Tropical forcing of North Pacific intermediate water distribution during Late Quaternary rapid climate change?

Quaternary Science Reviews 22 (2003) 673–689

Tropical forcing of North Pacific intermediate water distribution during Late Quaternary rapid climate change? Ingrid L. Hendya,*, James P. Kennettb b

a Department of Earth and Ocean Sciences, University of British Columbia, Vancouver, BC, Canada Department of Geological Sciences, and Marine Science Institute, University of California, Santa Barbara, CA, USA

Abstract Evidence is presented demonstrating intermediate water (B500 m) temperature variability at ODP Hole 893A in Santa Barbara Basin during submillennial climate change (11–60 ka). Benthic d18O oscillations are considered to result primarily from shifts in intermediate water temperature at the site. Detailed comparison of both benthic and planktonic records from the basin provide crucial evidence for differing surface and intermediate water mass temporal responses to rapid climate change. Gradual warming of intermediate water compared to abrupt cooling suggests mechanistic differences between processes controlling North Pacific Intermediate Water expansion and contraction relative to ‘southern component’ intermediate waters. Comparisons suggest intermediate water warming preceded (by 60–200 years) the most rapid interval of surface warming inferred to be associated with North Pacific atmospheric reorganization. Tropical forcing of sea level anomalies in the eastern Pacific via trade wind strength may control California Undercurrent flow (300–500 m) and be the cause of early intermediate water warming in Santa Barbara Basin. r 2002 Elsevier Science Ltd. All rights reserved.

1. Introduction Recent high-resolution studies of late Quaternary paleoclimate have demonstrated inherent instability in the cryosphere/climate system on submillennial to multidecadal time-scales (Bond et al., 1993; Grootes et al., 1993). Significant, abrupt warmings and coolings (Dansgaard/Oeschger or D–O events) during the last glaciation imply little inertia in the Earth’s climate system to prevent rapid switching between different climate states. However, processes forcing climate change remain poorly understood. Understanding the temporal relationship between climate change processes will improve our ability to predict climate change, which may prove significant in an age of anthropogenic climate forcing. The recovery of relatively rare geologic records with extremely high depositional rates and reliably preserved climate proxy signals of atmospheric/oceanic processes is crucial for determining temporal relationships. One such record, Santa Barbara Basin ODP Hole 893A (341 *Corresponding author. Present address: Department of Geological Sciences University of Michigan, Ann Arbor, MI 48104, USA. Tel.: +1-734-615-6892; fax: +1-734-763-4690. E-mail address: [email protected] (I.L. Hendy).

17.250 N, 1201 2.20 W) provides the highest known resolution late Quaternary paleoclimatic sequence of surface and intermediate water response. Here we demonstrate for the first time, the great variability of intermediate water temperature in the North Pacific during the last 60 ka, confirming intermediate water mass changes at the site (Kennett and Ingram, 1995; Behl and Kennett, 1996; Cannariato and Kennett, 1999). Comparison of multi-proxy records within the core at unprecedented detail also allows a unique and critical insight into the different temporal behaviour of thermohaline circulation and atmospheric reorganization, in response to rapid climate change. These observations may be inferred to be the possible result of an early tropical response during rapid warmings. That intermediate waters can be extremely responsive to climate forcing has been demonstrated with monitoring of both intermediate water warming (Levitus et al., 2000) and freshening (Wong et al., 1999) over the last few decades. This alongside emerging evidence of methane hydrate instability (Kennett et al., 2000, 2002) at intermediate water depths has raised the question of the significance of intermediate water temperatures in producing greenhouse gas amplification of late Quaternary climate change. The potentially significant role of intermediate water temperature warming in

0277-3791/03/$ - see front matter r 2002 Elsevier Science Ltd. All rights reserved. PII: S 0 2 7 7 - 3 7 9 1 ( 0 2 ) 0 0 1 8 6 - 5

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destabilizing methane hydrates supports serious consideration of benthic paleotemperature records between 400 and 1000 m depth. 1.1. Regional setting Santa Barbara Basin is located in a highly productive, environmentally sensitive region and provides an exceptional setting for climate change studies. The modern basin (B600 m deep with a B450 m sill) containing oxygen minimum zone waters below sill depth, is further depleted of oxygen by organic material degradation from surface productivity (Emery and Hulsemann, 1962). Switches in seasonal sedimentary sources are preserved as annual laminations through dysoxic conditions, which prevent bioturbation by benthic macrofauna (Emery and Hulsemann, 1962). Massive, well-oxygenated sediments occurring during cool intervals of the last glacial have been suggested to result from a proximal well ventilated NPIW source, and laminated, low oxygen sediments occurring during warm intervals, a distal, poorly ventilated source (Keigwin and Jones, 1990; Kennett and Ingram, 1995; van Geen et al., 1996; Cannariato and Kennett, 1999). Previous work at ODP Hole 893A has demonstrated that the oxygen content of the basin fluctuated during the D–O cycles of the last glacial due to some combination of changes in surface water productivity and intermediate water ventilation (Kennett and Ingram, 1995; Behl and Kennett, 1996; Cannariato and Kennett, 1999). Santa Barbara Basin bottom water is controlled by the interaction of sill bathymetry with intermediate water circulation (Emery, 1954). Intermediate water enters the basin from the northwest from the continental slope (Emery, 1954). A Santa Barbara Basin flushing event recorded in spring suggested to be in response to upwelling (Sholkovitch and Gieskes, 1971) is similar to better documented events occurring in other Southern Californian basins, which appear to have no simple relationship with upwelling, subtidal or interannual (ENSO) variability (Hickey, 1991). Flushing events within these basins occur abruptly during a period of less than a month and with a frequency of less than once a year (Hickey, 1991). 1.2. Undercurrent flow Subsurface poleward flow occurs along all five major oceanic eastern boundaries, opposing the surficial equatorward currents at mid-latitudes. A poleward undercurrent (California Undercurrent) flows just off the shelf break of the North American continent carrying warm, salty, low-oxygen water north from Baja California to Vancouver, Canada (Fig. 6c; Reid, 1965; Hickey, 1979; Chelton, 1984; Tisch et al., 1992).

