Evidence of climatic changes in North Atlantic deep-sea cores, with remarks on isotopic paleotemperature analysis

EVIDENCE NORTH

OF

CLIMATIC

ATLANTIC

WITH

CHANGES

DEEP-SEA

REMARKS

ON

PALEOTEMPERATURE

IN

CORES,

ISOTOPIC ANALYSIS

ERIC OLAUSSON

Oceanografiska Institutionen, G6teborg, Sweden

INTRODUCTION

The following discussion concerns the stratigraphy of cores (Nos. 217-299) collected during the cruise of the Albatross in 1948. In connection with this some other problems will also be considered. The stratigraphy is based on foraminiferal analyses made by PHLEGER, PARKER and PE1RSON (1953), by SCHOTT (1952; Core 227) and by ANNIE ANDERSSON (in OLAUSSON, 1960d (Cores 297 and 299) and unpublished); on isotopic paleotemperature analyses (EMILIANI, 1955b. 1956, and 1958); on the chemical and lithological data published in the text and diagrams by OLAUSSON(1960d).

T H E C H R O N O L O G Y OF T H E N O R T H A T L A N T I C C O R E S

The author has tried to divide the core sequences discussed in this paper into alternating warm and cool phases. Every whole glacial and interglacial has received a certain time unit (I-VI) and a corresponding time-rock unit (1-6) ill accordance with that used earlier (OLAuSSON, 1961a, p. 348). Age I is considered to be the time since the last glaciation, i.e. the Postglacial (Holocene, Recent). Age II corresponds to the last ice age (Warm). No attempts have been made to distinguish temperature variations during Age lI. Age llI is considered as the Last Interglacial (Eem or Riss Wtirm). Whether the boundary ll/IlI corresponds to the boundary Eem/Wtirm (ZAGWIjN, 1961) or is slightly younger is left open. Age IV is intended to correspond to the Riss glaciation and the last phase of it is given the number IV I~ (stage 42). The reason for this interpretation is given below. For the Eastern Mediterranean the author assumed that Stage 4 221

222

OLAUSSON, ERIC

was deposited during Riss "provided that Riss does not consist of two more or less separate ice ages" (OLAuSSON, 1961a, p. 349). The author later came to believe in a "Riss complex". In a later paper (OLAuSSON, 1961b, p. 24) two peaks in some Pacific cores were interpreted as corresponding to two substages in the Riss complex with an interstadial between the cool phases, and it was added that "Perhaps it should not be excluded that the climatic development during this sub-age may have been more complicated than is assumed

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FIG. 1. The figure shows the distribution of sapropelitic mud in the Eastern Mediterranean cores, the percentage o f " c o o l " forams according to PARKER(1958), and the time-stratigraphic units used earlier (left column) and the modification suggested here in the text (right column). The figure is composed of Stage 1 from Core 194 and stages 2-5 from Core 189.

EVIDENCE OF CLIMATIC CHANGES IN DEEP-SEA CORES

223

today (cf. the horizons m and m / n in Cores 59, 60, and 74 in Fig. 4)". Age IV in the Eastern Mediterranean sequences is equivalent to Age IV rr for the Pacific and the Atlantic. Furthermore, Va-Vc (Fig. 1) should be included in the Riss complex. No attempt is made to distinguish further the sub-ages (sub-stages) in the Riss complex, except for the last phase IV II (42). Age V is tentatively presumed to correspond to the Holstein interglacial and Age VI to the Mindel Glaciation. The boundary between two stages (ages) is given as a fraction of them.

THE I N F L U E N C E OF THE F O R M A T I O N AND MELTING OF GLACIERS ON CHANGES IN THE ISOTOPIC C O M P O S I T I O N OF SEA-WATER The paleotemperature analysis, devised by UREY, has been used by EMmIANI on several cores. The errors have been dealt with by EMILIANI(1955b).* One of these errors, the changing influence of melting and formation of ice sheets on isotopic composition must be discussed again. The relationship between temperature (T), the per mil difference between the O18/O 16 ratio in the sample and in a fixed standard (3'), and the per mil difference between the water of ancient composition applicable to the tests being examined and the present day marine water of average composition (3) is T = 16.5 -- 4.3(3' -- 8) + 0.14(3' -- 3)z Due to the glacial storage effect there has been a change in the average value of 8. We must here calculate the interval of this change in 818 value for average sea water during the last epochs. The volume of the Antarctic ice sheet is, according to FAIRBRIDGE(1961) 33 × 106 km 3, according to HOLLIN (1962), and SCHYTT (1963) 30 X 106 km 3. The volume of the Greenland ice sheet is 2.6 7, 106 km 3 (FLINT, 1957, p. 52). Using the smaller value for the Antarctic ice sheet, the present value of the ice on Earth would be about 32-6 × 106 km 3 which, if melted would be equivalent to a water layer of about 85 m over the present sea surface. Due to a more or less intense vertical sinking of ice to great depths in ice sheets it must be a decrease with depth of O is in non-central parts of ice sheets However, if in the following calculation we assume for the sake of simplicity that the isotopic ratio is independent of depth and if we use the known data * The possible error introduced by a formation of crystalline crust on foraminiferal tests at a level below their general depth habitat as well as by ion exchange by xoanthelles will not be discussed here. The former error is discussed by ERICSONet al. (1961, p. 225-228). According to them Ptd[eniathza obliquiloculata lacks such a crust while it occurs in Globorotalia tumida. Judging from Cores 58 and 60 (EMIt.IANr,1955b, p. 556) the paleotemperature curves of these two species have the same trend.

224

OLAUSSON, ERIC

for the present surface of ice sheets, we may arrive at an upper limit for 3 in glaciers. I f there is a more intense vertical sinking, which is likely, the absolute value of the 318 will be somewhat larger than those calculated here. The relations and data given by DANSGAARD (1961, p. 56--64), especially his equation (personal communication) are used: 3 (annual precipitation) = 0-70 t (mean annual air temperature)-- 14.2 per rail According to HOLLIN (1962) the mean thickness of the Antarctic ice sheet is 2300 m. Using the calculation for the lapse rate by LOR1US (1962, Fig. 3) we arrive at a mean surface temperature of 34°C which gives a mean 318 of --38 per mil. Similarly we get --39 per mil if we use the temperature data from Dronning M a u d Land given by SCHYTT (1960, p. 197). Another method is to calculate the 318 value on the basis of the contour and temperature maps published by SCHYTT (1963). By means of DANSGAARD'Srelation the isotherms are converted to give 318 values. The contour map can give approx, thickness of the ice. For the latter I have assumed that the rock surface level is equal to sea-level; however, the major channel below sea level between the Ross and the Bellingshausen seas, recorded by U . S . I . G . Y . investigations (BENTLEY et al., 1960) has been consiaered. The thickness of the ice cap in the area within the 3600 m:s equi-distance is taken as 3600 m. By this method I have obtained a 318 value for the Antarctic ice sheet of --40 per mil. The mean ice surface temperature of Greenland is assumed to be --22°C. (DANSGAARD, personal communication) which gives a mean 318 value of --29 per rail. In the same way as for the Antarctic I have calculated the 318 value on the basis of a contour and temperature map, published by DIAMOND and redrawn by DANsGAARD (1961, p. 77). This calculation gives a mean value of --30 per mil. A weighted 318 mean of --39 per mil is thus likely for the two present ice sheets. A volume of ice amounting to 32.6 × 106 km 3 corresponds to a water volume of 29.3 × 10 n km 3 which is equal to 2.1 per cent of the volume of the sea water. These ice sheets, if melted, would lower the present average 3 value in the sea by 0-84 per mil. During ice ages a water layer of up to 100 m (or more, e.g. 120-160 m; see DONN et al., 1962) was transformed into ice caps. The main portion of this water was bound into continental ice sheets in the northern hemisphere. We must calculate a mean 3 for the most important ice caps (those larger than 10 ~ km3). For the Laurentian ice sheet we assume a mean sea level temperature of --10 ° to --13°C. (FLOHN, personal communication). Further, we use an isotopic altitude effect of --1.3 ppm/100 m (the same as for present Greenland; see DANSGAARD, 1961, p. 63--64), and a mean thickness of the ice of 2000 m (DONN, et al., 1962). These figures give for the two temperatures a 3

