Structure and timing of the last deglaciation: Oxygen-isotope evidence

Quaternary Science Reviews, Vol. 4, pp. 5 9 - I [|8.1985.

(1277- 3791/85 $ll.(l(I + . 5( I C o p y r i g h t (~) 1985 P c r g a m o n Press Ltd.

Printcd in G r c a t Britain. All rights rcscr,,cd.

STRUCTURE AND TIMING OF THE LAST DEGLACIATION: OXYGENISOTOPE EVIDENCE l

A l a n C. Mix*'l" a n d W i l l i a m F. R u d d i m a n * * Department of Geological Sciences, Columbia University, and Lamont-Doherty Geological Observatory, Palisades, N Y 10964, U.S.A. "~Present address: College of Oceanography, Oregon State University, Corvallis, OR 97331,

U.S.A. Received 2 September 1985

Foraminiferal oxygen-isotope data from 24 tropical Atlantic sediment cores, constrained by 77 ~4C dates, are stacked to form a composite record of isotopic Termination 1..This record indicates that most of the isotopic transition at the end of the last ice age occurred between 14 ka BP and 6 ka BP. Minor isotopic expression of deglaciation is permitted as early as 16 ka BP, but the most rapid rate of change occurred between 14 ka BP and 12 ka BP. Three "steps" of maximum change are present. Although they are close to the statistical limits of detection in the composite record, the clear presence of the steps in individual records suggests that they are real. We estimate their timing at 14-12 ka BP (Termination l-a), 10-9 ka BP (Termination l-b), and 8-6 ka BP (Termination I-c). Centering of the termination near 11 ka BP is consistent with the "Milankovitch" hypothesis that high summer insolation caused deglaciation, In detail, however, maximum rates of change prior to the 11 ka BP insolation extreme, and the inferred steps require additional mechanisms controlling the tempo of glacial-interglacial climate change. Steps equivalent to those in 8~0 have not been detected in ice-margin retreat data. Steps in the isotopic transition, if real. may record thinning of the ice sheets not accompanied by loss of area. Alternation between near-equilibrium and near-stagnant ice-sheet profiles during deglaciation is hypothesized, perhaps due to calving and unstable "down-draw" of the ice sheets followed by partial re-equilibration. Significant problems remain. The effects of temperature on the isotope record are only partially constrained. Presently available data allow only semi-quantitative intercalibration of ice volume, sea level, and isotopic estimates of glaciation.

I N T I ~ O D U CTI ON From earh' studies of the non-isotopic deep-sea record, the Pleistocene-Holocene transition was interpreted to have occurred rapidly (within a few thousand years), about 11 ka BP (Ericson et al.. 1956: Broecker et al.. 1960). More recenth', following Shackleton (1967). many paleoclimatologists have used the marine oxygen-isotope record as tin icevolume proxy to constrain the timing o | glaciation on hind. The existence of a rapid isotopic ~Lamont D o h c r ~

Geological O b s c r v a t o r 3 Contril~ution No. 3S~5.

59

60

A.C. Mix and W.F. Ruddiman

'termination' (named by Broecker and Van Donk. 1970) ending the last ice age 11 ka BP has until recently been widely accepted in the paleoclimatic literature. The timing of deglaciation has important consequences regarding the mechanisms of climate change. Milankovitch (1941) hypothesized that the end of the last ice age was caused by slight changes in the earth's orbit. Because of these orbital changes, caloric summer insolation peaked in the northern hemisphere 11 ka BP (A. Berger, 1978). If net ablation rate responded directly to insolation, the maximum rate of melting should also have occurred near 11 ka BP (Imbrie and Imbrie, 1980). Indeed, the 11 ka BP "rapid melt" chronology of Broecker et al. (1960) was used by Broecker (1966) in support of the Milankovitch climate theory, which was out of favor at the time. A challenge to the traditional view of deglaciation came from Duplessy et al. (1981), who analyzed ~SO and radiocarbon in deep-sea sediment cores from the North Atlantic. They concluded that deglaciation occurred in two discrete steps; the first (Termination I-a) between 16 and 13 ka BP and the second (Termination l-b) between 10 and 8 ka BP. This interpretation put a pause in deglaciation exactly at the time that Ericson et al. (1956) and Broecker et al. (1960) inferred the most rapid melting. If high rates of deglaciation occurred as early as 16 ka BP, a direct link between ablation rate and insolation, as required by the Milankovitch climate theory, is unlikely. Alternatively, additional climate mechanisms may exist, either as feedbacks within the climate system, or as external forcing unrelated to orbital variation. If rising insolation acted as the sole trigger for deglaciation, the chronology of Duplessy et al. (1981) implies enormous sensitivity of the earth's climate system to small perturbations, at least when ice sheets are large. In order to minimize the smoothing effects of bioturbation, Duplessy et al. (1981) analyzed high sedimentation rate cores from sediment drifts and continental margins in the North Atlantic. These areas have deposition rates > 10 cm/ka (higher than typical pelagic rates of 2 to 4 cm/ka) because fine-grained sediments reworked by bottom currents and/or down-slope transport accumulate there. This is not a serious problem if only contemporaneous materials are reworked, because the contaminants would have a radiocarbon age near zero at the time of deposition. If older sediments are included, however, the radiocarbon ages would be too old. High-latitude North Atlantic sites are also contaminated with ice-rafted carbonate derived from pre-Pleistocene continental rocks (Bramlette and Bradley. 1941). Although Duplessy et ul. (1981) attempted to minimize this contamination by dating only the >44 p,m fraction after removing ice-ra~ted detritus larger than 150 ~m, the ~4C dates on these samples may be too old. In addition, the high-latitude North Atlantic is prone to large local temperature and meltwater effects (Ruddiman and McIntyre, 1981a), which may affect ~'~O signals directly and mask the global ice volume effect. Given the implications of the deglacial chronology for mechanisms of climate change, we test the chronology of Duplessy et ul. (1981). Our strategy is to analyze ~xO in cores containing normal pelagic sediments in order to avoid contamination of ~4C dates with reworked material. Many records are "stacked' or superimposed, to make a composite. This improves the signal-to-noise ratio in the data, because "local" random errors cancel each other, and signal common to all cores is enhanced. Stacking has the disadvantage that

61

Structure and Timing of the Last Deglaciation

inevitable slight miscorrelations cause artificial smoothing of high-frequency structures that are real. We studied tropical Atlantic sediments for four reasons. First, relatively high pelagic sedimentation rates in this area (typically 3-5 cm/ka) allow high-resolution studies and minimize (but do not eliminate) the effects of bioturbation. Second, abundant and wellpreserved carbonate sediments minimize potential dissolution problems. Third, the area is far enough away from meltwater influx that biotic barren zones do not exist. Although W.H. Berger (1978) proposed that meltwater affected planktonic ~ ( ) here (and throughout the global oceans), Jones and Ruddiman (1982) disagreed. Fourth, sea-surface temperature variability is relatively small in the tropical Atlantic (Mclntyre et al., 1981). Thus, possible temperature effects on ~JSo are minimized. In any case, the possible effects of bioturbation, carbonate dissolution, and temperature are evaluated, and an attempt is made to interpret the structure and timing of the last deglaciation as recorded by ~JSO.

METHODS Hand-picked samples of sized planktonic and (where present) benthonic foraminifera from 24 cores (locations in Fig. 1 and Table 1) were analyzed for ~180 on a VG-Isogas 903E mass spectrometer at Lamont-Doherty Geological Observatory using standard techniques (Shackleton and Opdyke, 1973). In all cores analyzed for ~]80, planktonic foraminiferal species were counted (taxonomy of Kipp, 1976; Be, 1977) and weight-percent calcium carbonate was measured (techniques of Hulsemann, 1966; Jones et al., 1983). Although these data will be discussed fully in other papers, we use them here to assess the degree to which radiocarbon and 8~80 are affected by bioturbation, carbonate dissolution, and temperature. In order to compare and stack the ~ O records, age models were generated for each core by assuming constant sedimentation rates between '4C dates (Table 2). Radiocarbon -89 , ~

- "/8 ......

-68

• . . . . . . . . . .

-58 • . . . . . . . .

-I;8

• . . . . . . . . .

-$e

• ......

II.w

•

-20 . . . . . . . . .

- !8

e . . . . . . . . .

8

• ....

V32-08

18

.'.'.:"l'.:'.'.:.":':

. . . . .

'..!~.. ". • ' "''.'/"

>- -:..: ..~...'~ ..:

.i"i'."'."

28 :

•

/:~.:....

v3o-nl

lo ~ •.' . -'

~ ~

~.::..'.'~V15-168 -

"

ilIR CIBIV3; ~8: lV 2 2 -mlm82~j _ n ~ V 2 9 1~.':.. _ :1

.' :'.'.:..':...'~..

-

22-1rr

=Re~-le

\: ~:1

-]8 -' oe6

-78

-68

-59

-g9

-$8

-re

- 18

FIG. 1. Core locations (also sec Tablc 1).

9

115

se

02

A.C. Mix and W.F. Ruddiman

T A B L E I. Core locations

CORE

LATITUDE

LONGITUDE

DEPTH

EN66-10G

6°39'N

21°54'W

3527

M-12392

25°10'N

16°51'W

2575

RC9-49

ll°ll°N

58°36'W

1851

RC13-189

I°52'N

30°00'W

3233

RC24-01

0°34'N

13°21'W

3850

RC24-07

i°21'S

II°55'W

3899

RC24-16

5002'8

I0°12'W

3543

V15-168

O°12'N

39°54'W

4219

V22-38

9°33'S

34°15'W

3797

V22-177

7°45'S

14°37'W.

3290

V22-182

0°33'S

17°16'W

3776

V22-222

28°56'N

43°39'W

3197

V23-II0

17°38'N

45°52'W

3746

V25-56

3°33'S

35°14'W

3512

V25-59

I°22'N

33°29'W

3824

V25-60

3°17°N

34°50'W

3749

V25-75

8°35'N

53°i0'W

2743

V29-144

0°12'S

6°03'E

2685

V30-36

5°21'N

27°19'W

4245

V30-40

0°12'S

23°09'W

3706

V30-AIK

0°13'N

23°04'W

3874

V30-49

18°26'N

21°05'W

3093

V30-51K

19°52'N

19°55'W

3409

V32-08

34°47'N

32°25'W

3252

63

Structure and Timing of the Last Deglaciation

TABLE 2. Chronologies (* = '4C date corrected for 6'3C - 400 yr) (All dates are on total CaCO3 unless otherwise indicated) CORE

DEPTH (cm)

EN66-10

M-12392

AGE + ERROR

REFERENCE/COMMENT

0

1500

Core top

8.5

6000

Event I.I

26.5

14,000

Event 2.02

38

24.000

Event 9,0

0-7 25 105-110 194-200 232 306-312

710 (± 230) 6000 15,145 (± 400)

'19,000 (+i170,-I020) 24,000 28,250 (+1200,-1060)

14C organic Koopmann

(1979)

Event I.I I~C organic Koopmann (1979) 14C organic Koopmann (1979) Event 3.0 '4C organic Koopmann

(1979)

474

53.000

0

1500

Core top

13.5-16.5

6770 (±350)

14C B~ et aZ.

(1976)

17-22

7540 (±400)

14C B~ et a~,

(1976)

23-27

8260 (±600)

*'C Ruddiman and Mix (~n pre88)

32-37

9900 (±525)

14C Ruddiman and Mix (~n press)

71-75

14,160 (±490)

14C this study * GX-IOIII

23.920 (+3000)

a4~ ~

RC9-49

101-109 RC13-189

0

et oZ.

(~976)

1500

core

21-23

8540 (±150)

14C this study * GX-9532

31-33

11,240 (±450)

*'C this study * GX-9533

37-39

12,630 (±490)

14C this study * GX-IOII2

60-63

16,160 (±600)

14C this study * GX-10113

85

24,000

.Event 3.0

210

59,000

Event 4.0

~5,760 (+760)

14C this study (discarded) ,

46-4~

top

64

A.C. Mix and W.F. Ruddiman

TABLE 2. (continued) CORE

DEPTH

RC24-01

core top

0-5

3460 (±190)

14C this study * GX-10114

12.5-15.5

7235 (±300)

14C this study * GX-10115

21-24

I0,000 (±370)

14C this study * GX-10116

33-36

12,690 (±405)

14C this study * GX-IOII7

49-52

15,680 (±720)

14C this study * GX-10118

76.5-79.5

18,520 (±820)

14C this study * GX-10119

105

24.000

Event 3.0

0

1500

core top

31-35

7950 (±320)

14C this study * GX-10120

71-75

15,270 (±720)

14C this study * GX-IOI21

115

17,5OO

Event 2.1

175

24,000

~veBt ~,O

0

1500

core top

18

6000

Event I.i

60

14,OOO

Event 2.02

96

24.000

Event 3.0

0

1500

core top

23-29

7425 (±275)

t4C B~ e~ u~.(1976)

86-94

12,050 (±490)

14C this study * GX-9537

17,740 (±2500)

14C B~ e~ a~.(1976)

24,000

Event 3.0

15.7OO (+20007

14C B~ et a~.

RC24-16

V15-168

142.5-147.5 165 67-78 V22-38

REFERENCE/COMMENT

1500

RC2A-07

0

AGE + ERROR

0

1500

core top

17.5

6000

Event i.i

32.5

14,000

Event 2.02

45

24.000

Event 3.0

(197~) discarde d

Structure and Timing of the Last Deglaciation

65

TABLE 2. (continued) CORE V22-177

DEPTH

AGE + ERROR

0

1500

core top

13-16

5025 (±220)

x4C this study * GX-9538

21-23

7940 (±250)

x4C this study * GX-10122

31-34

11,130 (±440)

a4C this study * GX-9539

44-46

12,650 (±525)

14C this study * GX-9540

62-64

17,610 (±710)

14C this study * GX-10123

~4,000

Event 3.0

83 V22-182

V23-110

V25-56

V25-59

COMMENT/REFERENCE

0

1500

core top

34-36

9740 (±2~0)

*4C this study * GX-9541

45-49

12,710 (±190)

14C this study

54-56

14,950 (±720)

'4C this study * GX-10124

85-89

21,000 (+770)

14 C this study

0

1500

core top

6

6000

Event i.i

18

14,000

Event 2.02

$8

~4,000

~v~nt 3,0

QC-216

OC-215

0-3

2750 (±170)

14C this study * GX-10125

26-28

8450 (±355)

*4C this study * GX-10126

46-49

11,540 (±350)

14C this study * GX-10127

62-65

14,330 (±600)

a4C this study * GX-10128

120

24,000

Event 3.0

260

59,000

~v~nt 4,0

0

1500

core top

15-18

6490 (±130)

14C B4 et aZ.(1976)

33-38

11,950 (±280)

14C B4 e~ aZ.(1976)

45-48

15,485 (±680)

*4C-this study * GX-10129

64-68

17,240 (±530)

'4C B4 e~ uZ.(1976)

76-79

21,790 (±680)

a4C this study * GX-IOI30

$~.450 (+1350)

14C B4 et ~Z.(~97~)

101-~Q8

66

A.C. Mix and W.F. Ruddiman

TABLE 2. (continued) CORE

DEPTH

V25-60

0

core top

9-11

4350 (±190)

14C this study * GX-IOI31

22-24

9025 (±260)

'4C this study * GX-IOI32

28-32

10,210 (±370)

14C this study * GX-10133

41-44

15,610 (±560)

14C this study * GX-IOI34

~7,~

~4,000

Event 3,0

0

1500

core top

51-59

8400 (±280)

14C this study * GX-8621

100-104

14,685 (±570)

