The till of the Toruń Glaciation (till BII) overlies the Lower Vistula formatian in almost the whole area, but pre-glacial sediments are included in this formation [...]."[39]

Skjonghelleren GlaciationsEdit

Skjonghelleren is a cave on the island of Valderøy. Credit: ElekTrond.

The Ebersdorf Stadial may correspond to the earlier two glaciation (I & L) of the Skjonghelleren Glaciations of Scandinavia where ice crosses the North Sea between 50-40 ka BP.

"Two radiocarbon dates on bones [from the Skjonghelleren (cave)] and three Uranium series dates on speleothems from this bed all cluster around 30,000 B.P. [Bed G: 29,600 ± 800, 32,800 ± 800], i.e., the end of the Ålesund interstadial. Above the uppermost laminated bed, bone fragments of birds, fish and mammals, deposited between c. 12,000 and c. 10,000 B.P. [Bed B: 10,360 ± 170, 11,510 ± 190] were found."[40]

"Three sequences of laminated clay [are Beds F, I & L], suggesting that the cave has survived at least three glaciations since its formation. Four blocky units were formed in ice-free periods prior to [Block K], between [Block G], and after the deposition [Block B] of the laminated sequences."[40]

Bed A is travertine, Beds C, H & J are silt, Bed D is granulated clay, Bed E is clay with intraclasts, Block M is above the Bedrock.[40]

Beds A and B "were deposited after the last deglaciation. The date 11,510 ± 190 B.P. [...] gives a minimum age for the last deglaciation in the Skjonghelleeren area. Previous work (Mangerud et al. 1981a), however demonstrates the the deglaciation occurred some time before 12.3 ka."[40]

The sequence from young to old is

  1. A - travertine,
  2. B - c. 10,000 - 12,000 a, ice-free period,
  3. C - silt,
  4. D - granulated clay,
  5. E - clay with intraclasts,
  6. F - glaciation, ~ 20,000 a,
  7. G - Block G, c. 29,000 - 34,000 a,
  8. H - silt,
  9. I - glaciation, ~ 40,000 a,
  10. J - silt,
  11. K - ice-free period,
  12. L - glaciation, ~ 50,000 a,
  13. M - ice-free period,
  14. N - Bedrock.

Devensian glaciationsEdit

"The results of paired terrestrial cosmogenic nuclide analyses 26
constrain the timing of this extensive glaciation and provide, for the first time, an age for the exposure of Lundy granite following deglaciation. The results from nine paired samples yield 26
exposure ages of 31.4-48.8 ka (10
) and 31.7-60.0 ka (26

"Bowen et al. (2002) used 36
to provide exposure ages for glacial landforms around Ireland and demonstrated that the most extensive phase of glaciation in many areas dates from the Early Devensian."[41]

" A 10
exposure age of 19.8 ka from an upturned boulder on the Isles of Scilly (McCarroll et al., 2010), at the southernmost limit of an Irish Sea Ice Stream (ISIS), may suggest a Late Devensian age close to the global Last Glacial Maximum (LGM; 21 ka) a date which is supported by radiocarbon and thermoluminescence dates (Wintle, 1981; Scourse, 1991, 2006)."[41]

"Furthermore, this age for the glacial and related sediments on the Isles of Scilly is independently supported by ice rafted debris of Celtic Sea sources in continental margin cores from the Goban Spur that date to Heinrich Event 2 (Scourse et al., 2000, 2009; Haapaniemi et al., 2010). This pulse may well represent the Celtic Sea advance to Scilly during the LGM, as discussed in Scourse and Furze (2001)."[41]

Wisła GlaciationsEdit

"The whole post-Eemian complex has been previously referred by me to the North Polish (Baltic, Wisła, Vistulian) Glaciation; at present the lower tills BI and BII are connected with the Toruń Glaciation and only the upper tills (BIII, BIV and BV - with the Wisła Glaciation (A. Makowska, 1986b, 1992)."[39]

"The ice sheet of the Toruń Glaciation occupied a smaller area in comparison with the two younger ice sheets of the Wisła Glaciation."[39]

The Wisła Glaciation consists of four climatic variations:[39]

  1. Late Wisła: Oldest Dryas--Younger Dryas,
  2. Upper Wisła: Leszno-Pomorze Stadial,
  3. Middle Wisła: Grudziądz (Łęcze) Interstadial, and
  4. Lower Wisła: Świecie Stadial.

These were after the Krastudy Interglacial.[39]

The Toruń Glaciation preceded the Krastudy Interglacial.[39]

Eem Interglacial preceded the Toruń Glaciation.[39]

The Middle Polish Glaciation (decline) or Sztum Warming preceded the Eem Interglacial.[39]

North Polish GlaciationsEdit

The North Polish Glaciations include the Baltic, Vistulian and Wisła Glaciations.[39]

Karmøy glaciationsEdit

The Karmøy stadial begins in the high mountains of Norway about 58 kyr B.P. and expands to the outer coast by 60 kyr B.P.[42]

The Schalkholz Stadial in North Germany is equivalent.

In "the period 52-70 ka when the [ice rafted debris] IRD curves [...] show the highest peak during the entire Weichselian [...] the ice sheet [may have] reached the continental shelf during [the Karmøy glaciation]."[43]

It "appears likely that Scandinavia was colonised by European lemmings (Lemmus sp.) during an interstadial period sometime between the Karmøy glaciation, which ended ~ 60 kyr BP (Mangerud et al. 2011), and the last glacial advance ~ 30 kyr BP [...]."[44]

"The principal stadial during MIS 4 is assumed to be the Karmøy glaciation of Norway as it is correlated with the Ristinge and Old Baltic Ice Advance of Denmark, Göteborg I of Sweden, and Vuoddasjarvi of Finland. Ice coverage during this event peaked at c. 64 kyr BP, with its eastern extent over Russia and the Kara Sea as defined in [Lambeck et al. 2006]. Subsequent ice retreat was substantial but not without interruption and re-advance, culminating in the Ålesund interstadial."[45]

Toruń GlaciationsEdit

"In general stratigraphical schemes (A. Makowska, 1986b, 1992) the pre-glacial part of the Toruń Glaciation forms a relatively short time span between the end of the Eemian Interglacial determined at about 115 ka and the beginning of ice sheet advance during the Malbork Phase of the Toruń Glaciation, defined at 110 ka."[39]

At "present the lower tills BI and BII are connected with the Toruń Glaciation".[39]

"The till of the Toruń Glaciation (till BII) overlies the Lower Vistula formation in almost the whole area, but pre-glacial sediments are included in this formation [...]."[39]

The Toruń Glaciation begins with the Herning Stadial, Brørup interstadial, the Rederstall Stadial, the Odderade interstadial, the Karøy stadial, the Oerel interstadial, the Ebersdorf Stadial, and ends with the Glinde interstadial (or, Krastudy interstadial).[39]

Würm glaciationsEdit

Violet is the extent of the Alpine ice sheet in the Würm glaciation. Blue is earlier ice ages. Credit: Lencer.
Würm glaciation, shown in ice core data from Antarctica and Greenland. Credit: Leland McInnes.
Moraines and gravel beds formed in the Würm glaciation near Leutkirch, Westallgäu, Germany, Zeil castle can be seen on the left. Credit: Rhmaster.

The Würm glaciation or Würm stage or ice age, in the literature usually just referred to as the Würm,[46] often spelledt "Wurm", was the last glacial period in the Alpine region.

The Würm ice age may be dated to the time about 115,000 to 11,700 years ago, the sources differing depending on whether the long transition phases between the glacials and interglacials (warmer periods) are allocated to one or other of these periods. The average annual temperatures during the Würm ice age in the Alpine Foreland were below −3 °C (today +7 °C), determined from changes in the vegetation (pollen analysis) as well as differences in the facies.[47]

The corresponding ice age of North and Central Europe is known as the Weichselian glaciation. Despite the global changes in climate that were responsible for the major glaciations cycles, the dating of the Alpine ice sheet advances does not correlate automatically with the farthest extent of the Scandinavian ice sheet.[48][49] In North America the corresponding "last ice age" is called the Wisconsin glaciation.[50]

Wisconsin glaciationsEdit

The Wisconsin glaciation was the most recent glacial period of the North American ice sheet complex that included the Cordilleran Ice Sheet, which nucleated in the northern North American Cordillera; the Innuitian ice sheet, which extended across the Canadian Arctic Archipelago; the Greenland ice sheet; and the massive Laurentide Ice Sheet,[51] which covered the high latitudes of central and eastern North America.

On Kelleys Island in Lake Erie, northern New Jersey and in New York City's Central Park,[52] the grooves left in rock by these glaciers can be easily observed.

