Productivity changes in the Bering Sea during the late Quaternary

ARTICLE IN PRESS

Deep-Sea Research II 52 (2005) 2150–2162 www.elsevier.com/locate/dsr2

Productivity changes in the Bering Sea during the late Quaternary Yusuke Okazakia,, Kozo Takahashib, Hirofumi Asahic, Kota Katsukib, Joichi Horib, Hisato Yasudac, Yuko Sagawad, Hidekazu Tokuyamaa a

Ocean Research Institute, the University of Tokyo, Minamidai 1-15-1, Nakano-ku, Tokyo 164-8639, Japan Department of Earth and Planetary Sciences, Graduate School of Science, Kyushu University, Hakozaki 6-10-1, Fukuoka 812-8581, Japan c Center for Advanced Marine Core Research, Kochi University, Monobe B200, Nankoku 783-8502, Japan d Marine Works Japan Ltd, LivePier Kanazawahakkei 1-1-7 Mutsuura, Kanazawa-ku, Yokohama 236-0031, Japan

b

Received 2 November 2004; accepted 27 July 2005 Available online 21 October 2005

Abstract Changes in biological productivity of the Bering Sea have been evaluated for the late Quaternary based on biogenic opal, calcium carbonate, and microfossils in two piston cores. Biological productivity increased during Marine Isotope Stage (MIS) 1 after the last glaciation. During the last deglaciation, two pronounced peaks of calcium carbonate content were observed at ca. 2400 m water depth, which can be explained by the ‘‘CaCO3 compensation hypothesis’’ and coccolithophore blooms. Simultaneous with the CaCO3 peaks, oxygen-poor deep water was apparently distributed in the Aleutian Basin based on benthic foraminiferal assemblages. The low-O2 events seem to be related to deep-water circulation and/or elevated productivity in the northwest Pacific. After the CaCO3 peak events, biogenic opal contents and diatoms increased gradually, associated with enhanced vertical mixing with an inflow of the Alaskan Stream through the eastern Aleutian Arc passes. r 2005 Elsevier Ltd. All rights reserved. Keywords: Bering Sea; Biological production; Calcium carbonate; Low oxygen; Quaternary

1. Introduction The Bering Sea is the third largest marginal sea in the world linking the Pacific Ocean with the Arctic Ocean (Hood, 1983; Stabeno et al., 1999; Fig. 1). The Bering Sea can be divided into two areas of Corresponding author. Current address. Institute of Observational Research for Global Change, Japan Agency for MarineEarth Science and Technology, 2-15 Natsushima-cho, Yokosuka 237-0061, Japan. Tel.: +81 46 867 9515; fax: +81 46 867 9455. E-mail address: [email protected] (Y. Okazaki).

0967-0645/$ - see front matter r 2005 Elsevier Ltd. All rights reserved. doi:10.1016/j.dsr2.2005.07.003

different depth character: the broad continental shelves (o200 m water depth) mostly in the east and the deep basins in the west (e.g., Aleutian Basin: 43500 m). Seasonally sea-ice covered areas are located mainly over the northern continental shelf (Niebauer et al., 1999). In the Bering Sea, two ridges exist: the Bowers Ridge extending north from the Aleutian Island Arc and the Shirshov Ridge extending south from Kamchatka into the Aleutian Basin. The subarctic circulation system in the North Pacific, which influences the circulation in the

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65˚ N

SIBERIA Bering Shelf

ALASKA

500 m

60˚ N 3500 m Shirshov Ridge

Aleutian Basin BS BOW-8A

C

UMK-3A

tra rS

BOW-9A

im

Amchik

a Pass

s rP as ldi Bu

s

s

ES

s

Ala

as aP

50˚ N

utk Am

883

am

tre

S kan

s Pa

N

ak

ea

Un

it

55˚ N

Bowers Ridge

Emperor Seamounts

45˚ N 170˚ E

180˚

170˚ W

160˚ W

Fig. 1. Topographic map showing the locations of Cores UMK-3A, BOW-9A, and BOW-8A in the Bering Sea and Core ES in the subarctic Pacific obtained during KH99-3 Cruise. ODP Site 883 is also shown. Contour intervals in solid lines are 500 m. Arrows show the directions of major surface currents (Map drawn by ‘‘Online Map Creation’’).