'4C this study * GX-8622

150

~4,ooo

~ven~ ~,0

o

1500

core top

30

6000

Event i.i

70

14,000

Event 2.02

ii 0

~4,000

Event ~,0

0

1500

core top

11-14

7070 (±175)

14C this study * GX-8266

19-21

10,320 (±330)

t4C this study * GX-10135

36-40

16,500 (±880)

14C this study * GX-10136

46

~,000

~veDt $,0

0

1500

core top

3.5-5.5

2010 (±170)

14C this study * GX-9273

21.5-23.5

7685 (±300)

14C this study * GX-9274

29-31

9980 (±325)

14C this study * GX-IOI37

43-46

12,420 (±400)

14C this study * GX-IOI38

91.5

24,000

Event 3.0

183

59,000

Event 4.0

12.A05 (+735)

14C this study (discarded)

v29-144

V30-40

COMMENT/REFERENCE

1500

V25-75

V30-36

AGE + ERROR

54.5-56.5

67

Structure and Timing of the Last Deglaciation TABLE 2. (continued) CORE V30-41

v30-49

V30-51

DEPTH 0

AGE IERROR

.. COMMENT/REFERENCE

1500

core top

2.5-4

1700 (±120)

14C Jones & Ruddiman (1982)

4-5

2020 (±200)

*4C Jones & Ruddiman (1982)

i0-Ii

3350 (±230)

14C Jones & Ruddiman

16-17

4425 (±160)

14C this study * GX-8271

18-19

7490 (±350)

14C Jones & Ruddiman (1982)

24-26

1 0 , 4 7 0 (±360)

x4C J o n e s & Ruddiman (1982)

31-32

1 2 , 2 1 0 (~610)

14C J o n e s & Ruddiman (1982)

38-39

1 4 , 8 3 0 (±840)

14CJones

44-46

1 8 , 6 8 0 (±920)

*4C J o n e s & Ruddiman (1982)

55

24,000

~v@nt $,0

0

1500

core top

17-19

6100 (±220)

14C this study * GX-10139

42-44

11,780 (±440)

14C this study * GX-IOI40

65-68

15,000 (±680)

14C this study * GX-10141

88

19,000

Event 2.2

102

24,000

Event 3 0

224

~9,000

~v@nt 4,o

(1982)

& Ruddtman (1982)

1500

c o r e top

4.5-5.5

1770 (±160)

14C this study(<63pm)

* GX-10142

4.5-5.5

1765

14C this study(>63~m)

* GX-8274

0

±135)

5

1768 (±210)

CaCO3-weighted mean

24.5-25.5

8310 (±255)

14C this study(<63pm)

* GX-10143

24.5-25.5

7825 (+195,-170)

14C this study(>63pm)

* GX-8275

25

8083 (±320)

CaCO 3 weighted mean

37.5-38.5

12,220 (±470)

'4C this study(<63pm)

* GX-10144

37~5-38.5

9200 (±310)

14C this study(>63pm)

* GX-8276

38

10,743 (~560)

CaCO 3 weighted mean

54.5-55.5

19,420 (±430)

14C this study(<63pm)

54.5-55.5

16,025 (+375,-310)

55

18.578 (+37Q)

14C this study(>63~m) CaCOl weighted mean

* GX-I0145 * GX-8277

A.C. Mix and W.F. Ruddiman

68

TABLE 2. (continued) CORE

DEPTH

AGE ! E R R O R

COMMENT/REFERENCE

V32-08

0

1500

core top

12

6000

Event I.i

60

14,000

Event 2.02

96

~,QQQ

Event 3.0

dates were obtained from - 1 0 g samples of bulk carbonate (unless otherwise indicated) and were calculated using the Libby half-life of 5570 years. Although not all cores have dates at their tops, in those that do, mixed-layer ages range from 700-3500 years. Core-top ages are not held constant through an assumed mixed-layer thickness, because piston cores commonly lose a few centimeters from their tops during coring. It is rarely possible to know what has been lost. For cores lacking mixed-layer ~4C data, we assume a top age of 1500 years. Although the assumption of a 1500-year core-top age is probably not correct for all the cores used here, it does not affect interpretation of structure and timing of the deglaciation if age control is present below the mixed layer and above (younger than) the deglaciation. If no dates exist >20 ka BP, the isotopic stage 2/3 boundary is assumed to be 24 ka BP (Imbrie et al., 1984). Other correlation points in specific cores are discussed below and listed in Table 2. Before stacking multiple records into a composite, all data are expressed as anomalies relative to the modern by subtracting the mean Late Holocene (0-5 ka BP) 8~80 value in each core from all down-core data points. The data series are then stacked by merging at the ages assigned by the age models. Smoothing is done .with a 2000-year Gaussian convolution filter. This filter is better than a 'boxcar' moving average filter because no phase inversion occurs in the high frequencies (e.g. Hamming, 1983). As the filter operates on data sampled at constant time intervals, the stack must be interpolated before smoothing. Various interpolation schemes were tested (linear, cubic spline) to see if the final result is dependent on interpolation. It is not, with the exception that a broad interpolation step can introduce aliasing. Results shown here were interpolated linearly at an interval that gives roughly the same number of interpolated points as original data points. See Davis (1973) for examples of interpolation routines. The error envelope of + one standard deviation (cr) is calculated by:

or, =

V ~ j=i--N i+E N( Y j 1

-

y/)2

(1)

1/n

where Yj is the interpolated data (at each time step), ?j is the smoothed estimate, and n = 2N + 1 is the number of filter elements. Our stacking and smoothing procedures differ somewhat from those used by other authors. For example, Berger et al. (1977) stacked 8)~O data from six Pacific cores in the depth domain, after adjusting depths by correlating to one well-dated reference core.

69

Structure and Timing of the Last Deglaciation

Depth-domain data were interpolated to the depths sampled in the reference core, and the interpolated values were averaged. An age model for the reference core was then used for the stacked record. This procedure has the disadvantage of placing a great deal of trust in the reference core. If the reference is anomalous, the stack will reflect that anomaly. In contrast, our stacking procedure minimizes assumptions about sedimentation rates, needs no arbitrary reference core, and avoids artificial smoothing and aliasing induced by interpolating before stacking.

RESULTS We have made 976 ~SO analyses from 23 tropical Atlantic cores. In addition, we have used published g180 values from benthonic foraminifera in core M12392 (Shackleton, 1977). Age models (Fig. 2) in the 24 cores are constrained by 77 14C dates (Table 2). Of these, 22 dates are from the literature, and 55 are new dates obtained for this study. Some of the dates were corrected for ~3C. This removes the effects of isotopic fractionation in a sample by adjusting its ~4C activity for the ~3C offset between the sample and the oxalic acid 14C standard (e.g. Faure, 1977). For typical marine carbonate samples (~13C = 1-2%o vs PDB) this adjustment adds approximately 400 years to the measured age. Another adjustment to measured radiocarbon dates must be made for the apparent age of surface ocean water (the reservoir age). The average reservoir age for the modern (preanthropogenic) tropical Atlantic is about 400 years (Stuiver and Ostlund, 1980). In the absence of data to constrain variations in this adjustment, we assume that it remains constant and subtract it from each ~]3C-corrected date. Because the reservoir age roughly cancels the original ~3C correction in the tropical Atlantic, no adjustments were made to dates that were not ~13C-corrected.

(a)

H-12392

..

DEPTH

i

V15-168

~

*~

•

,

"

*

DEPTH

."

,

'

~

~,

.

'

'

'

DEPTH

~

i

.

.

' i'

V29-144 ,

•

'

1

~D

2. 0

"I

" o i

tO.O

--

20.0

30.

\

0

i

RC24-07

DEPTH

:

....

t

J

i

V2S-75

:

;

DEPTH

;

;

....

FIG. 2. Age-depth diagrams for all cores used in this paper. Boxes indicate correlation points, all other control is radiocarbon, shown with error bars (also see Table 2). (a) Average sedimentation rates >5 cm/ka. (b) Average sedimentation rates 3-5 cm/ka. (c) Average sedimentation rates 1.5-3 cm/ka.

70

A.C. Mix and W.F, Ruddiman

I

1 +0.

1

1

I

I

0

I0.

"-

:

:

:

120.

:

0,. I,i,I I~

s]

41-

il0°

I

,

,] e

LlO.

I

,,o'P

40.

120. 40.

4D O I Psi

~

IL uJ D

IO,

Ii

a,

! !

120, 40,

oo

X k,0,,. kiJ It) !

O0,

~

pep

O0,

O,

"r~,-

O.

•

-r. i

,0

120.

120.

O.

~

,0

qP

:

oo. 0.

an"

120.

,',~'

.~

,,,"

O.

o,.

+4I

~ oc

O0.

[.I.

O.

0.

i,,o.

120.

,0

~

L20.

,0 J

sjj"

!

~

,....,.'~ /

~

,,,-~ 0. +

I

d,

>~

X l-.O.

• e"IP

L,~O.

:

co.

"

e0.

I

0.

~

40,

T. I-O. IlJ 40. *"'*

I:PO.

os

o0

I

00.

~ n-

O0. 0.

O. x

~3

..rt-.. 40.

1:3

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40.

U0 ItS

i

RGE

(]0O0-'lPI

OPI

c +-

40. ~;3

'

II0.

! I

40. 140

,,,it.

1 'I"

,m

120.

~

]]

00. 0o

+0

o

o

o ROE

o

•

(1000-TR

°

o 8PI

! ~1 •

p,

71

Structure and Timing of the Last Deglaciation

(c)

[N66-10

DEPTH

V23-110 DEPTH

~

o

,=

o

0.0

°

V30-36 DEPTH

*

o

o

o

o

,

o

o

o

o

*

i

o

o

!

|

o

*

m

0

LO.C o !

--

20.0

\ )0.0 P o

o

Q

V22-38 DEPTH . . . .

o

o

o

o

o

V25-60 DEPTH . . . .

FIG. 2. (continued).

Figure 3 illustrates the ~]So data plotted versus age and grouped by mean sedimentation rate as in Fig. 2. Four different composite 8]So records based on these data are shown in Fig. 4. The first, Fig. 4(a), contains 326 planktonic data points from 12 cores (Table 3) containing four or more ]4C dates each (63 tqtal). The mean sedimentation rate of this stack is 4.2 cm/ka. If viewed as an ice-volume proxy, this 'best dated' stack suggests that full glacial conditions existed until 14 ka BP, and that full interglacial conditions were reached by 6 ka BP. Within the transition period, variations in the rate of isotopic change are present. Maximum rates of change occur at 13 ka BP (just at the end of Termination 1-a of Duplessy et al., 1981), 9.5 ka BP (equivalent to Termination 1-b of Duplessy et al., 1981), and at 7 ka BP (here referred to tentatively as Termination l-c). The error envelope surrounding this structure, however, permits the interpretation of a constant rate of change between 14 ka BP and 6 ka BP. For the second stack, in Fig. 4(b), an additional 6 cores with less than four ~4C dates were included (Table 3). The total number of data points (633) is roughly double that of the first stack. Only seven additional ~4C dates were available, however, bringing, the total to 70. If radiocarbon dates were not available, age models for the additional cores were augmented by fixing the last 'full glacial' point (not necessarily the maximum) at 14 ka BP and the first 'full interglacial' point at 6 ka BP. Because of this, the dating errors in this stack are larger than in the first stack. The mean sedimentation rate of the second stack is 4.6 cm/ka. In this composite [Fig. 4(b)], some isotopic expression of deglaciation is permitted starting as early as 16 ka BP, but Termination 1-a remains at 13 ka BP. Termination l-b, at 10 ka BP, is slightly older than in the first stack, and Termination l-c, at 7 ka BP is similar to that of the first stack. Again, however, the error envelope allows a constant rate of ~]So change between 14 and 6 ka BP. Stacks one and two are listed in Table 4. The third stack, in Fig. 4(c), is composed of data from two species of benthic foraminifera, Cibicides wuellerstorfi and Uv!gerina peregrina in seven cores (Table 3). For cores in which both species were analyzed, species offsets were removed by adding 0.64 to

72

A.C. Mix and W.F. Ruddiman

+m ,?

- 0 . SO

,

OC l,bl I-.., 0

4

i

• *0. SO

@

o

a,,, ILl I""

m

i

~= E

I:1

2,00

=+

-t.TS I m

d

:1,,

2.00 4,

-~. 00 I,J In

O. ? S

d -2,00

0. H

0+ :)

'*+? 3.00

"~t~

O',-' O. SO

II0 +

3.00

I m

"JrI,3[ ILl n~*

S. S0

~E

3. O0 I,J

~>~

,,sip

S0 SO

IiJ GO

@

+ P-

I 3C

S. S0

i

o (:l

i o

o AGE

IIO00*TR

i

i o

~ OPJ

i

i o

~

,r-

73

Structure and Timing of the Last Deglaciation

~.4~

T

:

:

:

:

*l.Oe

IC uJ

+,k

* 0 . S0

*l,.1O

.

.

.

.

.

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d

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m q. o

i lid

i

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.-*

O.-

~1

r~

C

~

d

*~

c,J

I..00 -,1. 00

b'qL

d

*.J ¢C

0/

'qP

0. ?S

d ° 0 ° SO

-l.?5

m I ~

r~ Oo n

0.S0

- t . "INI,,

S.00,

MJ in 0)

:~.00

,4

2. ?S

r,,.

O. ?S

u') ~n

L$O

klJ OD

2. ?S,

I/')

d lid

-L.?S

kz m i

c~ n7 S. ~ l i o

I

!

0

O

!

0

RGI[ ( | 0 0 0 - ¥ R 6P)

o

I

0

O'?lo

:

.a

i

.a ~r~

:

i

~

tlOOO-TR SPJ

:

~m

sl

74

A . C , Mix and W.F. Ruddiman

!

i

i

i

1

- L, 7S

J

h 4'

"~'

i¢

o

-I,. ?S i rn

d~d

t.J 4Z'

O. 7 $ ID

-LTS

d B"

O. 7S ~. o . , , Y

- i . TS

(ID I U')

vl

O. ?S

-I.75

CO ~'! C] ca

t.I *.J O,

v

ul

L

?S

u~

-(I. 76

,dO "

O, 7S

Z C3

-L.?S

!

tO

@

w

Z

1. ?6 i

,L O0 ..I

aD In Z Ill

23 Z

O, 75

J o

"T c,o ¢10 z

:

4. SO °

: cl

:

: o

o

o

;IGIE

tLO00-~

. 8P~

:

: It

o

75

Structure and Timing of the Last Deglaciation

(a)

(b) 6too - I-Iol~ne

2-ka

t

dt~O -

Smooth, t t o

it

Holoceem

2--ka Smooth. t I o

2

ii

./

~

2" o; .$"

--

lC

e* .-* - - -

lt)

lb

!

oJ

....

