The presence of the Brunhes-Matuyama magnetic reversal in deep-sea core sediments makes possible an alternative to the usual K/Ar radioistope method of dating the reversal as found in rocks. The alternative method uses correlations of Northern Hemisphere summer insolation with oxygen-isotope ratios from tropical cores. The latitude-dependent insolation variations are calculated from planetary mechanics and thus provide a highly accurate astronomical time scale. The insolation variations strongly influence glacial-ice volume fluctuations that dominate the oxygenisotope ratio changes recorded in core sediments. The summer half-year insolation variations are identified with corresponding isotope-ratio changes in cores from the present through glacial Stage 20. Misleading effects of discontinuities or major nonuniformities of sediment deposition are avoided by an analysis of the uniformity of V28-238 and V28-239, the principal cores studied, and by comparisons with other cores. The top section (Stages 1 to 10) of V28-238 is uniformly deposited, and for this section an isotope-ratio time scale is chosen that agrees with the thorium-uranium date for the high sea stand of the last interglacial extreme. Over this interval, major glacial extremes (strong isotope-ratio minima) coincide consistently with major insolation minima at times of low orbital eccentricity. In addition interstadials are directly associated with precessional insolation peaks, and the envelope of isotope-ratio peaks resembles the envelope of precessionally dominated insolation peaks. The assumption that the glacial extremes depended similarly on insolation minima during Stages 10 to 20 permits minor age shifts of strong isotope-ratio minima in the two cores (relative to ages based on uniform overal deposition) to match the ages of low-eccentricity insolation minima. The age shifts reflect residual nonuniformities of deposition. The validity of this matching procedure is supported by a resulting consistent identification of principal isotope-ratio peaks with high- and low-latitude coincident insolation maxima. The Brunhes-Matuyama reversal is found intinterglacial Stage 19, and is dated on the astronomical time scale at 790,000 ± 5000 yr B.P.

L. V. Worthington (1968, Meteorological Monographs 8, 63–67) hypothesized that a low-salinity lid covered the entire world ocean. By deconvolving isotopic curves from the western equatorial Pacific and equatorial Atlantic, W. H. Berger, R. F. Johnson, and J. S. Killingley (1977), Nature (London) 269, 661–663) and W. H. Berger (1978, Deep-Sea Research 25, 473–480) reconstructed “meltwater spikes” similar to those actually observed in the Gulf of Mexico and thus apparently confirmed the Worthington hypothesis. It is shown that this conclusion is unwarranted. The primary flaw in the reconstructed meltwater spikes is that the mixing intensity used in the deconvolution operation is overestimated. As a result, structure recorded in the mixed isotopic record becomes exaggerated in the attempt to restore the original unmixed record. This structure can be attributed to variable ice-volume decay during deglaciation, effects of differential solution on planktonic foraminifera, temporal changes in abundance of the foraminifera carrying the isotopic signal, and analytical error. An alternative geographic view to the global low-salinity lid is offered: a map showing portions of the ocean potentially affected by increased deglacial meltwater at middle and high latitudes and by increased precipitation-induced runoff at low and middle latitudes.

Approximately 8000 km of continuous seismic reflection profiles throughout Lake Superior were examined for evidence of recessional moraines and other ice-margin deposits associated with the retreat of late Wisconsin ice. These features are correlated with the record of glacial-lake evolution in western Lake Superior. An offlapping sequence of glacial and glacial-lacustrine sediments overlying bedrock is recognized in west-central Lake Superior that is progressively younger to the northeast. The sequence underlies more recent glaical-lacustrine and postglacial sediments. Four facies are recognized on the basis of geomorphologic and acoustic properties and are interpreted to represent a southwest-to-northeast assemblage of: proglacial stratified drift (facies A), drift in major end moraines (facies B), till deposited as glacial retreat resumed, or possibly late-stage ablation till (facies C), and basal till (facies D). The prominent moraines of facies B are unusually thick and are believed to mark the ice-margin shorelines of successive major proglacial lakes that formerly occupied parts of western Lake Superior. The moraines are tentatively correlated with Glacial Lake Duluth (unit 1), Glacial Lake Washburn (unit 2), and Glacial Lake Beaver Bay (unit 3), the most prominent of lakes drained via the progressively lower outlets via the Moose Lake/ Brule-St. Croix Rivers, the Huron Mountains, and the Au Train-Whitefish regions, respectively.

The upper Enchantment Lakes basin in the North Cascade Range of Washington displays two moraine belts, each recording an episode of glacier advance after the end of the last glaciation. The inner belt, the Brynhild, 0.1 to 0.5 km beyond existing glaciers, postdates Mount St. Helens Wn tephra (∼450 yr old), which lies only beyond the moraines. The morainal surface is only slightly weathered, is almost barren of lichens, and is devoid of soil, evidence suggesting that the Brynhild moraines are no more than a century old. The outer moraine, the Brisingamen, 0.3 to 0.7 km beyond existing glaciers, is weathered and is covered with large lichens. On and behind the Brisingamen moraine the Mazama ash (6900 yr old) is present beneath the Mount St. Helens Yn and Wn tephras. Despite more than 7 millennia of weathering, the rock surface behind the Brisingamen moraine is measurably less weathered than the surface beyond, which was last glaciated during the Rat Creek advance about 13,000 yr ago. The age of the Brisingamen moraine therefore is probably early Holocene. The Brisingamen moraine evidently correlates with moraines near Glacier Peak, near Mount Rainier, in northeastern and central Oregon, in the southern Canadian Rockies, and in the northern U.S. Rocky Mountains. These regional effects suggest that a climatic episode of cooling or increased snowfall affected the entire region some time during the early Holocene.