The core of this flow occurs at 200–300 m depth, but has been recorded as deep as 1500 m (Pierce et al., 2000; Noble and Ramp, 2000). The undercurrent is narrow (10–40 km), occurring inshore (25–40 km off the shelf break) of the much broader California Current (Noble and Ramp, 2000). Current measurements have recorded continuous flow along the California slope for distances between 400 and 1000 km (Collins et al., 1996; Garfield et al., 1999; Pierce et al., 2000). Undercurrent flow of up to 2.5 Sv has been recorded throughout the year although with significant seasonal variability, and appears unaffected by local phenomena such as upwelling. Instead undercurrent fluctuations off the Pacific Northwest and northern California have been correlated with fluctuations in alongshore pressure gradients (Largier et al., 1993). Although the origin of seasonal alongshore gradients is not well known, the effect of the pressure gradient on currents has been documented (Hickey, 1998). Maximum flow occurs in summer to early fall, when sea level pressure is low along the northern North American margin (Werner and Hickey, 1983). Alternately very strong California Undercurrent flow is recorded during El Nino events when trade winds slacken, the eastern tropical Pacific sea level pressure is anomalously high, and California Current flow unusually weak (Chelton et al., 1982). 1.3. Intermediate water masses in the Pacific The vertical salinity structure of the Pacific demonstrates the presence of several intermediate water masses. North Pacific Intermediate Water (NPIW) is defined as the subsurface salinity minimum centered on the 26.80s density surface (Talley, 1991). Ventilation of NPIW occurs both in the Bering Sea (Takahashi, 1998) and Sea of Okhotsk (Talley, 1991) through sea ice formation and vertical mixing, as well as in the Alaskan Gyre where shoaling of NPIW occurs close to the base of the mixed layer ( VanScoy et al., 1991). Geostrophic flow along this density surface is generally equatorward except in the subtropical eastern Pacific where the flow turns westward at 201N, leaving a large area of the eastern tropical Pacific not directly renewed by intermediate water formed in the north Pacific (Fig. 6c) (Reid, 1965; Fine et al., 2001). In the western Pacific, however, NPIW is also identifiable to 3.581N, flowing eastward within the North Subsurface Countercurrent (McCreary and Lu, 2001). Antarctic Intermediate Water forms another fresh tongue (AAIW; core density B27.2–27.3s), circulating around the South Pacific subtropical gyre, extending to about 1.581S (Tsuchiya, 1981). Some AAIW flows through the equatorial zone via western boundary undercurrents in the Papua New Guinea region, before spreading back eastward in the Northern Hemisphere tropics. This process results in substantial decrease of

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AAIW tracer properties (Oudot et al., 1999), such as its distinctive low salinity character. However, it is still detectable in tongues of relatively high oxygen that cross into the North Pacific as far north as 6.51N (Reid and Mantyla, 1978; Tsuchiya and Talley, 1996; Reid, 1997; Wijffels et al. 1998). One final distinctive intermediate water mass in the tropics is associated with the Equatorial Undercurrent (Tsuchiya, 1981). This water mass consists of lower-thermocline (upper intermediate) water with a density range similar to that of NPIW. It is formed by subduction at southern mid-latitudes (Tsuchiya, 1981; McCartney, 1982; Toggweiler et al. 1991) and is a mixture of subantarctic mode water and thermocline water referred to as South Pacific Mode Water (SPMW). In this contribution we will refer to eastern tropical north Pacific intermediate waters containing the last traces of AAIW and SPMW (Tsuchiya, 1991) as ‘southern component’ intermediate water.

2. Methods Ocean Drilling Program (ODP) Hole 893A, drilled in Santa Barbara Basin, California at 576.5 m water depth, recovered 196.5 m of sediment covering 160 ka to present. 840 benthic stable isotope analyses have been produced from 718 samples from the interval 10–60 ka in ODP Hole 893A adding to the 69 analyses already published (Kennett, 1995). Isotopic analysis was conducted on several different benthic species because major, rapid fluctuations in the oxygen content of the benthic environment resulted in complete changes in benthic assemblages. Benthic foraminifera (Uvigerina spp., Bolivina tumida, B. spissa, B. argentea, Buliminella tenuata, Rutherfordoides pacifica; 4–25 specimens per sample) were picked for stable isotopic analysis on a Finnigan/MAT 251 light stable isotope mass spectrometer using standard preparation techniques. Where possible, multiple analyses were undertaken using different species within the same sample. Instrumental precision is o0.09% for oxygen and carbon isotopes, with all data expressed as standard d notation in % relative to the Pee Dee belemnite, related by repeated analysis of NBS-19 and -20. A new age model has been developed for the upper 85 mbsf of ODP Hole 893A based on linear interpolation between calibrated 14C datums in the upper 45 mbsf. Below 45 mbsf, climatic events identified in both the ODP Hole 893A planktonic record and the annual layer counted GISP2 provide datums for linear interpolation (Hendy et al., 2002). A constant surfaceocean reservoir correction of 633 years is used for the calibration of 14C age to calendar years (the sum of the 400 year global surface water reservoir-age correction (Stuiver and Braziunas, 1993) and a regional reservoirage correction (DR) of 233760 years (Ingram and

675

Southon, 1996)). Radiocarbon dates younger than 22,000 14C age were converted to calendar ages using the reservoir-corrected age following Stuiver et al. (1998). Radiocarbon ages older than 22,000 14C age were converted to calendar ages using Bard et al. (1998). For the period 30–60 ka BP synchroneity has been assumed between Santa Barbara Basin surface water (d18O) and Greenland air temperature (d18O) so that initiations and terminations of interstadials recorded in the surface waters of Santa Barbara Basin have been tied to those in the GISP2 record.

3. The Santa Barbara Basin benthic isotope record The benthic d18O record (576.5 m water depth) reveals distinct oscillatory behaviour in association with climate change that resembles the Greenland Ice Sheet d18O record (GISP2) rather than Antarctic (Byrd; Fig. 1). The most negative benthic d18O shifts during MIS 3 are associated with interstadial episodes and show remarkable similarity to records of lamination preservation and SST. Although the extent of variability between stadials and interstadials was up to B0.75%, shifts of B0.5% can be seen at every initiation and termination (Fig. 1). Variability of benthic d18O not associated with stadial– interstadial climate change is also apparent (o0.3%) especially during interstadial intervals. Although this may be in part due to changes related to the absence of bioturbation during interstadials, it may also be related to highly variable intermediate water characteristics. Most significant discrete data point decreases in benthic d18O are associated with decreases in the planktonic d18O record and/or peaks in the N. pachyderma coiling ratio (i.e. the two largest d18O decreases during IS 7; Fig. 2), although the magnitude of these discrete events differs between records. Benthic d18O decreases at the beginning of interstadials were recorded in two benthic foraminifers (Uvigerina and Buliminella) and varied in duration between 100 and 400 years (Fig. 2). At the end of interstadials, decreases occurred more rapidly (50–70 years) (Fig. 2). Compared to the Greenland ice records, the benthic record does not show the strong ‘sawtooth’ pattern of cooling through interstadial events. 3.1. Lead, lag or synchronous benthic response? At all interstadial terminations (except IS 6), the increase in both planktonic and benthic d18O occurred abruptly and synchronously. However, the response was dissimilar at interstadial initiations (Figs. 2 and 5; Tables 1 and 2). Detailed comparison between benthic and planktonic records at interstadial initiations does not support a lag but may support either an early or synchronous response in the benthic d18O compared to planktonic depending on how these transitions are

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676

Benthic δ18O (‰ PDB) 3.5

4

3

2.5

0 Holocene

B

C

MIS 1

A

10 Younger Dryas

MIS 2

BØlling/ ÅllerØd

20

Age (Ka BP)

2

3 4

30 5

6 7 8 40 MIS 3

9 11

10 12 13

50

14

15 16/17

60 -42

-40

GISP2

δ18O

-38

-36

(‰ SMOW)

-40

-37.5

-35

Byrd δ18O (‰ SMOW)

Fig. 1. Comparison between (A) GISP2 d18O (SMOW) isotope time-series (Grootes et al., 1993); (B) the benthic paleoclimatic record at ODP Site 893A Santa Barbara Basin and (C) Byrd d18O (SMOW) isotope time-series (Blunier and Brook, 2001) for the last 60 ka. Average benthic foraminiferal d18O values are represented by the light line, while smoothed data with a 5 point running average is represented by the heavy line. Breaks in these lines represent core intervals with insufficient benthic foraminiferal species for analysis. Variability in these parameters clearly define ( the Dansgaard/Oeschger (D/O) climate oscillations (numbers 17-3) during OIS 3 and the Blling/Allerd. Gray bands represent warm intervals (interstadials and the Holocene). Interstadials (D/O events) are numbered according to GISP2 scheme. Chronologic position of the Marine Isotopic Stages (MIS) are shown on left.

