Late Pliocene-Early Pleistocene Ecologic Changes 

in the Arctic Ocean Borderland 

I 

U.S. GEOLOGICAL SURVEY BULLETIN 2036 
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Ave. 
Late Pliocene-Early Pleistocene Ecologic Changes 

in the Arctic Ocean Borderland 

By CHARLES A. REPENNING and ELISABETH M. BROUWERS 

U.S. GEOLOGICAL SURVEY BULLETIN 2036 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON: 1992 
U.S. DEPARTMENT OF THE INTERIOR 

MANUEL LUJAN, JR., Secretary 

U.S. GEOLOGICAL SURVEY 

Dallas L. Peck, Director 

For sale by the Books and Open-File Reports Section 

U.S. Geological Survey 

Federal Center, Box 25286 

Denver, CO 80225 

Any use of trade, product, or finn names in this publication is for descriptive purposes only and 

does not imply endorsement by the U.S. Government 

Library of Congress Cataloging-in-Publication Data 

Repenning, Charles Albert, 1922-

Late Pliocene-Early Pleistocene ecologic changes in the Arctic Ocean Borderland I by Charles A. Repenning and Elisabeth M. 

Brouwers. 

p. cm.-(U.S. Geological Survey bulletin ; 2036) 

Includes bibliographical references. 

Supt. of Docs. no.: I 19.3:2036 

1. Paleoecology-Arctic regions. I. Brouwers, Elisabeth M. II. Title. III. Series. 

QE75.B9 no. 2036 

[QE720] 

557.3 s-dc20 92-23139 

[560'.45] CIP 
CONTENTS 

Abstract................................................................................................................................. 1 

Introduction........................................................................................................................... 2 

Age Assignments........................................................................................................ 5 

Isotope and Polarity Records...................................................................................... 6 

Aoral Terms................................................................................................................ 6 

3.0--2.0 Million Years Ago: The First Cooling Cycle.......................................................... 8 

Canada, Meighen Island: Beaufort Formation............................................................ 8 

Alaskan North Slope: Gubik Formation..................................................................... 9 

Colvillian and Bigbendian Transgressions ........................................................... 9 

Fishcreekian "Transgression"............................................................................... 10 

Yakutia: Begunov Suite and Kutuyakh Beds....................................................... 13 

Marine and Proxy Records of the First Cooling Cycle......................................................... 15 

Norwegian Sea: Scandinavian Ice Sheet.................................................................... 15 

Central Arctic Ocean: Eastern Alpha Ridge............................................................... 17 

Western Arctic Ocean: Mendeleyev Ridge . .............. .. ... .......... ..... ....... .. .... .. ............ .. 18 

Summary of the First Cooling Cycle.................................................................................... 18 

2.0--0.85 Million Years Ago: The Second Cooling Cycle.................................................... 20 

Greenland: Kap KS?Sbenhavn Formation................................................................................ 20 

Y akutia: Krestovka Olyor Suite.................................................................................. 22 

Yakutia: Type Olyor Suite.......................................................................................... 24 

Alaska: Cape Deceit Section ........ .. . . ......... .. .. . . . .. ..... ....... ...... ............ ...... ....... .. ........... 25 

Canada: Yukon Territory-Locality CRH-47 ........................................................... 27 

Marine and Proxy Records of the Second Cooling Cycle ........ ... .. . .. . . . . . ...... ... . . .. ..... ..... ........ 27 

Norwegian Sea: Scandinavian Ice Sheet.................................................................... 28 

Microtine Immigrants: Western United States..................................................................... 28 

First Cooling Cycle..................................................................................................... 29 

Beginning of the Second Cooling Cycle . .. ..... .. ....... ........... ................ ............ .......... .. 29 

Eburonian Cold........................................................................................................... 30 

Beginning of the Ice Ages .. ..... ............. .. . . ........ .. ..... ...... ...... ....... ... .. . . ............ ......... .. .. 31 

Summary: Late Pliocene-Early Pleistocene Ecologic Change............................................ 32 

References Cited................................................................................................................... 34 

FIGURES 

1. Map showing localities discussed in text........................................................................................................................ 3 

2. Diagram showing correlation of localities and events discussed in text ............. .............. ... . .. . .... ...... ...... .... .................. 7 

TABLE 

1. Pollen from the Bigbendian deposits, Alaskan North Slope, compared with pollen from the Begunov Suite, Yakutia 13 

III 

LATE PLIOCENE-EARLY PLEISTOCENE 

ECOLOGIC CHANGES IN THE 

ARCTIC OCEAN BORDERLAND 

By CHARLES A. REPENNING1 and ELISABETH M. BROUWERS1 

ABSTRACT 

Between 3.0 and 0.85 million years ago the Arctic 

Ocean Borderland experienced two cycles of climatic deterioration, both of which ended in major Northern Hemisphere continental glaciation. The first cooling cycle ended 

2.1 million years ago; continental ice extended down the 

Mississippi River Valley as far south as Iowa. The cooling 

cycle was followed by a collapse of glaciation and an 

extremely warm climate that extended to the northernmost 

land, almost to the North Pole. This warmth at the beginning of the second cooling cycle was greater than that 

known at the beginning of the first cycle and cooling began 

again 300,000 years later, late during the Olduvai Normalpolarity Subchron and about 1.7 million years ago. The 

second cooling cycle culminated with the classic Ice Ages, 

beginning 850,000 years ago. 

At the beginning of the first cooling cycle, forests, 

dominated by mixed conifers and including minor but 

moderately diverse deciduous broadleaf trees, bordered a 

warm-subfrigid Arctic Ocean. Treeline at lat 79°39' N. 

indicated an arctic climate remarkably warm by comparison with today's conditions but not so warm as the climate 

at the beginning of the second cooling cycle. Two successive terrestrial records in Siberian Beringia, four separate 

but sequential marginal-marine records in Alaskan Beringia and northern Canada, and core records from the Arctic 

Ocean indicate that between 3.0 and 2.4 million years ago 

the Arctic Ocean Borderland progressively changed to a 

conifer forest-tundra ecosystem, then to fully developed 

tundra. With these floral changes came the first record of 

permafrost and a frigid Arctic Ocean. 

Two Alaskan and one Arctic Canadian shallowmarine records indicate that the Arctic Ocean lacked complete ice cover at least until 2.4 million years ago. Faunal 

1

U.S. Geological Survey, Denver, Colorado 80225. 

and depositional records from the central part of the Arctic 

Ocean indicate that this lack of ice cover may not have 

extended to the pole; from about 2.48 until about 2.1 million years ago perennial ice cover is thought to have 

formed intermittently in the central Arctic Ocean. There 

are no known terrestrial records in the Arctic Ocean Borderland for much of the time of this first intermittent perennial sea ice; however, both sediments and fauna from the 

beginning and from the end of this period indicate that ice 

cover did not extend to the shore. 

Continental glaciation began in both Scandinavia and 

Laurentia almost 2.6 million years ago while the Arctic 

Ocean Borderland was still being deforested and the first 

permafrost formed in northeastern Siberia and Alaska, but 

before the first intermittent perennial ice developed in the 

central ocean area. Continental glaciation in the Northern 

Hemisphere was most extensive between 2.2 and 2.1 million years ago; in North America ice then flowed down the 

lowlands of the Mississippi River as far south as Iowa. 

Two and a half million years ago, shortly after ice 

began to accumulate in Scandinavia and Laurentia, depositional records in the Arctic Ocean indicate an abrupt 

increase in ice-rafted debris, presumably from more extensive mountain glaciation, and a change to prominent density stratification, possibly caused by the influx of greater 

amounts of freshwater. Arctic marine depositional and faunal records between 2.4 and 2.2 million years ago suggest 

that the terrestrial climate of the Arctic Ocean region fluctuated. After glacial intensity peaked about 2.1 million 

years ago, a variety of evidence from the Arctic Ocean, 

North America, Scandinavia, and northwestern Europe 

indicates that Northern Hemisphere glacial conditions 

ended abruptly, thus ending the first cooling cycle. 

With the disappearance of glacial climate, mixed 

coniferous forests returned to the shores of an unfrozen 

Arctic Ocean, and about 2.0 million years ago treeline 
2 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

extended as far north as the northernmost point of Greenland, lat 82°30' N. and about 500 mi (800 km) south of the 

North Pole. Faunal and floral records at lower latitudes in 

North America and Europe indicate warmer temperatures 

and increased precipitation at this time. 

The beginning of the second cycle of cooling is not 

sharply marked. Near the end of the Olduvai Normalpolarity Subchron (about 1.7 m.y. ago), proxy evidence in 

Europe and North America indicates that the extreme 

warmth in the Arctic Ocean Borderland had ended. Faunal, 

floral, and depositional records from the Arctic Ocean, 

Europe, and North America indicate an oscillation of cool 

and warm periods between 1.7 and 1.2 million years ago. 

Both Arctic and Atlantic oceanic temperatures suggest that 

average conditions were similar to those of today, but the 

Arctic Ocean was not perennially ice covered. 

The severity of the cold periods increased, and about 

1.1 million years ago the Arctic Ocean again began to have 

periods of perennial ice cover. A cold and dry climate, 

more severe than today, is indicated in western Beringia at 

this time, and the abundance of arboreal plants was at a 

minimum. This cold climate was recorded as far south as 

the Netherlands; however, less than 100,000 years later, 

during the Jaramillo Normal-polarity Subchron, the pollen 

spectrum of eastern and western Beringia was similar to 

that of today, if not somewhat warmer. About 850,000 

years ago, ice again flowed down the Mississippi River 

Valley, beginning the Ice Ages in the United States, and 

bottom sediments in the Arctic Ocean indicate that perennial ice cover was more frequent. Although the Ice Ages 

began 850,000 years ago, climatic deterioration apparently 

did not reach its extreme until about 400,000 years ago. 

INTRODUCTION 

This synthesis of ecologic changes in the Arctic 

Ocean Borderland during the late Pliocene and early Pleistocene draws upon published information for Europe westward across North America to eastern Siberia and for areas 

to the south where information has a bearing on the interpretation of high-latitude sites. Our discussion concentrates, however, on information from Beringia, a vast area 

covering the Arctic Ocean Borderland in northeastern 

Siberia and northwestern North America (fig. 1). 

More or less by definition Beringia has never been 

covered by continental glaciation. The region thus extends 

eastward almost to the Mackenzie River and its delta in 

northern Northwest Territories, Canada, the western limit 

of continental glaciation; the western limits of Beringia are 

beyond the Kolyma Lowlands of Y akutia but have never 

been clearly defined. 

A.V. Sher (1986) applied the name "Beringida" to this 

area, apparently restricting the name "Beringia" to the area 

immediately adjacent to the Bering Sea and the Bering 

Strait, but his failure to provide a geographic description 

makes the synonymy uncertain. He also provided the 

"Beringidian" mammalian age names, the Kutuyakan, the 

Capedeceitan, the Olyoran, the Akanan, and the "upper 

Paleolithic of Siberia"; these lack typification and characterization and also suffer from a lack of geographic constraints, and they are not used herein. The need for age 

names to aid discussion of these extremely provincial 

Arctic Ocean Borderland mammalian faunas is discussed 

by Repenning (in press), although he does not use Sher's 

terms either. 

Throughout the Tertiary and Quaternary, both northwestern North American and Siberian Beringia were a single faunal and floral province; minor east-west differences 

resulted from ecologic restrictions created by either the 

Bering Strait or the Bering Land Bridge. About three million years ago the Bering Strait opened (Gladenkov, 1981) 

and separated eastern and western Beringia, possibly for 

the first time since the mid-Cretaceous (Marincovich and 

others, 1990), although recent marine evidence from Kamchatka (Gladenkov and others, 1991) indicates that there 

may have been an earlier opening about 4.2 million years 

ago. The Bering Strait has been sporadically present since 

3.0 million years ago. 

Evidence from the Arctic and North Atlantic Oceans 

indicates that the opening of Bering Strait 3 million years 

ago introduced North Pacific marine invertebrates to the 

Arctic and North Atlantic Oceans. These Pacific invertebrates provide a useful, but somewhat imprecise, maximum 

age wherever found in deposits of the Arctic Ocean. The 

use of Pacific invertebrates for other correlation is limited 

by problems in identifying and dating faunal events. The 

first appearance of Pacific taxa in Arctic and North Atlantic deposits marks the oldest marine events discussed in 

this synthesis. 

This opening of the Bering Strait is not so clearly 

marked in terrestrial deposits; it occurred during climatic 

deterioration of the Northern Hemisphere that began about 

3.5 million years ago according to oxygen isotope records. 

Nonmarine records discussed here are not that old; the earliest, from Meighen Island in Arctic Canada, are inferred to 

be between 3.0 and 2.9 million years old on the basis of 

molluscan fauna and paleomagnetic evidence (Fyles and 

others, 1991), evidence that does not convincingly exclude 

the possibility that the deposits may be between 4.3 and 3.9 

million years old. 

The age of about 3.0 million years is supported by 

Alaskan mammal faunas that cannot be this old but are 

associated with deposits apparently representing a nearcontinuous record of both marine and terrestrial climate 

change with some mollusk and ostracode species that are 

identical to those found in the Meighen Island locality; 

these Alaskan faunas shortly postdate the opening of the 

Bering Strait. The age interpretation of the Meighen Island 

deposits that is preferred by Fyles and others (1991) is also 
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4 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

preferred by us. These deposits from Meighen Island of 

Arctic Canada are the oldest known post-Bering Strait 

record in the Arctic Ocean basin. 

Our reconstruction of climatic changes in the Arctic 

Ocean Borderland ends with the initiation of the traditional 

Ice Ages, marked by the first North American ice advances 

down the Mississippi River Valley that have been called 

the "Nebraskan" glaciation. This is a term abandoned by 

Hallberg (1986); Hallberg referred all former named North 

American Quaternary glaciations before the Illinoian glaciation to the "pre-Illinoian." The first ice advance of the Ice 

Ages is generally correlated with oceanic isotope stage 22, 

the peak of which has been dated about 850,000 years ago. 

All terrestrial evidence regarding the age of this first ice 

advance of the Ice Ages is in agreement with this date. 

Some authors have selected a pre-Nebraskan (late 

Pliocene) glaciation for the beginning of the Ice Ages, but 

this definition is of recent usage and it is not followed here. 

Neither usage equates with the beginning of the Pleistocene, which was typified in Italy and, in recent literature 

(Aguirre and Pasini, 1985), is defined as having begun 

about 1.67 million years ago. 

The temporal limits of our synthesis are dictated by 

the current limitations of field evidence. Work in progress 

in the Arctic Ocean Borderland will extend this history 

prior to the opening of the Bering Strait and will also identify and precisely date localities that are younger than 

events considered in the present discussion. At present, 

however, the ages of these localities are not known with 

sufficient certainty to incorporate into a history of climatic 

change; significant climatic changes may occur in as few 

as 20,000 years. 

As herein reconstructed, the Arctic Ocean Borderland 

underwent two major cycles of climatic deterioration during the time discussed, both of which terminated with continental glaciation that developed in and flowed from the 

European Scandinavian and the North American Laurentian centers. The reason for these two preferred locations of 

major continental ice accumulation is oceanic; that is, the 

locations are the result of major latitudinal shifts in North 

Atlantic current patterns. Behind this immediate cause is a 

wide spectrum of contributing factors including adjustment 

to changes in the global oceanic density and solar energy 

input that results from orbital, atmospheric, and albedo 

differences. These factors must combine in critical relationships to trigger the onset of glaciation, and no single factor 

is the cause of glaciation on a continental scale. 

If a single cause must be named, it is the gradual rise 

of Central America between 15 and 3 million years ago. 

Repenning and Ray (1977) briefly discussed some of the 

consequences of this shoaling and its age, but the literature 

on this subject is very large. The shoaling and ultimate 

emergence of this area completely disrupted former equatorial current patterns in the World Ocean and forced north 

and south transequatorial flow to make density adjustments 

that had previously been accomplished by less complex circulatory cells within individual oceanic basins, with interbasinal differences being adjusted by equatorial currents 

between the Atlantic and the Pacific. The process of density adjustment has recently been illustrated by computer 

modeling (Maier-Reimer and others, 1990). Transequatorial flow also varies with other factors such as equatorial 

heat and geographic configurations; atmospheric climate 

varies, in turn, with oceanic conditions. 

Although the loss of equatorial mixing initiated global 

climatic deterioration, the regions most sensitive to this 

change were at polar latitudes. A warming of the North 

Atlantic produces more atmospheric moisture over adjacent 

lands such as Laurentia and Scandinavia. If these lands are 

subject to freezing temperatures, ice will accumulate. Temperatures must be below freezing for the accumulation of 

ice but need not be much below freezing. Today's minimum temperatures vastly exceed the requirement, but 

today's precipitation is not adequate for ice accumulation. 

An increase in atmospheric moisture also produces greater 

cloud cover, which in turn reduces summer ice melt. 

As ice accumulates, other factors come into play, such 

as decreased insolation and the regional development of 

high-pressure areas above the ice. These alter global climatic patterns in proportion to the size of the ice field. 

They are also self-defeating with respect to the growth of 

continental ice fields; increased solar insolation is needed 

to provide the energy needed to remove oceanic water for 

the production of an ice field and greater insolation is 

needed to allow its growth to a continental scale, but the 

existence of ice fields reduces insolation and produces 

atmospheric conditions that deflect moist air masses away 

from the ice field. Hypothetically an ice field of continental 

scale should be shaped like a doughnut because the ice 

field creates a climatic situation in which no precipitation, 

only evaporation, can reach its center and precipitation, and 

ice accumulation, is greatest at its margins. 

A better understanding of the history of climatic 

change in the Arctic Ocean Borderland is, consequently, of 

more than provincial interest. Arctic climatic history is 

reflected in global climatic history. Furthermore, it is 

expressed in a region where, because it is cold enough to 

accumulate ice, this history has had most dramatic effects. 

High-latitude climatic deterioration is a reflection of 

changes of oceanic current patterns, and therefore what 

happens in the Arctic Ocean region correlates with what 

happens in lower latitudes. This relationship is justification 

for the use of proxy data, data from extra-Arctic regions 

that are contemporaneous with Arctic events and can be 

used to interpret Arctic conditions. But, to be used in interpretation, the proxy data must correlate in time with events 

from the Arctic region. 
INTRODUCTION 5 

For more than a decade it has been generally recognized that climatic changes approximately correlate with 

orbital factors of the Earth, which influence total solar 

insolation (Hays and others, 1976); how other factors modify this influence is still a subject of much consideration. 

Orbital parameters vary over relatively brief periods of 

geologic time, the shortest of which is about 20,000 years. 

To use proxy data in reconstructing the Arctic Ocean 

Borderland, some means of temporal correlation must be 

employed that can differentiate time periods as brief as 

20,000 years. 

Because the late Pliocene through early Pleistocene (3 

million to 850,000 years ago) is beyond the useful limit of 

carbon-14 dating (about 40,000 years) and because opportunities for the use of other radiometric dating methods are 

few in the Arctic region, our ability to correlate rocks of 

this time interval in the Arctic region has not kept pace 

with our ability to correlate such rocks in the rest of the 

world. Other than paleomagnetic stratigraphy, little evidence of geologic age is available in the Arctic to support 

biochronologic age inferences. Further, biochronologies in 

the Arctic region are drastically influenced by the extreme 

latitudinal provinciality of Arctic faunas and floras. 

Attempts to project biochronologies developed in lower 

latitudes into the Arctic region have consistently met with 

the conflicting aspects of latitudinal provinciality, which 

causes the unique biotic associations of the Arctic. 

Except for pioneering work attempting to determine 

the age of the opening of the Bering Strait (Hopkins, 1967; 

Gladenkov, 1981; Gladenkov and others, 1991}, success in 

determining geologic age in the Arctic has been minimal. 

The recognition that late Pliocene and early Pleistocene terrestrial faunal events are recorded in western Beringia has 

only existed for the last 20 years (Sher, 1971) and in central and eastern Beringia only for the last 8 years (Repenning, 1983, in press; Repenning and others, 1987). 

Terrestrial faunal events, which demonstrate the passage of 

time, have been recognized in temperate latitudes for more 

then 200 years, and their records are abundant. Only seven 

or eight localities of known age recording the terrestrial 

biota of the Arctic Ocean Borderland are discussed in this 

report. 

AGE ASSIGNMENTS 

The chronology used herein is based on the biochronology of the microtine rodents of the Northern Hemisphere (Repenning and others, 1990). This biochronology 

is new and, therefore, has not been successfully applied by 

others to faunas of the Arctic Ocean Borderland. Its use is 

limited to the past five and a half million years (the time 

during which microtine rodents have been a significant part 

of the world fauna). 

Within this time span, the biochronology can distinguish time differences with a minimum accuracy of 

250,000 years in most areas of the Northern Hemisphere; 

using adequate faunal samples of the better known lineages 

of microtine rodents, faunal differences can be identified 

that are as little as 5,000-10,000 years different in age 

(Repenning, in press). This is possible because microtine 

rodents evolved very rapidly, but it must also be admitted 

that adequate faunal samples of those microtine rodents 

having significance are not yet common in the Arctic 

region. 

This time discrimination is sufficiently precise to 

identify paleomagnetic and climatic events of known age. 

Age assignments based on microtine rodent evolution and 

dispersal history can thus be refined by correlation with 

these events, and climatic events can be identified with sufficient certainty to be used in support of correlation. For 

example, the two normal-polarity excursions within the 

Reunion Normal-polarity Subchron of the Matuyama 

Reversed-polarity Chron are separated by 80,000 years,· a 

span of time too brief to be detectable in much of the presently known microtine history of that period of time. Positive climatic evidence indicates, however, that continental 

glaciation was extensive in North America during the older 

of these two excursions and that climates were warm during and following the younger excursion. If the microtine 

biochronology can identify a polarity event in an outcrop 

as being one of the two Reunion excursions, it is considered reasonable to identify which excursion is represented 

on the basis of climatic evidence. 

Mammals other than microtine rodents lived in or 

have dispersed through Beringia, but either their rates of 

evolutionary change were too slow or their history too 

imprecisely known for similarly accurate age interpretation. Correlations in the Arctic Ocean region that are based 

on marine invertebrates are basically the application of 

marine paleoclimatic stratigraphy with an often misleading 

assist from the history of the invertebrates in lower latitudes, just as floral evidence is an application of terrestrial 

paleoclimate stratigraphy. Neither biologic group has an 

evolutionary history that is well enough understood in the 

Arctic to be used in a precise biochronologic sense, 

although both are most helpful in reconstructing past environments. 