Two related movements have been termed Wisconsin: Early Wisconsin and Late Wisconsin.[53]

The first Wisconsin period erased all the Illinoian glacial topography that it extended over.[53] The Late Wisconsin ice sheet extended more towards the west than the earlier movements, perhaps due to changes in the accumulation center of the ice sheet, topographic changes introduced by the Early phase or by pressure changes in the ice mass in the north.[53]

Estimated Age of Glacial Episodes (Leverett)[53]
Age Years before Present (YBP)
Culmination of Late Wisconsin 50,000
Culmination of Early Wisconsin 100,000
Beginning of Wisconsin 150,000
Culmination of Illinoian 300,000
Beginning of Illinoian 350,000
Culmination of Pre-Illinoian, i.e., old Nebraskan[54][55] 550,000
Beginning of Pre-Illinoian 1,200,000
Ice Caps[56]
Keewatin ice sheet Laurentide Ice Sheet Nova Scotia Ice Cap Newfoundland Ice Cap Greenland Ice Cap

The Labrador Ice Sheet centered east of Hudson Bay, expanding towards the southwest, into the eastern edge of Manitoba and across the Great Lakes to the Ohio River, upwards of 1,600 miles (2,600 km) from its source, with its eastern lobes covering New England and reaching south to Cape Cod and Long Island, New York.[57]

Glacial lobes and sublobes of the southern Laurentide Ice Sheet during the late Wisconsin Glaciation.[58]
Major Lobes Minor Lobes
Des Moines Grantsburg St. Louis Rainey
Lake Superior[56] Wadena Chippewa[56] Wisconsin Valley[56] Langlade[56]
Green Bay[56]
Lake Michigan[56] Delavan Harvard-Princeton Peoria Decatur
Minor lobes: Milwaukee, Two Rivers; Straits of Mackinac
Lake Huron[56] East White[56] Miami[56] Scioto[56]
Lake Erie[56]
Lake Ontario[56] Lake Champlain[56] Hudson River[56]
unnamed lobe in Quebec – New England Connecticut Valley[56] Buzzards Bay[56] Cape Cod[56] Georges Bank[56]

The Keewatin Ice Sheet began west of Hudson Bay in the Canadian Territory of Keewatin, moving south some 1,500 miles (2,400 km) into Kansas and Missouri, west 1,000 miles (1,600 km) to the foothills of the Rocky Mountains.[57]

The Cordilleran Ice Sheet left remnants throughout the Northern Rocky Mountains, unlike the other two ice sheets was mountain based covering British Columbia and reaching into northern Washington and Montana, with more of an Alpine style of many glaciers merged into a whole and striations made by the ice field in moving over the bedrock show that it moved principally to the west through the passes of the coast range.[57]

Maxima of the Wisconsin ice sheets[59]
Western Ice Eastern Ice Proximate years ago Position of ice border
Mankato Valders 25,000 Northern Washington, Idaho, and Montana to the Continental Divide – north of Edmonton – 65 miles east of Edmonton – northwest corner of North Dakota – Des Moines – west end of Lake Superior – Milwaukee – Port Huron – Buffalo – Schuylerville – St. Johnsbury.
(Great reduction of ice) Cary 27,500 Minneapolis – north Wisconsin – south of Chicago – Central Ohio – 50 miles south of Buffalo – Binghamton - Northampton
Tazewell 40,000 Rockford, Ill. – Peoria – south of Indianapolis – north of Cincinnati – northwestern Pennsylvania – central Long Island
Iowan No known ice 65,500 Northern Washington, Idaho, and Montana – northwest North Dakota – east central Iowa - Minneapolis

Middle Polish GlaciationsEdit

The Middle Polish Glaciations end with the beginning of the Eem Interglacial.[39]

South Polish GlaciationsEdit

The South Polish Glaciations occurred before the beginning of the Middle Polish Glaciations.[39]

Riss GlaciationsEdit

Extent is of the Mindel and Riss glaciation (blue) in comparison with that of the Würm period. Credit: Lencer.
Alpine Riss glaciation (in the north: the Saale) is compared with the later Würm glaciation (in the north: the Weichselian). Credit: Juschki.

In the Riss stage, there were several advances of the ice sheet, so that it can be divided into interstadials (ice retreats) and stadials (ice advances), and at least one hitherto unnamed warm period.[60]

The Riss ice age is roughly contemporaneous with the Saale glaciation of the North German glacial sequence. The Riss is paralleled by Marine isotope stage (MIS) 6, 8 and 10, which would therefore place it about 350,000 and 120,000 years ago.[61]

Excluded from the Riss glaciation is the so-called Old Riss (Ältere Riß),[62] the time of the greatest ice advance in the Alpine region: today it is referred to as the Haslach-Mindel complex (in Bavaria and Austria), Hoßkirch complex (in Baden-Württemberg) or Great Glaciation in Switzerland.

Saale GlaciationsEdit

Maximum extent (Drenthe stadium) of the Saale complex (yellow line). The red line shows the greatest extent of the younger Weichselian glaciation. Credit: Christian Fischer.

The Saale complex is currently estimated, depending on the source, as existing from around 300,000 to 130,000 years ago or 347,000 to 128,000 years ago (duration: around 219,000 years), roughly contemporaneous with the glaciation of the Riss Glacial in the Alpine region.[63]

The Saale complex may be divided into a lower (also Saale Early Glacial[64]) and an upper section (also Middle and Upper Saale Glacial,[64] or Younger Saale glaciation[65]), with glacial advances into Northern Germany.

The "Late Saalian (c. 140 kyr BP) [the stadials of marine isotope stage 5 (MIS-5d and 5b) and MIS-4 is] when the ice sheets were larger than at any time during at least the last two glacial cycles, to the final retreat of the MIS-4 glacier ice from the arctic Russian plain at c. 60 kyr BP."[38]

"The Late Saalian corresponds to a prolonged cold period for Europe during which the ice extended further south than for any subsequent period (e.g. Svendsen et al. 2004) and the advance occurred in at least two phases: the Drenthe and the Warthe."[38]

"Interglacial conditions existed at c. 210 kyr BP, with global ice volumes similar to those of today. Ice growth commenced primarily in the Kara Sea area, similar to the development during the early part of the last cycle. An oscillatory increase in ice volume occurs from c. 195 kyr BP up to the Drenthe advance with ice volumes growing in the same ratio as the ice- volume equivalent-sea-level change. By 180 kyr BP the ice sheet has expanded over the Barents􏰀Kara Sea, the Taymyr and Putorana areas of arctic Russia, and over Norway, northern Sweden and Finland. The ice margins at this time are assumed to have been similar to those that occurred later during the stadials MIS-5d and 5b. The Drenthe maximum occurs at c. 155 kyr BP and has a duration of c. 5 kyr. Some ice retreat occurs between the Drenthe and Warthe, consistent with the sea-level rise inferred at c. 150 kyr BP. This is followed by a readvance to the Warthe maximum at c. 143 kyr BP. The Warthe maximum lasts until 140 kyr BP and is followed by rapid melting. The penultimate glacial maximum over Scandinavia ends at c. 135 kyr BP, corresponding to the midpoint between the onset of the Warthe deglaciation and the time sea levels globally reached their present level in the subsequent interglacial. By 135 kyr BP the Russian ice has retreated to the Kara Sea."[38]

"Two recent compilations have been used to establish the ice margins for the Warthe phase of the Late Saalian (Ehlers & Gibbard 2003, 2004; Svendsen et al. 2004) (Fig. 2A). At this time, the Barents Sea was glaciated with an ice sheet extending out to the shelf edge west of Svalbard and Bear Island (Mangerud et al. 1998) and into the Arctic Ocean (Spielhagen et al. 2004). The southern margin in Siberia lies some 1400 km south of the arctic coastline. In the west, the ice sheet extends across the North Sea and joins up with the British ice sheet, the ice margin of which is assumed to have been similar to that for the Late Devensian 􏰀 corresponding to the Late Weichselian. In so far as the model predictions will not be used for sites in the British Isles and because the volume of ice over the British Isles represents only a few percent of the volume of the MIS-6 Eurasian ice, this approximation is adequate."[38]

Illinois Episode glaciationsEdit

"Ages of sediments immediately beneath the oldest till (Kellerville Mbr.) in the bedrock valley average 160 ka and provide direct confirmation that Illinois Episode (IE) glaciation began in its type area during marine isotope stage (MIS) 6. The oldest deposits found are 190 ka fluvial sands on bedrock in the deepest part of the valley. These correlate to earliest MIS 6. We now correlate the lowest deposits to the IE (Pearl Fm.)."[66]

"Illinoian [is] (ca. 220,000-430,000 yr BP)".[67]

MIS Boundary 7/8 is at 243 ka.[68]

Mindel glaciationsEdit

The Mindel glaciation is commonly correlated to the Elster glaciation of northern Europe, but the more precise timing is controversial since Mindel is commonly correlated to two different marine isotope stages, MIS 12[69] (478-424 thousand years ago[70]) and MIS 10[71] (374-337 thousand years ago[70]).

Penultimate Ice AgesEdit

The glaciation complex between the end of the Holstein interglacial and the beginning of the Eem interglacials is referred to as the Penultimate Ice Age and the Great Glaciation.[72]

Elster glaciationsEdit

Schematic diagram shows the maximum glaciation of the last three cold periods on the North German Plain:
red line = extent of the Weichselian glaciation;
yellow line = extent of the Saale glaciation;
blue line = extent of the Elster glaciation. Credit: Botaurus.

The glacial period is named after the White Elster, a right tributary of the Saale.[73]

Elster was correlated with the Mindel glaciation of the Alps and the Anglian glaciation of Great Britain and Ireland, but analysis in the 1950s of oxygen isotopes in deep sea core samples introduced a global glacial history, with warm and cold phases identified by marine isotope stages (MISs), identifying two glacial stages in the time slot of the Elster/Mindel/Anglian, namely MIS 12 and MIS 10:[74]

  • MIS 12, 478-424 ka ago,[70] is globally the stronger of the two glacials and long the preferred correlation to Elster/Mindel/Anglian. There is strong evidence for widespread glaciation of Great Britain during MIS 12, and only disputed and weak signs of glaciation during MIS 10.[75] Thus, the correlation of the Anglian glaciation to MIS 12 is uncontroversial. The glacial history of Europe is much simplified if also Elster and Mindel are correlated to MIS 12. The Subcommission on Quaternary Stratigraphy (SQS) of the International Commission on Stratigraphy (ICS), a scientific organisation within the International Union of Geological Sciences (IUGS), correlates both the Anglian and the Elsterian to MIS 12 in the 2011 version of its correlation table.[69]
  • MIS 10, 374-337 ka ago,[70] is globally a weaker glacial. A correlation of Elster to MIS 10 implies a complicated glacial history in Europe, with various geographical areas showing evidense from different glacials. However, different chronologies of separate ice strems during the late Weichselian glaciation gives credence to such a scenario.[76] In the Netherlands there is evidense for correlations of Elster to both MIS 12 and MIS 10.[77] A correlation to MIS 10 is given in the comprehensive review of the glaciations of northern Germany by Litt et al. (2007).[78] Version 2016 of the correlation table by the German Stratigraphic Commission correlates both Elster and Mindel to MIS 10.[71]

Haslach glaciationsEdit

Haslach Ice Age was not included in the traditional glacial schema of the Alps.[79] The glacial stage was first described[80] from its type region is the Haslach Gravels (Haslach-Schotter) in the area of the Riß-Iller-Lech Plateau.