Voluminous and widespread tephras were produced frequently during the last 36,000 yr of volcanic activity at Mount St. Helens. Numerous tephra sets have been defined by D. R. Mullineaux, J. H. Hyde, and M. Rubin (1975, U.S. Geological Survey Journal of Research, 3, 329–335) on the basis of field relations, Fe–Mg phenocryst assemblage, and 14C chronology and are valuable marker beds for regional stratigraphic studies. In this study modal abundances and mineral compositions were determined (via petrographic and electron microprobe techniques) for numerous samples of individual layers within tephra sets W and Y to evaluate the degree of compositional variability within and between tephra layers and criteria by which to distinguish among Mount St. Helens and other Pacific Northwest tephras. Although individual layers within a set (e.g., We, Wn) cannot be distinguished from each other on the basis of mineralogic characteristics examined, mineral compositions allow distinction among layers W and Y and other Pacific Northwest tephras (e.g., Mazama, Glacier Peak). Fe–Ti oxide compositions and T–fO2 estimates derived using coexisting magnetite—ilmenite are especially useful due to the compositional homogeneity of these minerals both within and between samples of a given unit over a wide geographic area. The silicates show more compositional variability than the oxides, but SiO2/Al2O3 contents in hornblende and Fe/Mg ratios in hypersthene aid in distinguishing among Pacific Northwest tephras.

The bedrock valley of the Pecatonica River north of Freeport, Illinois, contains a thick valley-fill complex of alluvium and drift. Within the valley, loess-capped benches surround hills of silty Illinoian drift. Beneath these benches lie thick deposits of poorly sorted stony silt interbedded with thin lenses of silt, sand, and organic-rich loam. Channel deposits and peat cap the diamicton in places. We interpret the stony silts as solifluction debris shed from silty slopes within the valley-fill during the Early or Middle Wisconsinan (Altonian). Top and bottom radiocarbon dates from a 2.5-m section of peat overlying the diamicton are 26,820 ± 200 and 40,500 ± 1700 yr B.P., respectively. We informally refer to the stony silts, channel sediments, and peat as the “Martintown unit.” Geomorphic position, sediment input, and macrofossils suggest that the dated peat was deposited in a floodplain pond (oxbow?). The pollen record from the peat indicates that a boreal forest dominated this area during the Middle Wisconsinan (late Altonian and Farmdalian). Two pollen zones are recognized: a basal Zone I with Pinus slightly more abundant than Picea and with few herbs and shrubs, and an upper Zone II dominated by Picea and with a larger representation of herbaceous and shrub taxa. Little displacement of vegetation zones is indicated, even though ice advanced to within 100 km of the site during the time of peat accumulation. Because of the problems involved in clearly defining Middle Wisconsinan forest-tundra in mid-latitudes by using analogs of Holocene forest-tundra in high latitudes, caution is required in making geomorphic inferences solely from vegetation data. Together, though, pollen and sediment data indicate that during the Middle Wisconsinan, Pecatonica hillslopes progressed through a sequence of instability-stability-instability related to climatic fluctuations.

The Christensen Mastodon Site, located in central Indiana, contains a rich assemblage of vertebrates (including mastodon, caribou, and giant beaver), invertebrates, and plant macrofossils in situ in lake and bog sediments of late-glacial age. Studies of pollen and plant macrofossils suggest the existence of open, white spruce-dominated boreal forests from > 14,000 yr B.P. to ca. 13,000 yr B.P. The regional decline of spruce, local occurrence of black spruce, white spruce, and larch, immigration of many hardwood taxa (e.g., ash, oak, elm), and the initiation of bog development are recorded beginning about 13,000 yr B.P. Recent reconstructions of late-glacial and early postglacial vegetational changes provide a context for understanding the disappearance of mastodons. The dramatic and rapid restriction of boreal forests along the retreating ice margin from 11,000 to 9000 yr B.P. may have caused a substantial reduction of mastodon populations. A diminished population would be more susceptible to small-scale, stochastic events such as short-term extremes of weather, outbreaks of disease, or predation pressure from paleoindian hunters.

The Lockport Gulf site near Lockport, New York, is a 1.9-m sequence of organic-rich marls having a basal date of approximately 10,920 yr B.P. Six bulk samples with a combined weight of 48 kg produced over 780 individual fossil insects representing five orders, as well as molluscs and abundant plant macrofossils. Coleoptera were represented by 24 families. Several major beetle groups (Carabidae, ground beetles; Hydrophilidae, water scavenger beetles; Elmidae, riffle beetles; Staphylinidae, rove beetles; and Scolytidae, bark beetles) indicate a riffle-and-pool stream, surrounded by marsh, with open riparian habitats and nearby trees. Two zones were recognized based on the Coleoptera assemblages. The Zone 1 fauna (ca. 10,920–9800 yr B.P.) was dominated by boreal forest taxa with abundant bark beetles indicating the presence of spruce trees. In Zone 2 (ca. 9700-9100 yr B.P.) the combination of species with a restricted modern distribution in the Great Lakes-St. Lawrence Forest region and pine and deciduous tree inhabitants suggests a change in vegetation by 9700 yr B.P. Thermal estimates from a faunal analysis indicate that the climatic conditions were stable across the spruce-pine transition, with the mean July temperature in the range of 16° to 18°C.