Deep Ocean Circulation (Elsevier Oceanography Series)

DEEP OCEAN CIRCULATION PH YSlCAL AND CHEMICAL ASPECTS FURTHER TITLES IN THIS SERIES Volumes 1-7, 11, 15, 16, 18, 19, ...

Author: T. Teramoto

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DEEP OCEAN CIRCULATION PH YSlCAL AND CHEMICAL ASPECTS

FURTHER TITLES IN THIS SERIES Volumes 1-7, 11, 15, 16, 18, 19, 21, 23, 29 and 32 are out of print. 8 E. LlSlTZlN SEA-LEVEL CHANGES 9 R U PARKFR STU'DY OF BENTHIC co MMuNITIES 10 J.C.J. NIHOUL (Editor) MODELLING OF MARINE SYSTEMS 12 E.J. FERGUSON WOOD and R.E. JOHANNES TROPICAL MARINE POLLUTION 13 E. STEEMANN NIELSEN MARINE PHOTOSYNTHESIS 14 N.G.JERLOV MARINE OPTICS 17 R.A. GEYER (Editor) SUBMERSIBLES AND THEIR USE IN OCEANOGRAPHY AND OCEAN ENGINEERING 20 P.H. LEBLOND and L.A. MYSAK WAVES IN THE OCEAN 22 P. DEHLINGER MARINE GRAVITY 24 F.T. BANNER, M.B. COLLINS and K.S. MASSIE (Editors) THE NORTH-WEST EUROPEAN SHELF SEAS. THE SEA BED AND THE SEA IN MOTION 25 J.C.J. NIHOUL (Editor) MARINE FORECASTING 26 H.G. RAMMING and Z. KOWALIK NUMERICAL MODELLING MARINE HYDRODYNAMICS 27 R.A. GEYER (Editor) MARINE ENVIRONMENTAL POLLUTION 28 J.C.J. NIHOUL (Editor) MARINE TURBULENCE 30 A. VOlPlO (Editor) THE BALTIC SEA 31 E.K. DUURSMA and R. DAWSON (Editors) MARINE ORGANIC CHEMISTRY 33 RHEKINIAN PETROLOGY OF THE OCEAN FLOOR 34 J.C.J. NIHOUL (Editor) HYDRODYNAMICS OF SEMI-ENCLOSED SEAS 35 B. JOHNS (Editor) PHYSICAL OCEANOGRAPHY OF COASTAL AND SHELF SEAS 36 J C J NIHOUL (Editor) HYDRODYNAMICS OF THE EQUATORIAL OCEAN 37 W. LANGERAAR SURVEYING AND CHARTING OF THE SEAS 38 J C J NIHOUL (Editor) REMOTE SENSING OF SHELF SEA HYDRODYNAMICS 39 T.ICHIYE (Editor) OCEAN HYDRODYNAMICS OF THE JAPAN AND EAST CHINA SEAS 40 J.C.J. NIHOUL (Editor) COUPLED OCEAN-ATMOSPHERE MODELS 41 H. KUNZENDORF (Editor) MARINE MINERAL EXPLORATION 42 J.C.J NIHOUL (Editor) MARINE INTERFACES ECOHYDRODYNAMICS 43 P. LASSERRE and J.M. MARTIN (Editors) BIOGEOCHEMICAL PROCESSES AT THE LANDSEA BOUNDARY 44 I.P. MARTINI (Editor) CANADIAN INLAND SEAS 45 J C.J. NIHOUL (Editor) THREE-DIMINSIONAL MODELS OF MARINE AND ESTUARIN DYNAMICS 46 J C J NIHOUL (Editor) :3?,A,L&-SCALE TURBULENCE AND MIXING IN THE

THE

47 M.R. LANDRY and B.M. HICKEY (Editors) COASTAL OCENOGRAPHY OF WASHINGTON AND OREGON 48 S.R. MASSEL HYDRODYNAMICS OF COASTAL ZONES 49 V.C. LAKHAN and A.S. TRENHAILE (Editors) APPLICATIONS IN COASTAL MODELING 50 J.C.J. NIHOUL and B.M. JAMART (Editors) MESOSCALE SYNOPTIC COHERENT STRUCTURES IN GEOPHYSICAL TURBULENCE 51 G.P. GLASBY (Editor) ANTARCTIC SECTOR OF THE PACIFIC 52 P.W. GLYNN (Editor) GLOBAL ECOLOGICAL CONSEOUENCES OF THE 1982-83 EL NINO-SOUTHERN OSCILLATION 53 J. DERA (Editor) MARINE PHYSICS 54 K. TAKANO (Editor) OCEANOGRAPHY OF ASIAN MARGINAL SEAS

Elsevier Oceanography Series, 59

DEEP OCEAN C IRC ULATI0N PHYSICAL AND CHEMICAL ASPECTS

Edited by

T. Teramoto Department of Information and Computer Sciences, Kanagawa University, Tsuchiya, Hiratsuka, Kanagawa 259-12, Japan

ELSEVIER

Amsterdam - London - New York -Tokyo

1993

ELSEVIER SCIENCE PUBLISHERS B.V. Sara Burgerhartstraat 25 P.O. Box 21 1, 1000 AE Amsterdam, The Netherlands

ISBN: 0 444 88961 2

0 1993 Elsevier Science Publishers B.V. All rights reserved. No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without the prior written permission of the publishers, Elsevier Science Publishers B.V., Copyright & Permissions Department, P.O. Box 521, 1000 A M Amsterdam, The Netherlands. Special regulations for readers in the U.S.A.- This publication has been registered with the Copyright Clearance Center Inc. (CCC), Salem, Massachusetts. Information can be obtained from the CCC about conditions under which photocopies of parts of this publication may be made in the U.S.A. All other copyright questions, including photocopying outside of the U.S.A., should be referred to the copyright owner, Elsevier Science Publishers B.V., unless otherwise specified. No responsibility is assumed by the publisher for any injury and/or damage t o persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the materials herein. This book is printed on acid-free paper. Printed in The Netherlands.

V

Preface This book comprises the final report of the project research titled the Dynamics of the Deep Ocean Circulation. The project was organized and conducted under the chairmanship of myself for four fiscal years from 1987 to 1990, with financial support of the Ministry of Education, Science and Culture of Japan under the grant in aid of Priority Area Programme. As will be described in Chap. 1 in more detail, the project needed to involve research groups who had sufficient experiences in long-term current measurements, accurate chemical analyses of dissolved heavy metals and isotope tracers, numerical and hydraulic modellings, sediment trappings, and so on. These experiences had been gained through two preceding research projects, the first of which bore the title, “Fundamental Studies on Preservation of Ocean Environment”, and the second, the “Ocean Characteristics and their Changes.” The former was carried out for four years from 1975 to 1978 under the chairmanship of Prof. Y. Horibe and the latter for four years from 1981 to 1984 under the chairmanship of Prof. K. Kajiura. Both were financially supported by the Ministry of Education, Science and Culture, too. I am greatly indebted to these supports and express my sincere gratitude to these chairmen. Throughout the period of research, I had arranged symposia several times, in which ocean scientists who were not directly committed to the project also joined. Comments and criticisms given there were very helpful for the promotion of the project. Here, I would like to express my sincere gratitude to Dr. Susumu Tabata, the former senior scientist of the Institute of Ocean Sciences in Sidney, B. C., CANADA, for reviewing whole the papers involved and giving valuable advices and appropriate comments.

Toshihiko TERAMOTO Kanagawa, University, 1992

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vii

Table of Contents Preface T.Teramoto

.................................

v

Chapter 1: Motivation of the Project Research and a Hypothetical Circulation model presented prior to the Research Motivation of the Project Research and a Hypothetical Circulation Model Presented Prior t o the Research T. Teramoto . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

3

Chapter 2: Structure of Subsurface Water Circulation and Associated Distribution of Water Properties Deep Water Properties and Circulation in the Western North Pacific M. Kawabe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Density Field along 12"N and 13'N in the Philippine Sea I<. Uehara, I(. Taira and A. Masuda . . . . . . . . . .

...

17

..........

39

Abyssal Boundary Current along the Northwestern Perimeter of the Philippine Basin M. Chaen, M. Fukasawa, A. Maeda, M. Sakurai and M. Takematsu . . . 51 Deep Circulation in the Shikoku Basin Measured with the SOFAR floats K. Taira and D. Yanagimoto . . . . . . . . . . . . . . . . . . . . . . . .

69

Chapter 3: Circulation and Mixing of Water as Deduced from Distribution of Chemical Tracers Chemistry and the Oceans:* An Overview Y. Nozaki . . . . . . . . . . . . . . . .

...................

83

Philippine Sea Abyssal Waters in the Northwestern Pacific: Characterization from Tracer-Tracer Diagrams

T. Gamo

...................................

91

Dynamics of the Japan Sea Deep Water Studied with Chemical and R a d i e chemical Tracers S. Tsunogai, Y. W. Watanabe, I(. Harada, S. Watanabe, S. Saito and M. Nakajima . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 Instrumental Development for Measurement of Phosphate in Seawater and Some Discussion of Nutrient Distributions in the North Pacific K. Fujiwara and H. Tsubota . . . . . . . . . . . . . . . . . . . . . . . 121 Actinium-227: A Steady State Tracer for the Deep-sea Basin-wide Circulation and Mixing Studies 139 Y. Nozaki . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

...

Vlll

Distributions and Mass-balance of 2399240Pu and 137Cs in the Northern North Pacific 157 Y. Nagaya and K. Nakamura . . . . . . . . . . . . . . . . . . . . . .

.

Trace Metals in the North Pacific - Recent Development of Clean Techniques and their Applications t o Ocean Chemistry H. Tsubota, S. Nakamura and K. Shitashima . . . . . . . . . . . . . . . 169 Distribution of Dissolved Organic Nitrogen in the North Pacific Ocean Y. Maita and M. Yanada . . . . . . . . . . . . . . . . . . . . . . . . .

.

185

Determination of Some Oxyacid Elements and Manganese in Seawater and their Distributions in Some Unique Environments of the North Pacific E. Nakayama, Y. Sohrin and K. Isshiki . . . . . . . . . . . . . . . . . . . 199

Chapter 4: Vertical Flux of Chemical Substances and their Behaviours in the Deep Oceans Seasonal Variation of Lithogenic Flux in Japan Trench Continental Slope Measured by Sediment Trap S. Noriki, R. Saito, C. Saito and S . Tsunogai . . . . . . . . . . . . . . 21 1 Vertical Fluxes of Organic Materials in the Northern North West Pacific and Breid Bay, Antarctica, with Special Reference in the Effect of Phytoplankton Bloom N. Handa and T. Nakatsuka. . . . . . . . . . . . . . . . . . . . . . . . 221 Organic Composition of Sinking Particles (JT-01) in Japan Trench as Revealed by Pyrolysis Gas Chromatography/mass Spectrometry R. Ishiwatari, S. Yamamoto, N. Handa and Y. Nozaki . . . . . . . . . 235

Chapter 5: Modelling of Subsurface Water Circulation and Associated Dynamical Processes Numerical Modelling of the Philippine Sea M.Kubota . . . . . . . . . . . . . . . .

.................

255

The Occurrence of Vacillation in a Model Ocean Driven by Wind, Heat and Salinity Flux K. Takeuchi and Y. Kashino . . . . . . . . . . . . . . . . . . . . . . . . 273 Modelling of Western Pacific Abyssal Circulation - Preliminary Experiment N. Suginohara, S. Aoki and M. Nakata. . . . . . . . . . . . . . . . . . Diagnostic Approaches on Deep Ocean Circulation N. Imasato, H. Nagashima, H. Takeoka and S. Fuji0

285

. . . . . . . . . . . 307

Laboratory Experiments of Dense Water Descending on Continental Slope Y. Nagata, R. Kimura, H. Honji, Y. Yamazaki, K. Kawaguchi and T. Hosoyamada . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. 333

ix Chapter 6: Development of Acoustic Technology for Ocean Measurement Target Parameter Estimation by Linear-Period h4odulated Signal W. Mitsuhashi and H. Mochizuki . . . . . . . . . . . . . . . . . .

....

353

A Low Frequency Underwater Sound Source and Control Technique of Transducer Directivity .365 H. Hachiya and M. Okujima . . . . . . . . . . . . . . . . . . . . . . . Development of Multipaths Inverted Echosounder T. Takeuchi and K. Taira . . . . . . . . . . . .

..............

375

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Chapter 1 Motivation of the Project Research and a Hypothetical Circulation model presented prior to the Research

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Deep Ocean Circulation. Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

3

Motivation of the Project Research and a Hypothetical Circulation Model Presented Prior to the Research Toshihiko TERAMOTO*

1.

Motivation of the Project Research

The subtropical area of the western North Pacific is much different in topographic feature from that of the western North Atlantic. The former is characterized by the Izu-Ogasawara-Mariana Ridges running almost longitudinally across the western part of it and the latter, by the Grand Banks occupying the north of the western part of it (Fig. 1). The difference, however, does not actually exert any serious influence upon the surface subtropical circulation, and the circulations in the both oceans are almost similar in configuration. In another words, the Kuroshio flowing along the continental boundary and the subsequent Kuroshio Extension flowing into the central ocean of the Pacific correspond well in their combination t o the Florida Current and the subsequent Gulf Stream of the Atlantic. In contrast to this, the characteristic bottom topographies described above seem t o affect greatly t o the deep circulations of the oceans. According t o the deep circulation model hypothetically delineated by Stommel and Arons (1960), the deep western boundary current in the North Pacific flows along the eastern side of the above-stated ridges and not along the continental boundary but that in the North Atlantic flows closely along the continental boundary (Fig. 1). Thus, the Philippine Sea, which is bounded by the continental boundary in the west and with the ridges in the east, has long been considered to be isolated in the deep layer from the main part of the North Pacific (Fig. 1). This leads t o an understanding that the deep water of the sea is almost stagnant and is old in age. The age determination with 14C indicated the deep water of the Philippine Sea is younger by about 200 years than that of the main North Pacific (Gamo, personal communication). This fact suggests that the deep western boundary current branches into the sea through the Yap-Palao Trough on the Mariana ridge soon after the current flows into the northern hemisphere upon passing over the equator (Fig. 2). This implies the deep Philippine Sea water is also involved in the deep circulation of the North Pacific. What is the configuration of the deep circulation of the western North Pacific? This question motivated the author t o programme the present project research. *School of Science, Kanagawa University Tsuchiya 2946, Hiratsuka, Kanagawa 259-12, Japan

4

Fig. 1. The bottom topography of the world ocean (lower panel) and the configurational model of the deep circulation by stommel and Arons (upper panel) (Open university/Pergamon). The programme is naturally expected t o give a solution to the question on what the flow configuration under the Kuroshio is, a question t h a t has long been borne in minds of many Japanese oceanographers.

2. 2.1.

The Subsurface Circulation Model Presented to be Verified through the Project Research The mid-depth circulation model

Through long-term measurements of vertical profile of velocity in the Kuroshio area south of Japan, the level of no motion is determined to be a t the depth around

5

Fig. 2 . The detailed bottom topography of the Philippine Sea and western North Pacific (Mogi, 1972). 2,000 m (Fukasawa and Teramoto, 1986). Thus, on taking the reference level at the isobaric surface of 2,000 db, the contour maps of geopotential anomalies on the isobaric surfaces of 1,000 and 3,000 d b for the Philippine Sea and western North Pacific (Fig. 3) are drawn on the basis of accurate hydrographic observations (Kaneko and Teramoto, submitted to J. Oceanogr.). The figure gives the configurational model of the mid-depth circulation, which has a two-layer structure separated at the isobaric surface of 2,000 db. The mid-depth used here is defined t o involve depths below 1,000 m and above 3,000 m. The most characteristic feature of the configuration is the elliptic eddy occupying the whole area of the Shikoku Basin (Fig. 2 ) . The eddy is commonly of the size of about 1,200 km in longer diameter and 600 km in shorter one both in the upper and lower mid-depth layers but is rotating clockwise in the upper layer and anticlockwise in the lower one. The geostrophic velocity profile associated with the eddy is almost equal t o t h a t from the direct measurements stated above. Through the water-mass analysis, the mid-depth water of the Philippine Sea and western North Pacific is understood to be classified typically into three water types; that is, the South Philippine Sea Water, Shikoku Basin Water and Northwest Pacific Water (Kaneko and Teramoto, submitted t o J. Oceanogr.). These water types, each of which is identified here by a combination of S (Salinity), DO

6

Fig. 3. Contour maps of geopotential anomaly on isobaric surfaces of 1,000 d b (upper panel) and 3,000 d b (lower panel) for the Philippine Sea and western North Pacific. The reference level is put at, 2,000 d b surface. (Dissolved Oxygen) and DSi (Dissolved Silica) of water, are defined as averages of

S,DO and DSi of water samples from ten stations in regions of 10'00' to 13"02'N, 125'37' t o 137'00'E in the South Philippine Sea, 32'30' to 33'32'N, 135'49'to 138'05'E in the Shikoku Basin and 29'55' to 35"56'N, 146'34' to 160'00'E in the western North Pacific, respectively. Actually, waters of S, DO and DSi in ranges of twice the standard deviations in distributions of S, DO and DSi over the ten stations are regarded as those of water types under consideration. Thus, the water

7

31.61

-

- 31.61 JI 6234.61-

SBW :31.6a 34.59-

'

'

,

'

I

NwPW

.

a

SBW

A

SPS\\' - 0 3 1

DO (mi

Fig. 4. S-DO diagramme (left upper panel) and DSi-DO diagramme (left lower panel) on the isopycnal surface of 0-0 = 27.70 for water samples in the Philippine Sea and western North Pacific. Positions for these water samplings are shown by the use of six kinds of symbol, which are used to identify the six areas illustrated on the right panel of the figure. Squares with symbols of SPSW, SBW and NWPW on the diagrammes indicate the domains for water types of south Philippine Sea Water, Shikoku Basin Water and Northwest Pacific Water, respectively. types are specified not at points but within square domains on S-DO and DO-DSi diagrammes as illustrated for the isopycnal surface of cr0 = 27.70 in Fig. 4. This isopycnal surface almost corresponds with isobaric surface of 2,000 db. The geographic distribution of these three water types on the isopycnal surface of = 27.70 is drawn in Fig. 5 (Kaneko and Teramoto, submitted t o J. Oceanogr.). Upon comparison of this figure t o Fig. 3 it is clearly indicated that the area of the Shikoku Basin Water in the mid-depth layer around 2,000 d b coincides very well with those occupied by the elliptic eddy described above. As suggested from S-DO and DO-DSi relations in Fig. 4, the Shikoku Basin Water cannot be formed through the linear horizontal mixing of waters of the other two water types. And so, the formation seems to be made through vertical mixing associated with the eddy motion of the intense vertical shear of velocity.

8

. . . .

m m m m m m

Fig. 5. The geographical distributions of water types of SPSW SBW (A)and NWPW ( W ) on the isopycnal surface of ~0 = 27.70.

2.2.

(0):

The deep circulation model

As seen on S-DO-DSi diagramme in Fig. 6, the square domains for the water type of the Shikoku Basin Water overlap with those of the South Philippine Sea Water for the deep layer around the isopycnal surface of 0 0 = 27.77. This surface corresponds to the isobaric surface of 3,000 db. The deep layer mentioned here implies the layer between 3,000 and 4,000 ni in depth. Thus, the deep layer of the whole Philippine Sea including the Shikoku Basin is known t o be filled with the South Philippine Sea Water. Then, the question arises on where the origin of this water is. For the purpose of tracing the characteristics of the water in contrast t o the North Pacific Water, Kaneko and the author draw S-DO (upper panel) and DSi-DO (lower panel) diagrammes in Fig. 7 for water samples from three groups of station in the following cruises; that is, the Ryofu-Maru Cruise along 137"E in the Philippine Sea, the Hakuho-Maru Cruise along 158"E in the western North Pacific and GEOSECS EXPEDITION almost along 180" in the equatorial area over the northern and southern hemispheres (Fig. 7). These diagrammes for the isopycnal surface of ce = 27.77 indicate that the deep Philippine Sea Water can be traced t o branches 08 from the deep equatorial water at around sta. 4 of the GEOSECS EXPEDITION near 10"s and t h a t the deep Northwest Pacific Water, on the contrary, from the deep equatorial water at around sta. 1 of the GEOSECS EXPEDITION near 5"N. These facts suggest that a part of the deep western boundary current, which flows northwards after passing through the Samoan Passage (Fig. 7), branches eastwards into the Philippine Sea and that the rest of the current flows further northwards

9

31.G5

t

1301

2.8

3.0

3.2

3.4

DO ( i d P ) Fig. 6 . S-DO diagramme (upper panel) and DSi-DO diagramme (lower panel) on the isopycnal surface of uo = 27.77 for water samples in the Philippine Sea and western North Pacific. Symbols used are the same as in Fig. 5.

into the western North Pacific along the Izu-Ogasawara-Mariana Ridge. As delineated in Fig. 8 (Kaneko and the author), the isolines for DSi and DO on the isopycnal surface of U Q = 27.77 are almost of the longitudinal, configuration, which is clearly different from the latitudinal one on the isopycnal surface of CTQ = 27.55 in the same figure. The latter surface agrees almost with the isobaric surface of 1,500 db, where the circulation is delineated as in Fig. 9(b). The current associated with the circulation on this surface is zonal and the isolines is parallel t o the current except for the western boundary region. On taking into consideration the fact that in the deep layer the isolines for DSi and DO are longitudinal and also the fact that the deep current directly measured flows westwards in the northern boundary region, the author modelled hypothet,ically the deep circulation of the Philippine Sea as illustrated in Fig. 9(d) on referring t o the world deep circulation

10 3 4 . 701

I

1.0

I. 5

2.5

2.0

3:5

3.0

D I S S O L V E 0 O X Y G E N (ml/l)

22!

%-\ 5'6

1

-3 .4

.5 1.0

1.5

2.0

nissnLvEo

2.5

3.0

-6

3.5

O X Y G E N rml/li

Fig. 7. S-DO diagramme (left upper panel) and DSi-DO diagramme (left lower panel) on the isopycnal surface of = 27.77 for water samples from stations of three cruises; Ryofu-Maru Cruise along 137"E (A), Hakuho-Maru Cruise along 158"E ( 0 )and GEOSECS EXPEDITION almost along 180" ( W ) . Station maps for the three cruises are shown in the right panel. model by Stommel and Arons (1960), in which the boundary current, interior flow and the associated upwelling play the essential role in combination. On summarizing these results the author, prior to commencing our project research, presented the layered model for the subsurface circulations of the Philippine Sea and western North Pacific as illustrated in Fig 9(b), (c) and (d) in contrast to that for the surface circulation (Fig 9(a)). The verification of the model through measurements as well as the understanding of the mechanism to produce the circulation provided the most important subjects to be worked out in our project research. In the modelling, the bottom circulation is omitted because of its serious dependency upon the bottom topography, which is very complicated.

3.

Flow Diagramme of the Project Research The project consists of research groups of the following specific activities:

11

i?b.F' "

~

"

'136'

"

'

' 'llti' ' "

"

' " lio' " " " "160' ' ""I

'

170''

" "

Fig. 8. The geographical configurations of isolines for DSi and DO distributions on the isopycnal surface of o g = 27.77 (left panel) and those on the isopycnal surface of ot)= 27.55 (right panel). (1) long-term Eulerian measurements of subsurface boundary current by the use of moored arrays of current meters. (2) Lagrangian measurements of subsurface interior flow by the use of SOFAR Float system. (3) accurate hydrographic observations with CTD and Rosette Water Sampler to obtain the distribution of geopotential as well as concentrations of dissolved

oxygen and silica. (4) accurate measurements of distributions of chemical tracers such as natural and artificial radionuclides, dissolved heavy metals and so on. (5) subsequent trapping of drifting sediments for long time t o reveal the transporting and diffusing processes of chemical substances in the deep ocean. ( 6 ) simulation and assimilation of subsurface circulation processes by means of numerical and hydraulic models.

12

Fig. 9. T h e configurational model of circulations in the surface layer between 0 t o 1,000 m (left upper panel), the upper mid-depth layer between 1,000 t o 2,000 m (left lower Panel), the lower mid-depth layer between 2,000 t o 3,000 m (right upper panel) and the deep layer between 3,000 t o 4,000 m (right lower panel) for t,he Philippine Sea and western North Pacific.

(7) development of new technologies t o make possible the more accurate measurements, quicker measurements and measurements of new qualities. These research activities have been organized in close mutual combination in accordance the flow diagram shown in Fig. 10.

References Fukasawa, M., and T. Teramoto, 1986. Characteristics of deep circulation off Cape Shionomisaki before and after formation of large meander of the Kuroshio in 1981. J. Oceanogr. SOC.Japan, 42(1), 53-68. Kaneko, I., and T. Teramoto, 1985. Water exchange between Shikoku-Philippine Basin and the North Pacific Ocean. In: Ocean Characteristics and their Changes, K. Kajiwara (ed.), Koseisha-Koseikaku, Tokyo, pp. 54-77 (in Japanese). Mogi, A., 1972. Bathymetry of the Kuroshio region. In: The Kuroshio, Its Physical aspects, H. Stommel and K. Yoshida (eds.), University of Tokyo Press, Tokyo, pp. 54-

1,

Numerical modeling Hydrauric modeling Data assimilation-----l

Exploratlon of mechanisms of ocean circulation

>-<

)- ,

Eulerian current measurements i&Lagranqian current measurements LCTD measurements

Delineation of deep ocean circulation

<

1

Preliminary presention of a hypothetical circulation model

Establishment of deep circulation model h

I Sediment trapping

of acoustic technology for ocean measurement

A

I

Exploration of substance cycle

chemical tracer distribution in the deep ocean

Fig. 10. T h e flow diagramme of the project research.

Water sampling and analysis

CL

w

14 80. Stommel, H., and A. Arons, 1960. On the abyssal circulation of the world ocean 11. Deep-sea Res., 6, 217-233.

Chapter 2 Structure of Subsurface Water Circulation and Associated Distribution of Water Properties

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

17

Deep Water Properties and Circulation in the Western North Pacific Masaki KAWABE*

Abstract Circulations of deep water in the Philippine Sea and the western North Pacific are examined from distributions of water properties, mainly c ~ 3and dissolved oxygen. 3 41.52 and 41.48 (around 4000 db), On the isopycnal surfaces of ( ~ between western boundary current, which carries high-oxygen water from the South Pacific, passes the Central Pacific and Melanesian Basins, and then flows northward through the central region of East Mariana Basin: it does not enter the Philippine Sea directly. On the other hand, the current on the 41.45 surface (about 3000 db) flows into the Philippine Sea and brings highoxygen waters directly, since the sea fully opens to the south of 12'30". Water circulation is greatly different between 41.48 and 41.45. The western boundary current seems to be weak on the 41.30 surface (about 2000 db). Waters between 41.51 and 41.48 flow into the Philippine Sea through the deepest gap of the Izu-Ogasawara-Mariana Ridge around 12'N, 139'E. Water at deeper than the gap does not enter a t all. The inflow water has clearly less oxygen than the water carried by the western boundary current. It spreads from the West Mariana Basin to the Shikoku and Philippine Basins. Oxygen on deep isopycnal surfaces in the Philippine Sea is much higher than that in the western North Pacific. This is probably caused by not only the inflow water having relatively high oxygen but also a convergence of upward oxygen flux.

1.

Introduction

Deep water cannot be produced in the North Pacific (Stommel, 1958; Warren, 1983), and it comes from the Antarctic after a long trip along the western boundary of the South Pacific. The mixed water of the Antarctic Bottom Water and the Circumpolar Water goes northward east of New Zealand and along the TongaKermadec Ridge, and enters the Central Pacific Basin through the Samoan Passage (Reid and Lynn, 1971; Warren, 1973; Reid and Lonsdale, 1974; Mantyla, 1975). The water originally shows a clear salinity maximum which is a remnant of the North Atlantic Deep Water, but the maximum becomes deeper and less clear due to mixing as it moves northward (Mantyla and Reid, 1983). Then, the core-shaped salinity maximum disappears north of the Samoan Passage, and the maximum *Ocean Research Institute, University of Tokyo, Minamidai, Nakano-ku, Tokyo 164, Japan

18 salinity of deep water is observed at the sea bottom in the North Pacific (Reid and Lynn, 1971). Concentration of dissolved oxygen also attains its maximum at the bottom. After passing through the Samoan Passage, flow paths are different by depth. The flow well examined so far is that of the Pacific Bottom Water which is steered by 5000-m isobaths. Part of the bottom water in the Central Pacific Basin flows eastward, passes through the Horizon and Clarion Passages, turns northward at the southeast of the Hawaiian Ridge, and finally enters the Northeast Pacific Basin (Edmond et al., 1971; Reid and Lonsdale, 1974; Mantyla, 1975). Other part flows out through the Wake Island Passage, and extends t o the Northwest Pacific and East Mariana Basins (Mantyla, 1975). The bottom water finally reaches the Northeast Pacific Basin through the northern margin of the Pacific and the passage north of the Hawaiian Ridge. Pioneer works of deep circulation in the North Pacific such as Wooster and Volkmann (1960) and Knauss (1962) also treat the near-bottom circulation controlled by 5000-m isobaths. At this depth, the Philippine Sea is perfectly separated from the main part of the North Pacific by the Izu-Ogasawara-Mariana Ridge. In the present study, water properties and circulation at lesser depths, mainly about 4000-m depths, will be examined. The difference in bottom topography by depth causes a difference in path of the western boundary current and in water exchange through the Izu-Ogasawara-Mariana Ridge (A to B in Fig. la) and the Kyushu-Palau Ridge (a t o b in Fig. la). The word “western North Pacific” will be used for the region except the Philippine Sea, although the word usually includes it. Mantyla and Reid (1983) concluded the inflow into the Philippine Sea through the deep gap north of the Yap Island and described briefly the water circulation in the Philippine Sea, on a basis of the global maps of water properties near the sea bottom. The near-bottom distributions are effective t o study a bottom-water move, but a slow adiabatic motion like a deep-water circulation actually occurs on a neutral surface, any motion along which does not change potential energy. A neutral surface at an arbitrary pressure coincides with the isopycnal surface referred t o the pressure (McDougall, 1987). Hence, the isopycnal surfaces with a large reference pressure must be used for the study of deep-water circulation (Reid and Lynn, 1971). The potential density referred t o 3000 db, ~ 7 3 ,will be used here. The deep inflow into the Philippine Sea was concluded also from the vertical distribution of water properties in the Philippine Sea (Sudo, 1982; Uehara and Taira, 1990), but it has not been confirmed with direct measurements of water current. I will examine the inflow by comparing the distributions of water properties with bottom topography and between the inner and outer regions of the Philippine Sea. It is well known that dissolved oxygen on an isopycnal surface in the Philippine Sea is much richer than that in the western North Pacific, as it is shown, for example, in the figures on isopycnal surfaces a t depths of about 1000 m, 2000 m and 3000 m in Kaneko and Teramoto (1985). This is a unexpected phenomenon t h a t the isolated marginal basin, where waters must be stagnating, is filled with

19 oxygen-rich, namely young waters. Many oceanographers are interested in this phenomenon, but it has not been described concretely and exactly. This is another issue in this paper.

2.

Data

Many hydrographic data were used in this study. The data sources are ocean expeditions INDOPAC (Scripps Institution of Oceanography, 1978), GEOSECS (National Science Foundation, 1982), EURYDICE (Scripps Institution of Oceanography, 1986) and WEPOCS (Joint Institute for Marine and Atmospheric Research, 1986) and Japanese cruises by R / V Hakuho Maru (Ocean Research Institute, 1971, 1975,1981, 1983,1988), Kaiyo Maru (Sudo, 1978, 1979), Ryofu Maru (Japan Meteorological Agency, 1982, 1989) and Takuyo (Japan Maritime Safety Agency, 1987). Moreover, from the historical hydrographic dataset kept by the Japan Oceanographic Data Center (JODC Reference 87-099), the data after 1964 taken down t o deeper than 3000 m were retrieved and used after checking d a t a quality. Data of isobath and water depth (5’ x 5’ mesh) were provided from the Japan Oceanographic Data Center (JODC Reference 84-037, 87-097), and were used t o draw isobaths and bottom topographies.

3.

Features of Massive Ridges

Bottom topographies steer deep-water current, and, as a result, have a decisive influence on distributions of water properties in deep layers. In particular, the ridge shown with the thick line of A to B in Fig. l a is greatly important to water exchange between the western North Pacific and the Philippine Sea. This main ridge is composed of some lesser ridges, i.e., the Shichito-Iojima Ridge, the West Mariana Ridge, and ridges connecting the islands of Yap, Palau and Sonsorol. It is generally called the Izu-Ogasawara-Mariana Ridge. A gap deeper than 4000 m exists only between 11’35” and 11’55’N, having a maximum depth of about 4450 m (Fig. l b ) . This is the deepest passage connecting the Philippine Sea with the western North Pacific. Gaps deeper than 3500 m increase to five. Three of them, including the deepest one, are located between 10’15” and 12’10” in the area which is often called the Yap-Mariana Junction. Other two gaps around 8’20” and between 1’50” and 2’30” connect the West Caroline Basin with the West Mariana and Philippine Basins, respectively. They are, however, expected to play much smaller roles in water exchange between the Philippine Sea and the outer regions than the gaps in the Yap-Mariana Junction, because of small width of the former gap and obstruction by many seamounts around the latter gap. At a depth of 3000 m, the deep gaps increase their widths, and some gaps add in the south of 8’40”. As a result, the ridge south of 12”30’N is much less effective as a barrier of water exchange, whereas the ridge north of 12”30’N is still a complete barrier. Passages of water increase more at a depth of 2000 m. At this depth, the ridge south of 30”N does not significantly prevent the water exchange between the Philippine Sea and the western North Pacific.

20 The Kyushu-Palau Ridge (a t o b in Fig. la) plays an important role in water exchange between the basins in the Philippine Sea. This ridge has four gaps deeper than 4000 m, which are located from 17'45" to 18"15'N, 19"N to 19"25'N, 22"N t o 23'20" (this divides into two gaps at deeper than 4050 m), and 23'55" t o 24'20" (Fig. lc). Another deep gap is shown around 30"30'N in Fig. lc, but it shoals t o about 4000 m just t o the west. Therefore, water exchange at deeper than 4000 m is only possible between the northern parts of the West Mariana and Philippine Basins. At depth less than 4000 m, passages of water much increase, and waters can also exchange between the southern parts of the West Mariana and Philippine Basins and between the Shikoku and Philippine Basins.

4.

Deep-Water Circulations Deduced from the Distributions of Water Properties

Figures 2 and 3 show vertical distributions of potential density referred t o 3000 d b ( a 3 ) ,potential temperature referred t o the sea surface, salinity and dissolved oxygen along about 13"N (S1 in Fig. la) and 23ON (S2). Density in the

~,

Rise

" Q

Fig. 1. (a) Map of the Philippine Sea and the western North Pacific. Thin lines are 4000-m isobaths. Thick lines of A to B and a t o b show the Izu-Ogasawara-Mariana Ridge and the Kyushu-Palau Ridge, respectively. The lines S1 and S2 indicate the sections, water properties on which are shown in Fig. 2 and 3. Dots on the lines are stations of hydrographic observation. (b) Section of the Izu-Qgasawara-Mariana Ridge (A t o B in Fig. la). (c) Section of the Kyushu-Palau Ridge (a t o b in Fig. la).

21 A

B 0

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10

,131

132

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.

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,

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8

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35N

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136

Fig. 1. (continued) western North Pacific shows simply layered stratification at all depths, while in the Philippine Sea a similar stratification is seen only at shallower than the isopycnal surface of u3 = 41.48 (about 3700 db), and the density stratification between 41.48 and 41.52 is different by basins and latitudes. 3 41.30, 41.45, On a basis of these characteristics, six isopycnal surfaces ( ~ = 41.48, 41.49, 41.50 and 41.52) are selected, and dissolved oxygen on these surfaces is mapped in Fig. 4. They are believed to reflect flow distributions. On the other hand, salinity is highly uniform on an isopycnal surface of 41.45 and more, and t h e uniformity increases with depth. Hence, salinity only on the 41.30 surface is shown in Fig. 5.

22 4.1. Deep-water circulation in the western North Pacific The isopycnal surface of C T ~= 41.52 exists between 4100 d b and 5100 d b in the western North Pacific (Fig. 2a, 3a). Dissolved oxygen on this surface shows that high-oxygen waters coming from the South Pacific flow northwestward through the Central Pacific and Melanesian Basins, then proceed northward in the central region of the East Mariana Basin, and reach around 25"N, 156"E (Fig. 4a). This is expected t o be a path of the western boundary current around 4000 db. Although the further downstream path cannot be supposed exactly from the oxygen map, distribution of oxygen higher than 3.8 ml1-I suggests that the current changes the direction t o eastward and northwestward. Part of the waters flowing eastward appears t o enter the Northeast Pacific Basin through the north of the Hawaiian Ridge, like the Pacific Bottom Water, while other parts turn northward along the Emperor Seamount Chain, and turn again westward along 35"N. The waters likely proceed north and south of the Shatsky Rise, and encounter the high-oxygen waters flowing northwestward from 25"N, 156"E. The concentration of oxygen much decreases as approaching t o the Izu-Ogasawara Ridge, and a further continuation of the circulation cannot be inferred. Similar circulation is suggested also on the isopycnal surfaces of C T ~ = 41.50, 41.49 and 41.48 (Fig. 4b,c,d). 4.2.

Inflow to the Philippine Sea in deep layers

The maximum density in the Philippine Sea is 41.517 and is found near the bottom in the region close t o the deepest gap in the Yap-Mariana Junction (Fig. 2a). The densest water is 1.080 or 1.099"C in potential temperature, 34.688 or 34.693 in salinity, and 3.73 ml 1-1 in oxygen (Fig. 2b,c,d). Waters having almost the same properties are found just outside the Philippine Sea at a depth of bottom of the deepest gap (INDOPAC Station at 10"52'N, 140"ll'E): 0 3 = 41.514, PT = 1.103"C, S = 34.693 and 0 2 = 3.74 ml 1-l. The depth of the gap is 4454 m according t o the depth data averaged in 5' x 5' grids, whose value corresponds t o 4522 db. Thus, the outside waters around the gap bottom occupy the deepest layer of the Philippine Sea near the gap, and the deeper waters in the western North Pacific, which are denser than 41.52, do not exist in the Philippine Sea. This strongly suggests that water in the western North Pacific flows into the Philippine Sea, but the water is limited to that at shallower than the bottom of the deepest gap. No other inflow at deeper than 4000 m is permitted by bottom topography (Sec. 3). The existence of the deep inflow into the Philippine Sea has been concluded by many authors (e.g., Sudo, 1982; Mantyla and Reid, 1983; Uehara and Taira, 1990), and they described that the water coming from the South Pacific flows into the Philippine Sea. However, the distribution of oxygen does not show a direct inflow of the oxygen-rich water coming from the South Pacific (Fig. 4b, c). The entrance region of the West Mariana Basin is occupied by relatively low-oxygen waters, which can be traced t o the western region of the East Mariana Basin. The inflow water may be carried by southward current along just east of the Mariana Ridge. Such current is actually shown in the map of geopotential anomaly referred

23 to 1200 d b in Moriyasu (1972). Another inflow at shallower than 4000 m occurs through the southernmost deep gap between the West Caroline and Philippine Basins. This is inferred from the existence of water of C T ~ = 41.49 in the southeast end of the Philippine Basin (Fig. 4c) and low oxygen in the south end of the basin on the 41.48 surface, which can be traced t o the West Caroline Basin (Fig. 4d). The distributions of water and oxygen do not suggest a significant outflow from the Philippine Sea on the 41.48 and deeper surfaces. Probably only inflow, at least net inflow, occurs through the gaps of the Izu-Ogasawara-Mariana Ridge. The inflow water must go t o the upper layer due t o the mass conservation, and, in other words, the upward mass flux occurs in the Philippine Sea. 4.3.

Deep-water circulation in the Philippine Sea

Waters entering through the deepest gap in the Yap-Mariana Junction occupy the West Mariana Basin on the surface of cr3 = 41.50 (Fig. 4b). The extension of water increases on the 41.49 surface; part of the water extends northward and enters the Shikoku Basin, while the remainder enters the northern region of the Philippine Basin through the deep gaps of the Kyushu-Palau Ridge between 17"N and 25"N (Sec. 3), and moreover extends t o the west and south in the Philippine Basin (Fig. 4c). This shows a water circulation along the sea bottom in the Philippine Sea. Similar description was yielded by Mantyla and Reid (1983). The waters denser than 41.50 spread northward along the western slope of the Mariana Ridge in the eastern region of the West Mariana Basin (Fig. 2a, 3a, 4b). This is consistent with the northeastward flow near the entrance shown by Uehara and Taira (1990), and seems to coincide with the flow along isobaths calculated numerically by Kubota and Ono (1992). Topographic /3 effect of the ridge slope may be essential to the bottom flow in the West Mariana Basin. The isopycnal surface of C T ~ = 41.49 deepens greatly over the Kyushu-Palau Ridge towards the Philippine Basin (Fig. 3a).l The 41.51 surface similarly deepens over the deepest gap of the Izu-Ogasawara-Mariana Ridge towards the West Mariana Basin, as the 41.51 surface in Fig. 2a is deeper than 4400 d b whose value is the depth of 41.51 at the just outer station. The 41.50 and 41.49 surfaces deepen downstream also in interior regions of the West Mariana and Philippine Basins, respectively (Fig. 2a, 3a). These show a large downstream gradient of horizontal pressure over the ridges and in the interior regions of the basins, since water motions are probably smaller at lesser depths. The pressure gradient in the bottomattached layer may balance with bottom friction, and may be essential t o the water movement of spreading over the ridges and on the sea bottom. Accompanied by the deepening of isopycnal surfaces, waters with C T ~more than 41.49 disappear gradually along the circulation. In other words, the density stratification at deeper than the 41.48 surface weakens in the downstream direction, and finally disappears in the southern part of the Philippine Basin. Potential temperature, salinity and oxygen also show similar changes (Fig. 2 and 3). As a result, the lNote that the Kyushu-Palau Ridge in Fig. 3 is a small mountain at 137'E, not a large one at 135'E. The large mountain is the southern end of the Oki-Daito Ridge.