EVIDENCE OF C L I M A T I C C H A N G E S IN DEEP-SEA CORES

225

value of --32 and --34 per rail respectively or --33 -- 1 per rail (cf. the 3 value of the Greenland ice sheet !). If we use 2500 m here as a tentative mean thickness only to show its influence on the 318 value, we get - - 35 and -- 37 per mil respectively. The Cordilleran ice sheet was quantitatively of less importance. A mean 3 of --25 per rail seems reasonable (cf. EPSTEIN and MAYEDA, 1953, p. 218 and the isotopic ratio for the Scandinavian ice sheet below). For the Scandinavian ice sheet we assume a mean sea level temperature well below the zero point. WOLDSTEDV(1958 Fig. 86) gives a temperature of --4 ° as probable for Middle Europe during the Wiirm maximum. A still lower temperature for the Scandinavian ice sheet must be considered because we have to take into account the ice-bound low-level inversion. According to FLOHN (personal communication), --7 ~ to --10°C is likely. Furthermore, we use an isotopic altitude effect o f - 1-3 ppm/100 m. These figures give for a mean thickness of 1750 m (see _FLINT, 1947, p. 434) a 3 value of --28 and --30 per rail respectively or 29 _ 1 per rail. (For a thickness of 2000 m the 3 values are --30 and --32 per rail.) The Siberian ice sheet may be assumed to have had a mean isotopic ratio of --25 per rail (or lower). The increase of the Antarctic ice sheet during glacials has been caused by a thickening of the ice cap (especially at the periphery) and by a more intense glaciation in the Ross and Weddell seas (HOLLIN, 1962, p. 192). The more peripheral addition to the ice sheet has probably not changed the average 318 value for the whole Antarctic ice sheet. This is because a general lowering of temperature and the increased continentality may have compensated for the peripheral addition of ice with moderate O~8-values. I can therefore use --40 per rail for the additional ice during the glacial ages. Judging from FLINT (1947, p. 434; 1957, p. 53) and from the discussion above we arrive at the following volumes and mean 318 values for the following ice sheets: Volume, l0 s km3 Cordilleran ice sheet L a u r e n t i a n ice sheet S c a n d i n a v i a n ice sheet Siberian ice sheet A n t a r c t i c ice sheet lexcess over the present v o l u m e )

Mean 3

2 25 7 1 5.5

--25 --33 --29 --25 --40

40

-- 33

The other, smaller ice caps and the increase of the Greenland ice sheet will not change the mean 8 calculated here. We take therefore a value of --33 per mil as a representative 818 value for the water transferred to ice sheets

226

OLAUSSON, ERIC

during glacials. 40 × 106 km 3 of ice corresponds to 2.6 per cent of the sea water. Such a large extraction would change the isotopic ratio in the sea water by 0.88 per mil. The highest sea-level position, relative to the land, may have been ~ 18 m during the Eem interglacial. This means that the volume of the ice sheets. compared with the present one, was reduced to about 4/5. For the Holstein interglacial a water level of up to + 3 0 m has been reported, indicating a still larger reduction in the volume of the ice caps (to about 2/3 of the present one). Only for the Late Tertiary have sea-level stands been recorded (see FAIRBRIDGE, 1961) pointing to a more or less total absence of ice sheets. A complete deglaciation of Greenland would raise the sea-level only 6 m. Therefore, most of the additional sea water during the Eem and Holstein interglacials (/>2/3 and 4/5 respectively) must have come from the Antarctic ice sheet. We cannot assume an equal thinning all over the ice sheets. A comparatively more intense thinning in the peripheral parts could be possible. In this case it may be of interest to know the mean isotopic composition of the ice which yearly enters the sea and melts. I assume a steady state for the Antarctic ice, and that the ice which melts yearly is equal in amount and in isotopic composition to the annual precipitation on the Antarctic continent. By the means of a map of the yearly net accumulation in the Antarctic, kindly placed at my disposal by Dr. V. SCHYTT, and by the additional information about the precipitation given by LILJEQUIST(1963, Fig. 2) and the temperature data mentioned above, I have arrived at a mean ~18 value for the precipitation in the Antarctic of --35 per rail. Bearing the Greenland ice sheet in mind we use --33 per mil for all meltwater added to the sea during the Eem and Hotstein interglacials. F r o m the discussion above we can deduce the following range of 3 for average sea water through the ages mentioned and the apparent isotopic

Ages Eem to Recent Holstein to Recent Late Tertiary to Recent

Range of 3 change*

Apparent temp. change

Apparent sea-level changes

--0.17 to +0.88 --0.26 to ~- 1.05 --0.84 to + 1.05

4.4oc 5.5-~C 8-C

--18 to --100 m +30 to -120 m ~-85 to - 120 m

* A sea-level lowering of 120 m means probably a change of about 1-05 per mil, and a lowering of 160 m about 1-4 per mil (cf. FLINT,1947, p. 435 ; DONNet al., 1962, p. 212; and ROaIN, 1964, p. 105). The former may be valid at least for the Riss glacial. At a lower sealevel than -- 100 m the proportion between the ice volumes may have changed. However, the change in 8 for the extracted water was not large. If Donn's values are used the apparent temperature change is 6.9°C during Upper Pleistocene and 9.4°C during Late Tertiary to RePent.

EVIDENCE OF CLIMATIC CHANGES IN DEEP-SEA CORES

227

temperature change which the glacial storage effect may have caused. The present average 3 value for sea water -----O. A consequence of this calculation is that the isotopic paleotemperature curves should always show the temperature variation given above, even in regions where the water temperature throughout the age interval mentioned has been constant. Furthermore, isotopic temperature variations larger than those given above can be assumed to indicate real temperature variations. The isotopic temperature curves from the Pacific (EMILIANI 1955b), show the following temperature amplitudes: Core 58 58 (above 611 cm) 60 61

4.Y~C (Upper Pliocene-Recent) 3°C (Pleistocene) 4-3 ~C (Mindel ?-Recent) 3-6°C (Holstein-Recent)

Judging from these data I am inclined to believe that the water temperature in the equatorial Pacific did not change discernibly during the Pleistocene. This statement concerns only the regions from which the cores are obtained and for the water depth where the analyzed foraminifera lived. The isotopic temperature amplitude for Core 58 is much smaller than the eight degrees discussed above. It suggests the presence of a large Antarctic ice sheet as early as the Upper Pliocene. The isotopic paleotemperature curves show therefore in these cores variations only in the volume of glaciers (or sea water). On the other hand the 818 changes in sea water are ultimately climatically determined. , The curve is therefore of interest in correlating the cores. The discussion of 3 in the temperature equation was based on an assumption of average values for all oceans. Deviations from the value of 3 could temporarily have been much larger in certain basins due to non-average conditions. As examples the paleotemperature curves from the Mediterranean and the North Atlantic cores will now be considered.

MEDITERRANEAN

In the Eastern Mediterranean there have been repeated changes between stagnant and non-stagnant states (see Fig. I). The stagnant phases are due to a high contribution of river effluents and especially from the Black Sea (see KULLENBERG, 1952 and OLAUSSON, t960C, 1961a). This means that the surface salinity and the isotopic composition during such phases have been quite different from the non-stagnant phases even if the temperature of the water was the same. The most pronounced stagnant phase occurred at the boundary IV/III (Riss/Eem; Ot.AUSSON, 1961a). From this stagnant series in Core 189 the

228

OLAUSSON, ERIC

following table can be drawn up (temperature in degrees Celsius, quoted from EMILIANI, 1955a): TABLE 1. Depth cm 312 322 332 342 353 362 372

Deposit Distinctly oxidized deposit --Oxidized deposit Greenish mud Sapropelitic mud Blue mud Light bluish mud Oxidized deposit

Globigerina Globigerina eggeri inflata

Globigerinoidesrubra gomituhts rubra 17"3 21 "2 26"8 32"8

10 6.7

12 8"5 7-5

20 12'6 12

The sample at 312 cm represents the Early Eem, and that from 362 the Late Riss, as indicated by the proportion o f " w a r m " foraminifera to "cooler" ones. The real temperature should therefore be expected to rise throughout this series. However, the isotopic "temperature" was different. The sapropelitic mud, supposed to have been formed during the time of maximal stratification of the water body, shows the maximal deviation in the table. Such a high temperature is impossible, as is also admitted by EMILIANI et al. (1961, p. 686). In addition, the increase for Globigerinoides rubra gomitulus from sample "362" to sample "353" is more than twice that of Globigerina inflata. Basing his conclusions on the literature the author earlier assumed that the former species lives at the surface during the summer while the latter is most frequent during the winter (OLAUSSON, 1961a. p. 377). This difference might be due either to a greater temperature contrast between summer and winter or to more meltwater at the surface during the summer. The latter alternative seems more likely. Also in connection with the other more or less stagnant phases, unexpected "temperatures" are given by Emiliani for the same core (189), for example at 232 cm (blue mud; 2 cm below sapropelitic mud), 271 cm (blue mud; 10 cm above sapropelitic mud), 552 and 561 cm (blue muds; sapropelitic mud between 556 and 561 cm), 698-5 cm (blue mud; 9 cm below sapropelitic mud), and at 772 cm (greenish mud). The most abnormal "temperatures" are thus recorded in sapropelitic mud (corresponding to maximal stratification of the sea water) while in adjacent blue muds (corresponding to a weaker stratification; the beginning and the end of an euxinic phase; see further OLAUSSON (1961a, pp. 350-351)) the "paleotemperature" is lower. Sample "372" represents the maximum of the Riss glaciation. If we compare the "temperature" from this sample with other cores we see that Globigerina iJ~ata occurred at a temperature in Core 189 about 3 ° lower than in