Eastern Beringia has maintained many faunal features 

of the northeastern Asian ecosystem because of its isolation 

from the rest of North America during continental glaciation. Because of increasing latitudinal differences in environment during climatic deterioration, these features were 

retained during periods in which no continental ice was 

present. Lower latitude North America had very little influence on the Beringian fauna. Because of increasing latitudinal zonation and provinciality during the late Pliocene 
6 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

and early Pleistocene, faunal influence from lower latitude 

Eurasia decreased. The Eurasian fauna filtered into Beringia and dispersed southward into North America, but very 

little of the North American fauna moved northward into 

Beringia. 

For the most part, similarities between Beringian terrestrial faunas and those of temperate regions have resulted 

from the southward dispersal of fauna during cold periods 

of the late Pliocene and early Pleistocene. That northward 

dispersal of temperate faunas into Beringia took place during warm periods is obvious, but records from the Arctic 

Ocean Borderland and from regions to the south strongly 

suggest that these elements became locally extinct with the 

return of more severe climates. This ecological constraint 

has made Beringia a "filter" dispersal route (Simpson, 

1940). The record of the passage of mammalian taxa 

through Beringia is mostly that of their appearance in temperate North America or Eurasia. Commonly there is no 

record of the dispersing organism in Beringian faunas 

because its passage through this filter dispersal route was 

brief and its discovery in the fossil record is, therefore, 

very fortuitous (Repenning, in press). 

The biohistorical aspects of Arctic North America 

must be considered provincial; they resulted from different 

dispersal and evolutionary histories, as well as from different ecosystems, than those that produced the faunal history 

of the United States. If "North American" mammal ages 

are applied to Beringian faunas, they must be redefined. 

Logically, an immigrant mammal, whose first appearance 

from higher latitudes defines a particular age in the contiguous United States, will be characteristic of an older age in 

the higher latitudes from whence it came. One should 

expect an Arctic fossil mammal fauna that seems characteristic of a North American Provincial Mammalian Age to 

represent an older age in Beringia. And, in fact, Arctic faunas in both the New and Old Worlds have proven to be 

older than conventional faunal definitions indicate. 

The North American Provincial Mammalian Ages of 

Wood and others (1941) were originally considered as representing the North American Tertiary mammalian fauna. 

They have come to be considered also as nonmarine biochronologic units, although Wood and others specifically 

included marine beds in which mammals are found. Possibly the last usage of such colonial thinking in the United 

States, the "provincial" part of their nomenclature was in 

distinction from "European ages," which were, in 1941, 

difficult to correlate to the United States. Although it 

would be unreasonable to think that they did not believe 

such existed, Wood and others (1941) expressed no awareness of latitudinal provincial distinctions, and they included 

Mexican and southern Canadian faunas and formations in 

their correlations. Their ages were assumed to be North 

American in scope but were defined primarily on the basis 

of the fossil faunas of the conterminous United States. 

Consequently, although mentioned herein a few times 

in explanation, the North American provincial mammalian 

ages of Wood and others (1941) are not used in this 

discussion of the Arctic Ocean Borderland and the Beringian faunal province; faunal ages are given only in radiometric years. 

ISOTOPE AND POLARITY RECORDS 

The oceanic cores used for correlation in this report 

are (1) site 552A (Shackleton and others, 1984; reevaluated by Raymo and others, 1989) on the western side 

of the Rockall Plateau south of Iceland; (2) site 644 

(Jansen and others, 1988) on the Voring Plateau in the Norwegian Sea off of Norway (fig. 1); (3) site CESAR-14 

(Scott and others, 1989) near the eastern end of the Alpha 

Ridge in the central Arctic Ocean (fig. 1); and (4) site T3-

67-12 (Herman and others, 1989) on the Mendeleyev 

Ridge off the Siberian Shelf in the western Arctic Ocean 

(fig. 1). Although used only occasionally, the numbers 

applied to oxygen isotope stages are those used by Ruddiman and others ( 1986) and extended by Raymo and others 

(1989); stage numbers mentioned in the text are shown in 

figure 2. 

We accept most of the radiometric age data for the 

late Pliocene and Pleistocene polarity events from the original presentation by Mankinen and Dalrymple (1979) 

because the biochronology here used is ( 1) based on radiometric age determinations rather than presumed constant 

rates of sea-floor spreading, which have been used to modify radiometric ages, and (2) was constructed on the 

Mankinen and Dalrymple calibration of the paleomagnetic 

stratigraphy. 

Several age assignments differ from those in 

Mankinen and Dalrymple. (1) The beginning of the Olduvai Normal-polarity Subchron of the Matuyama Reversedpolarity Chron is adjusted to 1.9 million years ago because 

the very securely dated KBS Tuff of Kenya is 1.88±0.02 

million years old, and this tuff lies at least 20 m stratigraphically above the base of the Olduvai Subzone 

(McDougall, 1985). (2) The age of the Cobb Mountain 

Normal-polarity Subchron of the Matuyama Chron is 1.10 

million years (Mankinen and others, 1978). (3) The beginning of the Brunhes Normal-polarity Chronis close to 0.79 

million years before present (lzett and Obradovich, 1991). 

FLORAL TERMS 

Use of the terms "mixed-conifer forest," "boreal 

forest"( taiga), and "tundra" is conventional. Pollen percentages cited herein that include some tundra plants as part of 

the flora could fall into the classification "forest-tundra" of 

some authors; however, such detail has little significance in 

the present discussion. The term "tundra" is intended to 

indicate extensive tundra well removed from any arboreal 
0 Q)Ol NORTH ATLANTIC NORWEGIAN SEA . --- ..... -

E~ () 180 Ice-rafted debris (percent >63 f.lm) Events c: NORTH AMERICA EASTERN ASIA ARCTIC OCEAN ARCTIC 0 LOW-LATITUDE 

!=>:'-

cr: 

Site 607 Site 644 and time e SECTIONS SECTIONS EVENTS BORDERLAND ::J PROXY DATA 

E 

w 

- (m.y. ago) .r::. EVENTS 5 EVENTS 

u z 

EXPLANATION I I I (1;;,;cc1 Marine deposits 

~ 

- c::J Marginal-marine deposits 

.:Jf£2 c::J Nonmarine deposits Perpetual 

8{~ w Age-significant evidence ice cover Cl) 

·~ Approximate position of fauna 

Cl) Microtus to Cl) 

........ 

Q) Q) Q) 

Polarity ~ ~ eastern U.S. ~ 

~ 

Q) Q) 

\ Normal (.J (.J ] 

M,;m,m gl";''' I z 

Reversed 1 

<1: 

.;;;:: ii: 

LJ.J 

===e- ~ 

~ 0.79 

0 

-=¥ cr: 

u leo to low' 1 

0.90 begin Lasiopodomys and 

~ ~ OLDCROW s Terricola to U.S. 

~- I~ CRH-47 Treeline I at 70° N. 

0.97 Forests 

r---=- Central ocean 

ice covered Tundra 

Treeline I at ±66° N. D CAPE DECEIT 

lila KRESTOVKA UJ l OLYOR SUITE ~ 

Treeline lat 

No ice cover 66°-70°N. South Dakota 

Tundra and forests warm and wet 

KL0-8 (alternating) z 

~ I~ l n•nii\IAI - ~~ ~~A Phenacomys to U.S. 

1.90 KAP K0BENHAVN 

_$- 2.01 

I Treeline I at 84o N.l ~ I Allophaiomys to U.S. 

REUNION{ 

e20my,go 

Forests § 

82~ 2.1 m.y. ago 

Maximum subfrigid Lepus and Hypolagus 

temperatures 

oo__J 2.14 Maximum glaciation I"? I Ice to Iowa 

FISH CREEK begins 

SECTION KUTUYAKH BEDS Central ocean Tundra 

100~101 

ice covered Treeline I at ±66° N. 5 2.40 m.y. ''' Frigid temperatures First permafrost 

! 2.48 2.48 m.y. ago 

r=-

Density stratification 

Prominent ice rafting I Treeline lat 70o N. I Bog lemmings to U.S. 

~ D BEGUNOV SUITE 

begins 

COLVILLIAN 

llllr""l,.,llr-a.l I 

Cold-temperate to I Forests 

subfrigid temperatures Treeline I at 79_5o N. 

Bering Strait open 

Figure 2. Correlation of localities and events discussed in text. Data for site 607 from Ruddiman and others (1989) and Raymo and others (1989); data for site 644 from Jansen and 

others (1988). Data for sites 607 and 644 have not been corrected for differences in polarity age assignments. Sources for other data are given in the text. Only those oxygen isotope 

I 

~ ::t' 

0 

0 

c: 

B 

~ 

stages that are mentioned in the text are numbered. -...J 
8 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

plants. In addition, many pollen samples cited have a high 

percentage of grass, such as in the younger part of the 

Kutuyakh Beds (Pliocene) of Sher and others (1979), and 

these floras might better be referred to as steppe or foreststeppe mosaic than as forest-tundra mosaic, but the distinction is not emphasized here. 

Acknowledgments.-This summary of the late 

Pliocene and early Pleistocene history of the Arctic Ocean 

Borderland is a consolidation of information from many 

sources that has accumulated over the past 10 years. During this time we have benefited from the work of many 

people, and we would like to particularly thank Thomas 

Ager (U.S. Geological Survey, Denver), Ole Bennike 

(Geological Museum, University of Copenhagen), Julie 

Brigham-Grette (University of Massachusetts, Amherst), L. 

David Carter (U.S. Geological Survey, Anchorage), Oldrich Fejfar (Charles University, Czechoslovakia), Rolf Feyling-Hanssen (Aarhus University, Denmark), Svend Funder 

(Geological Museum, University of Copenhagen), Niels 

Jorgensen (Institute of Historical Geology and Paleontology, University of Copenhagen), Louie Marincovich, Jr. 

(U.S. Geological Survey, Menlo Park), John Matthews, Jr. 

(Geological Survey of Canada, Ottawa), David Penney 

(Aarhus University, Denmark), Andrei Sher (Academy of 

Sciences of the U.S.S.R., Moscow), Charles Schweger 

(University of Alberta, Edmonton), Jack Wolfe (U.S. Geological Survey, Denver), and Vladimir Zazhigin (Academy 

of Sciences of the U.S.S.R., Moscow). Thomas Ager, Tony 

Bryant, Oldrich Fejfar, Art Grantz, Patricia Holroyd, John 

Matthews, Hugh Rieck, Richard Tedford, and Jack Wolfe 

have reviewed the manuscript and offered suggestions for 

its improvement. 

Repenning was supported by the G.K. Gilbert Fellowship of the U.S. Geological Survey from April 1991 to 

April 1992, and without this support this synthesis, which 

was begun in 1986, might not yet be completed. 

3.0-2.0 MILLION YEARS AGO: 

THE FIRST COOLING CYCLE 

As herein considered, the maximum age for the opening of Bering Strait is about 3.0 million years. This date is 

based on a dated lava flow that is interbedded with sediments containing the first occurrence of Pacific mollusks in 

Iceland (Gladenkov, 1981). Recently, evidence has been 

described (Gladenkov and others, 1991) that suggests an 

opening of the Bering Strait about 4 million years ago, but 

it contains some discrepancies between the biostratigraphic 

and radiometric dates. In addition, we are aware of no evidence for the dispersal of North Pacific invertebrates into 

the Arctic and North Atlantic Oceans during this time 

period. Known relationships between marine and terrestrial 

climatic trends and associated mammal faunas all favor the 

interpretation of an opening of the Bering Strait about 3.0 

million years ago, as has been discussed. 

CANADA, MEIGHEN ISLAND: 

BEAUFORT FORMATION 

The commingling of Pacific-derived mollusks and 

Atlantic-derived mollusks that were displaced from the 

Arctic Ocean in later faunas has been described from the 

Beaufort Formation (Pliocene) of Meighen Island, Canadian Arctic Archipelago (lat 79°30' N.) (Fyles and others, 

1991) (fig. 1). The molluscan fauna is believed to represent 

a rather brief time immediately after introduction of Pacific 

forms into the Arctic Ocean and prior to extirpation of 

some endemic Arctic-Atlantic species. The deposits are 

reversely polarized and bounded above and below by normally polarized deposits. 

The interpretation preferred by Fyles and others 

(1991) is that the reversed polarity associated with the molluscan fauna represents one of the two reversed-polarity 

events of the Gauss Normal-polarity Chron and indicates 

an age between 3.1 and 2.9 million years. This interpretation is in agreement with other evidence discussed herein 

and so is also preferred by us. The other possible interpretation is that some time within the Gilbert Chron is represented. 

Available faunal data indicate a cold-temperate 

marine climate with temperatures comparable to the modem North Atlantic south of Newfoundland, water temperatures not present in the modern Arctic Ocean. Plant and 

insect fossils in the overlying, normally polarized organic 

deposits indicate a treeline flora that differs from modem 

treeline floras (now about 1,230 mi ( 1,980 km) to the south 

at lat 64° N.) only in having a richer variety of taxa, with 

as many as six types of conifers, some now extinct (Fyles 

and others, 1991). 

Four species of shallow-marine ostracodes ( Cytheretta 

teshekpukensis, Acanthocythereis dunelmensis, Sarsicytheridea bradii, and Thaerocythere aff. T. crenulata) are 

present. These forms of Atlantic origin permit little chronologie interpretation. Their earliest known Arctic Ocean 

record is in the pre-Bering Strait deposits of the Nuwok 

Member (Oligocene to Pliocene, but see Brouwers, in 

press) of the Sagavanirktok Formation (Cretaceous to 

Pliocene) from the eastern Alaskan North Slope. Ecologically, they indicate a cold-temperate to warm-subfrigid 

marine climate, compatible with the environment indicated 

by the mollusks. 

No ostracodes of Pacific origin are present. Present 

records from the Arctic Ocean suggest that ostracodes disperse more slowly than mollusks, presumably because of 

their benthic habits throughout their life cycle, as contrasted with the readily mobile and dispersed pelagic habit 

of molluscan veligers. 
3.0--2.0 MILLION YEARS AGO: THE FIRST COOLING CYCLE 9 

ALASKAN NORTH SLOPE: 

GUBIK FORMATION 

In the Alaskan North Slope area (fig. 1), at about lat 

70° N. and 9:5° farther south than the locality on Meighen 

Island, deposits younger than the opening of the Bering 

Strait are included in the Gubik Formation (Pliocene and 

Pleistocene). This formation was deposited across a stable 

coastal plain, the Alaskan North Slope, that is bounded by 

the Brooks Range to the south and the Arctic Ocean to the 

north. A number of eustatic sea-level rises have resulted in 

marine transgressions across this plain that deposited a 

series of interbedded shallow-marine and nonmarine 

deposits comprising the Gubik Formation. The transgressions have been identified by geomorphic characteristics, 

faunal suites, and stratigraphy; they have been informally 

named by Carter and others (1979). Originally thought to 

be of middle and late Pleistocene age, those now known to 

be of late Pliocene age (Repenning, 1983; Repenning and 

others, 1987; Carter and Hillhouse, 1992) have been called 

the Colvillian, Bigbendian, and Fishcreekian transgressions, listed in depositional order. 

COLVILLIAN AND BIGBENDIAN TRANSGRESSIONS 

Diverse ostracode faunas, comprising 54 species, are 

present in deposits of the Colvillian and Bigbendian transgressions, the first deposits of the Gubik Formation of the 

Alaskan North Slope (Brouwers, in press). The oldest recognized transgression of the Gubik Formation, the Colvillian transgression, is marked by a mixture of Pacific and 

Atlantic species and therefore is younger than the Bering 

Strait. 

Geographically widely separated localities in the eastem Alaskan North Slope (eastern Beringia) are uniform in 

species composition and contain comparable ostracode 

species associations that include Acanthocythereis dunelmensis, Elofsonella concinna, Heterocyprideis sorbyana, 

Palmenella limicola, Pterygocythereis vannieuwenhuisei, 

Rabilimis paramirabilis, and Robertsonites tuberculatus 

(Brouwers, in press). These assemblages are dominated by 

Atlantic-derived genera (Sarsicytheridea, Rabilimis, Pterygocythereis, and Heterocyprideis) but contain some 

Pacific-derived genera (Finmarchinella, Palmenella, Hemicythere, Howeina, Pectocythere, and Robertsonites). 

These oldest post-Bering Strait assemblages of the 

Alaskan North Slope indicate a stable paleoenvironment 

with little to no fluctuation of the physical-chemical environment. These conditions are typical of deeper inner shelf 

to middle shelf water depths. Based on the ostracodes, the 

shallow Arctic Ocean was much warmer than today, typical of North American northern cold-temperate to southern 

subfrigid marine climatic zones (northern Nova Scotia and 

southern Labrador) and the Transitional and Norwegian 

biogeographic provinces of Europe (Hazel, 1970). 

The ostracode faunas of these two earliest transgressions (Colvillian and Bigbendian) are indistinguishable 

(Brouwers, in press). The mollusk assemblage from the 

lower, normally polarized part of the Bigbendian deposits 

indicates that the deposits are younger than the about 3.0-

million-year opening of Bering Strait and that the Arctic 

Ocean was not frozen close to the then-existing shore during the winters (Repenning and others, 1987; L. Marincovich, Jr., unpublished data). The presence of an archaic sea 

otter also indicates that the ocean was not frozen to its 

shore at any time during the year and that the deposits are 

less than 3 million years old (Repenning, 1983). 

Deposits of both transgressions were first recognized 

in the vicinity of Ocean Point, along the Colville River, 

where deposits of both the Colvillian transgression and the 

older part of the Bigbendian transgression are normally 

polarized. The upper part of the Bigbendian deposits have a 

reversed magnetic polarity that is believed to represent the 

beginning of the Matuyama Reversed-polarity Chron 

(Carter and Hillhouse, 1992). The evolutionary stage of the 

archaic sea otter found at Ocean Point in a Bigbendian 

mollusk bed having normal polarity and possibly 16 ft (5 

m) below the polarity reversal (stratigraphic position indicated by Carter and Hillhouse, 1992; specimen collected by 

Carter in 1975) indicates an age older than the Olduvai 

Subchron of the Matuyama Chron and younger than 3 million years (Repenning, 1983). 

Pollen from the Colvillian and Bigbendian deposits at 

the Ocean Point section indicate a terrestrial climate considerably warmer than today, possibly warmer than modem 

Anchorage, Alaska, which is north of the modem range of 

some of the trees recorded (Nelson and Carter, 1985, 

1992). Based on the flora, permafrost could not have been 

present. Nelson and Carter (1992) suggested that the flora 

of the Bigbendian transgression represents a more "open 

spruce-birch woodland or even parkland" than the flora of 

the Colvillian transgression. They noted the absence of 

hemlock and a reduction of pine in the Bigbendian flora 

and suggested that the climate was slightly more severe 

than that during the earlier Colvillian transgression. 

There is no record of extensive tundra or permafrost 

in these earliest post-Bering Strait deposits at about lat 70° 

N.; today, the vegetation of the Alaskan North Slope consists of full tundra and extensive permafrost. 

Although the mixed conifer and deciduous forests 

present in Alaska during the Miocene (Wolfe and Tanai, 

1980) had mostly withdrawn southward from the Arctic 

Ocean Borderland, spruce-birch forests with some pine, 

hemlock, larch, and poplar were still present less than 3 million years ago in the latitude of Beringia (Nelson and Carter, 

1985, 1992). To judge from the pollen records, particularly 

those of the Bigbendian transgression, these remnants of the 

early Pliocene forests included open areas that supported 

tundra plants and, at times, significant percentages of grass; 

a forest-tundra or forest-steppe mosaic is indicated. 
10 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

Amino acid racemization (alle/Ile) suggests a significant difference in age between deposits of the Colvillian 

and Bigbendian transgressions (Brigham-Grette and Carter, 

1992). The similar marine ostracode faunas indicate, however, stability and great geographic extent (as far north as 

Meighen Island) of marine temperatures during the time 

represented by these transgressions. Also indicated by 

comparison between Meighen Island and the deposits at 

Ocean Point is the fact that late Pliocene terrestrial climates, in contrast with shallow-marine climates, differed 

greatly with latitude, as they do today. 

Although the first two marine transgressions of the 

Alaskan North Slope represent a marine climate similar to 

that farther north at Meighen Island, they are believed to be 

younger because they lack some Atlantic mollusks, as indicated by the Bigbendian fauna, that are present in the 

marine beds of Meighen Island. Subsequent to the Meighen 

Island fauna, these Atlantic mollusks were no longer 

present in the Arctic Ocean. The Colvillian and Bigbendian 

transgressions may span the time from shortly after 3 million years ago until shortly after the end of the Gauss 

Chron, 2.48 million years ago. As has been noted, the 

Meighen Island flora represents a treeline environment; the 

differences between the Colvillian and Bigbendian floras 

and those from Meighen Island can be explained by the 

difference in latitude. 

At present there is no way of estimating the amount of 

time between the Colvillian and Bigbendian transgressions, 

but the similarity of the ostracode faunas and the terrestrial 

floras during the two events suggests little change in 

marine and terrestrial temperatures. Differences in the terrestrial flora, including the disappearance of hemlock and 

reduction in percentages of pine, suggest some lowering of 

air temperature and possibly a drier climate. 

Oxygen isotope ratios from the World Ocean indicate 

many fluctuations in sea level during the time ascribed to 

the Colvillian and Bigbendian transgressions (fig. 2), 

whereas the deposits of these transgressions indicate fewer. 

Although this relationship may indicate a lack of adequate 

depositional records on the Alaskan North Slope or inadequate detail of their examination, the Colvillian and Bigbendian records and others in the Arctic Ocean Borderland 

suggest that shallow-marine deposits and, particularly, 

terrestrial deposits dampen the sensitivity of the record 

found in marine isotope ratios; minor eustatic sea-level 

changes may result in little lithologic change if the locality 

is not crossed by a shifting strand line, and some of the isotope fluctuations may reflect Antarctic events. 

FISHCREEKIAN "TRANSGRESSION" 

A third Pliocene marine sequence was informally 

named the Fishcreekian transgression (Carter and others, 

1986). The type section is exposed along Fish Creek, 16 mi 

(26 km) northwest of Ocean Point. Its faunal and floral 

character were discussed in greatest detail by Repenning 

and others ( 1987). The Fishcreekian type section is composed of reversely magnetized marginal-marine deposits 

that contain a fauna of coldwater ostracodes much different 

than those of the Colvillian and Bigbendian transgressions 

and essentially identical to the living ostracode fauna of the 

inner sublittoral Arctic Ocean. 