24

I00

200

300

400 db

500

600

700 1

1000

2000

5000

4000 d b

5000

6000

7000

Fig. 2. Vertical distributions of water properties along S1 (about 13"N). (a) a 3 ( K P and M in the figure show the Kyushu-Palau and Mariana Ridges respectively), (b) potential temperature ("C), (c) salinity, (d) dissolved oxygen (ml 1-'). waters of a3 between 41.48 and 41.49 occupy the largest volume in the Philippine Sea. They have potential temperature between 1.2"C and 1.3"C, salinity between 34.68 and 34.69, and dissolved oxygen between 3.5 ml 1-' and 3.6 ml 1-'. The 41.48 surface is quite shallower than 4000 db, and the waters of v3 = 41.48

25

1000

2000

3000

4000 db 5000

6000

7000

Fig. 2.

(continued)

can exchange through many gaps of the Kyushu-Palau Ridge, not only between the northern parts of the West Mariana and Philippine Basins but also between their southern parts and between the Shikoku and Philippine Basins. In fact, westward flow from the West Mariana Basin to the Philippine Basin through the gap around ll"20'N is suggested by the low-oxygen belt probably made of the inflow water (Fig. 4d). This coincides with the westward flow at 12.5"N in the Philippine Basin shown by Uehara and Taira (1990) using a geostrophic calculation referred

26

120E

Fig. 3.

130

140

150

160

170

180

Same as Fig. 2 except for S2 (about 23"N). The arrows of

K P and OM in Fig. 3a show the Kyushu-Palau and Ogasawara-Mariana Ridges, respectively. to 1500 db. The extension of low oxygen suggests that the flow continues from the West Mariana Basin, and is confined meridionally. The flow is strong and almost barotropic below 3000 db according to the geostrophic calculation, while the oxygen on the 41.45 surface suggests a similar westward flow (Fig. 4e). The entire water between 41.45 and 41.49 appears to flow westward. Although oxygen

27 I20E

i30

i40

150

160

170

130

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150

160

170

I000

7000

3000

4000 db

5000

6000

1000 1 20E

180

1000

2000

3000

4000 d b

5000

6000

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Fig. 3. (continued) does not clearly show other flows, the waters less than 41.49 can exchange through many other gaps in the Kyushu-Palau Ridge. Water circulation at shallower than the 41.49 surface may be modified from that in the deeper layer close t o the bottom.

28 4.4.

Water circulation at shallower depths much less influenced by the ridges

At a depth of the surface of t73 = 41.45 (about 3000 db), the Izu-OgasawaraMariana Ridge opens south of 12'30". The Philippine Sea, the West and East Caroline Basins and the Melanesian Basin are linked directly. As a result, the waters with high oxygen above 3.2 ml l-l, coming from the South Pacific, flow directly into the Philippine Sea through these basins, without entering the East Mariana Basin (Fig. 4e). This is a large difference from the circulation on the deeper surfaces. Less-oxygen waters between 3.1 ml 1-1 and 3.2 ml 1-1 exist at the outer edge of the high-oxygen waters. Parts of them extend northward in the East Mariana Basin, while others enter the Philippine Sea and occupy its southeastern region. This suggests that a northward flow exists in the central area of the East Mariana Basin, as in the deeper layers, and a westward flow separates and enters directly the Philippine Sea. The less-oxygen inflow water passes over the ridge between 10"N and 12O30'N. The oxygen distribution, thus, suggests only inflow in the whole of passage south of 12"30'N. Absence of significant outflow from the Philippine Sea is the same as in the deeper layers. This concludes that a net vertical flux of water in the Philippine Sea is upward, at least, at deeper than the 41.45 surface, and the upward transport increases at a lesser depth. Since the area of the 41.45 surface in the Philippine Sea is little different from that of the 41.48 surface, the upwelling velocity must increase upward at least between these surfaces. This implies that the current is northward in the interior region of the Philippine Sea in which the linear steady balance of vorticity, ,!?v = f . dw/dz, is dominant, where f and ,!? are the Coriolis parameter and its latitudinal change rate, u and w are meridional and vertical velocities, and t is the upward coordinate. On the 41.30 surface (about 2000 db), the western boundary current is suggested from the distribution of oxygen higher than 2.6 mil-' (Fig. 4f). However, the highoxygen waters occupy only a little part of the Philippine Sea, and less-oxygen waters with large range between 2.3 ml 1-1 and 2.6 mi 1-1 occupy the wide area. The western boundary current may be less strong on this surface, and water conversion due t o lateral mixing may be effective. Since the Izu-Ogasawara-Mariana Ridge on this surface opens widely south of 30"N, it is possible that low- and high-oxygen waters flow into the Philippine Sea from the north and south respectively, and mixed waters with medium oxygen flow out in between.

5.

High Concentrations of Dissolved Oxygen on Isopycnal Surfaces in the Philippine Sea

Oxygen on a n isopycnal surface in the Philippine Sea is much richer than t h a t in the western North Pacific (Fig. 4). This characteristic is especially remarkable on the surfaces of t73 = 41.48 and 41.45. The region north of about 10"N in the western North Pacific on the 41.48 surface is occupied by water with oxygen lower than 3.4 ml l-', except a small region east of the north Japan (Fig. 4d). On the other hand, most region on the surface in the Philippine Sea is occupied by oxygen

29 higher than 3.4 ml 1-1 and partly by higher than 3.5 ml 1-l. It should be noted that high-oxygen waters more than 3.5 ml 1-1 do not come from the South Pacific, and they exist only in the almost closed basins: the Philippine Sea and the West and East Caroline Basins. Low-oxygen waters less than 3.2 ml 1-1 spread in the northeastern region of the western North Pacific and the region just east of the IzuOgasawara Ridge. Great difference in concentration of oxygen is found between east and west of the Izu-Ogasawara Ridge. Similar distribution is seen on the 41.45 surface (Fig. 4e). High-oxygen waters more than 3.2 ml 1-' occupy most of the Philippine Sea, while low-oxygen waters less than 3.1 ml 1-1 occupy most of the western North Pacific, and oxygens lower than 3.0 ml 1-1 exist in the northern region including just east of the Izu-Ogasawara Ridge. Oxygen and salinity on the 41.30 surface show large meridional variations decreasing northward (Fig. 4f, 5 ) . The isopleths of salinity extend almost zonally, and the distribution is not different between the Philippine Sea and the western North Pacific. On the other hand, the distribution of oxygen is different between them; oxygen in the Philippine Sea is rather uniform between 2.4 ml 1-1 and 2.6 ml l-l, while in the western North Pacific it shows a similar zonal structure t o salinity. The isopleths of oxygen between 1.8 ml 1-' and 2.3 ml 1-1 run meridionally over the Ogasawara Ridge, and they do not enter the Philippine Sea. Oxygen in the Philippine Sea is much higher than that in the western North Pacific on the 41.30 surface also. It should be noted that the difference in concentration of oxygen between the east and west of the Izu-Ogasawara Ridge is larger on shallower isopycnal surfaces. The difference reaches more than 0.6 ml 1-1 on the 41.30 surface. The characteristic of high oxygen in the Philippine Sea is seen also on the 41.49 and 41.50 surfaces (Fig. 4b, c) and on an isopycnal surface at about 1000 d b (Kaneko and Teramoto, 1985). This characteristic holds true in the thick layer down to the sea bottom, at least from 1000 db. Roughly describing, oxygen in the western North Pacific is high in the south region and low in the north. For example, just east of the Izu-Ogasawara-Mariana Ridge, it is low north of about 24"N and relatively high in the south (Fig. 4b to 4e). The main entrance to the Philippine Sea at deeper than the 41.48 surface, i.e., the deepest gap of the ridge, faces the south region filled with high-oxygen water. This is a first reason for the characteristic of high oxygen in the Philippine Sea. However, it is not all. We must notice the facts that oxygen in inner regions of the Philippine Sea is higher than that around the entrance, and that higher-oxygen waters than that coming from the South Pacific exists in the Philippine Sea on the 41.48 surface. These cannot be explained by lateral mixing because of no oxygen source in the Philippine Sea. Consumption of oxygen must result in a downstream decrease of oxygen, contrary t o the facts. The high oxygen may be formed and maintained by convergence of oxygen flux, which is mainly due to the advection of upward velocity. The upward flux from the lower layer is probably effective on this process for the layers around and less than 41.48. However, it is not available for deeper layers, e.g., the layers between 41.49 and 41.50 and between 41.50 and 41.51, because they

30

Fig. 4. Lateral distributions of dissolved oxygen (ml l-l) on isopycnal surfaces. (a) cr3 = 41.52, (b) 41.50, (c) 41.49, (d) 41.48, (e) 41.45, (f) 41.30. touch the sea bottom without a lower layer in most region. A mechanism of the high oxygen on the 41.49 and 41.50 surfaces may not be simple. Higher oxygen must be provided from the deeper layer at the entrance of the West Mariana Basin

31

l"2 0

130

140

150

160 E

I70

180

whiich disappears in a slightly downstream region. Effective transfer of oxygen may be accomplished through a bottom Ekman layer. On the 41.45 and shallower surfaces, the large inflow of high-oxygen water from thc2 South Pacific, carried by the western boundary current, must play a great role

32

N

30

in high oxygen in the Philippine Sea. This may cause the large difference in oxygen between the east and west of the Izu-Ogasawara Ridge on the shallow surfaces.

33

Fig. 5. = 41.30.

6.

Lateral distribution of salinity on the isopycnal surface of

g3

Summary

Characteristics and circulations of deep water in the North Pacific (0-40°N, 120"E-180") are examined using vertical distributions of water properties along 13"N and 23"N and lateral distributions of oxygen on six isopycnal surfaces of ( ~ 3= 41.30, 41.45, 41.48, 41.49, 41.50 and 41.52. Salinity is highly uniform on the 41.45 and deeper isopycnal surfaces. The degree of uniformity increases with depth. Therefore, paths and characteristics of deep flow are studied mainly using distributions of dissolved oxygen. Density stratification in the Philippine Sea is almost common everywhere at shallower than the 41.48 surface (about 3700 db), while it is spatially different in deeper layers between 41.48 and 41.52. The deep-layer stratification gradually weakens downstream along the circulation of bottom water, and finally the density becomes homogeneous in the south region of the Philippine Basin. As a result, the water with C J ~ between 41.48 and 41.49 occupies the largest volume in the Philippine Sea. Potential temperature, salinity and dissolved oxygen show almost the same stratifications as density. The respective properties of the largest volume water are between 1 2 ° C and 1.3"C, 34.68 and 34.69, and between 3.5 ml 1-1 and 3.6 ml 1-l. Since flows along the sea bottom steered by 5000-m isobaths have been examined well previously, shallower flows are examined in this study. On the 41.52 surface (a little deeper than 4000 db), the western boundary current carrying high-oxygen waters from the South Pacific flows northwestward through the Melanesian and

Central Pacific Basins, and then proceeds northward in the central region of the East Mariana Basin. The flow separates t o eastward and northwestward branches around 25"N, 156"E. The eastward flow appears t o separate again: one proceeds further eastward and enters the Northeast Pacific Basin through the north of the Hawaiian Ridge, while the other turns t o north along the Emperor Seamount Chain, turns again t o west along 35"N, and, after passing the Shatsky Rise, joins the other branch flow from 25"N, 156"E. Largest difference from the bottom flow results from the difference of topographic restriction that the current on the 41.52 surface can pass the Melanesian Basin, which closes for the bottom water. Main circulation of water between 41.48 and 41.51 is almost the same as that of 41.52, since the western boundaries in these layers are little different. However, there is a difference in terms of inflow into the Philippine Sea. Major entrance for waters between 41.48 and 41.51 is the deepest gap of the Izu-Ogasawara-Mariana Ridge located around 12"N, 139"E. The waters at deeper than the gap, which are denser than 41.52, do not enter the Philippine Sea, and less dense waters flow into there. Although less important, a gap at the southern end of the ridge is another entrance for the 41.48 and 41.49 waters. It should be noted that the high-oxygen waters coming from the South Pacific do not directly enter the Philippine Sea on these surfaces. Less-oxygen waters, which are probably formed by lateral mixing between the high-oxygen water from the South Pacific and low-oxygen water from the north, seem t o be carried t o the Philippine Sea by southward flow along just east of the Mariana Ridge. The inflow waters denser than 41.50 spread northward along the bottom slope in the west flank of the Mariana Ridge, and water of 41.50 spreads almost everywhere in the West Mariana Basin. Water of 41.49 extends to the Shikoku Basin, while it enters the Philippine Basin through the gaps of the Kyushu-Palau Ridge between 17"N and 25"N. It moreover extends to the west and south in the Philippine Basin. This is the water circulation in the Philippine Sea creeping along the sea bottom. Shallower circulations are expected to be modified by flows through many gaps of the Kyushu-Palau Ridge, such as a westward flow through the gap at ll"20'N. The 41.49 and deeper surfaces change the depth spatially; horizontal pressure in the bottom-attached layer has a downstream gradient over the gaps of the IzuOgasawara-Mariana and Kyushu-Palau Ridges as well as in interior regions of the West Mariana and Philippine Basins. The pressure gradient may balance with bottom friction, and may be essential to the inflow through the gaps and water spread in the basins along the sea bottom. Largest change of water circulation, including the western boundary current, is seen between the isopycnal surfaces of 41.48 and 41.45. On the 41.45 surface (about 3000 db), the entrance of the Philippine Sea greatly enlarges and fully opens south of 12"30'N, so that the western boundary current carrying high-oxygen waters flows directly into the Philippine Sea from the South Pacific by the way of the Melanesian Basin and the West and East Caroline Basins. The inflow transport on the 41.45 surface must be much larger than that on the 41.48 and deeper surfaces. Oxygen maps suggest only the inflow into the Philippine Sea for the 41.45 and deeper surfaces. This results in a upward flux of water in the Philippine Sea which

35 increases as a depth decreases. This leads t o the conclusion t h a t the upwelling velocity increases upward, at least, between the surfaces of 41.45 and 41.48, since the area of surface in the Philippine Sea is little different between them. This implies that the current at least in this layer is northward in a wide interior region of the Philippine Sea, where vorticity is balanced between the p and horizontaldivergence terms. Path of the western boundary current on the 41.30 surface (about 2000 db) is similar to that on the 41.45 surface, but distributions of oxygen and salinity are clearly different between these surfaces. On the 41.30 surface, they show large meridional variations with a northward decrease, and, in particular, salinity remarkably shows a zonal band structure. In the Philippine Sea, high-oxygen waters from the South Pacific occupy little part, and concentration of oxygen has a large range. The inflow flux by the western boundary current must not be dominant, while water conversion due to lateral mixing may be effective. Outflows from the Philippine Sea and other inflows may occur on the 41.30 surface, unlike the 41.45 and deeper surfaces. High oxygen in the Philippine Sea and low oxygen just east of the Izu-Ogasawara Ridge are common characteristics on every isopycnal surfaces between 41.30 and 41.50; so that, the concentration of oxygen is greatly different between the east and west of the ridge. The difference is larger a t lesser depth, reaching more than 0.6 ml 1-l on the 41.30 surface. This is probably due to the large oxygen flux into the Philippine Sea carried by the western boundary current on the 41.45 and shallower surfaces. Two reasons are considered for higher oxygen in the Philippine Sea. One is that the inflow waters have relatively high oxygen, because the entrance of the Philippine Sea is located in the south region filled with high-oxygen waters. The other is a convergence of oxygen flux caused by advection of upward velocity. The latter reason possibly explains t h a t oxygen in inner regions is higher than that near the entrance, and that higher oxygen than that from the South Pacific exists in a part of the Philippine Sea on the 41.48 surface.

Acknowledgment

I thank Dr. H. Sudo for helpful discussion and providing files of the Kaiyo-Maru data, Dr. I. Kaneko for providing d a t a files which he compiled and used in his study, the Japan Oceanographic Data Center for providing historical hydrographic data, and Mrs. M. Kawabe for kind help t o compile d a t a files, draw useful figures and so on. I also thank Drs. T. Teramoto and K. Taira for giving me a chance t o study deep-water circulation. References Edmond, J. M., Y. Chung, and J. G. Sclater, 1971. Pacific Bottom Water: Penetration east around Hawaii. J. Geophys. Res., 76, 8089-8097. Japan Maritime Safety Agency, 1987. Results of oceanographic observations for WESTPAC 1986. Data Report of Hydrographic Observations, No. 2, 79 pp. Japan Meteorological Agency, 1982. The Results of Marine Meteorological and Oceanographical Observations, No. 70, 261 pp.

Japan Meteorological Agency, 1989. The Results of Marine Meteorological and Oceanographical Observations, No. 79, 360 pp. Joint Institute for Marine and Atmospheric Research, University of Hawaii, 1986. Hydrographic Observations from the First U.S. Cruise of the Western Equatorial Pacific Ocean Circulation Study (WEPOCS), June-July 1985. R. Lukas, and M. Tsuchiya (ed.), 350 pp. Kaneko, I., and T . Teramoto, 1985. Sea water exchange between the Shikoku-Philippine Basin and the western North Pacific Basin. In: Ocean Characteristics and their Changes, K. Kajiura (ed.), Koseisha-Koseikaku, pp. 54-77. Knauss, J. A., 1962. On some aspects of the deep circulation of the Pacific. J. Geophys. Res., 67, 3943-3954. Kubota, M., and K. Ono, 1992. Abyssal circulation model of the Philippine Sea. Deep-sea Res., 39, 1439-1452. Mantyla, A. W., 1975. On the potential temperature in the abyssal Pacific Ocean. J. Mar. Res., 33, 341-354. Mantyla, A. W., and J. L. Reid, 1983. Abyssal characteristics of the World Ocean waters. Deep-sea Res., 30, 805-833. McDougall, T. J., 1987. Neutral surfaces. J. Phys. Oceanogr., 17 , 1950-1964. Moriyasu, S., 1972. Deep waters in the western North Pacific. In: Kuroshio - Its Physical Aspects, H. Stommel and K. Yoshida (ed.), Univ. of Tokyo Press, Tokyo, pp. 387-408. National Science Foundation, 1982. Geochemical Ocean Sections Study (GEOSECS) Pacific Expedition, Vol. 3, Hydrographic Data 1973-1974. w . S. Broecker, D. w. Spencer and H. Craig (ed.), 137 pp. Ocean Research Institute, University of Tokyo, 1971. Preliminary Report of the Hakuho Maru Cruise KH-70-1, 3 February - 5 March 1970. Y. Horibe (ed.), 46 pp. Ocean Research Institute, University of Tokyo, 1975. Preliminary Report of the Hakuho Maru Cruise KH-73-1 (IBP Cruise), 10 January 6 February 1973. A. Hattori (ed.), 45 PP. Ocean Research Institute, University of Tokyo, 1981. Preliminary Report of the Hakuho Maru Cruise KH-75-3 (Lyra Expedition), 14 May -- 2 June 1975. Y. Horibe (ed.), 76 PP. Ocean Research Institute, University of Tokyo, 1983. Preliminary Report of the Hakuho Maru Cruise KH-80-2 (CYGNUS Expedition), 25 April - 18 June 1980, and KH-82-1 (CEPHEUS Expedition), 22 January - 17 March 1982. Y. Horibe (ed.), 78 pp. Ocean Research Institute, University of Tokyo, 1988. Preliminary Report of the Hakuho Maru Cruise KH-87-1 (WESTPAC), 22 January - 10 March 1987. T. Teramoto and K. Taira (ed.), 139 pp. Reid, J. L., and P. F. Lonsdale, 1974. On the flow of water through the Samoan Passage. J. Phys. Oceanogr., 4, 58-73. Reid, J. L., and R. L. Lynn, 1971. On the influence of the Norwegian-Greenland and Weddell seas upon the bottom waters of the Indian and Pacific oceans. Deep-sea Res., 18, 1063-1088. Scripps Institution of Oceanography, University of California, 1978. Data Report, Physical, Chemical and Biological Data, INDOPAC Expedition, 23 March 1976 - 31 July 1977. SIO Reference 78-21, 424 pp. Scripps Institution of Oceanography, University of California, 1986. Data Report, Physical, Chemical and Current Meter Data, EURYDICE Expedition Leg X, 10 May - 1 2 June 1975. SIO Reference 86-20, 30 pp. Stommel, H., 1958. The abyssal circulation. Deep-sea Res., 5, 80-82. ~

37 Sudo, H., 1978. Physical data from the hydrographic survey on marine organism in relation to deep-sea disposal of radioactive wastes in the central western North Pacific aboard R. V. Kaiyo Maru in 1972 to 1974. Datum Collect. Tokai Reg. Fish. Res. Lab., 7, 1-40. Sudo, H., 1979. Physical data from the hydrographic survey on marine organism in relation to deep-sea disposal of radioactive wastes in the central western North Pacific aboard R. V. Kaiyo Maru in 1978 and 1979. Datum Collect. Tokai Reg. Fish. Res. Lab., 8, 1-48. Sudo, H., 1982. Hydrographic surveys in the central western North Pacific on the deep-sea disposal of radioactive wastes - Deep water characteristics and circulation. Japan Fisheries Agency, 174 pp. (in Japanese with English abstract and legends) Uehara, K., and K. Taira, 1990. Deep hydrographic structure along 12'N and 13'N in the Philippine Sea. J. Oceanogr. SOC.Japan, 46, 167-176. Warren, B. A , , 1973. Transpacific hydrographic sections at Lats. 43's and 28's: The SCORPIO Expedition - 11. Deep water. Deep-sea Res., 20, 9-38. Warren, B. A., 1983. Why is no deep water formed in the North Pacific? J. Mar. Res., 41, 327-347. Wooster, W. S., and G. H. Volkmann, 1960. Indications of deep Pacific circulation from the distribution of properties at five kilometers. 3. Geophys. Res., 65, 1239-1249.

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39

Density Field along 12"N and 13"N in the Philippine Sea Katsuto UEHARA, Keisuke TAIRA and Akira MASUDA*

Abstract

Hydrographic casts down to the bottom along two zonal sections at 12'N and 13'N (from 144'E to 127'E) were made with a CTD. Their analysis verified the existence of cold and saline abyssal water between the Mariana Ridge and the Kyushu-Palau Ridge. This result provides evidence of flow into the Philippine Sea through the gap called the Yap-Mariana Junction. The properties of deep water are variable in the West Mariana basin but quite homogeneous in the Philippine Basin, indicating the transitional nature in the West Mariana Basin and the existence of older bottom water in the Philippine Basin.

1.

Introduction

The Philippine Sea is the sea which locates to the south of Japan and comprises the westernmost part of the North Pacific Ocean (Fig. 1). For the depth below 4000 m, the Izu-Ogasawara and Mariana Ridges separate the Philippine Sea from the main Pacific Ocean and restricts the horizontal exchange of abyssal water except through the several narrow gaps. In reality, the map of the horizontal distribution of temperature, salinity, and dissolved oxygen compiled by Moriyasu (1972) using CSK data shows an apparent difference between the Philippine Sea and the main Pacific Ocean for the levels below 3000 m. In the North Pacific Ocean, including the Philippine Sea, the surface water cannot be dense enough t o form an existing abyssal water (Warren, 1983) and the abyssal water characterized by low salinity and high oxygen content are considered t o be supplied from the south (Stommel and Arons, 1960; Warren, 1981). Accordingly, abyssal water of the Philippine Sea must be supplied from outside. Additional data, including those obtained from the Indopac Expedition, have provided some information on this subject. The distribution of water characteristics shown by Mantyla and Reid (1983) and by Kaneko and Teramoto (1985) suggest a flow of bottom water into the Philippine Sea through the gap to the southwest of Guam, which is called the Yap-Mariana Junction. However, Mantyla and Reid (1983) used hydrographic data at the bottom layer only and the depth is different from station t o station while Kaneko and Teramoto (1985) analyzed the data mainly in upper 3000 m. The inflow through the Yap-Mariana Junction is not fully delineated. *Ocean Research Institute, University of Tokyo, 1-15-1, Minamidai, Nakano-ku, Tokyo 164 Japan

40 In order t o detect this inflow more clearly, hydrographic observations were carried out in the southern Philippine Sea along 12"N and 13"N, from 144"E t o 127"E during R/V Hakuho-Maru Cruise KH-87-1 leg 2 of 8-17 February 1987. In this manuscript, we made an analysis of the CTD dataset obtained from this cruise. Emphasis is placed on depicting the hydrographic structure of abyssal and bottom layers in the southern Philippine Sea. The existence of distinctly cold, saline and dense bottom water in the West Mariana Basin is shown in cross sections of 12"N and 13"N. We also examined the vertical profile of each station for the depth below 3000 db. We found homogeneous distribution in the Philippine Basin and rather variable structures in the West Mariana Basin. Also, in the West Mariana Basin, water tends t o get warmer, less saline, and less dense as it gets farther from the Yap-Mariana Junction. These results provide strong evidence for the inflow of cold, saline, and dense bottom water through the Yap-Mariana Junction into the West Mariana Basin. In the

N -

30"

20"

10"

0"

120"

130"

140"

E

150"

Fig. 1. Topography around the Philippine Sea and the hydrographic stations of the Hakuho Maru Cruise KH-87-1 leg 2. Numerals represent station numbers.

41 30.

25.

20.

W

a

13

?

15.

W

Q

2

W

t-

J

2

10.

2

W

-t

2

5.

0.

34.2 34.4

34.6 34.8

35.0 35.2

SALINITY(PSU)

Fig. 2. T-S diagram of the CTD data. Temperature is shown in potential temperature referred t o 0 db pressure. Isopycnal lines of potential density referred t o 0 d b is shown as contour lines. Philippine Basin, the bottom water was slightly colder and denser in the western part than in the eastern part. This may be indicating the existence of a broad western boundary current toward the south along the eastern rise of the Philippine Trench.

2. 2.1.

0bservat ion Measurement and data analysis

The topography of the Philippine Sea and the location of the hydrographic stations are shown in Fig. 1. The Izu-Ogasawara Ridge and the Mariana Ridge

42 separate the Philippine Sea from the North Pacific Ocean at a depth greater than 3000 m. The Kyushu-Palau Ridge along 135"E divides the Philippine Sea into the Philippine Basin and the West Mariana Basin. There is a gap with a sill depth of 5000 m t o the south of Guam called the Yap-Mariana Junction, which connects the Philippine Sea with the North Pacific Ocean. Hydrographic stations were occupied with a Neil Brown Mark IIIb C T D and a Rosette multi-sampler. At each station, CTD d a t a were recorded t o the bottom layer of 3000-6000 d b and water sampling was made at 1 2 levels. A zigzag course was adopted to examine the hydrographic structure at both 12"N and 13"N. Dense CTD casts were made around 135"E to examine possible effects of the KyushuPalau Ridge on the bottom water distribution. Calibration of CTD d a t a were made by referring the d a t a obtained from sampled water and reversing thermometer to reduce the systematic error of the C T D sensor. Afterwards, the d a t a are recompiled into 2 d b interval data. For station 209 (below 4000 db) and station 222, we could not get enough number of valid d a t a for station 209 (below 4000 db) and station 222, so they are not included in the dataset. Estimated error of the dataset is less than 0.01"C for temperature and 0.0037 psu (practical salinity unit) for salinity.

2.2.

Results

Surface layers In the surface layer of upper 200 db, strong baroclinic feature is found. Geostrophic calculation shows a surface confined westward flow with intensity exceeding 20 cm/s (not shown), which indicates the North Equatorial Current (e.g., Toole et al., 1990). There is a salinity maximum exceeding 35.0 psu at about 110-130 d b in depth (Fig. 2). Beneath this thin layer, there is a salinity minimum a t the depth of 600-800 d b or around 26.5 in sigma-theta (potential density referred to 0 pressure). In the succeeding section, we will deal with the deep and bottom layers below the salinity minimum.

Cross sections of 12"N and 13"N Figures 3(a), (b), and (c) show the cross-sectional distributions along 12"N of potential temperature theta-4 referred to 4000 db, salinity S, and potential density sigma-5 referred to 5000 db, respectively. Figures 4(a), (b), and (c) show the crosssectional distributions along 13"N. Above 3500 db, which is roughly the depth of the Kyushu-Palau Ridge, no apparent difference can be seen between the water properties in the West Mariana Basin and those in the Philippine Basin. Obvious differences, however, appear below 3500 db; the most prominent feature is the marked contrast of water properties between both sides of the Kyushu-Palau Ridge around 135"E. Colder, more saline, and denser abyssal water exists in the West Mariana Basin (eastern side of the ridge), whereas it is absent in the Philippine Basin (western side of the ridge). Vertical and horizontal distributions are highly variable in the West Mariana Basin. This is an indication that the West Mariana Basin is nearer t o the source of abyssal water. Dissolved oxygen content measured

43

-4--

3000

c

4000

5000

6000

1

125"

LONGITUDE

u5

130"

135' LONGITUDE

140'

E

(12"N)

1000

2000

125'

130'

135" LONGITUDE

140'

E

Fig. 3. The distributions along 12"N cross-section: (a) potential temperature theta-4 (referred t o 4000 db), (b) salinity S, and (c) potential density sigma-5 (referred t o 5000 db). The numerals at the top denote the station numbers. from sampled water shows higher values at the bottom layer in the West Mariana Basin (Watanabe et al., 1988), indicating that abyssal water is relatively young there. In the Philippine Basin, water of slightly lower temperature and large density is found in the western bottom layer as indicated by isotherms of 1.565"C and 1.57"C in Figs. 3(a) and 4(a) along both 12"N and 13"N. This water may spread over few

44

---

-

3000

D

a

__--

E

r

2

3000

-

D

4000

2 4000

5000

5000

nil-----

6000

6000 125'

~

130"

135'

E

140'

125"

130"

135"

140'

E

LONGITUDE

LONGITUDE

u5 (13'N)

---

3000

Iz( 9

P

4000

5000

6000

125'

130"

135'

140"

E

LONGITUDE

Fig. 4. The same as Fig. 3 except along 13"N hundred kilometers, although the westernmost part of the basin was not observed. It is not clear whether this water is significantly colder, because the temperature difference is less than 0.005"C which is as large as half the measurement error. T h e bottom water in the western Philippine basin has properties closer to the bottom water in the West Mariana Basin. These features indicate the existence of a broad and weak southward flow in the western part of the Philippine Basin.

45 Vertical profile of all CTD stations Figures 5(a), (b) show the vertical profiles of theta-4, S, and sigma-5 for depth below 3000 db, respectively. The profiles are drawn with different types of lines according t o the following geographical classification: 3000

3500

a: 3

4500

0)

0 W

a: a

5000

5500

Fig. 5. The vertical distributions of (a) theta-4, (b) S, and (c) sigma5 below 3000 db: - -, on the south-eastern flank of the Mariana Ridge; - _ - , in the West Mariana basin; - - -,in the eastern part of the Philippine Basin; and - in the western part of the Philippine Basin. An error bar is inserted in each figure.

46

I44E-127E

P0

TE

NT I A

Fig. 5 .

L D E N S I T Y ( k g Irn

1 )

(continued)

(1) the southeastern flank of the Mariana Ridge (stations 202 and 204; outside of the Philippine Sea), (2) near the shallowest portion of the Mariana Ridge (stations 201 and 2031, ( 3 ) the West Mariana Basin (stations 205-210), (4) near the Kyushu-Palau Ridge (stations 211-213, 008, and Oll), which is about 3500 m deep, ( 5 ) eastern part of the Philippine Basin (stations 214-219), (6) western part of the Philippine Basin (stations 220-223).

For stations 201 and 203, profiles are not included in the figures because depths are shallower than 2000 db. The profiles for stations 222 and 209 (below 4000 db) are not available because the raw data are erroneous. Above 4000 db, deep water southeast of the West Mariana Ridge (stations 202 and 204) is the coldest and most saline when compared at the same depth. The water retains the characteristics of the abyssal water from the south. The vertical profiles of region (1) are distinguished from those inside the Philippine Sea (regions 3 t o 6) at depths shallower than 4000 db. In the Philippine Sea, the profiles for stations in the West Mariana Basin (regions 3 and 4) are separated from those in the Philippine Basin (regions 5 and 6) below 3500 db. The depth corresponds t o the sill depth of the Kyushu-Palau Ridge, which divides the Philippine Sea into the two basins. The water in the Philippine Basin is remarkably homogeneous below 4000 db. It is warmer, less saline, and less dense compared t o the water in the West Mariana Basin. This implies that the Philippine Basin is far from the source of abyssal water. On the other hand,

47 2.0

1.5

1.0

0.5

0.0

- 0.5

Depth (db)

1000

2000

3000

4060

5000

6000

Depth ( d b )

Fig. 6. The vertical distributions of stability for the stations (a) 205210, (b) 214-223 for the depth below 1000 db, respectively. Stability is defined in terms of the Brunt-V ais a1 a frequency (cph). Five adjacent data (10 d b in depth) were fitted t o calculate a density gradient. Positive value represents the stable state, i.e., the density gets denser as the depth increases.

48 the water in the West Mariana Basin has a large variation within a relatively small zonal scale. In the West Mariana Basin, water properties are changing rapidly from those in the North Pacific Ocean to those in the Philippine Basin. It is to be noted that station 206 rather than 205, has a vertical distribution similar t o those at stations 202 and 204 in the western North Pacific Ocean. This is most clear for salinity profiles as shown in Fig. 5(b). Station 206 is nearest to the Yap-Mariana Junction which is located at around 11"N and 139"E (Fig. 1). Accordingly, we can state that abyssal water flows into the Philippine Sea through the Yap-Mariana Junction. When we compare the vertical profiles for regions (5) and (6) in the Philippine Basin, rather small though systematic difference can be found in the temperature and density profiles. The western side of the Philippine Basin is slightly colder and denser than the eastern side as discussed in the previous section.

3. Summary and Discussion CTD measurements were made down to the bottom along 12"N and 13"N in the Philippine Sea. The most important result is that the existence of cold and saline water was verified in the deep region of the West Mariana Basin below 3500 db. Water of these properties has been inferred t o enter the Philippine Sea through the Yap-Mariana Junction in previous data analyses by Mantyla and Reid (1983) and Kaneko (1984). At station 206, which is the station nearest to the Yap-Mariana Junction, we observed the coldest, most saline, and densest bottom water in the Philippine Sea. The marked contrast of water properties between the West Mariana Basin and the Philippine Basin can be ascertained from the instrumental accuracy. Also both the horizontal and vertical gradients are compared between the east and west sides of the Kyushu-Palau Ridge. The deep water has a variable structure in the West Mariana Basin, but is homogeneous in the Philippine Basin. We may say, therefore, that the present result gives evidence of flow into the Philippine Sea through the Yap-Mariana Junction. The vertical profiles of stability of the water in the West Mariana Basin and the Philippine Basin for the depth below 1000 d b are shown in Figs. 6(a) and (b), respectively. In the Philippine basin, the stability is almost neutral below the depth of 4200 db. This suggests that although the difference of the water properties between the West Mariana Basin and the Philippine Basin become apparent from 3500 d b (Fig. 5a and b), there may be a gap at the Kyushu-Palau Ridge with a depth of 4200 d b and the water underneath retains the characteristic of the sill depth. In the West Mariana Basin, on the other hand, we can see some signals showing the variable nature of water properties: the variations of the stability between the stations are large compared with those in the Philippine Basin, and the stability does not come to zero value (i.e., there are potential density gradients at most depths) except for the depth below 5000 d b at station 206. It is not clear from our dataset whether the homogeneous distribution of density below 5000 d b reflects the sill depth of the Yap-Mariana Junction because there is only one station which has d a t a deeper than 5000 d b in the West Mariana Basin. The water density in the western part of the Philippine Basin increases slightly

t o the west (Figs. 3, 4 and 5 ) . This might be interpreted as a broad western boundary current flowing equatorward. It is impossible t o substantiate this speculation at present. At least, t h e variation is quite small, beyond t h e present precision of data.

Acknowledgement We would like t o thank Prof. T. Teramoto, Dr. M. Kawabe, a n d Mr. S. Kitagawa for their support and discussions during t h e study. We also thank the crews a n d t h e participants of t h e cruise of t h e R/V Hakuho Maru.

References Kaneko, I., 1984. Structure of mid-depth water in the Philippine Sea. Ph. D. Thesis, University of Tokyo, Tokyo, 97 pp. Kaneko, I., and T. Teramoto, 1985. Sea water exchange between the Shikoku-Philippine Basin and the western North Pacific Basin. In: Ocean Characteristics and their Changes, K. Kajiura (ed.), Koseisha-Koseikaku, pp. 54-77. Mantyla, A. W., and J. L. Reid, 1983. Abyssal characteristics of the World Ocean Waters. Deep-sea Res., 30, 805-833. Moriyasu, S., 1972. Deep waters in the western North Pacific. In: Kuroshio - Its Physical Aspects, H. Stommel and K. Yoshida (ed.), Univ. of Tokyo Press, Tokyo, pp. 387-408. Stommel, H., and A. B. Arons, 1960. On the abyssal circulation of a world ocean - 11. An idealized model of the circulation pattern and amplitude in oceanic basins. Deep-sea Res., 6, 217-244. Teramoto, T., and K. Taira, 1988. Preliminary Report of The Hakuho Maru Cruise KH-87-1, Ocean Research Institute, University of Tokyo, 38 pp. Toole, J. M., R. C . Millard, Z. Wang, and S. Pu, 1990. Observations of the Pacific North Equatorial Current Bifurcation at the Philippine Coast. J. Phys. Oceanogr., 20, 307-318. Warren, B. A., 1981. Deep circulation of the world ocean. In: Evolution of Physical Oceanography, B. A. Warren and C. Wunsch (ed.), MIT Press, Cambridge, pp. 6-41. Warren, B. A., 1983. Why is no deep water formed in the North Pacific? J. Mar. Res., 41, 327-347. Watanabe, S., S. Noriki, C. Saitoh, and Y . Watanabe, 1988. Distribution of chemical tracers in the Philippine Sea. In: Preliminary Report of The Hakuho Maru Cruise KH-87-1, T. Teramoto and K. Taira (ed.), Ocean Research Institute, University of Tokyo, pp. 30-34.

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Deep Ocean Circulation, Physical and Chemical Aspects

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51

Abyssal Boundary Current along the Northwestern Perimeter of the Philippine Basin Masaaki CHAEN: Masao FUKASAWA: Akio MAEDAt Masahito SAKURAIf and Masaki TAKEMATSUf

Abstract Abyssal currents along the northwestern perimeter of the Philippine Basin were measured by moored current meters at three sites, which are located southeast of Okinawa, southeast of Taiwan, and the northern extremity of the Kyushu-Palau Ridge. Total length of observation was about 550 days at each site. Southeast of Okinawa and southeast of Taiwan, mean currents with magnitudes of 4.3 cm/sec and 1.3 cm/sec, respectively, were observed a t 400 m above the sea bottom at stations where water depth is greater than 4,000 m. The mean currents were parallel to the local bottom contour and flowing with the upslope on their right. Intensification of the mean current toward the bottom is clearly found on the slope of southeast of Okinawa. At the northern extremity of the Kyushu-Palau Ridge, the observed current does not show a stable abyssal current with a predominant direction, although the site is located at the narrow, deep passage between the Shikoku Basin and the northern portion of the Ryukyu Trench. Strong currents with opposite directions occurred intermittently and alternately. These results suggest that a steady abyssal current exists at depths greater than about 3,000 m along the northwestern perimeter of the Philippine Basin, and that the deep water exchange between the Shikoku and Philippine Basins occurs intermittently. Existence of an abyssal circulation system is also suggested in the northern part of the Philippine Basin, separated from the circulation in the northern part of the Shikoku Basin as suggested in previous studies.

1.

Introduction

T h e Philippine Sea is a marginal sea, separated from the Pacific Ocean by t h e Shichito, Ogasawara and Mariana Ridges. Figure 1 shows a schematic m a p of the topography around the Philippine Sea adapted from Mogi (1972). At depths greater than 3,000 m, the Philippine Sea is connected t o the North Pacific Ocean only through a few trenches caved at the eastern and western ends of the Yap Ridge, *Faculty of Fisheries, Kagoshima University 'Faculty of Marine Science and Technology, Tokai University *Faculty of Engineering, Kagoshima University §Research Institute for Applied Mechanics, Kyushu University

52 where Yap Islands are located. In the center of the Philippine Sea, the north-south oriented Kyushu-Palau Ridge separates the Philippine Sea into the eastern and western parts at depths greater than 3,500 m. The northern and southern basins of the eastern part of the Philippine Sea are the Shikoku and West Mariana Basins, respectively. The basin of the western part is called the Philippine Basin. The Philippine Sea is a favorable field for us t o carry out an experiment concerning the abyssal circulation. Stommel and Arons (1960a; 1960b) proposed the

Fig. 1. Topographic features of the Philippine Sea including basins, ridges, and trenches after Mogi (1972; 1979) and location of current meter mooring sites (black circles). Basins A: Amami Basin, B: Kita-Daito Basin, C: Minami-Daito Basin. Ridges 1: Shichito-Iwoto Ridge, 2: NishiShichito Ridge, 3: Ogasawara Ridge, 4: Mid-Mariana Ridge, 5: East Mariana Ridge, 6: West Mariana Ridge, 7: Yap Ridge, 8: Kyushu-Palau Ridge, 9: Amami Plateau, 10: Daito Ridge, 11: Oki-Daito Ridge. Trenches a: Nankai Trough, b: Ryukyu Trench, c: Philippine Trench, d: Yap Trench, e: Palau Trench.