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Core 280 (34 ~ 57' N; 44 ° 16' W). Further Globigerina eggeri occurred at a " t e m p e r a t u r e " about 11 ° higher in the same stage in cores from the Caribbean than in Core 189. The author, therefore, believes that the isotopic temperature in sample "372" is somewhat too low when compared with the other cores. This may be due to the fact that the Mediterranean contained more than the average of 018 in the water during that phase and/or that other regions were poorer in O 18 than expected, thus indicating too high a temperature. As appears from Table 1 the isotopic temperature varies from sample "342" to "312" by 15.5 C . However, this difference is determined by calculations between two subspecies. There is, as is discussed elsewhere, some difference in their ecology. According to analyses at horizons in Core 189, where both species were analysed together, Globigerinoides rubra rubra shows either the same temperature or only 1-3°C higher than G. rubra gomitulus. We may, therefore, reduce the difference between "342" and "312" by 3°C. We arrive thus at an isotopic temperature decrease of 12.5°C from "342" to "312". According to the composition of the foraminiferal assemblages, the water temperature corresponding to sample "312" should nevertheless be higher than that of the water corresponding to the stagnant phase, sample "342". Judging from the foraminifera in Core 189 and in adjacent cores by KULLENBERG (1952) and PARKZR(1958) the interglacial temperature maximum should correspond to the section 310-320 cm. Accordingly, the foraminiferal analysis and the isotopic analysis contrast very strikingly with one another. We will, therefore, assume as a compromise that the temperatures are equal, and that the discrepancy is caused by changes (not dependent on evaporation) in the isotopic composition of the sea water during the two corresponding phases. Judging from the relationship between temperature and 8, given above the value of the analysis of "342" (--2.18 per mil) should be increased by about 2.6 per rail in order to overcome the temperature gap of 12'5°C assumed above. Furthermore we suppose at first that the additional water was meltwater with a 618 of --20 per mil (cf. above p. 225). From this assumption we make the following calculations: "312. . . (non-stagnant state) Paleotemperature, :C Assumed temperature, :C Isotopic composition, per rail Salinity, per mil

17"3 17'3 a 38 a - - 2"6=a(I - - x ) - - 2 0 x 2.6 20--a t-

.

342"' i Meltwater (stagnant state) i addition 29.8 t - 17.3 a - 2.6 S

i -20

230

OLAUSSON, ERIC

x = the proportion of fresh water added to sea water. If a ~-~ 1 it follows that x ~ 0.12 which gives S = 0.88 x 38 or ~ 3 3 per rail. The figures arrived at indicate that the surface water during the stagnant phase contained 12 per cent meltwater. At present there is a freshwater deficit of about 10 per cent. The meltwater entering the Eastern Mediterranean from rivers must be assumed to be very small compared with that discharged from the rivers of the Black Sea. This is clear from estimations of the volumes of the ice caps (FLINT, 1947, p. 434) and the geographical distribution of these ice sheets. The amount of meltwater from the glaciers--in excess of the yearly water supply due to precipitation--into the Black Sea by rivers and eventually also from the Caspian Sea via the Manych, is difficult to estimate. Judging from the calculation of the amount of water bound in the Scandinavian Ice Sheet (up to about t5 times the present volume of the Black Sea (FLINT, 1947)) it seems likely that at least twice the volume of the Black Sea is the probable amount discharged into the Black Sea during the deglaciation phase of Riss. Such water from the Black Sea may have had an isotopic composition of the assumed value. If the additional water was chiefly due to an increase in the river effluent discharged into the Eastern Mediterranean, and/or to an increase in precipitation, a higher O18/O 16 ratio must be assumed for our calculations. In such a case a 8 value of --5 seems reasonable which gives a salinity decrease for the surface water nearly 4 times that calculated above, or 21.7 per mil. This salinity seems to the author to be too low. This is because although we arrive at a salinity of the surface water nearly equivalent to that of the present Black Sea, the fossil contents contradict such a large change. Thus, the Black Sea deposits lack foraminifera (with very few exceptions from older ages). Judging from the foraminiferal analysis of Eastern Mediterranean cores (PARKER, 1958), they occur in large quantities and shallow-water forms were also present. Further, the diatoms in the corresponding euxinic horizon of Core 192 (579-580 cm) are predominantly pelagic marine species indicating a smaller change in salinity (KOLBE, personal communication). Besides, it must be remembered that the "temperature" decrease from "342" to "312" has probably been larger than assumed above ( < t -- 17.3 c in "342") so that the estimated figures for salinity changes at the two assumed 3 of the additional water may be looked upon as minimal values. A large meltwater contamination ( = l o w 01s/01~ ratio) is therefore assumed, supporting the opinion of Black Sea water being the principal cause of this stagnant phase, (see OLAUSSON, 1961a, p. 357). The criticism of the paleotemperature curve of Core 189 could also be extended to the isotopic analysis of the Plio-Pleistocene section from Calabria (EMILIANI et al., 1961).

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Analyses o f percentages of Globigerinoides rubra rubra in relation to G. rubra gomitulus show for Core 196 the following values: Depth, cm

/o°1 G. rubra rubra

Deposit

499-500 501-502 502.8-503.5 503.5-504.3 504.3-505.5

blue mud sapropelitic mud sapropelitic mud sapropelitic mud blue mud

99.5 100 100 100 37

For Core 190 the following results can be given: Blue mud Depth in core, crn

328 %1

G. rubrarubra G. rubra gomituhts G. rubra rubra G. rubra gomitulus

Sapropelitic mud 341

~

%

348 "-:

55 i 3 9 I00 101 45 ~ 32 0 , 0 100 0

161 0

Blue mud!

%

~ 361.5 '-:

%

100 137 0"6 1

I00 ; 163 I00 i 167 0 0 0 0 i

i

"-: 8 : 0

32 68

Size fraction

70 151

150-250 !.t > 250 !z

Judging f r o m these analyses G. rubra gomitulus disappeared during the stagnant phases. W h e n it re,turned after an euxinic time the size of the species is at first smaller. This must be due to an unfavourable influence o f restricted environment on this species. The frequency curve of Globigerinoides rubra rubra in Fig. 2 of the paper by EMILIANI et al. (196l) is. therefore, supposed to represent to an appreciable extent, or exclusively, changes in hydrography, The abnormally high isotopic temperatures are obtained f r o m horizons when G. rubra gomitulus is absent. The Calabria section consists o f alternating layers of clay and diatom mud. In some cases the changes in lithology may be due to changes in h y d r o g r a p h y and productivity. R e m e m b e r i n g that diatomaceous layers only occur in connection with euxinic phases in the Eastern Mediterranean cores (OLAUSSON, 1961a, diatom mud) and that such layers gave the most abnormally high " t e m p e r a t u r e s " the author assumes that some of the diatomaceous horizons in the section from Calabria are related to more pluvial times (~,,~less surface salinity). This is supported by the high isotopic temperature in connection with some o f those layers, and the high percentages of Globigerinoides rubra rubra (sample 79 and 61). The author considers it, therefore, difficult or impossible to state anything about the real water temperature changes on the basis of EMIL1ANI'S analysis o f the section f r o m Calabria.

232

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ATLANTIC

OCEAN

(a) Advance Phase The amount of water transferred yearly from the North Atlantic to the ice caps is too small compared with the large surface inflow from the South Atlantic. Together with the strong possibilities of convective sinking in the North Atlantic, these make it likely that "average" isotopic composition has been maintained during the advance phases.

(b) Deglaciation Phase The deglaciation occurred step-wise after the Wiirm. Judging from investigation in the Great Lakes region (FLINT, 1957, p. 326) there was a rapid retreat of the ice margin during the Bolling (,-~13,400-12,600 B.P.) and Allerod (,-~11,900-10,800 B.P.) phases of about 500 miles, and a third retreat in the Pre-Boreal subage (,-~10,000-8700 B.P.). Such a trend was also recorded in Europe. At 9000 B.P. only about 15 per cent of the ice sheets remained, in excess of the present glaciers (FAmBR~DGE, 1961). This figure assumes that the sea level was at --100 m at the Wiirm maximum and that no correction is made for isostatic adjustment. The present rise of sea level was estimated by Mosby at 1.12 mm per year (FA1RBRIDGE, 1961, p. 103). During the warm subages Boiling, Allerod, and Pre-Boreal sea level rose rapidly. Two to three centimeters per year is estimated by FA1RBRIDGE (1961, p. 156); between 11,000 to 9000 B.P. 2.2 cm/yr (BROECKER, 1961). TO this should be added the yearly amount of precipitation. Probably more than 3/4 of this meltwater reached the North Atlantic. Taking 2 era/year we arrive at a yearly discharge of glacial-water to the North Atlantic and adjacent seas corresponding to a water layer of about 10 cm. Owing to the surface circulation in the North Atlantic, the meltwater was distributed to all parts of this ocean. The Mediterranean water left the Strait of Gibraltar as a surface current, at least at the end of the Riss glaciation (as stated above) and this reduced or stopped the sinking in that region. This water was instead added to the surface water of the North Atlantic. All this occurred during a period of increasing temperature and a general lowering of the salinity (1 per rail), factors which temporarily lessened the formation of deep water, or bottom water, in the North Atlantic. Further, there must have been an excess outflow during the deglaciation into the North Atlantic amounting to approximately 20 x 106 km z. The renewal of the North Atlantic water takes place by means of a surface inflow (South Equatorial Current), by convective sinking of surface water which through cooling obtained sufficient density to be able to sink, and by outflow in the lower part of the water body. During the deglaciation phase the second link in the renewal chain mentioned here was insufficient for