I.L. Hendy, J.P. Kennett / Quaternary Science Reviews 22 (2003) 673–689 Rutherfordoides pacifica Buliminella tenuata Bolivina tumida Uvigerina sp. Benthic δ18O 3.6 31

N. pachyderma δ18O

3.35

(a)

677

2.5

3.1

(b)

2

1.5

(d)

(c)

1

0.5

(e)

32

5

33 6 Age (ka)

34 7

35 36 37

8

38 39 4

3

2

0

1

100

Benthic δ18O 3.6

3.4

2

2.5

1.5

G. bulloides

N. Pachyderma coiling ratio (%)

Bioturbation index

(A)

50

1

0.5

δ18O

N. pachyderma δ18O 2.5

3.2

2

1.5

1

0.5

44 12 45 46 47 48 49 14

Age (ka)

50 51 52 53

15

54 55

16

56 57 58 59 4

(B)

3

2

1

Bioturbation index

0

50

100

N. Pachyderma coiling ratio (%)

2.5

2

1.5

1

0.5

G. bulloides δ18O

Fig. 2. Comparison between benthic and planktonic paleoclimatic records at ODP Site 893A Santa Barbara Basin from (A) 40 to 30 ka and (B) 60 to 45 ka. These time intervals represent the most complete sections of the core with the greatest climate variability. (a) The bioturbation index (bold), where on a continuum, 1 indicates laminated sediment facies and 4 indicates massive bioturbated sediment facies (Behl, 1995). (b) Benthic foraminiferal d18O (PDB) with species indicated, the fine line represents averaged data. The benthic foraminiferal species are as follows; Rutherfordoides pacifica (shaded diamonds), Buliminella tenuata (open squares), Bolivina tumida (open triangles) and Uvigerina ssp (black circles). (c) The relative abundance of dextral to sinistral coiled N. pachyderma (shown by % dextral N. pachyderma) (Hendy and Kennett, 2000). (d) d18O (PDB) records of thermocline planktonic foraminifera N. pachyderma, and (e) surface water planktonic foraminifera G. bulloides (Hendy and Kennett, 1999; Hendy and Kennett, 2000). Interstadials (light grey) are numbered according to GISP2 scheme and defined by the planktonic response. Dark gray bands represent the time between benthic d18O decrease and indications of surface temperature warming (both faunal and geochemical evidence).

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Table 1 r2 results from x2y plots comparing two planktonic species with the benthic average at the initiation (53.50–55 m) and termination (52.5–54 m) of IS 7 Steps moved

5 4 3 2 1 0 +1

r2 Results for the Initiation of IS 7

r2 Results for the Termination of IS 7

G. bulloides relative Benthic average to N. pachyderma relative to G. bulloides

Benthic average relative to N. pachyderma

G. bulloides relative Benthic average to N. pachyderma relative to G. bulloides

Benthic average relative to N. pachyderma

— — 0.37 0.47 0.65 0.93 0.73

0.42 0.54 0.57 0.57 0.33 0.30 0.11

— — 0.42 0.54 0.57 0.57 0.33

— — 0.1 0.02 0.06 0.41 0.23

0.39 0.61 0.50 0.45 0.31 0.31 0.06

— — 0.17 0.14 0.1 0.46 0.06

Steps moved represents the number of sample intervals that one record was offset from the other.

Table 2 r2 results from x2y plots comparing the two planktonic species with the benthic average at the initiation (76.50–78 m) and termination (72.5–74 m) of IS 14. Steps moved

2 1 0 +1 +2

r2 Results for the Initiation of IS 14

r2 Results for the Termination of IS 14

G. bulloides relative Benthic average to N. pachyderma relative to G. bulloides

Benthic average relative to N. pachyderma

G. bulloides relative Benthic average to N. pachyderma relative to G. bulloides

Benthic average relative to N. pachyderma

0.38 0.23 0.61 0.08 0.26

0.29 0.42 0.29 0.37 0.20

0.07 0.22 0.34 0.12 0.08

0.41 0.35 0.50 0.20 0.03

0.47 0.56 0.53 0.54 0.35

0.43 0.41 0.42 0.23 0.003

Steps moved represents the number of sample intervals that one record was offset from the other.

defined. The initiation of an interstadial could be defined as the first small decrease (0.05–0.5%) in planktonic d18O prior to most warm events. Based on this definition then the benthic d18O record is mostly synchronous with the planktonic record except at IS events 8 and 12 where there is an apparent lag. A characteristic of this definition is that a greater proportion of the benthic d18O shift occurs before the planktonic shift. Thus, prior to the most rapid jumps of surface water warming associated with interstadials, 30–100% of the benthic d18O decrease has occurred compared to only 8–40% of the planktonic d18O decrease. Although it may be argued that the minor planktonic and benthic d18O decreases should be considered equal responses to climate forcing, a counter-argument would state that as the planktonic response had not yet exceeded stadial variability the initiation of the interstadial had yet to occur. Thus, a second definition of an interstadial initiation would be the time of the first major shift in planktonic d18O outside stadial variability accompanied by a jump from sinistral to dextral N. pachyderma. In this case detailed examination of the planktonic and benthic

records reveals a difference in response to rapid climate change between surface and intermediate waters. A slow decrease in benthic d18O began 60–200 years before abrupt (50–70 years) change in surface water proxies occurred (demonstrated by both planktonic d18O and the ratio of sinistral to dextral N. pachyderma; Fig. 2 and 5). A statistical solution would be preferable. However, the data sets do not lend themselves well to statistics. A simple solution is to produce for the short intervals over the interstadial initiations and terminations, x2y plots of the various records and compare r2 values as the records are offset from each other stepwise (each step being equivalent to one sample interval). Several restrictions disqualify certain events, i.e. where the results may be biased by the existence of one very light benthic d18O value, and where one of the data sets contains a significant (more than one sample) data gap due to paucity of specimens. Thus, two interstadial events qualify for this approach. The results (Tables 1 and 2) support three observations previously discussed. First the high correlation between planktonic species at initiations and terminations decreases dramatically

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when the records are offset demonstrating the synchroneity of these records. Secondly no offset is required to improve the correlation between all data sets at the terminations (except IS 14 between benthic and G. bulloides; Table 2) as the benthic and planktonic records were also synchronous during these intervals. Finally, in contrast, at the event initiations one to four step offsets of the benthic record toward younger ages or shallower depths are required to produce the strongest correlation between the benthic and planktonic records. This supports the hypothesis that the benthic record led the planktonic records at interstadial initiations. Additional examination at higher resolution (B1 cm) over the rapid climate event transitions would enhance both the magnitude of change and temporal relationship between surface and benthic proxies. Stratigraphic integrity, however, over these intervals is now poor due to sample depletion and possible post-drilling carbonate dissolution (gypsum growth driven by pyrite oxidation under the cool moist conditions of core storage).