53 model for the abyssal circulation in the world ocean. In their model, the Philippine Sea is completely neglected because it is almost separated from the North Pacific Ocean at abyssal depths and does not contain any source of deep water in it. However, their abyssal circulation model could be applied for the Philippine Sea, because it is large enough having the horizontal scale longer than 2,000 km, is deep enough having a distinct thermocline, and has a source of deep and bottom waters. If the abyssal circulation in the Philippine Sea is clarified, it is also useful to understand the abyssal circulation in the world ocean. In the past ten years, intensive studies have been made for the Philippine Sea. Concerning the abyssal boundary current associated witfh the abyssal circulation system, Fukasawa and Teramoto (1986) and Fukasawa e t al. (1986) observed the abyssal boundary current along the northern perimeter of the Shikoku Basin, by moored current meters maintained during about five years. Their results strongly suggest that a steady abyssal boundary current exists at depths greater than about 3,000 m along the northern and northwestern perimeters of the Shikoku Basin, suggesting the existence of an abyssal circulation system, a t least, in the Shikoku Basin. With respect t o the inflow of deep water into the Philippine Sea through the Yap and Palau Trenches caved at both sides of the Yap Ridge, Kaneko (1984) analyzed chemical tracer data as well as CTD (Conductivity/Temperature/Depth recorder) d a t a show that chemical properties of the deep water in the Philippine Sea are almost the same as those of the deep water south and southeast of the East Mariana and Yap Ridges. The results show the evidence of the inflow of the North Pacific deep water into the Philippine Sea through the trenches. On the basis of the results obtained in the Philippine Sea, the direct measurements of the abyssal current in the Philippine Sea were carried out in the research project, Priority Area Programme entitled “Dynamics of the Deep Circulation”, 1987-1989 represented by Prof. T. Teramoto. In the programme, two kind of direct measurements of mid and deep ocean currents were carried out in the Philippine Sea; one was the SOFAR float trackings - a Lagrangian measurement of current - by Taira et al. (1990), and the other was the current meter moorings - a n Eulerian measurement of the current - by our group. The objectives of the Eulerian measurement were t o measure the abyssal boundary current along the northwestern perimeter of the Philippine Basin, and to estimate the volume transport of the deep and bottom waters flowing into the Philippine Basin via the West Mariana Basin. For the first objective, three mooring sites, RT, BC and C T were maintained along the northwestern perimeter of the Philippine Basin, as shown in Figure 1. According t o the detailed description of the bottom topography of the northwestern perimeter of the Philippine Basin by Mogi (1979), a trench stretches along the Ryukyu Islands and the Philippine Islands, making the western boundary of the Philippine Sea. The Ryukyu Trench along the Ryukyu Islands is divided into three sections by the seamount south of Miyako Island and the western extension of Amami Plateau. Water depths of the deepest parts are 7,130 m for the southern portion, 7,881 m for the central portion, and 5,998m for the northern portion. The central portion is the deepest

54

fE Fig. 2. Locations of mooring stations at Site RT situated southeast of Okinawa. Bottom topography is shown with contour intervals of 1 km sketched from a chart No.6302 published by the Hydrographic Department, Maritime Safety Agency, Japan (M. S. A., Japan; Approval No. 030065). Areas deeper than 4,000 m are shaded. and the longest. The northern one is surrounded by the Amami Plateau and the Kyushu-Palau Ridge; it is deeper than the Nankai Trough by about 1,000 m. For the second objective, seven mooring lines were deployed on the slopes of the Yap and Palau Trenches. However, almost half of the current meter data were unusable, because of troubles associated with battery supply of d a t a recorders. Additional observation by two mooring lines was carried out at the same location on the slope of the Yap Trench from April 1990 to April 1991. In this paper, the results of the abyssal boundary current. obtained by current meter moorings along the northwestern perimeter of the Philippine Basin are reported. Detailed studies based on the current meter d a t a will be carried out by collaborators in near future. The results for our second objective will be reported elsewhere including those obtained by the additional observation.

55 Table 1. Summary of current meter moorings at sites RT, BC, and CT.

Station

Location

RT1

25'34" 128'09'E 25"29'N 128"12'E 25"24'N 128"18'E

RT2

RT3 BC1 BC2 BC3 BC4 BC5 BC6

CT5 CT6

Water depth

4,570

RT31 RT32

Meter depth (m) 2,130 3,130 2,650 3,650 3,170 4,170

2,500

BCll BC12 BC21 BC22 BC31 BC32 BC41 BC42 BC51 BC52 BC61 BC62

1,100 2,100 2,800 3,800 3,400 4,400 2,280 3,280 3,250 4,250 3,900 4,900

87/03/01-87/11/01 87/03/01-87/09/08 87102128-87/10/29 87102128-87/10/03 87102128-87/10/09 87102128-87/09/21 87/11/01-88/05/24 87/11/01-88/05/24 87111101-88/05/24 87111101-88/05/24 87/11/01-88/05/24 87111101-88/05/24

245 191 242 217 223 205 205 205 205 205 205 205

CT51 CT52 CT61 CT62

2,990 3,990 2,940 3.940

87111110-89/05/13 87/11/10-89/04/28 87111110-89/05/11 87/11/10-89/05/13

550 535 548 550

(4

21"34'N 121'49'E 21"36'N 1 21"59'E 21"40'N 122'07'E 21"36'N 1 21"54% 21'42" 122"04'E 21"42'N 122'09'E 30"07'N 132"2 2'E 30"13'N 132"26'E

Current meter

3,530

RTll RT12

4,050

RT2 1

RT22

4,200 4,800 3,680 4,650 5,300

4,390 4,340

Observational period

Duration (days) 87/11/03-89/05/12 556 258 88108127-89/05/12 229 88108127-89/04/14 235 88108127-89/04/19 552 87111103-89/05/08 540 87/11/03-89/04/26

Observations and Data

2. 2.1.

Mooring stations

Figure 2 shows the locations of the current meter moorings at site RT and a sketch of the bathymetry southeast of Okinawa. The three RT stations were maintained at depths of 3,530 m t o 4,570 m on the northwestern slope of the central portion of the Ryukyu Trench, and were named RT1, RT2, and RT3 from inshore to offshore. Summary of the moorings, including location of station, water depth, current meter depth and mooring period is shown at the top panel in Table 1. The observation was started in November, 1987; continued in August, 1988 after retrieving first mooring, and ended in May, 1989. The total length of observation is about 550 days for all stations except station RT2, where the first mooring was

not retrieved. Figure 3 shows the locations of current meter moorings at site BC and a sketch

56 120'E

125'E

I5'E Fig. 3. Same as Fig. 2 except for Site BC located in the Bashi Channel southeast of Taiwan. of the bathymetry southeast of Taiwan. Summary of the moorings is shown in the middle panel of Table 1. The observation was started in March, 1987 and ended in May, 1988. When the second mooring lines were deployed in November, 1987, the mooring lines were not deployed at the same location as the first mooring stations BC1, BC2 and BC3 because of existence of the strong surface current. As a result, the current observation at site BC was made at six different stations with a mooring period of about 200 days. The first stations BC1, BC2 and BC3 correspond somewhat t o second stations BC4, BC5 and BC6, respectively. Figure 4 shows the locations of the current meter moorings at site C T and a sketch of the bathymetry southeast of Kyushu. Summary of the moorings is shown at the bottom panel in Table 1. The observation was started in November, 1987 and ended in May, 1989. The total length of observation was about 550 days; it is as long as that for site RT also. The moorings for our second objective were located on the slopes of the Yap and Palau Trenches at the eastern and western ends of the Yap Ridge. The site was named Y P after the Yap Ridge.

2.2.

Mooring lines and data sampling

In this observation, most of the mooring lines were composed of two current meters, an acoustic release, glass buoys and a weight as shown in Figure 5 . The

57 total length of a mooring line was about 1,470 m. The current meters were set at about 1,400 m and 400 m above the sea bottom in most mooring lines. Current meters were numbered by increasing depth on the mooring. For example, the current meters at the upper and lower layers at station RT1 were named R T l l and RT12, respectively. The current meter used in this observation were the rotor type Anderaa cufrent meters, named RCM4, RCMS, RCMG, and RCM8. Deployment and retrieval of the mooring lines were made on board T / S Keiten-Maru of the Faculty of Fisheries, Kagoshima University. CTD casts necessary to clarify the density field in the abyssal layer were not carried out, during the current meter mooring. Current meter d a t a were sampled at one hour interval throughout the observational period. The raw data were pre-processed by a 48 hour half power Gaussian filter filtering high frequency components, and the output of this filter was subsampled every one hour to make the data set for the present analysis.

130'E

135'E

Fig. 4. Same as Fig. 2 except for Site C T located a t the northern extremity of the Kyushu-Palau Ridge southeast of Kyushu.

58

4

RADIO BEACON

GLASS BALLS

i

B 1I

b

GLASS BALLS

ROPE GLASS BALLS

i

ACOUSTIC RELEASZ

WEIGHT 525kg

Fig. 5. Sketch of a mooring line.

3. Results 3.1.

Site RT

Site RT is located on the northwestern slope of the central portion of the Ryukyu Trench. The objective of observation is to examine the existence of abyssal boundary currents along the northwestern boundary of the Philippine Basin. The stick diagrams of current are shown in Figure 6; here the current data is subsampled once a day after taking the 72 hours running mean for the d a t a set mentioned in the previous section. The variations of the current for RT31 and RT32 fairly resemble t o each other; both currents flow with almost the same predominant direction. Figure 7 shows the stick diagram of the current for RT32 minus that for RT31, which demonstrates t h a t the current at the lower layer are faster than that

59 meter de th

water deFth

I

t

I I S I I S I IS 1 1 s I I S I 1s I I S I IS I IS O C T NOV OEC JAN F E E f l A R APR H A T JUb

s I I S I I S I IS I I S I IS I I S I I5 I 1 5 I IS IUL AUG S E P O C T NOV O E C J A N FEE M A R APR flAT

I988

Fig. 6. Stick diagrams of the current at stations RT1, RT2, RT3. Each stick shows the speed and direction of the current which is subsampled once a day at 0O:OO time after smoothing by a 72 hours running mean as well as a 48 hour half-power Gaussian low-pass filter.

rl LE P-

I I S I15 I IS I I S I I S I IS I IS I IS I IS I IS I I S I IS I I S I I j I IS I IS I I S I I S I IS I O C T NOV OEC J A N F E E MAR APR f l A 1 JUN JUL AUG S E P O C T NOV DEC JlrN FEE M A R APR f l A T

115

1988

Fig. 7. Stick diagrams of the current for RT32 minus RT31.

at the upper layer, and that the difference is very stable in direction. Such a feature of the abyssal current observed a t station RT3 is almost the same as t h a t observed at the northern perimeter of the Shikoku Basin (Fukasawa et al., 1986). Similar characteristic features of the mean current and variations of current are also found for stations RT1 and RT2, which are located at lesser depths than station RT3, though the former is not so clear as those for station RT3. The offshore extent of the abyssal current at site RT cannot be determined because there are no data concerning the density field in this region as mentioned previously. The basic statistics of current data are shown at the top panel in Table 2;

60 they include speed and direction of mean current, variance of currents (major and minor axes and their orientations) and the ratio of kinetic energy of current variations (ke) to t h a t of the mean current (km). The statistics at station RT3 are especially noticeable among others; namely, the mean currents for RT31 and RT32 have almost the same direction and the current speed for RT32 (400 m above the sea bottom) is faster than t h a t for RT31 (1,400 m above the sea bottom) by about 3 cm/sec. Moreover, the major axes and their orientations of the variance of currents for RT31 and RT32 are almost the same. Those observed results suggest t h a t there exists an abyssal boundary current flowing along the local bottom contour and with the upslope on its right on t h e northwestern slope of the central portion of t,he Ryukyu Trench. The vertical extent of the bottom intensification of the mean current is estimated t o be about 1,800 m above the bottom by the linear extrapolation of the mean current speeds for RT31 and RT32.

3.2.

Site BC

Site BC is located on the slope in the Bashi Channel southeast of Taiwan. The objective of the observation is t o examine the existence of the abyssal boundary current along the western boundary of the Philippine Basin, where the basin has the broadest width in the east-west direction. The stick diagrams of current are shown in Figure 8. Currents for BC31 and BC32 show similar variation during three months from July to September, 1987. The mean current speed for BC32 (400 m above the bottom) was larger than that for BC31 (1,400 m above the bottom) by about 1 cm/sec as shown later, and the current direction for BC31 and BC32 was south-southeast, or parallel t o the local bottom contour. This result suggests t h a t an abyssal boundary current along the local bottom contour existed at t h a t time. T h e vertical extent of this abyssal boundary current at station BC3 is estimated t o be about 1,600 m by linear extrapolation of the d a t a for upper and lower current meters. Similar features are also found for station BC2, though they are clear only for two months from August to September, 1987. At station BC6, the deepest mooring station a t site BC, currents for BC61 and BC62 had almost the same current speed and direction with similar periodic variation. The basic statistics of current d a t a are shown in the middle panel of Table 2. T he variabilities of current are large for site BC compared with site RT. At both stations BC3 and BC6 (offshore stations where the water depth is greater than 4,800 m), the variability is small compared with other stations a t site BC, and the mean current speed for both stations are also weak. Therefore the kinetic energy ratio Ke/Km is large. These statistics suggest t h a t the abyssal currents observed at site BC are unstable, and t h a t such a clear abyssal boundary current as found a t site RT does not exist at site BC. Although some abyssal boundary currents along the local bottom contour seem t o be observed in short period at stations BC2, BC3 and BC6, it is difficult t o decide whether those currents are abyssal boundary currents or not, because those appeared only in shorter period and also because we have no density data.

61 me ler depth water depth

(a)

"'"r

-

U

2.500

m

-

N

m U

N U

mN N

"

-a

m v m C N l

U

a3

IS

I

IS

1

JAN

L

FEE

IS

IS

HAR

APR

I

IS

I5

1

1

JUN

HAT

IS

15

I

JUC

I

AUG

IS

SEP

I

IS

1

OCT

1987

meter depth water depth d

* u m

f.4 v

U m

-

In

m U (Y

In

m U

d

(D

m U UY f.4

m

U

I

15

OCT

I

IS NOV

I

IS OEC

I

IS JAN

I

IS FEE

I

IS HAR

I

15

APR

I

IS

MAT

Fig. 8. Same as Fig. 6 except for (a) stations BC1, BC2, and BC3, and (b) stations BC4, BC5, and BC6.

62 Table 2. Statistics of current meter data at sites RT, BC, and CT. Variability of currents is represented in the form of a variance ellipse; length and orientation of the major axis are tabulated, and also length of the minor axis is shown in parentheses. ke and km are the eddy kinetic energy and mean kinetic energy, respectively. Current meter number

Duration

Mean current speed direction

Variability Axis length direction (cm/sec)2 (deg.) 28 (10.0) 40 22 22 ( 3.0) 29 ( 8.4) 32 34 17 ( 3.9) 32 11 ( 3.5) 51 1 2 ( 4.1)

ke/km

(days) 556 258 229 235 552 540

(cm/sec) 2.0 1.9 2.4 3.1 1.2 4.3

(deg.) 37 200 53 227 247 224

BC12 BC21 BC22 BC31 BC32 BC41 BC42 BC51 BC52 BC61 BC62

245 191 242 217 223 205 205 205 205 205 205 205

2.5 4.9 5.3 1.3 0.5 1.3 5.5 2.3 0.4 0.2 0.9 0.7

189 348 360 142 53 175 177 6 352 143 341 289

125 (20.0) 123 ( 9.8) 214 (21.0) 24 ( 2.1) 19 ( 2.8) 18 ( 2.2) 82 (14.0) 86 ( 4.3) 5.4( 1.6) 30 ( 2.5) 15 ( 2.4) 15 ( 2.5)

1 3 172 166 170 171 10 4 178 168 178 166

23.7 5.5 8.3 14.7 105.0 12.3 3.2 16.7 42.4 611.0 21.9 39.8

CT51 CT52 CT61 CT62

550 535 548 550

0.9 0.9 0.6 2.0

5 234 305 150

11 ( 5.6) 17 ( 4.7) 27 (24.0) 15 ( 7.0)

19 14 65 92

19.1 28.8 153.0 5.2

RTll RT12 RT2 1 RT22

RT31 RT32 BCll

3.3.

10.0 7.2 6.6 2.2 10.3 0.9

Site CT

The site CT is located at the northern extremity of the Kyushu-Palau Ridge, where the junction of the Shikoku Basin and the northern portion of the Ryukyu Trench exists. The junction forms a narrow, deep passage as shown in Figure 4. There are no other junctions between them around this area. The objective of observation at site C T is t o examine the existence of the deep water exchange between the Shikoku and Philippine Basins via the northern portion of the Ryukyu Trench. The stick diagrams of current at site CT are shown in Figure 9. Those for

63 meter depth water depth

c _

lOcm/se

5%

53

"7 T

I IS I15 1 1 5 I IS I IS I IS I I S I I S I I S I IS I IS 1 1 5 1 1 5 I15 I 1 5 I 1 5 I15 I I S 1 IS I IS O C T NOV OEC JAN FEB nAR A P R R A T JUN JUL AUG SEP O C T N O V OEC JAN FEE R A R APR R A T JUN

1988

1989

Fig. 9. Same as Fig. 6 except for stations CT5 and CT6. CT52 and CT62 (400 m above the sea bottom) show t h a t strong currents toward the north or east and those toward the south or southeast appeared intermittently and alternately; namely, there was no stable current with a predominant direction as observed at station RT3. For CT52, very strong currents toward the southsouthwest continued for about 40 days in February and March, 1989. Similar strong current also appeared intermittently in November, 1987, and May through July, 1988. These strong currents contribute t o the mean current as shown later. On the other hand, mean current direction for CT62 is calculated t o be 150"T, which is the direction perpendicular to the axis of the passage. It is because strong currents toward the east and those toward the south-southwest alternately appeared, resulting in cancellation of the component parallel t o the axis of the passage, and leaving the component perpendicular t o the axis of the passage. The variation of current for the upper current meters was nearly the same as that for the lower meters; that is, the strong current also intermittently occurred at CT51 and CT61 throughout the observational period, except for latter half from September, 1988 for station CT6, when current directions for CT61 and CT62 were opposite t o each other. The basic statistics of current d a t a are shown at the bottom panel in Table 2. The mean currents for upper and lower current meters had different current direction, both for stations CT5 and CT6. It is curious because those current meters were moored within the narrow, deep passage. The mean currents for lower current meters (400 m above the bottom) at both stations had different current directions; current direction for CT52 was 234"T, parallel t o the axis of the passage, while that for CT62 was 150"T, perpendicular t o the axis of the passage.

64 These results show that no stable abyssal current existed along the passage between the Nankai Trough and Ryukyu Trench. This conclusion is consistent with that of Fukasawa et al. (1986), who showed that the abyssal current toward the Philippine Basin occurred intermittently at a site (station KPO in their paper) near the present passage in the Shikoku Basin. At the same time, it is suggested that there is a case where the vertical extent of the abyssal current is less than 1,400 m above the sea bottom; for instance, only the current for lower layer was strong at site CT from March t o May, 1989 as shown in Figure 9.

0E '

Fig. 10. Mean current vectors for RT32, BC32, CT52, and CT62. The current meters were moored 400 m above the bottom. The current meter number, the current meter depth, the bottom depth, and ke/km with parentheses are shown for each station. Also shown are results from similar current measurements in the Shikoku Basin by Fukasawa and Teramoto (1986) and Fukasawa et al. (1986). Bottom topography is shown with contour intervals of 2 km sketched from a chart No.6302 published by the Hydrographic Department, M. S. A., Japan.

65 4.

Discussion

It has been suggested by the determination of 14C in the deep and bottom waters that the abyssal circulation in the Philippine Sea has the time scale of several hundreds years (Watanabe, 1989, personal communication). There is always a question on the representativeness in time and space of the observed value by the current measurement as is the case in all the such observations. In this observation, our observed currents are decided t o be the stable abyssal current or not on the basis of the flow stability, the basis which is used by Fukasawa et al. (1986). And it is also assumed that the abyssal boundary current has the same spatial structure both in the Shikoku and Philippine Basins. This assumption seems to be reasonable, because both basins are filled with the deep and bottom waters having the same origin through the Yap and Palau Trenches, and waters below 3,000 m depth in both basins have almost the same water characteristics (Kaneko, 1984). Figure 10 shows the abyssal currents related to the abyssal boundary current observed in this study. Also shown are those observed in the Shikoku Basin in the past several years, including observations at stations CS3, CS4, CS5, CA and KPO, which were made from April, 1981 to November, 1985 (Fukasawa et al., 1968). The strong abyssal boundary currents flowing t o the west and southwest are found at the northern perimeter of the Shikoku Basin. Fukasawa et al. (1986) reported t h a t the mean currents being parallel to the local bottom contour and flowing with the upslope on their right were strongly bottom controlled at depths greater than 3,000 m near the shelf toe and the eastern foot of the Kyushu-Palau Ridge. These mean currents were so stable t h a t the ratio ke/km scarcely exceeded unity. In this observation, the southwestward abyssal boundary current was found along the local bottom contour on the northwestern slope of the central portion of the Ryukyu Trench; its mean current speed was 4.3 cm/sec and the ratio ke/km was 0.9. This abyssal current is concluded to be stable considering the value of ke/km, which is comparable t o those obtained in the Shikoku Basin by Fukasawa et al. (1986). At the toe of the western slope of the Philippine Basin in the Bashi Channel, the mean current toward the south was weak and the value of ke/km was large, and hence the abyssal current is concluded t o be unstable. The deep and bottom waters in the Shikoku and Philippine Basins seem to be exchanged through the narrow, deep passage at the northern extremity of the Kyushu-Palau Ridge, but, no stable abyssal current was observed in the passage. The deep and bottom waters in both basins seem to be exchanged by the strong but intermittent alternating current. The flow pattern of the observed abyssal boundary current is approximately reproduced in a numerical model by Kubota and Ono (1991), but the estimated current speeds in the numerical model do not agree with the observed magnitudes. Another discrepancy is the abyssal boundary current in the passage at site C T mentioned above; the numerical model shows a considerable stable flow there. We do not understand the reason for the difference, but it might be related with the surface current condition. At depth of 1,500 m near site C T , the current condition was sluggish with clockwise turnings several times in an area of 150 km in the northeast direction and 100 km in the east-west direction according t o a trajectory of

66 a SOFAR float for about 200 days after deployed in November, 1988 (Taira et al.,

1990). Further study is required on the current condition from surface t o abyssal depth at the western part of the Shikoku Basin.

5.

Conclusions and Remarks

We made the abyssal current measurements using current meter moorings at several stations in the Philippine Sea to describe features of the abyssal boundary current and to estimate the volume transport of the deep water flowing into the Philippine Basin. In this paper, we report features of the abyssal current along the northwestern perimeter of the Philippine Basin and at the junction between the Shikoku and Philippine Basins. Current meters were set at depths greater than 2,500 m on the northwestern slope of the central portion of the Ryukyu Trench southeast of Okinawa and a t the toe of the western slope of the Philippine Basin southeast of Taiwan. The obtained data at depths greater than 4,000 m indicated the existence of the bottom intensified mean current which are parallel to the bottom contour and flowing with the upslope on their right, similar to currents at the northern perimeter of the Shikoku Basin as described by Fukasawa et al. (1986). The existence of an abyssal boundary current is strongly suggested on the northwestern slope of the central portion of the Ryukyu Trench. The vertical extent of the bottom intensification of the mean current is estimated to be about 1,800 m above the bottom on it. In the narrow, deep passage between the Shikoku and Philippine Basins located at the northern extremity of the Kyushu-Palau Ridge, current meter data indicate no stable abyssal current with a predominant direction. It might be suggested that the deep water transport through the passage between the basins occurs intermittently by strong alternating currents. As mentioned in Introduction, one of our important purposes was not achieved; namely, we were not able to estimate the volume transport of deep water flowing into the Philippine Basin, because data analysis was delayed including data from additional observation. Estimating this volume transport is crucial to build up the abyssal circulation model and to discuss the distribution of deep water properties in the basins of the Philippine Sea. The present observation suggests that the abyssal circulation in the Philippine Basin is more or less separated from that in the Shikoku Basin. Two current observations mentioned below should be made. One is the observation on the southern slope of the Oki-Dait,o Ridge, which is located at t,he northern boundary of the Philippine Basin as shown in Figure 1. The other is the observation at the passage between the Kyushu-Palau and Daito Ridges. Through this passage, the deep water may be supplied from the Philippine Basin into the Kita-Daito Basin and into the Amami Basin north of the Amami Plateau. According t o the numerical model of the Philippine Sea by Kubota and Ono (1991), another abyssal circulation exists in the southern part of the Philippine Basin. Therefore the observation at the western slope of the Philippine Trench is also useful for clarifying the abyssal circulation in the Philippine Sea.

67 Acknowledgement We would like t o express our gratitude t o Professor T. Teramoto of Kanagawa University, Program Scientist of this research program for his valuable suggestions and encouragement throughout the present study. We wish t o express our sincere thanks t o Professor K. Taira of Ocean Research Institute, University of Tokyo for his useful advice and help throughout this study. We are very grateful t o the captain and the crew of T/S Keiten-Maru, Faculty of Fisheries, Kagoshima University, for their sincere help. We thank Messrs. S. Kitagawa of Ocean Research Institute, University of Tokyo, N. Kita of Research Institute for Applied Mechanics, Kyushu University, and T. Yamashiro of Faculty of Engineering, Kagoshima University for their assistance on board. We also thank Drs. S. Imawaki and H. Ichikawa of Faculty of Fisheries, Kagoshima University for improving the manuscript and their help throughout this study.

References Fukasawa, M., and T. Teramoto, 1986. Characteristics of deep current of Cape Shionomisaki before and after formation of the large meander of the Kuroshio in 1981. J. Oceanogr. SOC.Japan, 42, 53-68. Fukasawa, M., T. Teramoto, and K. Taira, 1986. Abyssal current along the northern periphery of Shikoku Basin. J. Oceanogr. SOC.Japan, 42, 459-472. Kaneko, I., 1984. Structure of Mid-depth Waters in the Philippine Sea. Doctor’s thesis, University of Tokyo, 92 pp. Kubota, M., and K. Ono, 1991. Steady abyssal circulation in the Philippine Sea (submitted for publication). Mogi, A., 1972. Bathymetry of the Kuroshio Region. In: Kuroshio - Its Physical Aspects, H. Stommel and K. Yoshida (ed.), University of Tokyo Press, Tokyo, 517 pp. Mogi, A., 1979. An Atlas of the Sea Floor around Japan, Aspect of Submarine Geomorphology. University of Tokyo Press, Tokyo, 92 pp. Stommel, H., and A. B. Arons, 1960a. On the abyssal circulation of the world ocean I. Stationary planetary flow pattern on a sphere. Deep Sea Res., 6, 140-154. Stommel, H., and A. B. Arons, 1960b. On the abyssal circulation of the world ocean - 11. An idealized model of the circulation pattern and amplitude in oceanic basins. Deep Sea Res., 6, 217-233. Taira, K., S. Kitagawa, K. Uehara, H. Ichikawa, H. Hachiya, and T. Teramoto, 1990. Direct measurements of mid-depth circulation in the Shikoku Basin by tracking SOFAR floats. J. Oceanogr. SOC.Japan, 46, 296-306.

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69

Deep Circulation in the Shikoku Basin Measured with the SOFAR Floats Keisuke Taira and Daigo Yanagimoto*

Abstract

Tracking of the SOFAR floats were made from April 1988 to January 1991 by releasing twelve floats at 1000 m, 1500 m, 2000 m, 3000 m and 4000 m in the Shikoku Basin. The receivers were moored in five periods and trajectories were obtained for nine floats. Deep circulations and eddy diffusivity inferred from the trajectories of the floats are described.

1.

Introduction

The Shikoku Basin is located south t o Japan (Fig. 1). The Kuroshio, the western boundary current of the North Pacific Ocean, is flowing in the upper layer of the basin entering through Tokara Strait and going out over Izu Ridge. Although the water depth of Tokara Straight and Izu Ridge is hundreds meters, the Kuroshio has a deep current structure in the Shikoku Basin (e.g., Worthington and Kawai, 1972; Taft, 1978). The circulation at mid-depth in the Shikoku Basin is much different from that at the surface measured with the GEK and the satellite-tracking of surface drifters. The direct current measurements in the recent years suggest a deep counter clockwise circulation in the Shikoku Basin. Taira and Teramoto (1981) showed a northward mean flow of 3 cm s-l at 1670 m depth under the northward flowing Kuroshio west to Hachijojima. Nishida and Kuramoto (1982) showed a weak northward deep flow along the western flank of the Izu Ridge. Taira and Teramoto (1985) showed that a north flow at near bottom layer turned t o the southwest along an isobath of Nankai Trough. Their measurements were made along the northeastern perimeter of the Shikoku Basin. A westward flow was observed under the eastward Kuroshio off Omaezaki (Ishizaki et al., 1983) and off Shionomisaki (Fukasawa et al., 1986). Along the western perimeter of the Shikoku Basin, a southward deep flow was observed in the east t o Tanegashima, Okinawa and Taiwan (Chaen et al., personal communication). Circulations in the Shikoku Basin are subject t o cold and warm water masses. South t o the Kuroshio path, a warm water mass is located off Shikoku. The center of the warm water mass shifts westward by about 300 km during the large meandering path of the Kuroshio appears. A large cold water mass off Kii Peninsula appears to the north of the Kuroshio when the Kuroshio path meanders. *Ocean Research Institute, University of Tokyo,

Minamidai, Nakano-ku, Tokyo 164, Japan

70

34

32

30

28

130

132

134

136

138

140

Fig. 1. Bathymetry of the Shikoku Basin, mooring stations of SOFAR receivers ( R l , R2 and R3). Current records at limited numbers of moored stations are subject to the bottom control and to the flows associated to the water masses. A Lagrangian measurement of current, or tracking of water parcel a t the mid-depth, is most requisite to elucidate the deep circulations in the Shikoku Basin. Tracking of neutral floats over several hundreds km distance for several hundreds days was planned as one of the core subjects of the Priority Area Programme “Dynamics of the Deep Circulation, 1987-1989” represented by Prof. T . Teramoto. Twelve SOFAR floats were released in the Shikoku Basin at the depth of 1000 m (2 floats), 1500 m (4 floats), 2000 m (2 floats), 3000 m (2 floats) and 4000 m (2 floats). The tracking was made from April 1988 to January 1991. The circulation inferred by the float trajectories and Lagrangian flow characteristics are described. A part of this study may be found in Taira et al. (1990) and in Yanagimoto (1992).

2.

Tracking Observation and Data Analysis

We used a SOFAR float with capability depth of 6500 m, manufactured by Webb Research Corporation. Two glass balls are used both for buoyancy and for pressure housing. One glass ball contains control unit, and the other battery unit. Outer diameter of the glass ball is 43.2 cm. The float transmits sound waves of

71 Table 1. Float number, drift depth and position, water depth, and date of the deployment. The float number shows the time of the first emission, i.e., FOO emitts signals at OOh and 12h on each day. No. FOO F01 F02 F03 F04

F05 F06 F07 F08 F09 F10 F11

Depth 1500 m 1000 m 1000 m 1500 m 2000 m 2000 m 1500 m 3000 m 3000 m 1500 m 4000 m 4000 m

Latitude 30’00’N 30’00’N 32’51” 29’52” 30’00’N 32’51” 30’00’N 30’00’N 32’50” 29’52” 30’00’N 32’50”

Longitude 135’59’E 134’59’E 136’31’E 135’25’E 135’00’E 136’31’E 136’00’E 134’59’E 136’32’E 135’25’E 135’00’E 136’33’E

w.d. 4410 m 4628 m 4161 m 3521 m 4623 m 4486 m 4390 m 4626 m 4498 m 3530 m 4618 m 4492 m

Deploy date 17 April 1988 27 June 1989 22 June 1989 3 Nov. 1988 27 June 1989 22 June 1989 17 April 1988 27 June 1989 22 June 1989 3 Nov. 1988 27 June 1989 22 June 1989

780 Hz for 80 seconds two times every day for tracking. The sound source is a piezo-electric, bending element with a resonator of aluminum tube 60 cm long and 30 cm in diameter. The 0oat is t o be deployed at drifting depth after buoyancy balance in the laboratory. Buoyancy balance of the SOFAR floats was made in a cylindrical tank 60 cm in the inner diameter and 170 cm long. The tank was filled with seawater. The procedure of balance is described in Taira et al. (1990). In the deployment from a ship, the SOFAR float was hung down to the sea surface by rope with a wedge stopper. The wedge was pulled away after the resonator tube was filled with seawater. The floats released from the R/V Hakuho Maru in April 1988, from the T / S keiten Maru in November 1988, and from the R / V Tansei Maru in June 1989. The release depth, position, and date are tabulated in Table 1. The float is named in terms of signal emission time. For example, float F01 transmits at 0100 and 1300 on each day. A SOFAR receiver is composed with a hydrophone array of eight elements, a signal processor and a digital cassette recorder. The hydrophone array of 3.5 m overall length is tuned for sound waves from horizontal direction and it is set in the SOFAR channel, i.e., at the sound speed minimum layer of about 1000 m depth in the Shikoku Basin. A wave train of 780 Hz with an unit amplitude is generated in the signal processor and it is multiplied with the received signal which is clipped t o have an unit amplitude. Output signal of the multiplier has a maximum amplitude when the 80 second signal from a SOFAR float is received. The largest four amplitudes within every 10 minutes and their arrival time are recorded on a cassette tape. The SOFAR receivers were moored at three stations in the Shikoku Basin with a separation of about 250 km (Fig. 1). The first sets were moored from April 1988

72 Table 2. Mooring data of the SOFAR receivers. %.No. R1-1 R2-1 R3- 1 R1-2

R2-2 R3-2 R1-3 R2-3 R3-3 R1-4 R2-4 R3-4 R1-5 R2-5 R3-5

Latitude 32’30” 28’59” 28’59” 32’30” 29’01” 29’0 1’N 32’30” 29’01” 29’01” 32’31’N 29’00” 29”02” 32’28” 28’58” 29’00”

Longitude 136’02’E 133’30’E 136’00’E 136’02’E 133’30’E 136’00’E 136’01’E 133’30’E 136’00’E 136’02’E 133’30’E 136’05’E 136’03’E 133’30’E 136’00’E

Depth 4640 m 3380 m 4430 m 4600 m 3340 m 4250 m 4600 m 3340 m 4250 m 4631 m 3369 m 4495 m 4582 m 3363 m 4417 m

Set date 16 April 1988 18 April 1988 17 April 1988 27 Oct. 1988 1 Nov. 1988 2 Nov. 1988 27 Oct. 1988 1 Nov. 1988 2 Nov. 1988 23 June 1989 11 Aug. 1989 27 June 1989 14 April 1990 19 April 1990 20 April 1990

Recovery date 27 Oct. 1988 1 Nov. 1988 2 Nov. 1988 27 June 1989 11 Aug. 1988 27 June 1989 23 June 1989 11 Aug. 1989 27 June 1989 14 April 1990 19 April 1990 20 April 1990 1 2 Jan. 1991 23 Jan. 1991 14 Jan. 1991

t o October 1988, the second from October 1988 to June 1989, the third from June 1989 t o November 1989, the fourth from November 1989 t o April 1990, and the fifth from April 1990 t o January 1991 (see Table 2). Records from the receivers were not always of high quality. Some of them were not obtained due t o malfunction of the recording unit or due to the water leakage of the hydrophone array. Fortunately, at least two records were available for each mooring period. The float positions were determined from the arrival time at two stations by the method described in Taira et al. (1990) for the period from 20 April 1988 t o 27 June 1989, and for the period from 19 June 1989 t o 12 January 1991 by the method described by Yanagimoto (1992). In the former method a Cartesian coordinate is assumed, and in the latter a spherical coordinate is used. The battery life of the receivers was about 240 days. A complete data set for 440 days from 1 2 April 1988 to 20 June 1989 was obtained for the listening stations at R1 and R2. There were data gaps between the mooring periods except for the period.

3. Trajectories of SOFAR Floats Drifting at 1500 m Depth Four floats were deployed at 1500 m in Shikoku Basin, two on 17 April 1988 at 30’N and 135”59’E,two on 3 November 1988 at 29’52” and 133’25’E (see Table 1). Fig. 2 shows the trajectories for 440 days from 12 April 1988 t o 20 June 1989. Circles on the trajectories show every 50 days after deployment. The trajectories of the floats showed flows at 1500 m depth in the Shikoku Basin classified into three: 1) clockwise circulation in the central part, 2) northward and southward flows along the western perimeter of the basin, and 3) sluggish flow in the western part of the basin. Floats F03 and F06 were tracked in the period from 19 June

73

28 130

-*

132

I

134

-

I

I

136

138

140

Fig. 2. Trajectories of the SOFAR floats at 1500 m from 20 April 1988 to 27 June 1989 for floats FOO and F06, and from 6 November 1988 t o 27 June 1989 for floats F03 and F09. 1989 t o 12 January 1991, and these trajectories are described in Section 4).

Clockwise circulation in the Shikoku Basin Fig. 3 shows trajectories of floats FOO and F09. Float FOO drifted eastward for several days after deployment, but it turned t o south and then t o t h e west. It 3.1.

turned to the north on 220 day after deployment. From 250 day t o 350 day, it made a cyclonic turn around at 31'30" and 133'30'E. After the turn, it flowed northeastward. Float F09 was deployed 200 days after float FOO, and it flowed northeastward after a cyclonic turning with about 10 km diameter at around 30'30" and 133'30'E. The float flowed southeastward, and it showed quick cyclonic turnings of about ten around at 30"20'N and 138"30'E. Two trajectories showed a clockwise circulation in the Shikoku Basin. A diameter of the circulation is estimated t o be about 500 km. A period of 400 days for one turn gives a mean speed of 4.5 cm s-l. Drift of F09 from 50 day t o 100 day gave a mean speed of 11 cm s-'. The trajectory of float F09 shows quick cyclonic turnings around 30"20'N and 138'30'E. The area is located on the western slope of the Izu Ridge where water depth is beween 1500 m and 3000 m. T h e elliptic trajecories are 80 km long in the northwest and 25 km in the northeast. T h e velocity amplitudes are about 80 cm s-' for the north component and 40 cm s-l for the east component, and the period is about 8 days.

74

PI

-

132

I

.

134

I

1

I

136

138

140

Fig. 3. Trajectories of the SOFAR floats at 1500 m from 20 April 1988 to 27 June 1989 for float FOO, and from 6 November 1988 t o 27 June 1989 for float F09.

/

28 130

Pl

-

132

I

134

-

I

I

136

138

140

Fig. 4. Trajectories of the SOFAR float F06 at 1500 m from 20 April 1988 t o 27 June 1989. 3.2.

Northward and southward flows along the western perimeter

Fig. 4 shows trajectory of float F06, which was deployed on the same day and at nearly the same position as float FOO. Two trajectories were identical for several

75

28 130

PI

I

-

132

F03 I

134

-

I

I

136

138

140

Fig. 5. Trajectories of the SOFAR float F03 at 1500 m from 6 November 1988 t o 27 June 1989. days, but float F06 flowed t o the west until 150 day. It flowed t o the north until 200 day, and then it flowed southward along an isobath of 1500 m. From 250 day to 380 day, it flowed into the Shikoku Basin with a clockwise turning. It flowed again southward until the end of tracking. The speed of the southward flow was about 4 cm s-'. The trajectories estimated t o be lying in the western area where water depth smaller than 1500 m. This was considered t o be caused by a delay of the clock of the float by about 20 seconds.

3.3.

Sluggish flow in the western part of the basin

Fig. 5 shows trajectory of float F03 which was deployed on 3 November 1988 at 30°N and 133'25'E. It stayed in a square of 100 km in the north direction and 150 km in the east direction. A mean drift speed was less than 2 cni s-'. The float made clockwise turnings of several times. 3.4.

Trajectories of the floats at 1500 m from 19 June 1989 t o 12 January 1991

Fig. 6 shows trajectories of floats F03 and F06 in the period from 19 June 1989 to 12 January 1991. Float F06 appeared to the east of Kyushu and entered again into the deep Shikoku Basin. This shows that the southward flow under the Kuroshio is not stable. When it crossed the bottom contour, clockwise turnings with a period of 4 days were observed. The float went northeastward along the isobath south to Shikoku. The float appeared east to the Kyushuu-Palau Ridge in the beginning of the fifth period, and it flowed northwestward. Float F03 moved northward, and it returned near t o the deploy position after counter clockwise turning.