EVIDENCE OF CLIMATIC CHANGES IN DEEP-SEA CORES

233

maintaining isotopic equilibrium with other oceans. This is assumed because of the large volume of the meltwater and a density lower than that of sea water. This prevented both mixing and sinking, Furthermore the meltwater spread to all parts of the ocean while the sinking occurred only within one or perhaps a few small areas. Therefore the surface water had an abnormally low O18/O 16 ratio. If some part of this isotopically light water remained in the surface water and became enriched during times of rapid deglaciation, an unusual isotopic ratio would have existed in the North Atlantic surface water. For example: if only 2 per cent of this meltwater was accumulated every year in the uppermost 100 m over a period of 2000 years, corresponding to a salinity decrease of about 2 per mil, it would be sufficient to cause a "temperature" increase of about 5°C (if 8 = --25 per rail). The geological reasons for this assumption of temporarily non-average isotopic composition in the North Atlantic water are as follows: 1. The paleotemperature curves of the North Atlantic cores have a peculiar trend with the highest temperature just at the very beginning of an interglacial and with the lowest at the end of the succeeding glacial with a more or less continuous temperature decrease between these points (see e.g. Cores A 179-4 or the generalized temperature curve, EMILIANI(1955b), ROSHOLTet al. (I 961)). 2. The age of the Post-Wiirmian isotopic rise has been discussed in several papers by EMILIANIand by BROECKERet al. (1960). The latter authors reached a conclusion in striking contrast with that reached here. Judging from Cores A 179-4, A 240-M 1 (RosHOLT et al., 1962), and A 254-Br-C (RosHOLT et al., 1962), the midpoint of the isotopic temperature rise was of an age of about 11.000-12,000 yr B.P. while the upper end may be about 9000-11,000 yr B.P. The author therefore believes that the isotopic temperature rise is isochronous with the most intensive phases of deglaciation. The Postglacial began with the Pre-Boreal time (~10,000 yr B.P.) when a large temperature rise commenced (IvERSEN, 1954, p. 98). The climatic optimum occurred about 6000 B.P. (see BUTZER, 1958; FLINT and BRANDTNER, 1961; IVERSEN, 1960; NILSSON, 1961 and OLAUSSON, 1957). If the datings of the isotopic temperature rise are acceptable, which they may be, the "paleotemperature" curve should indicate a chronological phase difference of some thousands of years between the North Atlantic and continental deposits as shown by climatological evidence. The C14-age of the Postglacial horizons with highest frequency of " w a r m " foraminifera (Globorotalia menardii and others) seems to be much younger than the 9000-11.000 years mentioned above in the cores investigated by ERICSON et al. (1961). 3. The paleotemperature analysis of the Pacific cores 58 and 61 gave a higher "temperature" after the slightly cooler sub-age within the warm age III (Eemian) than before it. In the North Atlantic cores the opposite seems to have

234

OLAUSSON, ERIC

been the case. Is this real or n o t ? I f it is then it m a y be due to too high a n a m o u n t of 0 is in the Pacific d u r i n g the Early Eem, giving eventually too low a t e m p e r a t u r e , a n d to the fact that isotopic e q u i l i b r i u m between all oceans was reached later (cf. the sea-level changes d u r i n g R i s s - W t i r m : M o n a s t i r I--~ + 18 m, M o n a s t i r I I - ~ -+- 8 m).

20 25°C

0

~b 20

3'0 %

2s-c

0

i~320 " ,

,

10

25°C

io

20%

=

.

3O'/.

4"1.

73

0 t~620'/,

'

4'o'/.

FIG. 2. The isotopic paleotemperature and the percentages of grains larger than 62 microns in the Eemian stage of Cores 234, 246, and A 180-73. The frequency of the Globorotalia menardfi-tumida group is given for Cores 234 and 246 (the curve with the letter a). For Core A 180-73 the ratios of numbers of Globorotalia menardii to weight of material coarser than 74 microns in each sample is given (the curve with the letter b). The curves are redrawn from EMILIANI (1955b), and EPdCSON and WOLLIN(1956); the foraminifera analysis of Cores 234 and 246 was performed by PHLEGERet al. (1953).

EVIDENCE OF CLIMATIC CHANGES IN DEEP-SEA CORES

235

4. F r o m Figs 2-3 we observe the following details : The rise in the isotopic curves o c c u r r e d earlier t h a n that o f the w a r m - w a t e r G l o b o r o t a l i a m e n a r d i i - t u m i d a group. These o b s e r v a t i o n s are valid for all cores at the t r a n s i t i o n 3/4 a n d for Cores 234 a n d 235 at the b o u n d a r y 1/2. A t the t r a n s i t i o n 3/4 the f o r a m i n i f e r a reach their m a x i m u m in c o n n e c t i o n with the 15

20°C

10 %

0

20

40

20

25°C

0

5%

20

40.

20

0

20 ,

40%

25°C .i

2~0

0

r

10'/o

z.,'OO/,

FIG. 3. The isotopic paleotemperature and the percentages of grains larger than 62 ~ in Cores 234, 235 and 246 according to EMtLtANI(1955b). The frequency of Globorotalia menardii (the curve with the letter c) and the G. menardii-tumida group is given (the curve with the letter a). The foraminifera analyses were made by PHLEGERet aL (op. cir.) and A. ANDERSON. second isotopic p e a k a n d this s i m u l t a n e o u s l y with the m a x i m a l f r e q u e n c y o f ( f o r a m i n i f e r a ) grains larger t h a n 62 ~. As discussed elsewhere these species were less sensitive to solution. T h e i r increase in n u m b e r a n d p r o p o r t i o n c a n n o t therefore be influenced so m u c h by the vertical m o v e m e n t o f the

236

OLAUSSON, ERIC

critical depth for calcium carbonate (roughly 500 m) as was the case with other species (see Fig. 7). This is clear from a calculation based on a constant number of Globigerinoides sacculifera through the rise of the peak; such an assumption will not give the interglacial percentages for the Globorotalia menardii-tumida group. In Core 234 (and 233 as well) the first peak of foraminifera at the stage boundary 3/4 was isochronous with the first isotopic temperature peak. This means a m o n g other things an appreciable increase in number of Globigerinoides sacculifera..Also Orbulina unirersa reached a maximum here. Both species are less resistant to solution, and being present also in the lower "cooler" stages, their increase may be more or less due to a decreased dissolution at the bottom. (Nothing is known about variations in their production; a decreased deposition of non-calcareous matter cannot cause such an increase of easily dissolved foraminifera). In Cores 246 and A 180-73 the first increase of foraminifera and carbonates happened after the first isotopic maximum. The percentage increase of " w a r m " foraminifera at the stage boundary 1/2 occurred in Cores 234 and 235 later than the rise in the isotopic curve. On the other hand the isotopic curve of Core 246 underwent a rise at the same time as the percentage increase of " w a r m " foraminifera. Remembering the location of the cores and the position of the present current boundaries we see that there is a phase difference in time where the water was North Atlantic in origin (Cores 234 and 235). This is also evident in Cores 246 and A 180-73 obtained below the present position of the South Equatorial Current (South Atlantic water) at the boundary 3/4 but not at 1/2. If we assume that the surface water at the station for Cores 246 and A 180-73 at the time IV/Ill was to a large extent North Atlantic water (e.g. due to a more southerly position of the current boundary) these observations are easily understood. The increase of " w a r m " foraminifera was isochronous in the cores. The end of the rise in the isotopic curve was synchronous in the cores at the boundary 3/4 but occurred later in Core 246 compared with Cores 234 and 235 at the boundary 1/2 due to differences in the origin of tile surface water. The rapid and too early rise in the isotopic temperature curve cannot be due to a rapid warming of the ocean. It must be due to meltwater contamination of the surface water as discussed above. The first isotopic temperature peak in Stage 3 and 1 is consequently supposed to be more or less due to the supply of meltwater poor in 018 to the Atlantic Ocean. The climatic optimum during the Eem interglacial is supposed to have fallen about the time of the formation of the second isotopic temperature peak. Whether it corresponds to the second m a x i m u m or to one of the adjacent minima is left open. Judging from these statements the O18/O 16 ratio of the North Atlantic water has been much lower than "average" during part of the deglaciation phases,