The mollusks in the Fishcreekian type section, however, are distinctly warmer water forms than the modern 

Arctic Ocean fauna: they represent a cold-temperate to subfrigid marine climate similar to that of the southern Bering 

Sea and the northern Gulf of Alaska (Marincovich in 

Repenning and others, 1987). The presence of an archaic 

sea otter indicates a similar marine climate. The indications 

of the marine climate overlap in the subfrigid climate 

range. 

The presence of some marine mollusks and an archaic 

sea otter in the Fish Creek fauna indicates that, as during 

the Bigbendian transgression, the ocean could not have 

been frozen to the shore during the winter. Small crystal 

clusters (about 5 mm in diameter) of glendonite (a calcite 

pseudomorph after ikaite) in the mollusk- and vertebratebearing deposits in the upper part of the section indicate 

that the water was not much above freezing temperatures 

(maximum 5°C) (Kennedy and Pickthom, 1987). 

Pollen floras in the Fish Creek section show increasing climatic severity upsection, ranging from shrub tundra, 

possibly with some larch, in the lower part of the section to 

herb tundra near the top. Pollen of pine, spruce, and larch 

are sufficiently abundant in samples from the lower part of 

section to suggest that the limit of the boreal forest was 

farther north than today. Decreases in the percentages of 

these elements upsection suggest a southward withdrawal 

of distant forests or a change in prevailing wind direction. 

As the percentage of arboreal pollen decreases upsection, the pollen spectrum changes from shrub tundra to wet 

(herb) tundra, suggesting the possible development of permafrost inland from the marginal-marine deposits of the 

Fish Creek section. This change in floral aspect is accompanied by the appearance of the pollen of yellow water lily, 

indicating that the area was not as cold as it is today (Ager 

in Repenning and others, 1987). 

The lowest part of the basal silty unit in the Fish 

Creek section (about the lowest 4 ft (1.2 m)) contains a 

sparse ostracode fauna, including Pterygocythereis vannieuwenhusei. This fauna indicates a markedly warmer 

ocean than does the overlying Fish Creek fauna. The 

warmwater ostracode fauna (localities 83EB 171 and 

83EB172 in Repenning and others, 1987, fig. 3) is comparable to the faunas of the Colvillian and Bigbendian transgressions. The presence of P. vannieuwenhusei, an extinct 

species normally found in association with the extinct 

Rabilimis paramirabilis, indicates a post-Bering Strait age; 

the warmth it indicates implies a pre-Fishcreekian age. 
3.0-2.0 MILLION YEARS AGO: THE FIRST COOLING CYCLE 11 

Coldwater ostracode faunas, typical of the Fishcreekian 

deposits, are present in the upper part of this basal silty unit 

(locality 78ACr60a in Repenning and others, 1987). The 

upper part of the lower silt unit also contains the oldest 

pollen flora, which indicates shrub tundra (Repenning and 

others, 1987, fig. 3). Amino acid ratios of marine mollusks 

presumed to be from the oldest part of the Fish Creek section also indicate that these basal silt deposits are older 

than the overlying part of the section; the amino acid ratios 

are similar to those from deposits of the Bigbendian transgression at Ocean Point (Brigham-Grette and Carter, 1992). 

Although Brigham-Grette and Carter (1992) assigned 

the entire lower silt unit of the Fish Creek section ( -13 ft 

(4 m) thick) to the Bigbendian transgression, the warmwater ostracode fauna was found only in the lowermost 

meter of this silt unit, and coldwater ostracodes of the Fish 

Creek fauna, as well as the oldest tundra pollen floras, 

were found in the upper part of this lowest unit. The climatic change from open forests and warm water of the 

Bigbendian transgression to shrub tundra and cold water of 

the Fishcreekian transgression occurred entirely during 

deposition of the lower silt unit of the Fish Creek section. 

The entire Fish Creek section is reversely polarized 

(Carter and Hillhouse, 1992). A single basal sample 

reported to be normally polarized (Repenning and others, 

1987) was recollected and the initial polarity determination 

found to be in error (Carter and Hillhouse, 1992). Thus, 

warmwater ostracodes are present later in the sedimentary 

record than the end of the Gauss Chron and are found in 

deposits having a reversed polarity that represents the earliest Matuyama Chron. Probably the climatic change from 

warm to cold water and from forest to tundra was (1) rather 

abrupt, (2) during the oldest part of the Matuyama Chron, 

(3) entirely within the lower silt unit of the Fish Creek 

section, and ( 4) not marked by any depositional change, 

occurring instead at the peak of the transgression because 

the remainder of the Fish Creek section consists of regressive deposits. 

The type section of the Fishcreekian "transgression" 

grades upward from flatbedded, silty marine deposits to 

strongly crossbedded sandy units considered to be either 

tidal channel deposits or nearshore wave-washed sands 

(Repenning and others, 1987). These sandy units contain 

shallow-water benthic foraminifers and mollusks, as well 

as intermixed terrestrial and marine mammals. Upsection, 

the ostracode fauna contains increasing numbers of nonmarine species. The marginal-marine deposits are overlain 

by nonmarine sands. 

The type section of the Fishcreekian is a regressive 

sequence and apparently records the end of the Bigbendian 

transgression as well as the first cooling period following 

the warm, transgressive phase of the Bigbendian. It also 

apparently records the first cooling period during the 

Matuyama Chron. How this regressive sequence relates to 

evidence of a Fishcreekian transgression elsewhere 

depends on field evidence that we do not have. 

The polarity change from normal to reversed in the 

Bigbendian deposits represents the Gauss-Matuyama 

boundary, an interpretation of Carter and Hillhouse (1992) 

that is not disputed. The Ocean Point and Fish Creek localities record a single history of regression from the peak of 

the Bigbendian transgression. The transgression began with 

a warmwater fauna and open forests on adjacent lands. The 

regressive deposits record a northward (seaward) retreat of 

the Arctic Ocean from Ocean Point to a point north of Fish 

Creek and a coldwater fauna with tundra on adjacent lands. 

The Bigbendian transgression culminated during the earliest Matuyama Reversed-polarity Chron, and the Arctic 

Ocean withdrew north of Fish Creek as the oceanic water 

cooled abruptly and terrestrial climatic deteriorated. 

During the withdrawal of the Arctic Ocean, its mollusk fauna remained about the same. Primitive sea otters 

are known from both the "warm" (Bigbendian) and "cold" 

(Fishcreekian) parts of the transgression, but the ostracode 

fauna changed to a modem coldwater fauna before the end 

of deposition the oldest unit of the Fish Creek section. The 

terrestrial vegetation changed dramatically, from the forest 

described by Nelson and Carter (1992) for the Colvillian, 

to an open forest lacking pine and hemlock during the Bigbendian transgression, to shrub tundra and evidence of distant forests during the early Fishcreekian regression, to a 

wet tundra and possible permafrost and less evidence of 

distant forests during the late Fishcreekian regression. The 

youngest deposits of the Fish Creek section are nonmarine. 

Subsequent to the publication of Repenning and 

others ( 1987), much has been learned about the ages of the 

depositional units shown on their figure 13 (Repenning and 

Grady, 1988; Repenning, in press). However, aside from 

raising the base of the Fish Creek section so that normal 

polarity (then thought to have been present) is not shown at 

its base, the age assignment of 2.4 million years for the 

mammal fauna in the upper part of the section is still most 

likely. The oxygen isotope record for North Atlantic site 

552 (Shackleton and others, 1984; modified by Raymo and 

others, 1989) indicates that the first and most prominent 

eustatic sea-level regression following the end of the Gauss 

Chron began about 2.44 million years ago and peaked 

about 2.40 million years ago (site 607, fig. 2). 

If the regression recorded by the Fish Creek section 

represented any later, and less extreme, sea-level regression, one would expect to see some fluctuation in 

depositional environments and oceanic temperatures as 

reflected in the ostracode faunas of the Fish Creek section, 

and none was found. If other transgressive cycles had taken 

place between the Bigbendian transgression and the Fishcreekian regression, more time would be involved and one 

would expect the warmwater ostracodes at the base of the 

Fish Creek section to differ from the Bigbendian ostracode 

fauna; this was the time of the most marked change of the 
12 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

ostracode fauna of the Arctic Ocean. One would also 

expect some dissimilarity in the amino acid ratios of the 

Bigbendian and basal Fish Creek deposits if the basal beds 

of the Fish Creek Section did not represent the transgressive phase of the Bigbendian. The terrestrial climate trends, 

as indicated by the pollen from Colvillian and Bigbendian 

deposits, show the same uninterrupted trend in environmental cooling, and no suggestion of fluctuation. That a 

climatic fluctuation would be detectable in the pollen profile is demonstrated by the record of the youngest part of 

the Kutuyakh Beds in Y akutia, to be discussed; this pollen 

record was deposited during the same time period but 

lasted a bit longer, in that it records the return of forests 

following the conclusion of the cold period about 2.4 million years ago. 

In brief, there is only one uninterrupted record in the 

Fish Creek Section showing regression, marine cooling, 

and faunal and floral change between the Bigbendian transgression and the Fishcreekian regression; no evidence suggests that the Fishcreekian regression represents anything 

other than a single (the first) sea-level drop following the 

Bigbendian transgression that began before the end of the 

Gauss Chron. 

Carter and Hillhouse (1992) suggested that considerable time elapsed between the reversely polarized part of 

the later Bigbendian deposits at Ocean Point and the 

reversely polarized Fishcreekian regressive deposits of the 

Fish Creek section, apparently because of a misunderstanding of the age significance of the fossil evidence. 

The primitive bog lemming Plioctomys mimomiformis 

is present in the Fish Creek fauna near the top of the section (Repenning and others, 1987; Repenning, in press). A 

more advanced species of this genus (worldwide, the last 

record of the genus) became extinct in the United States 

before the two normal-polarity excursions of the Reunion 

Subchron of the Matuyama Chron (that is, more than 2.14 

m.y. ago) (Repenning and Grady, 1988). The Fish Creek 

record of P. mimomiformis in east-central Beringia is the 

only record of this species outside of Eurasia; the youngest 

Eurasian record is in the Kutuyakh Beds of the Krestovka 

sections, Y akutia, in westernmost Beringia (in the same 

faunal province as Fish Creek), which were deposited during the first cooling period after the polarity change at the 

end of the Gauss Chron. 

In the Krestovka sections of west Beringia, the climatic trend of terrestrial floras is essentially identical in 

time and nature to that of the Colvillian-BigbendianFishcreekian sequence (Repenning, in press). Although 

there are no marine deposits in the Krestovka sections, the 

sequence records the change from open forests (Begunov 

Suite), to forest-tundra (older Kutuyakh Beds), and to the 

first permafrost and treeless tundra (younger Kutuyakh 

Beds, above the Gauss-Matuyama polarity boundary). 

In western Beringia, the entire record is present at one 

locality, and there is less uncertainty than in reconstructing 

its history from localities 16 mi (26 km) apart on the Alaskan North Slope. The principal difference between the 

Alaskan and Yakutian climatic records stems from the 

much greater (but unknown) distance of the Krestovka sections from the late Pliocene Arctic Ocean and the continentality of the Asian land mass; the compared floras of 

Y akutia are consistently drier in nature and have a foreststeppe or steppe composition; grass is much more abundant 

in the pollen record of Yakutia than in the Colvillian-Bigbendian-Fishcreekian climatic succession of Alaska. 

As mentioned, the Krestovka record extends a bit later 

in time than does the Fish Creek section, and it shows the 

first of the following return of forests (youngest Kutuyakh 

Beds) about 2.38 million years ago based on the calibration 

of the oxygen isotope record of sea-level change. The top 

of the Kutuyakh Beds is marked by an erosional unconformity that is overlain by deposits about 1.5 million years old 

(Repenning, in press). 

The following points summarize the age-significant 

data pertaining to the Fish Creek section. The 

transgressional-regressional history of the BigbendianFishcreekian sequence, exposed in two localities 16 mi (26 

km) apart, represents one sea-level cycle of the Arctic 

Ocean tied together by an unbroken continuum in (1) cooling trends of both marine and terrestrial climates, (2) depo-. 

sitional types, (3) amino acid ratios, (4) marine faunas 

other than ostracodes, and (5) paleomagnetic patterns and 

by the lack of irregularities in these patterns and trends that 

would suggest temporal gaps in the record which would 

eliminate evidence of more than one eustatic sea-level 

cycle. 

This interpretation is supported by an essentially 

identical record at one locality in Y akutia that is correlated with the Alaskan record by (1) paleomagnetic pattern, (2) unique mammalian faunal elements, (3) very 

similar flora history, and ( 4) very similar climatic 

changes. This Y akutian record lasts briefly into the next 

climatic (eustatic) cycle, recording the return of warmer 

climate and of boreal forests; it thus indicates that, if 

more than one cycle had been involved in the Fish Creek 

section, the climatic-floral change would have been recognizable, as presumably would the change from regressive 

to transgressive deposits, as has been mentioned. 

Thus, all available evidence suggests that the Fish 

Creek fauna in the upper part of the Fish Creek section of 

the Alaskan North Slope is within the regressional sea level 

indicated between isotope stages 100 and 101 (fig. 2) and 

is about 2.40 million years old. We see no reason to change 

our age interpretation of the Fish Creek section as being 

about 2.4 million years old (Repenning and others, 1987) 

and see unnecessary complication and additional uncertainty in a younger age interpretation, as suggested by 

Carter and Hillhouse (1992) and Brigham-Grette and 

Carter (1992). 
3.0-2.0 MILLION YEARS AGO: THE FIRST COOLING CYCLE I3 

Table 1. Pollen from the Bigbendian deposits, Alaskan North 

Slope (Nelson and Carter, 1985), compared with pollen from 

the Begunov Suite, Yakutia (Giterman and others, 1982). 

Pinus 

Picea 

Abies 

Larix 

Pollen Reported from Both Bigbendian Deposits 

and Begunov Suite 

Alnus and Alnaster 

Betula 

Poaceae 

Cyperaceae 

Artemisia 

Caryophyllaceae 

Chenopodiineae 

Ranunculaceae 

Sphagnum 

Lycopodium 

Salix 

Ericales 

Not Reported from Begunov Suite 

Sparganium Saxifragaceae 

Juncaceae Rosaceae 

Liliaceae Leguminosae 

Polygonaceae Scrophulariaceae 

Papaveraceae Valeriano 

Cruciferae Other trilete spores 

Onagraceae Monolete spores 

Primulaceae Selaginella 

Polemoniaceae 

Not Reported from Bigbendian Deposits 

Pinus pumila 

Betula sect. nannae 

Identified in Greater Taxonomic Detail 

from Bigbendian Deposits 

Bryales 

Filicales 

Although the faunal evidence at Fish Creek suggests 

that winter ice did not extend to the shore, Scott and others 

(1989) reported evidence (CESAR core I4, from the central Arctic Ocean) suggesting that perennial ice cover 

began about 2.48 million years ago (at the end of the Gauss 

Chron) and was intermittently present until 2.I million 

years ago (at the beginning of the Reunion Subchron). This 

is the earliest record of perennial ice in the Arctic Ocean, 

and the ice was present only intermittently. 

The next intermittent perennial ice cover of the central 

Arctic Ocean reported by Scott and others (1989) was 

about I million years later, about 1.2 million years ago 

(with a reinterpreted identification of one magnetic event, 

to be discussed). 

At 2.I million years ago glaciers flowed down the 

Mississippi River Valley as far south as Iowa (Easterbrook 

and Boellstorff, I984). This was the time of greatest cold 

in the Arctic Ocean Borderland, rather than 2.4 million 

years ago as suggested by oceanic isotope records; the isotope records must reflect Antarctic conditions in addition to 

Arctic conditions. 

The isotopic records of the World Ocean suggest that 

sea levels were intermittently lower beginning with the 

Fish Creek depositional record and continuing until 2.I 

million years ago. A lack of terrestrial records in the Arctic 

Ocean Borderland during this time suggests that lowered 

sea levels during this first glacial period made continental 

deposition in the Arctic Ocean Borderland rare. This was 

certainly true in other areas, such as the Atlantic Coastal 

Plain of the United States, where Quaternary deposits dating from glacial periods are rare (T.A. Ager, written commun., I992). 

YAKUTIA: BEGUNOV SUITE AND KUTUYAKH BEDS 

The Krestovka sections (fig. I) are in easternmost 

Y akutia, along a bluff of the Krestovka River near its 

junction with the Kolyma River (Sher and others, I979). 

The sections are in a single exposure at about lat 69° N ., 

near the average latitude of the Alaskan North Slope. The 

Krestovka sections are more distant from the present and 

former sea coast than the Ocean Point sections in Alaska, 

and no marine deposits are known from them. Unlike the 

northern coast of North America, that of eastern Asia is, at 

present, mostly covered by an extensive marine 

transgression of the Arctic Ocean that makes the Siberian 

Shelf the broadest continental shelf in the world; almost all 

of it is less than IOO m below sea level. 

Two formational units in the Krestovka sections represent the first cooling cycle of the Arctic Ocean region: 

the older Begunov Suite· and the younger Kutuyakh Beds 

(Sher and others, I979). No mammals have been found in 

the Begunov Suite; however, the upper part of the Kutuyakh Beds contains several microtine faunas. Fossils of 

large mammals are rare. 

Although not completely documented by paleomagnetic sampling, the Begunov Suite has been checked by 

two samples, both of which were normally polarized after 

cleaning. Sedimentary structures in the Begunov Suite are 

interpreted as indicating seasonal frost cracks but not permafrost (Sher and others, I979). The Begunov Suite contains at least 20 percent arboreal pollen (Sher and others, 

I979) and cones of spruce and of three extinct species of 

larch. An open spruce-birch forest with some pine and with 

a greater diversity than today is indicated (Giterman and 

others, I982). 

Except for its lower percentage of arboreal pollen, the 

floral list of the Begunov Suite (Sher and others, I979) is 

very similar to that reported from deposits of Colvillian 

and Bigbendian transgression of the Alaskan North Slope 

(Nelson and Carter, I985, I992); the differences are related 

to the detail of identifications (table 1) and to fewer tundra 

plants in the Begunov Suite flora. The lesser amounts of 

tundra plants reflect the greater amounts of grasses. Several 

aspects of the Krestovka pollen spectra suggest that the climate was drier, reflecting greater continentality of Asia. 

The Begunov pollen record suggests less tundra and 

greater steppe flora, intermixed with boreal trees, than the 

Colvillian and Bigbendian records. One or two grains of 
14 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

fir, hemlock, hazel, or linden in 70 percent of the samples 

(Sher and others, 1979) suggest the existence of more distant, cool-temperate vegetation and enhance the similarity 

of the Begunov and Colvillian floras. The Colvillian flora 

also contains rare ftr, hemlock, and hazel (Nelson and 

Carter, 1992). 

The Begunov pollen content differs little from the 

pollen suite in the lower part of the overlying Kutuyakh 

Beds except for a decrease in the percentage of arboreal 

pollen, a trend that continues upsection within the Kutuyakh; this difference is comparable to the floral difference 

between the Colvillian and Bigbendian deposits. 

A polarity change from normal to reversed, correlated 

with the Gauss-Matuyama polarity boundary on the basis 

of the younger Kutuyakh microtine fossils, is present in the 

lower part of the Kutuyakh Beds (Sher and others, 1979, p. 

63: paleomagnetic studies by E.l. Virina; microtine rodent 

studies by V.S. Zazhigin). It is difficult to follow the reasoning of the authors (from their description and illustrations), but, if we have correctly interpreted their 

statements, the polarity boundary is about 1.5 m (5 ft) 

above the base of the Kutuyakh Beds. Sher and others 

(1979, p. 63) mentioned that permafrost action in the upper 

part of the Kutuyakh Beds has distorted paleomagnetic 

directions so that they are anomalous, and the polarity 

boundary may fall within this zone of anomaly. 

Thus, the age-discriminating fauna and the earliest 

evidence of permafrost, as well as the transition from open 

forest to tundra (or steppe) in the Kutuyakh Beds, are 

younger than the polarity reversal and contemporaneous 

with the early part of the Matuyama Chron, as is indicated 

in the Bigbendian and Fishcreekian deposits. 

The arboreal pollen content of the younger Kutuyakh 

Beds decreases upward to essentially none in that part of 

the section containing most of the fossil microtines. There 

the "pollen" spectrum contains abundant pond weed (several extinct species) and seeds of yellow water lily (Sher 

and others, 1979), comparable to the Fish Creek flora of 

Alaska. The youngest samples from the Kutuyakh Beds, 

however, show an increased percentage of arboreal pollen 

of between 7 and 12 percent (Sher and others, 1979). The 

return of arboreal pollen is unknown in the Fish Creek section (Repenning and others, 1987), probably because no 

Fish Creek deposits are as young as the youngest Kutuyakh 

Beds. 

Paleomagnetic determinations for the Colvillian and 

Bigbendian transgressions and the Fishcreekian regression 

suggest that these deposits and the Begunov-Kutuyakh 

sequence of the Krestovka sections are of essentially identical age. Both the Alaskan and Y akutian sections were 

deposited during the Gauss and earliest Matuyama Chrons, 

have similar pollen records and trends in floral composition, and are at almost the same latitude. Both sections contain a distinctive bog lemming (Plioctomys mimomiformis) 

in the same part of the paleomagnetic and paleoclimatic 

record, and both are in the Beringian faunal province. The 

Kutuyakh Beds record the return of forests and are believed 

to include slightly younger deposits than the Fish Creek 

section in Alaska, about 2.38 million years old (fig. 2). 

The lesser percentage of arboreal pollen in the 

Begunov-Kutuyakh sequence indicates a somewhat more 

severe or more arid climate during the last of the Gauss 

Chron than that recorded for Bigbendian deposits, although 

both sections show a continuous decrease in arboreal flora 

upsection. Some hemlock is present in the Colvillian flora 

but apparently absent in the Bigbendian flora; it is very 

rare in the Begunov records and absent in the Kutuyakh 

flora. Considering the continentality of the Asian climate 

and the greater distance of the Krestovka sections from the 

ocean (Fish Creek section, 23 mi (37 km); Krestovka sections, 200 km (124 mi) today and possibly much farther 2.5 

m.y. ago), the similarity is rather remarkable. 