76

28 130

132 134 136 138 140 Fig. 6. Trajectories of the SOFAR floats F03 and F06 at 1500 m in the period from 19 June 1989 to 1 2 January 1991.

34

32

30

28 130

132

134

136

138

140

Fig. 7. Trajectories of the SOFAR floats F01 and F02 at 1000 m in the period from 19 June 1989 to 1 2 January 1991.

77

28" 130

U

d

132

I

134

*

I

I

136

138

140

Fig. 8. Trajectories of the SOFAR floats F04 and F05 at 2000 m in the period from 19 June 1989 to 1 2 January 1991.

period

4.

Trajectories of the SOFAR Floats at 1000 m, 2000 m and 3000 m Depth

Eight floats were released a t 1000 m, 2000 m, 3000 m and 4000 m in June 1989 (See Table 1). The signals from float F08 deployed at 3000 m, floats F10 and F11

78 Table 3. Eddy viscosity and relaxation time estimated from the SOFAR float tracking. The values for the western North Atlatic Ocean are given by Riser and Rossby (1983). Sea area

depth (m)

Shikoku Basin

1000 1500 2000 3000 700 2000

Western North Atlantic Ocean

Eddy diffusivity ( 106cm2s-1) E-W N-S 25 13 11 9.9 6.6 7.7 4.9 4.5 45 18 9.0 2.6

Relaxation (day) E-W 8.1 6.9 14 7.6 18 15

time

N-S 7.1 6.2 16

17 7.0 7.0

deployed at 4000 m depth were sporadically received by one of three receivers. T h e trajectories of these floats were not obtained. Fig. 7 shows trajectories of floats F01 and F01 deployed at 1000 m. Float F01 deployed off Kii Peninsula first flowed southward, and it showed a clockwise turning, then a counter clockwise turning, and a clockwise turning. T h e drifting speed was about 8 cm s-'. Fig. 8 shows trajectories of floats F04 and F05 deployed at 2000 m depth. Float F04 made clockwise turnings. Float F05 flowed southward, and then it turned to northeastward. A mean drifting speed was about 3.2 cm s-'. Fig. 9 shows the trajectory of F07 at 3000 m. Mean drift speed was about 3.3 cm s-'. The flow is similar to t h a t at 2000 m depicted by float F04.

5.

Eddy Diffusivity Estimated from Lagrangian Velocities

Tracking of a SOFAR float is by quasi-Lagrangian measurements of water movement on an isobaric surface since the float is equipped with buoyancy-adjusting gear t o keep the pressure constant. T h e correlation function of Lagrangian velocity gives a n eddy diffusivity and a relaxation time (e.g., Riser and Rossby, 1982). The eddy diffusivity k is estimated by

1 t

k=

R(r)dr

(1)

where R ( r ) is the auto-correlation function of Lagrangian velocity v l , i.e., R(7)= vl(t)vl(t - 7). Th e relaxation time I represents a time scale of velocity correlation, and is estimated by rt

The estimated eddy diffusivity and relaxation time are shown in Table 3. The eddy diffusivity is larger at 1000 m and it decreases with the depth. The eddy

79 diffusivity for the east component is two times as large as that for the north component. It is not clear why the difference comes from because the drift speeds and their root-mean-squares are of the same magnitude for both components. At the deeper layers, eddy diffusivity is similar for the east and north components. The relaxation time varies from 4.5 day to 17 day, and it has a tendency t o increase with the depth. Riser and Rossby (1982) estimated the eddy diffusivity and the relaxation time from the trajectories of the SOFAR floats at 700 m and 2000 m in the recirculation of the Gulf Stream in the western North Atlantic. Their results are also shown in Table 3. The eddy diffusivity in the North Atlantic is of the nearly same magnitude as in the Shikoku Basin. The eddy diffusivity for the east component is nearly 3 times as large as that for the north component. The relaxation time is from 7 day t o 18 day, which is similar t o that observed in the Shikoku Basin.

6.

Summary

Twelve SOFAR floats were released in the Shikoku Basin at the depth of 1000 m (2 floats), 1500 m (4 floats), 2000 m (2 floats), 3000 m (2 floats) and 4000 m (2 floats). The receivers were moored at three stations in the five periods from April 1988 to January 1991. The signals from the floats a t 4000 m and from one at 3000 m were not received at more than two stations. The trajectories were obtained for nine floats. Circulation at 1500 m depth measured from April 1988 t o June 1989 may be summarized as: 1) a clockwise circulation around the Shikoku Warm Water Mass centered around at 31"N and 136"E. Two floats made a clockwise turning around the water mass with a diameter of about 400 km at a mean drifting speed of 4.5 cm s-'. 2) a southward flow under the northward flowing Kuroshio in the east t o Kyushu. One float revealed the southward flow of 4 cm s-l speed. The float first flowed northward, and it went out t o central portion of Shikoku Basin on a midway of southward drift. Although the southward flow was not steady, it was suggested to be one part of a anti-clockwise circulation at deep layers along the perimeter of the basin, which had been shown by the records of moored current meters. 3) a sluggish flow in the western part of the basin. One float deployed at 29"52'N and 133"25'E remained in a limited area about 100 km in the east direction and 150 km in the north direction for 234 days. However, the trajectory of the float crossed several times those of the floats drifting around the Shikoku Warm Water Mass and of the float flowing under the Kuroshio. Some of the above flow characteristics were confirmed at other depth and in the remaining periods. The Kuroshio path was rather unstable in the latter half of the tracking observations, and a further analysis on the distribution of the water masses is requisite. The trajectories of the floats revealed several cyclonic and anticyclonic eddies in the Shikoku Basin. Among them, the cold eddies at 1500 m depth in the west to the Izu Ridge were the most remarking. The shape was elliptic about 80 km long in the northern-northwest and 25 km long in the western-southwest, and the period was 8 days. The cyclonic turns were observed from March t o June 1988.

T h e eddy diffusivity estimated from t h e drift speeds of t h e floats varies from

4 . 5 - 2 5 ~ 1 0cm2sP1, ~ a n d t h e relaxation time is 5.4-17 days.

References Fukasawa, M., T . Teramoto and K.Taira, 1986. Abyssal currents along the periphery of Shikoku Basin. J. Oceanogr. SOC.Japan, 42, 459-472. Ishizaki, H., 0 . Asaoka, S. Konaga, and M. Takahashi, 1983. Direct measurement of near-bottom current on the continental slope off Omaezaki, central Japan. Pap. Met. Geophys., 33, 257-268. Nishida, H., and S. Kuramoto, 1982. Deep current of the Kuroshio around the Izu Ridge - Large meander of the Kuroshio in 1975-1980 (IV). Rept. Hydrogr. Res., Tokyo, N0.17, 241-255. Riser, S. C., and H. T. Rossby, 1982. Quasi-Lagrangian structure and variability of the subtropical western North Atlantic circulation. J. Mar. Res., 41, 127-162. Taft, B. A, 1978. Structure of the Kuroshio south of Japan. J. Mar. Res., 16, 77-117. Taira, K., and T. Teramoto, 1981. Velocity fluctuation of the Kuroshio near the Izu Ridge and their relationship to current path. Deep-sea Res., 28, 1187-1197. Taira, K., and T . Teramoto, 1985. Bottom currents in Nankai Trough and Sagami Trough. J. Oceanogr. SOC.Japan, 41, 388-398. Taira, K., S. Kitagawa, K. Uehara, H. Ichikawa, H. Hachiya and T . Teramoto, 1990. Direct measurements of mid-depth circulation in the Shikoku Basin by tracking SOFAR floats. J. Oceanogr. SOC.Japan, 46, 296-306. Worthington, L. V., and H. Kawai, 1972. Comparison between deep sections across the Kuroshio and the Florida Current of Gulf Stream. In: Kuroshio, Its Physical Aspects, H. Stommel and K. Yoshida (ed.), University of Tokyo Press, pp. 371-385. Yanagimoto, D. (1992). Observation of mid-depth circulation in the Shikoku Basin by tracking the SOFAR floats. Master Thesis, University of Tokyo, 48 pp.

Chapter 3 Circulation and Mixing of Water as Deduced from Distribution of Chemical Tracers

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

83

Chemistry and the Oceans: An Overview Yoshiyu ki N 0Z AKI*

Abstract

Chemical oceanographers are concerned with the composition and distribution of elements and isotopes in ocean water. Their determination is a basic study of marine science which not only satisfies human’s curiosity but also can help us to understand various oceanographic processes. A focus is placed on global change.

1.

Introduction

In the last few decades of 20th century, it has become increasingly convincing that the ocean plays significant role on our global environment. This trend has led us to currently reinforce a need of comprehensive study of the world oceans through international cooperation. Along the line, two major international oceanographic programs are now going on, which are the World Ocean Circulation Experiments (WOCE) and the Joint Global Ocean Flux Study (JGOFS). WOCE deals with physical ocean circulation and JGOFS deals with the biogeochemical cycling in the ocean. Both of which are very important, because the ocean acts as a significant modifier of the world climate through exchange of heat, water and COz across the air-sea interface. There is no doubt that without having detailed knowledge of the oceanic behavior, it is impossible to precisely predict the future trend of global climatic change, which may be caused by human use of fossil-fuels and reduction of terrestrial biosphere. Chemical oceanographers are concerned with the composition and distribution of elements in seawater, and eight papers contained in this chapter cover a wide range of interests from analytical method t o tracer modelling. Nevertheless, it seems clear t o me that all these efforts are directed toward improving our knowledge of the global system. To demonstrate this, I briefly review the history of chemical analysis of seawater. Present status and some future directions of chemical oceanography are also described.

2.

A Historical Review

The composition of major elements was established by the end of 19th century, when the British Challenger expedition was carried out. Since then, significant efforts of marine chemists have been devoted t o determine trace elements in seawater. However, this was not an easy task for many of the heavy metals because their * O c e a n Research I n s t i t u t e , University of Tokyo, Minamidai, Nakano-ku, T o k y o 164, J a p a n

84 Table 1. Some heavy metal concentrations in seawater (nmol/kg) adopted in various review articles

Element Fe cu Ag Au Pb Bi

Goldberg (1963) 180 50 3 0.02 0.2 0.1

Brewer (1975) 36 8 0.3 0.02 0.2 0.1

Quinby-Hunt & Turekian (1983) 0.7 2 0.03 0.06 0.005 0.05

Nozaki (1992) 0.6 2 0.02 0.00015 0.005 0.00015

concentrations are extremely low, and the water samples are very easily contaminated. It is a famous story t h a t a German chemist, Fritz Haber tried to recover gold from seawater in order for his mother country t o be able t o pay the deficit t h a t resulted from the World War I. His attempt completely failed, however, because the concentration of gold in seawater was at least three orders of magnitude lower than the value given a t the time when the program was started, as he found after his ten-year elaborated work. There are a number of other stories showing t h a t the values reported for heavy metal concentration in seawater decreased with time and some examples for this trend are given in Table 1. Of course, these are not real but artifact. This tendency continued even when the last major oceanographic program, the Geochemical Ocean Section Study (GEOSECS) was initiated in around 1970. The GEOSECS was aimed at obtaining the section map of various geochemical parameters in the world oceans, but it could not include trace elements except for Ba because of the lack of analytical consistency, and therefore its focus went on radionuclides and stable isotopes. It was since 1976 t h a t reliable d a t a for some heavy metals such as Cu, Ni and Cd became available owing t o the careful improvement of methods. In the following 15 years to date, there has been an accelerated quantum leap in our knowledge of trace elements in seawater. Figure 1 shows the most recent d a t a set for trace element distribution in the North Pacific given in the form of periodic table. If this were made 5 t o 10 years ago, perhaps one third to a half could be incorrect or unfilled. Even now, there is still lack of d a t a for some platinum metals, R h and Os, and Nb and Ta etc., but there is a good hope that all the d a t a will be filled by the year 2000 so far as the current efforts are continued. Although determining trace element and isotopic composition of seawater is a very basic study of marine geochemistry, it should be realized t h a t it is not only for satisfaction of human’s curiosity but also very important in solving various oceanographic problems related to global change.

3.

Where We Stand Now

All agree that the ocean plays a significant role on the global environment through exchange of heat and materials with other reservoirs across the boundaries and its internal physical and biogeochemical cycling. However, the problem is t h a t

Trace elememt distributions - 1

c

m

a,

0

0

.-0 c .0

m n

a,

5 .-c

in the Pacific Ocean

- I

......

...

!U. ;17 .......

............

85

: ........r \ ir\ ..

Fig. 1 A part of periodic table showing the vertical profiles of elements in the North Pacific Ocean.

86 we do not understood the system well enough to precisely predict the future trend of global change. Toward resolving the problem, chemistry of the ocean is very important in the following three aspects.

3.1.

Their usefulness as tracers of ocean circulation

Various chemical tracers of steady-state and transient nature such as 14C, 3HHe, Freons etc. are being measured in the WOCE program following the tradition of GEOSECS and the Transient Tracer in the Ocean (TTO). The tracers listed in Table 2 are considered to be useful to estimate the rates of basin scale circulation, ventilation, and thermocline mixing, which are somewhat difficult to obtain by other means. One of the major advances since GEOSECS is that analytical techniques have been considerably improved. For example, tandem accelerator mass spectrometry is now successfully applied t o the determination of 14C and “Be and revolutionarily reduced the amount of water needed for measurement from 200 L to less than 1 L for 14C and from thousands to a few tens of liters for 1°Be. A new high-sensitivity gas chromatographic method made it possible to use anthropogenic chlorofluorocarbons as novel tracers of physical circulation, in particular for the study of thermocline ventilation. 3 H and 3He are now routinely measured by mass spectrometry rather than the conventional gas counting method employed during GEOSECS. Further improvement of analytical techniques of potentially valuable tracers seems likely. They include a resonance ionization mass spectrometry for 39Ar, a continuous flow a-scintillation detection for 2 2 7 Aand ~ a high abundance sensitivity mass spectrometry for Pu. Trace elements can also serve as water mass tracers provided that convenient and precise methods are available. For example, Mn and Al, and CH4 are useful to survey the deep-sea hydrothermal vents which are hardly traced by regular properties like salinity and nutrients. A1 is also useful to follow the fresh water discharge from rivers and the Mediterranean outflow in the North Atlantic (Measures et al., 1986). The ocean circulation is undoubtedly an important control of global climate. Extensive upwelling of nutrient-rich deep water to the surface, if it occurs more than usual, enhances the ocean productivity and thereby reduces atmospheric COZ. It also carries cold water up to the surface so that the surface temperature may be reduced. Thus, any change of ocean circulation can influence the global climate. Chemical tracers provide one of the useful tools t o study such physical oceanographic behavior. 3.2.

Their importance in biological processes

It has been well known that the micro-nutrients such as N, P, K, Fe etc. are essential for plant growth. For terrestrial plants, it is hard to believe that Fe can act as a limiting nutrient because it is very abundant in soil. However, in the ocean, Fe is highly reactive and has a short residence time. And therefore, it is likely that the lack of Fe (which is necessary to form enzyme “cytochrome”) can limit phytoplankton growth. Martin and Fitzwater (1988) collected seawater using a clean technique, and made the culture experiments with and without addition of Fe. They observed enhanced plant growth with Fe, suggesting that Fe indeed acts as a limiting nutrient. Martin et al. (1989) also made a careful seawater measurements

Table 2.

Transient and radioactive tracers useful for the study of ocean circulation

Tracer

Species

1) Chlorofluorocarbons 2) Tritium 3) Helium-3 (a) decay product (b) primordial 4) Radiocarbon (a) man-made (b) natural 5) Radium-226 6) Argon-39 7) Silicon-32 8) Actinium-227 9) Radium-228

Diss. gas unreactive HTO unreactive

10) Kripton-85 11) Strontium-90

12) Cesium-137

Reactivity

Time Range

< -3Oy < -3Oy

Source Function Atmosphere (gas exchange) Atmosphere (fallout)

Diss. gas unreactive one month-30 y Decay of tritium Diss. gas unreactive Earth’s interior reactive HCO, reactive HCO, reactive Ra2+ Diss. gas unreactive reactive H2Si03 reactive A&O; slightly Ra2+ reactive Diss. gas u n re act i ve Sr2+ slightly reactive reactive cs+ slightly reactive

< -30 y 20-50,000 y 50-5,000 y 20-1,000 y 30-700 y 5-100 y 0.5-30 y

Atmosphere Atmosphere Sediment Atmosphere Atmosphere Sediment Sediment

(gas exchange) (gas exchange)

Water Volume Needed for Detection

< l L < l L

< 0.1 L < 0.1 L 1L

(gas exchange) (fallout)

200 L 20 L 1500 L 1500 L 200 L 200 L

< 30 y < 30 y

Atmosphere (gas exchange) Atmosphere (fallout )

200 L 50 L

< 30 y

Atmosphere (fallout)

50 L 00

88 and found that the surface waters in the higher latitude regions are depleted in Fe, although some N and P are still remained in it. It has been further argued that, during the last glacial time, the high-latitude ocean productivity may be enhanced due t o increased wind-blown Fe supply, and this mechanism can explain why the atmospheric CO;? was approximately 90 ppm lower than the present day pre-anthropogenic level, as indicated by the ice-core C02 study (Neftel et al., 1982). Based on these, hot arguments are now going on, regarding whether or not the counter-balancing the increasing atmospheric C 0 2 due to fossil-fuel burning by the artificial Fe fertilization in the Antarctic region is feasible. The arguments are now under significant controversy and further experiments are needed t o resolve the questions. But what I want to point out here is t,hat measurement of Fe in seawater has been a real basic study in trace element geochemistry regardless of its biological significance and could not be accomplished until recently because of the difficulty of obtaining and handling water samples without contamination from the surroundings, i.e., rust of ship. Perhaps, Martin did not expect such an importance in the global change, when he started to work on Fe. This example clearly tells us that upgrading of broad ocean science is essential to our better understanding of the global system. 3.3.

Their usefulness as a paleoceanographic indicator

In the course of measuring some heavy metals in sea water, it was found t h a t Cd

is well correlated with P throughout the ocean. From this, Ed Boyle of MIT came up with an idea t h a t Cd in foraminiferra of deep-sea sediments may be used to deduce phosphate concentration in the paleo-ocean. After elaborated development of cleaning technique, he succeeded to obtain down-core Cd/Ca records which added important constraints on the geochemical cycling together with 0 and C isotope records. One of the important outcomes is a new view that, during the last glacial time, the rate of deep water formation in the North Atlantic was significantly smaller than today. Again, this change of ocean circulation and the resultant chemical redistribution may be related to the climatic change and various models are proposed to reproduce the lower atmospheric concentration of C02 during the last glacial time (see Boyle, 1988). Similarly, Ba/Ca and Ge/Si ratios in calcareous and siliceous tests may also serve as paleo-oceanographic indicators, because Ba and Ge are remarkably well correlated with alkalinity and dissolved Si throughout the ocean (Chan et al, 1977, Froelich and Andreae, 1981). These new techniques should help us to understand what happened on the ocean circulation and biogeochemical cycling in the past and t o predict what is going t o happen in the future.

4.

Concluding remarks

These three examples described above are all based on the fact t h a t we are now able t o measure chemical constituents of seawater well accurate enough t o use them as oceanographic tracers. Future development of techniques will provide more automated, precise and accurate methods. Using those techniques, we will be able to explore more sophisticated phenomena in the ocean. It has been a dream of marine

geochemists to obtain detailed three dimensional map of geochemical properties in the world oceans. A part of it may be achieved by WOCE hydrographic program. However, it is not planned t o cover any aspects of the geochemical cycling on a global scale. Originally, this must be a part of jobs of JGOFS, but unfortunately, it is now oriented to the process study and the global description is left behind. Therefore, we will need to have another world-wide ocean program in the near future which certainly requires international cooperation. One possible strategy is t o build a new vessel equipped with all analytical capabilities and let it do mapping along her track. It may take more than 5 years to complete the Pacific Ocean and perhaps 20 years for the world oceans. But certainly, it will contribute significantly to our better understanding of the global change.

Acknowledgement

I gratefully acknowledge the funding agency t o the priority area research program “Dynamics of the Deep Ocean Circulation”, the Ministry of Education, Science and Culture, Japan for support. References Boyle, E., 1988. The role of vertical chemical fractionation in controlling the late Quaternary atmospheric carbon dioxide. J. Geophys. Res., 93, 15701-15714. Brewer, P. G., 1975. Minor elements in sea water. In: Chemical Oceanography, Vol. 1, 2nd Edition, J. P. Riley and Skirrow (ed.), Academic Press, London, pp. 415-496. Chan, L. H., D. Drummond, J. M. Edmond, and B. Grant, 1977. On the barium data from the Atlantic GEOSECS Expedition. Deep-sea Res., 24, 613-649. Froelich, Jr., P. N., and M. 0. Andreae, 1981. The marine geochemistry of germanium: Ekasilicon. Science, 213, 205-207. Goldberg, E. D., 1963. The ocean as a chemical system. In: The Sea, Vol. 2 , M. N. Hill (ed.), Interscience Publ. John Wiley & Sons, New York, pp. 3-20. Martin, J. H., and S. E. Fitzwater, 1988. Iron deficiency limits phytoplankton growth in the north-east Pacific subarctic. Nature, 331, 341-343. Martin, J. H., R. M. Gordon, S. E. Fitzwater and W. W. Broenkow, 1989. VERTEX: Phytoplankton/iron studies in the Gulf of Alaska. Deep-sea Res., 36, 649-680. Measures, C. I., J. M. Edmond, and T. D. Jickells, 1986. Aluminum in the northwest Atlantic. Geochim. Cosmochim. Acta, 50, 1423-1428. Neftel, A., H. Oeshger, J. Schwander, B. Stauffer, and R. Zumbrunn, 1982. Ice core sample measurements give atmospheric COz content during the past 40,000 years. Nature, 295, 220-223. Nozaki, Y., 1992. Trace elments in sea water: Their mean concentrations and North Pacific profiles. Chikyukagaku (Geochemistry), 26, 35-39. Quinby-Hunt, M . S., and K. K . Turekian, 1983. Distribution of elements in sea water. EOS, 64, 130-131.

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

91

Philippine Sea Abyssal Waters in the Northwestern Pacific: Characterization from Tracer-Tracer Diagrams Toshitaka GAMO*

Abstract

In an effort to elucidate the origin of the Philippine Sea bottom water (PSBW), recent reliable chemical data of deep and bottom waters as well as bottom topographic features in the northwestern Pacific Ocean including the Philippine Sea were compiled and examined in detail. Although the PSBW is a branch from the bottom western boundary current flowing northward through the western Pacific Ocean, chemical characteristics of the PSBW were found to be significantly deviated from the trends of the main northwestern Pacific bottom water in Tp (potential temperature)-Oz, Tp-Si, and Tp-A14C diagrams. It was inferred that the PSBW is strongly influenced by vertical mixing between the PSBW and the overlying Philippine Sea deep water, whose characteristics are different from those of the main northwestern Pacific deep water, probably due to the topographic barrier of Izu-BoninMariana Ridge sequence.

1.

Introduction

The bottom water in the Pacific Ocean is replenished by the Antarctic Bottom water (AABW) flowing northward as a bottom western boundary current (Sverdrup et al., 1942; Stommel, 1958; Stommel and Arons, 1960; Knauss, 1962; Reid et al., 1968; Edmond et al., 1970; Craig et al., 1972). Wooster and Volkman (1960), Mantyla (1975) and Mantyla and Reid (1983) discussed the direction and branchings of the bottom current by mapping abyssal distributions of potential temperature and dissolved oxygen in the Pacific Ocean. The water below the depth of 5 km is defined as the bottom water in this paper. Figure 1 shows a generally believed bottom water circulation pattern connecting four major basins (A-D) in the western Pacific Ocean. Basins A, B, and C are the southern, central and northern basins, respectively, according t o the classification by Wooster and Volkman (1960). Basin D corresponds to the Philippine Sea. These basins are connected by narrow passages: the Samoan Passage between basins A and B with a sill depth of about 5,000 m (Reid and Lonsdale, 1974; Craig et al., 1981b), and the Wake Island Passage between basins B and C with a sill depth of >5,180 m (Chase et al., 1971). These passages are topographic constrictions against *Ocean Research Institute, University of Tokyo, Minamidai, Nakano-ltu, Tokyo 164, Japan

92

Fig. 1. A general bottom circulation pattern in the western Pacific Ocean connecting four basins A, B, C and D. The bottom topographic map was prepared by the GEOSECS Operating Group (Craig et al., 1981b). the northward bottom current. Although behaviors of the bottom current through basins A, B and C have been relatively well documented so far (e.g., Mantyla and Reid, 1983), little attention has been paid t o the behavior of bottom water from basins B to D as well as the origin of the Philippine Sea bottom water. Spatial variations of physical and chemical tracers in seawater give us useful information on water circulation and mixing processes. The purpose of this paper is t o characterize the Philippine Sea bottom water by comparing it with the other bottom waters in the Pacific Ocean using appropriate tracer-tracer diagrams. Emphasis will be placed on bottom water modification processes from basins B t o D in relation to bottom topographic features around basin D.

93 2.

Chemical Tracers

Tracers used in this study are potential temperature (Tp), salinity (S), dissolved oxygen ( 0 2 ) , silicate (Si) and carbon-14 (A14C). Data of Tp, S , 0 2 and Si are referred from those of the GEOSECS Pacific Expedition in 1973-1974 (Craig et al., 1981a), those of the INDOPAC Expedition Legs I1 and I11 in 1976 (Scripps Inst. Oceanogr., 1978), and those of VEGA Expedition in 1978 (Horibe, 1978). The locations of stations and their depths are listed in the Appendix. A14C data in the Pacific Ocean were referred from those of GEOSECS ("Ostlund and Stuiver, 1980), Horibe and Gamo (1980), Gamo et al. (1987), and Watanabe et al. (1987), and those in the Philippine Sea were from Horibe and Gamo (1980), Broecker et al. (1986), and Watanabe et al. (1987). In order t o evaluate d a t a consistency among the different expeditions, two pairs of stations with almost the same locations were selected for comparing S, T p , 0 2 and Si d a t a between GEOSECS and INDOPAC, and between INDOPAC and VEGA. The former location is a pair of GEOSECS station 223 (34"58'N, 151"51'E) and INDOPAC Leg-1 Station 87 (35"02'N, 151"55'E), and the latter is a pair of INDOPAC Leg-3 Station 24 (13"03'N, 136'30'E) and VEGA Station 4 (12"54'N, 136'27'E). Intercomparison of A14C values between GEOSECS d a t a and our measurements was described by Gamo et al. (1987). Systematic cruise-to-cruise shiftas of Tp, S, 0 2 , Si, and A14C were proved to be almost insignificant: less than the error of measurement for T p (f0.01"C) and S (f0.003), less than f 2 pmol kg-' for 0 2 , f 5 pmol kg-' for Si, and f 0 . 5 % for A14C.

3.

Geographical Settings and Previous Works

Figure 2 is a simplified depth chart of the Philippine Sea and its surroundings. The Philippine Sea is a lozenge-shaped basin, the depth of which is roughly in the range of 5000-6000 m, except for the Ryukyu and Philippine Trenches (maximum depth: 10,030 m) along the western margin of the Philippine Sea. To the east of the Sea, a sequence of Izu, Bonin, Mariana, Yap and Palau islands forms an almost continuous topographic high. This topographic sequence is called "South Honshu Ridge" according t o the topographic map by Chase et al. (1971). The South Honshu Ridge is a topographic barrier surrounding the eastern and southern sides of basin D. The westward bottom water (depth > 5 km) from basin B is supposed to pass through the East Mariana Basin, being greatly hindered from entering the Philippine Sea by the topographic barrier. Judging from the detailed topographic features, only one route is available for the bottom water to enter basin D: a narrow passage between the Yap and Mariana Islands (Yap-Mariana Junction) as indicated by the thick arrow in Fig. 2. Its sill depth is 5,000 m according to Uehara and Taira (1990). Detailed CTD-hydrocasts along 12"N and 13"N in the Philippine Sea have verified the existence of cold and saline bottom water in the southern Philippine Sea between the South Honshu Ridgz and the Kyushu-Palau Ridge, which evidences bottom water inflow from the East Mariana Basin to the southern Philippine Sea (Uehara and Taira, 1990).

94

Fig. 2. A bathymetric map of the Philippine Sea and its surroundings. Thin lines represent the 4 km contours. The thick arrow between the Mariana and Yap Islands shows the most probable route for the bottom current entering into the Philippine Sea (Yap-Mariana Junction). Filled circles and squares stand for INDOPAC and VEGA stations, respectively. Previous works on water chemistry around the Philippine Sea have suggested that the topographic barrier of the South Honshu Ridge might be responsible for unique characteristics of the Philippine Sea abyssal waters. Moriyasu (1972) showed that Tp, S and 0 2 values below 4,000 m depth are significantly different between

95 the Philippine Sea and the northwestern Pacific. Kaneko and Teramoto (1985) mapped the horizontal distributions of T p , S, 0 2 and Si along isopycnals of ug = 27.30, 27.70, and 27.77 (approximate depths of 1000, 2000, and 3000 m, respectively), showing that the isopycnal water exchanges between the Philippine Sea and the northwestern Pacific are more greatly restricted by the bottom topography as the isopycnals become deeper.

4.

Characterization from Tracer-Tracer Diagrams

As a bottom water flows away from its origin, its characteristics are gradually modified from the original ones. This is mainly due t o the following three processes: (i) filtering out of the deepest (densest) water by the sills, (ii) vertical mixing with overlying deep waters, and (iii) in situ physical and chemical processes. The modification is reflected in the distribution of appropriate physical and chemical tracers in seawater. For example, bottom water T p increases due to (i) filtering out of the coldest water, (ii) vertical mixing with overlying warmer deep waters, and (iii) geothermal heating, while 0 2 decreases due t o (i) filtering out of the 02-rich bottom water, (ii) vertical mixing with deep waters which are poorer in 0 2 , and (iii) 0 2 consumption by in situ decomposition of organic matter. A X-Y plot of an appropriate pair of tracers (a tracer-tracer diagram) is useful for bottom water characterization. Figure 3(a)-(d) show the tracer-tracer diagrams (S, 0 2 , Si, A14C versus T p ) for the western Pacific bottom waters below a depth of 5 km. It is apparent that the tracer-tracer characteristics vary from basin t o basin, reflecting bottom water modification processes. As expected above, the Tp0 2 diagram (Fig. 3(b)) indicates that an increase of bottom T p is accompanied by a decrease in 0 2 according t o the northward bottom flow path shown in Fig. 1. It is noteworthy that the relationship between T p and S, 0 2 , Si or A14C is not a simple straight line, but has two remarkable discontinuities as described below: (1) a marked bending point between basins A and B, and (2) significant standoff of basin D characteristics from a smooth trend connecting basins B and C. A proper interpretation for the first discontinuity may come from the vertical mixing between bottom water and overlying deep water in basins A and B in addition t o the sill filtering effect. In basin A, the northward AABW is modified by vertical mixing with the overlying Circumpolar deep water (CDW). On the contrary, basin B bottom water is not significantly affected by CDW but rather modified by basin B deep water. Figure 4(a)-(c) compare the tracer-tracer diagrams (Tp-S, Tp-02, and Tp-Si) of deep and bottom waters below 500 m depth among the four basins. As shown in Fig. 4, CDW is significantly higher in 0 2 and S values, and lower in Si than the deep water in basin B. These difference of deep water characteristics between basins A and B is thought t o play a significant role in producing the A-B bendings shown in Fig. 3(a)-(d). Before discussing the second discontinuity, we define the following abbreviations for convenience. The northwestern Pacific excluding the Philippine Sea, i.e., basin C, is simply called “the northwestern Pacific”, and the bottom water (depth

96

1

t 0.5

1.o

Potential Temperature

("C)

Fig. 3. Tracer-tracer diagrams below the depth of 5 km in the western Pacific Ocean. (a) Tp-S, (b) Tp-02, (c) Tp-Si diagrams (data sources are Craig et al. (1981a), Scripps Inst. Oceanogr. (1987), and Horibe (1978)), and (d) Tp-A14C diagram, where open circles are from "Ostlund and Stuiver (1980), filled squares from Horibe and Gamo (1980) and Gamo et al. (1987), an open triangle from Broecker et al. (1986), and filled triangles from Watanabe et al. (1987).

> 5 km) in the northwestern Pacific is called NWPBW (North-Western Pacific Bottom Water). The Philippine Sea bottom water is abbreviated t o PSBW. The origin of NWPBW and PSBW must be commonly basin B as shown in Fig. 1. Why do the characteristics of the PSBW and NWPBW show so significantly different trends from each other as shown in Fig. 3(b)-(d)? Since the Tp-S diagram

97

0.5

1 .o

Potential Temperature

0.5

("C)

1.0 Potential Temperature

Fig. 3.

("C 1

(continued)

(Fig. 3(a)) shows little standoff of PSBW, the idea of higher heat flow in basin D than basin C may be rejected.

5.

Modifications of the Source Bottom Water by Vertical Mixing

Since vertical mixing with an overlying deep water is an important process t o alter the tracer-tracer characteristics of a bottom water as already shown in the case of the A-B bending in the previous section, deep waters in basins C and D may provide a clue to solve the problem. As mentioned in Section 3, it has been known that the deep waters in basins C and D have different characteristics due t o the topographic barrier of the South Honshu Ridge. Then it should be examined

98 whether the difference between the PSBW and NWPBW is related t o the difference of the deep water characteristics between the two basins. In Fig. 4(a)-(c), Tp-S, Tp-02, and Tp-Si diagrams of the deep and bottom waters below 500 m depth were compared among basins A-D. These diagrams clearly demonstrate that the deep water characteristics are significantly different between basins C and D. The Tp-S diagram of basin D is simply sandwiched by those of the basins B and C (Fig. 4(a)), while the Tp-02 and Tp-Si diagrams (Fig. 4(b) and (c)) are remarkably different in their shapes between the two basins. Basin C has relatively a sharp 0 2 minimum layer with 0 2 concentration of 20-50 pmol kg-' centered at T p = 3-4OC, while basin D has a broader 0 2 minimum with higher O2 values (60-100 pmol kg-') at T p = 446°C. As for Tp-Si diagram, basin C deep water has a well developed Si maximum layer centered at T p = 1.52.0°C, while basin D shows much less Si maximum layer centered at slightly less T p range. It is of interest that the Tp-S diagram (Fig. 3(a)) shows less standoff of the PSBW than the T p - 0 2 , Tp-Si, and Tp-A14C (Fig. 3(b)-(d)). This is probably because the depth of salinity minimum, which occurs at approximately 700 m depth in the midlatitudinal northwestern Pacific Ocean, is shallower than the depths of

I

0

I

--

I

I

I

I

I

I

I

5

I

10

Potential Temperature ("C) Fig. 4. (a) Tp-S, (b) Tp-02 and (c) Tp-Si diagrams of deep and bottom waters below the depth of 500 m in basins A (+), B (A), C ( 0 ) and D (0) (Craig et al., 1981a; Scripps Inst. Oceanogr., 1978; Horibe, 1978).

99 01

0

I

I

I

I

I

I

,

5 Potential Temperature ("C) Fig. 4.

I

10

(continued)

the 0 2 minimum (at about 1,200 m depth), Si maximum (2,300-2,500 m), and A14C minimum (2,000-2,500 m). The barrier effect of the South Honshu Ridge is much smaller at the salinity minimum depth (700 m) than those at the maximum or minimum depths (1,200-2,500m) for the other components. The salinity minimum water can therefore easily exchange between the basins C and D, while it may be more difficult for the 0 2 minimum, Si maximum and A14C minimum waters in basin C to enter basin D. The vertical distributions of 0 2 and Si in the northwestern Pacific are characterized by well developed mid-depth minimum and maximum, respectively (see Craig et al., 1981b). In contrast, the magnitudes of 0 2 minimum and Si maximum in the basin D are much less than those in basin C. This is probably because the supply of 02-poor and Si-rich intermediate water from the northwestern Pacific to the Philippine Sea is restricted by the barrier of the South Honshu Ridge as mentioned above. This barrier effect may contribute t o the higher 0 2 values in the basin D deep water below the 0 2 minimum layer than those in the basin C deep water as shown in Fig. 4(b). On the other hand, Reid and Mantyla (1978) showed that the mid-depth water circulation in the western Pacific carries 02-rich intermediate water from the south Pacific along the western boundary current. This process may also play a role in maintaining high 0 2 values in basin D deep water. Figure 5 schematically illustrates the bottom water modification processes in basins C and D. When the westward branched bottom water from basin B passes

100

0

200

,

I

I

I

I

I

I

I

I

I

I

I

5

I

I 10

Potential Temperature ("C) Fig. 4.

(continued)

through the Yap-Mariana Junction, it must be modified not only by the sill filtering effect but also by vertical mixing with basin D deep water which is richer in 0 2 , poorer in Si and probably has higher A14C value than the northwestern Pacific deep water. On the other hand, the northward bottom current through the Wake Island Passage is modified t o the NWPBW under the influence of basin C deep water. This is probably the reason why the bottom Tp-02, Tp-Si, and Tp-A14C characteristics (Fig. 3(b)-(d)) are significantly different between basins C and D. It should be noted that the one-way steady bottom currents as simply illustrated by the branched arrows in Fig. 5 are rather conceptional and qualitative ones. In the real ocean, there may be active isopycnal mixing, repeated cyclonic circulations in a basin as well as the steady bottom currents. In the course of such complicated water movements with various time scales, bottom water in a basin will be gradually modified by the three main processes (filtering effect, vertical mixing and in situ chemical process) as discussed in Section 4.

6.

Concluding Remarks

Bottom Tp-02, Tp-Si and Tp-A14C diagrams in the western Pacific Ocean demonstrated that the PSBW has different characteristics from those of the NWPBW, in spite of the fact that both bottom waters are branched from a common source of the central Pacific Basin bottom water. This difference is attributable

101 South Honshu Ridge

, ,

\M v

Bottom Current

Basin B

Fig. 5. A schematic illustration showing modification processes of the bottom waters in basins C and D. The arrows represent the movement and branching of the bottom water originated from basin B, and the two cylinders indicate the deep waters in basins C and D. The width and thickness of the arrows and cylinders are not quantitative. to bottom water modification processes, particularly t o vertical mixing between the source bottom water and the overlying deep waters. The tracer-tracer diagrams show t h a t the Philippine Sea deep water is richer in 0 2 and poorer in Si than the deep water of the northwestern Pacific Ocean, which may cause t h e difference between the PSBW and NWPBW characteristics. It is suggested t h a t the topographic barrier of the Izu-Bonin-Mariana ridge sequence (South Honshu Ridge) plays a dominant role in characterizing the Philippine Sea deep water by obstructing horizontal deep water exchange between the Philippine Sea and the northwestern Pacific Ocean.

Acknowledgement The author wishes t o thank Prof. Y. Horibe for his superior direction t o tracer studies. Thanks are also due t o Drs. M. Fukasawa and I. Kaneko for valued discussion. This study was financially supported by the grant-in-aid “Dynamics of the Deep Ocean Circulation” (Nos. 62610504 and 62610002) from t h e Ministry of Education, Science and Culture t o the University of Tokyo.

References Broecker, W. S., W. C. Patzert, J. R. Toggweiler, and M. Stuiver, 1986. Hydrography, chemistry and radioisotopes in the southeast Asian basins. J. Geophys. Res., 91, 14345-14354. Chase, T. E., H. W. Menard, and J. Mammerickx, 1971. Bathymetry of the North Pacific. Scripps Institution of Oceanography and Institute of Marine Resources. Craig, H., Y . Chung, and M. Fiadeiro, 1972. A benthic front in the south Pacific. Earth Planet. Sci. Lett., 16, 50-65.