3.2. Benthic interspecies d18O differences Significant d18O shifts do not appear to be the result of benthic foraminiferal interspecies vital effects based on the following observations. Multiple species were analyzed in 159 samples (22% of the total), allowing for comparisons between species. These comparisons demonstrate that mean d18O differences between species ranged between 0.019 and 0.121, with standard deviations varying between 0.09 and 0.219 (Table 3). However, it is unlikely that species preferring oxic conditions lived contemporaneously with those preferring dysoxic conditions. The presence of a very large porewater oxygen gradient existing under conditions of enhanced carbon rain and oxidation might explain the presence of oxic species at the sediment–water interface and subsoxic species at depth (Stott et al., 2002a). However, from a much smaller dataset, rose-bengal stained epifaunal B. argentea and infaunal B. tenuata from Santa Barbara and Santa Monica Basins show d18O differences of 0.01 (70.17; Stott et al., 2002a).

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Another explanation for the existence of oxic and dysoxic species within the same sample might be that Uvigerina specimens collected from a sample (representing 14 year of sedimentation) dominated by bolivinids may have grown during a long-lived (>1 year) flushing event within the basin. Though modern flushing events are associated with increased oxygen content and decreased bottom water temperatures and salinity (Sholkovitch and Gieskes, 1971) they are also of brief duration (o 1 month; Hickey, 1991). In this case, Uvigerina specimens would be recorders of a cooler (heavier d18O) bottom water mass in the basin compared to the bolivinids. It is not surprising therefore that greatest mean difference and standard deviations occurred between species (e.g. Uvigerina and B. tumida) known to tolerate very different environmental conditions. We suggest these differences largely reflect the environmental effects of different water masses bathing the site such as described by Sholkovitch and Gieskes (1971) rather than vital effects. Another statistical approach to determine if interspecies offsets occur is to compare the median and range of the data for each species (Fig. 3). However, as individual species are clustered during time intervals when bottom water environmental conditions were favourable for their growth, long-term trends such as global ice volume must be removed from the data. A box-and-whisker plot was made of the detrended results for each species (B. spissa was excluded as only two samples existed during this time interval; Fig. 3). Fig. 3 demonstrates that none of the median d18O values of the different species vary more than 0.1% from each other. Once again species preferring oxygenated conditions record slightly heavier d18O values than dysoxic and suboxic species, although the range of d18O values of all species (except Buminella) falls within those of Uvigerina spp. (the only species that existed in the basin that is commonly used elsewhere in paleoceanographic studies). Fig. 3 shows that although d18O variability is large all the measured species have a similar d18O range. The most compelling evidence that benthic d18O shifts seen in ODP Hole 893A are independent of interspecies d18O offsets due to vital effects during calcification occurs at the onset of interstadials. At these times

Table 3 Comparison of the d18O differences with standard deviations between benthic foraminifera from ODP Hole 893A, which occur within the same sample

Uvigerina spp. Buliminella tenuata Rutherfordoides pacifica Bolivina argentea Bolivina spissa

Bolivina tumida

Bolivina spissa

Bolivina argentea

Rutherfordoide pacifica

Buliminella tenuata

0.12170.219 0.10170.119 0.137 0.05570.156 0.08170.205

0.06070.170 0.11770.097 n/a 0.07970.130

0.02370.090 0.12570.068 0.093

0.01970.136 0.08370.155

0.06270.167

Duplications of more than one species were run on 159 samples, and more than two species were run on 47 samples.

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680

Uvigerina spp.

Rutherfordoides pacifica

0.2

Buliminella tenuata

0.3

Bolivina argentea

Difference between benthic and normalised δ18Oxygen (‰)

0.4

Bolivina tumida

0.5

0.1

0.0

0

-0.1

50

100

Number of analyses -0.2

-0.3

-0.4

-0.5

Dysoxic

Suboxic

Oxic

Fig. 3. Box and whisker plot of the detrended benthic d18O values (subtraction of individual benthic d18O values from polynomial fit) for the 6 benthic species analyzed in this study. Values greater than zero represent isotopic values heavier than the trend, and values less than zero are lighter than the trend. The box and whisker plot is a graphical display of a five number summary. The shaded box represents values falling between the 25 and 75 percentile with the heavy line in the middle of the box representing the median. The width of the boxes represents the number of data points comprising the box plot. The thin lines drawn away from the box represent the maximum and minimum values. The circles represent outliers (individual data points more than 1.5 box lengths). To remove long-term trends a polynomial curve was fit to the average benthic data (r2 of 95%). A poor fit of the curve during deglaciation resulted from rapid shifts in benthic d18O. To avoid this interval only data between 20 and 65 ka was used.

simultaneous d18O decreases generally occurred in several different oxic and suboxic taxa (including Buliminella tenuata & Uvigerina spp.), after which d18O values of dysoxic species (B. tumida) continued with similarly light values. Yet when the oxic species (Uvigerina spp.) reappeared at event terminations, they recorded d18O values similar to the prior stadial and higher than those of the same species at the onset of interstadials (Fig. 2). For example, the d18O decrease recorded both by B. tenuata (0.35%) and Uvigerina spp. (0.3%) at IS 14 was followed by similar d18O values recorded by B. tumida with variability almost as great as the stadial-interstadial shift (Fig. 2). Thus, benthic d18O variability within ODP Hole 893A must reflect environmental conditions rather than interspecies vital effect. 3.3. Ice volume effects It is possible that changes in benthic d18O seen at ODP Hole 893A result from global ice volume changes.

In this contribution we will use the conservative estimate of 1.1% (Chappell and Shackleton, 1986) for total global ice volume effect and a maximum sea level drop of 130 m (Fleming et al., 1998; Peltier, 1998) for the last glacial. Hence we estimate a 0.008% shift in d18Ow for every meter of sea level change. Changes in global ice volume were muted during interstadial warmings. Recent sea level estimates suggest 10–15 m of change occurred during the Bond cycles of MIS 3 (except IS 17 at the MIS 3-4 boundary; Chappell, 2002). This implies that only 0.09 to 0.13% of the benthic d18O shift during these events can be attributed to injection of meltwater into the global ocean assuming constant ice d18O values. There are several notable features in the ODP Hole 893A benthic d18O record at the last termination (Fig. 4). First, the full shift in benthic d18O from MIS 2 to the Holocene was 1.65%. When ice volume effects are removed the remaining 0.55% suggests at least a 2.31C temperature shift (assuming no salinity change; Epstein et al., 1953) occurred in bottom waters at the