Gunz glaciationEdit

Deep sea core samples have identified approximately 10 marine isotope stages (at least MIS 21 to MIS 11); i.e., 5 glacial cycles of varying intensity during Gunz.[71][81]

The Günz was thought to follow the Danube-Günz interglacial and was ended by the Günz-Haslach interglacial.[82][83][72]

The German Stratigraphic Commission puts the start of Gunz in the late Calabrian (approximately one million years ago, earlier than MIS 19) and shows a continuity of glacial cycles with the following Mindel stage, with the border arbitrarily put at the start of MIS 10 (approximately 374 000 years ago), corresponding roughly to the Cromerian stage in the glacial history of Northern Europe.[71]

During Gunz the 41,000 year glacial cycle of previous stages (Biber and Danube) had been replaced by a dominance of a 100,000-year cycle (Mid-Pleistocene Transition), where the most intense glacials of Gunz (MIS 16 and MIS 12) reached similar extents to those of the more recent Riss and Wurm glaciations.[84][85] These have not been easy to identify in the geological record of the Alps, but MIS 16 has been identified with the Don Glaciation of Eastern Europe.[86]

The strong glacial MIS 12 has been problematic, and has sometimes been identified with the Mindel glaciation, which would imply an end to Gunz already after MIS 13 (480 000 years ago).[87]

Biber ice ageEdit

The base of the marly layer overlying sapropel MPRS 250, located at 62 m in the Monte San Nicola section, is the defined base of the Gelasian Stage. Credit: D. Rio, R. Sprovieri, D. Castradori, and E. Di Stefano.

Some number of N tills occurred during the Olduvai subchron.[88]

The magnetic field reversal to the present geomagnetic poles (Olduvai subchron) occurred at 2,000,000 yr BP.

The oldest till group, R2 tills, consists of till units with a reversed polarity and >77% of sedimentary clasts. Low amounts of expandable clays, substantial amounts of kaolinite, and the absence of chlorite characterize the clay mineralogy of R2 tills. The mineralogy of the silt fraction of R2 tills is rich in quartz and depleted in calcite, dolomite, and feldspar. This till group includes a till unit that underlies the 2.0-Ma Huckleberry Ridge ash, thus indicating deposition sometime between ~2.5 Ma (onset of Northern Hemisphere glaciations) (Mix et al., 1995) and 2.0 Ma.[88]

During the Gelasian the ice sheets in the Northern Hemisphere began to grow, which is seen as the beginning of the Quaternary ice age. Deep sea core samples have identified approximately 40 marine isotope stages (MIS 103 – MIS 64) during the age. Thus, there have probably been about 20 glacial cycles of varying intensity during the Gelasian.

In the regional glacial history of the Alps this age is now called Biber. It corresponds to Pre-Tegelen and Tegelen in Northern Europe.[89]

During the Gelasian, the Red Crag Formation of Butley, Suffolk, the Newbourn Crag, the Norwich Crag Formation and the Weybourne Crag Formation (all from East Anglia, England) were deposited. The Gelasian is an equivalent of the Praetiglian and Tiglian stages as defined in the Netherlands, which are commonly used in northwestern Europe.

Biber or the Biber Complex is a timespan approximately 2.6–1.8 million years ago in the glacial history of the Alps. Biber corresponds to the Gelasian age in the international geochronology, which since 2009 is regarded as the first age of the Quaternary period. Deep sea core samples have identified approximately 20 glacial cycles of varying intensity during Biber.[71]

In 1953, Schaefer defined the Biber glaciation, Biber Glacial, or Biber Ice Age from gravel landforms of the Stauden Plateau in the area of the Iller-Lech Plateau and in the Aindling river terrace sequence, by grouping together the so-called Middle and Upper Cover Gravels or Deckenschotter. This corresponded to the Staufenberg Gravel Terrace on the Iller-Lech Plateau, identified in 1974 by Scheunenpflug, and the so-called High Gravels of the Aindling region.[90] The rich crystalline sedimentary facies, that Löscher distinguished in 1976 in the area of the Rhine Glacier of the western Riß-Iller Plateau may also be paralleled with these glacial landforms.[91] The gravels in the Iller-Lech region ascribed to the Biber glaciation are generally heavily weathered and originate from the Northern Limestone Alps. Löscher's Kristallinreiche Liegendfazies, by contrast, originates from the bedrock of the molasse zone.

The term Biber glaciation was not part of the traditional four-stage glaciation schema of the Alps by Albrecht Penck and Eduard Brückner, but was named after the Biberbach river north of Augsburg in 1953 by Ingo Schaefer, based on the naming system of the traditional Penck schema.[92][93] Its type locality or type region is the Stauden Plateau in the Iller-Lech Plateaux and the Staufenberg Gravel Terrace in the area of Aindling. The Biber glaciation was thought to be followed by the Biber-Danube interglacial and the Danube glacial.

The absolute timing and the connexion with the glacial classification of North Germany and the Netherlands has been problematic. The Biber glacial was fought to correlate either to the Eburonian complex or the Pre-Tiglian complex in the Netherlands. In the former case it would correspond to Marine isotope stage (MIS) 56 to 62, which would place it in the period between 1.6 and 1.8 million years ago,[94][95] in the latter case it would roughly correspond to MIS 96 to 100, and would therefore have taken place about 2.4 to 2.588 million years ago.[95][96][97] The correlation was fraught with problems however due to the fact that the corresponding depositions in the Netherlands were probably not governed by climatic changes. Similar doubts on climatic grounds for the depositions assessed as Biber-related also exist in the Alpine region. It is possible that there were tectonic influences perhaps in the wake of the uplift phases of the Alps. The succession and appearance of the gravel bodies makes it possible that during their formation there were several periods of alternating fluvial erosion and accumulation. The Biber cold period at least corresponds partly with the Swiss cover gravel glaciations.[98]

The 2016 version of the detailed stratigraphic table by the German Stratigraphic Commission firmly places Biber in the Gelasian and gives a correspondence to Pre-Tegelen and Tegelen in the glacial geology of northern Europe. There is continuity between Biber and the glacial cycles of the following Danube stage[71]

Deep sea core samples have identified approximately 40 marine isotope stages (MIS 103 – MIS 64) during Biber.[71] Thus, there have probably been about 20 glacial cycles of varying intensity during Biber. The dominant trigger is believed to be the 41 000 year Milankovitch cycles of axial tilt.[99][100]

Gravels ascribed to Biber (also called the Highland Gravel or Oldest Gravel occur northwest of Augsburg as the Stauffenberg Gravel, as well as northeast as the Hohenried Gravel and southwest of Augsburg as the Stauden Plateau Gravel. Also included are isolated gravels of the Hochfirst near Mindelheim and the Stoffersberg near Landsberg am Lech.[101]


Lithostratigraphy and magnetostratigraphy of the Patapatani West section is correlated with the geomagnetic polarity time scale and oxygen isotope record, and compared with contemporary glacial records. Credit: Nicholas J. Roberts, René W. Barendregt & John J. Clague.{{fairuse}}

Recurrent "large-scale glaciation in the Bolivian Andes [is] based on stratigraphic and paleomagnetic analysis of a 95-m sequence of glacial sediments underlying the 2.74-Ma Chijini Tuff. Paleosols and polarity reversals separate eight glacial diamictons, which we link to cold periods in the benthic oxygen isotope record."[102]

"The mid-Piacenzian (3.265–3.025 Ma8) is the most recent of Earth’s warm periods".[102]

"Common striated and faceted clasts show that matrix-supported, generally massive, diamicton units below the [Chijini Tuff] (units 3–7) are glacial in origin."[102]

"The Chijini Tuff (unit 8) comprises 20–45 cm of loose friable ash at its base, transitioning upward into ~9.5 m of cliff-forming, weakly cemented ash."[102]

"A massive diamicton overlies the [Chijini] tuff and is similar to diamicton units below it [...]."[102]

Unit "1 [is likely] proglacial outwash. Unit 2 may be the deposit of a glaciogenic debris flow, suggestive of a more ice-proximal setting. Oxidized zones capping diamicton units 3, 4, 5, and 6 are paleosols, indicating subaerially exposed surfaces that were stable for thousands to tens of thousands of years [...]."[102]

"Common striated and faceted clasts show that matrix-supported, generally massive, diamicton units below the tuff (units 3–7) are glacial in origin. They were deposited either beneath or at the margin of glaciers during separate glaciations. Strong clast alignment is consistent with high subglacial shear stresses26,27, suggesting ice flow initially to the south-southwest and later to the south-southeast [...]."[102]

"In light of the long periods of landscape stability indicated by each of the four paleosols, these glaciers formed during at least five separate glaciations."[102]

Glaciers "extended at least 14 km from the high Cordillera Real."[102]

"Two polarity reversals within unit 3 indicate that it comprises three subunits (3a, 3b, 3c [...]), possibly separated by hiatuses."[102]

Till "overlying the [Chijini Tuff] (Calvario Drift) records [an early] glaciation of the Cordillera Real to the northeast."[102]