(d)

(c) 6180 -- Ho]ocene

6taO

2 - - k o Smooth. = 1 o

-

2 - k a Smooth. t 1 o

t

}

1

"::"

•

"~=- lC

1C

lb

t;

eD .~

FIG. 4. Stacked and smoothed oxygen-isotope data. On each plot, crosses are stacked data points, solid line is the composite record, smoothed with a 2000-yr gaussian filter. Dashed lines show + 1 cr error envelope. (a) STACK 1: Planktonic data only from best dated cores. (b) STACK 2: all planktonic data from cores with sedimentation rates >2.5 cm/ka. (c) STACK 3: benthonic data. (d) STACK 4: planktonic data from same cores as benthonic data (no planktonic data from M-12392). See Table 3 for specific cores included in each stack.

measured GtSO values from C. wuellerstorfi (Shackleton, 1973). In this stack, the mean sedimentation rate is 5.0 cm/ka. Steps are present on the termination, but they occur slightly earlier than in the best dated planktonic stack, 1-a at 14 ka BP, 1-b at 10 ka BP and 1-c at 8 ka BP. In addition, a possible fourth step is apparent between Terminations 1-a and 1-b. In part this may reflect the lower density of data (and thus a higher noise level) in this stack (195 G~80 analyses and 23 ]4C dates). In sensitivity tests with random groupings of cores, stacks generally did not stabilize until they contained about 300 analyses and 30 dates. To test for the significance of the older ages in the benthonic stack, a fourth stack, composed of planktonic data from the same cores as the benthonics excepting M-12392

76

A.C. Mix and W.F. Ruddiman

T A B L E 3. Composition o f t h e 8 ~ 0 stacks

STACK ONE: BEST DATED PLANKTONICS. CORE

SPECIES

SIZE

0-Sk 61aO

COMMENT/REFERENCE

RC9-49

G. 8acc~lifer

355-415~m

-1.70

This study (appendix I)

RC13-189

G. 8acculifer

355-415~m

-1.46

This study (appendix i)

RC24-01

N. d u t e r t r e i

355-415~m

-0.07

This study (appendix i)

V15-168

G. sacculifer

355-415~m

-1.73

This study (appendix i)

V22-177

G. sacculifer

415-500~m

-1.31

This study (appendix I)

V22-182

G. sacculifer

355-415~m

-I.ii

This study (appendix i)

V25-56

G. sacculifer

355-415~m

-1.77

This study (appendix I)

V25-59

G. sacculifer

415-500~m

-1.33

This study (appendix i)

V25-60

G. sacculifer

355-415~m

-1.43

This study (appendix 1)

V30-40

G, sacculifer

415-500~m

-1.24

This study (appendix I)

V30-41

G. sacculifer

250-415~m

-1.35

This study (appendix I)

V30-51

N. dutertrei

355-415~m

-0.ii

This study (appendix i)

(all G. 8acculifer are 'with sac')

n - 326. 63 DATES USED.

STACK TWO: ALL PLANKTONICS, SEDIMENTATION RATES >2.5 cm/ky All data from stack one, plus: CORE

SPECIES

SIZE

0-5k 6180

COMMENT/REFERENCE

RC24-07

N.dutertrei

355-415~m

-0.18

This study (appendix I)

RC24-16

N. dutertrei

355-415~m

-0.44

This study (appendix I)

V25-75

G. sacculifer

250-500~m

-1.64

This study (appendix i)

V29-144

G. sacculifer

355-415~m

-1.37

This study (appendix I)

V29-144

N. dutertrei

355-415~m

-0.12

This study (appendix I)

V30-49

G. saceulifer

355-415~m

-I.08

This study (appendix i)

V32-08

G. tuber (w)

250-355~m

-0.76

This study (appendix i)

(all G. 8aeculifer are 'with sac', G. ruber is 'white' ) n - 633.

70 DATES USED

Structure and Timing of the Last Deglaciation

77

TABLE 3. (continued) STACK THREE: ALL BENTHONICS. CORE

0-5k 6*80

SPECIES

EN66-10

C. ta~ellerstorfi

M-12392

U. peregrina, & C, wuelleratorfi

+0.64

COMMENT/REFERENCE

2.37

Curry and Lohmann (1983)

3.20

Shackleton (1977)

V25-59

C. wuellerstorfi

2.53

This study (appendix i)

V25-75

U. peregrina

3.00

This study (appendix I)

V29-144

U. peregrina

3.27

This study (appendix i)

V30-49

C. wuellerstorfi

2.29

This study (appendix i)

V30-51

C. wuellerstorfi

2.30

This study (appendix I) n - 195.

23 DATES USED

STACK FOUR: PLANKTONICS FROM SAME CORES AS BENTHONICS. CORE

SPECIES

SIZE

0-5k 6*80

COMMENT/REFERENCE

EN66-10

N. dugertrei

355-415~m

-0.35

This study (appendix i)

EN66-10

G. 8acculifer

355-415~m

-1.43

This study (appendix I)

V25-59

G. 8acculifer

415-500~m

-1.33

This study (appendix i)

V25-75

G. sacc~lifer

250-500~m

-1.64

This study (appendix I)

V29-144

G. sacculifer

355-415~m

-1.37

This study (appendix 1)

V29-144

N. dutertrei

355-415~m

-0.12

This study (appendix i)

V30-49

G. 8acculifer

355-415~m

-1.08

This study (appendix i)

V30-51

N. du~ertrei

355-415~m

-0.ii

This study (appendix I)

(all G. 8acc'ulifer are 'with sac')

n - 271. 19 DATES USED

(Table 3, 271 8180 analyses and 19 dates) is shown in Fig. 4(d). The isotopic transition in this planktonic stack is similar in timing to that in the benthic stack. This suggests that the apparent lead of the benthonic 8~SO transition relative to that of planktonic stacks one and two is an artifact of a few anomalously old dates in the cores composing the benthonic stack. Within individual cores, the benthonic lead is inconsistent. In V30-51 and V25-59, the benthonic ~lSO records lead those of the planktonics slightly [Fig. 3(b)]. In V30-49, benthic ~lSO lags planktonic ~lSO. Other studies have yielded similar inconsistencies. Sarnthein, Erlenkeusser et al. (1982) found a lag of G. sacculifer ~tSO behind benthonic ~ s O by as much as 1500 years in one core. Duplessy et al. (1981) measured a lead of one planktonic ~tSO record (G. bulloides) and a lag of another (N. pachyderrna) relative to the benthonic record in North Atlantic core CH731392. Given this lack of agreement, we

78

A.C. Mix and W.F. Ruddiman TABLE 4. 8180 Stacks STACK-I

STACK-2

STACK-I

STACK-I

STACK-2

STACK-2

61e0

6*sO

AGE (ka)

61s0

8XSo

12.5

1.23

1.27

23.5

1.51

1.53

0.02

13.0

1.39

1.39

24.0

1.38

1.39

0.02

0.00

13.5

1.61

1.55

24.5

1.29

1.29

3.0

0.05

0.02

14.0

1.67

1.55

25.0

1.28

1.22

3.5

0.02

0.00

14.5

1.67

1.58

25.5

1.20

1.23

4.0

0.02

0.02

15.0

1.67

1.65

26.0

1.27

1.33

4.5

0.08

0.04

15.5

1.70

1.70

26.5

1.31

1.32

5.0

0.09

0.02

16.0

1.68

1.73

27.0

1.30

1.34

5.5

0.05

0.06

16.5

1.67

1.70

27.5

1.25

1.40

6.0

0~Ii

0.08

17.0

1.68

1.65

28.0

1.26

1.36

6.5

0.21

0.19

17.5

1.65

1.68

28.5

1.20

1.24

7.0

0.26

0.27

18.0

1.58

1.59

29.0

1.21

1.25

7.5

0.41

0.38

18.5

1.56

1.54

29.5

1.20

1.35

8.0

0.50

0.46

19.0

1.57

1.53

8.5

0.51

0.49

19.5

1.59

1.51

9.0

0.64

0.57

20.0

1.60

1.60

9.5

0.73

0.66

20.5

1.56

1.6,2

i0.0

0.81

0.84

21.0

1.56

1.67

10.5

0.87

0.92

21.5

1.54

1.62

ii.0

0.94

0.98

22.0

1.53

1.64

11.5

1.01

1.03

22.5

1.56

1.60

12.0

1.14

1.14

23.0

1.56

1.60

AGE (ka3

6*sO

61SO

1.5

0.06

0.07

2.0

0.00

2.5

AGE (ka)

hesitate to make any conclusions regarding slight differences between the benthonic and planktonic ~1~O stacks. It is possible, however, that future high resolution data will reveal consistent patterns with implications for relative changes between deep- and surface-water temperatures and/or isotopic signatures.

DISCUSSION The four stacked ~1~;0 records are reproduced together in Fig. 5 to facilitate comparison. The total amplitude of all the planktonic and benthonic stacks is very similar

Structure and Timing of the Last Deglaciation

79

S t a c k summary 0.0

m

lO.O

0 0 0 I q<

20.0

.,

"0 ,-

J

I

30.0

FIG. 5. Summary of the four stacks in Fig. 4, plotted with a small offset between records to facilitate comparison. (a) STACK 1: Planktonic data only from best dated cores. (b) STACK 2: all planktonic data from cores with sedimentation rates >2.5 cm/ka. (c) STACK 3: benthonic data. (d) STACK 4: planktonic data from same cores as benthonic data (no planktonic data from M-12392).

(1.7 ___ 0.1%o). Following Shackleton (1967) the first-order similarity between the benthonic and planktonic stacks argues that the major signal is that of continental ice volume, as it is unlikely that tropical surface waters and deep waters would experience identical temperature and/or salinity fluctuations. This is addressed in more detail below. The stacks allow the possibility of steps in the isotopic termination. The error envelopes around these steps, however, permit the alternate interpretation that the isotopic termination was smooth. Further, the timing of the steps is not completely consistent in the various stacks. The clear presence of steps in many of the individual cores in Fig. 3, and in higher sedimentation rate cores (e.g. Duplessy et al., 1981), however, suggest to us that the steps are real. If they are real, our best estimate of their timing is that Termination 1-a .occurred between 14 and 12 ka BP, Termination 1-b occurred between 10 and 9 ka BP, and Termination 1-c occurred between 8 and 6 ka BP. Before we can be certain that our deglacial 6~O chronology is correct, we must evaluate potential problems induced by bioturbation, carbonate dissolution, and local temperature and salinity effects. In all cases, rather than studying hypothetical examples, real cases from the present data set are examined. Bioturbation

Bioturbation acts as a low-pass filter by smoothing, but also as a high-pass filter by introducing noise. Although, in theory, the data can be processed to remove the effects of smoothing, it is impossible to remove the random noise of sub-mixed layer burrows from a single core (Berger et al., 1979). Because burrows cannot always be recognized or avoided during sampling, the best way to minimize this effect is to average many records as in our stacking procedure.

80

A.C. Mix and W.F. Ruddiman

Smoothing, however, remains a problem. When coupled to abundance changes of the signal carrier, it can cause large depth offsets in tracers such as ~ 8 0 (Hutson, 1980). This is especially true in cores containing biotic 'barren' zones, such as those in the high-latitude North Atlantic (Ruddiman and Mclntyre, 1981a). Although these zones in North Atlantic sediments are n o t entirely barren, essentially all biogenic carbonate present was mixed in from adjacent sections. This can create false steps in a smooth transition, even in cores with relatively high sedimentation rates. To illustrate possible problems induced by abundance bioturbation in our tropical Atlantic data set, we have modeled bioturbation of foraminifera and bulk CaCO3 in a core used in the stacked records. Following Peng et al. (1979), mixing intensity is defined by a diffusivity of 120 cm2/ka within a 2 cm mixed layer. Below this layer, diffusivity decreases exponentially with a half-depth of 1 cm. In the absence of abundance change, this mixing model yields a core-top age of 1750 years for a 3 cm/ka sedimentation rate, consistent with our data. For comparison to simpler mixing models (e.g. Berger and Heath, 1968), this is similar to the effects of a well-mixed layer 6 cm thick, which produces a core-top age of 1700 years for similar input. Using this mixing model, we first illustrate possible dating errors in a core with large changes in carbonate content. Core V25-59 is typical for tropical Atlantic cores, having lower %CaCO3 in glacial than in interglacial time (Damuth, 1975). If abundant Holocene carbonate mixed downward and contaminated the glacial section, dates for the isotopic Termination 1 could be too young. Relative to the other cores used for our study, V25-59 contains large variations in %CaCO3 (from 45% to 85%) and has a low sedimentation rate (3 cm/ka), so it constitutes a worst-case test for this effect within our data set. Input to the bioturbation model is a hypothetical unmixed carbonate curve (Fig. 6, left, dashed line). Approximately 5cm lost from the top of V25-59 during coring is accounted for. The input curve is numerically mixed, yielding an output curve (Fig. 6, left, solid line). Iterative adjustments were made to the input curve until the output curve resembled the %CaCO3 data (Fig. 6, left, circles). The solution is not unique, because actual mixing intensity is not well known, and the degree of fit to the data is subjective. Activity of the ~4C 'tracer' is related to abundance bioturbation of the carbonate 'carrier' because low-carbonate sections are preferentially contaminated with calcite of different ages mixed from the high-carbonate sections. For the model input, the 14C activity of surface waters (and thus the calcite rain) is assumed constant. As carbonate accumulates through time, the radiocarbon within it decays. The dashed age line in Fig. 6 (right) gives the 14C ages that would have been measured if no bioturbation had occurred. The solid age line in'Fig. 6 (right) gives the lac ages after bioturbation. The differences between the two curves are generally less than 100 years, within the error of ~4C measurement. Thus, in the cores we have used, mixing of bulk carbonate has a negligible effect on the radiocarbon dates. This does not mean, however, that all effects of mixing on the deglacial ~ 8 0 chronology are negligible. Bioturbation of individual foraminiferal species can induce age offsets in the ~'~O signals they carry if the species abundance variations are different from those of total carbonate, which carries the ~4C age signature. To illustrate this effect, ~lSO signals from two planktonic species with different abundance histories were measured in core V25-59.

Structure and Timing of the Last Deglaciation V25-59

=+C AGE ( l O 0 0 - y r

%CoCo3

•1~

01

~

-,4

Oa

0

b

o

b

b

o

O.

.o o

BP)

__

10

o

o

.

o

b

o

.l=

.o 0 I

'

'

'

:

" i_~_~-2-1'

I. . . . . . . E~ I"11 "0

81

~. . . .

2

I

.

40.

"1"

L. . . . . .

1

n

o

;I} m

80.

3

120.

FIG. 6. Effects of bioturbation on bulk-carbonate radiocarbon dates in core V25o59. This core contains the largest variations of %CaCOa in the cores we used, and has a relatively low sedimentation rate (3 cm/ka). Thus, it constitutes a worst-case analysis for this effect within our data set. Left: dashed line gives hypothetical %CaCO3 model input. Solid line gives the smoothed output. Dots are measured data. Right: Dashed line gives hypothetical age-model input. Solid line gives the smoothed output. Despite large changes in %CaCO3, effects on radiocarbon ages are negligible.

As with the bulk carbonate example from the same core (V25-59) above, this example approximates a worst-case analysis for bioturbation within our data set. Globigerinoides sacculifer [Fig. 7(a), left] is most abundant in V25-59 in interglacial time, and Neogloboquadrina dutertrei [Fig. 7(b), left] is most abundant in glacial time. Using the age model and bioturbation intensity from Fig. 6, abundance input functions for these species (dashed lines) yield good model approximations (solid lines) of the measured abundance data (crosses). To model the ~ 8 0 tracer in these species, we must assume an input function, which is then mixed. Of course, this ~tSO input function is not well known, One goal of this paper is to find it. The input curve used here is a simplification of the 'best dated' stacked isotope record, with Terminations l-a, l-b, and 1-c occurring instantaneously at 13, 9.5, and 7 ka BP respectively. This was transformed to a depth scale using the age model from Fig. 6 ]dashed ~180 line in Fig. 7(a), (b) right]. By using the same ~180 input function for G. sacculifer and N. dutertrei, we assume that both species experienced the same signal. That is, if local temperature and/or salinity variations are present, they are the same for both species. Mixing of this hypothetical ~180 input curve in combination with the different abundance curves for (~. sacculifer and N. dutertrei yields different output curves that are reasonably consistent with the data [5180 crosses in Fig. 7(a), (b) right]. For G. sacculifer [Fig. 7(a), right], the total amplitude of the ~180 output is 1.7%o, with a slight ~lSO minimum within stage 2, and a hint of a pause between Terminations la and lb. For N.