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0.4‰

0.3‰

0.52‰

681

(total ice volume) =1.22‰

0

MIS 1

Holocene

Younger Dryas warming

10 Age (Ka BP)

0.55‰ 0.5‰

Younger Dryas cooling

Younger Dryas BØlling/ÅllerØd

0.4‰ MIS 2

BØlling warming Total benthic δ18O shift = 1.65‰

20

IS 2

4

3.5

3 Benthic

2.5

δ18O

Fig. 4. The benthic paleoclimatic record at ODP Site 893A Santa Barbara Basin for the last 25 ka. Average benthic foraminiferal d18O values are represented by the light line, while smoothed data with a 5 point running average is represented by the heavy line. Breaks in these lines represent core ( intervals with insufficient benthic foraminiferal species for analysis. Dark gray bands represent warm intervals ( the Holocene and Blling/Allerd). Chronologic position of the Marine Isotopic Stages (MIS) are shown on left. Light gray boxes represent intervals of gradual d18O increase, while arrows represent magnitude and direction of rapid d18O shifts over deglaciation. Bold numbers represent total d18O value change represented by both bars and boxes.

site (Fig. 4). Second there were two abrupt decreases in d18O of 0.4–0.55% at the beginning of the Blling, and the termination of the Younger Dryas, as well as an abrupt increase (B0.5%) at the beginning of the Younger Dryas (Fig. 4). Finally, there were gradual decreases in benthic d18O from 17 to 14.7 ka, 14.6 to 12.8 ka and 11.2 to B7 ka (Fig. 4). Combination of the gradual positive d18O shifts over these three intervals is similar in magnitude (1.22%) to full ice volume (1.1%; Chappell and Shackleton, 1986), before the abrupt d18O shifts are even included. Therefore, these abrupt shifts at the termination can only be explained by temperature and/or salinity changes in bottom waters at the site. Nevertheless, estimating the d18Ow decrease in the global ocean resulting from meltwater injection is complicated by the time it takes for the ice volume

signal to arrive at the site. During intervals of ice sheet melting, the ocean below the surface mixed layer (upper 50 m) was unlikely to be isotopically homogenous. The d18O-depleted meltwater probably initially influenced the Atlantic Basin, as it was closest to the outflow of former ice sheets. A number of authors have suggested that the deep Pacific d18O record could have lagged Atlantic records significantly (B1000 years), depending on conveyor circulation rates (Broecker et al., 1988; Bard et al., 1991; Duplessy et al., 1991). Mix et al. (1999) suggest deep-sea warming in the Pacific preceded the arrival of meltwater pulse 1A in the region (as dated by Bard et al., 1990). In this scenario a meltwater signal propagating from the North Atlantic would arrive at Santa Barbara Basin well after the initiation of the rapid warming event. As most interstadial events are relatively brief (o1000 years), it is unlikely that a meltwater signal

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could arrive in Santa Barbara Basin in time to play a significant role in the benthic d18O shifts during MIS 3. Another possibility is that the meltwater d18O signal was instantaneously transported through the atmosphere via precipitation. Such a process is feasible because complete exchange of the upper 50 m of the tropical ocean with the atmosphere may occur within 100 years (Anderson and Thunell, 1993). However, it would take much longer to exchange the upper 50 m of the tropical ocean with intermediate waters. This process would occur from the ‘top-down’, so that the surface waters would receive isotopically light d18Ow prior to deeper waters, and thus ice volume effects in intermediate water should lag those in surface waters. The benthic d18O shifts during interstadial events appear to have occurred synchronously or earlier than surface water d18O change and hence could not have resulted from atmospherically driven meltwater injections. 3.4. Temperature and salinity We suggest the d18O shifts were produced mainly by temperature and in part by salinity. Assuming no salinity change occurred in intermediate waters, then changes in benthic d18O imply an average 1–21C temperature shift (based on the Shackleton, 1974 equation for Uvigerina peregina) between stadials and interstadials. However, these shifts coincide with evidence for decreased ventilation of intermediate waters, which in turn indicates changes in water mass and possibly therefore salinity (Behl and Kennett, 1996; Keigwin, 1998; Cannariato et al., 1999). If a wellventilated intermediate water mass entering the basin had characteristics similar to intermediate water formed in the Alaskan Gyre (salinity34%, d18Ow=0.2%; Epstein and Mayeda, 1953) or the Sea of Okhotsk (salinity = 33.7%, d18Ow = -0.4%, temperature= 2.91C; Schmidt et al., 1999), absolute bottom water temperatures would range between 31C and 41C (Shackleton, 1974). Similarly, assuming poorly ventilated intermediate waters during interstadials were the more saline (34.5%; d18Ow=0.04%, temperature=8.31C; GEOSECS Ostlund et al., 1987) ‘southern component’ water mass, then absolute temperatures during interstadials were B6.51C (similar to present day; Shackleton, 1974). The maximum modern d18Ocalcite change that can be attributed to shifting the source of intermediate water from a ‘southern source’ to the Okhotsk Sea is 1.24% for a 5.41C temperature difference (assuming 0.23% isotopic shift=1 1C temperature change; Epstein et al., 1953). However, the maximum d18Ocalcite change will be less than this (0.8%) due to the d18Ow differences between the two sources. Although it is known that the d18Ow (as well as salinity and temperature) of these two end-member water masses (NPIW; Keigwin, 1998;

AAIW; Lynch-Stieglitz et al., 1994) changed during the last glacial, it is unlikely differences between the North and equatorial Pacific were reversed. Therefore, temperature shifts between stadials and interstadials might have been as much as 3.51C when the maximum potential salinity/d18Ow differences between NPIW and ‘southern component’ water masses are incorporated. 3.5. Bioturbation It might be argued that the differences between planktonic and benthic d18O shifts (Figs. 2 and 5) resulted from bioturbation. The presence of laminations indicates an absence of bioturbation and hence the integrity of stratigraphy during interstadial events. During bioturbated intervals, disruption of the sediment fabric by burrows was observed to occur no deeper than 3 cm (R. Behl, pers. com.), much less than the 7 cm sampling resolution of this study. Studies of living benthic foraminiferal assemblages in Santa Monica Basin have demonstrated that B. argentea inhabits in the top 0.1 cm, while B. tenuata is most abundant from 0.3 to 0.6 cm depth and thus vertical migration within the sediment is insignificant (Stott et al., 2002a). Fig. 5 demonstrates that the lead in benthic warming prior to interstadial events occurs over a 20 cm interval at the initiation of IS 14 and over a 30 cm at IS 12. In contrast, the d18O increase in both planktonic and benthic records occurs abruptly in less than 7 cm at the interstadial terminations. As bioturbation at interstadial terminations destroys, and dysoxic conditions at initiations preserves, the integrity of the sediment below, gradual changes produced by bioturbation should occur at terminations not initiations where they are observed. Furthermore, the slow total benthic d18O decrease at the beginning of IS 14 occurs over an interval of 56 cm, while the rapid planktonic d18O decrease occurs over only 7 cm (Fig. 5). Mixing by bioturbation should produce similar patterns in both planktonic and benthic records.

4. Discussion The Santa Barbara Basin benthic d18O record demonstrates that intermediate water temperatures were unstable on a submillennial time scale, supporting the conjecture that shallow thermohaline circulation switches occurred in the North Pacific during the last glacial (Behl and Kennett, 1996; Van Geen et al., 1996; Keigwin, 1998; Cannariato and Kennett, 1999). The absence of a distinct sawtooth pattern, such as that observed in the Greenland ice cores, is similar to other records of submillennial climate change where bimodal changes occurred in either atmospheric or oceanic circulation (Sarnthein et al., 2000). This strongly

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Depth (meters)

53

54

Benthic δ18O (‰) 3.4 3.2

55

73

G. bulloides δ18O (‰) 2 1.5 1

2.5

Interstadial 7

3.6

683

3

3.6

Benthic δ18O (‰) 3.4

2.