EVIDENCE OF CLIMATIC CHANGES 1N DEEP-SEA CORES

237

and the first rapid increase in the isotopic temperature curve was influenced by meltwater contamination in the North Atlantic. This was probably the case even during older deglaciation periods. If we consider the first isotopic peak in each non-glacial age as questionable, we have an isotopic temperature range of 5-8 ° for the North Atlantic. Subtracting 5 ~ from the figures for the reason mentioned above (p. 226) we get 0-3°C as the approximate real temperature variation in the subtropical part of the North Atlantic. The increase in the foraminifera content and the carbonates is not simultaneous in the cores from the two regions. It occurred earlier in the Sierra Leone Basin. This is very probably due to differences in the amount of dissolution possible in the bottom waters in the two regions. In the basin, where Core 234 was collected, most (70-90 per cent) of the bottom water is of Arctic origin (Wt)sr, 1936: Beilage XXII). Cores 246 and A 180-73 came from an area where most of the bottom water may be of Antarctic origin (probably up to 70 per cent). During the ice age a greater amount of Arctic water and of lower temperature can therefore be supposed to have been transported to the two regions in question. Owing to the warming of the ocean and the supply of meltwater discussed elsewhere a decrease in the sinking rate is supposed to have occurred during the transition from ice age to non-glacial age. Such a change in the rate of the influx and the temperature of the Arctic bottom water would be especially noticeable in regions where this component dominates (e.g. Sierra Leone Basin) while the effect would become less towards areas where it is only of minor importance (Core 246). Changes in the influx of Antarctic bottom water especially influence, of course, areas where this component dominates and decreases in direction away from its source. If the influx of polar bottom waters, or the dissolution of power in the regions discussed here did not undergo a simultaneous decrease, the result must be a phase difference in time between the two regions. The increase in the content of foraminifera as well as of carbonates may therefore not be simultaneous everywhere in the North Atlantic. It occurred first where the bottom water is in the main of Arctic origin and occurred last where it is principally of Antarctic origin. Another error in the isotopic paleotemperature determination might be due to considerable changes in atmospheric and oceanic circulations. An indication of this type of error might be the following. The paleotemperature curves of Cores 280 and 234 from Stage 5 show lower "temperatures" than during the Eem and Postglacial. At least for the former curve the "temperature" cannot be real. However, it is beyond the author's competence to state how an age with a dominance of an extreme high-latitude zonal circulation pattern ("high index") could influence the amount of precipitation in both regions and the distribution and the amount transported of upwelled water in the Central

238

OLAUSSON, ERIC

Atlantic. The recorded "temperature" in the Lamont cores, from the Caribbean, is higher in Stage 5 than in younger stages--omitting the first peak after every glaciation. Pre-Pleistocene isotopic paleotemperature curves should not be discussed in this connection, because of the difficulties in assuming the variable 8. This is shown by the following: the upper part of the Miocene section of Core A 164-30 (EM1LIANI, 1956) shows isotopic temperatures which are about equal to those in the Wiirmian stage of Core A 172-6.

REMARKS ON THE CALCIUM CARBONATE CURVES ARRHENIUS (1952) showed that the carbonate content of his Pacific cores was higher in the glacial stages than in the interglacial ones. Such a distribution of carbonate is also found in one core from the Indian Ocean (Core 153; OLAUSSON, 1960b) but is not observed, as far as the author knows, in North Atlantic deep-sea cores. Calcium carbonate curves, which do not show correlations with the climatic development, are probably those of Cores 93 and 152 (OLAUSSON, 1960a, 1960b). The North Atlantic cores from shallower depths (e.g. 267,290, 242 2) seem also to belong to this group (see OLAUSSON, 1960d). A higher content of calcium carbonate in non-glacial stages than in glacial ones occurs in those North Atlantic cores which were collected from moderate or great depths (e.g. 233,234, 236, 280, 288,289, 292, 296, 299; see OLAUSSON, 1960d). In the North Atlantic and the Central Pacific cores a carbonate minimum in one region is isochronous with a maximum in the other. It is likely that a simple relationship exists. When in equilibrium the rate of withdrawal of calcium carbonates in marine sediments can be supposed to be equal to the amount of carbonates from weathering on land discharged by rivers into the sea. If the deposition increases in one region, a decrease is likely in another, insofar as the river discharge and the amount of carbon dioxide in the sea do not vary markedly. The dissolving capacity of the water is dependent on the content of carbonic acid, the temperature (and the depth). A lowering of the water temperature is evident for the North Atlantic during glacials but it is somewhat questionable for the Equatorial Pacific at least during their earlier phases. Up to a glacial maximum around 3 per cent of the water is transferred to ice sheets. An equal percentage of the carbon dioxide in the sea water, which is about 1.5 times its content in the present atmosphere, must have been precipitated somewhere if a portion of it did not remain dissolved in the water where the temperature decreased (as in the North Atlantic).

EVIDENCE

OF C L I M A T I C

CHANGES

IN DEEP-SEA

CORES

239

T h e b o t t o m water o f the N o r t h A t l a n t i c d u r i n g a glacial can dissolve m o r e c a l c a r e o u s m a t t e r f r o m the sea floor. As the b o t t o m water moves ( t o w a r d the Pacific) the u n d e r s a t u r a t e d water u n d o u b t e d l y takes up m o r e and m o r e c a l c i u m c a r b o n a t e a n d can therefore dissolve less and less. F i n a l l y it m a y b e c o m e s a t u r a t e d with calcium c a r b o n a t e . F o r a m i n i f e r a l tests a n d rods a n d discs o f C o c c o l i t h o p h o r i d a e m a y b e c o m e buried there w i t h o u t b e i n g dissolved

30000_

Cat03 40 60

20

i

m

8o

i

T

l I I

'

I I I ,

3500 x

:

.."

dO

O .'A~

I I

I

O.

-234

~242

i

0

-243

D•

~233 0

I I I I I I

40001 •0

x

4500~ IX

D

A

~ 235

/ / / / / 5000L

/

FIG. 4. The figure shows the percentage of CaC03 in some substages of cores from the central part of the Atlantic Ocean. Maximal values are used for the nonglacial substages and minimal ones for the glacial substages. The carbonate contents are given in relation to the coring depths (see Fig. 5 for symbols). by the water• This p r o b a b l y occurred in the e q u a t o r i a l Eastern and C e n t r a l Pacific d u r i n g the glacials. It m e a n s in reality a c a r b o n a t e t r a n s p o r t f r o m the A t l a n t i c to the Pacific d u r i n g these periods. The percental increase o f Globigerina a n d a decrease o f Globorotalia and Pulleniatina in glacial stages o f

240

OLAUSSON,

ERIC

equatorial Pacific cores are thus a result of decreased solubility in the water (see next section). During non-glacials the deposition of carbonates increased in the North Atlantic and, as a consequence, it decreased in the Pacific. C a CO3 2000 m

0

20 • • x [] 0

40

60

80 "/. • • /~ /17 ]

Postg|acia| WormlT Worm T Eem Riss ~' Upper Holstein

290

/ I

/ I/

f

/ / / /

2500 ~

/o,?

t, •

/

I ! I I ! ! / ! t I

'

3000 -

289 288

• i ~ • [3

x 0

I @

= 0

~

292

x I I I I I

~

~ 293

b 3500

,-

O "

0~.

4000

~ 296

.-

~x /, / /t /

' I © !

r~'" X 299

4500

I~

/ /

/

0

•

286

112

/

d

I

'28~

II~CIZ~'

=L280

299 ~

I `299

I

FIG, 5. The percentage of CaCOa in s6me cores from North Au~erican Basin-West European Basin. Maximal values are used for the non-glacial substages and minimal for the glacial ones. The carbonate contents are given in relation to the coring depths.

During the Upper Holstein (Stage 5) it reached its maximum in the North Atlantic (see below) simultaneously with minimal numbers of Globigerina,

EVIDENCE OF CLIMATIC CHANGES IN DEEP-SEA CORES

241

number of entire foraminiferal tests, and carbonates ill equatorial Pacific cores (index horizon l; OLAUSSON, 1961b, Fig. 5). The changing carbonate content may reflect climatically controlled changes in the intensity of the deep-water circulation. The carbonate content varies more from glacial to non-glacial stages in cores from great depths than in shallower depths (Figs. 4-5). This is in agreement with the statement above. Judging from Fig. 4 the carbonate content in that region decreases in the following order: Upper Holstein > Postglacial > E e m > Wfirm II = Riss II (Age IV II) > W~irm I The carbonates in Cores 233-234 seem to decrease more than in Core 243. During the Early W~irm the dissolution of carbonates was more intensified than in other periods of the Upper Pleistocene, and/or there was an unusually high sedimentation of siliceous organic remains and clay. In the northernmost part of the Atlantic Ocean. the differences in carbonate content between the cooler substages mentioned above are small even if there is a tendency for the carbonate minimum formed during the Late Wtirm to be the lowest (Fig. 5). One difference, when compared with the other section, is that the carbonate content in the Eemian stages seems to be larger than in the Postglacial sections. However, in cores with an appreciable thickness of Stage 1 (Cores 288 and 296), there are smaller or no differences, showing that the carbonate contents in Postglacial and Eemia'n sections would be about equal in this region. The data is too sparse to trace differences between the basins on either side of the ridge. The distribution of carbonates is highest and very regular in Stage 5.