82

A.C. Mix and W.F. Ruddiman

V 2 5 - S g G.sacc./gm

(a)

!

dlleo G.Iiaccullter + 1.33

O,

•

.

V25-59 N.duU~'./gm

(b)

,

!

I

J

8too N.cluter + 0.60

!

J 4. :

ig

m.

FIG. 7. Effects of bioturbation on fi~xO stratigraphy, core V25-59. Left: Dashed lines give species abundance/gin input. Solid lines give the model output. Crosses are the measured abundance data. Right: Dashed lines give the hypothetical isotopic termination, with three steps, and with a total amplitude of 1.8',~,. Solid lines give the model output. Crosses give the measured ~ O data, plus a constant to give a value of 0.0 for the unmixed core top. (a) G. sacculifer. (b) N. dutertrei.

dutertrei [Fig. 7(b), right] the amplitude of the ~~O output is only 1.3%o, with a smaller ~1'~O minimum within stage 2 and a strong single pause within the deglacial transition. This pause was also generated in another model run (not shown), which used N. dutertrei abundances but assumed a constant rate of deglaciation. This, then, is an example of false structure induced by bioturbation coupled to abundance change.

Given the simplicity of the input function and uncertainties in bioturbation intensity, our ability to model the data reasonably well is encouraging. Other models were attempted with

Structure and Timingof the Last Deglaciation

83

different input functions and bioturbation intensities, without success. This suggests that (1) the input function used here, although oversimplified, is reasonably close to reality, and (2) for this core at least, the intensity of bioturbation remained relatively high throughout the deglaciation. This analysis of bioturbation cautions that isotope records from species with large abundance changes should not be included in a composite record. For that reason, the only planktonic-species included in the stacked record are G. sacculifer (which is consistently abundant in the warmer regions of the tropical Atlantic), N. dutertrei (which is consistently abundant only in the cooler eastern-boundary regimes), and G. tuber (which is consistently abundant in the central subtropical gyres). Records strongly affected by abundance changes, such as that of N. dutertrei from western Atlantic core V25-59, were not included in the stacks, because they'may contain artificial steps induced by bioturbation. Thus, although we have not eliminated the effects of bioturbation from our isotope stacks, we have attempted to minimize them by (1) using relatively high sedimentation rate cores (>2.5 cm/ka), and (2) using species that do not demonstrate large abundance variations within a core. Carbonate Dissolution

One of the reasons tropical Atlantic cores were studied is that carbonate sediments are well preserved there. All of the cores used are above the interglacial lysocline, and most of them are above the glacial lysocline. Thus, the effects of dissolution on our data have been minimized. Possible effects of carbonate dissolution on both radiocarbon dates and oxygenisotope analyses, however, require some discussion. Although carbonate dissolution does not directly change the radiocarbon content of a foraminiferal test, age offsets could result from the coupled effects of bioturbation and dissolution (Broecker and Peng, 1982). Benthic mixing causes parcels of sediment from a depth horizon to contain a spectrum of ages. Carbonate dissolution is assumed to occur within the bioturbated layer (Broecker and Peng, 1982). Particles with longer residence times within this layer have a greater probability of dissolving, because of their longer exposure to corrosive waters. Through preferential removal of these older particles, dissolution would compress the age-spectrum within a sample toward its younger end (the particles that are preserved), resulting in bulk-carbonate 14C ages younger than would occur at the same depth if no dissolution had occurred. This does not necessarily mean, however, that the measured age of a ~1'~O record is too young in a partially-dissolved sample. The foraminifera analyzed for ~ O in this study were well preserved, complete specimens. If Broecker and Peng's (1982) model is correct that fragments from partially dissolved samples are on average older than well-preserved specimens, a bulk carbonate date would be older than the correct date specifically applicable to the well-preserved foraminifera. Further complications arise because different species dissolve at different rates (Ruddiman and Heezen, 1967). If a species were immune to dissolution, compression of the age spectrum for the rest of the carbonate sediment could yield a bulk carbonate date that was too young for the immune species. One way around these problems is to date individual species of foraminifera using accelerator techniques. First attempts at this have been made by Andree et al. (1984), but

84

A.C. Mix and W.F. Ruddiman

much needs to be learned before this technique will be generally applicable. For now, we can say that effects of dissolution on radiocarbon dates used in this paper are minimal. If they occur, they make dates on the ~tSo chronology too old. Dissolution may also have direct effects on the 8180 data. Berger and Killingley (1977) inferred that relatively large dissolution effects (crossing the lysocline) may change ~ 8 0 values o f G. sacculifer by as much as 0.3%0. They stated that this occurs because dissolution selectively removes the smaller, thin-walled, porous specimens that calcify in shallower (warmer) water. Berger and Killingley (1977) analyzed all specimens >295 Ixm. Curry and Matthews (1981) reported large size-fraction effects (as much as 0.8%0) in G. bulloides between the sizes 212-250 p~m and 355-425 Cm. Our data from G. sacculifer in core V2556 indicate a ~tSO offset of 0.2%0 to 0.4?'00 between the size-fractions 355,415 Ixm and 415-500 Ixm (Appendix 1). Thus, part of Berger and Killingley's dissolution effect on ~L80 could be a size effect. In our data set, each species was sized when possible to within a narrow (50-85 p,m) range, and only well-preserved, mature forms were analyzed. Because of this careful sampling strategy, and the location of the cores above the lysocline, dissolution effects should be negligible in our data set.

Temperature Variations Temperature variations of tropical Atlantic surface waters are small relative to those at high latitudes (CLIMAP, 1976; McIntyre et al., 1981). Even temperature changes of a few degrees, however, would affect the ~ 8 0 record. Sea-surface temperatures were estimated from the foraminiferal faunas in all cores studied using the 'transfer function' method of Imbrie and Kipp (1971). The temperature estimates will be discussed fully in another paper (Mix et al., in press). Two examples, however, are given here in Fig. 8. In column 1 (left), mean annual seasurface temperatures are plotted vs age. At western equatorial Atlantic core V15-168, glacial sea-surface temperatures were 1-2°C warmer than at present, while at eastern equatorial Atlantic core V29-144, peak Holocene temperatures were 2°C warmer, and peak glacial temperatures were 2°C cooler than at present. In column 2 of Fig. 8 (center), the ~lsO data from G. sacculifer (with 'sac', 355-415 ~m) are plotted for both cores. The ~ s O trends are similar between cores, except that V15-168 has a slightly higher amplitude 31SO record than V29-144. In column 3 of Fig. 8 (right), we estimate temperature-free isotope records in the two cores by subtracting the isotopic effect of the temperature estimate using the mollusk equilibrium equation of Epstein et al. (1953). The form of the estimated temperature-free ~lSO signals is quite different in the two cores. In V15-168, the amplitude of this estimated signal is over 2.5%o, while in V29-144, the amplitude is only 1.2%o, with a very gradual (and noisy) glacial-to-interglacial transition. Thus, attempts to correct for temperature overprints on the ice-volume ~ O signal have made the ~tsO records more dissimilar, rather than more similar. There are four possible explanations for the difference in amplitudes between these estimated temperature-free 8~SO signals. First, the temperature estimates may be wrong. The transfer function technique assumes that foraminiferal assemblages are in some way

Structure and Timing of the Last Deglaciation

85

V15-168 $ST

¢~ b 2

61 80

mean ©

b

~

6, o

i b

corr.

"

"

a.o

o

o

16.0

6180 corr.

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? m

2

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V29-144 FIG. 8. Effects of temperature on ~ O in cores V15-168 (top) and V29-144 (bottom). Left: Transfer-function temperature estimates of sea-surface temperature. Center: 8 ~ O data measured on G. sacculifer (with 'sac', 355-415 I~m). Right: Temperature-corrected ~5~O data. Removing temperature estimates from two similar ~mO curves yields dissimilar corrected 8t~O curves. This suggests that attempts to remove temperature from the ~ O signal using transfer-function temperature estimates do not improve estimates of changing oceanic ~1~O.

(not necessarily linearly or directly) related to temperature. If this assumption is incorrect, then temperature estimates in the past may be artifacts of other ecological changes unrelated to temperature. Arguments against this are that an equation calibrated in the North Atlantic can successfully estimate modern temperatures in the South Atlantic (Imbrie and Kipp, 1971), and different fossil groups from the same samples generally yield concordant paleotemperature estimates (Molfino et al., 1982). Second, the temperature estimates may be correct, but local salinity (and therefore water ~]sO) changes related to river outflow overprint the temperature effect in these cores, which are near the continents. If true, this would require high outflow from tropical African rivers to the site of V29-144 in glacial time, and/or from South America to the site of V15168 in interglacial time. This is consistent with studies inferring wet interglacial conditions in South America (Bradbury et al.+ 1981), but not with studies finding wetter conditions in Africa during deglaciation than during glacial maxima (Street and Grove, 1979). Third, the temperature estimates may be correct for surface waters, but G. sacculifer changes its depth and/or season of preference to track a specific temperature. In cultures,

86

A.C. Mix and W.F. Ruddiman

G. sacculifer calcifies at a range of temperatures from 14°C to 30°C (Erez and Luz, 1983). It is unknown, however, whether G. sacculifer can reproduce over this range. The culture studies may not apply to foraminiferal populations in the ocean, where competition between species may restrict given taxa to their optimum temperatures. Williams and Healy-Williams (1980) argued that modern surface-water temperature effects are present in ~ 8 0 measurements of core-top foraminifera, strongly for G. ruber and O. universa, but weakly for G. sacculifer. Fourth, the transfer-function temperature estimates may apply to subsurface (thermocline) waters, rather than surface waters. Fairbanks and Wiebe (1980) show that many species of foraminifera live within the thermocline where nutrients (and phytoplankton) are most abundant. Changes in thermocline depth could change foraminiferal species abundances, and thus the temperature estimate, without affecting surface-water temperature. The surface-dwelling G. sacculifer would thus incorporate little or no temperature signal into its down-core isotope record, even though temperature effects on the foraminiferal population as a whole were significant. The solution to this problem is not clear at present. It is apparent, however, that removing transfer function temperature estimates from foraminiferal 81SO measurements does not improve the proxy for ice-volume fluctuations. Thus, for the present, the ~lSO stacks without temperature corrections will be used as a measure of the changing isotopic composition of ocean water related to ice-volume change. This does not prove, however, that temperature effects are absent from the isotope records. Extraction of these effects from isotope records remains a major uncertainty in the interpretation of ice volumes from ~lsO measurements, which will limit our conclusions. The major constraint remains the first-order covariance of benthonic and planktonic records. Because benthonic records are limited by the freezing point of sea water, they constrain temperature effects to be less than 2°C, or 0.4%o ~tsO in many sites.

DATING THE ISOTOPIC TERMINATION Our results from stacking many ~tSO records are more similar to the traditional chronology, which called for smooth and rapid deglaciation about 11 ka BP, than to the newer 'two-step' chronology of Duplessy et al. (1981) (Fig. 9). Although hints of steps exist in our chronology, they are more subdued than those of Duplessy et al. (1981). The error envelopes on our stacks allow the interpretation of a smooth isotopic transition. Slight miscorrelations during stacking, however, would suppress the expression of steps. Thus, the steps may be more severe than we can show here. The presence of steps in individual cores (Fig. 3) suggests that they are real. Assuming the steps are real, we estimate Termination 1a at 14 to 12 ka BP, Termination 1-b at 10 to 9 ka BP and a possible third step (Termination l-c) at 8 to 6 ka BP. Minor isotopic expression of deglaciation is permitted by our data as early as 16 ka BP. Slight differences in the timing of deglaciation obtained from different stacks (Fig. 5) indicate that the errors on these dates is about + I000 years. These dates are in radiocarbon years. Transforming them to calendar years is not possible at present, because changes in the radiocarbon content of surface waters due to changes in ocean ventilation rate (Broecker et al., 1984) or variations in atmospheric production of 14C

87

Structure and Timing of the Last Deglaciation (a)

(c)

(b)

. :

:

E

e

W "D

FIG. 9. Comparison of three ~JsO chronologies. Dotted lines for reference represent the traditional view of rapid deglaciation centered at 11 ka BP. (a) Best dated planktonic stack from this paper. (b) Average of benthonic and planktonic data from high-latitude North Atlantic core CH73139C. the original "two-step" model (Duplessy et al., 1981). (c) A stack of six Pacific cores by Berger (1982). Note the large amplitude differences between the records. Berger's low amplitude of 1.2%o reflects attenuation due to bioturbation in the low sedimentation-rate cores used. Duplessy's high amplitude of 2.2%0 is too large to be accounted for by ice volume alone.

(e.g. Suess, 1970) are not known within this time range. These effects are likely to be small, however, on the order of hundreds of years, rather than thousands. Other views of the termination exist. As discussed earlier, Duplessy et al. (1981) inferred steps of deglaciation. To check their chronology, they ~ompared pollen data from a deepsea core with dated pollen records from land. This generally agreed with their J4C-dated oxygen isotope record, but their best g~sO data and the pollen data were from different cores. The two planktonic foraminiferai species analyzed for ~lSO in the same core as the pollen data yielded inconsistent results. The major differences between our chronology and that of Duplessy et al. (1981) are that: (1) we suggest three steps to the deglaciation, while Duplessy et al. (1981) inferred two (more severe) steps, and (2) our chronology places most of Termination 1-a between 12 and 14 ka BP, while Duplessy et al. (1981) place it between 13 and 16 ka BP. Ruddiman and Mclntyre (1981a) proposed (based on evidence for suppressed plankton productivity) that a large quantity of glacial ice and meltwater entered the North Atlantic between 16 and 13 ka BP. This chronology (and the 8~SO chronology of Duplessy et al., 1981) are constrained by a volcanic ash layer considered by Ruddiman and Mclntyre (1973) to be 9.3 ka BP and by Ruddiman and Mclntyre (1981a) to be 9.8 (+1.2 - 1.3) ka BP. More recently, this ash layer has been dated On land at 10.60 +__0.05 ka BP (Mangerud et al., 1984). Ruddiman and Mclntyre's (1981a) chronology was based on extrapolation from this ash zone. using 9.8 ka BP for the age of the weighted mean ash concentration and assuming constant sedimentation rates. That is, the inferred reduction in biological productivity' during deglaciation was assumed to be compensated for by an increase in icerafted sediment flux.