Bering Sea, has a large scale counterclockwise surface circulation, with water masses characterized by relatively low salinity (o34.0 psu) and relatively low temperature (ca. 4–12 1C) and a sharp halocline (ca. 150–200 m; Favorite et al., 1976). There are four gyres in the system: the Alaskan Gyre, the Bering Sea Gyre, the Okhotsk Sea Gyre (OSG), and the Western Subarctic Gyre (WSAG) (Favorite et al., 1976; Nagata et al., 1992). Circulation in the Bering Sea is strongly affected by the Alaskan Stream, which flows westward along the Aleutian Islands and enters the Bering Sea through many passes in the Aleutian Arc (Fig. 1; Stabeno et al., 1999). The

Bering Sea is located in the terminal region of global deep-water circulation. The North Pacific deepwater enters the Bering Sea through the Kamchatka Strait (sill depth: ca. 4000 m), and the bottom water slowly displaces the deep-water upward (Coachman et al., 1999). The Bering Sea is characterized by high silica concentrations in the water column (Tsunogai et al., 1979) and high biological production mainly due to the high contributions of siliceous plankton such as diatoms (Taniguchi, 1999; Takahashi et al., 2002). Along the continental shelf in the Bering Sea, there is an area named the ‘Green Belt’, which represents

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the local highest primary production (Springer et al., 1996). The ‘Green Belt’ is attributed to the Bering Slope Current (BSC), which flows northwestward along the eastern continental slope (Kinder et al., 1975). A part of the surface waters in the Bering Sea flows into the Chukchi Sea through the Bering Strait (present sill depth ca. 50 m). This is the only Pacific origin water that flows into the Arctic Ocean, which consequently presently exits into the Atlantic Ocean. Honjo (1990) suggested that this water mass process influences the present oceanographic conditions, the silica ocean in the Pacific and the carbonate ocean in the Atlantic. The water mass circulation during the glacials was significantly different from the present pattern because of the shutdown of the Bering Strait and several passes in the Aleutian Arc due to lowered sea-level with intensification of continental ice-sheets (Takahashi, 1998). The vast continental shelves in the Bering Sea also were aerially exposed (Beringia) and many works have dealt with the Quaternary history of Beringia (e.g., BrighamGrette, 2001 and references therein). Previous paleoceanographic studies in the Bering Sea mainly employed siliceous microplankton such as diatoms (Sancetta, 1983; Sancetta et al., 1985) and radiolarians (Morley and Robinson, 1986) due to the low calcium carbonate concentrations in the underlying sediments. Gorbarenko (1996) indicated that there were two events with lighter shift of planktonic foraminiferal d18O values during the deglaciation in the subarctic Pacific and the Bering Sea. With these events, CaCO3 contents showed abrupt increases in each sediment core, suggesting changes in biological productivity and pore water chemistry of surface sediments. Kienast et al. (2004) reviewed Bering Sea paleoceanography and showed that opal export was higher during the Holocene compared to the last glacial period. However, our Table 1 Summary information for samples used in this work: location, water depth, and core length, of four piston cores in the Bering Sea and northwestern North Pacific Core ID

Latitude

UMK-3A BOW-8A BOW-9A ES

541 25.220 541 46.990 541 02.230 491 44.700

N N N N

Longitude

Water depth (m)

Core length (cm)

1701 13.380 W 1761 54.990 E 1781 40.580 E 1681 18.930 E

1892 884 2391 2388

1335 894 850 872

current knowledge and available records about Bering Sea paleoceanography have remained limited. In this paper we present changes in biogenic opal, calcium carbonate, and organic carbon data from two piston cores obtained in the southern Bering Sea as well as microplankton (diatom and planktonic foraminifer) and benthic foraminifer records. We will discuss changes in biological production in the southern Bering Sea during the late Quaternary, which are related to the rise and fall of Beringia. 2. Materials and methods Four piston cores were obtained from the southern Bering Sea: Cores UMK-3A (Umnak Plateau), BOW-8A, and BOW-9A (Bowers Ridge) and from the western subarctic Pacific: Core ES (Emperor Seamount) during Cruise KH99-3 of R/V Hakuhomaru, University of Tokyo (Fig. 1; Table 1). The recovered sediment cores consist mainly of dark green-grey or olive-grey diatomaceous silt with several volcanic ash layers. Each core was continuously sliced every 2.0 cm in thickness. 2.1. Age model Age models for these cores were constructed by Okada et al. (2005) based on d18O stratigraphy, paleomagnetic intensity, and the last occurrence datum of radiolarian Lychnocanomma nipponica sakaii. First, age models for Cores ES and BOW8A were constructed on the basis of planktonic foraminiferal d18O (Neogloboquadrina pachyderma sinistral coiling) and comparisons of magnetic susceptibilities between Core ES and the neighboring ODP Site 883. The age models of Site 883 were well established earlier based on benthic foraminiferal d18O (Keigwin, 1995) and 14C data (Kiefer et al., 2001). Descriptions on the age models for Cores ES and BOW-8A have been noted also in Katsuki et al. (2003) and Ratnayake et al. (2005). The paleointensity of Cores UMK-3A and BOW-9A were compared with NAPIS75 (Laj et al., 2000) and Sint800 (Guyodo and Valet, 1999). The relative paleointensity provides a high-resolution global correlation tool for studying climate change, which is not available from oxygen isotope stratigraphy or any other geochronometer (Tric et al., 1992; Channel et al., 1997). Changes in biogenic opal contents in each core were used to define the MIS 1/2 boundary because their variation patterns in the