"The new 40
age on sanidine −2.74 ± 0.04 Ma [...] – confirms that the Chijini Tuff is latest Pliocene in age and allows polarity data to be confidently correlated to the geological timescale [...]. The polarity sequence constrains all diamicton units within the Gauss Chron [...], which spans the period 3.588–2.608 Ma (late Pliocene). Two polarity zones in the sequence (normally magnetized mid-Gauss Chron, C2An.2n; and reversely magnetized Kaena subchron, C2An.2r) are completely within the mid-Piacenzian warm period. The oldest normal magnetozone in the section (N1 [...]) is likely early Gauss".[102]

"K-Ar23 and 40
24 ages on potassium feldspar in the Chijini Tuff, which range from 2.650 ± 0.012 to 2.8 ± 0.1 Ma, are considered to be reliable and confirm that underlying sediments are pre-Pleistocene."[102]

"Of the nine pre-Chijini glacial units defined by polarity reversals and unconformities, six correspond unambiguously to five specific marine isotope stages: MG2, M2, KM2, G22 and G10 [...]. Two glacial units (3a and 3b) record the single strong cold peak within their respective subchrons (M2 during C2An.2r and KM2 during C2An.2n). Formation of a paleosol between the two tills that fall within the Kaena subchron (units 3c and 4) requires glaciation during both of that subchron’s strong cold peaks (KM2 and G22, respectively); MIS KM2 is thus recorded by two tills of opposite polarity, reflecting its occurrence at a polarity reversal. Units 2 and 7 are most likely to have been deposited during the strongest and latest of multiple cool peaks (MG2 and G10, respectively) preceding the start of the Mammoth subchron and deposition of the Chijini Tuff. Each of the other three glacial units is constrained to a small number of similar-magnitude cold peaks [...] during a given polarity subchron."[102]

Units "1–3 record either three or four glaciations, depending on whether unit 1 was deposited shortly before unit 2 during MIS MG2 or during an earlier glaciation in MIS MG1 [...]. The presence of well-developed paleosols, marking interglacial conditions, between units 4–7 indicates that these tills record four subsequent glaciations. The measured sequence below the tuff thus records either seven or eight glaciations. The expansion of Andean ice indicated by the shift from a proglacial environment (unit 1) to an ice-marginal or subglacial environment (unit 3a) records climate deterioration directly preceding and during the globally recognized12,29 MIS M2 (ca. 3.3 Ma) cooling event."[102]

"An Andean ice cap formed again during each of the two coolest parts of the mid-Piacenzian warm period (KM2 and G22) and in three subsequent glaciations (ending with G20) during climatic deterioration between ca. 3.0–2.8 Ma."[102]

Ice ages Glaciations Chron Stadials Interstadials
Name Name Name Name Timeline Name Timeline
Holarctic-Antarctic, Late Cenozoic, 33.9 million years ago Piacenzian, 3.6 Ma to 2.58 Ma Gauss
Piacenzian 104 2.62 Ma to 2.58 Ma Galasian 103 2.58 Ma to 2.57 Ma
G2 2.66 Ma to 2.63 Ma G1 2.63 Ma to 2.62 Ma
G4 2.70 Ma to 2.68 Ma G3 2.68 Ma to 2.66 Ma
Calvario Drift, G6 2.74 Ma to 2.71 Ma G5 2.71
G8 2.78 Ma to 2.77 Ma G7, Unit 8 2.77 Ma to 2.74 Ma
G10, Unit 7 2.83 Ma to 2.79 Ma G9 2.79 Ma to 2.78 Ma
G12, Unit 6 2.86 Ma to 2.84 Ma G11 2.84 Ma to 2.83 Ma
G14 2.90 Ma to 2.88 Ma G13 2.88 Ma to 2.86 Ma
G16 2.974 Ma to 2.972 Ma G15 2.972 Ma to 2.90 Ma
G18 2.98 Ma to 2.975 Ma G17 2.975 Ma to 2.974 Ma
G20 3.025 Ma to 3.00 Ma G19 3.00 Ma to 2.8 Ma
Mid-Piacenzian Warm Period, 3.6 Ma to 2.58 Ma
G22, Unit 4 3.05 Ma to 3.045 Ma G21 3.045 Ma to 3.025 Ma
cooling K1 3.08 Ma to 3.07 Ma K1 3.07 Ma to 3.05 Ma
K2 3.1 Ma to 3.09 Ma K1.5 3.09 Ma to 3.08 Ma
KM2, Unit 3c 3.11 Ma to 3.11 Ma K3 3.11 Ma to 3.1 Ma
KM2, Unit 3b 3.15 Ma to 3.127 Ma K5 3.127 Ma to 3.11 Ma
cooling 3.19 Ma to 3.18 Ma K7 3.18 Ma to 3.15 Ma
cooling 3.20 Ma to 3.19 Ma K9 3.19 Ma to 3.19 Ma
stadial 3.211 Ma to 3.210 Ma K11 3.21 Ma to 3.20 Ma
KM6 3.23 Ma to 3.211 Ma K13 3.211 Ma to 3.211 Ma
M2, Unit 3a 3.31 Ma to 3.265 Ma M1 3.265 Ma to 3.23 Ma
MG2, Unit 2 3.34 Ma to 3.34 Ma MG1 3.34 Ma to 3.31 Ma
MG4, Unit 1 3.375 Ma to 3.37 Ma MG3 3.37 Ma to 3.34 Ma
MG6 3.47 Ma to 3.47 Ma MG5 3.47 Ma to 3.375 Ma
MG8 3.53 Ma to 3.52 Ma MG7 3.52 Ma to 3.47 Ma
MG10 3.57 Ma to 3.55 Ma MG9 3.55 Ma to 3.53 Ma
MG12 3.59 Ma to 2.58 Ma MG11 3.58 Ma to 3.57 Ma

Antarctic Eocene-Paleocene glaciationEdit

The "Ross Sea Embayment area has been considered a laboratory of growing interest for the [...] onset of Antarctic Eocene-Paleocene glaciation".[103]

"The broad over-deepened and landward sloping Ross Sea outer continental shelf occupies a 1,000 km wide embayment on the present Antarctic margin. It is delimited to the west by the Victoria Land Coast (East Antarctic Craton), to the south by the Ross Ice Shelf, to the east by the Marie Byrd Land (West Antarctica) and to the north by the continental slope and rise. Its bathymetry and sedimentary cover were formed and shaped during the continental rifting phases of Late Cretaceous and Cenozoic West Antarctic Rift System (Behrendt et al., 1991; Rocchi et al., 2002, 2005), and by the associated marine and glacio-fluvial processes."[103]


Global average land (above) and deep sea (below) temperatures throughout the Cenozoic

The Paleocene climate was, much like in the Cretaceous, tropical or subtropical,[104][105][106][107] and the poles were temperate[108] and ice free[109] with an average global temperature of roughly 24–25 °C (75–77 °F).[110] For comparison, the average global temperature for the period between 1951 and 1980 was 14 °C (57 °F).[111] A 2019 study identified changes in orbital eccentricity as the dominant drivers of climate between the late Cretaceous and the early Eocene.[112]

Global deep water temperatures in the Paleocene likely ranged from 8–12 °C (46–54 °F),[113][114] compared to 0–3 °C (32–37 °F) in modern day.[115] Based on the upper limit, average sea surface temperatures at 60°N and S would have been the same as deep sea temperatures, at 30°N and S about 23 °C (73 °F), and at the equator about 28 °C (82 °F).[114] The Paleocene foraminifera assemblage globally indicates a defined deep-water thermocline (a warmer mass of water closer to the surface sitting on top of a colder mass nearer the bottom) persisting throughout the epoch.[116] The Atlantic foraminifera indicate a general warming of sea surface temperature–with tropical taxa present in higher latitude areas–until the Late Paleocene when the thermocline became steeper and tropical foraminifera retreated back to lower latitudes.[117]

Early Paleocene atmospheric CO2 levels at what is now Castle Rock, Colorado, were calculated to be between 352 and 1,110 parts per million (ppm), with a median of 616 ppm. Based on this and estimated plant-gas exchange rates and global surface temperatures, the climate sensitivity was calculated to be +3 °C when CO2 levels doubled, compared to 7° following the formation of ice at the poles. CO2 levels alone may have been insufficient in maintaining the greenhouse climate, and some positive feedbacks must have been active, such as some combination of cloud, aerosol, or vegetation related processes.[118]

The poles probably had a cool temperate climate; northern Antarctica, Australia, the southern tip of South America, what is now the US and Canada, eastern Siberia, and Europe warm temperate; middle South America, southern and northern Africa, South India, Middle America, and China arid; and northern South America, central Africa, North India, middle Siberia, and what is now the Mediterranean Sea tropical.[119]

Karoo Ice AgeEdit

Approximate extent of the Karoo Glaciation is shown in blue, over the Gondwana supercontinent during the Carboniferous and Permian periods. Credit: GeoPotinga.{{free media}}
Ice flow in the Karoo basins over southern Africa during maximum glaciation is indicated. Credit: J. N. J. Visser.
North-south section across the Kalahari and Karoo basins illustrating the relief of the basin floor and the lithostratigraphic units. Credit: Visser.