(a)

2.5 2 1.5 N. pachyderma δ18O (‰) 3.2

1

G. bulloides δ18O (‰) 2 1.5

2.5

1

(c)

(b)

75

76

Interstadial 14

Depth (meters)

74

77

78 3

2.5

2

1.5

N. pachyderma

δ18O

1

(‰)

Fig. 5. Comparison between benthic and planktonic paleoclimatic records at ODP Site 893A Santa Barbara Basin from (1) 52.5 to 55 m and (2) 72.5 to 78 m. These core sections represent Interstadials 7 and 14 and demonstrate the depth interval over which the benthic-planktonic lead/lag occurs. (a) Benthic foraminiferal d18O (PDB) with species indicated, the fine line represents averaged data. The benthic foraminiferal species are as follows; Bolivina argentea (shaded squares), Buliminella tenuata (open squares), Bolivina tumida (open triangles) and Uvigerina ssp (black circles). (b) d18O (PDB) records of thermocline planktonic foraminifera N. pachyderma, and (c) surface water planktonic foraminifera G. bulloides (Hendy and Kennett, 1999, 2000). Interstadials (light grey) are numbered according to GISP2 scheme and defined by the planktonic response. Dark grey bands represent the time between benthic d18O decrease and indications of surface temperature warming (both faunal and geochemical evidence).

supports the hypothesis that Santa Barbara Basin intermediate water temperature change was driven by switches in intermediate water mass origin. The cool, fresh, well-oxygenated water mass bathing ODP Hole 893A (Fig. 6a) during stadials was probably

North Pacific Intermediate Water (NPIW). Expansion of NPIW in the North Pacific may have been a result of either increased production of NPIW and/or decreased northward advection of ‘southern component’ water.

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4.1. Processes driving NPIW production Much potential exists for increased intermediate water formation in the North Pacific. Presently seasonal sea ice formation in the Sea of Okhotsk results in the production of dense, saline, cold water (Honjo et al., 1996; Takahashi, 1998). When the Siberian High dominates the region, resulting geostrophic winds transport cold air from the Siberian interior causing extensive sea ice formation (Parkinson, 1990). Production of intermediate water through sea ice formation may be affected by both changes in strength of the Siberian High or freshwater flow into the Sea of Okhotsk. Shutdown of North Atlantic thermohaline circulation has been shown through climate modeling to intensify cool winter Siberian airflow over the North Pacific. This extraction of North Pacific surface ocean heat produced a tongue of cool fresh water (up to 3 Sv) at intermediate depths in the North Pacific (Mikolajewicz et al., 1997). 4.2. Processes driving ‘southern component’ water In addition to expansion of NPIW, contraction of the ‘southern component’ water mass may have occurred. Kennan and Lukas (1996) suggested that high salinity water interlaced with modern NPIW near Hawaii is related to variability in the northern extent of Antarctic Intermediate Water (AAIW). The ability of AAIW to cross the equator would have effected intermediate water temperatures at ODP Hole 893A. Northward flow of AAIW across the equator may have been impeded by increased thermoclinal depth in the western equatorial Pacific (Andreasen and Ravelo, 1997) produced by possible stronger glacial tradewinds (Molina-Cruz, 1977; Pedersen, 1983; Sicre et al., 2000). However, if trade winds during cool intervals of the last glacial were weaker as suggested by some recent paleoceanographic studies (Little et al., 1997; Koutavas et al., 2002; Stott et al., 2002b) then flow through the equator would not have been restricted, although residence time would have increased. Reduced contribution of North Indian Intermediate Water to AAIW and/or changes in the production of AAIW at the Subantarctic Front (LynchStieglitz et al., 1994) may have been another factor in the northward extent of AAIW in relation to NPIW. Transport of AAIW over the equator into the North Pacific is also associated with Indonesian seaway outflow strength (Fine, 1985; Hirst and Godfrey, 1994). This outflow (5–15 Sv) is a sink of NPIW entering the tropical ocean (Bingham and Lukas, 1994; Fine et al. 1994) and is balanced by an inflow of AAIW into the North Pacific equatorial region. Modeling suggests that when outflow through the Indonesian seaway is greater than 4 Sv most tropical NPIW exits the North Pacific via the outflow, and is replaced by SPMW and AAIW

(3.8 Sv), which crosses the equator to fill the northern tropical ocean (McCreary and Lu, 2001). If outflow is small (>4 Sv) most tropical NPIW recirculates in the northern Tropics, no SPMW enters the region and AAIW transport into the Northern Hemisphere is an order of magnitude less (0.37 Sv; McCreary and Lu, 2001). Restricted flow through the Indonesian seaway due to lowered sealevels, particularly during MIS 2, would have reduced AAIW transport into the North Pacific. However, sea level changes between interstadials and stadials would probably have occurred too slowly (Chappell, 2002) to have produced the intermediate water mass shifts seen in Santa Barbara Basin. Presently ‘southern component’ water is transported poleward by an undercurrent. This current is driven by changes in alongshore pressure gradients or sea level height on the North American Margin, such as may occur with changes in tradewind strength. If a steeper thermoclinal gradient across the Pacific (Andreasen and Ravelo, 1997) and intensified trade winds occurred during cool intervals of the last glacial, relatively low sea levels may have resulted in the eastern tropical Pacific. Diminished coastal upwelling on the California Coast during stadial events (Hendy et al., in prep) (higher coastal sea levels) and a weaker California Current (Doose et al., 1997; Herbert et al., 2001) (reduced transport of surface water into the Eastern Tropcial Pacific) would also have contributed to a longshore sea level pressure gradient unfavourable to undercurrent flow. A weaker California Undercurrent during the last glacial has been suggested by Kienast et al. (2002). An alternate hypothesis remains that the benthic d18O shifts seen at ODP Hole 893A were not produced by temperature, but rather by the increased contribution of an intermediate water mass with a high d18Ow. If ‘southern component’ intermediate water maintained or increased present d18Ow values (excluding ice volume), was significantly cooler, and was characterized by high oxygen concentrations, then the benthic d18O results might be explained by a hypothesis of increased California Undercurrent flow. Stronger undercurrent flow during the cool intervals of the last glacial might result from weaker trade winds (Little et al., 1997; Koutavas et al., 2002; Stott et al., 2002b) raising relative sea level in the eastern tropical Pacific, coupled with stronger California Current flow (Hendy and Kennett, 2000) (increased transport of surface water into the Eastern Tropical Pacific) combining to create a favourable longshore pressure gradient. 4.3. Phasing of climate change responses Temporal differences between the benthic and planktonic temperature records at ODP Hole 893A may suggest the possible existence of lead/lags within the surface-ocean, and intermediate water response