The Colvillian and Bigbendian floras, before the end 

of the Gauss Chron, have a greater percentage of arboreal 

pollen at the beginning of their record and do not show as 

great a decrease upsection as the Begunov and Kutuyakh 

sequence, but a decrease is indicated (Nelson and Carter, 

1992), and the nearby and slightly younger Fish Creek section, which records full and extensive tundra during the 

time of the early Matuyama Chron, certainly indicates that 

this decrease in arboreal pollen continued. The two Alaskan localities have a less complete pollen sampling 

throughout the section, and the paleomagnetic record of the 

Krestovka sections is ambiguous at the time of the polarity 

change. These differences frustrate comparisons of the 

respective floras at the only identifiable point of contemporaneity, the Gauss-Matuyama boundary, but as far as comparisons can be made the timing and nature of floral 

changes are similar in both areas. 

Within the perspective of rapidly decreasing forests in 

the Arctic Ocean Borderland, the similarities in pollen content suggest that the Begunov Suite is only slightly older 

than the Kutuyakh Beds. The Begunov Suite had been considered to be early Pliocene in age because of its floral similarity to late Miocene floras of Siberia (Sher and others, 

1979). Based on the above relationships, the Begunov Suite 

more likely was deposited during the late part of the Gauss 

Chron, slightly before the Kutuyakh Beds and after the 

opening of the Bering Strait. No reversed magnetic polarity 

has been reported in either Alaska or the Y akutia, a polarity that would suggest that the Begunov or the Colvillian 

transgression was as old as the Kaena Subchron of the 

Gauss Chron, ending 2.92 million years ago. 

Permafrost is not indicated in the sediments of the 

Colvillian or Bigbendian deposits or the Begunov Suite; 

the spruce-birch forest in both areas suggest a lack of permafrost. Evidence of permafrost was first reported by Sher 

and others (1979) in the youngest part of the Kutuyakh 

Beds, upsection from the Gauss-Matuyama polarity reversal. Permafrost may first have appeared in northern Alaska 
MARINE AND PROXY RECORDS OF THE FIRST COOLING CYCLE 15 

at the time of the change of tundra types in the Fish Creek 

section underlying the mammal fauna but also above the 

Gauss-Matuyama polarity reversal. The absence of physical evidence of permafrost in the Fish Creek section could 

easily be due to the marine depositional environment. 

Plioctomys mimomiformis makes its only North 

American appearance in the Fish Creek fauna, but it is an 

abundant part of the reversely polarized Kutuyakh faunas 

of Yakutia. Phenacomys gryci is the other microtine from 

the Fish Creek fauna; it is the most primitive known species of the genus and is inferred to have been derived from 

the Asian genus Cromeromys (Repenning and others, 

1987). A species of Phenacomys that is inseparable from P. 

gryci, associated with Cromeromys, is known from a locality 4 km (2.5 mi) upstream from the Krestovka sections. 

This upstream locality is believed to be somewhat younger 

than the preserved Kutuyakh Beds (Sher and others, 1979). 

The specimen of Phenacomys is the only Old World record 

of the genus, which is known only from Beringia at this 

early date. Thus the Fish Creek fauna of Alaska and the 

late Kutuyakh faunas (including the isolated locality) share 

two unique records for their respective continents. 

The similarities between the Colvillian-BigbendianFishcreekian sections of central Beringia and the BegunovKutuyakh units of western Beringia are remarkable and 

suggest close contemporaneity of deposition. The differences suggest greater aridity and a steppe-forest mosaic 

flora in the more continental environment of Y akutia. 

MARINE AND PROXY RECORDS OF 

THE FIRST COOLING CYCLE 

We are not aware of any terrestrial deposits in the 

Arctic Ocean Borderland younger than the Kutuyakh Beds 

ofYakutia (about 2.38 m.y. ago) and older than 2.0 million 

years-the time of the culmination of the late Pliocene 

cooling cycle of the Arctic Ocean Borderland. 

The absence of terrestrial records is so widespread 

throughout the borderland that the question of a cause 

comes to mind. The depositional lacuna existed during a 

period of Northern Hemisphere continental ice accumulation, and, possibly, the resulting lower sea level made erosion, rather than deposition, the principal process in the 

Arctic Ocean Borderland. A similar lack of deposits during 

glacial periods is evident in the Atlantic coastal area of the 

United States (T.A. Ager, written commun., 1992). Marine 

records of this time period in the deposits of the Arctic 

Ocean and adjacent seas bear directly on the history of this 

coldest and final part of the first cooling cycle. 

Marine records show (1) the activity of the Scandinavian ice sheet as indicated by evidence from the Norwegian 

Sea; (2) deposition within the central Arctic Ocean north of 

Canada at the eastern end of the Alpha Ridge; and (3) the 

same record in the western Arctic Ocean from the Mendeleyev Ridge, north of Y akutia and the Siberian Shelf. The 

dating of these records is supported by lower latitude climatic records in Europe and North America. 

NORWEGIAN SEA: 

SCANDINAVIAN ICE SHEET 

Three Ocean Drilling Program cores were collected on 

the Voring Plateau in the Norwegian Sea, at sites 642, 643, 

and 644 (Jansen and others, 1988). Site 644 (figs. 1, 2) has 

the best paleomagnetic control and the least core disturbance, and it is the closest to the coast of Norway; it thus 

provides the best record of ice-rafted debris resulting from 

activity of the Scandinavian ice sheet (fig. 2). In the following summary we discuss temporal correlations to other 

events of the late Pliocene and early Pleistocene cooling 

cycles of the Arctic Ocean Borderland and elsewhere in the 

Northern Hemisphere. 

Site 644 records the last 2.85 million years of history 

in the Norwegian Sea. Activity of the Scandinavian ice 

sheet is indicated in the cores by the percentages of icerafted debris derived from this ice sheet when it flowed to 

the sea and dispersed as icebergs (fig. 2). Jansen and others 

(1988) reported that major expansion of the Scandinavian 

ice sheet began about 2.56 million years ago, before the 

end of the Gauss Chron and before the earliest permafrost 

in Y akutia and, by inference, Alaska. This was about 

120,000 years before the major lowering of sea level indicated in the oxygen isotope records of the World Ocean 

(stage 100, fig. 2). Stage 100 must reflect Antarctic, as well 

as Arctic, changes because it does not appear to represent 

the greatest accumulation of continental ice in the Northern 

Hemisphere during the first cooling cycle. 

About 2.56 million years ago the microtine bog lemming genera Plioctomys and Mictomys (a genus deriving 

from Plioctomys) dispersed southward from Beringia into 

the United States (Repenning and Grady, 1988). In the 

United States the date of this dispersal has been approximated (by interpolation between the Kaena Subchron and 

the end of the Gauss Chron) at about 2.6 million years ago 

on the basis of the earliest records of these immigrant 

genera between 10 and 15 ft (3.05--4.57 m) downsection 

from the Gauss-Matuyama magnetostratigraphic boundary 

in Texas, Arizona, and Idaho (Repenning, 1987). Because 

the dispersal probably was climatically stimulated, the 

2.56-million-year date of initial dramatic Scandinavian ice 

activity is considered a closer approximation. 

In the core from site 644, peaks of major glaciation 

(marked by high percentages of ice-rafted debris, comparable in size to those recorded during the late Pleistocene but 

not as frequent) persisted until about 2.0 million years ago, 

with relatively strong peaks between 2.2 and 2.0 million 

years ago (by interpolation between the Olduvai Subchron 

and the end of the Gauss Chron). The strongest peak 
16 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

correlates exactly with the "pre-Nebraskan till" in Iowa, 

which includes the 2.14-million-year-old beginning of the 

Reunion Subchron (Easterbrook and Boellstorff, 1984), 

and with oxygen stage 82 in oceanic records (fig. 2). 

At the eastern end of the Alpha Ridge, north of Ellesmere Island, Scott and others (1989) reported the intermittent presence of calcareous benthic foraminifers in CESAR 

core 14 between the Gauss-Matuyama boundary and just 

before the first Reunion normal-polarity excursion. These 

foraminifers were interpreted as indicating perennial ice 

cover. In that the Fish Creek fauna indicates that ice cover 

did not extend to the shore as recently as 2.4 million years 

ago, the presence of these calcareous foraminifers in 

CESAR core 14 suggests that perennial ice formed on the 

central Arctic Ocean long before it extended to the shores 

in the winter. The appearance of these foraminifers also 

suggests that the intermittent presence of perennial ice 

cover began shortly after initial activity of the Scandinavian ice sheet. 

Following the first period of glacial activity, a quiet 

period in Scandinavian ice activity (correlative with a lack 

of ice cover in the central Arctic Ocean at CESAR site 14) 

is indicated at site 644 immediately following the Reunion 

Subchron between 2.0 and about 1.7 million years ago 

(ending during the late part of the Olduvai Subchron). During this period, Scandinavian ice-debris percentages are at 

near-background levels; two minor exceptions of increased 

debris are at the beginning of the Olduvai Subzone (Jansen 

and others, 1988, fig. 8) (fig. 2). 

Before the Olduvai Subchron the site 644 record of no 

ice activity correlates closely with a time of warm climate 

from northern Greenland (Funder and others, 1985). During the first of the two minor peaks of rafted debris recording the brief break in warmth at the beginning of Olduvai 

Subchron, the microtine genus Allophaiomys dispersed 

southward from Beringia into the United States. Allophaiomys is also thought to have entered Beringia from lower 

latitude Asia at this time; it was a rodent that lived in temperate to subtropical climates and was not likely to have 

entered Beringia when arctic climate prevailed during maximum Northern Hemisphere glaciation (Repenning, in 

press). In North America, Allophaiomys is first found at, or 

just below, the base of the Olduvai Subzone in Colorado, 

immediately preceding in time the brief increases in icerafted debris in the Norwegian Sea. The genus appears in 

Holland (from the Black Sea area to the south) immediately after this minor peak of ice-rafted debris in the Norwegian Sea record, first occurring in Holland with late 

Tiglian warm floras in the older part of the Olduvai Subzone (van Kolfschoten, 1990a). 

The end of the period of minor to no ice activity of the 

Scandinavian ice sheet correlates well with the beginning 

of the Eburonian cold floral age of Holland (beginning 

about 1.7 m.y. ago late in the Olduvai Subchron) (van 

Kolfschoten, 1990a). At this time, the microtine genus 

Phenacomys (first known from the 2.4-million-year-old 

Fish Creek fauna of central Beringia) dispersed southward 

into the United States, where its oldest record is below the 

top of the Olduvai Subzone (about 1.7 m.y. ago) in Idaho 

(Repenning, in press). These ages are approximations 

based on the relative positions of the increase in ice-rafted 

debris of the Norwegian Sea, of the oldest Eburonian flora 

in Holland, and of Phenacomys in Idaho within the Olduvai 

Subzone. 

The record of ice-rafted debris between the beginning 

of the Eburonian and 1.2 million years ago is relatively 

minor at site 644 (Jansen and others, 1988). Only a few 

modest peaks are concentrated in the older part of this core 

segment, between 1. 7 and 1.5 (or perhaps 1.6) million 

years ago (fig. 2), and these are separated by longer periods 

of near-background percentages. These peaks would seem 

to correlate with the Eburonian cold flora of Holland, 

although the termination of this floral age is not dated. The 

remaining period, between 1.5 (or perhaps 1.6) and 1.2 

million years ago, has lower peaks and longer periods of 

near-background percentages of debris; it probable correlates with the Waalian warm floral zone of Holland, 

although again this floral zone is not dated except as following the Eburonian and between the Olduvai and 

Jaramillo events. 

Jansen and others (1988) noted that, during this quiet 

period, the marine climate was probably more latitudinally 

zoned: the water of the Norwegian Sea was more stable 

and less mixed with North Atlantic water than at present, 

and the thermal gradient from this sea to the North Atlantic 

was much steeper. The southward restriction of the North 

Atlantic surface current pattern and reduced flow to the 

Atlantic of Arctic water may have accentuated the warmth 

of the Waalian floral age in Holland and may have chilled 

the surface of the Arctic Ocean, permitting intermittent 

perennial ice to develop for the second time on the central 

Arctic Ocean. 

At about 1.2 million years ago the ice-rafted debris 

percentage peaks became greater in core at site 644 and 

periods of near-background percentages became very brief. 

Although poorly dated by interpolation between the Olduvai and Jaramillo Subchrons, this activity probably 

correlates with the beginning of the Menapian cold flora 

age of the Netherlands. As reviewed by Jansen and others 

( 1988), these peaks in ice-rafted debris persisted until 

600,000 years ago, after which very few of the low percentage points are at near-background levels and none persisted for a recordable length of time; despite evidence 

from more closely spaced samples, there were only very 

brief times of low glacial activity. 

· In Holland, Zagwijn and de Jong (1984) introduced 

the term Bavelian for a "warm" floral age between the 

Menapian and the Cromerian (the first glacial advance of 

the Ice Ages, not well marked by extensive ice records in 

Europe). The Bavelian includes, from the base upward, an 
MARINE AND PROXY RECORDS OF THE FIRST COOLING CYCLE 17 

"inter-glacial" (warmer flora), a "glacial" (colder flora), a 

second "inter-glacial," and a second "glacial." 

The Bavelian began just before the Jaramillo Subchron and ended shortly after it. Although not appearing 

very "warm" in the ice-rafted debris record of the core 

from site 644 or in the immigrant microtines of Holland 

(which, according to van Kolfschoten (1990b), include the 

first Microtus, a cool climate indicator), the climatically 

alternating floral events are represented in the debris percentages of core at site 644 in proper relationship to the 

Jaramillo Subchron. 

The difference in prominence of Jaramillo warm climates between the floras of Holland and the Scandinavian 

ice-rafted debris of the Norwegian Sea could have resulted 

from the exaggerated temperature gradient between the 

Norwegian Sea and the North Atlantic. This steep temperature gradient did not begin to diminish until after the 

Jaramillo Subchron (Jansen and others, 1988, fig. 9). In 

addition, the difference between the warm Bavelian flora 

of Holland and the prominence of ice-rafted debris in the 

Norwegian Sea may actually represent a correlation 

between warmer climate and ice accumulation (warm climate in Holland=warm North Atlantic=greater atmospheric 

moisture=greater precipitation on the Scandinavian and 

Laurentian ice sheets). 

The end of the Bavelian was the beginning of the Ice 

Ages and is marked by the first of many prominent peaks 

in ice-rafted debris in the core of site 644 at about 850,000 

years, although the age indicated by Jansen and others 

(1988) is slightly different because they used a different 

age for the beginning of the Brunhes Chron (730,000 years 

ago; we use 790,000 years). 

At this time the Cromerian of northwestern Europe 

began, the microtine genera Lasiopodomys and Terricola 

dispersed from Beringia into the United States (Repenning, 

in press), and the frrst of several ice advances called 

"Nebraskan" in North America is recorded in the Mississippi River drainage basin. This time also correlates with 

oxygen stage 22 in records of the World Ocean but is not 

evident in the isotope record of site 644 because of the lack 

of full mixing with Atlantic waters and because of the 

effects of carbonate dissolution in the Norwegian Sea and 

Arctic Ocean. 

The history of the Scandinavian ice sheet, as revealed 

in core 644 on the Voring Plateau of the Norwegian Sea, 

correlates, in varying degrees of precision, with the history 

of the two cooling cycles of the Arctic Ocean Borderland. 

Climatic events recorded in the Colvillian-BigbendianFishcreekian sequence of Alaska, the Begunov-Kutuyakh 

sequence of Y akutia, and southward microtine dispersals 

from Beringia into the United States correlate exactly, to 

the extent that the ages of events shown at site 644 can be 

interpolated. 

The missing history in the borderland records between 

about 2.38 and 2.0 million years ago is one of repeated ice 

peaks that reached a maximum in frequency and degree at 

about 2.1 million years ago. At 2.14--2.12 million years 

ago, the frrst normal-polarity excursion of the Reunion 

Subchron was recorded in glacial till in Iowa and adjacent 

regions of the United States. Before this excursion, bottom 

sediments on the eastern part of the Alpha Ridge provide 

evidence indicating the frrst development of perennial ice 

cover on the central Arctic Ocean. 

Just after the second Reunion normal-polarity excursion, about 2.0 million years ago, continental glaciation in 

Iowa was lacking and the climate in northernmost Greenland was extremely warm (Fonder and others, 1985); the 

first cooling cycle of the Arctic Ocean Borderland was at 

an end. 

The last break in the pattern of ice-rafted debris at site 

644 noted by Jansen and others (1988) was at about 

450,000 years ago, which is when the microtine Microtus 

dispersed southward into the United States. Microtus had 

dispersed as far south as southern Canada much earlier 

(Repenning, in press). Scott and others (1989) suggested 

that permanent perennial ice cover did not develop on the 

Arctic Ocean until about 300,000 years ago on the basis of 

their correlation of isotope stages. Due to very slow depositional rates (1 mm/1,000 yr) and the prevalence of samples 

too small for isotope measurement, their isotope stage 

identifications are questionable. This time period is later 

than the events discussed in this report. 

CENTRAL ARCTIC OCEAN: 

EASTERN ALPHA RIDGE 

Scott and others (1989), in their study of cores taken 

by the Canadian Expedition to Study the Alpha Ridge 

(CESAR), reported an abrupt replacement of agglutinated 

benthic foraminifers by calcareous benthic foraminifers at 

about the Gauss-Matuyama boundary. This replacement 

ended about 2.15 million years ago, before the older excursion of the Reunion Subchron as recorded in CESAR core 

14 on the eastern end of the Alpha Ridge of the central 

Arctic Ocean. 

This replacement is in the upper part of lithostratigraphic unit AB of Mudie and Blasco (1985) and consists 

of about three peaks of calcareous benthic foraminifers corresponding, in a general way, with the pattern of oxygen 

isotope fluctuations in the World Ocean between 2.48 and 

2.15 million years ago. Scott and others (1989) inferred 

that this change represents the first appearance of perennial 

ice cover of the Arctic Ocean. If so, one might assume that 

some of the lower sea-level peaks of the oxygen isotope 

records in the World Ocean would correlate with ice cover 

of the central part of the Arctic Ocean following the Fish 

Creek-Kutuyakh deposition in Beringia. With the exception of isotope stage 96, about 2.4 million years ago, these 
18 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

isotope stages show less eustatic lowering than isotope 

stage 82 at the base of the Reunion Subzone. 

The next record of periodic perennial ice cover of the 

Arctic Ocean is associated with a normal paleomagnetic 

event that Scott and others ( 1989) attributed to the Gilsa 

event (possibly shortly following the Olduvai). Unusual 

changes in the rate of sedimentation are implied (but not 

discussed in Scott and others) by identifying this normal 

event as the Gilsa. If it does represent the Gilsa, it also is 

not clear how this change to calcareous microfauna correlates with the core from site 644 in the Norwegian Sea. We 

believe that the normal event referred to the Gilsa by Scott 

and others (1989) correlates best with the increase in icerafted debris 1.2 million years ago at site 644, and thus perhaps the normal event reported by Scott and others is the 

Cobb Mountain event, 1.1 million years old (Mankinen and 

others, 1978), rather than the Gilsa event (see discussion in 

Repenning, in press). 

If the Cobb Mountain event is represented, then rates 

of sedimentation in CESAR core 14 are more uniform than 

those suggested by the correlation with the Gilsa event, and 

correlation with site 644 and with the coldest climate in 

Yakutia (to be discussed) is exact. Thus, an inferred correlation with the Menapian cold flora of Holland is plausible. 

However, the identity of the normal event called Gilsa in 

CESAR core 14 is in doubt. 

On the basis of changes in depositional character, 

Herman and others (1989) suggested that at about this time 

(about 1.1 m.y. ago and just before the Jaramillo Subchron, 

although they did not detect the Cobb Mountain Subchron), 

the Arctic Ocean became permanently ice covered. 

WESTERN ARCTIC OCEAN: 

MENDELEYEV RIDGE 

Herman and Hopkins (1980) and Herman and others 

(1989) identified three widespread depositional units in the 

bottom sediments on the Mendeleyev Ridge in the core 

from site T3-67-12, north of Yakutia and at the northern 

edge of the Siberian Shelf of the Arctic Ocean. The oldest, 

unit III, is characterized by oxygenated deposits of red 

clays containing iron and manganese micronodules. 

Deposition of unit II began during the latest part of 

the Gauss Normal-polarity Chron, presumably at about the 

time of the first activity of the Scandinavian ice sheet and 

during the reduction in forests of the Kutuyakh and 

Bigbendian deposits. Unit II is characterized by poorly 

oxygenated gray and tan clays and abundant coarse icerafted debris. 

Deposition of unit I began approximately at the beginning of the Ice Ages; it is believed by Herman and others 

(1989) to show the first development of ice cover on the 

Arctic Ocean. 

Using paleomagnetic determinations of core T3-67-

12, Herman and others (1989) approximated the age of the 

unit III-unit II boundary at 2.5 million years and attributed 

the lithologic change to the establishment of salinity stratifi.cation caused by an increased surficial contribution of 

freshwater from rivers and from melting icebergs. They did 

not believe that development of salinity stratification in the 

Arctic Ocean during deposition of unit II indicated perennial ice cover and stated that the water temperatures indicated by the dominant planktonic foraminifers are 

incompatible with the presence of a perennial ice cover. 

Core TI-67-12 is just off of the Siberian Shelf and may 

have been much closer to shore 2.5 million years ago than 

now; the suggestion of no ice cover may not be contradictory to the record from the central Arctic Ocean. 

The conclusions of Herman and others ( 1989) are 

strengthened and enlarged upon by Scott and others (1989), 

although the latter authors disagree about the lack of ice 

cover. The CESAR results suggest a period of intermittent 

perennial ice cover on the Arctic Ocean 2.48-2.15 million 

years ago and the first common, but still intermittent, 

perennial ice cover questionably beginning about 1.2 million years ago. Unlike Herman and others, Scott and others 

suggested that permanent perennial ice cover did not begin 

about 850,000 years ago but rather about 300,000 years 

ago (isotope stage 8). 

The temporal similarity is obvious in (1) the beginning of deposition of unit II at site T3-67-12, (2) the 

beginning of deposition of Arctic Ocean unit AB at the 

CESAR sites, (3) the beginning of glacial activity of the 

Scandinavian ice sheet as indicated in records from site 

644, and (4) the dispersal of bog lemmings southward into 

the conterminous United States, which implies cooler climates in lower latitude North America. All of these events 

took place shortly before the end of the Gauss Chron. 