ON THE S E L E C T I V E D I S S O L U T I O N FORAM INIFERA

OF PLANKTONIC

There are differences in the rate of dissolution among foraminifera. As was stated by SCHOTT (1935, pp. 102 and 108) the percentage of Globigerinoides saccul(fera decreases with increasing bottom depth. This is also the case with Globigerinoides rubra and Orbulina universa. On the other hand, the thicker-walled Globorotalia species dissolve more slowly and are therefore proportionately more abundant with increasing depth (SCHoTT, 1935, abb. 35. 36 and 41). PHLEGER et al. (op. cir. p. 117) stated that "many or most of the Glob(eerina species appear to be attached before, the Globorotalia species: Ptdleniatina obliquiloculata seems to be the last to be destroyed". According to PHLEGER et al. (1953) the highest percentage of Globorotalia tumida in our surface samples was found in the pilot Core 230A, (depth 5500 m) namely 61 per cent (1400 tests). Another rather high percentage is noted for Core 261 (depth 5128 m) and for Core 240A (depth 7500 m). In

242

OLAUSSON~ ERIC

these surface samples the percentages of Globigerhroides sacculifera were very low. Cores 230 and 230A are worthy of discussion because they give some information about the changes in the dissolution rate. Core 230A contained, as stated above, foraminifera in its upper layers, while the carbonate content decreased deeper down. Core 230 contained no, or only very few, foraminifera and the carbonate content was very low. The "Meteor" core from Station 264 (CoRRENS, 1937), collected near Core 230, consisted of a limy layer which graded downward into a red clay (0-3-I cm = 22~/~ CaCO3; 65-69 cm = 0"5 ~ CaCOa). The red clay in Core 230 does not therefore represent the modern condition at this place, as stated by PHLEGER et al. (1953, p. 80). No higher content of lime is registered from later interglacials, as the red clay down to ca. 1093 cm is supposed to represent Stage 2 and some other earlier stages. This is in agreement with observation from other cores in the Central Atlantic where the highest carbonate content occurred in the Postglacial stage. At the present sea surface, the " M e t e o r " Expedition collected much more Globigerinoides saccul!fera than Globorotalia tumida and menardii. At the bottom the opposite was true (ScHoTT, 1935). This observation is also valid for older stages (Fig. 6). In four cores from the Equatorial Atlantic, 233, 234, 236 and 246, the highest frequency for the Globorotalia menardii-tumida group in Stage 3 was found in Core 236. Its 246 (3202) 0

10

234

233

(3579)

(4125)

~0 20

2. . . .

30

0

10

20 30

236 (5o 65 ) ~0

0

~0 20

30

&O i

50 i

i

I~ - ~ [

/ FIG. 6. The percental distribution of Globorotalia menardii-tumida group in stages 2--4 of Cores 246, 234, 233 and 236. The coring depths are given within parenthesis.

EVIDENCE OF CLIMATIC CHANGES IN DEEP-SEA CORES

243

percentage decreases then with decreasing coring depths, reaching its minimum in Core 246. This also seems true for the lower part of Stage 4. In Stage 2 this group is absent. Looking at the distributions there is a decreasing percentage in the above mentioned cores: Stage 3 > Stage 41 > Stage 2 This trend is interpreted as indicating that Age IV I except for fairly brief intervals was a sub-age with a climate somewhat between glacial and real interglacial. The same result is also suggested by discussions of the Riss complex. Even the isotopic paleotemperature data seem to support this view.

STRATIGRAPHY

OF SOME CORES

The following contains some notes on the stratigraphy and other features in the cores from the North Atlantic, which are of special interest.

Core 227. The numbers of mineral grains larger than 6 ~ are generally lower in the "cooler" sections of the core (about half those in the " w a r m " stages). The layers with a higher carbon content coincide fairly well with the parts of the core deposited during the Riss-Wiirm and the Postglacial. The last point is interesting. The rate of sedimentation was higher (approx. twice) during the warm ages than during the interjacent cool one (WiJrm) and the carbon content is 4-8 times higher in the non-glacial stages. This may chiefly depend on differences in the rate of organic production. A different circulation pattern may also have predominated during the cooler age. During such times the upwelling rate in this region can be assumed to have decreased or temporarily stopped. This hypothesis explains the difference in the sedimentation rate as well as the change in the carbon content. Cores 233-236, 243, 246 and 280. With carbonate, foraminiferal, and paleotemperature curves stages 1--42 are easily distinguished in these cores. The sequence of Core 236 differs by the presence of two carbonate minima in the upper section of Stage 2. However, the foraminiferal curves show no differences in their trends. In Cores 243 and 246 the curves of CaCO3/P20~ have maxima in Stages I and 3. tn order to interpret Core 235 some foraminiferal analyses were carried out on the pilot Core 235A. A comparison between them and the paleotemperature curve, given by EMILIANI (1955b, p. 555), indicates that there is for this core a good, positive correlation between the temperature and the number and percentage of Globigerinoides sacculifera. Almost as good a correlation was found for the Globigerinoides rubra and the Globorotalia menardii-tumida group, while Globigerina eggeri showed a negative correlation. From the trends of these curves and the carbonate content, Core 235

244

OLAUSSON, ERIC

has been interpreted. Except for a more marked carbonate minimum in Stage 3 the trend of this carbonate curve is similar to those in Cores 233 and 234. Also the foraminiferal curves of Cores 235 in these stages are well correlated with those in the other Central Atlantic cores. Age IV represents, as stated above, a period with complicated climatic development. This age is clearly delimited from contiguous ages in some Pacific cores (OLAuSSON, 1961b, Fig. 5). Judging from this figure and from the paleotemperature curves of Atlantic Cores 234, 280 and the Lamont Cores A 172-4, Age IV had an initial and a finishing cool phase with an interstadial between them. This interstadial seems to have been interrupted by one cooler phase, and, judging from Figs. 7, 8 and 10, is, on the whole, characterized by a changing climate. This interpretation is in agreement with the data from continental deposits. Much evidence has accumulated in recent

280

233 40

23/* ?o

lo,/. ~

234

7o

235 lo% ~

236 70

lo% 4o

2 ,I ]

15-

i

4o,i. ~

i

:i

T3i

243 ~

2&6

246

40%

:

? ! ii~

?,

(

iI

FIG. 7. The distribution in calcium carbonate and paleotemperature curves of some Central Atlantic cores. Stages and stage boundaries are marked.

EVIDENCE OF CLIMATIC CHANGES IN DEEP-SEA CORES

245

years to support the theory that the Riss (as well as the Mindel) consisted o f cool sub-ages with temperate phases between them. Thus, Riss is recorded in Europe as consisting o f two phases Drenthe and Warthe with probably a fairly long (?) interstadial between the t w o (WOLDSTEDT, 1958; p. 22). The corresponding Illinoian in the United States is also supposed to be divided into two (or three) phases (FRYE, 1962). Bearing in mind that the isotopic curve shows phases within the W a r m there seems to be no reason to correlate the Riss with a single isotopic m i n i m u m of short duration (~,30.000 years).

20

289

288

280

280 Ce

50

80%

50

80%

50

-,

1"

2

BO%

292 50

e,O°/.

296.

293

50

aO°l*

3

,

50

BO'/,

299 20

50

p

i

3" &.

7-

4

ii

1t-

I

"i

'L 17-t

. i

''i

i

!,Ii

?

4

I~ I

FiG. 8. The distribution of calcium carbonate in some cores from North American Basin-West European Basin and the paleotemperature curve of Core 280. Stages and stage boundaries are marked.

EO%

246

OLAUSSON, ERIC

The topography of the Sierra Leone Rise is, as appears from K o c z v (1956) and from bathymetric charts, rough. Thus, there are sea-mounts, probably of volcanic origin, between and south of the stations where Cores 233 and 234 were taken. At about the time of the transition between Age V and Age IV

280

O0

288 10 %

0

289 10"/,

0

292 10%

0

293 , 10"/.

0

296 10"I.

299 10%

010"I.

2 2 2

3

8

)

9 10 11 12

7"

\

13 14

15 16 17

18

FIG. 9. The percentage of Globorotalia truncatulinoides in some cores. Analysts: PHLEGER et al. (op. cir.) and A. ANDERSSON.