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North Pacific were synchronous with d 18O records (Narita et al., 2002). In our study, biogenic opal variations in Cores BOW-8A and ES are clearly synchronous with d 18O records (Okada et al., 2005). The last occurrence datum of radiolarian Lychnocanomma nipponica sakaii is ca. 0.05 Ma in the subarctic Pacific and its marginal seas (Morley and Nigrini, 1995). The datum in Core BOW-8A is at 154.5 cm mbsf (meters below sea floor), which is interpolated as 49.5 ka. The interpolation is also adopted for Core BOW-9A and the datum is 49.7 ka, although there are associated errors of
0.3 kyr due to the sampled intervals (Okada et al., 2005). The associated errors of our age models based on graphic fitting of d 18O standard and paleointensity curves cannot be estimated by a statistical method. However, the graphic fitting methods probably have a considerable error of several kyr. Further, ages of each core-top are assumed to be zero-year-old in our study. This assumption has a considerable error because surface sediments may have lost during the coring process. Therefore, our age models of the piston cores have associated errors of several kyr in the top most portion, but the errors estimated for the MIS 1/2 boundary age are small considering the characteristic opal records.

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was analyzed using the CHN Elementary Analyzer. The difference between total carbon and organic carbon was defined as inorganic carbon (IC). Calcium carbonate (CaCO3) was calculated as follows: CaCO3 ¼ IC  8:333 The relative errors of biogenic opal, CaCO3, and OC were better than 6%, 8%, and 8%, respectively, including consideration for heterogenuity of the samples. The contribution of lithogenic matter was calculated using the following equation because marine sediments are mainly composed of the following four components: biogenic opal, calcium carbonate, organic matter, and lithogenic matter (Kawahata et al., 1998). Lithogenic matter ¼ Total-Opal-CarbonateOrganic matter ( ¼ 1.8  OC) Dry bulk density (DBD) was determined from weights of each of 10 cc sediment samples taken by a polypropylene cube, which is customarily used for geomagnetic measurements. The samples were weighed before and after drying in an oven at 50 1C for 24 h. Water contents (%) of each sample were also determined from weights of wet and dry samples. Selected samples in Core UMK-3A were used for planktonic foraminifer analysis (see Section 2.3.).

2.2. Chemical analyses 2.3. Microfossil analyses Chemical analyses were carried out on selected samples at 8 cm intervals in Cores UMK-3A and BOW-9A. Biogenic opal was analyzed by the modified technique of Mortlock and Froelich (1989). Sediment samples were crushed into fine powder after being freeze-dried at 45 1C for 24 h. After being dried, solutions of 10% H2O2 and 1NHCl were added to 30 mg of sample in a polypropylene centrifuge tube to remove organic material and calcium carbonate. Biogenic opal contents were determined through extraction using 20.0 ml of 2 M Na2CO3 solution at 85 1C for 5 h, followed by molybdate-yellow spectrophotometry with a Shimazu UV Mini-1240 Spectrophotometer. Employing the method of Takahashi et al. (2002), organic carbon (OC) was analyzed using a Perkin Elmer 2400 CHN Elementary Analyzer. Powdered dry samples (3 mg) were wrapped with tin foil and total carbon (TC) concentrations were measured. For the analysis of organic carbon contents, each sample was acid leached with concentrated HCl to remove the inorganic fraction. The leached residue