The oldest records of a marine ostracode fauna comparable to the living fauna of the Arctic Ocean, of extensive tundra pollen in the Fish Creek and late Kutuyakh 

sections, and of earliest permafrost in the Krestovka sections (all shortly following the Gauss Chron) are also synchronous and a somewhat later part of the same climatic 

change. Changes in the ecosystem, particularly to colder 

ocean temperatures, seem to have been rather abrupt at 

about the Gauss-Matuyama polarity boundary. 

SUMMARY OF THE FIRST COOLING 

CYCLE 

At the time of the opening of the Bering Strait (3.0 

m.y. ago) the Arctic Ocean Borderland was forested and 

lacked permafrost and the Arctic Ocean had a coldtemperate marine climate that was free of ice cover. The 
SUMMARY OF THE FIRST COOLING CYCLE 19 

Arctic Ocean was at a high eustatic level and covered 

much of the Alaskan North Slope. Treeline was at lat 

79°30' N ., farther north than most of the shoreline of the 

Arctic Ocean Borderland, and there is no indication of ice 

cover in western or central Arctic Ocean cores. Sprucebirch forests along the coast were more diverse than today 

and contained both extinct taxa and taxa that now live only 

far to the south. 

In temperate latitudes this time equates with the 

Rebielice small-mammal age of Europe, the Reuverian floral age of northwestern Europe, the late MN 16 zone of the 

Mediterranean region, the late Kujalnikian north of the 

Black Sea (Tatiana Svetlitskaya, written commun., 1992), 

the early part of the Kisichian in Western Siberia, the Itantsian in western Transbaikalia, Chinese zone III as correlated by Zheng and Li (1990), and the Blancan IV 

microtine age of the United States (Repenning and others, 

1990). Each of these areas had its own local climate, but 

all of the areas were climatically related to the Arctic 

Ocean and, in general, indicate warmth; in some cases, 

decreasing warmth is indicated. 

Toward the end of this time (toward the end of the 

Gauss Chron, 2.48 m.y. ago) the forests were thinning in 

Beringia to mosaics of forest-tundra or, in Asia, foreststeppe because grasses were more common in western Beringia than in the east. The Scandinavian ice sheet began to 

develop about 2.56 million years ago; at this time the Arctic Ocean became density stratified as eustatic sea level 

began to lower, although much of the Alaskan North Slope 

was still submerged. About 100,000 years after increased 

activity of the Scandinavian ice sheet was evident in the 

Norwegian Sea, evidence of ice-transported debris on the 

Arctic Ocean floor was more abundant and intermittent 

perennial ice cover began to form on the central Arctic 

Ocean; for at least 60,000 years, however, ice cover did not 

extend to the coast in the winter. 

The southward movement of some microtine rodents 

(bog lemmings) from Beringia to the United States about 

the time that Scandinavian ice growth became apparent 

(the arrival in lower latitudes of other northern mammals is 

not demonstrated at this time) suggests that increasingly 

cold weather extended far south in North America. The bog 

lemmings of today still inhabit northern forests of Canada 

and the northeastern United States. 

By perhaps 2.45 million years ago, shortly after the 

end of the Gauss Chron and the beginning of the 

Matuyama Chron 2.48 million years ago, treeline had 

shifted south of the average latitude of Beringia (70° N.), 

and full tundra was associated with permafrost in both 

eastern and western Beringia. About this time the ostracode fauna of the Arctic Ocean changed with dramatic 

abruptness to an assemblage characteristic of subfrigid or 

frigid temperatures, but the mollusk fauna changed little; 

in combination, the faunal evidence indicates a subfrigid 

marine climate that was still warmer than today but colder 

than earlier records. The shoreline of the Arctic Ocean had 

receded northward of lat 70° N. in Alaskan Beringia with 

global eustatic lowering of sea level. By 2.4 million years 

ago, ice cover had not yet developed on the Arctic Ocean 

to the extent of reaching the Alaskan coast, and marine life 

that was dependent on year-round ice-free water still survived. Intermittent perennial ice cover is believed, however, to have been present in the central part of the Arctic 

Ocean, and mountain glaciation and the Scandinavian ice 

sheet were contributing icebergs to the Arctic Ocean as 

well as to the Norwegian Sea from where they drifted into 

and around the Arctic Ocean. 

The remainder of the frrst cooling cycle of the Arctic 

Ocean region is not represented by deposits on land, 

possibly a result of lower sea levels that encouraged continental erosion rather than deposition. This last chapter of 

the first cooling period, during the earliest part of the 

Matuyama Chron, is recorded in the Norwegian Sea where 

activity of the Scandinavian ice sheet is indicated by icerafted debris. Intermittent increases in glacial activity climaxed at about 2.1 million years ago, when continental ice 

deposited till as far south as Iowa in the United States. 

In temperate latitudes the later part of the first cooling 

cycle equates with the European Villanyian age, the Praetiglian floral age of northwestern Europe, zone MN 17 of 

the Mediterranean area, the late part of the Kisichian of 

Western Siberia, the earlier Dodogolan of western Transbaikalia, Chinese zone IV as correlated by Zheng and Li 

(1990), and the Blancan V microtine age of the United 

States. During this time continental ice flowed south down 

the Mississippi River Valley to Iowa, and Blancan V 

microtine faunas both east and west of the Rocky Mountains are marked by the presence of bog lemmings (Plioctomys and Mictomys). In eastern Asia the ancestor 

(Allophaiomys) of the temperate to arctic micro tine voles 

with rootless cheek teeth (Microtus, Pitymys, Terricola) 

evolved out of a subtropical to temperate late Pliocene 

form with rooted cheek teeth (Mimomys). In North America the primitive, ubiquitous archaeolagine genus Hypo/agus was replaced by the modem, ubiquitous leporid Lepus. 

At the close of the first cooling cycle of the Arctic 

Ocean Borderland, treeline was well south of its modem 

limits in both eastern and western Beringia, and continental 

ice had spread southward into the United States. In the 

brief time interval of about 100,000 years, by 2.0 million 

years ago, a treeline flora and an absence of Arctic Ocean 

ice cover is known from the northern end of Greenland at 

lat 82°30' N. (Funder and others, 1985). This climatic 

warmth was greater than that known at the beginning of the 

first cooling cycle, about 3 million years ago, and it 

marked the beginning the second cooling cycle. 
20 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

2.0-0.85 MILLION YEARS AGO: 

THE SECOND COOLING CYCLE 

Records from Scandinavia (Jansen and others, 1988), 

the central United States (Easterbrook and Boellstorff, 

1984), and Greenland (Fonder and others, 1985) indicate 

that, by 2.0 million years ago, the late Pliocene continental 

glaciation of the Northern Hemisphere had ended and the 

Arctic Ocean and its borderland were much warmer. The 

ostracode faunas from the Kap K0benhavn Formation of 

northernmost Greenland at lat 82°30' N. indicate conspicuously warm waters in the Arctic Ocean and a treeline flora 

on land about 2.0 million years ago; it was much warmer 

than at lat 70° 16' N. about 2.4 million years ago, when full 

tundra covered the Arctic Ocean Borderland, or than the 

modem Arctic Ocean (Brouwers and others, 1991 ). 

To judge from the records on Meighen Island, the Kap 

K0benhavn fossil flora record indicates an Arctic Ocean 

climate warmer than that which existed at the opening of 

the Bering Strait a million years earlier. No perennial ice 

cover is suggested at any location in the Arctic Ocean at 

this time, and dispersing microtine rodents entering North 

America from Asia indicate that at the lower latitudes of 

Beringia a temperate climate was present 2.0 million years 

ago. 

In the conterminous United States winter freezes were 

rare across the northernmost Great Plains and in at least the 

southern half of inland California. To the east of the Rocky 

Mountains, on the northern Great Plains, summer rains 

were common, summer humidity was high, and maximum 

summer temperatures were lower than today (Repenning, 

1990). Northwestern Europe was characterized by the 

warm Tiglian flora. 

As the second cooling cycle progressed, Scandinavian 

glaciation again reached maximum activity between 1.2 

and 1.1 million years ago, as indicated in the core at site 

644, just before a normal-polarity excursion called the 

Cobb Mountain Normal-polarity Subchron (1.10 m.y. ago) 

(fig. 2). At this time full tundra returned to the Kolyma 

Lowlands of Y akutia, and, if paleomagnetic correlations of 

CESAR core 14 are as suggested by Repenning (in press), 

perennial ice cover was intermittent in the central Arctic 

Ocean. 

Between 2.0 and 1.2 million years ago the nonmarine 

climate of the Arctic region vacillated between cool and 

warm periods, beginning late during the Olduvai Subchron 

(about 1.7 m.y. ago) at site 644 with the return of moderate 

peaks of ice-rafted debris from the Scandinavian ice sheet. 

There are no records of the ecosystem of the Arctic 

Ocean Borderland during the Olduvai Subchron, and the 

history must be inferred from records of the Arctic Ocean 

and from lower latitude proxy records. The next record 

younger than the initial warm period 2.0 million years ago 

is from reversely polarized deposits before the Jaramillo 

Subchron in Yakutia (Virina and others, 1984) and from 

the Cape Deceit Formation on the Seward Peninsula of 

Alaska that may be 1.5 million years old, deposited shortly 

following the Olduvai Subchron. The age of these deposits 

is inferred primarily from an incompletely known history 

of the microtine rodents (Repenning, in press). Based on 

these microtine records, the climatic record in the Arctic 

Ocean Borderland can be confidently, but approximately, 

correlated with more complete oceanic records and with 

continental records in lower latitudes. 

During the Jaramillo Subchron (970,000-900,000 

y.b.p.) of the Matuyama Chron, two relatively brief warm 

periods interrupted this second cooling trend, as indicated 

in Holland, in Yukon Territory of Canada, and in Y akutia. 

These two warm periods are not clearly detectable in Norwegian Sea or Arctic Ocean records, for uncertain reasons. 

Following the warm periods, the cooling continued, and the 

Ice Ages began 850,000 years ago with the advance of continental glaciation down the Mississippi River Valley in the 

United States. This first ice advance, along with several 

subsequent ones, has been referred to as the "Nebraskan" 

glaciation. 

The only presumed record of the ecosystem in the 

Arctic Ocean Borderland at the close of the second period 

of climatic deterioration and the beginning of the Ice Ages 

is in the type section of the Olyor Suite in Yakutia, and it 

has not been studied in detail. The Y akutian record may 

not be complete, but it does contain reversely polarized 

deposits younger than the Jaramillo Subchron, a situation 

possibly unique in the Arctic Ocean Borderland. 

Scott and others (1989) suggested that permanent 

perennial ice cover of the Arctic Ocean was not established 

until perhaps 300,000 years ago and that before this time 

perennial ice cover of the Arctic Ocean was intermittent. 

Intermittent perennial ice cover began just before an S-cmthick normal-polarity event in CESAR core 14, here 

believed to record the Cobb Mountain event. 

GREENLAND: 

KAP K0BENHA VN FORMATION 

Following the first Northern Hemisphere glacial maximum 2.1 million years ago, a spectacular return of forests 

to the shores of the Arctic Ocean is indicated in the record 

of the Kap K0benhavn Formation of northernmost Greenland (Fonder and others, 1985; Bennike, 1990), which is 

the northernmost terrestrial fossil record in the world. This 

warming is also recorded in (I) the abrupt reduction of 

activity of the Scandinavian ice sheet; (2) lower latitude 

North American faunas that indicate subtropical climates 

near the Canadian border (Repenning, 1990); (3) the St. 

Erth Beds of Cornwall, England, that indicate subtropical 

temperatures for the North Atlantic (Jenkins and Houghton, 

1987); (4) the Tiglian warn floras of northwestern Europe; 
2.0-0.85 MILLION YEARS AGO: THE SECOND COOLING CYCLE 21 

and (5) immigration of the subtropical to warm-temperate 

microtine rodent to the United States. 

Fonder and others (1985) described a sequence of 

shallow-marine and marginal-marine deposits at Kap 

K~benhavn in the northernmost part of Greenland, at lat 

82°30' N. and about 525 mi (845 km) south of the North 

Pole (fig. 1). The older of two members of the Kap K~benhavn Formation, member A (fig. 2), records shallow 

glaciomarine deposition, and the younger, member B, is 

composed of high-energy marginal-marine and coastal 

deposits. Member B, representing as much as threequarters of the thickness of the formation, is subdivided 

into three units, referred to as B 1, B2, and 83, in order of 

deposition. 

The oldest quarter, member A, contains a modern 

Arctic Ocean mollusk fauna, with no extinct species, and 

dropstones thought to indicate nearby glaciation (Bennike, 

1990). Deposition was in the absence of wave action presumably because of extensive ice cover. The possibility of 

ice cover suggested in CESAR core 14, about 550 mi (886 

km) to the west-northwest, began about 2.48 million years 

ago and ended about 2.15 million years ago (Scott and 

others, 1989). 

The mollusk fauna of member A is the oldest known 

Arctic Ocean mollusk fauna containing no extinct taxa; the 

Fish Creek fauna (2.4 m.y. old) contains one extinct form 

(Marincovich in Repenning and others, 1987). The fauna 

and the glaciomarine environment of member A indicate 

that it probably is younger than the Fish Creek fauna and is 

equivalent to unit II in the western Arctic Ocean (Herman 

and others, 1989) and to unit AB of CESAR core 14 (Scott 

and others, 1989). 

Fossils and sedimentary structures of member B indicate that the Arctic Ocean could not have been perennially 

ice covered during deposition. Unit B3 contains evidence 

of a treeline biota with a mixture of boreal trees (trunks as 

much as 7.1 in (18 em) in diameter and 15 ft (4.5 m) long) 

and tundra shrubs with a preponderance of forest beetles 

(with lesser numbers of tundra beetles) and temperate to 

low arctic coastal moss (Bennike, 1990; Bennike and 

Bocher, 1990). This biota indicates that the land adjacent to 

the ice-free ocean had a climate similar to that of today 27° 

farther south (1,900 mi, 3,000 km) (Bennike, 1990). An 

upsection warming trend is indicated by these fossil 

records in the Kap K~benhavn Formation. Limited 

paleomagnetic study of member B indicates reversed polarity and that the section is not young enough to include the 

Olduvai Normal-polarity Subchron (Abrahamsen and Marcussen, 1986). The younger of the two events of the 

Reunion Subchron (ending 2.01 m.y. ago) may be present 

in member B, to judge from the paleomagnetic results of 

Abrahamsen and Marcussen (1986). 

Unit B3 (the youngest) is mostly composed of 

marginal-marine deposits. An ostracode fauna in its basal 

part (sample 11514, Geological Museum of Copenhagen 

University, locality 68) contains warm water forms ( Cytheromorpha macchesneyi, Cytheromorpha sp., Leptocythere 

spp., and Bythocythere sp.). This fauna is comparable to 

faunas from cold-temperate to warm-subfrigid marine 

provinces such as modern coastal Nova Scotia (Brouwers 

and others, 1991). 

Underlying unit B3, unit B2 is a marine silt containing 

mollusks now living in warmer Atlantic waters (0. Bennike, written common., 1987). Ostracodes in unit B2 indicate subfrigid water temperatures, which today are found in 

the Atlantic Ocean off of the coast of Norway and Newfoundland. The ostracode fauna indicates inner to middle 

shelf depths. 

The oldest unit in member B, unit B1, is a coastal 

deposit that contains a flora similar to that of unit B3 but 

with no trees. Except for the lack of arboreal plants, the 

dwarf shrub heath present in unit B 1 is similar to that of 

unit B3, and a subarctic climate near treeline is inferred (0. 

Bennike, written common., 1988). Ostracodes were recovered from only one sample from unit B1 (sample 11494, 

Geological Museum of Copenhagen University, locality 

74); the assemblage represents an inner shelf environment 

and is similar to the ostracode fauna from unit B2. 

Unit B3 (the youngest) contains an unusual association of modem and primitive rabbits; the Pliocene archaeolagine Hypolagus and the Pleistocene (and living) 

leporine Lepus are present. Lepus was not mentioned by 

Fonder and others (1985), but its provenance was confirmed by 0. Bennike (oral common., 1988). Only three 

other faunas in North America (Anita, Arizona; Borchers, 

Kansas; Big Springs, Nebraska) have this unusual association; two of these localities are 2.01 million years old, and 

the Anita fauna is not dated. Hypolagus is common in most 

North American localities older than about 2.0 million 

years, but Lepus is unknown, whereas Lepus is common in 

most fossil localities younger than 2.0 million years, but 

Hypolagus is unknown, these time relations suggest that 

unit B3 is close to 2.0 million years old. 

The relative warmth recorded in member B suggests 

that it postdates the time of till deposition in Iowa, peak 

glacial growth in the Scandinavian ice sheet, and the earliest period of ice cover on the central Arctic Ocean (all ending 2.1 m.y. ago). Because the normal polarity of the 

Olduvai event was not recorded in the paleomagnetic study 

of the Kap K~benhavn Formation, deposition must predate 

1.90 million years ago. The unusual association of rabbits 

confirms the age range of 2.1-1.9 million years and further 

suggests that unit B3 is constrained to an age of about 2.0 

million years. 

The progressive warming upsection indicates that the 

younger part of the Kap K~benhavn Formation records the 

transition from the first glaciation into a warm period; that 

is, the transition from the first cooling cycle into the second. The warming is also recorded in Holland, the Norwegian Sea, Great Britain, and the central United States as 
22 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

described in the beginning of this section. Its cause may 

relate to northward displacement of marine climate in the 

North Atlantic Ocean (Repenning, 1990), but, if so, a similar situation existed in the North Pacific, as indicated by 

the dispersal of microtine rodents, to be discussed. 

YAKUTIA: KRESTOVKA OLYOR SUITE 

In the Kolyma Lowlands of Y akutia, deposits of the 

Krestovka sections that are between 2.0 and 0.85 million 

years old are placed in the Olyor Suite of Sher (1971), and 

this record does not begin until about 1.4 million years ago. 

These deposits are reversely polarized in their older parts 

and normally polarized in their younger parts. The contained fauna of large mammals is a typical Ice Ages fauna 

(Sher, 1971) and was initially considered to represent the 

beginning of the Ice Ages; the Krestovka Olyor Suite was 

believed to represent deposition during the latest part of the 

Matuyama Reversed-polarity Chron and the earliest part of 

the Brunhes Normal-polarity Chron, perhaps 0.85-0.7 

million years ago (Sher and others, 1979; Sher, 1987) 

although there was no control of the younger time limit. 

Subsequently, Azzaroli (1983) indicated that the 

Olyor large mammal fauna was representative of his Galerian (early Ice Age) large mammal fauna, which dispersed 

throughout Eurasia during what he called the EndVillafranchian dispersal event. 

The only lapse of Azzaroli (1983) was not to consider 

the possible age of the source of the Galerian fauna, which 

he recognized as being the Arctic Ocean Borderland. The 

age problem came into focus in 1984, when Virina and 

others published their paleomagnetic study of the type 

Olyor Suite. They indicated that the Olyor (Galerian) fauna 

was known before the Jaramillo Subchron, whereas 

Azzaroli believed that his End-Villafranchian dispersal 

event, presumed to initiate the Galerian fauna throughout 

Eurasia, took place during or slightly after the Jaramillo 

Subchron. Recognition of the Jaramillo Subzone in the 

type Olyor Suite indicated that, in its source area, the Galerian fauna was older than the dispersal event. 

The microtine fauna of the Olyor Suite of the Krestovka sections had not been described in 1984, when 

Repenning, misled by differences in the use of taxonomic 

nomenclature in published faunal lists, indicated that the 

microtine fauna was much older than had been inferred in 

publications; the only critical microtine fossils from the 

faunas of the Olyor Suite illustrated in publication are 

those in Repenning (in press); attention is called to the differences in nomenclature in this report. On the assumption 

that he understood the meaning of the taxonomic names 

used, Repenning ( 1984) indicated that the climatic history, 

magnetic polarity pattern, and microtine history reported 

for the unit strongly suggested that the Olyor Suite in the 

Krestovka sections was between 2.4 and 2.0 million years 

old, representing the earliest part of the Matuyama Chron, 

during and before the Reunion Subchron. This opinion was 

restated in Repenning and others ( 1987) and has been cited 

in other publications. 

Unfortunately, the climatic history and magnetic 

polarity pattern present in the Olyor Suite are similar to 

those during both the first and the second cooling cycles of 

the Arctic Ocean Borderland, and correct identification of 

the microtine rodent fossils was the only key to separate 

these two similar segments of the history of the Arctic 

Ocean Borderland. Both interpretations were based upon 

reversed polarity in the older part of the section and normal 

polarity in the upper part. Both interpretations recognized a 

cooling trend with major glaciation and the coldest climate 

at the beginning of the normal polarity. The biochronologic 

significance of the microtine fossil called Allophaiomys 

was obscured by different definitions. 

Agreement between A.V. Sher (the principal investigator in the study of the Krestovka sections) and Repenning on the interpreted age of the Olyor Suite could not be 

achieved by correspondence. In 1990 Sher invited 

Repenning to visit Moscow and review the undescribed 

fossil faunas. Major differences in the application of systematic nomenclature were thus revealed and corrected. 

The revisions in the nomenclature and in the interpretation 

of the age of the Olyor Suite of the Krestovka sections are 

explained in Repenning (in press). 

The significant microtine faunas from the Olyor Suite 

in the Krestovka sections are now recognized as having lived 

during the Matuyama Chron before the Jaramillo Subchron 

(Repenning, in press), rather than before the Olduvai Subchron, as published faunal lists had suggested to Repenning 

(1984 ). The approximate age of the oldest part of the formation is now known to be about 1.4 million years, based on the 

dated records in the United States of lineages derived from 

Asia; the oldest part of the formation cannot be as old as the 

1.67-million-year-old termination of the Olduvai Normalpolarity Subchron because normal polarity of this event is 

not recognized in the sections. The youngest, temporally significant fauna lived immediately before the Cobb Mountain 

Normal-polarity Subchron (1.10 m.y. old) (Repenning, in 

press). These age limits apply to the older member of the 

Olyor in the Krestovka sections (called Olyor lila); the 

younger member (Olyor Illb) is normally polarized (a factor 

in the original age interpretation) and may or may not 

include the Jaramillo Subchron, as well as part of the 

Brunhes Chron. Sample sizes of the microtine faunas from 

the younger member (lllb) are small and do not permit close 

age approximations. 