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(B500 m) to climate forcing at interstadial transitions (Fig. 2). The most abrupt changes in Santa Barbara Basin surface water temperature (Hendy and Kennett, 1999, 2000) demonstrate close interdependence between atmospheric and surface ocean circulation and have been linked to contraction of the Aleutian Low and repositioning of the North Pacific High with associated jet streams in response to the D/O event climate forcing (Hendy and Kennett, 1999). If intermediate water warming led (up to 200 years) rapid Santa Barbara Basin surface water warming, then changes in intermediate water masses bathing the site preceded the rapid North Pacific atmospheric reorganization that strongly affected surface water temperature at interstadial initiations (Hendy and Kennett, 1999, 2000). Any process affecting NPIW temperatures related to atmospheric reorganization could not have driven intermediate water warming (such as sea ice formation), otherwise Santa Barbara Basin surface and bottom water warming would have been synchronous. In addition, intermediate water temperatures increased gradually relative to the abrupt surface water warmings (60–95% of the planktonic d18O shift occurs in less than 70 years) at onsets of interstadials (as defined by abrupt changes in planktonic d18O and fauna). Intermediate water warming continued for several hundred years after rapid climate change had occurred, requiring a gradual warming/return to ‘southern component’ intermediate water properties. This contrasts with the interpretation of abrupt synchronous cooling in both intermediate and surface water at the interstadial terminations requiring rapid NPIW expansion. Inferred changes in oxygen content of Santa Barbara Basin bottom water were also synchronous with surface water warming suggesting a decoupling between the oxygen concentration and temperature of bottom waters at this time. There are a number of paths by which intermediate waters could respond to climate change forcing as mentioned in the previous section. If initial intermediate water response truly leads surface water warming, it must have occurred in the absence of any significant direct North Pacific atmospheric teleconnection as seen in the surface water response. If intermediate water warming was synchronous with surface water warming, then atmospheric teleconnections need to be considered. Although these observations alone cannot verify any single mechanism, they may help constrain an explanation. 4.4. Freshening of the North Pacific Freshening of North Pacific surface waters may have resulted in NPIW contraction, either occurring because of changes in the North Pacific hydrological cycle or as the result of ice sheet melting. Cooler, drier conditions

685

recorded in the subarctic Pacific during the last glacial (Sabin and Pisias, 1996; de Vernal and Pedersen, 1997) resulted from the intensification of the Aleutian Low and southward deflection of the jet stream. A contraction of the Aleutian Low may increase precipitation and decrease NPIW production (Andreev and Kusakabe, 2001), but would also effect the North Pacific High and consequently surface waters in Santa Barbara Basin. Disruption of North Atlantic thermohaline circulation would also affect the precipitation–evaporation balance between oceans through atmospheric water vapour transport patterns (Manabe and Stouffer, 1988). Resulting reduction of North Atlantic salinity and salt accumulation in North Pacific is inferred to strengthen North Pacific thermohaline circulation. However, meltwater discharge in the North Atlantic (and subsequent reduction in evaporation) during Heinrich events (van Kreveld et al., 2000) would in effect reinforce NPIW production rather than slow it, and therefore cannot be the process causing early intermediate water warming in Santa Barbara Basin. It is possible that only minor amounts of meltwater would be required to prevent shallow thermohaline circulation in the North Pacific, resulting in contraction of NPIW. Meltwater runoff from either southern Alaska, the Aleutian Islands or north of the Bering Strait, has been suggested by pulses of ice rafted detritus (IRD) (Kotilainen and Shackleton, 1995) and d18Olocal minima (McDonald et al., 1999). If North Pacific meltwater inputs resulting in NPIW contraction were coeval with ice sheet collapse in the North Atlantic (Van Kreveld et al., 2000), it is possible that high latitude cyrospheric events led the climatic warming at interstadials. In this scenario the greatest North Pacific meltwater inputs should have occurred during the most pronounced ice sheet collapses (such as H1 and H2), although, possible lead/lags were less distinct during these intervals. Thus, the behaviour of surface and intermediate water masses (see Fig. 6), and the Northern Hemisphere cryosphere suggests that any relationship between NPIW contraction and high latitude processes was complex. 4.5. Tropical forcing of intermediate waters Another hypothesized mechanism complies with the following four aspects of intermediate water response to rapid climate change assuming benthic d18O shifts lead the reorganization of North Pacific atmospheric circulation. First the speed and magnitude of bottom water temperature change in Santa Barbara Basin can be produced through both expansion and contraction of NPIW and by changing the transport mechanism to the site (as described in previous sections). Secondly the slow B140 year lead of intermediate water warming compared to the rapid warming of surface waters can be

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686

(A) Stadials

ODP 893A

?

NPIW

?

? AAIW

(B) ~140 years before Interstadial onset

ODP 893A NPIW

?

?

?

AAIW (C) Interstadials

NPIW

AAIW

ODP 893A 'S o I n ut h ter e r me n C dia om te po Wa ne ter nt

Fig. 6. Maps of the North Pacific representing the different intermediate water masses on the 27.807 st (g/l) density surface for the following time slices: (A) Stadial or cool intervals; (B) B140 years before initiation of Interstadials; and (C) Interstadial or warm intervals based on geochemical properties (based on Reid, 1965). This density surface is found at depths of 400 m along the margin of North America, but may be as deep as 800 m in the western Pacific and as shallow as 200 m in the northwestern Pacific (Reid, 1965). Two sources of intermediate water in the North Pacific are displayed: In dark grey, Antarctic Intermediate Water (AAIW) on a circuitous route through equator creates the poleward flowing, poorly ventilated ‘southern component’ water mass (pale grey); and in mid-grey, water ventilated in the northwestern Pacific (North Pacific Intermediate water, NPIW). Flow direction based on acceleration potential (Reid, 1965) is represented by dashed lines and arrows.

explained by commencing northward flow of the California Undercurrent before major atmospheric reorganization over the North Pacific. Perhaps this occurred as a result of a favourable longshore pressure gradient due to diminishing trade wind strength. Thirdly the decoupling of intermediate water oxygenation from temperature may result from the initial water transported north being warm, but not suboxic (Fig. 6b). It follows that although AAIW became warmer and more saline due to mixing at the equator, the undercurrent