Diatom analysis was carried out on 74 selected samples (approximately every 16 cm) for Core UMK-3A. Wet samples of 10 mg were weighed and treated with hydrochloric acid, hydrogen peroxide, and Calgons (hexametaphosphate), and microslide preparation was carried out following the method described by Katsuki et al. (2003). Observations using a compound light-microscope were made at 400-600X magnification, at least 500 valves were counted per sample, then the number of diatom valves per gram of sediments was estimated. The weights of dry samples were calculated by wet sample weight and water content. Planktonic foraminiferal analysis was conducted on 36 selected samples in Core UMK-3A. The samples were the same ones employed for the DBD analysis. Samples were treated with 10% hydrogen peroxide and washed through a sieve with 63-mm mesh and dried in an oven at 50 1C for 24 h. Each dried sample was split into 1/2 aliquots for picking and sieved through a stainless steel screen with

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125 mm mesh. All planktonic foraminifers were picked under a stereo microscope and identified and mounted in 60 square faunal slides, and the number of planktonic-foraminiferal shells per gram of sediments (total PF; No. shells g1) was estimated based on count, sample volume, and DBD. Benthic foraminiferal analysis was carried out on 35 selected samples for Core BOW-9A. Sediment samples with a quantitative volume (15–29 cc) were taken by a polypropylene syringe. The samples were treated with
1% hydrogen peroxide for 24 h and washed through a sieve with 63-mm mesh and dried in an oven at 601C for 24 h. Each dried sample was split into appropriate aliquots for microscopic observation. At least 200 specimens of benthic foraminiferal shells were picked and identified under a stereo microscope and mounted in 60 square faunal slides. The number of benthic foraminiferal shells per gram of sediments (total BF; No. shells g1) was estimated based on count number, sample volume, and DBD. A total of 14 smear slides were prepared for microscopic observation for coccoliths in Cores UMK-3A and BOW-9A. Wet sediment samples of

approximately 2–3 mg were used for the slide preparation. Seven strata in each core were selected corresponding to CaCO3 concentrations. The abundance of coccoliths in each slide was determined by microscopic observations under a light microscope with polarized lenses at X400 magnification. We observed 42 microscopic fields and computed coccoliths abundance per slide. Our estimation has
50% relative errors, but it is possible to show the order of magnitude of coccolith abundance. Abbreviations for coccoliths abundances are R ¼ rare (o1000 coccoliths/slide), F ¼ few (1000–10,000 coccoliths/slide), C ¼ common (10,000–50,000 coccoliths/slide), and A ¼ abundant (450,000 coccoliths/slide). 3. Results Down-core profiles of biogenic opal (wt%), CaCO3 (wt%), OC (wt%), lithogenic matter (wt%), and linear sedimentation rate (LSR; cm/ kyr) in Cores BOW-9A and UMK-3A are shown in Fig. 2. Each of these components showed similar patterns in Cores UMK-3A and BOW-9A. The

Biogenic opal Calcium carbonate Organic carbon Lithogenic matter LSR (wt%) (wt%) (wt%) (wt%) (cm/kyr) 0 20 40 60 0 4 8 12 16 20 0 2 4 40 60 80 100 0 10 20 30 40

(A)

MIS

0 1 10 2

Age (ka)

20 30 40

3

50 60 0

(B)

20 40

60 0 4 8 12 16 20 0

2

4

40 60 80 100 0 10 20 30 40

MIS

0 1 10

Age (ka)

20

2

30 40

3

50 60

Fig. 2. Down -ore profiles of biogenic opal (wt%), CaCO3 (wt%), OC (wt%), lithogenic matter (wt%), and linear sedimentation rate (LSR; cm/kyr) in Cores (A) UMK-3A, and (B) BOW-9A. MIS 1-3 are labeled; MIS 2, which is the glacial period, is shown as medium gray areas. Light-gray area represents the subglacial period (MIS 3).

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Total diatoms (X106 No. valves g-1)

(A)

0

4 8 12 16 20

%Neodenticula seminae

Total planktonic %Neogloboquadrina foraminifers pachyderma (X103 No. shells g-1) sinistral 4 8 12 0 20 40 60 80 100 0 20 40 60 80 100 0

2155

%Neogloboquadrina pachyderma dextral 0 1 2 3 4 5

MIS

0 1 10 2

Age (ka)

20 30 40

3

50 60 (B)

Total benthic foraminifers (X103 No. shells g-1) 0 2 4 6

%Alabaminella weddellensis 0

20 40 60 80

%Rutherfordoides mexicana 0 10 20 30 40 50

0

%Bulmila pacifica 0

4

8

12

MIS 1

10

Age (ka)