The revised age interpretation is in satisfactory agreement with the inferred age for the type section of the Olyor 

Suite, located some 93 mi (150 km) north of the Krestovka 

sections. The type Olyor was deposited before, during, and 

after an event of normal polarity assigned to the Jaramillo 

Subchron (Virina and others, 1984; Sher, 1987). In a 
2.0--0.85 MILLION YEARS AGO: THE SECOND COOLING CYCLE 23 

second section in the same area, the type Olyor extends 

later, into the older part of the Brunhes Chron. 

Near the base of the Olyor Suite of the Krestovka sections, the proportion of shrub and arboreal pollen is about 

40 percent of the pollen spectrum (Giterman and others, 

1982), suggesting a flora similar to, but somewhat more 

diverse than, the boreal forest at the site today. Upsection 

in the upper part of the lower member (Ilia) of the Krestovka Olyor, just below and within deposits having normal polarity, the proportion of shrub and arboreal pollen is 

only 5 percent (Giterman and others, 1982); arboreal pollen 

is "almost absent" (Sher and others, 1979). Tundra, possibly near treeline, is indicated. This is the most severe environment recorded in the Krestovka Olyor and is more 

severe than that represented by the taiga at the locality 

today. Permafrost was present during deposition of Olyor 

lila (Sher and others, 1979). 

An inconsistency is evident when this decrease in 

shrub and arboreal pollen in the Krestovka sections is compared with the pollen record of the type Olyor Suite 

(outcrop 21 of Sher, 1971). Unlike the Krestovka sections, 

the type section of the Olyor Suite is in tundra today. 

Virina and others (1984) reported that the Jaramillo 

Subchron is in the middle of the type Olyor section. Yet, in 

the type section, Giterman and others (1982) reported that 

the percentage of shrub and arboreal pollen is no lower 

than 20 percent of the spectrum; in most of the section it is 

well above this value. The type section contains a pollen 

record of open forests and tundra, possibly near treeline, in 

an area now occupied by extensive tundra. It does not contain evidence of a cold period more severe than that indicated by the modem flora. 

The absence of evidence for a severe cold period during the Jaramillo in the type section, similar to that 

recorded in and just below the oldest normally polarized 

deposits of the Krestovka sections, casts doubt on correlation of the Jaramillo event to the Krestovka sections. Faunal considerations support this doubt (Repenning, in press). 

As noted by Sher and others (1979), Sher (1987), and 

Repenning (in press), the youngest significant microtine 

fauna in the Krestovka Olyor, locality KL0-9 (0.8 ft (0.25 

m) below the normally polarized deposits), contains several 

specimens of the genus Microtus that are the earliest record 

of this genus in the Krestovka sections and, as far as is 

known, in Beringia. They are of a primitive form and are 

characteristic of the earliest morphotypes of the genus 

found around the world, although these records differ 

greatly in age (Repenning, in press). This earliest record, 

and its primitive evolutionary stage, is inconsistent with the 

age for the beginning of the Jaramillo Subchron or younger 

normal-polarity events. 

In the Old Crow Basin of Yukon Territory, Canada, 

within the Beringian faunal province but in its eastern part, 

a large microtine fauna with a good number of Microtus 

morphotypes, some quite advanced, is known (R.E. 

Morlan, written commun., 1991) from deposits judged to 

represent the Jaramillo Subchron on faunal, climatic, and 

magnetic polarity criteria (Repenning, in press). Thus, in 

Beringia the genus was well established by the time of the 

Jaramillo event, which lasted 70,000 years. The deposits 

containing the advanced morphotypes of Microtus in the 

Old Crow Basin also contain a pollen flora indicating temperatures as warm as, or warmer than, today (J. V. 

Matthews, written commun., 1987), similar to the flora in 

the type Olyor and contrasting equally with the flora below 

and within the oldest normally polarized zone of the Krestovka Olyor. 

In the well-documented evolution of Microtus paroperarius in the United States it took 140,000 years to 

evolve "quite advanced" forms out of early morphotypes 

like those from locality KL0-9 of the Krestovka sections 

(Repenning, in press). If it took as long to evolve advanced 

morphotypes in Beringia, locality KL0-9 in the Olyor 

Suite of the Krestovka sections, western Beringia, should 

be a minimum of 70,000 years older and a maximum of 

140,000 years older than the Jaramillo event; however, 

locality KL0-9 is only 0.8 ft (0.25 m) beneath the base of 

the zone with normal polarity in the Krestovka sections. In 

addition, climatic evidence from this part of the section, 

discussed above, suggests that the oldest normal polarity in 

the Krestovka Olyor cannot represent the Jaramillo 

Subchron. 

Sher and others (1979) reported a brief change to 

reversed polarity about 1.6 ft (0.5 m) above the base of the 

normally polarized deposits in the Krestovka sections. This 

change is indicated by a single polarity determination in 

uncleaned samples but is indicated by three samples after 

cleaning (Sher and others, 1979, fig. 6). These three polarity determinations of cleaned samples suggest that the lowest normal-polarity deposits in the Krestovka sections may 

be less than 1.6 ft (0.5 m) thick. The presence of these 

higher reversed samples has not been considered significant, but it is possible that the lowest normal polarity of the 

Krestovka sections represents a brief event (its zone possibly less than a foot (0.3 m) thick) before the Jaramillo 

Subzone. 

This lowest normal polarity recorded in the Krestovka 

sections may represent the Cobb Mountain Subchron, 1.10 

million years old. If the first record of Microtus in the 

Krestovka sections was slightly below the Cobb Mountain 

Subchron, rather than slightly below the Jaramillo Subchron, a plausible length of time (comparable to that documented for the evolution of Microtus paroperarius in the 

United States) would be indicated for the evolution and 

diversification of Microtus found during the Jaramillo Subchron in eastern Beringia, as represented in the Old Crow 

Basin locality. 

The assignment of this earliest normal polarity to the 

Cobb Mountain Subchron would indicate a correlation of the 

most severe climate recorded in the Krestovka sections with 
24 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

(1) the second beginning of intermittent perennial ice cover 

on the Arctic Ocean, as reported from CESAR core 14 on 

Alpha Ridge (Scott and others, 1989) (the first beginning of 

intermittent perennial ice cover was between 2.4 and 2.1 million years ago, which was Repenning's 1984 interpreted correlation) and (2) the beginning of intense activity of the 

Scandinavian ice sheet recorded in the Norwegian Sea at site 

644 (Jansen and others, 1988, fig. 2). On the other hand, 

assignment of this short normal-polarity event to the 

Jaramillo Subchron would indicate a correlation of the most 

severe climate recorded in the Krestovka sections with (1) 

the Bavelian warm floral zone of Holland and (2) the Old 

Crow locality CRH-47 and the type Olyor, both of which 

contain evidence of a climate as warm or warmer than today. 

Assignment of this event to the base of the Brunhes would 

indicate a correlation with warm oxygen stage 19 and a correlation of the primitive and undiversified morphotypes of 

Microtus of locality KL0-9 with the advanced and diverse 

morphotypes of Beringia at the beginning of the Brunhes 

Chron, as illustrated by the fauna from the type Olyor (Sher, 

1987) and from locality CRH-47 in Canada that lived during 

the Jaramillo Subchron. 

Accordingly, the magnetically reversed and oldest part 

of the Olyor Suite in the Krestovka sections of Yakutia, as 

well as less than 1.6 ft (0.5 m) of overlying normally polarized deposits, is considered to be about 1.4-1.1 million 

years old. 

The presence of an essentially treeless pollen flora just 

before and during the 1.1-million-year-old Cobb Mountain 

Subchron in the Krestovka sections is correlated with other 

indications of cold climate from the central Arctic Ocean 

and Scandinavia and with the Menapian cold flora of the 

Netherlands. This cold period, ending about 130,000 years 

before the Jaramillo Subchron, is not recorded in the type 

section of the Olyor Suite and may be missing. 

The remaining normally polarized younger part of the 

Krestovka Olyor, perhaps 8.2 ft (2.5 m) thick, may represent the Jaramillo Subchron that is recorded in the type 

section, or it may represent part of the Brunhes Chron, or 

both. The microtine rodents are too poorly represented in 

this youngest part of the Krestovka Olyor to evaluate their 

age (A.V. Sher, oral commun., 1990); they do include more 

advanced types of Microtus (Sher, 1987), however, and 

cannot be older than the Jaramillo event insofar as knowledge of Beringian microtine history is known at present. In 

the younger, normally polarized part of the Krestovka 

Olyor the percentage of arboreal pollen is greater than, and 

the climatic implications are similar to, that of pollen from 

the Jaramillo and younger parts of the type section of the 

Olyor Suite and to the fauna of locality CRH-47 in eastern 

Beringia. 

Member Ilia of the Olyor Suite is interpreted (Repenning, in press) (1) as having been deposited between 1.4 

and, perhaps, 0.90 million years ago (including deposits 

above the Cobb Mountain Subchron); (2) to record the 

Cobb Mountain Normal-polarity Subchron of the 

Matuyama Chron; and (3) to record cooling climate 

before and during the Cobb Mountain Subchron, with subsequent warming leading into the Jaramillo Subchron. 

In the Krestovka exposures, there is no record of 

renewed cooling prior to the beginning of the Ice Ages and 

there are no younger reversely polarized deposits; unconformities may be present. 

YAKUTIA: TYPE OLYOR SUITE 

Sher (1971), Giterman and others (1982), Virina and 

others (1984), and Sher (1986, 1987) have provided information about the type sections of the Olyor Suite, on the 

Bol'shaya Chukoch'ya River about 93 mi (150 km) southwest of its (present) mouth at the Arctic Ocean and about 

an equal distance north of the Krestovka sections (fig. 1 ). It 

is somewhat difficult to correlate these several reports 

because different aspects of the section are emphasized in 

each. 

Two sections were used to define the type Olyor 

Suite. The older parts of the formation are present in the 

"primary" type section (Sher, 1971) and include the 

Jaramillo Subchron (Virina and others, 1984). In the 

"referred" type section the polarity pattern is continued 

upward from the youngest reversed polarity of the 

Matuyama Chron into the lower part of the Brunhes Chron 

(Virina and others, 1984). The microtine rodent faunas 

reported from these sections (Virina and others, 1984; 

Sher, 1987) are from all parts of the section, as subdivided 

by the polarity determinations, but are all represented by 

faunal samples too small for population analysis (A.V. 

Sher, oral commun., 1990). The morphotypes present, as 

named in Virina and others (1984) and Sher (1986, 1987), 

are, however, in agreement with the assignment of paleomagnetic events. 

The last occurrence of Lasiopodomys (listed as Allophaiomys in Virina and others, 1984, and Sher, 1986, 

1987) is just below the base of the Brunhes Zone, in agreement with its time of extinction in the United States, and 

the species of Microtus listed from this horizon (Virina and 

others, 1984; Sher, 1986, 1987) are advanced and are not 

known to be older than the Jaramillo event. They are in 

agreement with the paleomagnetic assignments. 

Giterman and others (1982) presented the pollen profile of the older (primary) type section; it does not match 

the floral trends of the Krestovka sections record and may 

contain gaps. Giterman and others ( 1982) conducted no 

paleomagnetic studies, and their thickness measurements 

do not agree with those in other publications. It is not possible, from published information, to identify the position 

of the Jaramillo Subchron in their pollen sequence except 

in approximate terms. 

Giterman and others (1982) indicated a general similarity of the flora throughout the primary type section and 
2.0-{).85 MILLION YEARS AGO: THE SECOND COOLING CYCLE 25 

suggested that the spruce and pine reported (only in trace 

amounts of less than 2 percent) were not locally derived. 

The same conclusion would seem to apply to larch and tree 

alder (shrub alder, Alanaster, is present throughout). The 

environment indicated is a shrub tundra. 

Willow, grass, sedges, and arctic dock (Rumex arcticus) are the only pollen that suggest some climatic variation; they are most common in the middle of the section, in 

that part of the section that probably includes beds deposited during the normal polarity of the Jaramillo event, as 

determined by comparison of the section published by 

Virina and others ( 1984) with that by Giterman and others 

(1982). Birch and the Compositae decrease accordingly, 

and the slight change suggests a somewhat more humid climate. The lowest percentages of tree and shrub pollen are 

in this part of the type section, but the change (from as 

much as 50 percent to as little as 20 percent) is not comparable to the progressive change to full tundra recorded in 

the Olyor of the Krestovka sections. 

In contrast to the suggestion of greater humidity in 

that part of the type section that presumably represents the 

Jaramillo Subzone, Sher and others (1979) noted that 

unusually arid climate plants (Selaginella sibirica spores 

predominate) are in the coldest part of the Krestovka sections associated with the Cobb Mountain event. 

In the referred section, which typifies the younger part 

of the type Olyor, the microtine Lasiopodomys is present in 

reversed-polarity deposits of the latest Matuyama Chron at 

a point slightly below the base of the Brunhes Chron (Sher, 

1987; reported as Allophaiomys), as was just mentioned. 

This age for the last appearance of Lasiopodomys is in 

approximate agreement with its time of extinction in the 

United States (-830,000 y.b.p.) (Repenning, in press), and 

possibly elsewhere, but is not in agreement with the record 

in the east Asian Chinese faunal province where the genus 

still lives, presumably as a sister lineage. 

In the Chinese province older records of this lineage 

of Lasiopodomys are small (Kudun fauna, Buryat; perhaps 

1.7 m.y. old). About 1.2 million years ago (Gongwangling 

fauna, Haanxi, China) Lasiopodomys was considerably 

more advanced morphologically than individuals of the 

genus elsewhere in the world, and only in the Chinese 

province does the genus still live, and it does so as much 

more evolved forms than are known as fossil anywhere 

else in the Northern Hemisphere. There are many other 

aspects of the fauna of the Chinese province that suggest 

isolation. Biochronology is not always the same in discrete 

faunal provinces (Repenning, in press). 

The earliest record of Microtus in the referred type 

section of the Olyor also is in the last reversed-polarity 

deposits of the Matuyama Zone (younger that the Jaramillo 

event) (Sher, 1987), but the species named are more 

advanced than the form from locality KL0-9 of the Krestovka sections, and the horizon between the Jaramillo and 

the Brunhes Chrons is demonstrably younger than the first 

record ef the genus elsewhere, including other Beringian 

records (Repenning, in press). The first record of Microtus 

in the type Olyor is thus a false "first record" (Sher, 1986, 

fig. 1 ), possibly resulting from the very small samples 

recovered from the type sections of the Olyor Suite. 

The several published studies of the type sections of 

the Olyor Suite leave many questions unanswered. For the 

purpose of the present synthesis, it is significant to note 

only that the microtine fauna and associated paleomagnetic 

polarity probably indicate an age range beginning before 

the Jaramillo Subchron of the Matuyama Chron and 

extending into the Brunhes Chron. A climate at least as 

warm as the present-day climate is indicated during the 

Jaramillo Subchron. 

ALASKA: CAPE DECEIT SECTION 

The Cape Deceit section is on the Seward Peninsula 

of Alaska near the village of Deering (fig. 1 ). The section 

is on the north side of Bering Strait but at about at the same 

latitude. Guthrie and Matthews (1971) described the Cape 

Deceit mammalian fauna from the upper (coldest) part of 

the Cape Deceit Formation (early Pleistocene). Matthews 

(197 4) made an analysis of the environments that are represented in the formation, and this analysis was restated in 

Giterman and others (1982). Repenning and others (1987) 

discussed aspects of the fauna and its age, but their age 

assignment was revised by Repenning (in press) because its 

interpretation is dependent on the age of the Krestovka 

Olyor, which was also revised. 

Two temporally important microtine rodents in the 

Cape Deceit fauna are the primitive meadow mouse 

Lasiopodomys deceitensis, which was described as Microtus deceitensis but has been assigned to the genus Allophaiomys by some, such as Sher ( 1986), or to Lasiopodomys by 

Repenning (1980), Repenning and Grady (1988), and 

Repenning (in press), and the primitive collared lemming 

Predicrostonyx lwpkinsi. Both of these species were named 

from this fauna. They provide slightly conflicting evidence 

of the age of the fauna when correlated with the older part 

of the Krestovka Olyor (Repenning, in press). 

Lasiopodomys deceitensis (called "Allophaiomys sp. 

cf. A. pliocaenicus" in Sher and others, 1979, and Sher, 

1987) is recognized with the first Microtus in locality 

KL0-9 of the Krestovka sections, where it is slightly older 

than 1.1 million years (0.8 ft, 0.25 m below the Cobb 

Mountain Subzone of the Matuyama Zone). No Microtus 

was found in the Cape Deceit fauna, which is a large 

enough sample to include individuals of Microtus had the 

genus lived in the area. Therefore, the Cape Deceit fauna is 

believed to be slightly older than that of locality KL0-9 

(older than 1.1 m.y.). 

A species of Lasiopodomys, also called "Allophaiomys sp. cf. A. pliocaenicus" in Sher and others (1979) 

and Sher ( 1987), is also present in locality KL0-8, the 
26 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

oldest microtine fauna of the Krestovka Olyor, but no 

Microtus is present at this locality·. The species is different 

and possibly more primitive than L. deceitensis from the 

Cape Deceit fauna (Repenning, in press). Locality KL0-8 

is believed to be between 1.6 and 1.3 million years old 

(Repenning, in press) but cannot be as old as 1.67 million 

years because the normal polarity of the Olduvai Subchron 

is not present. Locality KL0-8 cannot be as young as the 

overlying locality KL0-9. If the difference noted in the 

Lasiopodomys from KL0-8 is significant, then perhaps 

locality KL0-8 is older than the Cape Deceit fauna. 

Predicrostonyx hopkinsi from the Cape Deceit fauna 

is the most primitive collared lemming known. The ancestry of the collared lemmings is unknown; the genus is 

thought to have originated in submerged parts of the Arctic 

Ocean Borderland now forming the Siberian Shelf. 

In the Krestovka Olyor Suite, Predicrostonyx hopkinsi 

is known only as reworked specimens derived from 

unknown pre-Olyor deposits; a more advanced species, P. 

compitalis, in the oldest microtine fauna of the Olyor at 

locality KL0-8 (Sher and others, 1989) suggests that Krestovka locality KL0-8 is younger than the Cape Deceit 

fauna, and Predicrostonyx hopkinsi from the Cape Deceit 

fauna suggests that the Cape Deceit fauna may be as old as 

the Olduvai Subchron, which ended 1.67 million years ago. 

Unfortunately, there are no paleomagnetic polarity determinations for the Cape Deceit fauna. 

In the youngest significant fauna of the Krestovka 

Olyor, from locality KL0-9, Predicrostonyx compitalis is 

replaced by its descendant, Dicrostonyx renidens, which is 

associated with Microtus and Lasiopodomys deceitensis at 

that locality. Dicrostonyx renidens is replaced by D. simplicior at locality CRH-47 in the Old Crow Basin of Canada 

(R.E. Morlan, written commun., 1991), which was deposited during the Jaramillo Subchron. This is additional reason to think that the first normal-polarity zone in the 

Krestovka section is older than the Jaramillo. 

The suspicion (Repenning, in press) that the 

Lasiopodomys from locality KL0-8 is more primitive than 

L. deceitensis from the Cape Deceit fauna conflicts with 

the older age suggested by Predicrostonyx hopkinsi from 

Cape Deceit. Both suggestions are weak: the supposed 

primitiveness of Lasiopodomys from locality KL0-8 is not 

supported by morphologic trends in any other fauna in the 

world and the identification of species of Predicrostonyx 

(and Dicrostonyx) is based on average population morphology, and the KL0-8 sample of the lemming is not very 

large, raising the doubt that it may not be representative of 

the population. 

As a result, the Cape Deceit fauna is also considered 

to have an age range from possibly 1.6 to 1.2 million years 

(Repenning, in press), similar to the possible age range 

assigned to locality KL0-8, and it is not certain which 

locality, if either, is older. 

If the Cape Deceit fauna is as old as the Olduvai 

Subchron, it would be the oldest record of Lasiopodomys in 

the world except for the record of L. praebrandti from the 

Lake Baikal area (Erbaeva, 1976). This record, from the 

Kudun fauna, is not dated by external evidence but on 

other faunal evidence is believed to be about the age of the 

Olduvai Subchron (Repenning and others, 1990). It is also 

believed to be a fauna of the Chinese faunal province of 

Asia, which may have a biochronology with a somewhat 

different time scale and history than those under 

discussion. 

In his analysis of the environments of the Cape Deceit 

Formation, Matthews ( 197 4, table 6) summarized six successive environments for the formation and for the periods 

both immediately before and immediately after its deposition. The environment before deposition of the formation 

was cold and the environment following deposition was 

much warmer than today. 

Within the formation, Matthews recognized four climatic units: two tundra environments, colder than today, 

and two forest or forest-tundra intervals, warmer than 

today. The oldest warm period was the warmer of the two 

recognized within the formation; at this time spruce grew 

at or near Cape Deceit. The youngest cold period was the 

colder of the two recorded in the sediments; at this time 

herbaceous tundra grew on Cape Deceit, and the deposits 

show evidence of permafrost. This was the period during 

which the majority of the small mammals of the Cape 

Deceit fauna lived. 

In reality, the lack of forest today on the Seward Peninsula (and at Cape Deceit) is a coastal feature; forests 

extend farther north farther inland. The evidence of former 

forests in the Cape Deceit Formation possibly represents 

low sea levels, with the coast more removed from the 

locality, and thus a "cold" period could be inferred from 

the presence of forests in the Cape Deceit Formation. 

The top of the Cape Deceit Formation, immediately 

above the unit containing most of the mammalian fauna, is 

truncated by an erosional unconformity believed to represent a much warmer time than the present. This warm 

period, which resulted in thawing, slumping, and erosion of 

the previously frozen sediments, is the youngest climatic 

period that Matthews (1974) associated with the Cape 

Deceit Formation. 

This history of alternating tundra and forest in the 

Cape Deceit Formation, and a climaxing peak of cold 

climate (followed by temperatures much warmer than 

today), is in agreement with the climatic history of the Arctic Ocean Borderland for several periods of time between 

about 1.8 and 1.1 million years ago. It conforms broadly 

with the nature of the relatively quiet period of glacial 
MARINE AND PROXY RECORDS OF THE SECOND COOLING CYCLE 27 

activity recorded at site 644 in the Norwegian Sea but does 

not obviously correspond to any particular part of it. No 

basis exists for refining the age of the formation by comparison with climatic events, but this climatic history does 

tend to support the general fluctuations of climate during 

the earlier part of the second cooling cycle that were 

inferred from the record of ice-rafted debris at site 644 in 

the Norwegian Sea. 