A "glacial marine facies [occurs] on the Falkland Islands [Frakes and Crowell, 1967]."[120]

In "a complex situation, like the Karoo Basin and adjoining highlands [...] a marine ice sheet bounded the highlands during the last phase of glaciation".[120]

The "influence of Gondwana topography on glaciation about 275 to 300 m.y. ago, [...] is preserved as thick (up to 700 m) glacial and proglacial sequences in the Karoo, Kalahari, and Warmbad basins as well as other smaller basins toward the north. These deposits, known as the Dwyka Formation, cover an area close to a million square kilometers in southern Africa".[120]

"Valleys that had been incised into the Windhoek Highlands attained lengths up to 250 km, had striated floors and walls, and contained roches moutonnées [Martin, 1961]."[120]

"The Whitehill Formation (White Band) was taken as a datum [in the stratigraphic diagram on the lower right]. The Permo-Carboniferous boundary on the platform is based on microflora assemblages [Anderson, 1977]. Ms is massive; St, stratified; dmt, diamictite; Drg, dropstone argillite; Cb, carbonaceous; mds, mudstone; Sst, sandstone; Lm, laminated; cgl, conglomerate; sh, shale; Fs, fossiliferous; Cc, carbonate; and conc, concretions."[120]

"Glaciation is known from all continents that were once part of Gondwana, including: Africa, South America, Antarctica, India, Arabia and Australia. Glaciation began in the early Carboniferous (360 Ma), reached a peak in late Carboniferous, continued into early Permian, and mostly came to an end by late Permian (260 Ma) time, thus spanning 100 million years. Multiple glacial centers were active; each experienced repeated glacier advances and retreats. Particularly well-known glacial strata include the Dwyka Tillite (Karoo basin) in South Africa, Talchir Boulder Beds in India, and Wynyard Formation of Tasmania. Overall, two major glacial cycles took place. Both expanded gradually over periods of about 20 million years each to reach their maximum extents in late Pennsylvanian and early Permian times. Each major cycle then ended abruptly during only 1-10 thousand years (Gastaldo et al. 1996)."[9]

"Although precise dating is not possible for many of the Gondwana glacial deposits, a general migration of glaciation through time is apparent. Carboniferous glaciation took place mainly in South America, southern Africa, India, and western Antarctica; whereas Permian glaciation was located mostly in Australia and eastern Antarctica. This migration corresponded to the drifting of Gondwana over the South Pole [...] The Karoo ice age is marked by cyclothems, cyclic sedimentary sequences in continental areas that were located in low latitudes. Pennsylvanian and Permian cyclothems are well known throughout the mid-continent of the United States, particularly eastern Kansas. The cyclothems were created by repeated marine transgressions and regressions over a stable continental platform. These cycles are interpreted as results of frequent changes in global sea level associated with glaciation in Gondwana. Glacial cycles and variations in sea level are documented in oxygen-isotope variations within fossils of Pennsylvanian cyclothems [...]. Late Pennsylvanian sea-level fluctuations were at least 80 m and likely greater than 100 m in amplitude (Soreghan and Giles 1999)."[9]

Andean-Saharan ice ageEdit

Number 4 is the Andean-Saharan glaciation. Credit: Pedros.lol.{{free media}}

The "glacial episodes that occurred on Earth during the Palaeozoic (the Andean-Saharan between 450 and 420 Ma, and the Karoo between 360 and 260 Ma) did not achieve a global extent."[121]

"Glaciation is known from Arabia, central Sahara, western Africa, the lower Amazon of Brazil, and the Andes of western South America. Spectacular erosion of underlying rocks took place over large areas of the Sahara; whereas a good sedimentary record is preserved in Arabia. Continental ice sheets were developed in Africa and eastern Brazil, while alpine glaciers formed in the Andes region. The center of glaciation appears to have migrated through time: Ordovician (450-440 Ma) in Sahara, and Silurian in South America (Brazil 440-430 Ma, and Andes 430-420 Ma). The two continents were joined as parts of Gondwana, which was located over the South Pole".[9]

Gaskiers glaciationEdit

An Ediacaran glacial episode (Gaskiers) also occurs within the wide-ranging Marinoan Epoch.[122]

The Gaskiers glaciation is a period of widespread glacial deposits (e.g. diamictites) that lasted under 340 thousand years, between 579.63 ± 0.15 and 579.88 ± 0.44 million years ago[123] – i.e. late in the Ediacaran Period – making it the last major glacial event of the Precambrian.[124]

Varanger glaciationEdit

The Varangian apparently spans 610 to 575 Ma.

Elatina glaciationEdit

The 'golden spike' (bronze disk in the lower section of the image) or 'type section' of the Global Boundary Stratotype Section and Point (GSSP) for the base of the Ediacaran System. Credit: Peter Neaum.
The 'golden spike' marks the GSSP. Credit: Bahudhara.
Elatina Formation diamictite is below the Ediacaran Global Boundary Stratotype Section and Point (GSSP) site in the Flinders Ranges National Park, South Australia. An Australian $1 coin is for scale. Credit: Bahudhara.{{free media}}

"The Elatina glaciation has not been dated directly, and only maximum and minimum age limits of c. 640 and 580 Ma, respectively, are indicated."[108]

"The Elatina glaciation is of global importance for several reasons:

  1. its diverse and excellently preserved glacial and periglacial facies represent a de facto type region for late Cryogenian glaciation in general;
  2. the Elatina Fm. has yielded the most robust palaeomagnetic data for any Cryogenian glaciogenic succession; and
  3. the recently established Ediacaran System and Period (Knoll et al. 2004, 2006; Preiss 2005) has its Global Stratotype Section and Point (GSSP) placed near the base of the Nuccaleena Fm. overlying the Elatina Fm. in the central Flinders Ranges [...]."[108]

"Feeder dykes for volcanic rocks near the base of the [Adelaide Geosyncline] sedimentary succession have been dated at 867 ± 47 and 802 ± 35 Ma (Zhao & McCulloch 1993; Zhao et al. 1994) and 827 ± 6 Ma (Wingate et al. 1998)."[108]

"No volcanism is known in the region during the Elatina glaciation."[108]

"The Neoproterozoic–early Palaeozoic succession in the Adelaide Geosyncline was deformed by the Delamerian Orogeny at 514 – 490 Ma (Drexel & Preiss 1995; Foden et al. 2006)."[108]

"The Yerelina Subgroup at the top of the Cryogenian Umberatana Group embraces all the glaciogenic formations of the Elatina glaciation (Preiss et al. 1998)."[108]

"The Yerelina Subgroup is unconformably to disconformably overlain by the Ediacaran Wilpena Group."[108]

"Deposition in the North Flinders Zone commenced, possibly following an erosional break, with the 1070-m-thick Fortress Hill Fm., which comprises laminated siltstone with gritty lenses and scattered dropstones, some faceted, marking the onset of glacial deposition (Coats & Preiss 1987; Preiss et al. 1998). Clast lithologies include granite, quartzite, limestone, oolitic limestone and dolostone. The Fortress Hill Fm. is typical of the dominantly fine-grained units of the Yerelina Subgroup that are interpreted by Preiss (1992) as outer marine-shelf deposits."[108]

"The Fortress Hill Fm. is sharply overlain by sandstone and conglomerate at the base of the Mount Curtis Tillite (90 m) that may record a lowering of relative sea level and mark a sequence boundary (Preiss et al. 1998)."[108]

"The Mount Curtis Tillite is a sparse diamictite with erratics of pebble to boulder size, some faceted and striated, in massive and laminated, grey-green dolomitic siltstone. Clast lithologies are mostly quartzite, limestone and dolostone, but also include granite and porphyry (Coats & Preiss 1987). Granite boulders attain 3 x 8 m."[108]

"The Mount Curtis Tillite is overlain by the medium-grained, feldspathic Balparana Sandstone (130 m), which contains interbeds and lenses of calcareous siltstone and pebble conglomerate."[108]

"The Balparana Sandstone is disconformably overlain by the Wilpena Group. The main source for the glaciogenic deposits may have been the Curnamona Province to the present east [...] and possibly the now-buried Muloorina Ridge immediately north of the North Flinders Zone (Preiss 1987)."[108]

"The lower-most, laminated siltstone facies of the Fortress Hill Fm. shows progressively greater amounts of scattered, ice-rafted granules and pebbles. The shallow-water Gumbowie Arkose (45 – 90 m) disconformably overlies these early deposits at a possible sequence boundary and is conformably succeeded by the Pepuarta Tillite (120 – 197 m), which is a sparse diamictite with scattered clasts up to boulder size in massive and laminated, grey calcareous siltstone. Faceted and striated boulders reach 2.5 m in diameter. Clast lithologies include pink granite, granite gneiss, grey porphyry, quartz-granule conglomerate, various quartzites, and vein quartz. The siltstone facies with scattered large clasts of extrabasinal provenance implies deposition from floating ice."[108]

"The widespread Grampus Quartzite (60 m) disconformably overlies the Pepuarta Tillite, possibly at a sequence boundary defining a third genetic sequence of the Yerelina Subgroup (Preiss et al. 1998)."[108]

"It is conformably overlain by the laminated to cross-laminated, calcareous, pale grey Ketchowla Siltstone (271 m) (Preiss 1992). The Ketchowla Siltstone contains scattered ice-rafted granules, pebbles and boulders up to 1 m across, and is ascribed by Preiss (1992) to outer marine-shelf deposition under generally waning glacial conditions. It is overlain disconformably by the Nuccaleena Fm., with any Ketchowla Siltstone deposited in the North Flinders Zone having been completely removed by erosion at this sequence boundary (Preiss 2000)."[108]

"The outer marine-shelf successions of the Fortress Hill Fm. and Ketchowla Siltstone record the waxing and waning of glacial conditions, respectively. The Pepuarta Tillite and the correlative Mount Curtis Tillite mark the glacial maximum of the Elatina glaciation (Preiss et al. 1998)."[108]

"A U–Pb age of 657 ± 17 Ma was obtained for a zircon grain of uncertain provenance from the Marino Arkose Member of the underlying Upalinna Subgroup (Preiss 2000). Re – Os dating gave an age of 643.0 ± 2.4 Ma for black shale from the Tindelpina Shale Member at the base of the Tapley Hill Fm., which overlies glacial deposits of Sturtian age in the Adelaide Geosyncline (Kendall et al. 2006). Zoned igneous zircon from a tuffaceous layer near the top of the Sturtian-age glaciogenic succession gave a SHRIMP U – Pb age of c. 658 Ma (Fanning & Link 2006). Mahan et al. (2007) reported a Th–U–total Pb age of 680 ± 23 Ma for euhedral laths of monazite, interpreted as authigenic, from the Enorama Shale of the Upalinna Subgroup."[108]

"In the central Flinders Ranges the 4.5 km thick Umberatana Group encompasses the two main phases of glacial deposition (see Thomas et al., 2012). The carbonaceous, calcareous and pyritic Tindelpina Shale Member, of the interglacial Tapley Hill Formation, caps the Fe-rich diamictite and tillite formations of the Sturt glaciation. The upper Cryogenian glacials of the Elatina Formation are truncated by the Nuccaleena Formation at the base of the Wilpena Group and the Ediacaran System."[125]

"In 2004, the Global Stratotype Section and Point (GSSP) for the terminal Proterozoic was placed near the base of the Nuccaleena Formation in Enorama Creek in the central Flinders Ranges [in the image on the right], thus establishing the Ediacaran System and Period (Knoll et al., 2006). As the Nuccaleena Formation has not been accurately dated, a date of c. 635 Ma from near-correlative levels in Namibia and China is presumed for the base of the Ediacaran (Hoffmann et al., 2004; Condon et al., 2005; Zhang et al., 2005)."[125]

A glacial episode within the Marinoan Epoch is the Elatina glaciation after the 'Elatina Tillite' (now Elatina Formation).[126]

Nantuo glaciationEdit

The Nantuo glaciation apparently occurred 654 ± 3.8 Ma.