20

2

30 40

3

50 60 Fig. 3. (a) Down-core profiles of microfossils: (A) total diatoms (No. valves g1), %Neodenticula seminae, total PF (No. shells g1), %Neogloboquadrina pachyderma sinistral, and %Neogloboquadrina pachyderma dextral in UMK-3A; (B) total BF (No. shells g1), %Alabaminella weddellensis, %Rutherfordoides mexicana, and %Bolivina pacifica in UMK-3A.

biogenic opal content in the three cores showed lower values during the glacial periods and higher values (up to ca. 50 wt%) during Marine Isotope Stage (MIS) 1. Calcium carbonate content exhibited low values (mostly less than 1.0 wt%) during the glacial periods whereas two pronounced peaks were found during the deglaciation both in Cores UMK3A and BOW-9A. OC content ranged from 0.9 to 5.0 wt%, and increased during the deglaciation. The lithogenic material content in each core showed significantly high values (ca. 80%) during MIS 2 to

4, and decreased during MIS 1 due to an increase of other components such as biogenic opal and CaCO3. Fig. 3 represents changes in microfossils in Cores UMK-3A and BOW-9A. In Fig. 3A, total diatoms (No. valves g1), %Neodenticula seminae, total PF (No. shells g1), %Neogloboquadrina pachyderma sinistral, and %Neogloboquadrina pachyderma dextral are shown. In Fig. 3B, total BF (No. shells g1), %Alabaminella weddellensis, %Rutherfordoides mexicana, and %Bolivina pacifica are shown. Both

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Table 2 Census count for coccolithophores based on smear-slides observation at strata with various CaCO3 values Slide ID Core ID

Age (ka) CaCO3 (wt%) Coccolithophores

A B C D E F G

UMK-3A

0 7.4 9.1 9.6 10.9 11.8 12.4

0 0.8 3.8 0.8 10.4 10.2 0.9

R R F R F A R

H I J K L M N

BOW-9A

0.6 5.9 8.3 8.8 10.8 11.8 12.2

0.2 4.0 16.9 0.7 13.6 16.0 0

R F A C A C C

number of total diatoms and %N. seminae in Core UMK-3A have increased during MIS 1. Changes in %N. seminae showed a sharp peak (470%) at the MIS 1/2 boundary. Total PF in Core UMK-3A showed pronounced bimodal peaks during the deglaciation (ca. 12–8 ka) attributing to calcium carbonate contents. Neogloboquadrina pachyderma sinistral is a dominant taxon throughout, the core but its relative abundances decreased around the MIS1/2 boundary. Percent contribution of Neogloboquadrina pachyderma dextral were quite low (o4%). Likewise PF in Core UMK-3A, total BF in Core BOW-9A showed pronounced bimodal peaks during the deglaciation (ca. 12–8 ka), corresponding to the peaks of calcium carbonate content. In Core BOW-9A, Alabaminella weddellensis is a dominant taxon without regarding the total BF peaks. Rutherfordoides mexicana is a main contributor for the peaks (up to 40%) whereas %R. mexicana shows a few percent in other strata. In general, the abundance of coccoliths showed corresponding changes with CaCO3 contents in both Cores UMK-3A and BOW-9A (Table 2).

4. Discussion 4.1. General trend of paleoproductivity change In general, geochemical and microfossil records suggest an enhanced productivity during MIS 1. However, the components can be distinguished into

two groups: (1) silica components (biogenic opal; diatoms) with gradual increase after the last deglaciation; (2) carbonate components (CaCO3; planktonic foraminifers; benthic foraminifers) with pronounced two peaks during the last deglaciation. Today, the Bering Sea is known as an end member of the ‘‘Silica Ocean’’ by high diatom production (Honjo, 1990). Based on a sediment trap experiment, biogenic opal contributes a major portion of total fluxes today, due to the high production of diatoms such as Neodenticula seminae in the Bering Sea (Takahashi et al., 2002). Biogenic opal contents and OC contents showed a similar trend during the last glacial cycle. However, the variations of biogenic opal and OC contents showed different patterns during MIS 1; i.e. biogenic opal increased during the last 10 kyr, and OC showed their maxima during the deglaciation period. This result provides a possibility for enhanced productivity of calcareous microplankton/benthos abundance such as foraminifers and coccolithophores during the deglaciation. However, dissolution of calcium carbonate relating to CaCO3 peak events must be considered (see Section 4.2). 4.2. CaCO3 peak events during the last deglaciation Notable peaks of CaCO3 contents were observed during the deglaciation in Cores UMK-3A and BOW-9A (Figs. 3 and 4). This CaCO3 pattern is different from that observed in the eastern subarctic Pacific (Zahn et al., 1991), where CaCO3 contents showed higher values during the glacials than those during the interglacials. In the western North Pacific and its marginal seas (the Bering and the Okhotsk Seas), Gorbarenko (1996) reported two major CaCO3 peaks during the deglaciation and the timing of these events (North Pacific terminal events: TA1NP and TA2NP) as 12.6–11.7 ka and 9.8–8.2 ka ( 14C ages), respectively, which were almost synchronous with meltwater pulses in the North Atlantic. Such CaCO3 and OC peaks during the last deglaciation also were found in the western subarctic Pacific (Keigwin et al., 1992; Crusius et al., 2004) and the Okhotsk Sea (Ternois et al., 2001). In our study, such CaCO3 peak events also were observed during approximately 12–10 ka and 9–8 ka (calendar ages) in Core UMK-3A, and 12–10 ka and 8–7 ka (calendar ages) in Core BOW-9A, which seems to be coincident with the results from Gorbarenko (1996). Unfortunately, further comparison and evaluation of the timings