CANADA: YUKON TERRITORYLOCALITY CRH-47 

Locality CRH-47, along the Old Crow River in the 

northern part of Yukon Territory, Canada (fig. 1), is near 

the Canada-Alaska border and about 100 mi (161 km) 

south of the Arctic Ocean in easternmost Beringia. The 

locality was discovered by C.R. Harington and contains a 

very large microtine rodent fauna that is being studied by 

R.E. Morlan. Sediments at the locality have normal magnetic polarity and a pollen flora indicating boreal forest 

(Schweger, 1989); the climate during deposition was as 

warm as, or possibly somewhat warmer than, today. 

According to Morlan (written common., 1991), microtine rodents in the fauna include Lasiopodomys deceitensis, 

a variety of Microtus morphotypes, Phenacomys, Dicrostonyx renidens-Dicrostonyx simplicior (intergrading from 

bottom to top of the section), Mimomys (Cromeromys) sp. 

cf. M. (C.) virginianus, and other genera. 

As noted above, Dicrostonyx rem·dens replaced Predicrostonyx compitalis in the fauna from locality KL0-9 of 

the Krestovka sections in Y akutia, western Beringia, and 

only primitive morphotypes of Microtus were present at 

locality KL0-9. About 140,000 years were required to 

evolve the endemic Microtus paroperarius of the United 

States from a primitive morphotype comparable to the 

specimens at locality KL0-9 to a population morphologically varied, with advanced forms, comparable to those at 

locality CRH-47 (Repenning, in press). The presence of 

Microtus at this stage of evolution and of Dicrostonyx as 

more advanced species than those from KL0-9 indicates 

that the CRH-47 fauna is younger than the KL0-9 fauna. 

!he Cobb Mountain Subchron, immediately overlying 

localtty KL0-9, was perhaps only 10,000 years long (Mankinen and Gromme, 1982), and the next normal polarity 

event was the Jaramillo, which thus must be represented by 

the normal polarity during deposition of locality CRH-47; 

the time involved between the ages of localities KL0-9 

and CRH-47 (130,000-200,000 years) thus represents the 

time span of the species Dicrostonyx renidens and the 

evolution of varied and advanced morphotypes of Microtus 

in Beringia. This time span is comparable to that observed 

in the similar evolution of Microtus paroperarius in the 

United States (noted above) (Repenning, in press). 

That locality CRH-47 in Yukon Territory cannot be 

as young as the next normal polarity (Brunhes Chron) is 

indicated by the presence of Lasiopodomys deceitensis and 

Mimomys (Cromeromys) sp. cf. M. (C.) virginianus in that 

fauna (R.E. Morlan, written common., 1991). In the type 

locality of the Olyor Suite, western Beringia (Y akutia), 

Lasiopodomys (as Allophaiomys; Sher, 1987) became 

extinct before the Brunhes Chron, and the same date of 

extinction is recognized in the conterminous United States 

(Repenning, in press); lacking evidence to the contrary, it 

is reasonable to assume the same extinction date in eastern 

Beringia of North America. M. (C.) virginianus from the 

Cheetah Room fauna, West Virginia (Repenning and 

Grady, 1988), is older than the Brunhes Chron and is similar toM. (C.) dakotaensis from the much older (-1.5 million year old) Java fauna of South Dakota (Martin, 1989). 

No similar morphotype of Mimomys (Cromeromys) is 

known anywhere in the world during the Brunhes Chron, 

although the dissimilar but related M. (C.) savini (=M. (C.) 

intermedius) is. Finally, Dicrostonyx simplicior is known 

in Europe before the beginning of the Brunhes Chron, and 

it is not reasonable that its evolution out of D. renidens, 

recorded in locality CRH-47, could be an event of the 

Brunhes Chron. Thus the normal paleomagnetic polarity of 

locality CRH-47 in the Old Crow Basin was assigned to 

the Jaramillo Subchron by Repenning (in press). 

C.E. Schweger (1989) examined the pollen of locality 

CRH-47 and concluded that it indicates an environment as 

warm as today or possibly somewhat warmer. This assessment is in agreement with the floral record during the 

Jaramillo Subchron as described from the type section of 

the Olyor Suite (a low percentage of arboreal pollen in an 

area now occupied by tundra) and with the warm Bavelian 

floral of Holland as described by Zagwijn and de Jong 

(1984), also of Jaramillo age. This warm period is not evident in the cores from site 644 in the Norwegian Sea or 

CESAR site 14 on the eastern Alpha Ridge but neither is 

such a warm period evident during the Brunhes Chron. The 

CRH-47 fauna is older than the 840,000-year-old Cheetah 

Room fauna of West Virginia based on the history of 

Microtus and on the normal magnetic polarity of locality 

CRH-47 (Repenning and Grady, 1988; Repenning, in 

press); however, locality CRH-47 may not be inuch older. 

MARINE AND PROXY RECORDS OF 

THE SECOND COOLING CYCLE 

Similar to the last part of the first cooling cycle, there 

are few records in the Arctic Ocean Borderland that give 

evidence of climatic change during the second cooling 

cycle. The Kap K~benhavn Formation in northern Greenland provides evidence of treeline at lat 82°30' N. 2 million 

years ago at the beginning of the second cycle; this 
28 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOOIC CHANGES, ARCTIC OCEAN BORDERLAND 

position is supported by the marine record of little glacial 

activity in the Norwegian Sea and by other records that 

have been discussed. This was the Tiglian warm period as 

recognized in Holland. 

The remainder of the second cooling cycle was reconstructed by the intercorrelation of the few records from the 

Arctic Ocean Borderland that have been discussed and 

lower latitude proxy evidence. 

NORWEGIAN SEA: 

SCANDINAVIAN ICE SHEET 

Between 2.0 and about 1.5 million years ago there is 

little evidence of climatic change in the Arctic except that 

from site 644 in the Norwegian Sea. The part of this record 

that corresponds to the warm Tiglian flora of Holland indicates minor peaks in glacial activity at the beginning (1.90 

m.y. ago) and near the end (-1.7 m.y. ago) of the Olduvai 

Subchron (fig. 2). The first of these peaks is coincident 

with lower latitude proxy evidence of the dispersal of the 

distinctly temperate to subtropical microtine rodent Allophaiomys (1) from Beringia south to the United States at 

the beginning of the Olduvai Subchron (possibly somewhat 

more than 1.9 m.y. ago; Repenning, in press) and (2) from 

the Black Sea area northward into Holland during the earliest part of the Olduvai (van Kolfschoten, 1990a; somewhat 

less than 1.9 m.y. ago). The latter dispersal would be during the last half of the Tiglian warm floral age (Zagwijn 

and de Jong, 1984) and immediately following the minor 

peak in glacial activity indicated at the beginning of the 

Olduvai Subchron in the core from site 644 (fig. 2). 

During the Olduvai the average ice activity was low 

(mostly background) in the record of the Norwegian Sea; 

this period coincides, on the basis of paleomagnetic correlation, with the last half of the warm Tiglian flora age of 

Holland (Zagwijn and de Jong, 1984). Again on the basis 

of paleomagnetic correlation, the period of minor, but 

increased, glacial activity late during the Olduvai Subchron 

(beginning -1.7 m.y. ago) correlates with the beginning of 

the Eburonian flora of Holland and with the dispersal of 

Phenacomys southward into the United States. Both the 

appearance of Eburonian colder floras and the increase in 

ice rafting of debris occurred late in the period of time represented by the Olduvai Normal-polarity Subchron. 

The end of the Eburonian cold flora and the beginning 

and ending of the following Waalian warm and Menapian 

cold floras of Holland (Zagwijn and de Jong, 1984) all are 

within the 700,000-year-long reversed-polarity part of the 

mid-Matuyama Chron, which is a period of reversed polarity almost as long as the normal polarity representing the 

Brunhes Chron. Except for the very brief Cobb Mountain 

Subchron, the lack of polarity reversals during this 

700,000-year period provides no basis for approximate age 

interpretation between the Olduvai and Jaramillo Subchrons. The only posssible solution for an approximate age 

interpretation of these floral ages in northwestern Europe is 

in the correlation with oceanic records, where deposition 

rates were presumably more uniform and interpolation 

between age control points (polarity reversals) may be used 

with greater confidence than in continental deposits. 

This correlation of these floral ages with the record of 

Scandinavian ice sheet activity suggests that (1) the Eburonian cold floral age terminated about 1.58 million years 

ago; (2) the Waalian warm floral age ended about 1.40 million years ago; and (3) the Menapian cold floral age ended 

about 1.10 million years ago and the Bavelian warm floral 

age began shortly before the Jaramillo Subchron (at the 

Cobb Mountain Subchron, which marks the last cold 

period before the Jaramillo Subchron) (fig. 2). 

To judge from the record of ice-rafted debris at site 

644, Scandinavian ice sheet activity less than 1.2 million 

years ago was more continuous and is characterized by a 

lack of prolonged minima in debris accumulation. The 

greatest amounts of ice-rafted debris were about 400,000 

years ago, possibly correlating with the establishment of 

permanent perennial ice cover on the Arctic Ocean as 

inferred in CESAR core 14 (Scott and others, 1989), 

although Scott and others suggested a slightly later date. 

MICROTINE IMMIGRANTS: 

WESTERN UNITED STATES 

The distribution of microtine rodents is strongly constrained by climate; for the most part they are temperate to 

arctic animals. For 15 years it has been widely recognized 

that new appearances of microtines in various areas of the 

Northern Hemisphere coincide with colder climates, which 

allow more northerly forms to spread southward (Fejfar, 

1976). Until quite recently, however, it has not been 

generally recognized that some microtine taxa more 

adapted to temperate climates spread northward during 

warmer climates. Because generally cooling climatic patterns are discussed herein, the examples of northward dispersal during warmer periods are few (Repenning, 1990, in 

press) in the present discussion. These few, however, are 

informative and provide support for climatic age assignments in the Arctic Ocean Borderland. The following is a 

review of this evidence, of both cooling and warming periods, more fully discussed in Repenning (1987, 1990, in 

press), Repenning and others (1987), Repenning and Grady 

( 1988), and Repenning and others ( 1990). 
MICROTINE IMMIGRANTS: WESTERN UNITED STATES 29 

FIRST COOLING CYCLE 

The primitive bog lemming, Plioctomys, is first 

known about 4 million years ago in southern Russia and 

later in Hungary, Poland, and Mongolia. It is abundant in 

the upper Kutuyakh Beds of the Krestovka sections, where 

it is slightly younger than the Gauss-Matuyama polarity 

boundary; it is also known at this time from northern 

Alaska. Plioctomys dispersed southward from Beringia to 

the United States (1) just before the end of the Gauss 

Chron; (2) at the time that activity of the Scandinavian ice 

sheet began to increase; (3) with the density stratification 

of the Arctic Ocean; (4) when boreal forests were thinning; 

(5) as treeline was moving south from the shores of the 

Arctic Ocean; and (6) before full tundra and permafrost 

were established in Beringia. The genus persisted in the 

United States (as the more advanced species Plioctomys 

rinke ri) until sometime before 2.1 million years ago and 

then became extinct (Repenning and Grady, 1988). It was 

replaced, in the North American fauna, by the most primitive morphotype of the derived genus Mictomys. Mictomys, 

the living northern bog lemming, lives in boreal forests but 

avoids tundra. 

The two genera Plioctomys and Mictomys are first 

known in the United States from deposits about 15ft (-4.5 

m) below the Gauss-Matuyama polarity boundary in Texas 

(Plioctomys) and in Arizona and Idaho (Mictomys). The 

significance of the dispersal of the bog lemmings is not so 

much that it corroborates the timing of these events in the 

Arctic, because these events are rather well dated, but that 

it provides an idea of the effect of climatic deterioration on 

climates in lower latitudes of North America. 

Plioctomys became extinct before 2.1 million years 

ago (there are no records this young), following the dispersal of Mictomys around the southern end of the Rocky 

Mountains to the Great Plains. The earliest record of Mictomys in the Great Plains is in Kansas 2.0 million years ago 

and 100,000 years after the deposition of glacial till in 

Iowa; there are no known associations of early Mictomys 

with Plioctomys. The range of Mictomys has subsequently 

withdrawn northward with the Canadian forests, but, both 

east and west of the Rocky Mountains, much of this retreat 

has been after the end of the Pleistocene; during the Ice 

Ages it returned southward with several, perhaps all, ice 

advances, as presumably did the forests. 

A lineage called Synaptomys diverged from Mictomys 

in the Appalachian Mountains about 840,000 years ago 

(Repenning and Grady, 1988) and now lives in the northeastern United States and adjacent southeastern Canada; 

relict populations, surviving from a late Pleistocene distribution, live as far southwest as Kansas. 

Other than Plioctomys and Mictomys, no microtines 

are known to have dispersed southward in North America 

during the frrst cooling cycle. 

BEGINNING OF THE 

SECOND COOLING CYCLE 

The extinct microtine genus Allophaiomys had a distinct preference for temperate climates. From perhaps 2.2 

million years ago (Beijing, lat 40° N .) to 0.83 million years 

ago (Maryland, lat 40° N.) it was abundantly present 

throughout the lower latitudes of the Northern Hemisphere 

(Repenning, in press). Its known latitudinal range was 50°-

300 N. in North America and 52°-31 oN. in Europe and Asia 

Minor. Allophaiomys may have originated in China (perhaps 

2.2 m.y. ago) and dispersed to both Europe and North 

America, where its first record is about 1.9 million years old 

in both Holland (van Kolfschoten, 1990a) and southern Colorado (Repenning, in press). There are no records of Allophaiomys in the Arctic, although the genus must have arrived 

in North America by way of the Bering Land Bridge (about 

lat 63°-69° N.). One must assume thatAllophaiomys crossed 

the Bering Land Bridge during a period of temperate climate 

but was forced southward in North America by the return of 

cold climate. 

As discussed above, 2.1 million years ago a continental 

glacier flowed down the Mississippi River Valley as far 

south as Iowa, ice cover was present on the central Arctic 

Ocean, and ice-rafted debris from the Scandinavian ice sheet 

was as great as it was later during most of the Ice Ages (fig. 

2); this was hardly a time of temperate climate on the Bering 

Land Bridge. 

In North America Allophaiomys is first known in Colorado, where it was recovered from rocks recording the 1.9-

million-year-old beginning of the Olduvai Subchron or from 

just beneath them (Repenning, in press). One must assume 

that between 2.1 (cold in Beringia) and 1.9 million years ago 

Allophaiomys crossed the Bering Land Bridge and was 

forced southward in North America during the brief cold 

period at the beginning of the Olduvai event that is indicated 

by the ice-rafted debris record of the Norwegian Sea (site 

644) and by the isotope records of the World Ocean (site 

607, fig. 2). If both assumptions are reasonable, one must 

then assume that Beringia had a temperate climate between 

2.1 and 1.9 million years ago. 

The coincidence of this constraint on the age of a temperate climate on the Bering Land Bridge with the age, 

inferred earlier, of the treeline forest in member B3 of the 

Kap K~benhavn Formation of northernmost Greenland is 

mutually supportive of these interpretations. Thus the 1.9-

million-year date of the oldest Allophaiomys from North 

America substantiates the age assignment that has been 

inferred for member B3 of the Kap K~benhavn Formation, 

even though Allophaiomys is not known from Greenland or 

Beringia. It also indicates that the warm climate recorded in 

the Kap K~benhavn Formation was not, as might be supposed, a local climatic extreme caused by conditions in the 
30 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

North Atlantic Ocean but was also evident in the Bering 

Strait region of the North Pacific Ocean and thus may have 

been present throughout the Arctic Ocean basin. 

This extreme warm period between 2.1 and 1.9 million 

years ago that is recorded in north Greenland and that probably existed in Beringia marked the end of the first cooling 

cycle and the beginning of the second. The rather widespread 

ability to correlate this time in the contiguous United States 

is due to the fortuitous eruption of the Huckleberry Ridge 

Tuff in the Yell ow stone National Park area of Wyoming and 

Idaho (Christiansen and Blank, 1972) 2.01 million years ago, 

to the spread of its ash in still-visible deposits across much 

of the United States from the Pacific Ocean to the Mississippi River (lzett and Wilcox, 1982), and to the coincidence 

of this eruption with the younger normal-polarity event of 

the Reunion Subchron. 

Perhaps 500,000 years later than its oldest record in the 

United States, there is no record of Allophaiomys in Beringia 

as indicated by the oldest fauna of the Krestovka Olyor Suite 

(locality KL0-8; possibly 1.4 m.y. old; Repenning, in press) 

nor is the genus known in younger Olyor faunas, in the Cape 

Deceit fauna, or the Old Crow Basin during the Jaramillo 

Subchron (0.97-0.90 m.y. ago). 

In the United States Allophaiomys is not known west of 

the Rocky Mountains; for some unknown reason the dispersal route west of the Rocky Mountains was not open at 

this time. Yet before the end of the Olduvai Subchron the 

genus Phenacomys immigrated, west of the Rocky Mountains, from Beringia to Idaho; failure of Allophaiomys to 

accompany Phenacomys southward at this time suggests that 

Allophaiomys may have become extinct in the Arctic Ocean 

Borderland before the end of the Olduvai Subchron (Repenning, in press). This hypothesis is consistent with the inference that Allophaiomys was present in Beringia only during 

a very brief and exceptionally warm period, apparently about 

2.0 million years ago. 

During most of the Matuyama Chron Allophaiomys 

occupied the territory in North America from southern 

Saskatchewan to southern Texas and from the Rocky 

Mountains to the Atlantic coast. For more than a million 

years the genus thrived in the United States but it probably 

was able to endure Beringian climates for less than 100,000 

years. One must assume that the temperate climate of 

Beringia between 2.0 and 1.9 million years ago was most 

unusual and, by inference, that the warmth 500 mi (805 

km) from the North Pole at Kap K~benhavn was equally 

unique. 

Allophaiomys, possibly older than the Olduvai Subchron, is also known along the northern shore of the Black 

Sea and in southern Europe (Alexy Tesakov, Moscow, oral 

commun., 1990). However, it was not until after the 

beginning of the Olduvai Subchron that the genus first 

appeared in Holland (van Kolfschoten, 1990a). In the same 

record (Zuurland Boreing; van Kolfschoten, 1990a) this 

warm climate genus possibly became extirpated from Holland with the beginning of the Eburonian, late during the 

Olduvai Subzone. The history of the distribution of Allophaiomys not only supports the age assignment of the 

forest-tundra in the Kap K~benhavn Formation of North 

Greenland, and provides some insight into the probable climate in Beringia during an extremely warm 200,000 years 

that started the second cooling cycle of the Arctic Ocean 

Borderland, but it also suggests that this warming may 

have been influential as far away as northern Europe. 

It is universally agreed that the meadow mouse Microtus is derived from Allophaiomys. During this evolution, 

Microtus developed arctic, as well as temperate tolerances 

(but lost subtropical tolerances), and it now lives throughout the Arctic Ocean Borderland and southward (in highlands) into Mexico and to the Mediterranean coastal area of 

North Africa. Microtus does, however, have climatic limits 

(Repenning, 1990). Although it was present in Beringia 1.1 

million years ago (in the Krestovka Olyor locality KL0-9), 

was morphologically varied during the Jaramillo Subchron 

(0.97-0.90 m.y. ago), and was present in southern Canada 

after the Jaramillo Subchron (Repenning, in press), Microtus did not enter the United States (east of the Rocky 

Mountains) until 450,000 years ago (possibly as late as the 

development of permanent perennial ice on the Arctic 

Ocean according to Scott and others, 1989). The absence of 

Microtus in the United States during the second cooling 

cycle suggests a subtropical climate and (or) excessive 

rainfall (Repenning, 1990), an indication of the climate 

preferred by Allophaiomys, which was present in the 

United States during much of this time. 

The replacement of temperate to subtropical Allophaiomys by temperate to arctic Microtus happened many times 

in various parts of the Northern Hemisphere. Because the 

derivation of Microtus from Allophaiomys is widely 

accepted, claims have been made of the origin of Microtus 

in France, Ukraine, Western Siberia, Buryat, Yakutia, 

China, and the United States. Most claims result from the 

ecologic replacement of Allophaiomys by Microtus during 

southward climatic shifts. The place of origin (or origins) 

of Microtus has not yet been identified, with the exception 

of the species Microtus paroperarius, which evolved 

endemically in the United States and never left that country 

(Repenning, in press). 

EBURONIAN COLD 

The warm spell at the beginning of the second cooling cycle began about 2.0 million years ago and lasted, 

with minor interruption, until peaks of ice-rafted debris 

from the Scandinavian ice sheet again appeared in abundance at site 644 in the Norwegian Sea, shortly before the 

end of the Olduvai Normal-polarity Subchron. This same 
MICROTINE IMMIGRANTS: WESTERN UNITED STATES 31 

time constraint marks the beginning of the cold-climate 

Eburonian flora of Holland. Other than these two records, 

there is no further indication of the beginning of this cold 

event, although the southward dispersal of the heather vole 

Phenacomys to the United States is contemporaneous and 

suggestive. There are no deposits of this age in the Arctic 

Ocean Borderland, and there are no obvious signals suggesting the beginning of the Eburonian in the cores of the 

Arctic Ocean. The Eburonian return of colder climate 

(recorded in Holland) may have been a North Atlantic 

feature caused by a temporary collapse of the strong thermal gradient between the Norwegian Sea and the North 

Atlantic. Jansen and others (1988) felt that this gradient 

existed until 1.2 million years ago. Microtine dispersal 

provides one further clue. 

The heather vole, Phenacomys, is first known about 

2.4 million years ago in Beringia; the genus is of Beringian 

origin. Only one tooth representing it has ever been 

reported from the Old World, and this record is from eastern Beringia in the Kolyma Lowlands of Y akutia (Sher and 

others, 1979, p. 38). 

In the contiguous United States Phenacomys first 

appears in the Froman Ferry fauna of western Idaho 

(unpublished data, Repenning). The associated fauna is 

otherwise completely characteristic of the older Grand 

View fauna of that area (Blancan V age); the presence of 

immigrant Phenacomys is considered as indicative of the 

next younger microtine age, Irvingtonian I. The species of 

Phenacomys from Froman Ferry is P. gryci, named from 

the Fish Creek fauna of the Alaskan North Slope. Preliminary paleomagnetic study of the Froman Ferry faunal 

sequence (J.T. Sankey, written commun. 1990) indicates 

that Phenacomys first appears in the section shortly below 

the top of the Olduvai Subzone. The age indicated would 

approximate that of the beginning of the Eburonian cold 

period and the increase in Scandinavian ice activity. 