Ice Brook glaciationEdit

The Ice Brook glaciation apparently spans 651 to 659 Ma.

Ghaub glaciationEdit

"Dropstone-bearing glaciomarine sedimentary rocks of the Ghaub Formation within metamorphosed Neoproterozoic basinal strata (Swakop Group) in central Namibia contain interbedded mafic lava flows and thin felsic ash beds. U-Pb zircon geochronology of an ash layer constrains the deposition of the glaciomarine sediments to 635.5 ± 1.2 Ma, providing an age for what has been described as a “Marinoan-type” glaciation. In addition, this age provides a maximum limit for the proposed lower boundary of the terminal Proterozoic (Ediacaran) system and period. Combined with reliable age constraints from other Neoproterozoic glacial units—the ca. 713 Ma Gubrah Member (Oman) and the 580 Ma Gaskiers Formation (Newfoundland)—these data provide unequivocal evidence for at least three, temporally discrete, glacial episodes during Neoproterozoic time with interglacial periods, characterized by prolonged positive δ13C excursions, lasting at most ∼50–80 m.y."[127]

"Dropstones are ubiquitous within the finer-grained (Ghaub) lithofacies, and their presence, along with the facies context for subglacial and near grounding-line deposition, indicates a glacigenic origin for the Ghaub Formation, despite its subtropical paleolatitude and distal foreslope setting."[128]

Marinoan glaciationsEdit

This diamictite is from the Neoproterozoic Pocatello Formation, a 'Snowball Earth'—type deposit. Credit: Qfl247.{{free media}}

The term Marinoan glaciation has been applied globally to any glaciogenic formations assumed (directly or indirectly) to correlate with the Elatina glaciation in South Australia.[129] Recently, there has been a move to return to the term Elatina glaciation in South Australia because of uncertainties regarding global correlation and because an Ediacaran glacial episode (Gaskiers) also occurs within the wide-ranging Marinoan Epoch.[122]

The Marinoan glaciation was a period of worldwide glaciation that lasted from approximately 650 to 635 Ma and may have covered the entire planet, in an event called the Snowball Earth, where the end of the glaciation may have been sped by the release of methane from equatorial permafrost.[130][131] Great uncertainty surrounds the dating of pre-Gaskiers glaciations: the status of the Kaigas is not clear; its dating is very tentative and many researchers do not recognize it as a glaciation.[132]

During the Marinoan glaciation, characteristic glacial deposits indicate that Earth suffered one of the most severe ice ages in its history: glaciers extended and contracted in a series of rhythmic pulses, possibly reaching as far as the equator.[133][134]

Apparently the major glacial period the Marinoan occurred during the Cryogenian.[135]

A similar period of rifting, to the break up along the margins of Laurentia, at about 650 Ma occurred with the deposition of the Ice Brook Formation in North America, contemporaneously with the Marinoan in Australia.[136]

The Marinoan glaciation ended approximately 635 Ma, at the end of the Cryogenian.[137]

The Marinoan glaciation was a period of worldwide glaciation that lasted from approximately 650 to 635 Ma, where the end of the glaciation may have been sped by the release of methane from equatorial permafrost.[137][138]

The name is derived from the stratigraphic terminology of the Adelaide Geosyncline (Adelaide Rift Complex) in South Australia and taken from the Adelaide suburb of Marino to subdivide the Neoproterozoic rocks of the Adelaide area and encompass all strata from the top of the Brighton Limestone to the base of the Cambrian.[139] The corresponding time period, referred to as the Marinoan Epoch, spanned from the middle Cryogenian to the top of the Ediacaran and included a glacial episode within the Marinoan Epoch, the Elatina glaciation, after the 'Elatina Tillite' (now Elatina Formation).[126] The term Marinoan glaciation came into common usage because it was the glaciation that occurred during the Marinoan Epoch.[139]

The term Marinoan glaciation was applied globally to any glaciogenic formations assumed to correlate with the Elatina glaciation in South Australia.[140] The Elatina glaciation in South Australia and the Gaskiers also occurs within the wide ranging Marinoan Epoch.[122]

The Earth may have underwent a number of glaciations during the Neoproterozoic era.[141]

There were three (or possibly four) significant ice ages during the late Neoproterozoic, periods of nearly complete glaciation of Earth are often referred to as "Snowball Earth", where it is hypothesized that at times the planet was covered by ice 1–2 km (0.62–1.24 mi) thick.[142]

During the Marinoan glaciation, characteristic glacial deposits indicate that Earth suffered one of the most severe ice ages in its history, where glaciers extended and contracted in a series of rhythmic pulses, possibly reaching as far as the equator.[143][144]

The melting of the Snowball Earth is associated with greenhouse warming due to the accumulation of high levels of carbon dioxide in the atmosphere.[145]

Glacial deposits in South Australia are approximately the same age (about 630 Ma), confirmed by similar stable carbon isotopes, mineral deposits (including sedimentary barite), and other unusual sedimentary structures.[142]

Two diamictite-rich layers in the top 1 km (0.62 mi) of the 7 km (4.3 mi) Neoproterozoic strata of the northeastern Svalbard archipelago represent the first and final phases of the Marinoan glaciation.[146]

The Marinoan "is separated from the Sturtian by a thick succession of sedimentary rocks containing no evidence of glaciation. This glacial phase could correspond to the recently described Ice Brooke formation in the northern Canadian Cordillera."[136]


The Gucheng is apparently comparable to the Marinoan.


The Jiangkou spans the Chang'an through the Gucheng.


The Chang'an occurred about 715.9 ± 2.8 Ma.

Port Askaig glaciationEdit

The Port Askaig glaciation is above the Elbobreen-Wilsonbreen glaciation.

Elbobreen-Wilsonbreen glaciationEdit

The Elbobreen-Wilsonbreen glaciation in Svalbard occurred c. 720 Ma.

Cryogenian ice ageEdit

Earth is depicted during the Cryogenian as a snowball. Credit: たけまる.{{free media}}

The Cryogenian Ice Age, or the Stuartian-Varangian Ice Age, a "Late Proterozoic ice age was apparently the greatest of all. Glacial strata are known from all modern continents (except Antarctica) with an overall time range of about 950 to 600 million years old. Glacial strata from several intervals during this time are well preserved in Africa, China, Australia, Europe, Arabia, North America, and elsewhere. Multiple glaciations are the rule. In Scotland and Ireland, for example, three glacial episodes took place between 700 and 580 million years ago (McCay et al. 2006)."[9]

It apparently consists of

  1. glaciation of the Lower Congo region, Africa occurring 950-750 and 620-600 Ma,
  2. Stuartian glaciation, Australia, 800-780 Ma,
  3. Sinian glaciation, China, 800-760, 740-700, and 600 Ma,
  4. glaciation in Western Canada/U.S.A., 850-800 Ma,
  5. glaciation of the Saharan region, Africa, 730-650 Ma,
  6. Marinoan glaciation, Australia, 690-680 Ma, and
  7. Varangian glaciation, Norway, about 650 Ma.[9]

"Late Proterozoic glaciogenic deposits are known from all the continents. They provide evidence of the most widespread and long-ranging glaciation on Earth."[136]

Def. "a geologic period within the Neoproterozoic era from about [720] to 600 million years ago"[147] is called the Cryogenian.

Apparently two major glacial periods occurred during the Cryogenian: the Marinoan and the Sturtian,[135][123] formerly considered together as the Varanger glaciations, from their first detection in Norway's Varanger Peninsula.

The Cryogenian is a geologic period that lasted from 720-635 Mya.[148]

The Cryogenian period was ratified in 1990 by the International Commission on Stratigraphy.[149]

Several glacial periods are evident, interspersed with periods of relatively warm climate, with glaciers reaching sea level in low paleolatitudes.[136]

The deposits of glacial tillite also occur in places that were at low latitudes during the Cryogenian, a phenomenon which led to the hypothesis of deeply frozen planetary oceans called "Snowball Earth".[150][151]

"Most Neoproterozoic glacial deposits accumulated as glacially influenced marine strata along rifted continental margins or interiors."[136]

The base of the period is defined by a fixed rock age, that was originally set at 850 million years,[152] but changed in 2015 to 720 million years.[148]


The Sturtian glaciation was a glaciation, or perhaps multiple glaciations,[153] during the Cryogenian Period.[135][123]

The break up along the margins of Laurentia at about 750 Ma occurs at about the same time as the deposition of the Rapitan Group in North America, contemporaneously with the Sturtian in Australia.[136]

The Sturtian glaciation persisted from 720 to 660 million years ago.[137]

A Sturtian age was assigned to the Numees diamictites.[154]

The duration of the Sturtian glaciation has been variously defined, with dates ranging from 717 to 643 Ma.[155][156][153] Or, the period spans 715 to 680 Ma.[157]

"Glaciogenic rocks figure prominently in the Neoproterozoic stratigraphy of southeastern Australia and the northern Canadian Cordillera]. The Sturtian glaciogenic succession (c. 740 Ma) unconformably overlies rocks of the Burra Group."[136]

The Sturtian succession includes two major diamictite-mudstone sequences, which represent glacial advance and retreat cycles, stratigraphically correlated with the Rapitan Group of North America.[136]

The Sturtian is named after the Sturt River Gorge, near Bellevue Heights, South Australia.