102 Craig, H., W. S. Broecker, and D. Spencer, 1981a. GEOSECS Pacific Expedition, Vol. 3. U.S. Government Printing Office, 137 pp. Edmond, J. M., Y. Chung, and H. Craig, 1970. Pacific bottom water: penetration east around Hawaii. J. Geophys. Res., 76, 8089-8097. Gamo, T, Y. Horibe, and H. Kobayashi, 1987. Comparison of oceanic A'*C data with those of GEOSECS: vertical profiles in 1973 (GEOSECS) and in 1980 at (30°N, 170'E) in the northwestern Pacific Ocean. Radiocarbon, 29, 53-56. Horibe, Y., 1978. Preliminary report of Hakuho Maru Cruise KH-78-1: VEGA Expedition (unpublished). Horibe, Y., and T. Gamo, 1980. Chemical characteristics of the Cold Water mass of Kuroshio. In: The Kuroshio IV (Proc. 4th CSK Symp.), Tokyo, pp. 360-370. Kaneko, I., and T. Teramoto, 1985. Sea water exchange between the Shikoku-Philippine Basin and the western North Pacific Basin. In: Ocean Characteristics and their Changes, K. Kajiwara (ed.), Koseisha-Koseikaku, Tokyo, pp. 54-77 (in Japanese). Knauss, J. A., 1962. On some aspects of the deep circulation in the Pacific. J. Geophys. Res., 67, 3943-3954. Mantyla, A. W., 1975. On the potential temperature in the abyssal Pacific Ocean. J. Mar. Res., 33, 341-354. Mantyla, A. W , and J. L. Reid, 1983. Abyssal characteristics of the world ocean waters. Deep-sea Res., 30, 805-833. Moriyasu, S., 1972. Deep waters in the western north Pacific. In: Kuroshio: Its Physical Aspects, H. Stommel and K. Yoshida (ed.), University of Tokyo press, pp. 387-408. "Ostlund, H. G and M. Stuiver, 1989. GEOSECS Pacific Radiocarbon. Radiocarbon, 22, 25-53. Reid, J. L., H. Stommel, E. D. Stroup, and B. A. Warren, 1968. Detection of a deep boundary current in the western South Pacific. Nature, 217, 937. Reid, J. L., and P. F. Lonsdale, 1978. On the flow of water through the Samoan Passage. J. Phys. Oceanogr., 4, 58-73. Reid, J. L., and A. W. Mantyla, 1978. On the mid-depth circulation of the north Pacific Ocean. J. Phys. Oceanogr., 8 , 946-951. Scripps Institution of Oceanography, 1978. Data report, physical, chemical and biological data, INDOPAC Expedition. Univ. of California, San Diego, Scripps Inst. Oceanogr. Ref. 78-21, 424 pp. Stommel, H., 1958. The abyssal circulation. Deep-sea Res., 5, 80-82. Stommel, H., and A. B. Arons, 1960. On the abyssal circulation of the world ocean - 11. Deep-sea Res., 6, 217-233. Sverdrup, H. U., M. W. Johnson, and R. H. Fleming, 1942. The Oceans: Their Physics, Chemistry and General Biology. Prentice-Hall, New York, 1087 pp. and K. Taira, 1990. Deep hydrographic structure along 12 N and 13 N in the Uehara, K., Philippine Sea. J. Oceanogr. SOC.Japan, 46, 167-176. Watanabe, S., M. Nakajima and S. Tsunogai, 1897. 14C in abyssal waters of the Philippine Basin. Annual Meeting of the Oceanographical Society of Japan, Shimizu, October 1987, Abstr. No. 361 (in Japanese). Wooster, W. S., and G. H. Volkman, 1960. Indications of deep Pacific circulation from the distribution of properties at five kilometers. J. Geophys. Res., 65, 1239-1249.

103 Appendix. List of stations in basins A-D referred in this study. GEOSECS, INDOPAC, and VEGA data are according to Craig et al. (1981a), Scripps Institution of Oceanography (1978), and Horibe (1978). Station No. Basin A GEOSECS-257 -263 -269 -282 -290 -293 -296 -303 -306 Basin B GEOSECS-229 -231 -239 -241 -246 -251 Basin C GEOSECS-204 -212 -213 -214 -217 -222 -223 -224 -225 -226 -227 INDOPAC Leg-2- 6 -7 -8 -9 -11 -14 -15 -16 VEGA-11 VEGA-12

Location

Depth(m)

(10"1O'S, (16"39'S, (23"59'S, (570353, (58"00'S, (52"40'S, (440595, (38"22'S, (32'503,

169'58'W) 167'04'W) 174'26'W) 169"36'E) 174"OO'W) 178"05'W) 166'42'W) 17Oo04'W) 163'38'W)

5,180 5,713 5,999 5,213 5,316 5,335 5,342 4,841 5,624

(12"53'N, ( 14"07'N, ( 5"53'N, ( 4"33'N,

( o"oo's, ( 4"34'S,

173"28'E) 178"34'W) 172"OO'W) 179'00'E) 178"59'E) 178"57'E)

5,730 5,663 5,774 5,738 5,412 5,376

(31"22'N, (30"00'N, (31"00'N, (32"01'N, (44"40'N, (40"lo", (34"58'N, (34" 15'N, (32"37'N, (30"34'N, (25"00'N, (31"43'N, (3l"4 1'N, (28"58'N, (28" 16'N, (26"01'N, (22"30'N, (22"28'N, (22"32'N, (29"40'N, (34" 16'N,

150"02'W) 159"50'W) 168'27'W) 176"59'W) 177'03'W) 160"30'E) 151"50'E) 141'58'E) 161"55'E) 170'36'E) 170"05'E) 142" 12'E) 143'20'E) 143'48%) 142'49'E) 143O14'E) 145"02'E) 145"54'E) 147"31'E) 146'16%) 142"OO'E)

5,410 5,733 5,683 5,306 5,762 5,579 6,131 9,739 5,948 5,491 5,900 9,100 5,801 5,862 4,949 4,484 4,517 7,526 5,766 6,120 9.000

104 Appendix. Station No. Basin D INDOPAC Leg-3-23 -24 -25 -26 -27 -28 -29 -30 -31 -35 -36 -39 -40 -41 -42 -44 -45 -46 -47 -48 -49 -50 -51 -52 -54 -55 VEGA-1 VEGA-2 VEGA-3 VEGA-4 VEGA-5 VEGA-7 VEGA-8 VEGA-9

(continued) Location

(13"00'N, (13"03'N, (12"59'N, (13"01'N, (12"59'N, (13"01'N, ( 12" 59'N, (12"59'N, (13"02'N, (19"38'N, (20"00'N, (22"35'N, (22"35'N, (22"37'N, (23" 15'N, (22"28'N, (22"26'N, (19"20'N, (20"55'N, (22" 31'N, (22" 31'N, (22"32'N, (22"29'N, (22"33'N, (19"31'N, (18"26'N, (26"11'N, (22"00'N, ( 17" 35'N, (12"54'N, (18"11'N, (26"12'N, (30"31'N, (31"30'N,

138O59'E) 136'30'E) 133"58'E) 132"29'E) 130" 48'E) 128'59'E) 127"43'E) 126"40'E) 125"37'E) 123"53'E) 123"13'E) 122"ll'E) 123'06'E) 124"20'E) 124"43'E) 126"32'E) 129'34'E) 13l"06'E) 131"57'E) 132" 27'E) 134"OO'E) 135"59'E) 137'20'E) 139'25'E) 140"06'E) 141"14'E) 136'42'E) 137" 5 1'E) 137'14'E) 136" 27'E) 132" 25 'E) 129'36'E) 134'02'E) 137"OO'E)

Depth(m) 5,000 5,461 5,517 5,717 5,939 5,737 5,442 5,374 8,076 5,452 5,206 4,800 5,661 5,735 6,566 5,919 6,231 5,993 5,978 5,592 5,450 5,294 4,490 4,523 4,964 4,656 5,300 4,425 5,150 5,400 6,000 6,370 4,490 4,150

Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

105

Dynamics of the Japan Sea Deep Water Studied with Chemical and Radiochemical Tracers Shizuo TSUNOGAI, Yutaka W. WATANABE, Koh HARADA, Shuichi WATANABE, Shimei S A I T O a n d Michio NAKAJIMA"

Abstract

Chemical and radiochemical tracers have been used to make clear the dynamics of the Japan Sea deep water including its chemical alteration. They are the Si-0 combination, tritium, CFCs (trichlorofluoromethane and dichlorodifluoromethane), C-14 and Ra-226. The Si-0 diagram method has revealed a small areal variation of the Japan Sea deep water except water in a small basin surrounded by the Yamatotai Rise and its coastal upwelling off Japan. The tritium and Ra-226 combination method is applied to a box model with three layers and four boxes, which is based on the vertical profiles of tracer components including conservative and nutrient elements. The result shows that the turnover time of the Japan Sea water for vertical mixing is about 100 yrs and the residence time of the whole Japan Sea water except the warm Tsushima current water is about 1000 yrs, indicating that the Tsushima current flowing through the Japan Sea makes little interaction with the Japan Sea Proper Water. The difference between the apparent C-14 age of the Japan Sea deep water of 300 years and the turnover time is originated in the delay in the air-sea gas exchange equilibrium of carbon dioxide. The gas transfer velocity has been calculated to be 1.5 m/day. The vertical profiles of CFCs also support the estimated water exchange rates between the reservoirs, but our present data are insufficient to explain their small apparent gas exchange rates less than 0.1 m/day. The oxygen consumption rate of the Japan Sea deep water is nearly equal to that of the Pacific deep water indicating that the Japan Sea is not fertile, while the regeneration rate of silica in the Japan Sea is much larger than that in the Pacific.

1.

Introduction

Recently chemical and radiochemical tracers have been well used t o reveal t h e oceanographic problems especially in t h e dynamic aspects of water movement (e.g., Broecker and Peng, 1982) a n d will play a major role in t h e many international projects such a s W O C E (World Ocean Circulation Experiment). In this s t u d y we have tried t o apply t h e tracer method t o t h e J a p a n Sea water as a n example for making a future extensive study program in t h e North Pacific. Those tracers are *Department of Chemistry, Faculty of Fisheries, Hokkaido University, Hakodate 041, Japan

106 the dissolved silica-oxygen combination, Ra-226, tritium (H-3), radiocarbon ((2-14) and chlorofluorocarbons (CFCs). The Japan Sea ( 1 . 0 1 3 ~ 1 0km2) ~ is one of the largest marginal seas in the world (Menard and Smith, 1966). The Japan Sea is isolated from the Pacific except the four straits of which sill depths are shallower than 130 m (Fig. l), although its maximum and mean depth are, respectively, 4036 m and 1667 m. Thus the deep water of the Japan Sea should be formed within the sea, probably in winter in the northern area (Uda, 1934; Fukuoka, 1965). The Japan Sea water below 200-300 m is occupied by a remarkably uniform water mass, which is called the Japan Sea Proper Water having relatively low temperature of 0.1-0.3"C and low salinity of 34.0-34.1 psu as compared with the Pacific Deep Water. Since the Japan Sea deep water contains much dissolved oxygen, it is presumed t h a t the vertical mixing is active and its turnover time is short (Uda, 1934). Indeed the carbon-14 age of the bottom water has been determined t o be 300 years (Gamo and Horibe, 1983). On the other hand, Harada

Fig. 1. The Japan Sea. Isobaths of 200, 1000 and 3500 m depth are shown. Solid circIes with numbers are the observation stations occupied by Horibe (ed., 1981).

107 and Tsunogai (1986) have reported the relatively high concentration of Ra-226 in the deep Japan Sea water showing its longer residence time within the sea, whereas Watanabe et al. (1991) have found the bomb tritium in the deep Japan Sea water more than 2000 m depth indicating an active vertical mixing. In this study we have tried t o give a unifying interpretation t o these apparently incompatible results with a quantitative model calculation.

2.

Si-0 Diagram Method

A classical Si-0 method is applied t o the Japan Sea water for understanding its structure before the discussion on its dynamics obtained by using modern tracers. The Si-0 method has been successfully applied to the North Pacific water (Tsunogai, 1987). For example, one of the results clearly shows t h a t the Pacific Deep water in the western North Pacific flows along the Kuril-Kamchatka Trench and into the eastern North Pacific via the north end of the Emperor Seamount chain and the Aleutian Trench. Here we have treated the Japan sea water observed by Horibe (ed., 1981) in the same manner. The Si-0 diagram of the Japan Sea water is completely different from those

Pacific

200

0 -

2oo(uM/kg)

300

Fig. 2. Si-0 diagram for waters collected in the Japan Sea, the Antarctic Ocean (AAO), the western South Pacific (WSP), the eastern South Pacific (ENP), the western North Pacific (WNP), the central North Pacific and the Bering Sea (BS).

108

!ol

$,.,,*,

\

137'21' E

tN

Fig. 3. Concentrations of dissolved silica (Si in pM), nitrate (N in pM) and phosphate (P in pLM), and potential temperature ( T p in "C) at Sta. 18. in the other oceans (Fig. 2). This is evidence t h a t the Japan Sea is isolated and its deep water is formed independently of the neighboring Pacific and its marginal seas. It is interesting to note that the silica content in the Japan Sea deep water below the oxygen minimum layer is the highest among those containing the same amount of dissolved oxygen in the world ocean, although its absolute value (80 pM) high. This is probably due to the same cause for the relatively high concentration of Ra-226 in the Japan Sea deep water (Harada and Tsunogai, 1986), which is discussed later. The vertical profile of dissolved silica in the Japan Sea (Fig. 3 ) shows t h a t the concentration of silica increases markedly with depth down to about 1000 m depth and the increase rate largely steps down in the deep water below the depth, where those of temperature, nitrate and phosphate are almost constant with depth. T h e vertical profiles of silica between the 200-1000 m depth are different from station t o station but are parallel with each other (Fig. 4). The horizontal distribution of dissolved silica a t the 600 m surface is compared in Fig. 5 . We can see the highest value at the station nearest t o the Japanese coast, Sta. 13, indicating the upwelling of 100-200 m in the intermediate water a t Sta. 13 near the coast as compared to Sta. 1 4 which is 65 km apart to the offshore. The upwelling is confirmed by the Si-

109 20

40

60

Si

r

(m)

Fig. 4. Vertical profiles of dissolved silica (in pM) at Stas. 13, 14 and 15 occupied by Horibe (ed., 1981).

0 diagram given in Fig. 6a, where the S i - 0 curve for the water below 280 m depth or containing dissolved silica more than 40 pM at Sta. 13 (Fig. 6a) substantially coincides with those at other stations in the diagram. The Si-0 diagrams for the Japan Sea deep water below 1000 m at various stations (Fig. 6) coincide fairly well with each other except one station, Sta. 22, in a small basin surrounded by the Yamatotai Rise (Fig. 6b). The concentration of silica increases with depth even in the deep water at Sta. 22 (water depth, 2200 m). Therefore the highest concentration of dissolved silica is not found in the main Japan Sea basin deeper than 3500 m, depth, but found in the bottom of the Yamatotai basin (around Sta. 22 in Fig. 1). This may indicate t h a t the bottom water of the Yamatotai basin is more stagnant but the main basin water is vertically well mixed. In the surface 200 m layer the concentrations of dissolved silica and other components are variable from station to station. At Stas. 24 and 25 near the Tsushima Strait, the lower oxygen (< 200 pM) and higher silica(> 10 p M ) water at about 100 m depth (Fig. 6c) is clearly distinguished in the Si-0 diagram (as compared t o Fig. 6a,b). The water may have originated in the water flowing from the bottom of the continental shelf region of the East China Sea through Tsushima Strait. Based on the vertical distributions of dissolved silica and conservative elements such as water temperature described above. The Japan Sea water can be divided into three layers, the surface (0-200 m), the intermediate (200-1000 m) and the deep (below 1000 m depth) layers. According to Harada and Tsunogai (1986), the

110 surface layer is further divided into the wa, d cold surface reservoirs. The warm surface is covered by the warm Tsushima current which makes a steep thermocline preventing the water exchange with the deep water. The boundary between the warm and cold surface reservoirs is not clear and the area and depth of the cold surface region may be enlarged in winter when the surface is cooled. In the intermediate water below 300 or 400 m the concentration of silica increases almost linearly and largely with depth (e.g., Figs. 3 and 4). This indicates t h a t the vertical mixing is prevailing rather than the advective flow (upwelling) in the intermediate water of a large part of Japan Sea, which can not be clearly realized by non conservative dissolved oxygen and nutrients which are influenced by the decomposition of organic matter and by the conservative elements of which vertical variation is small. The usefulness of silica is due t o its regeneration at the bottom.

3. Transient Tracers and the Vertical Mixing in the Japan Sea 3.1.

Vertical profiles of transient tracers

We have determined the concentrations of tritium (Watanabe et al., 1991), CFCs (Watanabe et al., 1988), C-14 (Watanabe et al., 1989) and Ra-226 (Harada

45-

Fig. 5 . Horizontal distribution of dissolved silica (in pM) a t 600 m depth. Solid points indicate the observation stations.

111 and Tsunogai, 1986). The water samples for tritium, C-14 and CFCs were collected in 1987 and 1988, and their vertical profiles were determined at 7 , 1 and 7 stations for tritium, C-14 and CFCs, respectively. The more detailed report on the CFCs and C-14 data will be reported elsewhere. In Fig. 7 their mean values are shown, where their scales are designed to make the comparison easy. Tritium (half-life, 12.4 yrs) in the present seawater has been produced chiefly by the H bomb explosions in the 1960’s and transported t o the sea surface. The rapid decrease in its concentration with depth (Fig. 7) suggests a certain time

801

200

8Ot

\.

1( i\ 80 sot

,

60

I

200

300

Fig. 6. Si-0 diagram (in unit of pM) for the Japan Sea water at Stas. 13, 14 and 15 in Fig. 6a, a t Stas. 19, 22 and 23 in Fig. 6b, and at Stas. 24 and 25 in Fig. 6c.

112

8 " "-" "

Tritium(T.U.) 2

1

3000b

0.5

4000

'

-60 I

1 .o

4

3

-F - l l , ( p M ) , , 1.5

,

,

,

,

.......... c - 1 4 ("00) ' -46 ' -iO ' ( ! I ' 20 ' - - - - - Ra-226 (dpm(m3) 140

120

100

,

, ,

2.0

' 40' 1

80

Fig. 7. Vertical distributions of tritium, CFCs (F-11and F-12), C-14 (A1*C value) and Ra-226 (radioactivity). Their units and scales are given within the figure. The concentrations are in the level of 1/1/1988. necessary for the vertical mixing of water bearing the tritium, but its detection in the water below 1500 or 2000 m depth indicates an active vertical mixing and advection (upwelling and downwelling) as compared to that in the Pacific water. Watanabe et al. (1991) have estimated the turnover time of the Japan Sea deep water below 200 m depth t o be about 100 yrs, based on its vertical distribution and a crude model. The concentrations of CFCs, F-11 (trichlorofluoromethane) arid F-12 (dichlorodifluoromethane) in the Japan Sea water also decreased with depth and the decrease rate is larger than t h a t of tritium. This steeper decreases in F-11 and F-12 is probably due t o the fact that CFCs are released t o the atmosphere more recently with a large increase rate than tritium. Upon examining Fig. 7 closely, we can recognize tha t F-11 decreases more steeply with depth than F-12. This may be caused by the release rate of F-11/F-12 which increased with time until 1975 (Smethie et al., 1988). The A14C value of the Japan Sea deep water was nearly the same as that determined by Gamo and Horibe (1983) in 1977 (-74 ym). The rapid decrease rate in the C-14 values with depth seems to be due to the larger capacity of dissolved inorganic carbon (carbonate) in seawater, which induces the delay in the air-sea exchange equilibrium of the carbon system, as compared to sparingly soluble gases

113 such as CFCs. The concentration of Ra-226 inversely increases with depth. Since the distribution is probably stationary, this is due to its sources at the bottom, the dissolution from biogenic particles which contain Ra-226 removed from the surface water, and from the bottom sediments which contain Ra-226 produced from the decay of Th230. As discussed by Harada and Tsunogai (1986) the Ra-226 concentration in the Japan Sea deep water is large as compared to that in other oceans containing the same amount of C-14 or dissolved oxygen. This niay be due t o the fact the Japan Sea water is vertically well mixed and most of the surface water upwelled in winter does not flow out to the Pacific, but sinks again after dissolving the atmospheric oxygen and some radiocarbon at the surface. We therefore must distinguish the turnover time of the Japan Sea water for the vertical mixing and its residence time

Fig. 8. A box model with three layers and four boxes of the Japan Sea. The solutions for the case of T (the proportion of the cold surface) of 0.5 are also shown. The numerical values given are the bottom area, S , the volume of water, V , the production (or loss) rates of tritium and Ra-226, P , and the observed concentrations of tritium and Ra-226. T h e six unknowns solved for T = 0.5 and the residence times calculated from the solutions are underlined.

114 within the Japan Sea excluding the warm surface water (Tsushima current) flowing in and out like river water.

3.2.

Three layers model

We have calculated the residence time of the Japan Sea water using a box model with three layers and four boxes (Fig. 8), which has been stated in Section 2 for t:-e Si-0 diagram. The model is basically the same as t h a t used by our previous worb (Harada and Tsunogai, 1986; Watanabe et al., 1991) except for the following three points. 1) We add the intermediate water reservoir to the previous model. Then the boxes are the warm surface, the cold surface, the intermediate and the deep reservoirs. It is assumed t h a t the warm surface does not directly exchange water with the intermediate reservoir (Fig. 8). 2) More recent estimate of the hypsometry of the Japan Sea basin (Menard and Smith, 1966) shows somewhat different figures from the older ones used by the previous works. The area, volume, mean depth of the Japan Sea are, respectively, 1 . 0 1 3 ~ 1 0km2, ~ 1 . 6 9 0 ~ 1 0km3 ~ and 1667 m. We have also used the depth zones in the Japan Sea area estimated by them. The area and volume of each box are given in Fig. 8. 3) The particulate flux of Ra-226 (345 dpm/m2/yr) used in this study has been calculated from the observed ratio of excess Ra-226 t o opal in the sediment trap samples collected in the Atlantic (Brewer et al., 1980; Honjo, 1980), and the observed particulate flux of opal in the Japan Sea (Masuzawa et al., 1990). We assume the particulate flux of Ra-226 is proportional t o the surface area and the excess Ra-226 defined by Brewer et al. (1980) is completely dissolved within a few hundred years. We further assume t h a t the benthic flux of Ra-226 is proportional to the bottom area. The time change in the concentration of a tracer element in each box is expresses as follows:

where C , V and P are the concentration of a tracer element, volume of water, production (or loss) rate of C , respectively, and the suffixes, w, c, m and d denote the warm surface, the cold surface, the intermediate and the deep reservoirs. T h e rates of water exchange, K 1 , Kz and K3 are of those between the warm and the cold surface reservoirs, the cold and the intermediate water, and the intermediate and the deep waters, respectively. The proportion of the cold surface t o the whole surface, r , seems to be variable from season to season and we regard the proportion as one of parameters. Then V, = V,r where V, is the volume of the whole surface water. The residence times of the deep water (Td), of the intermediate and deep water (Tmd) and of the whole Japan Sea water excluding the warm surface water (Tcmd) are obtained from the following equations; if we can determine the K values,

115

I

/

_ _ _ _ _ _ _-_ _ _ _ _ k_ _ - - --------,191

m

I

1

‘0.4

/

0.5

0.6

r

0.7

0.8

0.9

Fig. 9. The solutions as a function of r , where T is the residence time of water in the respective reservoirs expressed with suffixes, K is the exchange rate of water, and k is the gas transfer velocity of carbon dioxide at the surface, which is calculated from the K values and the vertical profile of C-14. Tcmd

+ vm + vd)/Kl

= (6

+ [(vm+ vd)/K2][(vm+ v d ) / ( & + v m + v d ) ] + (Vd/K3)[Vd/(T/c+ Vm + %)I.

(6)

We have solved the three K values and other three unknowns (C, and Cd for tritium and C, for Ra-226) by introducing the observed mean concentrations of C, (3.6 T.U. in 1988) and C, (2.6 T.U.) for tritium and C, (81 dpm/m3), C, (134 dpm/m3) and c d (154 dpm/m3) for Ra-226 into Eqs. (1) t o ( 3 ) for tritium and Ra-226, namely six equations. The P value for tritium is the input rate from the atmosphere to the surface and the radioactive decay which is proportional t o its amount in the respective box. The P value for Ra-226 consists of the particulate and benthic fluxes which are assumed to be proportional to the area, and the radioactive decay. The benthic flux of Ra-226 from a unit area is the same as t h a t used by Harada and Tsunogai (1986). A steady state condition ( d C / d t = 0) can be applied for Ra-226, but not for tritium. The details of the calculation are given in Watanabe et al. (1991). The results obtained are expressed as a function of r (Fig. 9). The results do not highly depend on the T value which can be roughly regarded t o be 0.5 according t o H a t a (1962). In Fig. 8, therefore, the K values and other unknowns obtained for the case of T = 0.5 are also shown. The residence times calculated from the above A’ values and Eqs. (4) t o (6) for T = 0.5 are 100, 110 and 990 years for respectively, T d , T m d and Tcmd,which are nearly the same as those obtained by Watanabe et al. (1991). They have reported that the residence times of the deep water below 200 m and t h a t of the cold surface water are, respectively, about 100 and 1000 yrs. The calculated concentration in the deep water, C d = 0.5 T.U., is also reasonable as compared with the observed

116 mean concentration, 0.4 T.U. (Watanabe et al., 1991). Using t h e d a t a observed by Horibe (ed., 1981), we have estimated t h e mean concentrations of dissolved oxygen, which are 242.2 pM for C, and 224.6 pM for Cd. By introducing t h e K3 value obtained here ( 9 . 3 ~ 1 0 ' m3/yr), ~ t h e oxygen consumption rate in t h e deep water below 1000 m depth turns out to be 164 x lo9 moles/yr or 172 pmoles/m3/yr or 3.8 ml/m3/yr. T h e consumption rate in a unit volume is comparable with t h a t in t h e Pacific deep water, 3.1 ml/m3/yr (Tsunogai, 1972). T h e rate in a unit area is 0.26 moles 0 2 / m 2 / y r , which is 43% of t h e organic carbon flux directly observed in September by t h e sediment traps in t h e surface 1000 m, 0.61 moles C/m2/yr (Tsunogai and Noriki, 1987). T h e difference may be d u e to t h e decomposition of organic compounds i n t h e water above 1000 m depth or t h e larger particulate flux in the summer season. At any rate the Japan Sea is not so fertile. T h e mean concentrations of dissolved silica are 49.2 p M for C,, and 73.0 pM for Cd (Horibe, ed., 1981). T h e regeneration rate of silica in the Japan Sea deep water below 1000 m depth is calculated to be 221 x l o 9 moles/yr or 232 pmoles/m3/yr or 0.36 moles/m2/yr. T h e rate in a unit area agrees fairly well with t h e observed flux of opal, 0.33 moles/rn2/yr (Masuzawa et al., 1989). According to Tsunogai et a1.(1990), t h e particulate opal is not substantially dissolved during sinking through t h e water column b u t dissolved on t h e bottom. T h e atomic ratio of Si/O in the regeneration (or loss) term, P, is 0.67 in t h e Japan Sea, whereas t h a t in t h e Pacific deep water is 0.39 (Tsunogai, 1972). This enrichment in silica in t h e J a p a n Sea may be d u e to a larger input from t h e Asian continent via t h e atmosphere and rivers, and may be related to large fragile radiolarians found in t h e sediment traps (Noriki, personal communication). T h e apparent C-14 age of t h e Japan Sea deep water (A14C = -74 %), is 300 yrs, which is larger than t h e turnover time for t h e vertical mixing obtained in this study, Table 1. Concentrations of CFCs in t h e various reservoirs calculated from the estimated water exchange rates(k) for the case of T = 0.5 by assuming t h a t t h e atmospheric input rate is proportional to t h e difference from t h e equilibrium concentration and the C m values are equal to observed ones.

F-11 (pM) Calc. (Obs.) Warm surface Om

Cw

F-12 (pM) Calc. (Obs.1

(2.24) (1.84)

-

-

-

(1.10) (0.95)

1.90

( -)

1.07

( -)

0.72

(0.72)

0.42

(0.42)

0.05

(0.10)

0.03

(0.06)

-

Cold surface

CC Intermediate

cm

Deep

Cd

117 100 yrs, b u t shorter than its residence time within t h e J a p a n Sea, 990 yrs. T h e discrepancy between t h e apparent C-14 age and t h e turnover time is d u e to t h e delay in the gas exchange equilibrium at t h e surface for t h e carbonate system. If t h e li' values obtained here is applied to t h e carbonate system, we get t h e gas exchange rate or t h e gas transfer velocity of carbon dioxide to be 1.5 m/day (Fig. 9) almost regardless of t h e T value. This is about a half of t h e mean value of t h e world oceans, 3.3 m/day (Broecker and Peng, 1982). T h e small gas transfer velocity seems to be caused by the lower sea surface roughness of the northern Japan Sea as observed by t h e GEOSAT altimeter, although t h e coastal region along Japan is fairly rough in winter (Ebuchi e t al., 1992). We have also calculated t h e concentrations of CFCs in t h e cold surface ( C c ) and t h e deep (Cd) reservoirs, using t h e observed concentration of CFCs in t h e intermediate water ( C m ) t,h e estimated zi' values and a n assumption t h a t the input rates of t h e atmospheric CFCs are proportional to their atmospheric concentrations of Smethie et al. (1988). T h e calculation method is similar to t h a t of Watanabe et al. (1991) for tritium except their Eq.(9) mhich should be modified as follows; Cc(t+At)

= CC(t) + at[lil(C,(,)C,,t,,,/~~(,),,- C C ( t ) ) li'2(Cm(t) - G ( t ) ) + k ( C c ( t ) e g- CC(t))SCl/VC,

+

(7)

where C,(t)eqand Cc(t)eq are t h e equilibrium (satura.ted) concentrations at 20°C and 0°C with t h e atmosphere at the time, 1, and k is t h e apparent gas exchange coefficient in this case. T h e calculated concentrations of CFCs in t.he cold surfa.ce, C, (Table 1)) coincide fairly well with those observed in t h e warm surface, C,, including t h e water of 100-200 m depth and 1-2°C. T h e slightly larger observed concentrations in t h e deep reservoir may be d u e to the cont,amination during the sampling and chemical analysis. T h e calculated net input rates in 1987 in t h e Japan Sea are 9 ~ 1 0 - ~ ~ for F-12, which are smaller than rnoles/m2/yr for F-11 and 5 ~ 1 0 -moles/m2/yr 11x10-* moles/m2/yr for F-11 in t h e 1970 level calculated by Liss and Slater (1974). T h e calculated gas exchange coefficients k, are only 20 m/yr for F-11 and 30 m/yr for F-12, which are 20-30 times smaller than t hat estimated from C-14, which is 1.5 m/day. Of course t h e calcula.ted gas exchange coefficient is apparent, because the degree of equilibration of CFCs in t h e surface mixing layer j u s t below the surface film should be larger than t h a t i n t h e bulk surface of 200 m thick for the sparingly soluble CFCs, and t h e gas excha.nge should also occur in t h e warmer seasons a t the Lower solubilities of CFCs (i.e., the concentration difference between the atmosphere and t h e surface wa.ter is small). Indeed t h e concentration of CFCs i n the summer surface water (Table 1) is almost equilibrated with t h e atmosphere at the ambient temperature. Owing to t h e lack of d a t a on CFCs in northern Ja.pan Sea in winter, howeverjt is difficult to explain these differences although t h e smaller input rate may be partly d u e to the small a.ir-sea ga.s exchange coefficient in t h e Japan Sea.

118 Acknowledgement We would like to acknowledge valuable discussion with Dr. Shinichiro Noriki a n d the m e m b e r s of t h e Analytical Chemistry Laboratory, Faculty of Fisheries, Hokkaido University. We also t h a n k Ms. Keiko Moriya for her kind cooperation in making t h e manuscript. References Brewer, P. G., Y. Nozaki, D. W. Spencer, and A. P. Fleer, 1980. Sediment trap experiments in the deep North Atlantic: isotopic and elemental flux. J . Mar. Res., 38, 703-728. Broecker, W. S., and T. -H. Peng, 1982. Tracers in the Sea, Eldigio Press, Columbia Univ., New York, 690 pp. Fukuoka, J., 1965. Hydrography of the adjacent sea (1) - T h e circulation of the Japan Sea. J. Oceanogr. SOC.Japan, 21, 95-102. Ebuchi, N . , H. Kawamura, and Y. Toba, 1992. Growth of wind waves with fetch observed by the GEOSAT altimeter in the Japan Sea under winter monsoon. J . Geophys. Res., 97, 809-819. Gamo, T., and Y. Horibe, 1983. Abyssal circulation in the Japan Sea. J. Oceanogr. Soc. Japan, 39, 220-230. Harada, K., and S. Tsunogai, 1986. "'Ra in the Japan Sea and the residence time of the Japan Sea water. Earth Planet. Sci. Lett., 77, 236-244. Hata, K., 1962. Seasonal variation of the volume transport in the northern part of the Japan Sea. J . Oceanogr. SOC.Japan, 20th Anniversary Volume, 168-179. Honjo, S., 1980. Material fluxes and modes of sedimentation in the mesopelagic and bathypelagic zones. J . Mar. Res., 38, 53-97. Horibe, Y. (ed.), 1981. Preliminary Report of the Hakuho Maru Cruise KH-77-3 (Pegasus Expedition), Ocean Res. Inst., Univ. Tokyo, 55 pp. Liss, P. S., and P. G . Slater, 1974. Flux of gases across the air-sea interface. Nature (London), 247, 181-184. Masuzawa, T., Noriki, S., T . Kurosaki, S. Tsunogai, and M. Koyama, 1989. Compositional change of settling particles with water depth in the Japan Sea. Mar. Chem., 27, 61-78. Menard, H. W., and S. M. Smith, 1966. Hypsometry of ocean basin provinces. J. Geophys. Res., 71, 4305-4325. Smethie, Jr., W. M., D. W. Chipman, J . H. Swift, and K. P. Kolterniann, 1988. Chlorofluoromethanes in the Arctic Mediterranean seas: evidence for formation of bottom water in the Eurasian Basin and deep-water exchange through Fram Strait. Deep-sea Res., 35, 347-369. Tsunogai, S., 1972. An estimate of the rate of decomposition of organic matter in the deep water of the Pacific Ocean. In: Biological Oceanography of the North Pacific Ocean, A. Y. Takenouti (ed.), Idemitsu Shoten, Tokyo, pp. 517-533. Tsunogai, S., 1987. Deep-water circulation in the North Pacific deduced from Si-0 diagrams. J . Oceanogr. SOC.Japan, 43, 77-87. Tsunogai, S., and S. Noriki, 1987. Organic matter fluxes and the sites of oxygen consumption in deep water. Deep-sea Res., 34, 755-767. Tsunogai, S., S. Noriki, K . Harada, and K. Tate, 1990. Vertical-change index for the particulate transport of chemical and isotopic components in the ocean. Geochem. J., 24, 229-243. Uda, M., 1934. Hydrographical studies based on simultaneous oceanographic surveys made in the Japan Sea and in its adjacent waters during May and June, 1932. Rec.

119 Oceanogr. Works Japan., 6, 19--107. Watanabe, S., S. Saito, and S. Tsunogai, 1988. Chlorofluoromethanes in the Japan Sea. Proc. 1988 Fall Meeting of Oceanogr. SOC.Japan, 285. Watanabe, S., M. Nakajima, and S. Tsunogai, 1989. Carbon dioxide exchange rate in the Japan Sea estimated from radiocarbon. Proc. 1989 Annual Meeting of Geochem. SOC. Japan, 47. Watanabe, Y . W . , S. Watanabe, and S. Tsunogai, 1991. Tritium in the Japan Sea and the renewal time of the Japan Sea deep water. Mar. Chem., 34, 97-108.

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved

121

Instrumental Development for Measurement of Phosphate in Seawater and Some Discussion of Nutrient Distributions in the North Pacific Kitao FUJIWARA a n d Hiroyuki TSUBOTA*

Abstract

Several methods for the measurement of phosphate in sea water are presented based on the different principles. For classical colorimetry, a rapid automated method based on flow injection analysis is proposed where a mirrorcoated long-path cell was applied. Detection limit, of this system is about 0.05 p M . Conversion of phosphate to phosphine is also developed and coupled with gas-chromatography or ozone gas-phase chemiluiriinescence for total or dissolved total phosphorus measurement. Besides the analytical methods for phosphorus, the relationship between phosphate and nitrate (+ nitrite) in the western north Pacific is discussed.

1. Introduction Nutrients are t h e most important species for describing t h e chemical nature and status of seawater. They link t h e biological environment a n d dynamics of t h e ocean. In another words, t h e measurement of nutrients are indispensable in chemical oceanography. In general, nutrients are measured by colorimetry; however, t h e methods currently adopted d o not necessarily provide sufficient sensitivity for deternlining them accurately in seawater. Accordingly, for describing minute chemical structure in the deep sea, a more precise method is required. Bubble-segmentation system was introduced into t h e majority of t h e commercial autoriiated a n a l y x r s such as Technicon’s. However, t h e bubble-segnientatiori has some disadvantage as a ship-board technique: slow speed of analysis a n d unstableness against t h e niovement of ship. For t h e discussion of minute structure in the nutrient distribution, the present method lacks t h e precision in speed a n d operation of measurement. For overcoming this problem, a n application of flow injection analysis (FIA) may be effective, which is faster in measurement speed a n d of which stability increases during operation compared t o t h e analyzer based on bubble-segmentation. However, the colorimetric detection is made before reaching t h e stationary s t a t e in analytical reactions a n d diffusion of t h e solutions is serious. which means t h a t improvement in the sensitivity is required for colorimetric detector of FIA. Besides t h e colorimetric detection of nutrients by a FIA with enhancing detection power, t h e method based on a completely different chemical principle is *Facultyof Integrated A r t s and Sciences, Hiroshima University, Higashisenda-rnachi, Naka-ku, Hiroshima 730, Japan

122

mirrored kurface

: incident (sour'ce) 1 i g h t - - - - - - - - t ;reflected (transmitted) light

P

Fig. 1. Schematic illustration of the mirror-coated long-path cell with the bifurcated optical guide.

Table 1. Instruments used in this study. Light source Laser diodes

Sharp, type LTD030MD 3 niw (750 nm) 30 mw (830 nm) : Sharp, type 5HD-23 : Machida denshi, type LBS-750 :

LED Power source Detector (silicon photodiode) : Preamplifier : : Electrometer Optical guide : Flow injection system Analyzer : : X-Y Sampler Commercial colorimetric

Hamamatsu, type S780-5BQ-3L N F Electronic, type LI-76 Advantest, type TR-8652 Koyo, type EK-3036

Hitachi, type K-1000 Hitachi, type 105XY detector : Jasco, type 870-UV (for HPLC) Spectrophotometer for reference measurement of absorption : Shimadzu, type UV-120-01 Computer used in the flow injection analysis Computer : NEC, type PC-8801 Printer , : Epson, type FP-80

proposed in this paper because the analytical values obtained by the double or triple independent methods are sometimes indispensable for improving the reliability of the results.

2.

Developments of Analytical Methods for Phosphorus in Seawater

123

LED, LD

m-,v

PHOTO 7 DIODE

1

OPTtCAL

GUIDE

pzinting aluminum powder

N0.5

-20-

Fig. 2. Designs of the mirror-coated long-path cells.

2.1.

Highly sensitive colorimetric detector for flow injection analysis

Nutrient species are generally determined by colorimetric methods. Several automated systems have been developed for routine measurement of nutrients in seawater, and we adopted here a flow injection method. The method has several merits especially in field operations such as on board research vessels, that it is simple and rapid. One of its drawbacks is the low sensitivity because the detection was done before reaching a maximum of color-development. The dilution of the solution is another serious problem here. These weakness may be overcome by the enhancement of detection power of flow injection analysis. In light absorption spectrometry, extension of optical pathlength is one of the effective ways t o increase sensitivity. In our previous papers (Fuwa et al., 1984; Fujiwara and Fuwa, 1985; Fujiwara, 1988), we have discussed how a capillary com-

124

OPTICAL GUlOE

unnin

Fig. 3 . Illustration of flow injection system for phosphate nieasurement with a mirror-coated long-path cell. mercially sold for gas chromatography was transfornied to a colorirnet,ric cell for this purpose. I t was possible t o design cells of as long as 50 m for colorimetry. through which the light is transmitted in any geometric configuration by total reflection a t the inner cell wall. However, this tot>alreflection cell cannot, be applicd t o the nutrient measurements because water has low refractive index. We, therefore, developed a new type of linear long-path cell, in wliich introduction of source light and detection of reflected light in the cell are carried out. in the same direction by using a bifurcated optical guide. The source light is introduced froin one end of the flow cell, transmitted in the sample solution with reflecting a t the outer cell wall, and then collectd a t the same source-light inlet (see Fig. 1). Using this type of cell it is possible t o collect the reflected light passed via various pathways inside the cell. This mirror-coated long-path cell is used as the colorinietric cell in our flow injection system (Fujiwara et al., 1990). T h e instriiinents employed are listed in Table 1. The mirror-coated long-path cells of various designs are shown in Fig. 2. Five types of cells (hereafter referred t o as No. 1-No. 5 according t o the figure) were constructed. Cell No. 1 is a Pyrex tube with i.d. of 3 iiim and 0.d. of 5 mm and its outer cell wall is painted with aluminum powder. T h e outer walls of the other cells (No. 2-No. 5 ) were silver-coated by the ordinary met,liod. For the protection of the silver-coating, a cylindrical glass is placed over tlic cell. No. 2 and 3 cells are the same except for the orientation. Cell No. 3 was originally made for turbidity measurement, in which t,he suspended particles often settle downward, causing a drift of the signal in such a case. T h e leak of solut,ion a t the connection of the light, guide (optical fibers) and the cell was prevented by packing with a black polyethylene tube. T h e aperture of the light guide at the cell side is 3.00 m n ,and at the light source/detector side is 2.16 mm. In this light guide, the bundle of optical fibers (the 0.d. of each fiber is 50 um) from the cell side was divided into two branches, each of which was connected t o either the light source or the detector, thus creating bifurcation. The

125

5 0 " Z

P/rnL

8 rnin.