The immigration of Phenacomys to Idaho would have 

been along the coastal route from Alaska, west of the 

Rocky Mountains. East of the Rocky Mountains the oldest 

record of Phenacomys is in the Java fauna of South Dakota 

(Martin, 1989; listed as "Hibbardomys"). This fauna also 

has a high percentage of Blancan holdovers and is considered a very old Irvingtonian fauna (Martin, 1989). 

The fossil reptiles and amphibians of the Java fauna 

indicate greater rainfall than today and considerably 

warmer winter air temperatures (Holman, 1977). The presence of the cotton rat, Sigmodon, in the Java fauna indicates that winter freezes were uncommon (or lacking), and 

a subtropical climate similar to modern St. Louis, Missouri 

(or farther south), rather than modern South Dakota, is 

indicated. Subtropical Allophaiomys is also present in the 

Java fauna, but temperate to arctic Microtus is not (Martin, 

1989). 

This climatic pattern in South Dakota does not suggest 

a correlation with the Eburonian cold flora of Holland, or 

with the Froman Ferry fauna of Idaho, but instead a correlation with the preceding Tiglian warm flora or the following Waalian warm flora. The paleomagnetic polarity of the 

Java locality is unknown, but the fauna represents a wet 

and warm environment; a climate brought into this area by 

a northward shift of the North Atlantic Bermuda High 

(Repenning, 1990) that warmed the regions of Holland and 

South Dakota is indicated. 

If Phenacomys entered the United States on the east 

side of the Rocky Mountains at the same time that it 

entered on the west, it would be with a cooler winter, hotter 

summer, and drier climate more like today; the Bermuda 

High would have been in a southern position. Therefore, it 

is unlikely that the Java fauna could represent the time of 

the end of the Tiglian flora and the Froman Ferry fauna 

The Java fauna may postdate the arrival of Phenacomys 

east of the Rocky Mountains and correlate with the 

Waalian warm flora of Holland (possibly 1.58-1.40 m.y. 

ago); early Irvingtonian I faunas of the United States as 

young as 1.40 million years contain a high percentage of 

Blancan V taxa. Present evidence does not exclude, however, the possibility that, on the east side of the Rocky 

Mountains, Phenacomys entered the United States with 

Allophaiomys 1.9 million years ago, before the genus followed the route down the Pacific coast; in which case the 

warm Java fauna would correlate with the Tiglian warm 

period. 

BEGINNING OF THE ICE AGES 

At the beginning of the Ice Ages, the end of the second cooling cycle as here used, the microtine genera 

Lasiopodomys and Terrico/a first entered the United States 

from Beringia; this was about 850,000 years ago 

(Repenning, in press). Lasiopodomys is known for perhaps 

650,000 years before this time in Beringia, but Terricola 

was (and is) a southern form in Eurasia and, similar to 

A/lophaiomys, has no record in Beringia. It also presumably passed across the Bering Land Bridge quickly and 

then retreated southward in both Asia and North America 

with the next cold period. The genus still is a temperate 

form in the living fauna of Eurasia, but in North America it 

may survive only in Mexico as Terricola quasiater 

(Repenning, in press). Terrico/a meadensis, the immigrant 

species, survived in the United States at least until the 

Porcupine Cave fauna in south-central Colorado (Barnosky 

and Rasmussen, 1988), which, because it contains both the 

endemic Microtus paroperarius and the immigrant Microtus sp. cf. M. montanus, is believed to be the age of the last 

immigration from Beringia and about 400,000 years old. 
32 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

Endemic Microtus paroperarius apparently became 

extinct with the immigration of Beringian Microtus, and 

the climatic implications of the immigration of Beringian 

Microtus suggest that Terricola shortly moved southward 

to Mexico; thus the association of these taxa in the Porcupine Cave fauna suggests that the fauna preserves a relatively brief time of north-south faunal association. Unlike 

the record of climate assumed to have existed during the 

passage of Allophaiomys, there is no proxy evidence in the 

Arctic Ocean Borderland suggesting a warm period that 

might have permitted the Terricola passage across the Bering Land Bridge. 

It is not difficult to correlate the immigration event of 

Lasiopodomys and Terricola around the Northern Hemisphere because it was during a marked climatic change that 

permitted the fauna of the Arctic Ocean Borderland to 

move southward in both Old and New Worlds; Azzaroli 

( 1983) called it the "End-Villafranchian Dispersal Event" 

in Eurasia and summarized the faunal change there. Repenning (in press) summarized the age control of microtine 

dispersal during this event in North America; the dispersal 

was presumed to correlate with oxygen isotope stage 22 

(about 850,000 years ago) and with the first of the 

"Nebraskan" ice sheets. Dated microtine faunas in the 

United States constrain the time of dispersal to well after 

900,000 years ago and well before 820,000 years ago. If 

large mammals in North America were associated with this 

dispersal event, as they were in Eurasia, their time of dispersal has not yet been identified with sufficient certainty 

to suggest an age. 

During the Ice Ages there was one additional southward dispersal known in North America; Europe apparently experienced several (Repenning and others, 1990). 

The North American immigration was believed to have 

occurred about 450,000 years ago on the basis of the age of 

isotope stage 12 (Repenning, 1984), but evidence from the 

CESAR-14 site suggests that the immigration may have 

been during isotope stage 8. Both interpretations are based 

on the assumption that the dispersal occurred during an 

extreme cold period and not on any direct evidence of the 

faunal age. This is the dispersal that is recorded in Porcupine Cave, Colorado, which is not otherwise dated. 

SUMMARY: LATE PLIOCENE-EARLY 

PLEISTOCENE ECOLOGIC CHANGE 

About three and a half million years ago the average 

ratio of oxygen isotopes indicates that the volume of global 

ice stored on land (inverse=isostatic sea level) began to 

increase. Ice storage was greatest in the polar regions, but 

the records that indicate the development of ice storage in 

the Arctic region during the two cooling periods do not 

closely match the global average of isotope changes. Thus, 

differences in proportionate ice storage in the Arctic, in 

mountain glaciers at lower latitudes, and in the Antarctic 

are indicated. 

From 3.0 to 2.4 million years ago, evidence from both 

northeastern Siberia and northern Alaska and Canada indicates that the Arctic Ocean Borderland experienced a 

southward shift of treeline, changing from a coastline occupied by mixed coniferous forest, to treeline forest-tundra, 

and then to extensive tundra and the first record of permafrost. Treeline moved southward from almost lat 80° N. 

about 3 million years ago to less than lat 70° N. by 2.4 million years ago. 

The records of marine ostracodes across this wide latitude zone indicate that about 3.0 million years ago the 

adjacent sea was cold-temperate to subfrigid, similar to sea 

temperatures adjacent to modern Labrador and Nova 

Scotia. By 2.40 million years ago the marine ostracode 

fauna had changed to one indistinguishable from the modern fauna of the Arctic Ocean with a temperature tolerance 

of subfrigid to frigid; tundra developed on shore. Climatic 

indications derived from associated mollusks overlap the 

temperature indicated by ostracodes in the subfrigid range 

at lat 70° N. 

The Siberian record, in western Beringia, indicates 

that ice-wedge polygons, permafrost, and steppe-tundra 

first developed about 2.4 million years ago, contemporaneous with a similar change in northern Alaska that is 

inferred from the change in tundra type. Both invertebrate 

and vertebrate evidence in Alaska indicate that the Arctic 

Ocean was not yet frozen to the shore, even during the 

winter. 

Somewhat earlier (about 2.56 m.y. ago), Beringian 

bog lemmings dispersed southward to the conterminous 

United States, to as far south as central Arizona. At the 

same time the Arctic Ocean became density stratified and 

the number of icebergs markedly increased; the 

Scandinavian ice sheet began to enlarge, sending greater 

amounts of ice to the shores of and across the Norwegian 

Sea. 

From 2.5 to 2.1 million years ago, the climate of the 

Northern Hemisphere was dominated by the first significant continental glaciation of the late Tertiary. Records 

around the Arctic Ocean Borderland and in the central 

United States indicate that this glaciation developed to a 

maximum between 2.2 and 2.1 million years ago and then 

terminated. Direct evidence of increasing glaciation comes 

from (1) bottom sediments of the Norwegian Sea (Jansen 

and others, 1988), which indicate activity of the Scandinavian ice sheet; (2) North Atlantic site 552A (Shackleton 

and others 1984), which also indicates increases in icerafted debris; (3) the western Arctic Ocean site T3-67-12 

(Herman and Hopkins 1980; Herman and others, 1989; 

their Unit II), which indicates increased iceberg activity 

and salinity stratification; ( 4) the central Arctic Ocean site 
SUMMARY: LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGE 33 

CESAR-14 (Scott and others, 1989), which indicates that 

the central Arctic Ocean had intermittent perennial ice 

cover at this time; (5) the central United States, which indicates continental glaciation flowing southward to Iowa 1.1 

million years ago; and (6) the Kap Kfijbenhavn Formation 

in northernmost Greenland (Fonder and others, 1985), 

which records the end of this first period of ice cover. Mollusk faunas from the oldest part of the Kap Kfijbenhavn 

Formation indicate that marine temperatures were as cold 

as today, and these fauna are the oldest Arctic Ocean mollusk fauna with no extinct elements. 

For most of this first glaciation there are no known 

terrestrial deposits in the Arctic Ocean Borderland, possibly because of a lowered base level caused by lowering sea 

level. A similar lack of deposits during glacial periods is 

evident in the Atlantic coastal area of the United States 

(T.A. Ager, written common., 1992). 

The first significant late Cenozoic continental glaciation of the Northern Hemisphere, presumably as extensive 

as any later North American glaciation, reached a peak 

between 2.2 and 2.0 million years ago; the 2.1-millionyear-old first normal-polarity excursion of the Reunion 

event has been located within till in Iowa (Easterbrook and 

Boellstorff, 1984). This glaciation at the end of the first 

cooling cycle terminated abruptly, possibly within 50,000 

years; the 2.0-million-year-old second excursion of the 

Reunion event is present in interglacial deposits overlying 

the till in Iowa (Easterbrook and Boellstorff, 1984). By 2.0 

million years ago, climates much warmer than today were 

present in the Arctic Ocean Borderland, and treeline moved 

northward to within about 500 m (800 km) of the North 

Pole. The temperate to subtropical microtine genus Allophaiomys was then able to move as far north as the Bering 

Land Bridge, enter North America, and return southward to 

the United States. 

The warming of the Arctic Ocean Borderland to conditions similar to modem temperate climates at Beringian 

latitudes is also indicated by ( 1) the appearance of mixed 

conifer treeline forests on the northernmost shores of the 

Arctic Ocean at Kap Kfijbenhavn, lat 82°30' N. (Fonder and 

others, 1985); (2) the presence of much warmer water 

ostracodes (cold temperate to warm subfrigid) living at this 

locality on the northern shores of Greenland, considerably 

farther north than their latitudinal limits today (Brouwers 

and others, 1991); (3) the great reduction in activity of the 

Scandinavian ice sheet (Jansen and others, 1988); (4) the 

lack of further evidence of glacial activity in the midwestern United States until the end of the second cooling cycle 

(Easterbrook and Boellstorff, 1984); (5) the beginning of 

the warm Tiglian floras of northwestern Europe; and (6) 

the entrance of the microtine genus Allophaiomys into 

North America, which indicates temperate conditions in the 

Bering Strait region of Beringia. Subtropical sea-surface 

temperatures were indicated for this time in the vicinity of 

Cornwall, England (Jenkins and Houghton, 1987), and 

subtropical temperatures and greater rainfall are indicated, 

possibly for a somewhat later time, in north-central South 

Dakota (Holman, 1977). In Holland the warm Tiglian flora 

was present from 2.0 to about 1.7 million years ago (van 

Kolfschoten, 1990a). 

This unusually warm period marked the beginning of 

the second cooling cycle. The warmth persisted for about 

300,000 years with only one modest cold period at the 

beginning of the Olduvai Subchron (1.9 m.y. ago) when 

Allophaiomys entered the United States on the eastern side 

of the Rocky Mountains. About 1.7 million years ago ice 

activity of the Scandinavian ice sheet increased, the cold 

Eburonian flora became established in Holland, and the 

heather vole Phenacomys migrated down the western side 

of the Rocky Mountains from Beringia to Idaho in the 

United States. 

Allophaiomys persisted in the United States for more 

than a million years, until shortly after the beginning of the 

Ice Ages at the end of the second cooling cycle, about 

830,000 years ago (Repenning, in press). For at least part 

of this time, most of the United States appears to have had 

a subtropical climate. Jansen and others (1988) reported the 

presence of a steep temperature gradient between the Norwegian Sea and the North Atlantic that may have accentuated the warmth of this period and caused a northward shift 

of the subtropical North Atlantic Bermuda High, bringing 

warmth and greater rainfall to the United States east of the 

Rocky Mountains and to the Atlantic regions of Europe. In 

both Beringia and the contiguous United States, the genus 

Lasiopodomys also became extinct at this time. Phenacomys still lives, as several more advanced species, in 

Canadian forest environments and humid coastal forests of 

the Pacific Northwest of the United States. 

The return of stronger and more frequent peaks of icerafted debris from the Scandinavian ice sheet marks the 

beginning of the Eburonian cold floral age of Holland, 

which began late during the Olduvai Subchron (van 

Kolfschoten, 1990a). From the end of the Olduvai Subchron until the Jaramillo Subchron there is no paleomagnetic clue available to form a basis for correlation, and it is 

an assumption, made only from the record of core 644 and 

its temporal coincidence with those events that are dated 

elsewhere, that the pattern of variation in ice-rafted debris 

after the end of the Olduvai Subchron corresponds to the 

floral ages of Holland. With this assumption, the W aalian 

warm floras began about 1.58 million years ago (the Eburonian being shorter than generally portrayed) and the 

Menapian cold flora began about 1.19 million years ago. 

The peaks of Scandinavian ice influx decreased after 1.1 

million years ago (Cobb Mountain Subchron), and this 

time probably marks the end of the Menapian cold floras 

and the beginning of the Bavelian warm floras, which 

straddle the Jaramillo Subchron (Zagwijn and de Jong, 

1984). 
34 LATE PLIOCENE-EARLY PLEISTOCENE ECOLOGIC CHANGES, ARCTIC OCEAN BORDERLAND 

Thus, we suggest that the Cobb Mountain Subchron, 

discussed earlier in connection with the cold climate in the 

Krestovka sections, marks the end of a cold, ice-active 

period that correlates with the Menapian flora of Holland, 

although the Cobb Mountain Subchron is not recorded in 

core 644 nor in Holland. The subchron is believed to be 

present in CESAR core 14 of the central Arctic Ocean, 

where it is associated with the second development of 

intermittent perennial ice cover on the ocean (Scott and 

others, 1988). Just before the Cobb Mountain event the 

microtine genus Microtus first appeared in the coldest early 

Pleistocene environment known from the Krestovka sections of Yakutia. 

During the Jaramillo Subchron there probably was a 

slight warm spell, named the Bavelian in Holland, that is 

not clearly marked in oceanic cores 644 or CESAR-14. 

According to Zagwijn and de Jong (1984), the Bavelian 

consisted of a series of short warm and cold periods, and, 

although the ice-active periods are evident in core 644, the 

intervening warmer periods appear to have no time duration; they may have been too brief to be of recognizable 

duration in the plots of the core. Microtus first appeared in 

Holland, as advanced morphotypes, during the Bavelian; 

the genus represented a morphologically varied population 

in Yukon Territory during the Jaramillo Subchron (R.E. 

Morlan written common., 1991). 

Fifty thousand years after the end of the Jaramillo 

Subchron the first of several glacial advances that have 

been included in the "Nebraskan" glaciation, the first of the 

Ice Ages, spread across the north-central part of the contiguous United States, ending what is here called the second 

cooling cycle. 

With this first ice advance of the Ice Ages the microtine genera Lasiopodomys and Terricola entered the United 

States from Beringia. Although Lasiopodomys had been a 

resident of Beringia for 250,000 years before the first 

advance of "Nebraskan" ice, similar to the genus Allophaiomys there is no Beringian record of Terricola. Terricola is 

and has always been a temperate-climate animal of Eurasian origin; thus it is presumed that at least one of the 

brief warm periods during or shortly following the 

Jaramillo Subchron must have approached a temperate climate in Beringia (Repenning, in press). 

One further dispersal into the United States by microtine rodents has been mentioned, that of Beringian Microtus. This dispersal was believed to have occurred during 

isotope stage 12, about 450,000 years ago, on the basis of 

the climatic extreme suggested by isotope records, and 

this inference is weakly supported by age control of a 

similar dispersal in Eurasia (Repenning, 1984, 1987; 

Repenning and others, 1990); however, climatic extremes 

recorded in CESAR core 14 suggest that cold temperatures may have reached a maximum starting with isotope 

stage 8. Whether or not there is a correlation between 

these events has not been established. It might also be 

noted that current information suggests that Beringian 

Bison did not enter the United States until isotope stage 6 

(Repenning, 1987), so there may have been several immigrations during the late Ice Ages, stimulated by different 

cold periods in the Arctic Ocean Borderland. 

This correlation of micro tine biochronology, paleomagnetic signatures, and climatic trends as indicated by pollen and other records of the terrestrial flora and by marine 

fauna, as well as by the evidence of ice activity, has 

produced a chronology of events that we feel is convincing. 

There can be little question that between 3.0 and 2.0 million 

years ago the Arctic Ocean Borderland changed from a forested region to one in which forest-tundra and extensive tundra were dominant and in which permafrost first developed 

by 2.4 million years ago. During this same period the Arctic 

Ocean changed from cold-temperate to sub frigid; water temperatures became subfrigid about the time that tundra and 

permafrost developed in the latitudes of Beringia. 

The end of this first cooling period is dramatically 

marked by the return of warm climate in the Northern Hemisphere. In the Arctic Ocean Borderland the climate was 

extremely warm, warmer than any other late Tertiary or 

Pleistocene climate so far recorded. With this warming the 

second cooling period began, and it culminated with the Ice 

Ages. The climate of the second cooling period fluctuated 

more than that of the first, with a return to warmer climates 

about every 100,000-300,000 years. Including the colder 

fluctuations, the warm period persisted, intermittently, for 

about a million years during most of the second cooling 

cycle. And the warmth that it began with was extreme in the 

conterminous United States, as well as in the Arctic Ocean 

Borderland. 

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Circulars present administrative information or important scientific information of wide popular interest in a format designed for distribution at no cost to the public. Information is usually of short-term interest. 

Water-Resources Investigations Reports are papers of an interpretive nature made available to the public outside the formal USGS publications series. Copies are reproduced on request unlike formal USGS 

publications, and they are also available for public inspection at depositories indicated in USGS catalogs. 

Open-File Reports include unpublished manuscript reports, maps, 

and other material that are made available for public consultation at depositories. They are a nonpermanent form of publication that maybe cited in other publications as sources of information. 

Maps 

Geologic Quadrangle Maps are multicolor geologic maps on topographic bases in 7 1/2- or 15-minute quadrangle formats (scales mainly 1 :24,000 or 1 :62,500) showing bedrock, surficial, or engineering 

geology. Maps generally include brief texts; some maps include structure 

and columnar sections only. 

Geophysical Investigations Maps are on topographic or planimetric bases at various scales, they show results of surveys using geophysical techniques, such as gravity, magnetic, seismic, or radioactivity, which 

reflect subsurface structures that are of economic or geologic significance. Many maps include correlations with the geology. 

Miscellaneous Investigations Series Maps are on planimetric or 

topographic bases of regular and irregular areas at various scales; they 

present a wide variety of format and subject matter. The series also includes 7 1/2-minute quadrangle photogeologic maps on planimetric 

bases which show geology as interpreted from aerial photographs. The 

series also includes maps of Mars and the Moon. 

Coal Investigations Maps are geologic maps on topographic or 

planimetric bases at various scales showing bedrock or surficial geology, 

stratigraphy, and structural relations in certain coal-resource areas. 

Oil and Gas Investigations Charts show stratigraphic information for certain oil and gas fields and other areas having petroleum potential. 

Miscellaneous Field Studies Maps are multicolor or black-andwhite maps on topographic or planimetric bases on quadrangle or ~rregular areas at various scales. Pre-1971 maps show bedrock geology m relation to specific mining or mineral-deposit problems; post-1971 maps 

are primarily black-and-white maps on various subjects such as environmental studies or wilderness mineral investigations. 

Hydrologic Investigations Atlases are multicolored or black-andwhite maps on topographic or planimetric bases presenting a wide range 

of geohydrologic data of both regular and irregular areas; the principal 

scale is 1:24,000, and regional studies are at 1:250,000 scale or smaller. 

Catalogs 

Permanent catalogs, as well as some others, giving comprehensive 

listings of U.S. Geological Survey publications are available under the 

conditions indicated below from USGS Map Distribution, Box 25286, 

Building 810, Denver Federal Center, Denver, CO 80225. (See latest 

Price and Availability List.) 

"Publications of the Geological Survey, 1879-1961" may be purchased by mail and over the counter in paperback book form and as a set 

microfiche. 

"Publications of the Geological Survey, 1962-1970" may be purchased by mail and over the counter in paperback book form and as a set 

of microfiche. 

"Publications of the U.S. Geological Survey, 1971-1981" may be 

purchased by mail and over the counter in paperback b~ok form (two 

volumes, publications listing and index) and as a set of mtcrofiche. 

Supplements for 1982, 1983, 1984, 1985, 1986, and for sub~­

quent years since the last permanent catalog may be purchased by matl 

and over the counter in paperback book form. 

State catalogs, "List of U.S. Geological Survey Geologic and Water-Supply Reports and Maps For (State)," may be purchased by mail and 

over the counter in paperback booklet form only. 

"Price and Availability List of U.S. Geological Survey Publications," issued annually, is available free of charge in paperback booklet 

form only. 

Selected copies of a monthly catalog "New Publications of the 

U.S. Geological Survey" is available free of charge by mail or may be 

obtained over the counter in paperback booklet form only. Those wishing 

a free subscription to the monthly catalog "New Publications of the U.S. 

Geological Survey" should write to the U.S. Geological Survey, 582 National Center, Reston, VA 22092. 

Note.-Prices of Government publications listed in older catalogs, 

announcements, and publications may be incorrect. Therefore, the prices 

charged may differ from the prices in catalogs, announcements, and publications.