88

A.C. Mix and W.F. Ruddiman

Ruddiman and Mclntyre (1981a) did not infer steps in the deglaciation. Following Mercer (1970) they proposed that a minor productivity suppression at 11-10 ka BP reflected disintegration of an Arctic ice shelf, rather than ice on land. This interpretation was based on the lack of ice-rafted detritus in the North Atlantic at this time, in contrast to its abundance in the earlier (16-13 ka BP) barren zone. If our new isotope chronology is correct, it may require revision of the North Atlantic foraminiferal chronology of Ruddiman and Mclntyre (1981a). It is possible, however, that both chronologies are correct. This would require either that North Atlantic productivity suppression is not a function of meltwater input, or that it is a highly non-linear function, with a small amount of meltwater from 16 to 13 ka BP causing a large reduction in foraminiferal productivity. Berger (1982) stacked planktonic ~180 data from eight Pacific cores (Fig. 9)~ This composite isotope chronology begins to rise at about 15 ka BP, and peaks at about 10 ka BP with the most rapid change occurring between 12 and 10 ka BP. A hint of a step exists, pausing at about 13 ka BP, but Berger cautioned about over-interpreting his data. Average sedimentation rates in these cores is about 2 cm/ka (compared to 4.2 cm/ka for our 'best dated' stack), and thus they are strongly affected by bioturbation. This is reflected in the amplitude of Berger's stack, which is only 1.2%o. Berger stated that his data are consistent with rapid deglaciation at about 11 ka BP, but considering problems with radiocarbon dates, 'the real 14C age o f the transition very likely lies between 9 and 10 ka BP', (Berger, 1982), p. 276. In a recent paper, Berger et al. (1985) proposed a standard deglacial 8180 stratigraphy based on a composite of three Atlantic cores with 14C-calibrated sedimentation rates ranging from 1.6 to 1.9 cm/ka. The preferred composite of Berger et al. (1985) was derived by averaging ~SO analyses from the three cores at their measured depths (i.e., ignoring the 14C dates and assuming that all cores had identical sedimentation rates). A time scale was applied by assigning an age.of 11 ka BP to the midpoint of deglaciation (or the G. menardii biozone boundary) and extrapolating. Following this age assignment, 2000 years was subtracted from each age value, to account for assumed contamination with old carbonate. Using this composite, Berger et al. (1985) inferred that two steps to deglaciation occurred, the first near 13 ka BP, and the second at 10 ka BP. An isotopic minimum, or 'overshoot' (assumed to correlate with a similar feature in the Gulf of Mexico) was inferred at 9 ka BP, followed by "rebound' to greater 8~O values about 8 ka BP. A gradual transition to full interglacial values followed, lasting until about 2 ka BP, We have found no evidence for this overshoot/rebound effect in our data. An alternate stack, discounted by Berger et al. (1985), was composed of the same data, but was produced by assigning mean sedimentation rates to each of the three cores based on their radiocarbon dates. This yielded a different chronology, with a relatively smooth deglacial transition between 15 ka BP and 7 ka BP. Sarnthein et al. (1982b), who dated three ~ s O records from the northwest African continental margin, were the first to propose three steps to the deglaciation. Their chronology placed deglaciation earlier than ours, and that of Duplessy et al. (1981), with Termination l-a between 16 and 13.5 ka BP, an intermediate melting event between 12.4 and 11.1 ka BP, and Termination l-b between 10 and 8.5 ka BP. Keigwin et al. (1984) also suggested that three steps exist in the termination, but were unable to date the events due to

Structure and Timing of the Last Deglaciation

89

severe sediment reworking that caused |4C age reversals in the very high sedimentation rate core they used. More direct evidence of glacial melting has been sought in the preserved record of |~Odepleted meltwater events in the Gulf of Mexico (Emiliani et al., 1975; Kennett and Shackleton, 1975; and Leventer et al., 1982). This record has proven difficult to date. The Leventer et al. (1982) chronology inferred meltwater input from 16.5 ka BP to 11.6 ka BP. It was based on extrapolation from the G. m e n a r d i i biozone boundary, which has been dated at 11 ka BP in the Atlantic (Broecker et al., 1960) but is diachronous in some regions (Sarnthein et al., 1982b). Kennett and Shackleton (1975) also used this biozone boundary for dating and argued that meltwater entered the Gulf between 15 and 12 ka BP. The chronology of Emiliani et al. (1975) contains many |4C dates, and suggests maximum meltwater at about 11.5 ka BP. As noted above, Berger et al. (1985) believe that the Gulf of Mexico isotopic minimum correlates with a similar feature in their tropical Atlantic data, but infer its age to be 9 ka BP. They also proposed that this event does not necessarily represent glacial meltwater. Given the present uncertainty of dating and disagreement in interpretation of the Gulf of Mexico records, we feel that they cannot yet be used to constrain the timing of deglaciation. We believe that the data presented in this paper give the clearest picture yet derived of the timing of the isotopic expression of deglaciation. Both our chronology and that of Duplessy et al. (1981) permit steps in the deglaciation. Although the steps in our stacked isotope records are at the statistical limits of detection, this may be caused by artificial smoothing of fine structure during stacking. Because of this smoothing, the best expression of the severity of the steps will come from individual records. The best estimate of the dates, however, should come from the stacked record.

COMPARISON

TO GLACIAL GEOLOGY

Our chronology, although generated entirely independently of land-based chronologies, is generally consistent with glacial geological data. The presence of ice close to its maximum extent as late as 14 ka BP has been well documented. The extensive literature has been reviewed recently by Denton and Hughes (1981). Although the southern Laurentide maximum was between 21 and 17 ka BP (Fayette and Tazewell Stades, and equivalents), the Carey Stade (and equivalents) at - 1 4 ka BP returned to nearly maximum conditions. In Iowa, this time was the glacial maximum (Ruhe, 1969). Also, Cordilleran ice extent was significantly larger at 14 ka BP than at 18 ka BP (Waitt and Thorson, 1983). Our Termination l-a (14-12 ka BP) coincides with marine incursion into the Gulf of St. Lawrence (=12.5 ka BP) (Gadd, 1976). Termination 1-b (10-9 ka BP) may coincide with the downdraw of ice over Hudson Bay, if Mayewski et al. (1981) are correct that the marine incursion into Hudson Bav about 8.2 ka BP (Prest, 1969) just cleared thin ice from the area after an earlier period of ice-stream downdraw. In contrast, Andrews (1973, 1984) felt that the events at =8 ka BP represented significant ice-volume loss via calving and/or downdraw at that time. Rapid deglaciation of Foxe Basin occurred after 7 ka BP, and final dissipation of ice over Labrador occurred prior to 5 ka BP (Blake, 1966), consistent with final deglaciation during Termination 1-c (8-6 ka BP).

90

A.C. Mix and W.F. Ruddiman

In Europe, deglaciation was rapid after 13 ka BP (Andersen, 1981). The Younger Dryas glacial readvance between 11 and 10 ka BP (e.g. Mangerud et al., 1978) coincides with the isotopic pause between Terminations 1-a and 1-b at 12-10 ka BP. In Antarctica, less that disintegration of since 17 ka BP, but speculated that rapid 8-6 ka BP.

is known about the timing of deglaciation. Stuiver et al. (1981) infer the West Antarctic Ice Sheet has been generally slow and continuous that pulses of fast deglaciation may have occurred. These authors ice retreat occurred at 7-6 ka BP, similar to our Termination 1-c at

Although we have attempted to point out some of the similarities between our isotope chronology and deglacial chronologies from land, this is not a simple comparison. There are two reasons for this. First, the ~t80 record is a globally averaged record that integrates numerous events worldwide. The glacial geological data are inherently local. It is highly unlikely that all the events of deglaciation fit into a simple global model, whether stepped or smooth. Second, the data from land generally consist of areal, rather than volumetric data. For legitimate comparison to the ~ 8 0 record, inferences regarding ice volumes based on the areal data are required. This will be addressed in the following section.

VOLUMETRIC

DEGLACIATION

Assuming that the isotopic chronology presented here is close to reality, the next step is to determine how this picture of the isotopic 'termination' relates to volumetric deglaciation on a global scale. The total amplitude of 8~80 variation in our Atlantic data is =1.7%o. Although temperature effects may be embedded in the 6180 signal, they can not be detected using transfer-function temperature estimates (see above). Thus, in the following analysis we begin by assuming that temperature effects on the marine 81SO signal are negligible. The rationale for this is thatwe will attempt to explain the signal as simply as possible (i.e. with a minimum number of variables), and onl~¢ add variables as necessary. We realize, however, that this assumption may be incorrect, and we will limit our conclusions accordingly. If temperature effects are negligible, can reasonable estimates of glacial-maximum ice volume explain the entire 1.7%o amplitude measured in 8~O? A constraint in this calculation is the isotopic composition of ice. Most authors (e.g. Dansgaard and Tauber, 1969: Broecker, 1978; and others) agree that Laurentide ice 6tSO was probably - 3 0 to -35%0 (SMOW). Isotopic effects on the ocean are then generally calculated assuming this ice value remains constant through time. Mix and Ruddiman (1984), however, pointed out that time-varying isotopic composition of glacial ice may cause non-linear recording of ice volume by 6t'~O. The equations governing the isotopic effects on the ocean are the balance of water and its isotopic signature: Vtota I

----

Vicc +

V~,ccan

6,,,,.tV.,°.i = 6i~.~.Vi~. + 6 ........ ~V .........

(2) (_3)

Structure and Timingof the Last Deglaciation

91

where ~ is the 8~O value of the reservoir, and V is the water-equivalent volume. For the modern world ~occa, = 0%o (SMOW), and ~ice = -47%o (the weighted mean of 28 x l0 t' km 3 of Antarctic ice averaging -50%0, 3 z 106 km 3 of Greenland ice averaging -30%0 and 1 x 106 km 3 of other ice averaging -20%o; volumes from Hughes et al., 1981). Given the volume of the oceans as 1.350 × 109 km 3 (Menard and Smith, 1966), we estimate ~total to be - 1.1%o. Using this result, the isotopic composition of excess glacial ice needed to explain the marine record can be calculated, if ice volumes are known. We use two sets of ice volume estimates from Hughes et al. (1981) that were intended to bracket minimum and maximum possible values for glacial-maximum ice volume. Not all Quaternary geologists, however, accept these models as minimum/maximum brackets. For example, the Flint (1971) model, containing 77 x 106 km 3 of ice at the glacial maximum, is somewhat smaller than the Hughes et al. (1981) minimum model, which contains 84 x lff' km 3 of ice. This difference results largely from Flint's smaller estimate of Antarctic ice volume. The Peltier and Andrews (1976) model contained =70 × 106 km 3 of ice (our estimate based on their published value of 77 m mean sea level change for complete isostatic equilibrium). Their model contained no Barents Ice .Sheet, because at that time they disputed its existence. It also neglected ice volume changes in Antarctica. If this ice is added, the Peltier and Andrews'model is approximately the same as the Hughes et al. minimum model in total volume. The Hughes et al. models are used here because they are the most recent reconstructions available that include Antarctica. For their maximum reconstruction, Hughes et al. (1981) estimate a total ice volume of 98 × 106 km 3. To produce the 1.7%o oceanic value, this ice must have averaged -38%0 (equations 2 and 3, above). Assuming present Antarctic ice (28 x 106 km 3) had a ~lsO value of -50%o, and excess Pleistocene ice in Antarctica (10 x 106 km 3) had a ~180 value of -55%o (Barkov et al., 1977), the northern-hemisphere ice would have had a ~180 value of -28%0. Similarly, for the Hughes et al. (1981) minimum reconstruction, containing 84 x 106 km 3 of ice at the peak glaciation, a northern-hemisphere ~JSO ice composition of -37%0 would be required to explain the entire marine record. Both of these isotopic values for northern-hemisphere ice are reasonable. Thus, either the Hughes et al. (1981) maximum or minimum model (or something in between) could produce the entire 1.7%o oceanic amplitude without other contributing effects. Another potential constraint for this calculation comes from direct indicators of sea level. According to Hughes et al. (1981), their maximum and minimum models imply respectively 163 and 127 meters of sea level fall at the glacial maximum without isostatic compensation of the sea floor. With complete isostatic compensation (also assumed by Peltier and Andrews, 1976), Hughes et al. (1981) indicate that the models imply respectively 117 and 91 meters of sea level fall. Various estimates of sea level fall from dated shells and peats on continental margins generally fall on the low end of this range. Curray (1965) estimated about 125 m of sea level fall, with the maximum occurring about 19 ka BP. Milliman and Emery's (1968) maximum (130 m) agreed in amplitude with Curray's, but dated at 15 ka BP. Neither of these curves was constrained well at the glacial maximum. Dillon and Oidale (1978) corrected Milliman and Emery's curve for inferred continental shelf subsidence, and suggested that only 90 m of sea level fall occurred. The scatter around this line is at least 20 m, however, and the deepest 'fixed' sample measured (80 m), which dated at about 10 ka BP, was ignored. Thus, in our opinion, sea level curves generated from "direct' indicators in

92

A.C. Mix and W.F. Ruddiman

'4 ICE VOLUME

ICE VOLUME

.i

p

0.0

_~..~.~'~-" ,. ~.~-'~.~ .... • • .~'~,

0

10.0

° • ° -

•

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,°

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)

FIG. 10. Comparison of the best dated isotope stack from this paper with ice-volume estimates made using ice-area data. (A) Comparison to Bloom's (1971) estimates of Northern-Hemisphere ice volume, which assume a constant power-law relationship between ice area and ice thickness. Curve a: our best dated stsO stack. Curve b: ice area ~25 (i.e. equilibrium profile). Curve c: ice area L5 (Bloom's preferred model). Curve d: ice area 2-° ('downwasting' profile). (B) Comparison to Paterson's (1972) estimates of Laurentide ice volume, which use ice-sheet areas as input to a model of ice rheology. Curve a: our best dated ~ 8 0 stack. Curve b: equilibrium profile ice sheets. Curve c: stagnant profile ice sheets after 11.8 ka BP. Curve d: stagnant profile after 15 ka BP. Paterson preferred a model intermediate between curves b and c. The isotope chronology lags volumetric estimates, especially during late and early deglaciation. Possible isotopic steps near 13 ka BP, 9.5 ka BP and 7 ka BP and pauses (or reductions in rate) at 10.5 and 8 ka BP are not readily apparent in the ice-area data.

continental margin sediments are not yet sufficiently precise to constrain ice volume and isotope models. Hopefully further work will improve this situation. If anything, the sea level studies suggest that the Hughes et al. (1981) minimum model is closer to correct than the maximum model. Using ice area data from North America and Scandinavia, Bloom (1971) attempted to reconstruct the time sequence of volumetric deglaciation (without an absolute scale), assuming power-law relationships of ice-sheet thickness to area. Bloom's ice volume estimates are compared to our best-dated isotope stack in Fig. 10(A). The curves shown are (a) our 'best dated' isotope stack, (b) ice-sheet area ~25 (roughly equivalent to an equilibrium profile ice-sheet), (c) ice-sheet area 15 (Bloom's preferred curve), and (d) icesheet area 2° (Bloom's example of a 'downwasting' ice-sheet). Our ~180 stack lags Bloom's volume estimates somewhat, especially in the early and late phases of deglaciation. A slight pause in Bloom's volume estimates occurs at the Valder's readvance, 11.8 ka BP. McDonald (1971), and Paterson (1972), however, made similar compilations of ice-sheet areas, and did not find this pause. Wright (1971) suggested that minor ice margin readvances during deglaciation were local surges without implications for pauses in net areal deglaciation. Mickelson et al. (1983), however, took the alternative view that readvances generally do correlate, and that present dating uncertainties obscure real steps

Structure and Timing of the Last Deglaciation

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in areal retreat. Bloom (1971) noted that the assumptions regarding ice-sheet thickness were the major unknowns in his study. Many variables affect ice-sheet thickness, such as area, isostatic depression, marine or non-marine character of the-ice sheets (Denton and Hughes, 1981), thermal characteristics (Sugden, 1977), and bed properties (Boulton and Jones, 1979). Paterson (1972) made a more rigorous estimation of Laurentide ice volumes during deglaciation by considering some of these effects and using a model based on the rheology of ice, rather than just assuming an exponent to parameterize ice-sheet thickness. His volume estimates, although made just for the Laurentide, are normalized to 0-100% range for comparison to our best-dated isotope stack in Fig. 10(B). The curves shown are (a) our 'best dated' isotope stack, (b) Paterson's equilibrium profile model, (c) a model assuming a stagnant profile after 11.8 ka BP, and (d) a model assuming a stagnant profile after 15 k'a BP. Patersbn (1972) considered the most likely case to be intermediate between curves b and c. Two points requiring explanation emerge from Fig. 10. First, on average, the isotopic termination (solid lines) lags behind the volumetric deglaciation estimated from ice area data. Second, steps are not apparent in the areal retreat data, but are present (weakly) in the isotope stack. As discussed above, the stacking procedure underestimates the severity of steps. Thus, the differefices between a relatively smooth areal retreat of ice, and a discontinuous isotopic termination may be quite significant. Several explanations are possible for the isotopic lag. First, the chronologies for either the isotope record or the areal retreat reconstructions may be wrong. As discussed above, the errors on our 5~SO chronology are about +1000 years. No summaries of ice-sheet volumes during the deglaciation have been done since 1972, despite a significant increase in the available data. Second, the ice volume models based on areal retreat data are not global. They are based on North American and Scandinavian ice (Bloom, 1971), or just on Laurentide ice (Paterson, 1972). This represents only 60-70% of the total excess glacial ice volume. Other contributions to the total deglaciation (notably from Antarctica) may account for the offset. If this is the sole source of the mismatch, volumetric ice loss from Antarctica must lag Northern-hemisphere deglaciation. As noted above, this is possible, but not well constrained. Some evidence from Antarctica points to early ice retreat (of unknown volumetric significance), about 17 ka BP (Stuiver et al., 1981). Third, temperature changes could have contributed to the isotopic lag, but only if tropical Atlantic temperatureswere I°C colder during deglaciation than during the glacial and interglacial maxima. Although this may be true locally, the dominant pattern from the tropical Atlantic is cold glacial, warm interglacial (Mix et al., in press). As discussed above transfer function temperature estimates could not be detected in the isotope signals. Fourth, an isotopic contribution from a floating ice shelf could account for the observed offset between 51sO and ice volume estimates based on data from land, but only if the shelf volume was greatest during deglaciation. Broecker (1975) thought that the isotopic effect of a floating Arctic ice-sheet (at the glacial maximum) could be as much as 0.4%0, but Dodge et al. (1983) and Mix and Ruddiman (1983.) argued that its maximum effect was more likely 0.1-0.2%o. No one has proposed that ice shelves were largest during deglaciation.