Reusch's Moraine in northern Norway may have been deposited during this period.[158]

"Diamictites of the Negash Formation atop the Mariam Bohkahko Formation previously were reported only within the core of the Negash Syncline [...], but new mapping has led to the discovery of exposures of the diamictite and underlying strata near the town of Samre [...]. These exposures significantly expand known exposure of the formation, which is dominated by matrix-supported diamictite with intervals of conglomerate and sandstone. The contact with the underlying Mariam Bohkahko formation is distinct but seemingly conformable. However, in isolated areas within the carbonate-dominated facies, a carbonate conglomerate occurs along the contact, indicating some erosion of the carbonate lithologies did occur and some time may be missing. Clasts within the diamictite include carbonate lithologies likely sourced from the Tambien Group, volcanic lithologies likely sourced from the basement arc volcanics, and other lithologies such as granite and felsic gneiss that are likely extra-basinal. Striated clasts can be found within the diamictite, which along with stratigraphic arguments, have led to a glacigenic interpretation for the unit (e.g., Miller et al., 2003). The additional exposure near Samre has presented an opportunity to develop further chemostratigraphic and geochronologic constraints on the interval immediately preceding the Sturtian Glaciation."[159]

"Two samples (SAM-ET-03 and SAM-ET-04) of light-colored ~25-cm-thick tuffaceous siltstones in the upper Mariam Bohkahko Formation were collected. [Both] samples show distinct age clusters at ca. 719 Ma that overlap with uncertainties of ~1 m.y. [...]. Weighted mean ages of 719.68 ± 0.46 Ma (n = 8) and 719.58 ± 0.56 Ma (n = 3) were calculated for samples SAM-ET-04 and SAM-ET-03, respectively, using these age clusters [...]. [The] 719.68 ± 0.46 Ma and 719.58 ± 0.56 Ma dates [are likely] eruptive ages that constrain the depositional age of the strata [...]."[159]

"The two dates near the contact with the Negash Formation diamictite confirm the interpretation that the Negash diamictite is Sturtian in age as originally proposed by Beyth et al., (2003). These dates provide new constraints on the timing of initiation of the Sturtian glaciation in the Arabian-Nubian Shield to be after ca. 719 Ma."[159]


The Numees has a Sturtian age.


"The two glacial episodes that define the Cryogenian, the Sturtian and Marinoan, are both preceded by large negative 𝛅13
isotope excursions that have been identified in numerous sections globally (Halverson et al., 2005; Prave et al., 2009)."[159]

Preceding "both the Sturtian and Marinoan Snowball events, carbon isotopes recover from deeply negative values prior to glacial conditions (Halverson et al., 2005; Prave et al., 44 2009; Rose et al., 2012)."[159]


The Tereeken occurred < 727 ± 8 Ma.

Rapitan glaciationEdit

"The Rapitan Group (Cryogenian) of western Canada is similar to the Chuos Formation in both lithofacies and basin context, representing deposition in a paraglacial rift basin (Young, 1976; Eisbacher, 1985). An iron-rich, dropstone-bearing unit (the Sayunei Formation) is capped by a diamictite unit (the Shezal Formation) (Hoffman and Halverson, 2011). Measured sections (Fig. 3 of Eisbacher, 1985) illustrate that the most complete successions have a basal ferruginous shale sequence bearing occasional dropstones. These deposits pass gradationally upward, via 5–40 m jaspillite-hematite ironstone at the top of the Sayunei Formation, into diamictites. The ironstone is laterally persistent in depocentres (Eisbacher, 1985). Sea-ice removal may have triggered local grounding line advance, resulting in deposition of the Shezal Formation (Eisbacher, 1985): Hoffman and Halverson (2011) recognised this as a possible catalyst for ironstone precipitation. In addition to an abiotic “rusting of the seas” model, a biologically-mediated mechanism was also considered. Once “the ice cover thinned and finally disappeared, anoxic and oxygenic photosynthesis could have precipitated Fe2O3-precursor from anoxic Fe(II)-rich basin waters” (Hoffman et al., 2011). [...] Such a biogenic mechanism for ironstone precipitation, via for example photosynthetic stromatolites, would be in agreement with our observations in Namibia."[160]

Port NollothEdit

The Port Nolloth extends from the Kaigas formation upwards to the Murmees.

Kaigas formationEdit

The Kaigas glaciation was a hypothesized snowball earth event in the Neoproterozoic Era, preceding the Sturtian glaciation inferred based on the interpretation of Kaigas Formation conglomerates in the stratigraphy overlying the Kalahari Craton as correlative with pre-Sturtian Numees formation glacial diamictites;[161] however, the Kaigas formation was later determined to be non-glacial, and a Sturtian age was assigned to the Numees diamictites.[162]


The Vendian occurred about 740 Ma.

The Vendian concept was formed stratigraphically top-down, and the lower boundary of the Cambrian became the upper boundary of the Vendian.[163][164]

The Vendian in its type area consists of large subdivisions such as Laplandian, Redkino, Kotlin and Rovno Regional stages with the globally traceable subdivisions and their boundaries, including its lower one.

Chuos glaciationEdit

"The "grainstone prism" was a major submarine drainage system localized in a paleovalley carved during the Chuos glaciation, which was occupied by a transverse ice-stream that cut the Duurwater trough during the Ghaub glaciation."[128]

"Despite early indications of two distinct glaciations (Kröner and Rankama, 1972; Guj, 1974), the prevailing view of a single glaciogenic horizon that could serve as a basis for correlation throughout the Otavi Group (Hedberg, 1979; SACS, 1980; Miller, 1997) led to the former "Otavi Tillite" (le Roex, 1941) being assigned to the Chuos Formation of Gevers (1931), a glaciogenic diamictite with an intimately associated banded iron formation that is widely distributed within the orogens bounding the Otavi platform (Martin, 1965a, 1965b). More recently, two glaciations have been firmly established in the Otavi Group (Hoffmann and Prave, 1996; Hoffman et al., 1998a; Hoffman and Halverson, 2008), the older Chuos Formation and a younger glaciation represented by the "Otavi Tillite" (le Roex, 1941), and its correlative carbonate-clast breccia unit of the Fransfontein homocline (Frets, 1969; Guj, 1974). Hoffmann and Prave (1996) renamed this younger glaciogenic unit the Ghaub Formation, after a farm near the section originally described by le Roex (1941)."[128]

The "Chuos glaciation occurred during a period of active faulting, which is reflected by the diversity of its debris and a low-angle (1.5°) structural unconformity [...] that cuts out ~2 km of strata (Hoffman et al., 1998a)."[128]

The Rasthof Formation [is] the postglacial cap carbonate overlying the Chuos diamictite".[128]

Below the Chuos glaciation is the Naauwpoort dated at 746 ± 2 Ma giving an upper age limit to the base of the Chuos.[128]

"U–Pb ages from the Askevold Formation (Hoffman et al., 1996) [Nabis Formation 747 ± 2 Ma (Hoffman et al., 1996)] are from further west: this formation is not preserved beneath the Chuos Formation in [the Ghaub and Varianto farm areas of the Otavi Mountain Land]."[160]

"Earlier analyses of the Chuos Formation concentrated on meta-sediments in the vicinity of its type section south of Windhoek and in the Damara Belt (Gevers, 1931; de Kock and Gevers, 1933; Martin et al., 1985; Henry et al., 1986; Badenhorst, 1988). More modern stratigraphic analyses several hundred kilometres to the west of the Otavi Mountain Land demonstrate that the Chuos Formation is cradled in a rift-related, fault bounded palaeotopography (Hoffman and Halverson, 2008), and hence its substrate also changes along strike, across the southern flank of the Owambo Basin. In the area of Ghaub and Varianto farms, the study interval comprises the Nabis Sandstone Formation of the Nosib Group, overlain by the Chuos Formation and succeeded by the Berg Aukas Formation [...]. This particular area has been mapped at the 1:250,000 scale (Geological Survey of Namibia, 2008). Age constraints include 747 ± 2 Ma from the Naauwport volcanics, locally beneath the Chuos Formation (Hoffman et al., 1996) and 635 ± 1 Ma from ash beds in the younger Ghaub Formation (Hoffmann et al., 2004)."[160]

Beiyixi glaciatonEdit

The Beiyixi is later than 755 Ma.