EVIDENCE OF CLIMATIC CHANGES IN DEEP-SEA CORES

247

there was volcanic activity not far from the stations where Cores 233-235 were obtained. In the lowest portions of Stage 4 these cores contain volcanic matter and their sediments seem to have been disturbed. However, comparison with Core 280 makes it likely that the disturbances were not too important because the paleotemperature and carbonate curves for Cores 234 and 280 also show marked similarities at this boundary (and Stage 5), where there is no reason to assume important disturbances. Core 236 has disturbances of more local character. The lower boundary for Stage 4 is not defined, so that no interpretation of this core below 880 cm seems possible. This is in agreement with Phleger's statement. In Core 233 old foraminifera such as Globigerina venezuetana occur in the sample at 801-802 cm but then again only below 1013.5 cm. In Core 234 the old foraminifera occur rarely between 967-5 and 1054.5 cm and commonly down to 1112 cm. This indicates that the section of Core 234 with an even carbonate distribution (750-968 cm) represents an age of undisturbed sedimentation while the section between 968-1112 cm represents one with disturbances. The lower section also differs with the presence of volcanic matter. One further difference between these two cores is the distribution of the Globorotalia menardii-tumida group below Stage 4. In addition, the trend of the carbonate curves is different. Therefore the section of Core 233 below 8 m is considered to be Early Pleistocene with a hiatus at the level mentioned. The transition between Stages 4 and 5 is distinct in Cores 234, 235 and 280 (Fig. 10). Thus, in the uppermost part of Stage 5 the paleotemperature curve is comparatively high (for Lamont Cores A 172-6 and A 179-4 while Cores 234 and 280 display a moderate "temperature", a matter discussed later on p. 237) and especially more uniform (concerns Cores 234, 280. and the mentioned Lamont cores). Furthermore, the carbonate content is very high and evenly distributed for the Swedish cores (the carbonate content has not been published for the Lamont cores). A third characteristic is the absence of the Globorotalia menardii-tumida group (in Cores 234, 235 (in the Lamont cores their frequency is very low, as mentioned by ERICSON et al., 1961, p. 266), in Core 280 this group occurs very infrequently, probably about the same as in the L a m o n t cores). This temporary absence (or very low frequency) within a sub-age, in view of paleotemperature and carbonate curves and the presence of more easily dissolved foraminifera, must be due to other factors than intense dissolution (too low dissolving capacity of the bottom water ?) and low temperature. It seems to be a more or less regional phenomenon. Age V is, according to the author's interpretation of the Pacific cores, a considerably long and warm age. It extended from the end of a cool age (index horizon h, OLAUSSON, 1961b, p. 23) up to another with index m. The first is tentatively correlated with the Mindel (Mindel II) and the last with the Early Riss (Riss I). Three small climatic deteriorations are recorded (i, k

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and l). The indications for climatic changes observed within the Great Interglacial were considered earlier (OLAUSSON,1961a, p. 349-350). In the Pacific cores the continuous curve of Sphaeroidinella ends within this stage and Pulleniatina increases in frequency associated with the index horizon h ( = M i n d e l II). The latter indicates probably a very low rate of sedimentation. The Globorotalia menardii-tumida group within this stage is minimal in number per gram of sediment (cf. above). The changes in frequency of the foraminifera in the Pacific cores seem to be a result of selective dissolution, as was also pointed out by ARRHENIUS (1952). The division in Core 234 between 750-968 cm represents Stage 5, but probably only the upper part of this stage. The division in this core down to 1112 cm may also belong to this stage but, owing to disturbances in the sedimentation process and the supply of older material, its age must be left open. The section below 1112 cm is Miocene in age (PHLEGER e t al., 1953). Lower-Middle Miocene, according to EMILIANI, 1956, p. 282. There is thus a considerable hiatus at a depth of 1112 cm. In the Lamont Core (A 179-4) the Globorotalia menardii-tumida group is absent down to the lowest paleotemperature minimum (Fig. I0). They occur up to this minimum but disappear when the isotopic temperature increased. This layer may correspond to the level around 870 cm in Core 235 where the drop in the Globorotalia group occ/urs, where Globigerhm pachyderma reaches 5 per cent, where Globigerina eggeri has a frequency of 58 per cent and Globigerinoides saccul!ferd was only 8 per cent of the population (Fig. 10). Hence, this must represent a cooler sub-age. A further interpretation of Core 235 is difficult. At 1120 cm there was a drop in the curve of Sphaeroidinella (Fig. 10). Whether this has something in c o m m o n with a similar trend in the Pacific cores remains to be proved. Up to this level Pulleniatina is on the increase, a trend also noticed for Pacific cores. Below this horizon the Globorotalia menardii-tumida group has a very low frequency. Much the same trends as those mentioned for foraminifera may be noticed in the lowermost 3 m of Core 233 for carbonate curves. In both cores old foraminifera occur singly within the section discussed. Whether the horizon below 1120 cm in Core 235 can be correlated with the index horizon h in Pacific cores (Mindel II ?) is left open for the time being. In the L a m o n t cores mentioned above, Globorotalia truncatulinoides changes from a right coiling to left coiling shell during the first cool phase of Age IV. Such a change also occurs farther down during the two cooler phases in Age V (ERICSON and WOLLIN, 1956, p. 123). This change in the coiling direction within Stage 5 was also found in Core 280 associated with the frequency minima of the species, namely at around 1170 cm (92 per cent of the tests of this species were coiled to the left at 1169-1170 cm), and at 1300 cm (90 per cent coiled to the left at 1298-1299.5 cm). Adjacent horizons as

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well as others in the core have shown quite other percentages (30 per cent at 478.5--479.5 cm; 44 per cent at 543-5-545 cm; 58 per cent at 998-99.5 cm; 48 per cent at 1248-1249-5 cm). The horizon at 1170 cm coincides well with minima in the curves of the isotopic paleotemperature, the carbonate and grains larger than 62 tz (EMmIANI, 1958; OLAUSSON, 1960d, and Fig. I0). All these characteristics (except for the curve of grains > 62/~) can also be found around the 1300-cm level. Another similarity between the Lamont cores and Core 280 may be the occurrence of Globorotalia punctulata punctulata. In the two Lamont cores this species is "abundant" in Stage 5 but considerably less frequent in younger stages; in the three lowest samples (of Stage 5) from Core A 172-6 the foraminifera species occurs more sparsely (ERICSON and WOLLIN, 1956). This trend is also found in Core 280; it is most frequent in the middle and upper part of Stage 5, and lessens in frequency below 1350 cm. From the above it seems likely that the two lowest climatic deteriorations recorded in Lamont cores are also noticeable in the data from Core 280. At about 15 m there is another cooler phase (or two slightly separated phases). These three slightly cooler sub-ages may be of the same type as the index horizons i, k and l in the Pacific cores. The first real cool time below Stage 4 is the isotopic minimum at about 17 m in the core. The isotopic ratio is here about the same size as in the last phase of the Riss (Riss II). It is very unlikely that such a climatic deterioration (isotopic ratio change) occurred within the Great Interglacial and therefore it seems very probable that this corresponds to the last phase of the Mindel Glaciation (Mindel II). Eighty-five per cent of Globorotalia truncatulinoides showed left coiling at 1798-1799-5 cm. The paleotemperature curve at the bottom of the core reveals another cool phase with a temperate phase (?; interstadial ?) between them. This agrees with my earlier interpretation where the Mindel is clearly divided into two phases. As stated above this is also assumed based on the evidence of continental deposits.

Cores 255-258. All have a limy upper division of about 40 per cent CaCOa. The subjacent horizons are not or only slightly calcareous. This is due to a decrease in the amounts of carbonates dissolved during the Postglacial. The cores from the northern North Atlantic have been correlated on the basis of the carbonate curves and the foraminiferal analyses, and more especially on the trends of the Globorotalia truncatulinoides curves (Fig. 9). Furthermore the paleotemperature curve of Core 280 was considered. Cores 292-296. Maxima in the frequency of Globigerina pachyderrna coincide fairly well with the lime minima. Core 297. The high carbonate content recorded in this core is abnormal

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for such a depth o f 5980 m. The a u t h o r considers it therefore likely that the core consists of the redeposited matter. Core 299. The foraminiferal analyses, carried out by ANNIE ANDERSSON, indicate " w a r m " faunas in the two uppermost samples and between 335 and 485 c m ; at other levels there are " c o o l e r " assemblages. C o m p a r i n g this statement with the content o f grains > 6 5 ~z, we find that the minima in the occurrence of non-calcareous matter coincide with the horizons containing " w a r m " foraminifera and the maxima with " c o l d " stages. The presence o f only a very few benthonic foraminifera makes it more likely that the layers with high a m o u n t s o f detrital grains are formed during other processes than by turbidity currents (ice-rafted and eolian matter ?).

A.CKNOWLEDGEMENTS The a u t h o r is greatly indebted to Drs W. D a n s g a a r d and V. Schytt for discussions and suggestions a b o u t the 018 content in precipitation and glaciers and a b o u t glaciological problems. My thanks are due to Professor B. K u l l e n berg, Professor N. Jerlov, and Dr. L. Lybeck for stimulating discussions about hydrographical problems. F o r valuable suggestions, advice, data, and information the a u t h o r is m u c h indebted to Miss A. Andersson and the Professors R. F. Flint, H. Flohn, R. Gonfiantini, G. Hoppe, E. B. Kraus and I. I. Schell. REFERENCES A'RRHENIUS,G. (1952) Sediment cores from the East Pacific. Rept. Swedish Deep-Sea Exp.