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(A)

0

Organic carbon Calcium carbonate Total planktonic %Neogloboquadrina pachyderma (wt%) (wt%) foraminifers sinistral (X103 No. shells g-1) 4 8 12 0 20 40 60 80100 0 2 4 0 10 20 0

2157

Total diatoms %Neodenticula (X106 No. valves g-1) seminae 0 4 8 12 16 20

Biogenic opal (wt%)

0 20 40 60 80100 0

20

40

60

A 5 Age (ka)

B CD

10 G

E F

15 20

%Bolivina %Alabaminella %Rutherfordoides Organic carbon Calcium carbonate Total benthic pacifica foraminifers weddellensis mexicana (wt%) (wt%) 3 -1 (X10 No. shells g ) 4 8 12 2 4 6 0 20 40 60 80 0 10 20 30 40 50 0 0 2 4 0 10 20 0

(B)

0

Age (ka)

0

20

40

60

H

5 10

Biogenic opal (wt%)

I J

K L N

M

15 20

Fig. 4. Down-core profiles of geochemical components and microfossils during the last 24 kyrs: (A) OC (wt%), CaCO3 (wt%), total PF (No. shells g1), %Neogloboquadrina pachyderma sinistral, biogenic opal (wt%), and %Neodenticula seminae in Core UMK-3A; (B) OC (wt%), CaCO3 (wt%), total BF (No. shells g1), %Alabaminella weddellensis, %Rutherfordoides mexicana, and %Bolivina pacifica, and biogenic opal (wt%) in Core BOW-9A. The alphabets in the illustration for CaCO3 indicate strata where smear slide observations were performed.

are not possible at this stage due to our age model limitation (see detail in Section 2.1). CaCO3 variation patterns in Cores UMK-3A and BOW-9A can be explained by the ‘CaCO3 compensation’ hypothesis, which predicts CaCO3 dissolution and preservation events responding to CO2 3 content of the deep-sea waters, proposed by Broecker and Peng (1987). The box model employed in the hypothesis predicted a CO2 3 excursion, which lasted for several thousand years during the last deglaciation (Broecker and Peng, 1987). Broecker and co-workers have reconstructed a series of deepsea CO2 3 changes during the last glacial cycle based on planktonic foraminiferal weights (Broecker et al., 1999), and suggested CaCO3 preservation maximum during Termination I (Broecker et al., 2001). Recently, Marchitto et al. (2005) have reconstructed deep-sea CO2 changes in the equatorial Pacific 3 during the last glacial cycle based on Zn/Ca records from benthic foraminiferal shells, a new proxy for

deep-sea CO2 3 contents. The reconstructed deep-sea CO2 changes were compared with the predicted 3 deep-sea CO2 3 changes by box-models of ‘rain ratio hypothesis’ (Archer and Maier-Reimer, 1994) and ‘CaCO3-compensation hypothesis’. They indicated that the benthic foraminiferal Zn/Ca data agreed with the CaCO3-compensation hypothesis. Today, CaCO3 saturation horizon in the Bering Sea is significantly shallower (o500 m) than that in South Pacific (close to 3000 m) (Feely et al., 2002). In our study, CaCO3 components, including numerous coccoliths, were well preserved during the last deglaciation even at water depth below 2000 m (Core BOW-9A), significantly deeper than the CaCO3-saturation horizon. Our data strongly support for the CaCO3-compensation hypothesis in the terminal region of deep-ocean circulation. Another factor relating to CaCO3 dissolution/ preservation is a denitrification on the continental shelves, which generates alkalinity regardless of