Fig. 4. Example of flow injection signals for phosphate measurement. The signal as printed through the personal computer is shown. optical fibers from the light source and the detector sides are randomly mixed in the cell side. The detection system is shown in Fig. 3. The flow injection analyzer was connected to the mirror-coated long-path cell. The absorbance of the solution was calculated by a personal computer (instead of by a log amplifier), t o which the signal from an electrometer was sent via GP-IB. Laser diodes a t 750 nm ( 3 mW) and 830 nm (30 mW) operated in a cw mode were used as the light source. When the present mirror-coated long-path cells were used, the sample was introduced into the flow injection analyzer using an X-Y auto-sampler so t h a t u p to 100 samples can be sequentially measured. Since generation of the bubbles in the carrier disturbed the stability in the signal, degassing of the carrier was done for 2 h. For phosphate measurement, the color-developing reagent was made by mixing 14% (v/v) sulfuric acid (150 mL) 0.275% (w/v) antimony potassium tartarate (45 niL) 4% (w/v) ammonium molybdate (45 mL) in this order, with 1.76% (w/v) ascorbic acid (90 mL) and ethanol (50 niL) being added just before measurement. When measurement was made without ethanol, baseline gradually shifted. Ethanol was effective in diminishing this shift. It has been suggested that an addition of some surfactants such as SDS or Triton X-100 is helpful for stabilization of the baseline of the signal in various auto-analyzing systems coupled with CL coloriinetric detector. However, addition of these reagents made no difference in our system. The flow injection analyzer was a sandwich type, where the sample (30 pL) is sandwiched between the color developing agent (100 pL). The flow line of the carrier was heated at about 130 C for 3 m and then introduced into a mixing coil of 2-m length before going t o the detector (mirror-coated long-path cell). The 2-m length coil was also attached t o the outlet of the detector in order to create back pressure which prevent bubble formation. The flow rate of the carrier was about 2.0 mL/min. Absorbance observed by the present long-path cell is compared with that obtained by an ordinary cell of 1-cm length. Using cell No. 4, an enhancement of about 27 times in the absorbance was attained (in the range of less than 0.05 in the ordinary cell although the cell length is only 15 times longer than t h a t of the ordinary cell. Cell No. 5, which has the same cell length as t h a t of cell No. 4, provides poor

+

+

126 enhancement in absorption than did No. 4 cell. Cell No. 2 gave 16 X enhancement in the linear range of the calibration curve. On the other hand, the absorbance obtained by cell No. 1 was in 1 : 1 ratio to that obtained by the ordinary cell. The present system was used as the detector for flow injection analysis of phosphate. The source light used here was a laser diode at 750 nm emission. The laser diode is an effective tool as a light source for detecting molybdenum blue colorimetric detection of phosphate ion from the viewpoint of its stability and intensity. Fig. 4 shows the typical results obtained by the present system with cell No. 4 under the condition that the flow rate of the carrier was 1.53 mL/min. and the solution was heated at once at 130°C. In this case, 1 ng P/mL (about 0.03 pmol/L) can be measurable according to the computation of signal averaging, and 3 ng P/mL (about 0.1 pmol/L) is detectable by the direct reading of the recorder chart of Fig. 4. The detection power of the present system is sufficient for almost all the seawater samples including nutrient-depleted surface waters. Thus, introduction of the present mirror-coated long-path cell in flow injection analysis is quite useful in the field of nutrient geochemistry phytoplankton growth. In practice, the present system was operated during the R.V. Tansei-Maru cruise (KT 88-19, Nov. 8-11, 1988), and its feasibility and handling performance was confirmed. 2.2.

Measurement of phosphate by the methods associated with the phosphine generation technique

Here, the method of phosphate determination using the phosphine generation technique will be discussed. It is possible to apply various spectroscopic methods for measuring phosphine which exists in gas-phase at the room temperature because of the boiling point of -87.7"C. Since the methods for measuring phosphate in water sample are rather limited, the development of the method involving quantitative conversion from phosphate to phosphine may be useful for comparison. However, the major problem is the difficulty in the reduction of phosphate, because the redox potential of phosphate to phosphine is lower than that of water: For the reactions, P 3H20 3e = PH3 3 0 H - and 2Hz0 2e = Hz 20H-, -0.87 and -0.8277 V vs SHE, respectively, are given. This means that the generation of phosphine from phosphate is difficult under the presence of water. We have developed the procedure of quantitative generation of phosphine by the use of sodium tetrahydroborate as the reducing agent, which is popularly utilized in the hydride generation of metalloide elements such as As, Se, Sb, and Sn in atomic spectrometries. The procedure is outlined in Fig. 5 , where the sample phosphate is mixed with aqueous solution of sodium tetrahydroborate in a quartz vessel shown in Fig. 6 and is dried at about 40°C. The quantitative generation of phosphine was possible when this mixture was heated t o 460°C. This phosphine generation method expands the analytical methods available for phosphate. For example, gas-chromatography (Hashimoto et al., 1987), ozone gas-phase chemiluminescence detection (Fujiwara et al., 1989a), and ICP atomic emission spectrometry (Fujiwara et al., 1989b) can be applied to the phosphate analysis in combination with this phosphine generation technique in this report. Fig. 7 shows the distribution of the total and total dissolved phosphorus in the

+

+

+

+

+

Table 2. Determination of Phosphorus in Various Compounds"?*. method KH2P04 DTPS" ADP AMP IMP phytic acid molybdenum blue colorimetry 100f2.2 100f3.1 100f3.3 10052.6 100k3.4 100f3.4 with K2S2O8 digestion present method with 99.7f2.4 98.3f3.4 101.5f2.9 99.4f3.2 99.6f3.3 97.8f3.1 K2S2O8 digestion 101.4f2.3 101.lf2.7 98.6f3.4 98.4f2.6 99.5f3.2 68.9f4.8 present method " Sample amount was 100 ng of P/mL, 100 p L . Average value for five repetitive measurements. trans-1,2-Diphenyl-l,2,3,6-tetrahydro-1,2-diphosphorin 1,2-disulfide.

'

128 samp1e +mixing with NaBI-i4 and drying bmeasurement of TP by hydride-generation method ,digestion

with IC2s208

Leasurement of TP by molybdenum-blue colorimetry

+filtration with 0.2-pm Nuclepore filter

lL

mixing with NaBI14 and drying easurement of TDP by hydride-generation method

digestion with K2S208 Lmeasurement of TDP by molybdenum-blue colorimetry

Fig. 5. Procedure for phosphorus measurement. TP: total phosphorus, TDP: total dissolved phosphorus.

IOcm Fig. 6. Designs of sample container for generation of phosphine from phosphate. a type: half cylindrical, b type: boat. Japan Trench obtained by the phosphine generation-gas chromatography method.

It should be noted that the results obtained by the method correspond t o the concentration of the total phosphorus compounds in the sample, and do not agree with orthophosphate (P043-). The difference between the phosphine generation method and the ordinary molybdenum blue colorimetry can be attributable t o the phosphorus compounds other than orthophosphate. Table 2 shows the recoveries for various phosphorus compounds based on the phosphine generation-gas chromatography and potassium persulfate digestion-molybdenum blue colorimetry. Except for phitic acid, both the results show good agreement with each other. Gas-phase chernallurnanescence detectaon of phosphate

It is well-known that phosphine naturally burns in the atmosphere and emits visible light, which is called igunus flatuus or will’o-the-wisp. However, this luminescence was not well elucidated. Before establishing the gas-phase chemiluminescence measurement of phosphate, chemiluminescence spectra occurring between phosphine and ozone were observed. A silent discharge type ozone generator was used at the applied electric voltage of 100 V with flowing oxygen at the rate of

129 phosphorus (yrnol liter")

hydride-generation method 0: total phosphorus 0: t o t a l d i s s o l v e d phosphorus

90 01

10000~

Fig. 7. Distributions of dissolved total and total phosphorus at at the Japan Trench (29"05'N lat, 142'51'E long). 100 mL/min. The concentration of ozone was 13 g/m3 under this condition. For conversion of phosphate in the sample to phosphine, the sample was taken in a quartz vessel (either boat or half cylindrical types, see Fig. 6). A boat-type container is more convenient for retaining t,he sample and sodium tetrahydroborate solutions than the half cylindrical one. The generated phosphine was mixed with ozone in a chemiluminescence chamber. The chemiluminescence chambers of various design were made and used in this experiment. For quantitative analysis of phosphate, a photon counting system with a cooled photomultiplier was adopted where no spectral separation was carried out due to the small intensity of chemiluminescence emission. The analog output of the photon counter was recorded on a strip chart recorder. The equipments used are listed in Table 3. The recommended procedure for determination of phosphate is as follows: 102000 pL of the sample phosphate solution is placed in the quartz vessel and completely dried. Then, 100 ILLof 6% sodium tetrahydroborate is added t o the residue and dried again a t a temperature below 45°C in an oven for 2 h. This dried mixture is inserted into a phosphine generation tube made of quartz and then heated t o

130

810

650 570 WAVELENGTH, n m

730

490

b

800

700

600 WAVELENGTH

I

500

400

nm

Fig. 8. Chemiluminescence spectra. a: Spectrum for high concentration of phosphine (about 10%). b: Spectrum for low concentration of phosphine (108 ppm).

Fig. 9. Schematic diagram of phosphine generation and chemiluminescence detection systems: 02, oxygen cylinder; He, helium cylinder; (1)gas flow meter, (2) cooling water, (3) ozone generator, (4) chemiluminescence chamber, ( 5 ) aluminum case, (6) slide transformer, (7) phosphine generation tube, (8) digital thermometer, (9) heater, (10) phosphine trap wounded ny a Nichrom wire which heats trap to expel the condensed moisture, (11) photomultiplier housing, (12) cooling unit, (13) photon counter, (14) recorder. 480-500°C by a cylindrical Nichrom heater. The temperature is monitored and controlled by a digital thermometer. Helium is used as a carrier gas of the generated phosphine and continuously circulated

131 Table 3. Instruments used in this work. : Nippon Ozone Co., Type 0-3-2 Ozone generator Spectrometer : Jasco, Type CT-25-C Multichannel detector PC D image sensor : Hamamatsu, Type S2304-512Q Detector Type C2327 Dat a processing unit : Type C2890 DMA interface Type M2891 : NEC, PC9801 VX Personal Computer Printer PC-PR201F2 Scope : Trio, Type CS1577(30 MHz) Fourier transformation and general spectral observation Analyzer(Recorder) : Yokogawa, Type 36553 FFT module : Yokogawa, Model 3659 20A DC amplifier : Kiethley, Type 427 : N F Electronic, Type LI-76 DC(A/V) amplifier Function generator : N F Electronic, Model FG-121B PM power source : HTVC 488R PM(photomultip1ier) : Hamamatsu, R456 Determination of phosphate : Mitamura Riken, 400 W Electric furnace Digital thermometer : Nippo Electric, Type LM499 P M housing with electronic cooler : Hamamatsu, Type C659-A PM Type R649 Photon counter Type (22130 : Hitachi, Model 056 Strip chart recorder

through the phosphine generation tube at a flow rate 600 mL/min. Phosphine is trapped in a U-tube packed with a small amount of quartz wool by cooling it with liquid nitrogen for 2 min. After collecting phosphine, the cold trap is soaked (warmed) in a water bath. The gaseous phosphine is, then, introduced into the chemiluminescence chamber and mixed with ozone supplied from a generator of silent discharge type through which oxygen flows at a rate of 150 mL/min. T h e peak height of chemiluminescence signal fed into a strip chart recorder was read for calculating phosphate concentration. When phosphine at high concentration was exposed t o air, the color of emission changed from white to red and then reversed from red t o white before termination. Occasionally, a greenish emission appeared, which is presumably due to H P O or (PO), . This fact suggests t h a t the chemiluminescence spectrum changes according to the concentration of phosphine. Fig. 8a shows the chemiluminescence spectra given by phosphine at about 1015% (v/v) concentration. A broad peak of emission appeared at 670 nm with

132

Exhaust

N0.3 Exhaust

4

03 10 C r n

Fig. 10. Schematic illustration of the chemiluminescence chambers.

D i o d e

I

I

I

A r r a y

HRlOOO

J-Y

Mono.

H10

I a

V a l v e

# Z

a Col d

R e a c t i o n

r

Tr up

Fig. 11. Schematic diagram of optical system and hydride generation sample introduction system for ICP-AES.

a shoulder at 740 nm. This emission looked orange. Both oxygen and ozone can make luminescence at this concentration of phosphine. Fig. 8b shows the

133

I ~ G O " E

180'

160'W

Fig. 12. Sampling stations in the western North Pacific. Solid circle, open circle and open triangle are for the cruise of KH 80-2 (CY), KH 82-1 (CE) and KH 85-4 (DE), respectively. emission spectrum given by phosphine at the concentration of 103 ppm. Only ozone generated this emission and oxygen did not. The broad spectrum extending from 350-700 nm was obtained. Neither the sharp peaks in the visible region nor those due to P O at around 280 nm were observed. The small peak at 650 nm is the same as the one in Fig. 8a. The system illustrated in Fig. 9 was constructed for phosphate determination. The chemiluminescence chambers of type 1-3 in Fig. 10 were compared in regard t o the quantitativeness of the measurement. In chamber No. 1, phosphine was injected into the ozone flow. In chambers No. 2 and 3, the flows of carrier/phosphine and ozone were merged in each other from the opposite side. The chemiluminescence chamber No. 1 generated slightly stronger emission than chamber No. 3. The shorter chamber (No. 2) gave about 40% weaker chemiluminescence. Both No. 1 and 3 chambers give the same detection limits of 1 ng, and linear dynamic ranges are from 2 t o 5000 ng for chamber No. 1, and from 10 t o 5000 ng for chamber No. 3. Chamber No. 2 gave a detection limit of 2 ng with a linear dynamic range of 5-1000 ng. The narrow dynamic range of chamber No. 2 can be explained by its short length between the entrance and exhaust of gases. Phosphane generataon and ats detectaon by anductavely coupled plasma-atomac emassaon spectroscopy (ICP-AES) The application of ICP-AES will be briefly mentioned for phosphate analysis below. In the system for observing phosphorus atomic emission in the ICP-AES with phosphine generation, helium continuously flows into the I C P for sustaining its operation. Therefore, two lines were constructed:one is the pathway through the hydride generation tube (sample) and another is a bypass.

134

50 -

40

CE-13

-

d P e 0

30-

0 20 -

10-

0

o h

I

1

I

2

I

3

The system is shown in Fig. 11. The phosphiiie generation technique is the same as previously described. The detection limit was the same as one without hydride generation (direct introduction of sample solution into the ICP). This is due to the combined effect of the high efficiency of phosphorus atomization and the about 50% conversion efficiency. Thus, the merit of hydride generation is canceled in this case. The advantage of the hydride generation-ICP method is in the separation of phosphorus from the sample matrix, and hence the huge interference of NaC1, if any, in ICP-AES for phosphorus can be eliminated. Otherwise, seawater sample can not be determined directly by ICP-AES. The detailed results were shown in our paper (Fujiwara et al., 1989b).

135 Table 4 The slopes (dNO3/dPO4) in the whole water column (W), the upper layer (U) and the deeper layer (D) in the North Pacific. Station CY-3 CY-4 CY-5 CY-6 CY-8 CY-9 CY-10 CY-11 CY-13 CE- 1 CE-2 CE-3 CE-4 CE-5 CE-6 CE-8 CE-9 CE-10 CE-11 CE-12 CE- 13 CE-14 CE-15 CE-16 DE-2 DE-4 DE-7 DE-8

3.

The slopes W 14.1 14.8 15.0 14.6 14.9 14.4 14.6 14.5 14.4 13.9 14.2 14.0 14.8 13.0 14.0 14.0 14.1 13.2 13.4 13.5 15.3 14.3 14.9 14.8 23.4 24.2 14.4 14.0

(dN03/dP04) D U 13.6 10.5 13.8 11.8 14.4 12.3 12.2 14.1 14.4 13.2 14.1 12.5 14.4 12.9 12.9 14.0 14.0 13.0 13.6 12.5 13.7 11.5 13.7 12.3 13.8 12.0 12.3 9.4 13.4 12.1 13.5 11.3 13.4 12.4 12.5 10.5 14.2 12.9 6.2 12.2 6.4 14.0 13.8 13.9 11.7 15.4 14.1 11.2 23.0 19.5 23.6 17.1 14.0 11.0 13.4 9.2

Discussion on the Distribution of Nutrients in the Western North Pacific

Phosphate, nitrate, and silicate concentrations have been obtained a t various locations in the North Pacific (Horibe, 1980, 1982). The stations are shown in Fig. 12. The stations CE-5 and CE-15 are located in the mid-North Pacific gyre. The stations DE-4 and CY-5 are located in the subarct,icregion, where the nutrients are not entirely depleted in the surface layer. The slopes of the linear lines are less in the deeper layer than in the upper layer at all the stations in the North Pacific, as is shown in Fig. 13. These features are

136 considered t o be reflected from dissolution and/or adsorption of nutrients from/to settling particles and sea floor. In the dissolution stage in shallower layers, appearance of nitrate exceeds than t h a t of phosphate. In contrast t o the shallower layer, the decreasing trend of phosphate concentration towards the sea floor in the deeper layer is more distinctive than t h a t of nitrate, i.e., in the map presenting the relationship of nitrate and phosphate, d N / d P becomes smaller in the deeper layer below the point t h a t gave the highest concentrations of phosphate and nitrate. Larger and smaller d N 0 3 / d P 0 4 are commonly found for the upper and deeper water columns, respectively, although the degrees in this difference were found t o be variable for the stations. At some stations, the slopes ( d N 0 3 / d P 0 4 ) are rather close t o each other for upper and deeper layers, while clear separations can be found between the upper and deeper layers at other stations. For example, at station CE-13, the deeper layer is clearly separated from t h a t of the upper layer; however, difference of the slopes between them is a little at station DE-4. T h e slopes in the whole water column, the upper and deeper layers are shown in Table 4. I t is evident from the table t h a t at all the stations phosphate is more depleted than nitrate in the deeper layer. In our previous paper (Hayase e t al., 1988), it is stated that phosphate, nitrate, and silicate are more enriched than fluorescent organic matter in the deep layer. Those features may be related t o phosphate-nitrate diagrams with regard t o dissolution of settling particles. Further improvements in regard t o analytical speed and precision would be required for the nutrient measurements. This will help further interpretations regarding the minute structure of the distributions of nutrients in seawater.

References Fujiwara, K., and K. Fuwa, 1985. Liquid core optical fiber total reflection cell as a colorimetric detector for flow injection analysis. Anal. Chem., 57, 1012-1016. Fujiwara, K., 1988. Application of a wave-guide capillary cell in the determination of copper by flow injection analysis. Anal. Chim. Acta, 212, 245--251. Fujiwara, K., T. Kanchi, S. Tsumura, and T. Kumamaru, 1989a. Phosphine-ozone gasphase chemiluminescence for determination of phosphate. Anal. Chem., 61, 26992703. Fujiwara, K., M. A. Mignardi, G. Petrucci, B. W. Smith, J. D. Winefordner, 198913. Determination of phosphorus by ICP-AES using solid-phase hydride generation. Spectrosc. Lett., 22., 1125-1140. Fujiwara, K., T. Nakamura, T. Kashima, H. Tsubota, T. Solin, M. Aihara, and M. Kiboku, 1990. Mirror-coated long-path cells for colorimetry with the use of a bifurcated optical guide. Appl. Spectrosc., 44, 1084-1088. Fuwa, K., L. Wei, and K. Fujiwara, 1984. Colorimetry with a total-reflection long capillary cell. Anal. Chem., 56, 1640-1644. Hashimoto, S., K. Fujiwara, and K. Fuwa, 1987. Determination of phosphorus in natural water using hydride generation and gas chromatography. Limnol. Oceanogr., 32, 729735. Hayase, K., M. Yamamoto, I. Nakazawa, and H. Tsubota, 1987. Behavior of natural fluorescence in Sagami Bay and Tokyo Bay, Japan -- vertical and lateral distributions. Mar. Chem., 20, 265-276. Hayase, K., H. Tubota, I. Sunada, S. Goda, and H. Yamazaki, 1988. Vertical distribution

137

of fluorescent organic matter in the North Pacific. Mar. Chem., 25, 373--381. Horibe, Y. (ed.), 1981, 1982. Preliminary Report of the Hakuho Maru Cruise KH 802 (CYGNUS Expedition) and K H 82-1 (CEPHEUS Expedition). Ocean Research Institute, Univ. of Tokyo.

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

139

Actinium-2 27: A Steady State Tracer for the Deep-sea Basin-wide Circulation and Mixing Studies Yoshiyuki NOZAKI*

Abstract The presence of excess "'Ac (half-life, 21.8 years) relative to the parent 231Pain deep ocean waters indicates its usefulness as a novel tracer of basin-wide circulation and mixing studies on the time scales of less than 100 years. For this purpose, 227Acis probably best utilized by coupling with "*Ra (half-life, 5.75 years), both of which have their significant source in the bottom sediments. Most important is to develop a new method for simultaneous and accurate determination of 227Acand '"Ra in seawater. This may be achieved by combination of MnO2-fiber extraction and subsequent a-scintillation counting of '"Rn and '19Rn emanated from the fiber. New data on the vertical profiles of 227Acand '"Th (as an indicator of "'Ra) obtained in the deepest part of the Izu-Ogasawara Trench down to 9,700 m suggest that they are predominantly governed by isopycnal mixing. Within the trench, both nuclides are nearly constant with depth, suggesting that the water is uniformly mixed. This view is consistent with the distributions of other chemical species and radioisotopes whose concentrations are almost identical to those of the deep water at 5000-6000 m depths around this region.

1.

Introduction

Among the soluble members of U / T h decay series nuclides, 222Rnand 228Ra, having their significant source in the bottom sediments, have been exploited as tracers for the study of mixing and exchange processes of ocean waters. Because of its short half life of 3.83 days, application of 222Rnis limited to processes occurring on time scales of less than two weeks. Hence, the extensive studies using 222Rn(e.g., Broecker, 1965; Chung and Craig, 1972) have focused on the characterization of bottom boundary layer. By contrast, 228Rawith a. half life of 5.75 years can be used in the study of physical processes of decadal time scales, such as vertical mixing and circulation of ocean basins (e.g., Moore, 1969; Sarmiento et al., 1976). 228Rais particularly important for the deep sea environments into which the anthropogenic tracers such as artificial 3H and chlorofulorocarbons have not penetrated as yet. Recently, Nozaki (1984) has found excess 227Ac over its parent 231Pain North Pacific deep waters, suggesting t h a t it can serve as a deep sea tracer as well. 2 2 7 A ~ *Ocean Research Institute, University of Tokyo, Minarnidai, Nakano-ku, Tokyo 164, J a p a n

140 has a half life of 21.8 years (see Fig. l ) ,approximately four times longer than that of "'Ra and have a merit that its application to oceanographic problems is extended up t o -100 years. The usefulness of this relatively newborn tracer in the study of deep water circulation and mixing is not well explored as yet. In this paper, I describe the present status of knowledge on the oceanic distribution of 227Acand their geochemical and oceanographical significances as well as the development of sampling and analytical techniques.

2.

Background

Because actinium has no stable isotope, its physico-chemical nature is not well understood. Yet limited informations on the characteristics of Ac compounds (Latimer, 1952) and the radiotracer studies (Karalova and Myasoyedov, 1982) suggests that the chemical behavior of Ac resembles that of lanthanum. Until recently, the trivalent lanthanides had been thought t o belong to a group of the least soluble elements in seawater such as Fe, Al, Pa and T h whose concentrations are maintained at extremely low level by scavenging of particulate matter. Thus, if 2 2 7 A and ~ 231 Pa were virtually insoluble and were rapidly scavenged to particles, the oceanic 23sU-231Pa-227Acsystem could be used t o determine the mean residence time of marine particles in analogous t o the fact that the 222Rn-210Pb-210Bisystem in the atmosphere can be used to determine the mean residence time of aerosols (see Turekian et al., 1977). This idea was noted when the first PARFLUX Atlantic sediment trap samples were analysed radiochemically (Spencer et al., 1977). The Pa activity ratios in the Pacific particle unequivocal measurements of 227Ac/231 samples done by Anderson (1979) resulted in the values from less than 0.1 t o 0.6, and appeared to be consistent with the above idea. Unfortunately however, the situation was not so simple. Nozaki (1984) found that 227Acis present in excess over 231Pain Pacific deep waters indicating that 2 2 7 Ais~ supplied from somewhere

3.43 x

J" 3.92s

J Fig. 1.

A portion of decay scheme of 23sU series chain.

141 to the deep water and not effectively removed by particle scavenging as a whole. This clearly means that the analogy of marine 235U-231Pa-227A~ chain to the atmospheric zzzRn-210Pb-210Bisystem does not hold. Instead, investigation on the usefulness of 2 2 7 Aas ~ an oceanographic tracer in the context of physical mixing and circulation has begun. This is an aspect described in more detail below.

3. Methodology Application of 2 2 7 Aas ~ an oceanographic tracer strongly depends upon the availability of t h e conventional method of accurate determination and hence methodological development, is one of the most important topics on this subject. There are several ways of determining 227Acin seawater. Not all methods have been thoroughly tested so that their advantage and disadvantage are, in part only conceptually and theoretically, described. Because 227Acis a weak @-emitter( p energy, 0.0455 MeV) and present at extremely low concentrations of < 3 dpm/1031 in seawater, it is difficult to directly count the low-level activity of 227Ac. Therefore, the methods generally include extraction of 227Acfrom large-volume of seawater ( > l o 0 l), isolation and detection of some of its daughters.

-200

L

seawater

Carriers (Fe,

Be,Pb)

ion exchange column

U

Fig. 2. A diagram showing the analyt,ical procedure for radionuclides in large-volume water samples as currently been used on board R. V. Hakuho-Maru.

142 3.1. Extraction of Ac from large-volume seawater Coprecipitation with iron hydroxide

Here I describe the chemical procedure (Fig. 2), with focus on 2 2 7 A ~which , has currently been used on board of R. V. Hakuho-Maru for simultaneous determination of various radionuclides (Th, Pa, Pu, Am, P b , Be, Sr, Cs and C isotopes) using a -250 1 water samples obtained from various depths with a twin PVC bottle sampler (Horibe and Tsubota, 1977). Soon after sampling, unfiltered seawater of 200-250 1 was placed in a closed container, acidified with 500 g of concentrated H N 0 3 (one bottle of Wako Pure Chemical Industries Ltd.). Then, Np-gas was circulated through the water and a column containing 4 M NaOH solution for several hours t o extract dissolved C o p gas for 14C measurements (Horibe et al., 1986). The remaining water was transfered into a PVC container with a funnel-shape bottom. To the water, various carriers (1 g as Fe (111), 25 mg as P b , 1 mg as BeO, 20 mg as Cs) and yield monitors (appropriate amounts of 233Pa,234Th, 242Pu and 243Am) were added and thoroughly mixed. Then, 500 g of concentrated NH40H (one bottle of Wako Chemical Industries LTD) was added with stirring to precipitate iron hydroxide together with the actinides at pH 7.5. The precipitate was allowed t o settle for more than 24 hours and then separated from most of the water through a valve at the bottom of the container. The supernatant was further treated for "Sr and 137Csaccording t o the method of Nagaya and Nakamura (1981). The iron hydoxide precipitate was further separated from water by decantation and filtration, and then redissolved in ~ 5 ml 0 of 8 M HN03. The solution was allowed to stand for more than a week, meanwhile gelatinous precipitate of amorphous silica was formed. This precipitate which contained most of Pa was separated from the solution by centrifugation and set aside for Pa analysis (Nozaki and Nakanishi, 1985). The 8 M H N 0 3 solution containing 2 2 7 Aand ~ other radionuclides was then passed through a Dowex 1x 8 anion exchange column (12 c m x 1 cm, 100-200 mesh) preconditioned with 8 M HN03. The isotopes of Ac, Be, P b and Am passed through the column, whereas U, Th, P u and Pa were adsorbed on the resin. The column was rinsed twice with 20 ml of 8 M HN03. The time of this ion exchange separation was recorded. The 8 M HN03 effuluent and rinses were combined and the iron content in the solution determined by atomic absorption spectrometry t o determine the recovery of iron hydroxide up t o this step. The solution containing 2 2 7 Awas ~ stored in a polyethylene bottle and was set aside for further treatment. The isotopes on the resin were successively eluted with aliquots (40 ml each) of 10 M HC1 for Th, 10 M HC1 0.13 M H F for Pa, 10 M HC1 NH41 for P u and 0.1 M H N 0 3 for U. These nuclides were then analysed by a and ,B spectroscopy described in Nozaki et al. (1981) for Th, Nozaki and Nakanishi (1985) for Pa and Nagaya and Nakamura (1987) for Pu. N

+

+

MnOp -fiber extration

The acrylic fibers coated with manganese dioxide (called MnOp fiber, hereafter) prepared by the method of Moore (1976) and Nozaki (1983) are very effective for extracting R a from seawater ranging in volume from 20 t o several thousand liters.

143 Reid et al. (1979) demonstrated that the MnO2-fibers are also efficient adsorbers for Ac as well. They found the extraction efficiency of 99*1% for both Ra and Ac under the flow rate of 0.5 l/min using 30 1 of surface seawater. Thus, the use of MnO2-fiber in extracting Ac from seawater greatly reduces the labor needed for sampling large-volume waters and shipboad chemical extraction. Using the submersible pumps with a MnO2-fiber column as described in Krishnaswami et al. (1976), Orr et al. (1984), Nozaki et al. (1987) and Bacon (personal communication), Ac can be extacted from seawater in situ. More convieniently, some specific devices attached t o hydrowire or mooring rope in order for the M n 0 2 fiber t o be exposed t o seawater may also be used as described in Moore and Reid (1973), Moore (1976), Anderson (1979) and Nozaki (1983). In the latter, only the isotopic composition adsorbed on the MnO2 fiber can be determined. The effective volume from which radioisotopes were extracted may be estimated if any one of the isotopes is known for its concentration in the seawater. The equilibrium concentration of 234Thwith respect to 238Ucan be used for a-emitting T h isotopes (Moore, 1981; Nozaki and Horibe, 1983). The 226Raconcentration measured in a separate aliquot of seawater is used for 228Ra (Moore and Reid, 1973). No appropriate isotopes are available for 2 2 7 Aand ~ 231Pafor this purpose. 3.2.

Alpha-spectrometry of 227Th

The Mn02-fibers are leached in hot 6 M HCl for 12 hrs t o dissolve all the nuclides and manganese in the presence of appropriate carrier and spikes (Pb, 233Pa,229Thetc.), depending upon the nuclides of interest. The white fiber was removed from solution after squeezing the solution and washing. The solution was filtered to remove fabric debris using a Millipore HA filter and then evaporated t o dryness. The residue was redissolved in -40 ml 8 M HN03 and loaded on a Dowex 1 x 8 anion exchange column preconditioned with 8 M HN03. The column was washed twice with 20 ml of 8 M HN03. The effluent and rinses were combined and set aside for 227Acanalysis. The T h and Pa isotopes adsorbed on the resin were analyzed according t o the methods given above. 2 2 7 Ain ~ the 8 M HN03 solution obtained by either iron hydroxide or MnO2fiber methods can be determined by a-spectrometry of 227Th which was milked from the 227Acsolution, as follows: The 8 M H N 0 3 containing 227Acsolution was stored for more than 5 months during which 227Th grew into equilibrium with 227Ac. To the solution, an appropriate amount of 230Th(-0.2 dpm) was added as a yield monitor, and the solution was passed through an anion exchange column preconditioned with 8 M HN03. The column was washed twice with 20 ml aliquots of 8 M HN03, and the effluent and rinses were combined and stored. The time of the ion exchange separation between 227Th and 227Acwas recorded. T h on the resin was eluted with 40 ml of 8 M HC1 and the solution was evaporated t o dryness. The T h residue was taken in one drop of 8 M H N 0 3 and a few ml of deionized H 2 0 . The pH of the solution was adjusted t o 1.5-2.0 with dilute NH40H and then T h was extracted from the solution to 3 successive 0.5 ml aliquots of 0.4 M TTA-benzene solution. The T h was plated on a stainless steel disk and counted for its a-activities by a-spectrometry.

144 I

I

(a)

MnO,

150-

100

i

iiber Sample

TA06'82 5370m 919mm Detector N o 2

-

2

In +

5 =

3

0 50 -

0 200

460

mu"--

MnO, fiber sample T A 0 6 ' 8 2 . 4965m 4494min Detector N o 2

(b) 500-

9

400-

6

-

2

r

- 5 2

0

0 200-

4

100-

200 0

250

(C) 150-

300

350

400

Seawater sample Ce-13. 5 8 7 0 m 2873min

9

Fig. 3 . Examples of a-spectrum of T h isotopes, based on Mn02fiber samples (a and b) and a large-volume seawater ( c ) . Fig. 3c was obtained by the second milk of T h fraction grown in the 8 M H N 0 3 solution containing 2 2 7 A(230Th ~ was used as a yield monitor).

145 An example of @-spectrum is shown in Fig. 3. Apparent dual peaks due to 227Thwere registered centering at the energy of -5.7 MeV and -6.0 MeV. T h e count rates from 5.70 t o 5.80 MeV and from 5.95 to 6.05 MeV corresponds t o 39% and 54% of 227Th decay, respectively. It is also clear from Fig. 3b t h a t 224Ra( a energy, 5.686 MeV) and 212Bi (a-energy 6.051 MeV) contribute to the count rate in the same a-energy region of 227Th. The contribution of 224Rais corrected by subtracting the count rate of 220Rn,from a peak around 5.7 MeV, and t h a t of '12Bi is corrected by subtracting 35% of 220Rncount rate (based on a branching ratio) from the peak around 6.0 MeV. These contributions become larger as the activity of '"Th on the counting source is higher and the time elapsed between preparation of T h counting source and counting is longer. Since 228Thwas separated from 2 2 7 A ~ by the first 8 M H N 0 3 anion exchange, the correction is minimized for the second milked T h fraction as shown in Fig. 3c. The activity of 2 2 7 Ain ~ seawater sample ( A ~ + 2 2 7 ) 0 is calculated on the basis of the equation given by AAc)(AT'h-227)meATht2 1 ATh(e-XActl - eAThtl

(ATh (AAc-227)O

=

- (ATh-227)rne-AThtZ (1 - e-AThtl -

1

11

"7

(1)

"7

where tl is the time elapsed from the first to the second anion exchange in nitrate form to separate Ac and T h , t2 is the time between the second anion exchange and measurement of T h , and f is the chemical yield of Ac. Using 227Actracer and 20 1 aliquots of surface seawater, 2 2 7 Awas ~ confirmed t o be coprecipitated with iron hydroxide quantitatively (Nozaki, 1984). I t was also ~ not retain at all on the Dowex 1 x 8 resin in 8 M H N 0 3 form, found t h a t 2 2 7 Adoes and therefore only losses of Ac during processing and recovering of iron hydoxide precipitatc need to be corrected. This correction was made by monitoring the recoveries of iron for each sample (-80%). The only assumption here is that 227Ac is quantitatively coprecipitated with iron hydroxide even in expanded volume of seawater ( ~ 2 5 1). 0 This can be explicitly tested using 227Actracer. Also if 225Ac (half life, 10.0 days) is used as a yield monitor, the assumption is no longer required. These, however, remain t o be done in the future. As an independent check, one can calculate in sitii 227Th concentration as a measure of 2 2 7 Ain ~ seawater when sampled. Based on the studies of 234Th-238U disequilibrium in the ocean (Bhat et al., 1969; Amin et al., 1974), and the 24.1 day half life of 234Th which is longer than that of 227Th,the assumption seems t o be justified for most of the ocean except for the shallow waters ( < l o 0 m). Nozaki (1984) actually took this approach for some of Pacific deep waters. In Fig. 4, the ~ obtained as above. results of in situ 227Th are compared with the 2 2 7 Avalues Although agreement between the two is good, there is a fairly large uncertainty in the value of in situ 227Th. This is because correction for 227Thdecay became large due to the time (-1 month) elapsed from water sampling and measurement of T h activities for these samples. It is a disadvantage for this method t h a t rapid chemical separation and counting of T h are necessary. In practice, this is often difficult when a number of large-volume water samples need t o be processed. It is

146 I

I

/ 0

I

1 2 "'Th (dprn/103kg)

3

Comparison of in situ 227Th activity and 2 2 7 A activity ~ Fig. 4. determined for deep water samples during the Cepheus Expedition of R. V. Hakuho-Maru (KH-82-1).

Drying Agent

Rn Cell

L

a,

z LL I C

2

Fig. 5 . A dual cell counter for simultaneous determination of '"Rn and 220Rn (After Orr, 1988).

also noticed that the presence of 228Thin the sample often requires relatively large correction for the contribution of its daughters, 224Raand 'l'Bi (See Fig. 3b). For this reason, the in situ 227Thmethod was difficult to apply t o shallower waters in which the 228Thconcentration was significantly higher than t h a t of 227 Ac. For MnO2-fiber sample obtained after natural exposure t o seawater, it is difficult t o estimate the effective volume of seawater because of the lack of natural Ac isotope with a n appropriate half-life. Therefore, Nozaki and Yang (1987), in interpreting their radionuclide data, assumed t h a t the extraction efficiency of Ac is equal t o t h a t

147 of Ra. This is based on the laboratory experiment of Reid et al. (1979), but has not been tested in natural condition. Thus, intercomparison of the various methods for "?Ac determination should help confirming the validity of assumptions. 3.3.

Alpha-scintillation counting of 219Rn

An a-scintillation counter for 222Rn,consisting of a-scintillation cell inside of which is coated with ZnS powder and a photomultiplier, has been commonly used to determine 226Rain environmental samples (Lucas, 1957; Broecker, 1965). By coupling the a-scintillation methods and the MnO2 fiber on which R a isotopes were collected, Rama et al. (1987) developed a method of measuring 224Ra(half-life, 3.6 days) via counting "'Rn (half-life, 56 seconds) and its short-lived daughter. They suggested that the system might also be used t o measure 223Ra (half-life, 11 days), a ground-daughter of 227Ac, via counting 'lgRn (half-life, 4 seconds) and its daughters. Orr (1988) further considered the possibility, theoretically, and reached to the conclusion that the two cell counting system (Fig. 5) is necessary t o resolve the activities due t o 'lgR.n and 220Rnif both are present on the M n 0 2 fiber. Upon leaving the MnOz-fiber, the radon must have a short travel time (less than a few seconds) t o reach the first cell and a long travel time (-30 seconds) t o reach the second. In this way, the first cell counts both 219Rn and 220Rn,while the second counts mostly 220Rn,thereby allowing the resolution of individual activities with two simultaneous equations. Once the method is established, the simplicity (no chemistry) and greater sensitivity of supported radon emanation will make "?Ac measurements more practical on seawater samples. There is also an advantage in such a system of measuring 223Ra (as 2 2 7 A ~and ) 224Ra (as 228Th or 228Ra) simultaneously, if the radioactive equiliribria can be assumed in the sample.

4. Marine Geochemistry of 2 2 7 A ~ 4.1. Reactivity of actinium with respect to settling particles In order to use "?Ac as a tracer of physical oceanographic processes, we must understand the extent to which the 227Acconcentration might be affected by interaction with particles. The interactions of natural radionuclides between particles and solution, so called "scavenging" in t,he deep sea is best represented as a reversible process rather than the unidirectional transfer from solution to particles as exemplified by the case of T h isotopes (Nozaki et al., 1981). Thus, the reactivity of nuclides with respect to particles can be evaluated in terms of the equilibrium concept of distribution coefficient, K D defined as, hTD =

paticulate nuclide in dpm/g of dry particles dissolved nuclide in dpm/g of seawater

(2)

For one of the least soluble nuclides, 230Th in deep seawater, approximately 20% of total 230Th is associated with particles and K n is estimated to be 2 ~ 1 0 ~ (GESAMP, 1984). Nozaki and Nakanishi (1985) demonstrated that K D for Pa is -10 times smaller than KD for Th, i.e., KD(Pa) = 2x106 in the North Pacific. This implies that only a few percent of 231Pais in particulate form.

148 The only available d a t a for particulate 227Acare based on the sample collected by sediment traps. Anderson (1980) determined three samples from the PARFLUX Pacific site, southeast of Hawaii (Honjo personal communication). In Fig. 6, the results of the (227Ac/231Pa)activity ratio are compared with those obtained for seawaters of the western North Pacific (Nozaki, 1984). T h e (227Ac/231Pa)activity ratio in particulate matter is one order of magnitude smaller than that in seawater, indicating that KD(Pa)/Ko(Ac) = 10 (or KD(Ac) = 2 ~ 1 0 and ~ ) only a few tenth of percent for Ac is bound with particles. Another way of expressing the elemental reactivity with particles is the scavenging residence time which is inversely related t o KO. Nozaki (1984) has estimated the scavenging residence time of Ac to be ~ 3 0 0 0years which is 100 times longer than that of T h (Nozaki et al., 1981). This implies t h a t since the mean life of 2 2 7 Ais~ only 31 years, the particle scavenging can hardly influence the distribution of 2 2 7 Ain ~ the deep sea, and hence 227Accan be used as a steady-state “conservative” tracer (Craig, 1969).

Comparison of the (227Th/231Pa)activity ratios between Fig. 6. seawater and particulate matter. Data are based on Nozaki (1984) arid Anderson (1979).

149 4.2.