1947-1948, 5 (1), 227 pp. B~, A.. W. H. (1959) Fluctuations in the faunal boundary between temperature and subtropical planktonic foraminifera in the North A.tlantic. Preprints lnt. Oceanogr. Congr. 1959, AAAS, Washington, 134-136. BENTLEY,C. R., CRARY,A'. P., OSTENSO,N. A. and TmEL, E. C. (1960) Structure of West Antarctica, Science 131 (3401), 131-136. BROECKER,W. S. (1961) Radio-carbon dating of Late Quaternary deposits, S. Louisiana: A. Discussion. Bull. Geol. Soc. Amer. 72 (1), 159-161. BROECKER,W. S., EWING, M. and HEEZEN,B. C. (1960) Evidence for an abrupt change in climate close to 11,000 years ago. Amer. J. Sci. 258, 429--448. BtrrZER, K. W. (1958) Quaternary stratigraphy and climate in the Near East. Bonner Geographische Abh. (24). CO~ENS, C. W. (1937) Die Sediments des ,~quatorialen Atlantischen Ozeans. Wiss. Ergebn. Deutschen Atl. Exped. Meteor 1925-1927 3 (3), 298 pp. (Mit Beitr~igen yon Wolfgang Schott, Viktor Leinz und Otto-Ernst Radczewski). DANSGA^RD,W. (1961) The isotopic composition of natural waters, with special reference to the Greenland ice cap. Medd. om Gronland 165 (2), 120 pp. DONN, W. L., FARRAND,W. and EWING, M. (1962) Pleistocene ice volumes and sea-level lowering. J. Geol., 70, 206-214. EMtLZANt,C. (1954) Depth habitats of some species of pelagic foraminifera as indicated by oxygen isotopic ratios. Amer. J. Sci. 252, 149-158. EMtLIANI,C. (1955a) Pleistocene temperature variations in the Mediterranean. Quaternaria 2, pp. 87-98.

EMILIANI,C. (1955b) Pleistocene temperatures. J. Geol. 63, pp. 538-578.

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EMIUANI, C. (1956) Oligocene and Miocene temperatures of the equatorial and subtropical Atlantic Ocean..L Geol. 64 (3), 281-288. EMILIANI, C, (1958) Paleotemperature analysis of Core 280 and Pleistocene correlations. J. Geol. 66, pp. 264-275. EMILIANI, C., MAYEDA,T. and SELLI, R. (1961) Paleotemperature analysis of the PlioPleistocene section at Le Castella, Calabria, Southern Italy. Ball. Geol. Soc. Am. 72, No, 5, 679-688. EPSTEIN, S. and MAYEDA,T. (1953) Variation of 018 content of waters from natural sources, Geochim. et Cosmochim. Acta 4, 213-244. ERICSON, D. B., EWING, M., WOLLIN, G. and HEEZEN, B. C, (1961) Atlantic deep-sea sediment cores. Geol. Soc. Amer. Bull. 72, 193-286. ERICSON, D. B. and WOLLIN, G. (1956) Correlation of six cores from the equatorial Atlantic and the Caribbean. Deep-Sea Research 3, 104-125. FAIRBRIDGE, R. W. (1961) Eustatic changes in sea level. In: Physics and Chemistry of the Earth, Pergamon Press, London, 4, 99-185. FARRANO,W. R. (1962) Postglacial uplift in North America. Amer. J. Sci. 260 (3), 181-199. Ft.ltaT, R. F. (1947) Glacial geology and the Pleistocene Epoch. Wiley, New York, 589 pp. FLINT, R. F. (1957) Glacial and Pleistocene geology. Wiley, New York, 553 pp. Ft.1NT, R. F. and BRANDTNER,F. (1961) Climatic changes since the last interglacial. Amer. J. Sci., 259 (5), 321-28. FRYE, J. C. (1962) Comparison between Pleistocene deep-sea temperatures and glacial and interglacial episodes. Bull. Geol. Soc. Amer. 73 (2), 263-266. HOLLIN, J. T. (1962) On the glacial history of Antarctica. J. Glaciology. 4 (32), 173-195. HOPPE, G. and LIUEQtrlST, G. H. (1956) Det sista nedisningsf6rloppet i Nordeuropa och dess meteorologiska bakgrund (The course of the last glaciation in northern Europe and its meteorological background). Ymer, Stockholm, ( 1), 43-74 (Summary in English). IVERSEN, J. (1954) The late-glacial flora of Denmark and its relation to climate and soil. Danmarks Geol. Undersagelse, (2), 80. IVERSEN, J. (1960) Problems of the early Postglacial forest development in Denmark. Danmarks Geol. Undersagelse (4), 4 (3). KoczY, F. F. (1956) Echo soundings. Rep. Swedish Deep-Sea Exped. 1947-1948, V. IV, fasc, II. KULLENSERG, B. (1952) On the salinity of the water contained in marine sediments. Medd. Oceanografiska lnst. GOteborg. 21. LIIAEQUIST, G. (1963) Antarktisk meteorologi. In: Svensk Naturvetenskap, Stockholm, 22--40. LORIL'S, C. (1962) Contribution to the knowledge of the Antarctic ice sheet: A synthesis of gtaciological measurements in Terre Ad61ie. J. Glaciology, 4 (31), 79-92. MXLTHERS, V, (1950) Die Gliederung und Verbreitung der scandinavischen Vereisungen in Nordwesteuropa. Geol. FSren. FOrh, Stockholm, 72, 257-268. NILSSON, T. (1961) Ein neues Standardpollendiagramm aus Bj/irsj6holmssj6n in Schonen. Lunds Univ. ~rsskr. (n.f) (2), 56, 18, 34 pp. OLAt;SSON, E. (1957) Das Moor Roshultsmyren. Eine geologische, botanische und hydrolische Studie in einem SiJdwestschwedischen Moor mit exzentrisch gew61bten Mooselementen. Lunds Univ,, ~rssr. (n.f.), (53) (12), 72 pp, OLALrSSON, E. (1960a) Description of sediment cores from Central and Western Pacific with the adjacent Indonesian region. Rept. Swedish Deep.Sea Exped., 1947--48, 6 (5). OLAb'SSOY,E. (1960b) Description of sediment cores from the Indian Ocean. Rept. S,'edish Deep-Sea Exped., 1947-1948, 9 (2), 53-88. Ot.Aussoy, E. (1960c) Description of sediment cores from the Mediterranean and the Red Sea. Rept. Swedish Deep-Sea Exped., 1947-1948, 8 (3), 287-334. OLAUSSON, E. (1960d) Description of sediment cores from the North Atlantic. Rept. Swedish Deep-Sea Exped., 1947-1948, 7 (5), 229-286. OLAUSSOr<, E. (1961a) Studies of deep-sea cores. Sediment cores from the Mediterranean Sea and the Red Sea. Rept. Swedish Deep-Sea Exped., 1947-1948, 8 (4), 337-391.

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OLAUSSON, E. (1961b) Remarks on some Cenezoic core sequences from the Central Pacific. with a discussion of the role of Coccolithophoridae and foraminifera in carbonate deposition. G~teborgs Kungl. Vetenskaps- och Vetterhets-Samhiilles HandL, Sj~tte F/51jden, (B), 8 (10) (Medd. Oceanografiska Inst. Giiteborg, 29), 35 pp. OLAUSSON, E. (1961c) Remarks on Tertiary sequences of two cores from the Pacific. Bu//. Geol. Inst., Uppsala, 40, 299-303. PARKER, F. L. (1958) Eastern Mediterranean foraminifera. Rept. S, edish Deep-Sea Exped. 1947-1948, 5 (2). PHLEGER, F. B., PARKER, F. L. and PEmSON, J. F. (1953) North Atlantic Foraminifera. Rept. S;tedi h Deep-Sea Exped., 7 (1), 122 pp. ROBIN, G. de Q. (1964) Glaciology. Endeavour XXIII, 89, 102-107. ROSHOLT, J. N., EMIUANI, C., GEISS, J., KoczY, F. F. and WANGERSKY, P. J. (1961) Absolute dating of deep-sea cores by the Pa23VThZ3° method. J. Geology, 69, 162-185. ROSHOLT, J. N., EMIt.IANI, C., GEISS, J., KOCZY, F. F. and WANGERSKY, P. J. (1962) Pa TM, Th 2z° dating and O~8/O 1~ temperature analysis of core A 254-BR-C. J. Geophys. Res. 67 (7), 2907-2911. SCHO'rr, W. (1935) Die Foraminiferen in dem/i.quatorialen Tell des atlamischen Ozeans. Wiss. Ergeb. Deutschen Atl. Exped. Meteor 1925-1927, 3 (3), 43-134. SCHOTT, W. (1952) On the sequence of deposits in the equatorial Atlantic Ocean. G6teborgs Kungl. Vetenskaps- och Vitterhets-Samhiilles Handl., Sj/itte F61jden, (B), 6 (2) ( Medd. Oceanografiska Inst. G6teborg. 18), 15 pp. SCHY'rT, V. (1960) Snow and ice temperatures in Dronning Maud Land. Norwe,eian-

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