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CaCO3 dissolution (Chen, 2002). During the last deglaciation, rapid sea-level rise occurred until 6 ka with the late Quaternary ice sheet melting (Cronin, 1999) and the vast Beringia was submerged. The process must have enhanced the nitrification on the continental shelves and consequent alkalinity generation. However, Chen (1993) indicated the continental shelf water penetrated no more than 1000 m in the present Aleutian Basin, thus, the direct effect of the generated alkalinity on CaCO3 preservation is questionable during the deglaciation in Cores BOW-9A (2391 m) and UMK-3A (1892 m) but the indirect effect through biological production is considerable. Localized rain-ratio changes have a potential to impact regional CaCO3 dissolution history, even if the rain-ratio hypothesis cannot work on a global scale (Marchitto et al., 2005). Based on OC, CaCO3, and microfossil abundance changes, biological production in the Bering Sea seems to have been elevated during the last deglaciation mainly due to the contribution by CaCO3 shell-bearing plankton. Takahashi et al. (2002) presented the results on long-term particle flux changes in the Bering Sea during 1990–2000 (Station AB: sediment trap station occupied in the Aleutian Basin (53.51N, 1771W)). They showed that calcium carbonate contributes 13 wt% to the total particle fluxes at Station AB during 1990–2000. During the last deglaciation, CaCO3 contents were up to 17 wt% in Core BOW-9A. Considering significant CaCO3 dissolution on the sea-floor, we interpret that the high CaCO3 preservation cannot be simply explained by changes in dissolution effect. Nakatsuka et al. (1995) indicated that vertical mixing was suppressed and thus nutrient concentrations were low in surface waters during the last deglaciation in the Aleutian Basin due to melting water from sea-ice (Sancetta, 1983) and probably continental ice sheets (Grosswald and Hughes, 2002). During the last deglaciation, reduction of %Neogloboquadrina pachyderma sinistral, a cold-water dwelling planktonic foraminifer (Hemleben et al., 1989), suggests the presence of relatively high-temperature surface water. In the present Bering Sea, a massive coccolithophore bloom was observed over the eastern continental shelf in 1997 (Vance et al., 1998; Napp and Hunt, 2001; Iida et al., 2002). Napp and Hunt (2001) suggested that the reason for the bloom was nutrient depletion in the surface water with thermocline intensification due to calm winds, clear skies, and warm air temperature. This is

because coccolithophores are able to cause a bloom in lower nutrient conditions than diatoms (Margalef, 1978; Egge and Heimdal, 1994). Likewise in the Bering Sea, calcareous plankton such as coccolithophores was dominant rather than diatoms during the deglaciation in the Okhotsk Sea based on biomarker changes (Seki et al., 2004). Thus, the massive blooms of coccolithophores may have been a common phenomenon in the marginal seas in the subarctic Pacific, relating to sea-ice retreat during the deglaciation. 4.3. Oxygen-poor deep water during the last deglaciation Based on benthic foraminiferal assemblages, an oxygen-poor condition is suggested during the last deglaciation. During the two CaCO3 peak events, total BF also showed notable peaks in Core BOW9A. Simultaneous with the peaks, Rutherfordoides mexicana (Loeblich and Tappan, 1987) became the most abundant species, replacing Alabaminella weddellensis. Alabaminella weddellensis is known as a phytodetritus species, which is abundant opportunistically when phytodetritus deposition increases, reflecting seasonal phytoplankton bloom in surface water (Gooday, 1993; Thomas et al., 1995). Thomas et al. (1995) suggested that the phytodetritus species cannot dwell in high-productivity regions throughout the year because of oxygen depletion in the bottom water due to decomposition of organic matter. Genus Rutherfordoides is a closely related taxon with genus Bulimina and genus Fursenkoinak, whose species are generally living in oxygen-poor sediments (Sen Gupta and Machain-Castillo, 1993). During the CaCO3 peaks, relative abundance of Bolivina pacifica, living in the eastern Pacific oxygen minimum zones with the oxygen concentration of 0.3–0.5 ml/l (Douglas and Hamilton, 1979; Sen Gupta and Machain-Castillo, 1993), also showed its maxima. These suggest that oxygen-poor water was present in the Aleutian Basin ca. 2400 m depth. Today, oxygen concentrations near Site BOW-9A are 1.8 ml/l at 2500 m, and the oxygen minimum zone (OMZ) is located in the depth interval of approximately 500–1500 m with 0.6–0.9 ml/l oxygen contents (Conkright et al., 2002). Therefore, some dramatic changes relating to the low-O2 events are visible. Such kind of dysaerobic events are unusually widespread along the western margin of North America during the Bølling-A˚llerød interval (14.7–12.9 ka) (Zheng et al., 2000; van Geen et al.,