Diffusion of 2 2 7 Afrom ~ deep-sea sediments

The presence of excess 2 2 7 Ain ~ deep waters clearly indicates t h a t there is a n additional source to the water other than 231Padecay. The distribution pattern of 2 2 7 Aincreasing ~ toward (Nozaki, 1984) the bottom and the good correlation between 227Acand 228Ra(Nozaki and Yang, 1987) strongly suggests t h a t excess 227Acmay be derived from underlying sediments through porewater diffusion much like Ra isotopes (Cochran, 1979). This is not quite surprising because the rare earth elements are diagenetically mobilized through porewater of marine sediments (Elderfield and Sholkovitz, 1987). More directly, Nozaki et al. (1990) have measured the pore water profile of 227Acin Northwest Pacific deep-sea sediment (Fig. 7 ) . The station locations where 2 2 7 Ad~a t a are available are shown in Fig. 8. T h e 227Acconcentration decreasing toward the sediment-water interface implies t h a t 227Acis diffusing out of the sediment. The model calculation suggested t h a t the molecular diffusion alone can support only one half of the excess 2 2 7 A standing ~ crop in the water column and the remainder need t o be supplied by the aid of bioturbation. It is also estimated the Langmuir isotherm adsorption coefficient, K for Ac solid/solution interaction, t o be (3&1)x1O3 which is comparable with t h a t Ac-227 (dpm/kg) 0.00

0.05

0.15

0.10

20

0

PzO.30 crn2/s

5

10

0 15 CD

-2 5

h

2 20

I

v

25

30

? N

0 0

35

Fig. 7. The 227Th profile in the pore water of Northwest Pacific deep-sea sediment obtained at the AN-5 site of Fig. 8.

150

N

40

20

a Fig. 8. The station locations where seawater and sediment samples were collected for 227Acanalyses. Those were occupied by two cruises of R. V. Hakuho-Maru (Cepheus and Antares Expeditions in 1982 and 1984, respectively.)

for Ra (1000-5000) reported in Cochran and Krishnaswami (1980). T h e adsorption coefficient K is related to the distribution coefficient K D by the equation (Cochran, 1979) given by

' = 0.83 (Nozaki et al., 1990), h - D becomes -O.4x1O5, a Using ps = 0.35 and 3 somewhat lower value compared with K D = 2x105 in t,he water column. This difference, if significant, may be caused by one or combination of changes in surface properties of sediment particles due to decomposition of organic matter and MnO2 coating, and in solution chemistry of porewater due to complexation, low solution/solid ratio, etc. Noteworthy in the calculation is t h a t the reactivity of 2 2 7 A in ~ sediment porewater is approximately equal t o that of Ra, regardless of difference in the valency state (+3 versus $2) and thereby their different solution chemistry in laboratory experiment. However, this is not surprising because there is a marked resemblance between Ba (as analogue of Ra) and La (as analogue of Ac), whose oceanic distributions are strongly correlated t o that of reactive silica (Chan et al., 1976; Klinkhammer et al., 1983). Thus, so far, we have found any disernible difference between R a and Ac in their geochemical behavior in the deep-sea environment. 4.3.

2 2 7 Ain ~ coastal and surface waters

Little work has been made to establish the 227A~--231Pa disequilibrium relationship in the shallow water environment. This is because both 2 2 7 Aand ~ 231Pa concentrations are extremely low. Yet it is clear from the open ocean d a t a (see below) t h a t there is very little excess 227Acin the surface water in contrast t o

151 Ac-227 (dpm/lOOOL) 0.0

+

0.5

1 .o

1.5

0

)-9(

w

20001

m x m

AN-1

AN-1 mx 4000

P 2 5

3

6000

8000

+ 10000

Fig. 9. Trench.

X X

t--s-+

cw

10000

The vertical profiles of 2 2 7 Aand ~ "'Th

in the Izu-Ogasawara

the presence of large excess 228Ra. The reason for this is that, because of a lack of authigenic 231Pa,production of 227Acin the shallow water sediments is about a factor of 20 smaller than that in deep-sea sediments. This recoil production of 2 2 7 Ain ~ sediments simply depends upon the 231Pacontent which is 2.8f1.4 dpm/g for deep-sea sediments (Yang et al., 1986) and -0.1 dpm/g for shallow water sediments assuming -3 rnuglg of average U content in shallow water sediment and the radioactive equilibrium between 231Paand 235U.In addition, Ac might be more particle-reactive than R a i n the esturine-coastal environment. Li et al. (1975) found that R a is significantly desorbed from river-borne particles in the estuarine mixing zones. On the other hand, Sholkovitz and Elderfield (1987) showed that the rare earth elements are removed from water during estuarine mixing. Further study is clearly needed in these environment.

5.

Simultaneous Use of 2 2 7 Aand ~ 228Raas Tracers of DeepSea Processes

Until now, 227Acd a t a were available only in the western North Pacific (Nozaki, 1984; Nozaki and Yang, 1987). The profiles were obtained t o the bottom of up to 6000 meters. Here I report the first d a t a set reaching a maximum depth of 9750 m in the deepest part of the Izu-Ogasawara Trench (station AN-1 in Fig. 8). At the site, various natural and man-made radionuclides together with nutrients and trace metals were measured, but only the profiles of 2 2 7 A and ~ 228Th are shown in Fig. 9. My fundamental premise to show the 228Th data is t h a t they are presumably in equilibrium with the parent, 228Rain the deep waters below

152 ~ 3 0 0 0meters. The reason for this comes from the box-model type calculation for concordance of (228Th/228Ra) and (230Th/234U)ratio (Nozaki and Yamada, 1987) given by, (ARa-228) = (:Th-230) AU-234 (4)

(

ATh-228

Th-228

)

+

ATh-230

In deriving equation (4), the first-order scavenging rate constant is assumed to be equal for 228Thand 230Th. The measured (234U/230Th)for the deep waters is becomes 0.95 which is indistinguishable from about 2000, and hence (228Th/228Ra) the equilibrium value within the uncertainty of measurement. Thus, I refer "'Th in the deep waters as (228Ra)hereafter.

5.1.

Vertical profiles of 2 2 7 Aand ~ 228Th('"Ra) in the deep trench

The vertical profile of 2 2 7 Ain ~ general is sirnlar to that of (228Ra)except for the top -1000 in where ("'Ra) has a strong surface maximum. The disparity in the shallow water can be ascribed to the difference in the flux from coastal sediments and possibly in the scavenging rates in nearshore/offwhore mixing zone as discussed above. The profiles of 227Acand (228Ra)down t o -6000 m, the bottom depth of eastern abyssal plain, are very similar to those observed in the western North Pacific basin (Nozaki, 1984; Nozaki et al., 1981) showing an increase with depth. Below 6000 meters, both 2 2 7 Aand ~ (228Ra)show some scatter but are relatively constant with depth. The scatter of data within the trench for 2 2 7 Aand ~ (22sRa) is not consistent with any other nuclides and therefore I regard them to be artifact in the analysis rather than real. The vertical profiles obtained here are first interpreted by using a one dimensional steady-state diffusion model given by

where K Z is the vertical eddy diffusion coefficient, Z is the height above the bottom, X is the decay constant, C is the excess activity of radionuclide of interest over the parent. The choice of the model is not based on the reality of physical mixing but purely based on the fact that the available data set is one dimensional. Thus, its application to the d a t a will show the inadequancy rather than successfulness of the model. T h e solution of equation ( 5 ) for the boundary conditions of C = C0 at Z = 0 and C = 0 at Z -+ 00 is given by,

Some example of profiles for given li'z values are shown in Fig. 10, where the excess activity is normalized to 1 a t the bottom interface. Based on the equation (6), K z must be higher than lo4 cm2/s for the nearly constant 2 2 7 Aand ~ (228Ra) profiles within the trench, and 50 and 130 cm2/s, for the 2 2 7 A and ~ 228Radata respectively between 3000 m and 6000 m. The high apparent vertical diffusivity within the trench may be due to the narrow topography which tends to induce

153 rapid boundary mixing (Armi, 1978). It is probable that the water within the trench is very rapidly renewed with the deep water of 5000 to 6000 meters around this region. The inconsistency of the apparent vertical diffusion coefficient derived from the 2 2 7 Aand ~ (228Ra)d a t a suggests that the chemical distributions are largely governed by isopycnal mixing (Sarmiento et al., 1982). This is also supported by the simlarity in the vertical profile down to -6000 meters between the Izu-Ogasawara trench and the western North Pacific regardless of the difference in the bottom depth.

0.0

0.2

0.4

0.6

0.8

1.0

0.2

0.4

0.6

0.8

Ra-228iAc-227

0.0

0.2

0.4

0.6

0.8

1.0

1.2

Ra-228

Ac-227

0.0

1.2

1.0

Ac-227

Fig. 10. The calculated profiles of 2 2 7 Aand ~ 228Rabased on a simple diffusion model and given Eiz values. Note t h a t the vertical model 228Ra2 2 7 A plot ~ gives a concave curve which differ from the rapid isopycnal mixing model with negligible radioactive decay (the solid line of Fig. 10d).

154 3

T

0

1

2

3

Ac-227 (dpm/lOOOL) Fig. 11. The correlation diagram between 2 2 7 Aand ~ 228Rain the deep water. The d a t a (squares) deeper than 3000 m at the AN-1 station are plotted. The calculated curve of the simple vertical mixing model is shown by the diamonds which differs from the rapid isopycnal mixing model with (227Ac/228Ra)ratio of 1 and negligible decay (the solid line). One of the interesting outcomes of the simple vertical model is shown in Fig. 10d and 11, where the 22sRavesus 2 2 7 Aplot ~ shows the concave feature regardless of the value of K z . However, if the distribution of the two nuclides is governed by rapid horizontal mixing (hence, negligible decay) with a contstant 22sRa to 227Acsource ratio, it should approach t o the linear line as shown in Fig. 11. Thus, the difference between the vertical and horizontal models suggests t h a t the simultaneous use of 228Raand 227Actracers should, at least in principle, resolve the relative importance of diapycinal and isopycnal mixing in the deep sea. The published d a t a near the slope of Japanese Islands obtained by Nozaki and Yang (1987) appears to follow a linear trend rather than a concave curve, showing their distributions are largely governed by lateral mixing.

Acknowledgements This work was supported by the Ministry of Education, Science and Culture, Japan through the Grant-in-Aid Nos. 63610004 and 01610003 to the University of Tokyo. The manuscript was typed by Ms. K. Hasegawa.

References Amin, B. S., S. Krishnaswami and B. L. K. Somayajulu, 1974. z34Th/238U activity ratios in Pacific Ocean bottom waters, Earth Planet. Sci. Lett., 21, 342. Anderson, R. F., 1981. The Marine Geochemistry of Thorium and Protactinium. Ph. D. Thesis, M.1.T.-W.H.O.I. Joint Program, Woods Hole, Mass., 287 pp.

155 Armi, L., 1978. Some evidence for boundary mixing in the deep ocean. J. Geophys. Res., 83, 1971-1979. Bhat, S. G., S. Krishnaswami, D. Lal, Rama and W. S. Moore, 1969. Th234/U238ratios in the ocean. Earth Planet. Sci. Lett., 4, 483-491. Broecker, W. S., 1965. An application of natural radon to problem in ocean circulation. In: Symposium on Diffusion in Oceans and Fresh Waters, Lamont-Doherty Geological Observatory, Palisades, NY, pp. 116-145. Chan, L. H., J. M. Edmond, R. F. Stallard, W. S. Broecker, Y. C. Chung, R. F. Weiss and T . L. Ku, 1976. Radium and barium at GEOSECS stations in the Atlantic and Pacific. Earth Planet. Sci. Lett., 32, 258-267. Chung, Y., and H. Craig, 1972. Excess radon and temperature profiles from the eastern equatorial Pacific. Earth Planet. Sci. Lett, 14, 55-64. Cochran, J. K., 1979. The Geochemistry of 226Raand 228Rain Marine Deposits. Ph. D. Thesis, Yale University, New Haven, Conn. 260 pp. Cochran, J. K., and S. Krishnaswami, 1980. Radium, thorium, uranium and 'loPb in deep-sea sediments and sediment pore waters from the North Equatorial Pacific. Amer. J. Sci., 280, 849-889. Craig, H., 1969. Abyssal carbon and radiocarbon in the Pacific. J . Geophys. Res., 74, 5491-5506. Elderfield, H., and E. R. Sholkovitz, 1987. Rare earth elements in the pore waters of reducing nearshore sediments. Earth Planet. Sci. Lett., 82, 280-288. GESAMP, 1983. An Oceanographic Model for the Dispersion of Wastes Disposed of in the Deep Sea. IAEA, Vienna, 182 pp. Horibe, Y., and H. Tsubota, 1977. A large-volume water sampler. In: Preservation of the Marine Environment, Report of Co-operative Research, Vol. 1, Y. Horibe (ed.), University of Tokyo Press, 199 pp. Horibe, Y., Y. Takashima,, T . Gamo, Y. Nozaki, M. Sakanoue and Y. Nagaya, 1986. Mixing of seawater studied with chemical tracers. In: Dynamics of the Oceans, K. Kajiura (ed.), Koseikaku-Koseisha, Tokyo, pp. 9-77. Karalova, J . K., and B. F. Myasoyedov, 1982. Actinium, Science Academy, Moscow, USSR. Klinkhammer, G., H. Elderfield and A. Hudson, 1983. Rare earth elements in seawater near hydrothermal vents. Nature, 305, 185-188. Krishnaswami, S., D. Lal, B. L. K. Somayajulu, R. F. Weiss and H. Craig, 1976. Largevolume in-situ filtration of deep Pacific waters: mineralogical and radioisotope studies. Earth Planet. Sci. Lett., 32, 420-429. Latimer, W. M., 1956. Oxidation Potentials, 2nd Edition, Prentice-Hall, Inc., Englewood Cliffs, N. J., 392 pp. Li, Y. H., and L. H. Chan, 1979. Desorption of Ba and 226Rafrom river borne sediments in the Hudson Estuary. Earth Planet. Sci. Lett., 43, 343-350. Lucas, H. F., 1957. Improved low-level alpha-scintillation counter for radon. Rev. Sci. Instrum., 28, 680-683. Moore, W. S., 1969. Oceanic concentrations of "'Ra. Earth Planet. Sci. Lett., 6, 437-446. Moore, W. S., 1976. Sampling "'Ra in the deep ocean. Deep-sea Res., 23, 647-651. Moore, W. S., and D. F. Reid, 1973. Extraction of radium from natural waters using manganese-impregnated acrylic fibers. J. Geophys. Res., 78, 8880-8885. Nagaya, Y., and K. Nakamura, 1981. Artificial radionuclides in the western Northwest Pacific (I); "Sr and 137Csin the deep waters. J. Oceangr. Soc. Japan, 37, 135-144. Nagaya, Y., and K. Nakamura, 1987. Artificial radionuclides in the western Northwest Pacific (11); 137Csand 239,240Pu inventories in water and sediment columns observed

from 1980 to 1986. J. Oceanogr. SOC.Japan, 43, 345-355. Nozaki, Y., 1983. Determination of thorium isotopes in seawater by moored MnOz-fiber method. J . Oceanogr. SOC.Japan, 39, 129-135. Nozaki, Y., 1984. Excess "?Ac in deep ocean water. Nature, 310, 486-488. Nozaki, Y., and T. Nakanishi, 1985. 231Paand 230Thprofiles in the open ocean water column. Deep-sea Res., 32, 1209-1220. Nozaki, Y., and H. Yang, 1987. T h and P a isotopes in the water of the western margin of the Pacific near Japan: Evidence for release of "* Ra and "'Ac from slope sediments. J. Oceanogr. SOC.Japan, 43, 217-227. Nozaki, Y., Y. Horibe and H. Tsubota, 1981. The water column distributions of thorium isotopes in the western North Pacific. Earth Planet. Sci. Lett., 54, 203-216. Nozaki, Y., and Y. Horibe, 1983. Alpha-emitting thorium isotopes in northwest Pacific deep waters. Earth Planet. Sci. Lett., 65, 39-50. Nozaki, Y., H. S. Yang and M. Yamada, 1987. Scavenging of thorium in the ocean. J. Geophys. Res., 92, 772-778. Nozaki, Y., and M. Yamada, 1987. Thorium and Protactinium isotope distributions in waters of the Japan Sea. Deep-sea Res., 34, 1417-1430. Nozaki, Y., M. Yamada and H. Nikaido, 1990. The marine geochemistry of actinium-227: Evidence for its migration through sediment pore water. Geophys. Res. Lett., 17, 1933-1936. Orr, J. C., N. L. Guinasso and D. R. Schink, 1984. Device for Rapid in situ Extraction of Tracers from Large Volumes of Seawater. Tech. Resp. 84-T-1, Dep. of Oceangr., Texas A&M University, College Station, TX, 60 pp. Orr, J. C., 1988. Evaluation of counting methods for oceanic radium-228. J. Geophys. Res., 93, 8265-8278. Rama, J. F. Todd, J. L. Butts and W. S. Moore, 1987. A new methods for the rapid measurement of 224Rain natural waters. Mar. Chem., 22, 43-54. Reid, D. F., R. M. Key and D. R. Schink, 1979. Radium, thorium, and actinium extraction from seawater using an improved manganese-oxide-coated fiber. Earth Planet. Sci. Lett., 43, 223-226. Sarmiento, J. L., C. G. H. Rooth and W. S. Broecker, 1982. Radium 228 as a tracer of basin wide processes in the abyssal ocean. J. Geophys. Res., 87, 9694-9698. Sarmiento, J. L., H. W. Feely, W. S. Moore, A. E. Bainbridge and W. S. Broecker, 1976. The relationship between vertical eddy diffusion and buoyancy gradient in the deep sea. Earth Planet. Sci. Lett., 32, 357-370. Spencer, D. W., P. G. Brewer, A. Fleer, S. Honjo, S. Krishnaswami and Y. Nozaki, 1978. Chemical fluxes from a sediment trap experiment in the deep Sargasso Sea. J. Mar. Res., 36, 493-523. Turekian, K. K., Y. Nozaki and L. K. Benninger, 1977. Geochemistry of atmospheric radon and radon products. Ann. Rev. Earth Planet. Sci., 5, 227-255. Yang, H. S., Y. Nozaki, H. Sakai and A. Masuda, 1986. The distribution of 230Thand 231Pain the deep-sea surface sediments of the Pacific Ocean. Geochim. Cosmochim. Acta, 50, 81-89.

Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

157

Distributions and Mass-Balance of 239,240Pu and 137Csin the Northern North Pacific Yutaka NAGAYA and Kiyoshi NAKAMURA*

Abstract 239,240Pu and '37Cs,in t h e water and sediment columns of t h e northern North Pacific (including t h e Bering Sea) were determined in 1988, and their distributions were compared with those in the southern regions. T h e 137Cs water and sediment inventories in t h e area north of t h e Subarctic Convergence are remarkably lower than t h e estimated global fallout inputs. This suggests t h a t 137Cs is effectively transported southward from t h e northern upwelling regions t o the central Pacific gyre. T h e difference between t h e total inventory and t h e atmospheric input of 239,240 P u is relatively smaller than that of 137Cs. This is attributable to t h e fact t h a t P u is more particle reactive than Cs, and hence it is more actively involved in t h e vertical biogeochemical cycling. Penetration of t h e radionuclides into t h e sediment column occurs due mostly t o bioturbation. T h e particle mixing coefficients calculated based on artificial radionuclide distributions are comparable t o those estimated from excess zlOPb.

1.

Introduction

Artificial radionuclides released, largely due to the nuclear weapon testing with a maximum in 1963, into the marine environment have been regarded as useful tracers in the study of various marine processes. Hence their distributions and pathways in the marine environment have been investigated extensively for our understanding of not only the degree of radioactive contaniination but also natural geochemical and oceanographic processes occurring in the ocean (for example, see Pentreath, 1988). The recent studies of 2391240Puand 137Csin the North Pacific by Bowen e t al. (1980) and Nagaya and Nakamura (1981, 1984 and 1987) showed t h a t there is considerable disagreement between the inventories of the radionuclides in the water column and the estimated global fallout inputs t o the surface ocean. In the central North Pacific, the water column inventories were higher than t.he estimat,ed atmospheric inputs, and it was suggested t h a t the discrepancy is attributable t o the contribution of local tropospheric input which presumably occurred near the test site of the equatorial North Pacific and it might also be in part due to the contribution of the remobilized radionuclides from the test sites (Noshikin and Wong, *Division of Marine Radioecology, National Institute of Radiological Sciences, Isozaki, Nakaminato, 311-12 Japan

158 1980; McMurtry et al., 1985). On the other hand, the water column inventories were nearly equal or slightly lower than the fallout inputs in the northern and northwestern North Pacific. In our previous paper (Nagaya and Nakamura, 1987), we suggested that this latitudinal transition in reversing the ratio of standing stock t o fallout or the radionuclides occurs across the Subarctic Convergence around 40°N. However the observed data on the area north of the Subarctic Convergence were sparse for that time, and it was necessary to make further measurements t o confirm this. In this paper, we present the new data on 2393240Puand 137Cs in the northern North Pacific and their geochemical significance is discussed.

2.

Materials and Methods

All sea water and sediment samples were collected during the KH-88-3 cruise (Draco Expedition, 1988) of R/V Hakuho-Maru, the Ocean Research Institute, University of Tokyo. Sampling stations are shown in Fig. 1, together with those occupied by the previous cruises (Nagaya and Nakamura, 1981, 1984, 1987). Surface water samples ( w 100 1 each) were collected by using a pump, and subsurface samples ( w 250 1 each) were collected using a large volume sampler described in Horibe et al. (1977). Sediment samples were collected with a box

Fig. 1. Sampling stations. 0: this work, 0 : previous cruise (CY..1980; DE..1985; B,C..1986). a: surface water and sediment, b: surface water only, c: sediment only.

159

4

-

-

2t

4

0 U

-0-0.

0 CY-6

L

I

+ - u 0

DR-I0 -0-

6

Fig. 2. Vertical distribution of 239,240Pu in north (DR-10) and south (CY-6) of the Subarctic Convergence. corer (50 cm x 50 cm x 50 cm). After recovery a 12 cm x 12 cm square core was subsampled from each sediment, and then extruded and cut at 1 t o 2 cm intervals on board the ship. 2 3 9 , 2 4 0a ~ ~ nd 137Cscontents were determined for all samples, except 2391240Pu contents in subsurface water samples from 2 stations (DR-5 and DR-21). Analytical procedures for the radionuclides were described elsewhere (Nagaya and Nakamura, 1981, 1984, 1986; Nakamura and Nagaya, 1985). The natural radionuclide 210Pb in the sediment samples was determined by gamma-spectrometry, and calculated from its count rate at the gamma energy of 46.5 keV.

3.

Results and Discussion

The analytical data are shown in Tables 1 and 2. All the d a t a are corrected for radioactive decay t o the date of sample collection. The radionuclide inventories calculated from the d a t a are given in Table 3. 3.1.

Vertical profiles of 239,240Pu and 137Csin the water and sediment columns

The vertical profiles of 2393240Pu and 137Csin the water column at 50"N (DR-10) obtained here are compared with those at 40"N (CY-6) reported earlier (Nagaya

160 Table 1. Results of sea water analysis. DEPTH (m) 0

T ("C)

(Ym)

-

-

IJ7 cs (mBq/t) 4.94f0.29

DR- 5

0 103 257 401 598 789 947 2,000 3,486 4,999 6,239

8.2 8.32 6.02 4.59 4.02 3.45 3.40 1.88 1.51 1.57 1.70

33.360 33.048 33.884 33.892 34.100 34.279 34.385 34.599 34.681 34.691 34.689

DR- 7

0

-

DR-10

0 106 279 397 598 799 997 2,003 2,989 4,000 5,615

0 105 204 405 700 1,000 1,499 2,505 3,937

STN Ur- 2

DR-13

DR-18

0

DR-21

0 96 246 402 602 797 1,002 2,003 3,014 4,001 4,850

DR-23

0

S

(mBq/100l) 4.35-

PU/CS x102 (Activity ratio) 0.88fO.09

3.23f0.22 3.43f0.15 3.09f0.15 2.17f0.12 0.70f0.07 0.45f0.07 0.31f0.09 0.04f0.04 0.01f0.03 0.06f0.04 0.02f0.04

1.12f0.26

0.35f0.08

34.483

3.23f0.20

1.45f0.19

0.45f0.07

4.98 2.03 3.24 3.35 3.04 2.77 2.49 1.78 1.55 1.48 1.64

33.142 33.392 34.048 34.170 34.286 34.360 34.434 34.602 34.647 34.669 34.693

2.64f0.18 2.08f0.10 0.80f0.08 0.56f0.07 0.28f0.05 0.17f0.06 0.15f0.07 0.04k0.04 0.09f0.04 0.11f0.04 0.07f0.04

2.02f0.19 0.97f0.05 1.63f0.23

0.77f0.09 0.47f0.03 2.0f0.4

6.11 2.40 3.07 3.55 3.20 2.85 2.28 1.73 1.59

33.191 33.304 34.578 34.031 34.240 34.362 34.500 34.632 34.655

2.82f0.22 1.71f0.11 2.34f0.12 0.88f0.07 0.29f0.06 0.14f0.03 0.08f0.03 0.09f0.03 0.03f0.03

1.85f0.21 0.80f0.07 1.30f0.07 1.63f0.07 1.48f0.08 1.20f0.09 0.90f0.07 0.56f0.05 0.9150.07

0.66f0.09 0.47f0.05 0.56f0.04 1.9f0.02 5.1fl.l 8.6f2.0 11.3414.3 6.5f2.6

-

32.921

2.1150.16

1.23f0.28

0.6f0.1

8.51 4.45 3.96 3.80 3.45 3.13 2.84 1.96 1.57 1.52 1.58

32.763 32.853 33.877 34.049 34.210 34.309 34.383 34.588 34.647 34.681 34.684

2.7960.19 1.94f0.09 1.69f0.10 0.97f0.07 0.41f0.06 0.36f0.05 0.06f0.03 0.04f0.03 0.12f0.04 0.06f0.03 0.03f0.03

0.40f0.09

0.14f0.03

-

32.527

2.63f0.18

0.23f0.07

0.09f0.03

2syJ4uPU

-

-

2.25f0.24 l.ll*0.20 1.14f0.08 0.60f0.05 0.41f0.05 0.57f0.06 1.00f0.07

8.0f1.7 6.5f2.6 1.7f3.6

* Activity ratio is neglected because of large deviation.

*

4.8f2.4 5.2f2.0 14.3f8.2

*

161 Table 2.

(cm) 0- 2 2- 4 4- 6 6- 8 8-10 10-12 12-14 14-16

Water Contents @) 67.9 68.4 65.5 71.2 63.3 62.6 61.6 60.8

0- 3 3- 6 6- 9 9-12 12-15 15-18 18-21

73.1 70.5 67.2 66.6 64.5 63.5 63.7

5.83f0.34 1.63k0.22 0.42f0.15 0.37f0.14 0.15f.O.13 0.08f0.11 0.00f0.12

0.571f0.017 0.271f0.010 0.090f0.006 0.072f0.005 0.027f0.003 0.024f0.003 0.009f0.002

0.098f0.006 0.17f0.02 0.21f0.08 0.20f0.08 0.18f0.16

1,087k33 864f29 683f26 687f26 717f27 711f27 680f26

0- 2 2- 4 4- 6 6- 8 8-10 10-12 12-14 14-16

70.0 65.9 63.0 61.0 64.6 67.0 68.7 67.1

3.1l f 0 . 2 6 2.11k0.21 1.05k0.15 0.39f0.17 0.17f0.12 0.06f0.11 0.08f0.15 0.16f0.14

0.267f0.009 0.210f0.013 0.164f0.010 0.057k0.004 0.035f0.004 0.007k0.013 0.012k0.003 0.007f0.002

0.086f0.008 0.1OkO.01 0.16f0.02 0.15f0.07 0.21f.0.15

1,253f35 857f29 507f23 378f19 468f22 739f27 592f24 361f19

2 4 6 8 8-10 10-12 12-14 14-16

77.7 75.2 72.5 70.6 70.0 67.0 69.2 69.0

7.07h0.41 3.37f0.29 0.64f0.20 0.21f0.13 0.1160. 13 0.01f0.12

0.843f0.033 0.699f0.025 0.193f0.011 0.040f0.004 0.032k0.002 0.027k0.003 O.OOkO.ll 0.005f0.001 0 . O O f 0. 10 0.004f0.001

0.119f0.008 0.21f0.02 0.30f0.10 0.19k0.12

0- 2 2- 4 4- 6 6- 8 8-10 10-12 12-14 14-16

66.1 55.7 52.4 51.4 50.2 48.5 50.5 49.6

2.4df0.24 0.98f0.17 0.75k0.17 0.38f0.14 0.29f0.12 0.00f0.08 0.07f0.10 0.00f0.09

Depth

STN DR- 2

DR- 5

DR-10

DR-13

DR-21

0246-

cs

23y,24"

Pu

Pu/Cs

0- 2 2- 4 4- 6 6- 8 8-10 10-12 12-14 14-16

""Pb

(Ba/kg-dry) (Activity ratio) (Ba/kg-dry) 1.58f0.20 0.320f0.015 0.20f0.03 655f26 2.86f0.24 0.284f0.011 0.099f0.009 854f29 2.9950.25 0.388f0.015 0.1360.01 785f28 0.26f0.13 0.027f0.007 0.10f0.06 522f23 0.22f0.14 0.162f0.007 0.74f0.47 329f18 0.19f0.13 0.041f0.004 243f16 0.22f0.15 0.18f0.14 0.011f0.002 0.0660.05 227f15 0.30f0.17 0.011f0.003 226f15 0.04f0.02

0.511f0.017 0.303f0.014 0.144f0.007 0.065f0.004 0.031f0.004 0.016f0.002 0.008k0.001 0.008f0.002

* *

* *

0.04f0.04

*

* * * 0.21f0.02 0.3140.06 0.19f0.05 0.17f0.06 0.11f0.05

* * *

77.8 6.77f0.37 1.880f0.038 0.28f0.02 73.2 4.67f0.36 1.361f0.036 0.29f0.02 73.0 2.41f0.84 0.961f0.022 0.40f0.14 69.1 0.77f0.37 0.274f0.010 0.36fO.34 68.7 1.29k0.42 0.305f0.011 0.24f0.15 * 68.7 0.20k0.38 0.055f0.004 66.5 O.OOk0.42 0.021f0.002 65.1 0.31k0.25 0.016f0.002 0.054~0.04 Activity ratio is neglected because of large deviation.

DR-25

*

Results of sediment analysis.

*

2,726f52 1,7244142 601f25 592f24 662k26 622f25 427421 350f19 533f23 389620 348f19 379f20 429f21 410f20 462f22 394f20 1,846f43 1,341f37 723f27 649f26 449f21 340f18 245f16 253f16

162 IR7CsConc. (mBq/l) 0.01

0.1

1

10

E

Y

5n

3

--

4

--

B

5

--

6

--

0

CY-6

0

DR-10

Fig. 3. Vertical distribution of 137Cs in north (DR-10) and south (CY-6) of the Subarctic Convergence. and Nakamura, 1984) in Figs. 2 and 3, respectively. General pattern of profiles for each radionuclide are similar in the two stations, but the radionuclide contents are remarkably different between the stations, especially in upper part of the water column. Naturally this difference in the radionuclide contents resulted in the discrepancy in water column inventories for the southern and northern stations. 2391240Pu/137Cs ratio in water column tends t o increase with increasing depth, since the value is generally higher in the near bottom water than in mid depth. The ratio in the bottom water is nearly equal t o that in the surface sediment. These suggest that 239)240Pu, in comparison with 137Cs,is more actively scavenged from the upper layer of water column and tIansported by settling particulate matter t o the bottom, where a part of particulate nuclide are generated into the bottom water. The vertical profiles of the radionuclides in the sediment columns (Table 2) indicate approximately exponential decrease of the concentrations with depth except for station DR-2 where a subsurface maximum of 239)240Pu often found in the southern (< 40'N) area (Yang et al., 1986; Nagaya and Nakamura, 1987) was observed. The 239,240Pu/137Cs ratio in the sediment columns does not seem to show any systematic tendency with depth.

163 Table 3. Radionuclide Inventories. Stn

WATER Column

Sediment Column

3.2.

DR- 5 DR-10 DR-13 DR-21 DR- 2 DR- 5 DR-10 DR-13 DR-21 DR-25

239J4" P u (MBq/km2 1 1,990f90 1,150f80 47fl 1,070f50 37fl 1,220f50

137cs

66.3f4.4 78.3f4.8 61.6f4.3 61.9f3.7 44.9f4.9 103.3f13.4

9.5f0.2 10.1h0.3 6.8f0.2 10.550.2 9.9f0.2 30.4f0.4

Pu/Cs (Activity ratio) ~

0.041f0.003 0.026f0.001 -

0.14f0.01 0.13f0.01 O.llfO.O1 0.17f0.01 0.22f0.03 0.29f0.04

Radionuclide water and sediment inventories and atmospheric fallout input

The radionuclide inventories are latitudinally compared in Table 4. The results for station DR-25 were excluded from the table because the inventories are unusually high. The higher inventories at DR-25 are presumably caused by lateral input from nearby North American terrestrial sources (Koide et al., 1980; Nakamura and Nagaya, 1985, 1990). The global fallout inputs in the table are based on the UNSCEAR report (UNSCEAR, 1982). The radioactive decay of delivered 137Cs(half life 30.17 years) was corrected t o January 1, 1989, but no correction was needed for 2391240Pubecause of their long half lives ( 2 . 4 lo4 ~ and 6 . 5 7 lo3 ~ years, respectively). The radionuclide inventories in the water column are remarkably higher than those in the sediment column for the same location. Clearly most of the radionuclide delivered to the sea are retained in the upper layer of the water column. The mean ratios of the sediment t o water inventories are 2.2f1.2% for 137Cs and 8.9&3.2% for 2399240Puin between 30"N and 40"N, and 4.4f1.3% for 137Cs and 18.5f6.3% for 2391240Puin between 40"N and 55"N. These ratios indicate that 2391240Puis more effectively removed from the water column t o the sediment than 137Csdoes. The higher particle reactivity of P u is also indicated by the vertical profile of the 239,240Pu/137Cs ratio in the water column. Despite the relatively lower global fallout inputs, the water column inventories in the southern (30 t o 40"N) area are higher than those of the northern (40 t o 55"N) area. Since the sediment inventories are approximately equal in the two latitudinal zones, the total (water sediment) inventories in the southern area also become higher than those in the northern area. The ratios of the total inventory to the atmospheric input for 3O0N-40"N are 1.5f0.5 for 137Csand 3.4f0.7 for 2391240Pu. On the other hand, for 4OoN-55"N, the mean total inventory of 137Csis about 1 / 3 of the fallout input and that of 2391240Pu is nearly equal to the fallout input.

+

164 Table 4.

Stn.

Water Column

Sediment Column

DR- 5 CY- 5 CY- 6 CY- 8 CY-11 mean DR- 2 DR- 5 CY- 5 CY- 6 CY- 8 CY- 9

mean Fallout* *UNSCEAR (1982)

Comparison of radionuclide inventories.

3OoN-4O0N '5"Cs Z3YJ4"PU (MBq/km2) 1,990 2,980 87 3,380 118 4,460 135 4,560 156 3,480&1,0709 66 78 46 146 76 58 78f35 2,370

Stn.

DR-10 DR-13 DR-21 DE- 2 DE- 4

124f29 10 10 6 17 13 11

mean DR-10 DR-13 DR-21 DE- 5

llf3 40

mean

B C

40°N-55'N 13'(CS Z3YJ4"PU (MBq/km2) 1,150 47 1,070 37 1,221 1,095 35 1,490 49 1,205f171 62 62 45 68 54 30 53f14 3,260

42f7 6.8 10.5 9.9 8.3 4.9 5.6 7.7f2.3 56

These differences suggest that the lateral redistribution by oceanic mixing is significant and there seems t o be a net transport of the radionuclides from north t o south in the northern part of the Pacific. Especially, the lower 137Csinventory in the area north of the Subarctic Convergence suggests that the outflow of surface and subsurface water may be occurring to the southern area presumably due to upwelling of deep water, having very low 137Csconcentration.

3.3.

Penetration of the radionuclides into the sediment column

The vertical profiles of the artificial radionuclides in sediments have been interpreted as a consequence of biological sediment particle mixing, so called "bioturbation". This process is often described as the simple particle mixing analogous t o Fickian diffusion (for example, see Cochran, 1985). Here the vertical particle mixing coefficient DB was calculated using a pulse input and a continuous input models. According to Cochran (1985) and Yang et al. (1986), the equation of the pulse input model is given by A = exp(-z2/4~Bt) (1) A0

For the continuous input model, the equation becomes

.

165

where, erfc is the error function complement. As sample collection was done in 1988, t = 25 years is used for the calculation. The results are given in Table 5. Generally speaking, the profiles of the artificial radionuclides in 40 t o 55"N are better fitted by the pulse input model (Cochran, 1985; Yang et al., 1986; Lapicque et al., 1987). The D B values of 137Cs and 239,240Pu in the area north of the Subarctic Convergence (0.03 t o 0.39 cm2/y) are similar t o those calculated using the pulse input model in 30-40"N area, 0.02 t o 0.06 cm2/y (Yang et al., 1986) and those in the equatorial North Pacific, 0.03 to 0.36 cm2/y (Cohcran, 1985). The calculation was not made for the DR-2 sediment because the d a t a are not consistent with the simple model. The subsurface maxima and the deeper penetration of the radionuclides observed in DR-2 sediment might be caused by the preferential transport of biogenic sinking particles through burrows of benthic organisms (Cochran, 1985; Stordal et al., 1985; Smith et al., 1986; Yang et al., 1986; Nagaya and Nakamura, 1987). The sediment particle mixing ate was also calculated a natural radionuclide 210Pb with a half life of 22.3 years (Nozaki et al., 1977). The vertical distribution of excess 210Pb (relative to 226Ra)and the calculated mixing rates are shown in Fig. 4. The excess 'loPb is estimated by subtracting the constant 210Pb activity in the deeper layer from the measured 210Pb activity in the shallower depths. T h e equation used for calculation of D B is given by

A=

. exp(-

JT;-I'DB.Z )

(3)

In the calculation, some excess 'loPb d a t a were omitted, because they are not fitted well with the particle mixing model as shown in Fig. 4 by open circles. The results

1

DR-2

DR-5

Dg=OlI

DR-I0

Dg

E

DR-13

DB = 0 0 5

007

L Depth (cm)

DR-21

\I7

e

A

DB = 041,

DR-25 D~ = n 4 n

\

_mlT_

in

Fig. 4. Vertical profile and calculated particle mixing coefficient (DB, cm2/y) of excess 210Pbin sediment column.

166 Table 5. Calculated particle mixing coefficients (cm2/y). Stn

Pulse Input 137cs

DR- 5 DR-10 DR-13

DR-21 DR-25

0.07-0.29 0.10-0.22 0.05-0.15 0.04-0.30 0.11-0.39

Continuous Input

239,240~~ 1 3 7 ~ 2 ~3 9 , 2 4 0 ~ ~

0.12-0.39 0.16-0.31 0.11-0.21 0.08-0.23 0.12-0.35

0.1-0.5 0.3-0.5 0.1-0.2 0.1-0.5 0.3-0.7

0.3-0.6 0.4-0.8 0.2-0.8 0.1-0.5 0.3-0.7

show the particle mixing coefficients ranging 0.05 to 0.46 cm2/y for excess 210Pb, and the values are similar to those estimated form 137Cs (0.03 to 0.39 cm2/y) and 2391240Pu (0.08 to 0.39 cm2/y).

4.

Conclusion

T h e artificial radionuclide 13’Cs which originated from the global radioactive fallout, in the northern North Pacific have been laterally transported southward across the 40”N latitudinal zone. This transport is presumably due t o upwelling of the deep water in the northern region, which have very low 137Csconstant, and subsequent outtiow of the surface and subsurface waters top the southern area. The lower 137Cs inventory in the northern area than t h a t in the southern area is mostly resulted by this transport. 239,240Pu,also derived from the global fallout, is more particle reactive than 137Cs and therefore more actively involved in the vertical geochemical cycling, and hence its lateral transport seem less than that of 137Cs. The penetration of t h e artificial radionuclides into sediment column in the northern North Pacific occur due mostly t o bioturbation and the particle mixing coefficients calculated are comparable to those estimated from excess 210Pb.

Acknowledgement We wish to express our hearty thanks to Dr. Y. Nozaki and the staff of the Ocean Research Institute, University of Tokyo, and officers and crew of R/V Hakuho-Maru of the Institute for their support in collecting samples. We are also grateful t o the scientists who participated on the cruise for their cooperation in sample collection.