Makganyene glaciationEdit

"In its eastern domain, the Transvaal Supergroup of South Africa contains two glacial diamictites, in the Duitschland and Boshoek Fms. The base of the Timeball Hill Fm., which underlies the Boshoek Fm., has a Re-Os date of 2,316 ± 7 My ago (13). The Boshoek Fm. correlates with the Makganyene diamictite in the western domain of the Transvaal Basin, the Griqualand West region. The Makganyene diamictite interfingers with the overlying Ongeluk flood basalts, which are correlative to the Hekpoort volcanics in the eastern domain and have a paleolatitude of 11° ± 5° (14). In its upper few meters, the Makganyene diamictite also contains basaltic andesite clasts, interpreted as being clasts of the Ongeluk volcanics. The low paleolatitude of the Ongeluk volcanics implies that the glaciation recorded in the Makganyene and Boshoek Fms. was planetary in extent: a snowball Earth event (15). Consistent with earlier whole-rock Pb–Pb measurements of the Ongeluk Fm. (16), the Hekpoort Fm. contains detrital zircons as young as 2,225 ± 3 My ago (17), an age nearly identical to that of the Nipissing diabase in the Huronian Supergroup."[165]

The "Makganyene glaciation begins some time after 2.32 Ga and ends at 2.22 Ga, the three Huronian glaciations predate the Makganyene snowball."[165]

Huronian ice ageEdit

Proposed correlation is of the Huronian Supergroup and the upper Transvaal Supergroup. Credit: Robert E. Kopp, Joseph L. Kirschvink, Isaac A. Hilburn, and Cody Z. Nash.{{fairuse}}
The Earth is depicted during Huronian Glaciation. Credit: Oleg Kuznetsov.{{free media}}

The Huronian Ice Age is known "mainly from Canada and the United States in North America, where dated rocks range from 2500 to 2100 million years old. The Gowgonda Formation of Ontario is especially noteworthy for its excellent preservation of glaciogenic strata dated about 2300 million years old. Other glacial deposits are found in Wyoming, Michigan, Quebec, and the Northwest Territories. These rocks record extensive Early Proterozoic continental glaciation through a time span of about 400 million years, during which three or more glacial expansions took place. The configuration of the continents during this time is highly speculative."[9]

Gowganda glaciationEdit

"The period from 2.45 Ga until some point before 2.22 Ga saw a series of three glaciations recorded in the Huronian Supergroup of Canada (11) [in the above centered image]. The final glaciation in the Huronian, the Gowganda, is overlain by several kilometers of sediments in the Lorrain, Gordon Lake, and Bar River formations (Fms.). The entire sequence is penetrated by the 2.22 Ga Nipissing diabase (12); the Gowganda Fm. is therefore significantly older than 2.22 Ga."[165]

"The three Huronian glacial units, penetrated and capped by the Nipissing diabase, predate the Makganyene diamictite in the Transvaal. The uppermost Huronian glacial unit, the Gowganda Fm., is overlain by hematitic units, perhaps reflecting a rise in O2. The basal Timeball Hill Fm. contains pyrite with minimal [mass-independent fractionation] MIF (26), whereas the upper Timeball Hill Fm., which we suggest is correlative to the Lorrain or Bar River Fms., contains red beds. The Makganyene diamictite records a low-latitude, snowball glaciation (29), perhaps triggered by the destruction of a CH4 greenhouse. It is overlain by the Kalahari Mn Field in the Hotazel Fm., the deposition of which requires free O2."[165]

"In its eastern domain, the Transvaal Supergroup of South Africa contains two glacial diamictites, in the Duitschland and Boshoek Fms. The base of the Timeball Hill Fm., which underlies the Boshoek Fm., has a Re-Os date of 2,316 ± 7 My ago (13). The Boshoek Fm. correlates with the Makganyene diamictite in the western domain of the Transvaal Basin, the Griqualand West region. The Makganyene diamictite interfingers with the overlying Ongeluk flood basalts, which are correlative to the Hekpoort volcanics in the eastern domain and have a paleolatitude of 11° ± 5° (14). In its upper few meters, the Makganyene diamictite also contains basaltic andesite clasts, interpreted as being clasts of the Ongeluk volcanics. The low paleolatitude of the Ongeluk volcanics implies that the glaciation recorded in the Makganyene and Boshoek Fms. was planetary in extent: a snowball Earth event (15). Consistent with earlier whole-rock Pb–Pb measurements of the Ongeluk Fm. (16), the Hekpoort Fm. contains detrital zircons as young as 2,225 ± 3 My ago (17), an age nearly identical to that of the Nipissing diabase in the Huronian Supergroup. As the Makganyene glaciation begins some time after 2.32 Ga and ends at 2.22 Ga, the three Huronian glaciations predate the Makganyene snowball."[165]

"In contrast to the Makganyene Fm., the three Huronian diamictites are unconstrained in latitude. Poles from the Matachewan dyke swarm, at the base of the Huronian sequence, do indicate a latitude of ≈5.5° (18), but ≈2 km of sedimentary deposits separate the base of the Huronian from the first glacial unit (19), which makes it difficult to draw conclusions about the latitude of the glacial units based on these poles. Low latitude poles in the Lorrain Fm. (20, 21), which conformably overlies the Gowganda diamictite, are postdepositional overprints (22)."[165]

"Some of the earliest continental red beds were deposited in the Firstbrook member of the Gowganda Fm. and in the Lorrain and Bar River Fms. in Canada, as well as in the upper Timeball Hill Fm. in South Africa. The basal Timeball Hill Fm. has recently been dated at 2,316 ± 7 My ago (13). In our proposed correlation, all of the red bed-bearing units were deposited after the last Huronian glaciation and before the Makganyene glaciation. The formation of the red beds could involve local O2, although it does not demand it (34). Syngenetic pyrite from the basal Timeball Hall Fm. shows only slight MIF of S (26), consistent with the initiation of planetary oxygenation or enhanced glacial activity."[165]

There are two diamictites, one at the base and the other at the top of the Gowganda Formation.[166]

Upper Gowganda FormationEdit

"The presence of dropstones and micro-laminae within the mechanically and geochemically unaltered rhythmite sequences of the interbedded siltstone and claystone lithofacies supports the conclusion that the rhythmites represent varve deposits, confirming that the upper Gowganda Formation represents some of the oldest glaciogenic deposits found on Earth."[167]

"The Gowganda Formation is composed of two members: the lower Coleman Member, containing laminated paraconglomerate and mudstone with dropstones, and the upper Firstbrook Member, composed largely of interbedded sandstone and mudstone (Junnila and Young, 1995). Overall, at least two major glacial advances are recognized in the Gowganda Formation, and deposition is interpreted to have occurred beneath a continental ice sheet with periods of deep-water conditions during glacial retreat (Coleman member), followed up-section by a prodeltaic setting (Firstbrook Member) (Young and Nesbitt, 1985; Long and Leslie, 1986)."[56]

Coleman glaciationEdit

The "glaciogenic deposystem for the basal 180 m of the Coleman member, lowermost unit of the Gowganda Formation, Huronian Supergroup [occurred on the subsiding shelf of a south-facing passive margin.] Massive basal diamictites were deposited subglacially as extensive primary basal till and as restricted lee-side till that infilled paleotopographic lows and covered preglacial regolith."[168]

Bruce glaciationEdit

"The second Huronian glaciation is represented by the Bruce diamictite and is correlated with the Vagner diamictite in the Snowy Pass Supergroup (Wyoming) [49]."[166]

"Sandstones and arkoses of the Mississagi Fm. are succeeded by the glacially derived Bruce Fm., which contains dropstones [26] and is immediately overlain by a carbonate with carbon isotope characteristics that broadly resemble those described for Neoproterozoic cap carbonates [17]. Sedimentary rocks of the post-glacial Espanola Fm. constitute a transition from a lower carbonate–limestone member to a higher siltstone-heterolithic member, which is interpreted to have formed in a shallow-marine or restricted lacustrine environment possibly during a period of active continental fragmentation [15,27–29]. Rare columnar stromatolite features have also been reported in the Espanola carbonate [30]."[166]

Ramsey Lake glaciationEdit

"The glacial diamictite of the Ramsey Lake Formation [...] contains dropstones in siltstone interbeds, implies deposition in a glaciomarine setting at an ice margin."[166]

"A recent study of the organic geochemistry of the pre-glacial Matinenda Fm. has documented oil droplets with relatively high levels of 2α-methylhopanes and other biomarkers interpreted to be eukaryotic in origin [22]. It was also suggested that the oil probably migrated in the Matinenda Fm. from the McKim Fm. [which underlays the Ramsey Lake Formation] during post-depositional processes and consider unlikely the possibility that these cyanobacterial biomarkers are indigenous to the Matinenda Fm."[166]

Pongola glaciationEdit

General stratigraphy of the Mozaan Group of the Pongola Supergroup. Credit: Shuhei Ono, Nicolas J. Beukes, Douglas Rumble, and Marilyn L. Fogel.{{fairuse}}
Microscope photo of the sample PNG2-657.2 (plane polarized light). Credit: Shuhei Ono, Nicolas J. Beukes, Douglas Rumble, and Marilyn L. Fogel.{{fairuse}}

The Pongola glaciation is dated "at 2.9 Ga".[165] It extends to 2780 Ma.

"The oldest known midlatitude glaciation, recorded in the Pongola Supergroup diamictite, occurred at 2.9 Ga (10)."[165]

"As the geological expression of oxygen does not appear during the ~2.8 Ga Pongola glaciation or during the Huronian glaciations, when glacial weathering should have elevated these fluxes, oxygenic cyanobacteria may not have evolved and radiated until shortly before the Makganyene Snowball."[169]

The "Pongola glaciation in Southern Africa at 2.9—2.7 Gyr ago has at least three distinct units of diamictite that can be traced laterally over hundreds of square kilometers and that contain occasional small drop stones."[170]

"Sulfur isotope mass-independent fractionation (S-MIF) is a unique geologic record of Archean atmospheric chemistry that provides important constraints on the evolution of the early Earth’s atmosphere and its impact on early life."[171]

"[M]ultiple-sulfur (33
, 34
, and 36
) isotope ratios of sulfide minerals and carbon (13
) isotope ratios of organic carbon for shale in the ~2.96 to ~2.84 Ga Mozaan Group of the Pongola Supergroup, Southern Africa [have been measured]."[171]

On the right is a stratigraphic column for the general "stratigraphy of the Mozaan Group of the Pongola Supergroup (left) and the location of the Pongola (hatched area) and Witwatersrand (grey) basins (top right). Arrows indicate the stratigraphic location of samples used for this study. The figure is modified from Beukes and Cairncross (1991) and Nhleko (2003)."[171]

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