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Fifth, non-linear recording of ice volume by ~180, caused by non-constant isotopic composition of ice sheets, may account for the lag of the blSO transition behind those of the area/volume models. Mix and Ruddiman (1984) pointed out that the isotopic expression of deglaciation would lag the true volumetric deglaciation by a few thousand years if lowerlatitude ('warmer', less ~80-depleted) ice sheets melted first. This idea is consistent with glacial geological data that show melting of southern Laurentide, British, and Scandinavian ice pre-dated melting of the high-latitude ice (e.g" Bryson et al., 1969; Prest, 1969: Andersen, 1981). Mix and Ruddiman's (1984) model, which produced a lag of ~180 behind true icevolume change, "assumed a smooth 10 ka transition. The possibility of steps in the deglaciation has implications for this calculation. The ~ 8 0 lag for a shorter-period change would be less (zero for an instantaneous step). Thus, rapid isotopic steps, if real and related to ice volume, probably do not lag significantly behind ice-volumetric steps. The pauses, however, may be misrepresented. For example, the ice sheets preserved through the pause between Terminations 1-a and 1-b were the high-latitude ice sheets. If these ice sheets were more depleted in 180 (i.e. colder) than the lower-latitude ice sheets that melted early, their volume during this pause was over-represented in the marine isotope record. By overrepresenting the ice volume at the pauses, an average lag of ~ISo behind ice volume of a few thousand years may still be generated, even though the timing of the steps is nearly correct. In summary, comparison of the isotope stack with estimates of glacial maximum ice volumes made from ice-sheet models and sea level estimates allows for the possibility that the isotope record is entirely a result of ice accumulation on land, without contributions from temperature or floating ice shelves. This does not mean that additional contributions to the ~tsO record do not exist. It means that they cannot be unambiguously detected within the constraints of present data. Comparison of the time-sequence of ddglaciation based on ice area data with the isotopic termination shows that the ~tsO transition lagged behind volumetric deglaciation of North America and Scandinavia, if the models based on ice area data are correct. The most likely causes of this apparent mismatch are (1) contributions to the isotope signal by ice volume change in Antarctica (especially important because of the greater t80 depletion of Antarctic ice), and (2) non-linear recording of ice volume by 8t~O, because of selective preservation of more ~O-depleted ice within the northern hemisphere during deglaciation. Further work on isotopes, sea levels, and glacial geology must be done on a global scale for more rigorous intercalibrations to be made.

IMPLICATIONS FOR MECHANISMS OF DEGLACIATION To the first order, our isotope chronology is consistent with the traditional view that deglaciation (centered near 11 ka BP) was driven by the seasonal distribution of insolation. Closer examination of the data, however, reveals complications. Although our chronology is slightly younger than that of Duplessy et al. (1981), we agree that the most rapid isotopic change occurred before the maximum in caloric summer insolation at 11 ka BP. Removing the effects of nonlinear recording of ice volume due to the changing isotopic composition of ice would only strengthen the early (pre-ll ka BP) deglaciation inferred from the isotope record. Thus, net ablation rates are not directly or linearly linked to insolation. Possible

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explanations are that slow melting forced by rising insolation triggered a more rapid mechanism that is internal to the earth's climate system, or that other mechanisms unrelated to insolation contibute to the tempo of deglaciation. Several plausible feedback mechanisms have been proposed that could contribute to rapid rates of deglaciation. Isostatic adjustment of the earth to the weight of the ice sheets has been investigated by Birchfield et al. (1981), Oerlemans (1980), and Peltier and Hyde (1984). The effects of ice calving rapidly to the ocean (Hoppe, 1948; Blake, 1966; Andrews, 1973) or to proglacial lakes (Andrews, 1973; Pollard, 1984) could augment melting begun by insolation change. Grounding line recession, caused as rising sea level 'unpins' marine portions of ice sheets (Weertman, 1974; Thomas, 1979), may accelerate this mechanism. In addition, Hughes (1977) and Denton and Hughes (1981) speculate that 'downdraw' of icesheet centers may occur via unstable ice streams. Ruddiman and Mclntyre (1981b) suggested that moisture starvation of the ice sheets due to expanded North Atlantic sea-ice cover during deglaciation may accelerate deglaciation. This mechanism is not independent of ice calving or downdraw, however. Ruddiman and Mclntyre envisioned moisture starvation as feedback forced in part by glacial meltproducts. Shackleton and Pisias (1985), using a ~13C proxy from marine sediments for atmospheric CO2 concentration, suggested that rising atmospheric pCO2 as early as 20 ka BP contributed to the early forcing of deglaciation. Direct measurements of pCO2 from ice cores (Neftel et al., 1982), although less well dated than the marine ~13C record, suggest that the deglacial rise in pCO2 did not begin until 13.5 + 1.5 ka BP. Ice core 8180 data from Antarctica, however, bear a striking resemblance to the Shackleton and Pisias (1985) proxy for pCO2, suggesting a potential causal link between CO2 and climate (Lorius et al., 1985). We do not attempt here to distinguish between any of these models, as most have not been rigorously tested in a way that would allow direct comparison with our data. We do point out, however, that the model or models that ultimately gain acceptance must account for most-rapid deglaciation earlier than the maximum summer insolation in the northern hemisphere, followed by slower deglaciation as full interglacial conditions are reached. In addition, models of deglaciation must account for the possibility of steps. As we discussed above, although steps in the isotopic termination are weak in our composite records, this may be an artifact of stacking. Their clear presence in many of our individual records (and in other investigators' data) suggests that they are real. If the steps are real, it is unlikely that they are explained by temperature or ice shelf effects superimposed upon a smooth deglaciatiQn, as this would require rapid cyclic behavior of a magnitude not observed or proposed. The two most likely explanations of the isotopic steps are (1) that they reflect rapid pulses of globally integrated deglaciation, or (2) they represent transient meltwater 'spike' inputs into the ocean enhanced by incomplete mixing of the ocean. Either case implies variations in the rate of volumetric deglaciation. We consider the first possibility more likely. If the steps present in benthic foraminiferal isotope records are real, the meltwater spike hypothesis can be rejected. Using the analogy of modern freshwater flux from the Amazon River, Jones and Ruddiman (1982) argued that glacial meltwater would mix into the ocean too rapidly to be detected in tropical Atlantic surface waters. The relatively smooth character of the ice areal-retreat data seems to preclude the possibility of steps in deglaciation if the ice sheets maintained an equilibrium profile. Thus, unless uncertainties in dating the continental record mask global steps in the retreat of the

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ice margins, it is unlikely that isotopic steps indicate advances and retreats of active ice sheets superimposed on a relatively smooth deglaciation. Paterson's (1972) models, however, provide a simple mechanism for generation of volumetric steps during deglaciation without corresponding steps in ice area; that is, alternation between nearlyequilibrium and nearly-stagnant profiles (end member conditions probably never existed). Stagnation of the ice sheets could occur following catastrophic down-draw events, which would result in dramatic thinning of the ice sheets (Denton and Hughes, 1981). Although down-draw may have occurred in conjunction with sea level rise and marine calving of ice, Denton (pers~ c o m m u n . , 1985) emphasizes that the critical point for down-draw is the instability of ice streams, and not necessarily the flux of ice to the ocean. Partial reequilibration following down-draw would account for reduced rates of volumetric deglaciation between steps, even though areal retreat continued (on average) unabated. In support of this scenario, large moraine systems that have been interpreted as 'reequilibration moraines' (that is, moraines produced by restoration of local ice-sheet equilibrium profiles following rapid ablation) formed in North America around 8 ka BP (e.g. Cockburn and Sakami moraines) and 9.6-10.8 ka BP (e.g. Lac Daigle-ManitouMatamek and St-Narcisse moraines) (Andrews, 1973; Hillaire-Marcel et al., 1981). This is consistent with our proposed timing of pauses in the isotopic termination. Thus, if the steps in the isotopic expression of deglaciation are real, their apparent lack of expression in the areal retreat of the ice sheets suggests that the process generating the steps is a significant thinning of the ice sheets without cor~esl~onding loss of area. That is, ice sheets may have alternated between thick, nearly-equilibrium, and thin, nearly-stagnant profiles. A likely mechanism for this occurrence is calving of ice to the oceans and/or downdraw of the center of the ice sheets. Multiple steps in the deglacial transition would suggest that down-draw was triggered more than once, possibly in different places or ice-volume states. This may reflect different ice sheets reaching points of instability at different times, or it may reflect a single ice-sheet reaching multiple points of instability during the transition from a glacial to an interglacial world. CONCLUSIONS

,

We have produced a well-dated composite record of the oxygen-isotopic expression of the last deglaciation, using planktonic and benthic foraminifera from the tropical Atlantic Ocean. To the first order, centering of the termination near 11 ka BP is consistent with the 'Milankovitch' mechanism of insolation forcing deglaciation. The timing of maximum isotopic change prior to maximum summer insolation, and changes in the rate of deglaciation that are not present in insolation forcing, however, imply that other mechanisms control the tempo of glacial-interglacial climate change. Our chronology allows for (but does not prove) the possibility, suggested by Duplessy et al. (1981), that deglaciation occurred in discrete steps. The severity of steps in our isotopic composite is probably underestimated, however, because slight miscorrelations cause smoothing of high-frequency features. The appearance of steps in many individual records in this and other studies suggests that they are real. Although similar structures appear in various composites, slight differences in their ages imply that the errors of dating are about _+1000 years.

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We estimate that the major step of the isotopic transition, Termination l-a, occurred between 14 and 12 ka BP. The second step, Termination l-b, occurred 10-9 ka BP. We tentatively identify a third step, Termination l-c, which occurred 8 - 6 ka BP. Pauses or low rates of change in the isotopic termination occurred at about 11-10 ka BP, and again near 8 ka BP. The relatively smooth character of ice areal retreat data, if correct, suggests that highfrequency oscillations of active ice sheets are not responsible for the isotopic steps. To explain the steps, we hypothesize that discrete calving and/or down draw events were triggered as sea level rose. These events could have caused significant thinning of the ice sheets without greatly reducing their area. T o generate the multiple steps of deglaciation, points of instability may have been reached more than once, at various global ice volume states.

ACKNOWLEDGEMENTS The NATO meeting on the last deglaciation (Airlie House, Virginia. 1983), where a preliminary version of this paper was presented, stimulated much thought. Discussions with J.-C. Duplessy and N.J. Shackleton there were especially helpful. In addition, an early version of this manuscript was circulated to many scientists active in the field, Many thanks to J.T. Andrews, W.H. Berger, A. Bloom, W.S. Broecker, G.H. Denton and W.S.B. Paterson for returning spirited reviews. Assistance with the isotope analyses by D. King is appreciated. The L-DGO stable isotope lab is directed by R.G. Fairbanks, who was helpful throughout the project. This research was supported by NSF grants OCE80-18177 and OCE83-15237, and DOE contract DE-ACO2-79E~10097 (subcontracted from Brown University).

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Keigwin, L.D., Corliss, B.H., Druffel, E.M. and Laine, E.P. (1984). High resolution isotopic study of the latest deglaciation based on Bermuda Rise cores. Quaternary Research, 22. 383-386. Kennett, J.P. and Shackleton, N.J. (1975). Laurentide ice sheet meltwater recorded in Gulf of Mexico deep-sea cores. Science, 188, 147-150. Kipp, N.G. (1976). New transfer function for estimating past sea-surface conditions from sea-bed distributions of planktonic foraminiferal assemblages in the North Atlantic. In: Cline, R.M. and Hays, J.D. (eds) Investigations of Late Quaternary Paleoceanography and Paleoclimatology. Geological Society of America Memoir, 145, 3-42. Koopman, B. (1979). Saharastaub in den Sedimenten des subtropischen Nordatlantik wahrend der letzten 20,000 Jahr. Ph.D. dissertation, Christian-Albrechts University of Kiel, Kiel. FRG. Leventer, A., Williams, D.F. and Kennett, J.P. (1982). Dynamics of the Laurentide ice sheet during the last deglaciation: Evidence from the Gulf of Mexico. Earth and Planetary Science Letters, 59, 11-17. Lorius, C., Jouzel, J., Ritz, C., Merlivat, L., Barkov, N.I., Korotkevitch, Y.S. and Kotlyakov, V.M. (1985) A 150,000-year climatic record from Antarctic ice. Nature (Lond.), 316, 591-596. Mangerud, J., Larsen, E., Longva, O. and Sonstegaard, E. (1979). Glacial of western Norway 15,000-10,000 B.P. Boreas, 8, 179-187. Mangerud, J., Lie, S.E., Furnes, H., Kristiansen, I.L. and Lomo, L. (1984). A Younger Dryas ash bed in western Norway, and its possible correlations with tephra in cores from the Norwegian Sea and the North Atlantic. Quaternary Research, 21, 85-104. Mayewski, P.A., Denton, G.H. and Hughes, T.J. (1981). Late Wisconsin ice sheets of North America. In: Denton, G.H. and Hughes, T.J. (eds) The Last Great Ice Sheets, pp. 67-178. Wiley Interscience, NY. McDonald, B.C. (1971). Late Quaternary stratigraphy and deglaciation in eastern Canada. In: Turekian, K.K. (ed.) Late Cenzoic Glacial Ages, pp. 331-353. Yale Univ. Press. New Haven, Conn. Mclntyre, A. and CLIMAP project members (1981). Seasonal reconstructions of the earth's surface at the last glacial maximum. Geological Society of America Map and Chart Series, MC-36. Menard, H.W. and Smith, S.M. (1966). Hypsometry of ocean basins. Journal of Geophysical Research, 71, 4305-4325. Mercer. J.H. (1970). A former ice sheet in the Arctic Ocean? Palaeogeography, Palaeoclimatology, Palaeoecology, 8, 19-27. Mickelson, D.M., Clayton, L., Fullerton, D.S. and Borns, H.W., Jr. (1983). The Late Wisconsin glacial record of the Laurentide Ice Sheet in the United States. In: Porter, S.C. (ed.) Late Quaternary Environments of the United States. Volume l, The Late Pleistocene, pp. 3-37. Univ. of Minnesota Press, Minneapolis, MN. Milankovitch, M. (1941). Kanon der Erdbestrahlung und seine Andwendung auf das Eiszeitproblem. Royal Serbian Academy Special Publication. 133, Belgrade. Milliman, J.D. and Emery, K.O. (1968). Sea levels during the past 35,000 yr Science, 162, 1121-1123. Mix, A.C. and Ruddiman, W.F. (.1984). Oxygen-isotope analyses and Pleistocene ice volumes. Quaternarv Research, 21, 1-20. Mix, A.C., Ruddiman, W.F. and Mclntyre, A. (in press) Late Quaternary paleoceanography of the tropical Atlantic, Part I: Spatial variability of annual mean sea-surface temperatures, 0-20,000 yr BP. Paleoceanography. Molfino, B., Kipp, N.G. and Morley, J.J. (1982). Comparison of Foraminiferal, Coccolithophorid, and Radiolarian paleotemperature estimates: Assemblage coherency and estimate concordancy. Quaternary Research, 17. 279-3t3. Neftel, A., Oeschager, H., Schwander, J., Stauffer, B. and Zumbrunn, R. (1982). Ice core sample measurements give atmospheric CO2 content during the past 40,000 yr. Nature, (Lond)), 295. 22{I-223. Oerlemans, J. (198{)). Model experiments on the 100,000-.y,r glacial cycle. Nature, (Lond), 287, 43(I-432. Paterson, W.S.B. (1972). Laurentide Ice Sheet: Estimated volumes during late Wisconsin. Reviews of Geophysics and Space Physics, 10, 885-917. Peltier, W.R. and Andrews, J.T. (1976). Glacial-isostatic adjustment - - I. The forward problem. Geophysics" Journal, Royal Astronomical Society, 46, 6{)5-646.