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2003; Crusius et al., 2004). For this reason, there are two plausible explanations: (1) enhanced productivity along the continental margin by coastal upwelling (e.g., Mix et al., 1999; Stott et al., 2000; Ortiz et al., 2004); (2) reduction in intermediate-water ventilation of the North Pacific (e.g., Duplessy et al., 1989; Kennett and Ingram, 1995; Zheng et al., 2000). In our study, biological productivity may have been enhanced during the CaCO3 peak events mainly contributed by calcareous plankton in the Bering Sea. With the elevated productivity, OMZ expansion or migration to deep water may have occurred. However, surface nutrients were depleted due to intensified stratification. Thus, the biological productivity is important but probably not the sole factor for the oxygen-poor condition in the Aleutian Basin. As for intermediate-water ventilation, it is hard to adopt the scenario in the California margins because the water depth of Core BOW-9A (2391 m) is significantly greater than the distribution range of the North Pacific Intermediate Water, which ranges ca. 300–800 m at present (Talley, 1993). We suggest a possible reason to decipher the oxygen-poor events in the Bering Sea in regard to deep-water circulation. The source for Bering Sea bottom water is the North Pacific from
3500–4000 m depths, which enters through the Kamchatka Strait (sill depth: ca. 4000 m) and the bottom water slowly displaces the deep-water upward (Coachman et al., 1999). During the last glacial maximum, the deep-water circulation (below 2000 m depths) was similar to that of today (Keigwin, 1998; Matsumoto et al., 2002). According to Sigman and Boyle (2000), the OMZ migrated into the abyssal Pacific Ocean, which was caused by changes in patterns of nutrient supply and export production in the Southern Ocean. The oxygenminimum abyssal water may affect the Bering Sea bottom water depending on accelerated or sluggish deep-water circulation patterns. Another candidate is the high productivity in the northwest Pacific region during the last deglaciation. Recently, Crusius et al. (2004) suggested that elevated productivity in the northwest Pacific could have reduced intermediate-water O2 concentrations during the Bølling-A˚llerød interval along the western North America margin. We assume that such a high productivity must have reduced deep-water O2 concentrations in the northwest Pacific and thus the deep-water penetration into the Bering Sea during the last deglaciation could have produced oxygen-poor condition below 2000 m water depth.

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Further investigations of the timing and magnitude of the horizontal and vertical convective processes are required to clarify the CaCO3 peaks with oxygen-poor events. 4.4. The Holocene After the last deglaciation, both biogenic-opal contents and total diatom abundances have increased gradually. These are probably due to nutrient supply to the upper water with enhanced vertical mixing relating to weakened melt-water input (Nakatsuka et al., 1995) and Alaskan Stream injection through the Aleutian passes. Neodenticula seminae, the most abundant diatom taxon in flux assemblages of the Bering Sea (Takahashi et al., 2000, 2002), can be a good tracer for the Alaskan Stream water (Sancetta, 1982; Sancetta and Silvestri, 1984). The Alaskan Stream water flows into the Bering Sea through the Unimak Pass at present (Stabeno et al., 2002). The present maximum depth of the Unimak Pass is 160 m and thus the southeastern Bering Sea around Site UMK-3A was semi-closed during the last glacial. Elevation of relative abundances of N. seminae strongly suggests an enhanced inflow of the Alaskan Stream water into the southeastern Bering Sea. The Alaskan Stream water has relatively high temperature and salinity than those of the Bering Sea Gyre water (Favorite et al., 1976). Therefore, the increased injection of the Alaskan Stream water during the Holocene may have promoted the vertical mixing in the Bering Sea and thus this serves as the main reason why primary production changed from coccolithophores to diatoms. Acknowledgments We would like to express our appreciation to all the scientists, technicians, captain and crew of the R/V Hakuho-maru, for their able assistance in collecting the piston cores during Cruise KH99-3. Many thanks are extended to the Bering Sea paleoceanography working group members for providing valuable suggestions. The manuscript was greatly improved by constructive comments and suggestions from two anonymous reviewers and Dr. R. Jordan. Dr. Jordan also provided English editing of the manuscript. This study was supported by the following research programs: MEXT Grantsin-Aid-for Scientific Research B1 Project No. 13440152, B2 Project No. 15310001 and JSPS B Project No. 17310009.

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