flow may not have been suboxic during cool intervals due to suppressed productivity and decreased rain of organic material along the North American Margin (Ganeshram and Pedersen, 1998; Hendy et al, in prep). After initiation of intermediate water warming by global climate forcing, it would appear to take up to B140760 years or less for a threshold to be crossed causing extremely abrupt reorganization of atmospheric circulation over the North Pacific. When rapid atmospheric reorganization occurred, surface waters warmed as subtropical water was transported northward by the Davidson Current. High productivity along the eastern Pacific margin recommenced causing ‘southern component’ intermediate water to become suboxic as the organic material rain rate increased (Fig. 6c). It is possible that the residence time of AAIW at the equator also increased if tradewinds diminished. Fourthly, the rapid and synchronous cooling of intermediate and surface waters at times of abrupt interstadial collapse can be explained by the proximity of Santa Barbara Basin to NPIW. At interstadial terminations, undercurrent flow diminished and expansion of cool, well-ventilated proximal NPIW rapidly bathed the site, assisted by high latitude processes such as the expansion and intensification of the Aleutian Low pressure system (Fig. 6a). Although this mechanism describes only intermediate waters in the eastern North Pacific, it is most likely to have occurred along other eastern ocean boundaries in response to trade winds, suggesting intermediate waters are more responsive to climate change than previously thought. Finally, the possibility that tropical climates play a role in rapid climate change is not inconceivable since it is known that coupled tropical ocean–atmosphere interactions play a crucial role in modulating global climate on interannual, decadal and possibly glacial– interglacial timescales (Bush and Philander, 1999; Lea et al., 2000; Tudhope et al., 2001). In particular, small variations in equatorial Pacific Ocean temperatures are known to produce a trade wind response that affects * global climate (El Nino/Southern Oscillation or ENSO). Recent modeling studies have shown a response in ENSO to both orbital forcing (Clement et al., 1999) and Laurentide Ice Sheet dynamics (Bush and Philander, 1999). Furthermore, tropical paleoclimate studies suggest that there was an atmosphere–ocean–cryosphere linkage through changes in trade wind intensity and zonality derived from the insolation-linked precessional cycle which resulted in submillennial climate variability (McIntyre and Molfino, 1996; Little et al., 1997; Stott et al., 2002b). However, paleoclimatic studies conflict as to whether tropical circulation was invigorated (‘La Nina-like’; Bush and Philander, 1999; [19]) or sluggish * (‘El Nino-like’; Little et al., 1997; Koutavas et al., 2002; Stott et al., 2002b) during the cool intervals of the last glacial.

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5. Conclusions

References

The Santa Barbara Basin benthic d18O shifts appear to represent a continuum between two intermediate water circulation modes (a cool, ventilated mode of NPIW expansion and a warm, low oxygen mode of ‘southern component’ intermediate water dominance). Although intermediate waters are restricted to shallow depths (400–800 m) in the North Pacific, they affect large areas of the continental margin. Depending on how rapid climate change transitions are defined, intermediate water temperatures either lead or are synchronous with surface water temperature change. Intermediate waters appear to lead rapid atmospheric reorganization at or after the initiation of interstadials by 60–200 years. However, at event terminations, decreases in intermediate water temperatures appear synchronous with cooling of surface water. These observations can best be explained by mechanistic differences in intermediate water transport during processes of NPIW expansion and contraction. Cool NPIW was rapidly transported south in the North Pacific gyre system without modification of water characteristics during stadials, while during interstadials ‘southern component’ intermediate water (AAIW modified at low latitudes) was transported north by the California Undercurrent. This scenario supports the possibility through changes in trade wind strength that small changes in meridional or zonal tropical temperature gradients played a role in rapid climate change. This result increases in relevance in light of recent suggestions that intermediate water temperatures may play a role in greenhouse gas amplification of Late Quaternary climate change (Kennett et al., 2000).

Anderson, D.M., Thunell, R.C., 1993. The oxygen-isotope composition of tropical ocean surface water during the last deglaciation. Quaternary Science Reviews 12, 465–473. Andreasen, D.J., Ravelo, A.C., 1997. Tropical Pacific Ocean thermocline depth reconstruction for the last glacial maximum. Paleoceanography 12 (3), 395–413. Andreev, A.G., Kusakabe, M., 2001. Interdecadal variability in dissolved oxygen in the intermediate water layer of the Western Subarctic Gyre and Kuril Basin (Okhotsk Sea). Geophysical Research Letters 28 (12), 2453–2456. Bard, E., Hamelin, B., Fairbanks, R.G., Zindler, A., 1990. Calibration of the 14C timescale over the past 130,000 years using mass spectrometric U–Th ages from Barbados corals. Nature 345, 405–410. Bard, E., Arnold, M., Duplessy, J.-C., 1991. Reconciling the sea level record of the last deglaciation with the d18O spectra from deep-sea cores. Quaternary Proceedings 1, 67–73. Bard, E., Arnold, M., Hamelin, B., Tisnerat-Laborde, N., Cabioch, G., 1998. Radiocarbon calibration by means of mass spectrometric Th-230/U-234 and C-14 ages of corals: an updated database including samples from Barbados, Mururoa and Tahiti. Radiocarbon 40 (3), 1085–1092. Behl, R.J., 1995. Sedimentary facies and sedimentology of the Late Quaternary Santa Barbara Basin (Site 893). In: Kennett, J.P., Baldauf, J.G., Lyle, M. (Eds.), Proceedings of the Ocean Drilling Program, Scientific Results. Ocean Drilling Program, College Station, pp. 295–308. Behl, R.J., Kennett, J.P., 1996. Evidence for brief interstadial events in the Santa Barbara Basin, NE Pacific during the past 60 Kyr. Nature 379, 243–246. Bingham, F., Lukas, R., 1994. The southward intrusion of North Pacific Intermediate Water in the Mindanao Coast. Journal of Physical Oceanography 24, 142–154. Blunier, T., Brook, E., 2001. Timing of millennial-scale climate change in Antarctica and Greenland during the last glacial period. Science 291, 109–112. Bond, G., et al., 1993. Correlation’s between climate records from the North Atlantic sediments and Greenland ice. Nature 365, 143–146. Broecker, W.S., et al., 1988. Radiocarbon-based chronology for the 18 O/16O record for the last deglaciation. Paleoceanography 3, 509–515. Bush, A.B.G., Philander, S.G.H., 1999. The climate of the Last Glacial Maximum: results from a coupled atmosphere-ocean general circulation model. Journal of Geophysical Research 104 (D20), 24509–24525. Cannariato, K.G., Kennett, J.P., 1999. Climatically related millennial-scale fluctuation in the strength of the California margin oxygen-minimum zone during the past 60 k.y. Geology 27, 975–978. Chappell, J., 2002. Sea level changes forced ice breakouts in the Last Glacial cycle: new results from coral terraces. Quaternary Science Reviews 21, 1229–1240. Chappell, J., Shackleton, N.J., 1986. Oxygen isotopes and sea level. Nature 324, 137–140. Chelton, D.B., 1984. Seasonal variability of along-shelf geostrophic velocity off central California. Journal of Physical Oceanography 12, 757–784. Chelton, D.B., Bernal, P.A., McGowan, J.A., 1982. Large-scale interannual physical and biological interaction in the California Current. Journal of Marine Research 40, 1095–1125. Clement, A.C., Seager, R., Cane, M.A., 1999. Orbital controls on the El Nino/Southern Oscillation and the tropical climate. Paleoceanography 14 (4), 441–456.

Acknowledgements We appreciate the generosity of M. Sarnthein, P. Cooke and S. Kienast for valuable discussions on North Pacific Intermediate Waters. We thank personnel of the Ocean Drilling Program, K. Thompson and H. Berg for invaluable technical assistance, and K. Cannariato for valuable advice. We are grateful for the comments and assistance of the L. Stott and J. Lynch-Stieglitz for helping clarify this manuscript. We acknowledge support for J. Kennett from National Science Foundation grant EAR99–04024 (Earth System History) and by Biological and Environmental Research Program (BER), US of Energy, through the Western Regional Center of the National Institute for Global Environmental Change (NIGEC) under Co-operative Agreement no. DE-FC03–90ER61010. I. Hendy was also supported by a National Science Foundation Coastal Research Fellowship.

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