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P'eltier, W.R. and Hyde, W. (1984). A model of the ice age cycle. In: Berger, A., Imbrie, J., Hays, J., Kukla, G. and Saltzman, B. (eds) Milankovitch and Climate, Part 2. pp. 565-580. Reidel, Dordrecht. Peng, T.-H., Broecker, W.S. and Berger, W.H. (1979). Rates of benthic mixing in deep-ocean sediments as determined by radioactive tracers. Quaternary Research, 11,141-149. Pollard, D. (1984). Some ice-age aspects of a calving ice-sheet model. In: Berger, A., Imbrie, J., Hays, J., Kukla, G. and Saltzman, B. (eds) Milankovitch and Climate, Part 2. pp. 541-564. Reidel, Dordrecht. Prest, V.K. (1969). Retreat of Wisconsin and recent ice in North America. Geological Society of Canada, Map 1257A (with text). Ruddiman, W.F. and Heezen, B.C. (1967). Differential solution of planktonic foraminifera. DeepSea Research, 14, 801-808. Ruddiman, W.F. and Mclntyre, A. (1973). Time-transgressive deglacial retreat of polar waters from the North Atlantic. Quaternary Research, 3, 117-130. Ruddiman, W.F. and Mclntyre, A. (1981a). The North Atlantic Ocean during the last deglaciation. Palaeogeography, Palaeoclimatology, Palaeoecology, 35, 145-214. Ruddiman, W.F. and Mclntyre, A. (1981b). Oceanic evidence for amplific~ition of the 23,000-yr ice volume cycle. Science, 212, 617-627. Ruddiman, W.F. and Mix, A.C. (in press). The North and Equatorial Atlantic at 6000 and 9000 YBP. COHMAP Volume. Ruhe, R.V. (1969). Quaternary Landscapes in Iowa. Iowa State Univ. Press, Ames, Iowa. 255 pp. Sarnthein, M., Erlenkeuser, H. and Zahn, R. (1982a). Termination I: The response of continental climate in the subtropics as recorded in deep-sea sediments. Bulletin Inst. Geologique Bassin d'Aquitane, Bordeaux, 31,393-407. Sarnthein, M., Thiede, J., Pflaumann, U., Erlenkeusser, H., Futterer, D., Koopman, B., Lange, H. and Seibold, E. (1982b). Atmospheric and oceanic circulation off Northwest Africa during the past 25 million years. In: von Rad, U.. Hinz, K., Sarnthein, M. and Seibold, E. (eds) Geology of the Northwest African Continental Margin, pp. 545-604. Springer-Verlag, Berlin. Shackleton, N.J. (1967). Oxygen isotope analyses and Pleistocene temperatures, reassessed. Nature, (Lond.), 215, 15-17. Shackleton, N.J. (1973). Attainment of isotopic equilibrium between ocean water and the benthonic foraminiferal genus Uvigerina: Isotopic changes in the ocean during the last glacial. Colloques Internationaux du Centre National de la Recherche Scientifique, 219, 203-219. Shackleton, N.J. (1977). Carbon-13 in Uvigerina: Tropical rainforest history and the equatorial Pacific carbonate dissolution cycles. In: Andersen, N.R. and Malahoff, A. (eds) The Fate of Fossil Fuel CO_~in the Oceans, pp. 401-427. Plenum Publications, New York. Shackeleton, N.J. and Opdyke, N.D. (1973). Oxygen isotope and palaeomagnetic stratigraphy of equatorial Pacific core V28-238: Oxygen isotope temparatures and ice volumes on a 10: and 10" year scale. Quaternary Research, 3, 39-55. Shackleton, N.J. and Pisias, N.G. (1985). Atmospheric carbon dioxide, orbital forcing, and climate. In: Sundquest. E.T. and Broecker, W.S. (eds) The Carbon Cycle and Atmospheric CO,: Natural Variations Archean to Present. Geophysical Monograph 32. pp. 303-318. AGU. Washington, D.C. Street, F.A. and Grove, A.T. (1979). Global maps of lake-level fluctuations since 30.000 yr BP. Quaternary Research. 12, 83-118. Stuiver, M., Denton, G.H., Hughes, T.J. and Fastook, J.L. (1981). History of the marine ice sheet in West Antarctica during the last glaciation: A working hypothesis. In: Denton, G.H. and Hughes, T.J. (eds) The Last Great Ice Sheets, pp. 319-436. Wiley Interscience, NY. Stuiver. M. and Ostlund, H.G. (1980). GEOSECS Atlantic radiocarbon. Radiocarbon, 22, 1-24. Suess, H.E. (1970). The three causes of secular ~4C fluctuations, their amplitudes and time constants. In: Olsson, I.U. (ed.) Radiocarbon Variations and Absolute Chronology, pp. 595-605. Wiley Interscience. New York. Sugden, D.E. (1977). Reconstruction of the morphology, dynamics and thermal characteristics of the Laurentide ice sheet at its maximum. Arctic and Alpine Research, 9, 21-47. Thomas, R.H. (1979). The dynamics of marine ice sheets. Journal of Glaciology, 24, 167-177. Waitt. R.B.. Jr. and Thorson. R.M. (1983). The Cordilleran Ice Sheet in Washington. Idaho. and

102

A.C. Mix and W.F. Ruddiman

Montana. In: Porter, S.C. (ed.) Late Quaternary Environments of the United States, Vol. I. The Late Pleistocene, pp. 53-70. Univ. of Minn. Press, Minneapolis, MN. Weertman, J. (1974). Stability of the junction of an ice sheet and an ice shelf. Journal of Glaciology. 13, 3-11. Williams, D.F. and Healy-Williams, N. (1980). Oxygen isotope - - hydrographic relationships among recent planktonic foraminifera from the Indian Ocean. Nature, (Lond.), 283, 848-852. Wright, H.E. (1971). Retreat of the Laurentide Ice Sheet from 14,000 to 9,000 years ago, Quaterna O, Research, 1,316-330.

103

Structure and Timing of the Last Deglaciation

APPENDIX 1. 81~iO data V30-51K

V25-59

V25-75

V30-49

C. w~ellerstorfii

U. peregrina

C. wuellerstorfii

C. ~ellerszorfii

Depth

6 * '0

Depth

6 *"O

Depth

6 *'0

Depth

0.0 2.5

2.68 2.52

5.0 7.5 10.0 15.0 17.5 20.0 22.5 25.0 27.5 3O.O 32.5 35.0 40.0 42.5 45.0 47.5 50.O 55.0 60.0 67.5 70.0 72.5 75,0 77°5 80.0 82.5 85.0 87.5 90.0 92.5 95.0 97.5 100,0

2.41 2.63 2.56 2.78 2.87 2.75 2.94 2.98 2.83 3.50 3.08 3.77 3.60 3.51 4.30 4.07 4.23 4.28 4.29 4.35 4.18 4.15 3.93 3.90 3.84 3.76 3.80 3.90 3.86 3.56 3.72 3.61 3.62

0.0 0.0 10.0 20.0 30.0 40.0 60.0 70.0 70.0 80.0 90.0 100.0 I00.0

3.02 2.95 3.04 2.99 2.93 3.10 3.36 3.72 3.64 3.79 4.30 4.59 4.29

0.0 0.0 I .0 4.0 8.0 10.0 10.0 12.0 16.0 24.0 30.0 32.0 40.0 44.0 48.0 50.0 50.0 52.0 56.O 60.0 64.0 70.0 72.0 76.0 80.0 84.0 88.0 90.0 92.0 100.0 104.0 108.0 110.0 112.0 116.0 120.0 124.0 128.0 130.0 130.0 132.0 136.0 140.0 144.0 148.0 150.0

2.45 2.50 2.23 2.28 2.22 2.18 2.45 2.44 2.17 2.22 2.69 2.71 2.92 3.20 3.09 3.36 3 •49 3.36 3.45 4.10 4.08 3.99 4.20 4.06 3.84 4.04 3.74 3.88 3.77 3.98 3.26 3.25 3.54 3.46 3.39 3.64 3.55 3.25 3.59 3.48 3.75 3.57 3.36 3.59 3.72 3.33

3.0 8.0 10.0 1.0 2.0 3.0 4.0 6.0 6.0 17.0 18.0 20.0 24.0 26.0 28.0 30.0 31.0 32.0 33.0 35.0 35.0 38.0 38.0 38.0 39.0 42.0 44.0 45.0 46.0 47.0 48.0 49.0 50.0 52.0 54.0 56.0 58.0 61.0 63.O 65.0 67.0 69.0 76.O 78.0 82.0

v29-144

U. peregrina Depth 0.0 5.0 10.0 20.0 30.0 35.0 40.0 45.0 50.0 55.0 60.0 70.0 75.0 80,0 85.0 90.0 95.0 100.0 105.0 110.0

6 '"0 3.18 3.19 3.46 3.23 3.20 3.33 3.58 3.27 3.97 3.80 4.50 4.80 4.88 4.86 4,65 4.72 4.56 4.70 4.46 4.32

6 *'O 2.10 2.39 2.19 2.35 2.12 2.41 2.36 2.29 I .98 2.33 2.24 2.02 2.21 2.42 2.38 2.20 2.63 2.73 2.91 2.89 2.89 3.17 2.85 3.04 3.12 3.12 3.46 3.62 3.97 3.73 4.20 3.83 4.16 3.70 3.97 3.92 4.17 4.15 3.75 3.90 4.16 4'.03 3.73 3.33 3.59

104

A.C. Mix and W.F. Ruddiman

APPENDIX 1. (continued) EN66-I0

EN66-I0

RC13-189 (cont.)

RC24-07

G. sacculifer 355-415um Depth 6*'0

N. dutertrei 355-415um 6"O Depth

G. saoculifer 355-415~m

H. dutertrei 355-415~m

Depth

6*'0

Depth

6*'0

30.0 32.5 35.0 37.5 40.0 43.0 45.0 47.5 50.0 52.5 55.0 57.5

-0.64 -0.48 -0.06 -0.14 -0.03 0.30 0.28 0.24 0.39 0°17 0.25 0.21

60.0 62.5

0.28 0.07

65.0 67.5 ~'0.0

0o17 0.09 0.13

O.0 5.0 10.0 15.0 20.0 25.O 30.0 35.0 40.0 45.0 50 .O 55.0 60.0 65.0 70.0 75.0 80.0 90.0 95.0 100.0 105.0 110.0 115.0 120.0 125.0 130.0

-0.24 -O. 03 -0.29 -O.16 -O.03 0.03 o.19 0.54 0.70 0.93 0.92 I ,13 I .21 1.29 I . 45 1.86 I . 82 I .86 I .75 I .36 I .46 1.63 I .93 I .41 I .61 I .41

2.5 2.5 4.5 6.5 6.5 8.5 10.5 12.5 14.5 16.5 18.5 20.5 22.5 24.5 24.5 26.5 28.5 30.5 30.5 32.5 32.5 34.5 36.5 38.5 38.5 40.5

- I .45 - I .41 - I . 17 -I .20 -I .38 -I .23 -I .IO -0.85 -0.67 -0.57 -0.08 -0.06 0.27 O. 17 -0.27 O. 37 0.44 0.36 -0.01 0.38 -O. 10 O. 45 0.29 -0.02 o.13 0.19

2.5 4.5 6.5 8.5 10.5 12.5 14.5 16.5 18.5 18.5 20.5 20.5 22.5 22.5 24.5 26.5 28.5 30.5 32.5 34.5 36.5 38.5 38.5 40.5

-0.14 -0.35 -0.32 -0.38 -0.23 0.04 0.55 0.65 0.71 0.89 I .19 0.83 0.70 0.86 O.99 I. 46 I .47 I .22 I .31 I .33 0.87 I .22 1.18 0.85

RC13-189

G. sacaulifer 355-415um RC9-49

Depth

6 *'0

2.5 5.0 7.5 7.5 10.0 10.0 ~2.5 12.5 15.0 17.0 20.0 22.5 25.0 25.0 27.5 27.5 30.0

-I .44 - I .39 - I .47 - I .71 -I .48 - I .29 - I .32 -I .73 -I .19 -I .33 - I .12 -0.78 -0.39 -0.87 -0.28 -0.25 -0.59

C. saccu~ifer 355-415um Depth

6*'0

5.0 13.0 18.0 37.0 45.0 55.0

- I .70 - I .60 - I .43 -0.97 -0.92 -0.50 -0.10 -0.04 0.29 -0.30

65.0 75.0 85.0 95.0

RC24-01

N.dutertrei 355-41 5~m Depth

6 *lO

0.0 4.0 8.0 12.0 16.0 20.0 24.0 28.0 32.0 36.O 40.0 44.0 48.0 52.0 56.0 60.0 64.0 68.0 72.0 76.0 80.0 84.0 88.0 92.0 96.0 100.0

0.03 -0.19 0.05 0.11 0.35 O. 50 0.80 0.81 0.74 1.18 I .31 I .41 I .69 I .62 I .54 I .26 I .34 I .73 I .23 I .31 I . 30 I .58 I .52 i .55 I .33 i .57

Structure and Timing of the Last Deglaciation

105

APPENDIX l (continued) RC24-16

N.dutertrei 355-415pm Depth

61So

0.0 3.0 6.0 9.0 12.0 15.o 15.0 18.0 18.0 21.0 24.0 27.0 27.0 30.0 33.0 33.0 36.0 36.0

-0.21 -0.46 -o.01 -0.41 -o.13 -0.20 -0.07 -0.12 -0.37 -0.18 -0.08 O.O2 -0.13 -0.05 0.22 0.20 0.05

39.o 39.0 42.0 45. o 48.0 51 .'0 51.0 54.0 54.0 57.0 57.0 60.0 63.0 66.0 69.0 72.0 75.0 78.0 81.0 81.0 84.O 84.0 87.0 90.0 93.0 96.0 99.0 102.0 I05.0 108.0 111.0 114.0 117.0 120.0

0.03 0.o4 -0.05 0.84 O. 46 0.74 I .19 0.90 0.92 0.86 I .13 I .16 1.28 1.36 1.32 1.22 1.24 1.06 1.02 1.26 0.88 O.93 0.92 1.05 0.94 0.98 0.86 O.59 0.74 0.69 0.55 0.50 0.53 0.54 0.43

V15-168 G. sacculifer 355-415um Depth 6'IO 0.0 3-0 7.0 10.0 13.0 17.0 20.0 23.0 27.0 30.0 35.0 45.0 55.0 65.0 75.0 85.0 96.0 I05.0 115.0 125.0 135.0 142.0 155.0 165.0

-I .60 -I .98 -I .71 -I .68 -1.69 -1.51 -1.55 -1.61 -1.58 -0.82

-1.12 -1.15 -0.80 -I.00 -0.78 -0.94 -1.01 0.18 0.19 0.35 0.34 0.15 0.30 -0.02

V22-38

G. sacculifer 355-415~m Depth

O. 0 2.5 5.0 7.5 10.0 12.5 15.0 17.5 20.0 22.5 25.0 27.5 30.0 32.5 35.0 37.5 40.0 42.5 45.0 45.0 47.5 50.0 50.0 52.5 55.0 57.5 60.0

6110

- I .06 -I • 12 -0.96 -1.21 -1.20 -1.14 -1.06 -1.20 -0.80 -0.79 -0.70 -0.68 0.03 -0.05 0.03 0.10 -0.05 -0.07 -0.05 -0.19 -0.21 -0.18 -0.44 -0.52 -0.34 -0.57 -0.55

V22-182

V22-177

G. sacculifer 415-500um

G. saeculifer 355-415u

Depth

Depth

6*'0

0.0 O. 0 4.0 4.0 8.0 12.0 16.0 20.0 24.0 28.0 32.0 36.0 40.0 44.0 48.0 52.0 56.0 60.0 64.0 68.0 72.0 76.0 80.0 84.0 88.0

-O.93 - I . 01 -0.85 -0.90 -I .26 -I .27 - I .09 -0.99 -0.94 -0.76 -0.62 -0.69 -0.14 -0.15 0.13 0.43 0.44 O. 46 0.30 0.43 O. 48 0.39 0.32 0.24 0.21

2.5 5.0 5.0 5.0 7.5 10.0 10.0 12.5 12.5 15.0 17.5 20.0 22.5 25.0 27.5 27.5 27.5 32.5 32.5 35.0 37.5 40.0 42.5 42.5 42.5 45.0 45.0 45.0 47.5 52.5 52.5 55.0 57.5 60.0 63.0 63.0 66.0 66.0 69.0 72.0 75.0 78.0 81.0 84.0 87.0 90.0 93.0 96.O 99.0

6110 -I .31 -I .01 -0.98 -I .63 -1.20 -1.43 -1.66 -I.02 -1.55 -1.23 -1.14 -1.16 -1.13 -0.94 -1.02 -1.24 -1.22 -0.65 -0.34 -0.32 -0.16 -0.03 0.28 -0.24 -0.26 -0.01 -0.19 -0.23 0.08 0.12 0.09 0.30 0.33 0.50 0.11 0.11 0.48 0.00 0.29 0.08 0.26 0.14 0.20 -0.22 -0.35 -0.18 -0.23 -0,17 -0.22

V23-I I0

G. saoculifer 355-415~ Depth

6 ~aO

0.0 2.0 4.0 6.0 8.0 10.0 12.0 14.0 16.0 18.0 20.0 22.0 24.0 26.0 28.0 30.0 32.0 34.0 36.0 38.0 40.0

-0.72 -0.95 -I .28 -I. 05 -0.82 -0.66 -0.50 -0.76 -0.22 -O.01 -0.24 -0.17 -0.43 -0.42 -0.37 -0.57 -0.2O -0.27 -0.27 -0.41 -0.84

106

A.C. Mix and W.F. Ruddiman

APPENDIX 1. (continued) V25-56 G. saccu li f er 415-500um 6S,O Depth

V25-56 (cont.) G. sacculifer 355-415um Depth 6'I0

5.0 10.0 15.0 21.,0 25.0 30.0 35.0 40.0 45.0 50.0 50.0 55.0 60.0 65.0 70.0 76.0 8O.0 85.0 90.0 95.0 100.0 100.0

70.0 75.0 80.0 85.0 90.0 100.0 105.0 110.0 115.0 120.0 125.0 130.0 135.0 140.0 145.0 150.0

V25-56 G.sacculifer Depth 0.0 5.0 10.0 15.0 15.0 20.0 25.0 30.0 35.0 40.0 40.0 45.0 50.0 50.0 55.0 55.0 60.0 60.0 65.0 65.0

- I .35 - I .46 - I .45 -1.39 -1.25 -1.12 -1.01 -0.96 -0.58 -0.69 -0.57 -0.25 0.18 0.12 0.05 0.08 0.08 0.19 0.01 0.00 0.11 0.19

355-415~m 6Z'O -1.82 -1.80 -1.73 -1.78 -1.86 -1.65 -1.44 -1.49 -1.12 -0.95 -1.02 -0.80 -0.69 -0.90 -0.55 -0.16 -0.25 0.09 0.13 -0.08

V25-59 G.sacculifer Depth 2.5 5.0 7.5 I0.0 12.5 15.0 17.5 20.0 22.5 25.0 30.0 32.5 37.5 40.0 42.5 45.0 47.5 50.0 52.5 55.0 57.5 60.0 62.5 67.5 70.0 72.5 75.0 77.5 80.0

0.00 0.07 0.10 -0.03 -0.17 -0.26 -0.12 0.00 -0.16 -0.21 -0.29 -0.32 -0.29 -0.38 -0.28 -0.44

415-500um 6*'0 -1.33 -1.42 -1.25 -1.13 -1.28 -1.11 -0~63 -0.18 -0.67 -0.62 -0.40 -0.11 -0.16 0.10 0.16 0.15 0.41 0.13 0.32 0.03 0.26 0.35 0.14 0.18 0.25 0.28 0.23 0.00 o.16

V25-59 N.dutertrei Depth 2.5 5.0 7.5 10.0 12.5 15.0 17.5 20.0 22.5 25.0 30.0 32.5 35.0 37.5 40.0 42.5 45.0 47.5 47.5 50.0 52.5 55.0 57.5 60.0 62.5 67.5 70.0 72.5 75.0 77.5 80.0 82.5

415-500~m 6*'0 -0.18 -0.41 -0.27 -0.05 -0.16 -0.18 O. 06 0.39 0.15 O.77 0.44 0.58 0,73 0.76 0.61 0.99 0.95 0.82 I. O0 0.88 O. 98 0.78 0.88 0.83 0.84 0.76 0.58 O.60 0.70 0.64 0.59 0.58

V25-60 G. sacculifer Depth 5.0 7.5 11.0 12.5 15.0 17.5 20.0 22.5 25.0 25.0 27.5 29.0 32.5 35.0 37.5 41.0 42.5 45.0 47.5 50 .o 52.5 55.0 57.5 62.5 65.5 67.5 69.0

355-415~m 6''0

-1.08 -1.29 - I .34 -1.50 -1.57 -1.23 -0.79 -0.62 -0.76 -I .O2 -0.64 -0.82 -0.54 -0.26 0.12 0.19 O. 13 -o.oi -0.05 -o. o3 O. O2 0.09 -0.22 -0.12 0.38 -0.48 -0.18

107

Structure and Timing of the Last Deg|aciation

APPENDIX l. (continued) V25-75 G. sacculifer 250-500um Depth 6 *"0 0.0 7.O 1o.0 10.0 20.0 25.0 30 .o 30.0 40.0 45.0 50,0 60.0 65.o 70.0 75.0 75.0 80.0 85.0 9O.O 95.0 100.0 I00.0 105.0 110.0 115.0 115.0 120.0 125.0 1 30.0 135.0 140.0 145.0 150.0

-I .24 -I .51 - I .73 -I .75 -I .56 -I .62 -I .65 -I .78 - I .39 -I .48 -I .34 -1.16 -0.97 -o.79 -I. o8 -I. O6 - I .14 -0.37 0.09 O. 07 -0.25 -0.18 -0.04 -0.06 -0.30 -0.19 0.02 0.17 0.22 0.16 0.20 0.06 0.11

V29-144 N.dutertrei 355-415~m Depth 61'0

0.0 10.0 15.0 20.0 25.0 30.0 35.0 45.0 50.0 55.0 60.0 65.0 70.0 75.0 80.0 85.0 90.0 95.0 100o0

105.0 110.0

-0.29 -0.12 -0.05 0.18 0.16 -0.16 0.09 0.54 0.56 O.8O .05 .9O .47 .47 .43 .54 .34 .30 .34 .29 .10

V29-I 44 G. saccu filer 355-415~m Depth 6, '0 0.0 5.0 10.0 15o0 20.0 25.0 30. o 35.0 40.0 45.0 50.0 55.0 60,0 65.0 70.0 75.0 80.0 85.0 9O.O 9O. o 95.0 95.0 100.0 100.0 105.0 105.0 111.0 110,0

-1.26 -1.31 -1.38 -1.51 -1.09 -0.97 - I . 39 -1.11 -0.76 -0.73 -0.44 -0.70 -0.43 -0.18 -0.29 0.38 0.50 0.30 -o.01 -0.38 0.42 0.16 0.23 0.40

0.43 0.24 -0.12 -0.12

V30-40 G. sacculifer 415-500~m Depth 61sO 0.0 0.0 3.0 3.0 6.0 9.0 12.0 15.0 18.0 21 .0 24.0 27.0 30.0 33.0 36.0 36.0 39.0 42.0 48.0 51.0 54.0 54.0 57.O 60.0 63.O 66.0 69.0 72.0 72.0 75.0 78.0 81 o0 84.0 87.0 90.0 90.0 93.0 96.0 96.0 99.0 99.0 102.0 105.0

-I .22 TI .29 -I .22 -I .28 -I .26 -I .19 -0,80 -0.98 -0.96 -0.74 -0.61 -0.25 -0.34 -0.32 -0.42 -0.1 6 -0.15 0.oo o.27 0.52 0.45 O. 45 0.52 o. 40 0.41 O.3O 0.37 0.30 O.27 0.48 0.38 0.36 0.30 0.31 0.37 0.09 O. 05 o. 07 -0. oo -0.02 -0.09 -0.06 -0.12

V30-36 G. sacculi f er 250-5001Jm DepT.h 6, '0

0.0 0.0 2.0

-1.46 "1.38 -1.38

4.0

-1.27

6.0 6.0 8.0 10.0 12.0 14.0 16.0 18.0 20.0 22.0 24.0 26.0 26.0 28.0 32.0 34.0 36,0 38.0 40.0 42.0 44.0 48.0 50.0

-1.45 -I .29 -1.54 -I .32 -I .25 - I .08 - I .36 -I .03 -0.48 -0.58 -0.42 -0.26 -0.61 -0.21 -0.17 -0.13 0.42 0.14 -0.01 0.28 0.35 -o. 06 -0.15

108

A.C. Mix and W.F. Ruddiman

APPENDIX 1. (continued) V30-41

V30-51K

V30-51K (cont.)

G. sacculifer 250-415wm

N. dutertrei 355-415!Jm

N.dutertrei 355-4151~m

Depth

6''0

Depth

6 s60

Depth

6 *=0

4.5 6.5 8.5 10.5 12.5 12.5 14.5 16.5 18.5 20.5 22.5 24.5 26.5 28.5 30.5 32.5 36.5 38.5 40.5 48.5 52.5 54.5

- I .32 - I .35 - I .43 - I .23 - I .48 -I .43 - I .24 - I .32 - I .33 -I .40 -0.71 -0.60 -0.29 -0.35 -0.09 0.13 0.51 0.45 0.39

0.0 I .0 2.0 3.0 4.0 5.0 7.0 8.0 9.0 10.0 11.0 11.0

-0.02 -0.08 - o . 16 0.03 -0.02 -0.09 -0.18 -0.06 -0.I0 -0.13 -0.33 0.04

41.0 42.0 43.0 44.0 45.0 46.0 47.0 48.0 49.0 49.0 50.0 51.0 52.0 53 .O 54.0 55.0 56.0 57.0 58.0 59.0 60.0 61.0 62.0 63.0 64.0 65.0 66.0 67.0 68.0 69.0 70.0 71.0 72.0 73.0 74.0 74.0 75.0 76.0 77.0 78.0 79.0 80.0 81.0 82.0 84.0 85.0

I .03 0.69 0.89 I .65 1.46 I .33 I .40 I .48 I .76 I .69 I .37 I .32 I . 48 1.61 I .47 I .67 I .49 I .28 I .29 I .65 I .93 I .35 I .70 1.61 I .61 I .50 I .57 I .32 1.41 1.50 I .20 I .61 I .29 I. 41 1.37 I. 37 1.49 I .09 1.35 I .25 I .29 0.98 I .40 1.35 0.98 1.30

-0.02

0.17 0.12

V30-49

G. saceuLifer 355-415~m Depth 4.0 12.0 1 6.0

20.0 24.0 28.O 32.0 36.0 40.0 44.0 52.0 56.0 6O.O 64.0 68.O 72.O 76.0 80.0 84.0 88.0 92.0 96.0 100.0

6Z'O - I .14 - I .02 - I . 07 -0.95 -0.53 -0.64 -0.75 -0.42 -0.47 -0.27 -0.19 -0.33 -O.O6 0.50 0.53 0.62 0.73 0.36 0.48 0.59 0.64 O.67 0.38

12.0

0.05

13.0 13.0 14.0 15.0 1 6.0 17.0 18.0 18.0 19.0 20.0 21.0 22.0 24.0 25.0 26.0 26.0 27.0 27.0 28.0 29.0 30.0 31.0 32.0 32.0 33.0 34.0 35.0 36.0 37.0 38.0 39.0 40.0

-0.30 -0.09 -0.11 -0.36 0.03 0.14 -0.22 -0.42 -0.01 0.12 -0.02 0.12 0.06 0.27 0.12 0.39 0.10 0.09 0.26 0.37 0.24 0.45 0.81 0.14 0.68 0.98 0.67 0.92 0.97 0.76 0.87 0.95

V32-08 G.ruber (wl~Ite) 250-355~m Depth 6, so 1.0 1.0 4.0 8.0 12.0 12.0 16.0 16.0 20.0 20.0 24.0 28.0 32.0 36.0 40.0 40.0 44.0 44.0 48.0 52.0 56.0 60.0 64.0 68.0 68.0 72.0 76.0 76.0 80.0 80.0 84.0 88.0 88.0 92.0 96.0 I00.0 104.0 108.0 112.0

-0.86 -0.56 -0.70 -0.71 -0.81 -0.91 -0.32 -0.70 ,,0.11 -0.25 -0.12 0.21 o. 18 0.24 -0.10 -0.07 0.53 0.56 0.77 0.69 0.90 0.96 1.17 1.31 0.98 I .IO I .09 I .11 0.41 0.52 0.57 0.61 0.78 0.97 0.20 0.31 0.50 0.49 0.27