References Bowen, V. T., V. E. Noshkin, H. D. Livingston, and H. L. Volchok, 1980. Fallout radionuclides in the Pacific Ocean; Vertical and horizontal distributions, largely from GEOSECS stations. Earth Planet. Sci. Lett., 49, 411-434. Cochran, J. K., 1985. Particle mixing rates in sediments of the eastern equatorial Pacific: Evidence from 210Pb,239,240Pu and 13’Cs distributions at MANOP sites. Geochim. Cosmochim. Acta, 49, 1195-1210. Horibe, Y., K. Taira, T. Terashima, H. Tsubota, S. Imawaki, Y. Kodama, and H. Igarashi, 1977. Development and recovery of moored arrays of instruments, large volume sam-

167 pler, autoanalyzer. In: Environmental Marine Science, Y. Horibe (ed.), Univ. Tokyo Press, Tokyo, pp. 182-208 (in Japanese). Koide, M., E. D. Goldberg, and V. F. Hodge, 1980. 241PuandZ4'Am in sediment from coastal basins off California and Mexico. Earth Planet. Sci. Lett., 48, 250-256. Lapique, G., H. D. Livingston, C. E. Lambert, E. Bard, and L. D. Labeyrie, 1987. Inin Atlantic sediment with a non-steady state input model. terpretation of 239,240Pu Deep-sea Res., 34, 1841-1850. McMurtry, G. M., R. C. Schneider, P. L. Colin, R. W . Buddemeier, and T. H. Suchanek, 1985. Redistribution of fallout radionuclides in Eniwetok Atoll lagoon sediments by callianassid bioturbation. Nature, 313, 674-677. Nagaya, Y., and K. Nakamura, 1981. Artificial radionuclides in the western Northwest Pacific (I): "Sr and 137Csin the deep waters. J. Oceanogr. SOC.Japan, 37, 135-144. Nagaya, Y., and K. Nakamura, 1984. 239,240Pu, 137Cs and "Sr in the central Pacific. J. Oceanogr. SOC.Japan, 40, 416-424. Nagaya, Y., and K. Nakamura, 1986. 239,240Pu and 137Csconcentrations in some marine biota, mostly from the seas around Japan. Nippon Suisan Gakkaishi, 53, 873-879. Nagaya, Y., and K. Nakamura, 1987. Artificial radionuclides in the western Northwest inventories ' in water and sediment columns observed Pacific (11): 137Csand 239,240Pu from 1980 to 1986. J. Oceanogr. SOC.Japan, 43, 345-355. Nakamura, K., and Y. Nagaya, 1985. Accumulation of Cs-137 and Pu-239,240 in sediments of the coastal sea and the North Pacific. In: Marine and Estuarine Geochemistry, C. Sigleo and H. Hattori (ed.), Lewis Pub., Inc., Chelsea, pp. 171-180. in the sediment Nakamura, K., and Y. Nagaya, 1990. Distribution of 137Csand 239,240Pu of the Set0 Inland Sea. J. Radioanal. Nucl. Chem., Articles, 138, 153-164. Noshkin, V. E., and K. W. Wong, 1980. Plutonium mobilization from sedimentary sources to solution in the marine environment. In: Marine Radioecology, OECD-NEA, Paris, pp. 165-178. Nozaki, Y., J. K. Cochran, K. K. Turekian, and G. Keller, 1977. Radio-carbon and lead210 distribution in submersible-taken deep-sea cores from Project FAMOUS. Earth Planet. Sci. Lett., 34, 167-173. Pentreath, R. J., 1988. Sources of artificial radionuclides in the marine environment. In: Radionuclides. A Tool for Oceanography, J. C. Guary, P. Guegueniat and R. J. Pentreath (ed.), Elsevier Applied Science, London, pp. 12-34. Smith, J. N., B. P. Boudreau, and V. E. Noshkin, 1986. Plutonium and '"Pb distributions in northeast Atlantic sediments: Subsurface anomalies caused by non-local mixing. Earth Planet. Sci. Lett., 81, 15-28. Stordal, M. C., J. W. Johnson, N. L. Guinasso, and D. R. Schink, 1985. Quantitative evaluation of bioturbation rates in deep ocean sediments. 11. Comparison of rates determined by 'loPb and 239,240Pu. Mar. Chem., 17, 99-114. Yang, H.-S., Y. Nozaki, H. Sakai, Y. Nagaya, and K. Nakamura, 1986. Natural and man-made radionuclide distributions in Northwest Pacific deep-sea sediments: Rates of sedimentation, bioturbation and 226Ramigration. Geochem. J., 20, 29-40. United Nations Scientific Committee on the Effects of Atomic Radiation, 1982. Exposures resulting from the nuclear explosions. In: Ionizing Radiation: Sources and Biological Effects, 1982 Report, United Nations, New York, pp. 211-248.

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

169

Trace Metals in the North Pacific - Recent Development of Clean Techniques and their Applications to Ocean Chemistry Hiroyuki TSUBOTA: Seiji NAKAMURAt and Kiminori SHITASHIMA*

Abstract The well established clean techniques were applied to the study on trace metals in the ocean. A new sampler was developed for this purpose. Two different analytical methods were employed for analyzing trace metals. Isotope dilution mass spectrometry was used as a standard method, while a chelex100 column extraction-AAS method was used as a supplementary method. The analytical results from these two methods agreed with each other. Total and dissolved metals in water column were determined at several stations in the North Pacific and marginal seas. In the North Pacific, vertical profiles of Fe, Ni, Cu, Zn, Ag and Cd were nutrient-type, those of Mn, Co and P b were scavenging-type and those of V, Mo and T1 were conservative-type. The vertical distributions of total, dissolved and particulate metals in Japan Trench (about 10000 m deep) are discussed. A large portion of Fe, Mn and P b existed as particulate form in water column. Especially, particulate Fe and Mn increased markedly in deep trench (below 7000 m). These metals were strongly scavenged through the water column, but a large amount of particulate Fe and Mn was supplied from continental shelf and slope to deep trench. In the case of the others metals, particulate form was relatively small, and scavenging did not occur strongly. In particular, Cd, V, Mo and T1 were mostly in the dissolved form in the entire water column, so these metals were very stable in the dissolved form.

1.

Introduction

The shocking results of both intranational (e.g., Murozumi et al., 1976) and international (Brewer and Spencer, 1970) intercomparison exercises have appealed for an urgent need to develop new sampling and analytical techniques for the determination of trace metals in seawater. Since then, some Japanese marine chemists have currently made efforts to eliminate sources of contamination in analytical processes through repetitions of int,ercalibration (e.g., Sugawara, 1978). Meanwhile, Tsubota and Murozumi have felt a need of a good sampler and suitable containers which prevent contamination. Patterson (1974) also had the same opinion as above, and recommended the use of isotope dilution mass spectrometry (IDMS) rather than atomic absorption spectrometry (AAS). In order to obtain the exact 'School of Biosphere Sciences, Hiroshima University, Naka-ku, Hiroshima 730, Japan t Muroran Institute of Technology, Mizumoto-cho, Murora n 050, Japan

170 lead concentration in seawater, Shaule and Patterson (1981) have developed a special clean sampler (CIT sampler). Though the CIT sampler is quite useful, it has a disadvantage that a single hydrocast can collect only one sample from a desired depth. If one wishes to determine the detailed vertical profile of trace metal with this sampling method, more than 10 repetitions of hydrocast are required. Tsubota and his colleagues (1985), therefore, developed a new sampler (called TRAMS sampler hereafter) which can be used for serial operation in a single hydrocast. By employing clean sampling devices and analytical techniques, since 1982, the authors and their colleagues have been able t o determine the concentration of trace metals in oceanic waters. These techniques are described in detail, here. The results of trace elements obtained from the cruises of R / V Hakuho Maru and of other vessels in the North Pacific and its marginal seas are also reported and their oceanographic implications are briefly discussed. ?

5

6

Fig. 1 Schematic diagram of the new water sampler.

A) The sampler is lowering closed. 1 bag, 2 valve, 3 inner cylinder, 4 outer cylinder, 5 weight, 6 releaser of valve being opened. B) Inner cylinder hits the releaser 6 t o close the valve when the bag is filled with sample. C) The sampler proceeds in the downstream direction of the hydrowire. Cutting of the tubing is carried out on the upstream section of the wire. The tubing expands after cutting end. 7 sampler, 8 hydrowire, 9 cutting point.

171

Methods

2. 2.1.

Clean sampling and clean treatments

The TRAMS sampler is light and designed to avoid contamination from hydrowire and surrounding sources, so that it can be used for a serial operation in a hydrocast. It can be lowered closed, opened for sampling a t a desired depth and again closed for retrieval. This sampler is schematically shown in Fig. 1. Seawater is introduced into a bellows type bag made of low-density polyethylene by suction (such as filling a balloon). The materials of the bag were chosen after cautious and repetitive examinations. The polyethylene bag is folded before sampling and is protected inside the methacrylate cylinder. When the messenger hits the trigger, it releases the polyethylene tube (sample inlet). The sampler moves downstream of the hydrowire and the inlet tube upstream. Then, the tube extrendsabout 1 m. Thus, the seawater at about 1 m upstream of the wire is drawn into the bag by its downward extension resulting from the weight. The container to store the sample is also made from the same polyethylene. Surface water samples were taken with a specially designed all-polyethylene sampler (MIT sampler) at the bow of the ship running at 2 knots t o avoid contamination from the ship. Subsurface water samples were usually collected by TRAMS while a CIT sampler was sometimes used. A sample in the sampler was divided into several aliquotes and taken into containers in a clean room on board. Centrifuge, filtration and acidification were also carried out in the clean room. Further treatments for chemical analyses were conducted in land-based laboratories. Sample preparations for IDMS are all performed in ultraclean draft chambers and on benches of Class 0 set in pressurized Class 100 clean rooms build at Muroran Institute of Technology. All reagents and water were purified by repeated subboiling distillation of ulTable 1. Contaminants from reagents used for seawater analysis (ng). (ml) 9.25 NH4OH 5.25 HC104 0.4 Dz-CHC13 20.0 CHC13 15.0 H2 0 33.0 2% 20pl 0.015% $ 3 0 2 1101-11 From room environments Total Contamination level (%) When commercial Ultra Pure reagents used (%)

14 M 20% 60% 0.0013%

Reagent HN03

T1 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 1.8

Ag 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 1.7

Cu 0.002

0.007 0.00 0.003 0.003 0.00 0.00 0.00 0.05 0.09.6 0.45 184

Cd 0.00 0.00 0.00 0.00 0.00

Pb 0.00 0.001 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.021 0.00 0.25 15.3 40

Ni 0.01 0.11 0.00 0.32 0.03 0.02 0.00 0.00 0.06 0.55 0.72 3

Zn 0.001 0.02 0.00 0.10 0.01 0.003 0.00 0.00 0.03 0.166 1.7 167

172 trapure reagents using quartz or Teflon devices. Dithizone-CHC13 was purified by 10 times repeated extraction using purified water of pH = 9. Silica gel suspension water was prepared by the hydrolysis of SiC14. The concentrations of metal impurites in these reagents were determined by the use of isotope dilution surface ionization mass spectrometry, IDMS. Contaminants originating from reagents, water and room environmenh are listed in Table 1. Total contaminations are a t the levels of 0.00-1.7% relative to the lowest concentration in Pacific surface water. If commercial reagents are used, contamination level will rise up t o 184% of the amounts in the sample. Samples for column extraction-AAS method were treated in a booth of Class 0 set in a pressurized Class- 1000 clean room at Hiroshima University. All the liquid reagents and water used were purified by repeated subboiling distillation of ultrapure reagents. The contaminant levels were monitored by the use of IDMS. 2.2.

Analytical methods

Two methods were employed for analyzing trace metals. IDMS was used as a standard method, while a chelex-100 column ext,raction-AAS method was used as a supplementary method. Analytical results obtained by AAS were compared with those obtained by IDMS to confirm the consistency of both the methods. Isotope dilution m a s s spectrometry When the isotopic equilibrium is reached in sample solution with a known amount of spike added, the following relation exists between the molar amounts of the isotopes:

iN/i" = ( N , x

ifn

+ N, x

y S ) / ( N nx

iIfn

+ N, x

iIfs)

(1)

where i N and i'N represent the molar amounts of isotopes i and isotope 'i comprising the element under interest, and N , and N , are the molar amounts of the element in the sample and of the added spike. Terms, i f , and i f s are the isotopic abundance of the isotope i in the sample and the spike, respectively. Thus, the amount of N , can be determined by measuring the i N / i'N ratio if the other parameters are known. When a composite spike solution of 61Ni, 65Cu, 68Zn, lo7Ag, '16Cd, 203Tl and 206Pb is used, seven i N / i ' N equations hold a t the same time for 58Ni/61Ni, 63Cu/65Cu, 66Zn/68Zn, 1ogAg/107Ag,114Cd/'16Cd, 205Tl/203Tland 208Pb/206Pb ratios. By measuring these isotopic ratios, N N ~Nc,, , Nz,, A T ~ g Ned, , N T ~and Npb can be simultaneously determined. The spikes used were imported from the Oak Ridge National Laboratory, Tenn., U. S. A. Each spike was dissolved in a 5% H N 0 3 solution. The isotopic abundance and concentration of every stock spike solution were normalized and standardized t o the reference material of known isotopic abundance and concentration. Every working spike solution was prepared by diluting the stock solution with a 5% H N 0 3 solution. A composite spike solution was also prepared by mixing the stock solutions and its isotopic abundance and concentration were measured by mass

173 spectrometry. The mixed spike solution was used only when the composite spikes did not cause any mutual contamination. A Hitachi RMU-6 mass spectrometer equipped with a rhenium single filament as an ionization device can detect an emitter current of 1O-l' A emitted from 10-l' to g of each element. T h e 'N/"N ratio is measured with an error of less than 1% in the coefficient of variation. This means that less than ng amounts can be simultaneously determined within an error of 1%for those metals. The seawater samples spiked a t pH = 1 were kept stand for several weeks t o reach the isotopic equilibrium, then the pH of the samples was adjusted t o 2.0 with addition of NH40H. Silver and copper were extracted into 10 ml of a 0.0013% dithizone-CHC13. Extraction coefficient of silver checked by the double spike ("'Ag and "'Ag) method, was 99.5% for 1 kg of seawater treated. This CHC13 solution was transferred t o a second separatory quartz or Teflon funnel. Nickel, copper, zinc, cadmium, thallium and lead remaining in the acid solution were further extracted into another 10 in1 aliquot of dithizone-CHC13 solution under pH 8.5. T h e contents in both separatory funnels were combined and all the metals were extracted into 6.0 ml of a 7.0 M HN03 solution. The solution is evaporated t o dryness. Finally the residue is dissolved in a mixed solution of a 2% H3P04 and 60 pl of a 0.03% silica gel suspension water, an aliquot of the dissolved residue being applied to the mass spectrometer. For the simultaneous determination of nickel, copper, zinc, silver, cadmium, thallium and lead, emitted ion beams of Ag+, T1+, C u t , Cd+, P b + , Ni+ and Zn+ successively gain the maximum intensity with increasing temperature of the ioii-

5 a 10-l3 0 H

I

I

1.0

I

1

1.2

1.4

I

1.6

I

1.8

I

2.0

I

I

2.2

2.4

I

Filament Current (A)

Fig. 2 Emission of mono-valent ion beam from each 0.1 ug of metal loaded on the Re-ionization filament.

Table 2.

Silver isotopic abundance of lo9Ag and lo7Ag spikes, commercial silver metal, Japanese rock and plant standards, and N.B.S.f Orchard Leaves.

Ag laoded on the Filament ionization filament current Material (4 (A) 1.0 1.027 Commercial 0.50 1.270 silver matal 0.10 1.190 0.05 1.250 0.01 1.270

Granodiorite J G - 1 Basalt JB-1 Pepper-bush Orchard Leaves lo7Ag spike

lo9Ag spike

a

Ial"Ag+

lo7Ag+ ion beam (10-13 A) 5.7 4.3 4.8 7.5 1.5

0.9209 0.9196 0.9212 0.9204 0.9204

n

'

15 9 11 13 13

Isotopic abundance C.V.

(%I 0.1 0.1 0.1 0.2 0.2

0.127 0.207 0.097 0.104 1.0 0.50 0.10

1.270 1.240 1.360 1.300 1.290 1.220 1.260

7.8 14.7 2.2 5.1 13 45 3.3

0.9204 15 11 0.9217 0.9187 15 0.9245 9 0.0 1806 7 0.01797 11 0.01817 7 O.R.N.L." certified

0.2 0.2 0.2 0.1 0.5 0.2 0.4 value

1.0 0.50 0.10

1.260 1.190 1.330

160d 44d 13d

135.43 9 136.26 7 136.69 7 O.R.N.L. certified

0.2 0.3 0.5 value

lo7Ag

109ag

0.5206 0.5209 0.5205 0.5207 0.5207 Ave. 0.5207 f0.0002 0.5207 0.5204 0.5212 0.5196 0.9823 0.9823 0.9822 0.9822 f0.0005 0.0073 0.0073 0.0073 0.0074 f0.0005

0.4794 0.4791 0.4795 0.4793 0.4793 0.4793 f0.0002 0.4793 0.4796 0.4788 0.4804 0.0177 0.0177 0.0178 0.0178 f0.0005 0.9927 0.9927 0.9927 0.9926 f0.0005

denotes the intensity of lo9Ag+ ion beam observed. Coefficient of variation in measurements. Intensity of Iel"Ag+. Oak Ridge National Laboratory, Tenn., U. S. A. National Bureau of Standards. U.S.A.

' Number of I.'09Ag+/I.107Ag+ measured. f

I.'O9Ag+ a I.IU-'Ag+ -

4 cL

P

175 ization filament, as illustrated in Fig. 2. This means ' N l i ' N ratios for above seven elements can be measured without mutual isobar effect. The detailed procedure for these metals are described in Murozumi et al. (1978), Murozumi and Nakamura (1980) and Murozumi (1981). Here, analytical details for silver are only described. By loading 1-0.01 p g of silver nitrate on the Re ionization filament, 10-11-10-12 A of Agf ion beam was emitted. The 1ogAg/107Agisotopic ratio was determined with the accuracy of 0.1-0.2% for natural silver and 0.2-0.5% for the spike in the coefficient of variation as shown in Table 2. The detection limit is estimated t o be of the order of 10-13-10-15 g Ag. The added amount of the lo7Ag spike was kept constant for all measurements, so that from the intensity of lo7Ag mass spectrum, it can be judged whether or not the sample preparation procedure and mass running were properly performed. The 1ogAg/'07Ag ratios in the measurements were between 1.0 and 0.01 which enabled us to calculate the reliable value of N A based ~ on equation (1). In the duplicate analyses, there existed a linear relationship between the measured N A values ~ and the amount of sample taken, intercepting zero in the diagram. This verified that no contamination occurred during the analytical procedure. Silver concentrations in some environmental reference materials together with those in Muroran coastal sea water are listed in Table 3. Chelex-100 column extraction-atomic absorption spectrometry

Two aliquots were taken. One was filtered with Nuclepore filter of 0.4 p m pore size; the other was not: The difference in the metal concentrations between unfiltered and filtered samples was considered as the concentration of particulate metal. Chelex-100 resin is able t o retain selectively trace metals (Riley and Taylor, 1968), such as vanadium, manganese, iron, nickel, cobalt, copper, zinc, molybdenum, cadmium and lead from seawater. The procedure used here is almost the same as those by Kingston and his coworkers (1978), except that a thorough purification of the resin was carried out. Since the resin dose not retain colloidal and organically bound metals, it is necessary to decompose those forms of metals prior to preconcentration. For this purpose, seawater samples were acidified to pH lower than 2 with a nitric acid and kept for more than 2 months before analysis. To carry out constant-column operations, particle size of the resin was selected by elutriation with deionized water for greater than 210 pm portion from commercially available resin of 50-100 mesh. The resin was purified by successive column extraction method with a 2 M nitric acid, sub-boiled deionized water (SDW), 2 M sodium hydroxide and SDW for 20 times, and then with a 2 M nitric acid, SDW, 2 M aqueous ammonia and SDW for 20 times. By this washing method, the impurities in the resin were removed t o a level under the detection limit. The purified resin of 30 ml (ammonium form) was taken into a 164 Teflon tube of 240 mm in length (column dimension is 164x150 mm in ammonium form), and further washed with 200 ml of a 2 M nitric acid, SDW, a 2 M aqueous ammonia and SDW for 5 times. The column was preconditioned by passing 100 ml of a 2 M nitric acid, 200 ml of SDW and 100 ml of a 0.1 M ammonium acetate solution (pH 5.5)

Table 3. Determination of silver in some environmental reference materials and sea water. Sample lU7Agspike added (nmol) taken I.lo7Ag+ (g) lo7Ag "'Ag (10-13k) Orchard Leaves 0.1866 5.018 0.090 17 0.3414 3.360 0.061 1.3 Samplea

IS1"Ag+ I.'"''Ag+ measured 0.0226 0.0299

n 7 9

Silver C.V. found (%) (ng) 0.3 5.27 0.4 9.50

Silver concn. (PPb) 28.0 27.7

Average (ppbj 27.9h0.2

Pepper-bush

0.3238 0.6061

3.351 3.304

0.060 0.059

0.31 3.7

0.0321 0.0448

5 9

0.2 0.1

11.0 21.0

34.0 34.0

34.3&0.3

JG-1

0.1789 0.3507 0.6045 0.9848

1.199 1.663 1.194 1.158

0.022 0.030 0.022 0.021

1.8 3.4 1.4 0.34

0.0341 0.0406 0.0703 0.104

13

0.3 0.5 0.3 0.5

4.51 8.99 15.2 24.5

25.0 25.5 25.1 25.8

25.4h0.4

7 9 7

JB-1

0.1669 0.3160

1.020 1.114

0.018 0.020

3.9 1.0

0.0467 0.0671

13 7

0.4 0.1

6.95 13.3

41.4 41.2

41.3f0.1

Muroran coastal sea water

1837 759.2

1.236 1.366

0.022 0.025

0.1 0.1

0.0319 0.0251

7 9

0.3 0.3

4.09 2.24

2 . 2 3 ~ 1 0 - ~2 . 5 & 0 . 4 ~ 1 0 - ~ 2.94~10-~

177 successively. To the acidified 1 1 sample, 1 ml acetic acid was added and the pH was adjusted to 5.3-5.5 with aqueous ammonia. The solution was passed through the Chelex100 column a t a flow rate of 5 ml/min. The column was washed with 200 ml of 1 M ammonium acetate solution (pH 5.5) a t a flow rate of 2.5 ml/niin t o remove alkali and alkaline earth elements. followed by SDW. T h e metals adsorbed on the column were eluted with 80 ml of a 2 M nitric acid into a quartz dish at a flow rate of 1 ml/min. The eluate was evaporated to dryness on a hot plate in the nitrogen stream clean box, and then the residue was dissolved and diluted with a 2 M nitric acid to ca. 10 ml by weighing. The metals in this solution were determined by GFAAS and inductivity coupled plasma atomic emission spectrometry (ICP-AES). The column was regenerated for reuse in the preconcentration step, by successive loading of 200 ml of SDW, a 2 M aqueous ammonia, SDW, a 2 M nitric acid, SDW and 100 ml of 0.1 M ammonium acetate solution (pH 5.5) at a flow rate of 5 ml/min. The column was washed in the same manner as for a new column when it was not used for more than 3 days. Analytical results obtained by AAS were occasionally compared with those obtained by IDMS. The comparison of analytical results for copper in the Pacific water, taken by TRAMS samplers in 1982, is illustrated in Fig. 3 , as an example.

3.

Results and Discussion

Seawater contains micro-organisms and particulates, so t h a t complete dissolution of particulates is necessary for obtaining total concentration of trace metals in seawater. To achieve this, the seawater sample was divided into two or three aliquots. Each aliquot was treated by different methods as shown in Table 4. Those include direct extraction with dithizone-CHC13, addition of HN03 and HC104 and subsequent evaporation to dryness, and irradiation of 800 W UV lamp for 6 hours in the presence of HZ02. After these treatments, the aliquots underwent the chemical procedures described above. The solid/solution partitioning of metals in seawater was also examined. Particulate matters were separated either by filtration or centrifuge. The seawater sample was filtered through a Nuclepore filter, and the filter was dried and weighed. Aliquots of the filtrate were treated by three different methods as just mentioned above. Alternatively, particulates in the seawater sample were also separated by centrifuge using polycarbonate tubes, under 10000 rpm a t 15"C, for one hour. The particulates were dissolved into a mixture of HN03, HC104 and HF, and the trace nietals in the solution were determined. Three different pretreatments give the same values to each metal as for the total concentration. For all the metals, the sum of a concentration in the filtered particulates and that in the filtrate exactly agrees with the value of total concentration in the seawater sample. This is also true for the case of centrifuge. T h e metal partition coefficient between the solids and the solution is different for each other metal. The use of a centrifugal machine, under such a condition applied here, can collect more solids than with Nuclepore filter, Table 4. This fact is clearly supported by the calculation based on Stoke's Law. The most of particles larger than 0.4 p M (pore size of the filter) can be separated from the solution by a mild

Table 4. Determination of thallium, copper, cadmium and lead in particulates and solution in a Muroran coastal seawater sample taken on March 20, 1979, by IDMS

Pretreatment Without filtering

WPPt) WPPt) Cd(PPt) Pb(PPt) Measured Average Measured Average Measured Average Measured Average 1 2 .8

Direct extraction into H2Dz-CHC13 12.7

12.8f0.1

12.8

284 289 290

28763

45.8 45.8 46.0

45.9f0.1

69.4 67.1 69.0

68.5f0.9

Extraction after HN03-HC104 treatment and evaporation

12.7 12.8

12.8f0.1

309 288

299f9

45.7 46.3

46.0f0.3

69.0 72.3

70.7A1.6

Extraction after UV-irradiation

12.7 12.6

12.7f0.1

285 295

290f5

45.8 46.1

45.9f0.2

70.0 71.8

70.9f0.9

Average of above Filtered through a Nucleopore filter

12.8fO.l 12.7

Direct extraction into H2Dz-CHC13 12.4

12.660.2

12.6 12.8 12.5

12.6f0.2

Extraction after HN03-HC104 treatment and evaporation Extraction after UV-irradiation In particulates Centrifuge (10000 rpm for 1 hr) Solution Particulates

12.4 12.7

12.6f0.2

293f6 268 272 266 259 267 260 264

267f7 263f4

262f2

46.0f0.3 45.6 45.4 45.6 45.9 45.5

0.14

25

45.5 45.4 1.20

12.4 0.32

227 71

45.4 1.94

c

4 00

45.2f0.2 45.7f0.2

45.5f0.1

70 5 1 . 1 29.8 31.4 31.3 30.4 31.0 29.4 31.4 39.3

19 51.6

31.1f0.5 30.7f0.3

30.4f1.0

179 centrifugal condition of several thousands r.p.m. The calculating result is consistent with experimental results obtained by the use of radioactive tracers (Watari et al., 1966). Shitashima and Tsubota (1990) also discussed the comparison between centrifugal and filtration methods and their elemental differences for coastal and estuarine areas. IDMS, although it is matter time-consuming methods, was applied t o established the true concentration of metals for samples taken from most parts of the North Pacific and the marginal seas: between Japan and Hawaii, off Hawaii Islands, between Hawaii and Tahiti, off Guam and Ponape Islands, off Mexico, at Mariana and Japan Trench, Gulf of Alaska, the Sea of Japan, East China Sea and the Bering Seas. The details of these results and discussion will be reported elsewhere, and only the vertical distribution of dissolved and particulate metals in the Japan Trench are mentioned here as typical examples. The vertical distributions of thallium, silver, copper, cadmium, lead, nickel and zinc in the Japan Trench (AN-l), taken by the CIT sampler and obtained by the IDMS, are shown in Table 5. Those of iron, manganese, cobalt, cadmium, molybdenum and vanadium, taken by TRAMS samplers and obtained by the chelex-100 column extraction-AAS method, are displayed in Fig. 4. In all of the areas investigated, the vertical distributions of dissolved metals were Table 5. Vertical distribution profile of heavy metals in Japan Trench (29"05'N 142"51'E) determined by IDMS. ng.kg-'. Sampling depth (m) Surface

384

1453

1983

4880

9773

Seawater Particulates Solution Seawater Particulates Solution Seawater Particulates Solution Seawater Particulates Solution Seawater Particulates Solution Seawater Particulates Solution

T1

Ag

Cu

Cd

Pb

Ni

Zn

13.3 1.09 12.6 13.4 0.91 12.4 13.3 0.77 12.5 13.2 0.69 12.4 13.6 0.60 13.0 13.8 0.85 12.9

0.46 0.13 0.26 1.15 0.32 0.83 5.46 0.60 4.77 5.78 0.64 5.19 5.27 0.21 5.06 4.98 0.12 4.83

42.0 5.28 33.4 54.1 12.7 39.6 104 12.1 95.6 127 13.7 115 182 6.00 176 189 9.14 181

1.24 0.62 0.65 18.8 0.42 18.3 97.5 2.39 95.4 96.0 3.19 93.3 87.1 2.00 85.1 96.8 1.01 91.1

16.8 2.08 14.1 15.6 5.15 10.9 6.00 1.64 4.43 7.15 1.78 5.51 4.61 1.05 3.56 4.59 0.96 3.66

153 41.1 104 255 100 155 561 40.4 516 503 52.7 450 606 30.0 576 570 20.7 551

19.0 6.67 11.1 69.4 34.7 35.7 546 33.4 505 539 26.9 513 540 22.0 518 597 20.3 575

180 almost similar t o those in previous observations from the North Pacific (Bruland, 1980; Landings and Bruland, 1980; Schaule and Patterson, 1981; Collier, 1984, 1985; Flegal and Patterson, 1985; Martin et al., 1989). Vertical profiles of cadmium, nickel, zinc, copper and iron were nutrient-type, those of cobalt, manganese and lead were scavenging-type and those of molybdenum, vanadium and thallium were conservative-type. It was newly found that silver exhibited nutrient-type profile in the same manner as zinc. Among the nutrient-type metals, the proportion of particulate to dissolved metal was extremely small for cadmium and was fairly small for nickel, zinc and copper. This fact indicates that these metals are relatively stable in the dissolved form and are not so strongly scavenged through the water column. For unknown reason, an appreciable amount of particulate silver is distributed in the mid depth (around 2000 m). A large portion of iron (20-80%) existed as particulate form from the surface to the bottom reflecting the insoluble nature of this metal (Martin et al., 1989). Especially, a remarkable increase of particulate iron was observed below 7000 m. Yamada and his coworkers (1983) measured natural radioactive tracers and heavy metals in sediment collected at Izu-Ogasawara Trench, and reported that the sedimentation rate a t the trench was relatively fast (over 1 cm/kyr) compared to that a t oceanic basin. They concluded t h a t it was due t o the slipping-down of particles from the continental shelf. The particulate Fe/T\iln ratio in the Japan Trench water (about 9) was nearly equal to Fe/Mn ratio in the surface sediment at Izu-Ogasawara Trench (about 11; Yamada et al., 1983). Therefore, the high concentration of particulate Fe in the Trench water mass may be caused by the supply of a large amount of particulate matter along the continental shelf and slope to the deep Trench. In the case of scavenging-type metals, fair amounts of particulate manganese

0

100

IDMS

200

300

Fig. 3 Comparison of copper concentrations determined by IDMS and

AAS (ng/kg). Samples were collected with TRAMS in the North Pacific, 11"59'N, 152"30'E, Feb. 1982. The straight line indicates 1:l relation.

181

(

C

200(

2000

400C

E

+

W

s3 n

4000 0

BOW

10000

C

+

+++ 60

o

4000

4000

o

100

loo00

ot

so

0000

+

-a

+ O + O

0

8

8

8

0

2

BOW

8OOo

o

1

0

+ o

8

+ o

lOOOc

w

160 I

s

+

2000

4000 -

0

W

+0 so

0

2000

E

2000

0000

++ 8000

00

ot

4000

0

5

8 n

BOO0

BOO0

BOO0 ~

+

0000

0 0 0 10000

@

80 01

10000

so

eooo

+O

@ + O

so

10000

Fig. 4 Vertical profiles of iron, manganese, cobalt, cadmium, molybdenum and vanadium at the Japan Trench (29"05'N, 142"51'E), determined by AAS. Circle: unfiltered or total, Cross: filtered or dissolved.

3

182 and lead were distributed in the surface t o the mid depth and a large increase of particulate manganese observed in the deep Trench. Particulate lead is mainly supplied t o the surface water through atmospheric input. Murozumi and his coworkers have reported lead isotopic composition in the smog collected in Los Angeles as 0.8666 for 207/206 and 2.107 for 208/206 (Murozumi et al., 1969), and those in Hokkaido as 0.8606 for 207/206 and 2.071 for 208/206 (Murozumi et al., 1982). As is evident from Table 6, this input is considered t o be anthropogenic, especially showing aerosol origin. The isotopic composition of lead changes notably from the surface, where a large portion of lead exists as particulate form, to 4000 m depth, where the proportion of particulate t o dissolved metal is fairly small, and shows little variety from 4000 m t o the bottom and marine sediments. Bruland (1983) pointed out that Mn enrichment in surface water was mainly contributed by supply from particulates derived from river or shelf sediment, and reduction and/or diffusion from such particulate in oxygen minimum layer. The manganese concentration maximum at AN-1 was associated with the oxygen minimum layer, and relatively high concentration of particulate manganese observed in the same layer. It is considered that the source of manganese to surface water was attributable to river input or diffusion from shelf sediments. Particulate manganese in the deep Trench is probably supplied along the continental slope in the same manner as particulate iron. The vertical profiles of cobalt were similar t o that of manganese in northern North Pacific except in deep trenches (Shitashima and Tsubota, unpublished), but particulate cobalt hardly existed at all through the whole water Table 6. Vertical distribution profile of lead isotopic composition at Japan-Trench (29’05” 142’51’E). P b Isotope composition Depth 2071206 2081206 2061204 (In water column, m) 0.8544 2.120 18.0 Surface 389 0.8519 2.091 18.0 2.099 18.0 1453 0.8494 2.066 17.5 1983 0.8478 2.066 18.6 4880 0.8369 9773 0.8354 2.063 18.6 (In bottom sediment, cm) 3-4 0.8393 2.087 18.6 2.086 18.7 6-7 0.8384 2.084 18.6 8-9 0.8398 2.082 18.6 81-82 0.8389 2.083 18.6 139-140 0.8384 2.081 18.5 203-204 0.8401 2.096 18.6 476-478 0.8420 2.087 602-603 0.8404 18.6

183 column in all stations. It may be inferred t h a t t h e cause of cobalt profile was due t o dissolution of manganese oxide derived from river in oxygen minimum layer. I n t h e case of conservative-type metals, t h e proportion of particulate t o dissolved metal was negligibly small for molybdenum a n d vanadium, whereas a n appreciable portion of thallium existed as particu1at)e form. This is consistent with t h e fact t h a t molybdenum a n d vanadium are oxyacid elements which are soluble in seawater around pH 8 in their highest valency state. It has been reported that thallium exists as a simple monovalent cation in seawater a n d as insoluble trivalent s t a t e in manganese nodule (Flegal e t al., 1989). This is also estimated from t h e fact t h a t t h e concentration factor ([M mol kg-l]nodule/[M mol l-llseawater) of thallium in manganese nodule (usually proportional t o t h e hydrolysis constants of metals) is too high when assuming as monovalent a n d is too low when assuming as trivalent. Therefore, thallium should be fixed in particulate m a t t e r in t h e trivalent state.

References Brewer, P. G., and D. W. Spencer, 1970. Trace element intercalibration study. W. H. 0 . I. Report, No. 70-62. Bruland, K. W., 1980. Oceanographic distributions of cadmium, zinc, nickel and copper in the north Pacific. Earth Planet. Sci. Lett., 47, 176-198. Bruland, K. W., 1983. Trace elements in sea water. In: Chemical Oceanography, Vol. 8, J. P. Riley and R. Chester (ed.), Academic Press, London, pp. 157-221. Collier, R. W., 1984. Particulate and dissolved vanadium in the North Pacific Ocean. Nature, 309, 441-444. Collier, R. W., 1985. Molybdenum in the North Pacific Ocean. Limnol. Oceanogr., 30, 1351-1354. Flegal, A. R., and C. C . Patterson, 1985. Thallium concentration in seawater. Mar. Chem., 15, 327-331. Flegal, A. R., S. Sanudo-Wihelmy and S. E. Fitzwater, 1989. Particulate thallium fluxes in the northeast Pacific. Mar. Chem., 28, 61-76. Kingston, H. M., I. L. Barnes, T. J. Brady and T. C. Rains, 1978. Separation of eight transition elements from alkaline earth elements in estuarine and seawater with chelating resin and their determination by graphite furnace atomic absorption sepectrometry. Anal. Chem., 50, 2064-2070. Landing, W. M., and K. W. Bruland, 1980. Manganese in the North Pacific. Earth Planet. Sci. Lett., 49, 45-56. Martin, J. H., R. M. Gordon, S. Fitzwater and W. W. Broenkow, 1989. VERTEX: phytoplankton/iron studies in the Gulf of Alaska. Deep-sea Res., 36, 649--680. Murozumi, M., T. J. Chow and C. C. Patterson, 1969. Chemical concentration of pollutant lead aerosols, terrestrial dusk and sea salts in Greenland and Antarctic snow strate. Geochim. Cosmochim. Acta, 33, 1247-1294. Murozumi, M., S. Nakamura and K. Ito, 1976. Isotope dilution mass spectrometry of copper in sea water. Bunseki Kagaku, 25, 706-710 (in Japanese). Murozumi, M., S. Nakamura, T . Igarashi and H. Tsubota, 1978. Isotope dilution-surface ionization mass spectrometry of copper, cadmium, and lead in sea water. Nippon Kagaku Kaishi, 565-570 (in Japanese). Murozumi, M., and S. Nakamura, 1980. Isotope dilution mass spectrometry of copper, cadmium, thallium and lead in marine environment. In: Isotope Marine Chemistry, E. D. Goldberg, Y. Horibe and K. Saruhashi (ed.), Uchida Rokakuho, Tokyo, pp. 439-

471, Murozumi, M., 1981. Isotope dilution surface ionization mass spectrometry of trace constituents in natural environments and in the Pacific. Buiiseki Kagaku, 30, S19-S26. Murozumi, M., S. Nakamura and K. Yoshida, 1982. Impacts of aerosollead to natural ecosystems. Nippon Kagaku Kaishi, 1479-1484 (in Japanese). Patterson, C., 1974. Lead in seawater. Science, 183, 553-554. Riley, J. P., and D. Taylor, 1968. Chelating resin for the concentration of trace elements from sea water and their analytical use in conjunction with atomic absorption spectrophotometry. Anal. Chim. Acta, 40, 479-485. Schaule, B. K., and C. C. Patterson, 1981. Lead concentrations in the Northeast Pacific: evidence for global anthropogenic perturbation. Earth Planet. Sci. Lett., 54, 97--116. Shitashima, K., and H. Tsubota, 1990. Transport of heavy metals into and out of the Set0 Inland Sea, Japan. Geochem. J., 24, 283--293. Sugawara, K., 1978. Interlaboratory comparison of the determination of mercury and cadmium in sea and fresh waters. Deep-sea Res., 25, 323-332. Tsubota, H., 1985. The distribution and behavior of heavy metals in the ocean. In: Ocean Characteristics and their Changes, I<. Kajiura (ed.), Koseisha Koseikaku, Tokyo, pp. 225-236 (in Japanese). Watari, K., H. Tsubota, T. Koyanagi and M. Izawa, 19GG. Content,ration of radionuclides in sea water by “metal sulfide-ion exchange resins.” J. At. Energy SOC.,Jpn., 8, 182185. Yamada, M., H. Kitaoka and S. Tsunogai, 1983. A radiochemical study of sedimentation onto the Japan Trench floor. Deep-sea Res., 30, 1147-1156.

Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

185

Distribution of Dissolved Organic Nitrogen in the North Pacific Ocean Yoshiaki MAITA4and Mitsuru YANADA"

Abstract The concentration and composition of dissolved organic nitrogen (DON) in the North Pacific Ocean were investigated to determine the vertical and horizontal variations of DON. The significance of dissolved organic nitrogen in the geochemical cycle of nitrogen was discussed in relation to the hydrographical conditions, biological production and decomposition. Seawater samples were collected from the surface to 2,000 In depth a t 33 stations of the subarctic and subtropical regions of the North Pacific Ocean and the DON concentrations were investigated by high temperature catalytic oxidation method. The DON concentrations in the upper waters (0-200 m) of the subtropical region were slightly higher than those in the subarctic region, whereas the concentrations in the middle (200-1000 m) and deep (> 1000 m) waters were nearly the same. Moreover, the vertical DON profiles were almost similar to each other in all the stations: the average DON concentration gradually decreases with depth from 6.461.9 pg-at.N/l a t the surface to 2.9f1.9 pg-at.N/l at and below 200 m, where the DON concentrations were almost constant with depth. In the upper water, DON existed as labile nitrogen compounds which can easily be decomposed by oceanic bacteria, i.e., low molecular weight organic nitrogen compounds. In the middle and the deep waters, however, DON existed more as biochemically refractive nitrogen compounds, i.e., higher molecular weight organic nitrogen compounds. Despite the fact that dissolved inorganic nitrogen (DIN) shows the large horizontal and vertical variations in the ocean, the spatial variation of DON is small both vertically and horizontally. This suggests that a significant source of DIN is not, DON but particulate organic nitrogen (PON), and the DON may not be actively involved in t,he nitrogen cycle in the North Pacific Ocean.

1.

Introduction

The major nitrogen species present in seawater are: particulate organic nitrogen (PON), dissolved inorganic nitrogen (DIN) and dissolved organic nitrogen (DON). In order to better understand the nitrogen cycle in the sea, it is necessary to know the concentrations and behaviors of those nitrogen species. Recent studies *Division of Marine Biochemical Science, Research Institute of North Pacific Fisheries, Faculty of Fisheries, Hokkaido University, Minatocho, Hakodate 041, J a p a n

186

c" 65NI-

,I3p$,

,

, &OEr,

0

0888-23

A

0889-28

, ,

,15pE-60€

170E

180

l70W

I OW

----im--P

150W

140W

I3OW-

0887-18

Fig. 1. Location of sampling stations in the North Pacific Ocean. have shown the geochemical significance of PON and DIN as well as their behavior (Sharp, 1983). However, there is a significant controversy regarding DON in seawater in terms of distribution and geochemical significance. Suzuki et al. (1985) and Sugimura and Suzuki (1988) claimed that the DON concentration in open ocean waters is high (- 40 pM), if it is determined by the high temperature catalytic oxidation (HTCO) method. The value is several times higher than those obtained by the traditional persulfate oxidation (wet digestion) method. Their results suggest not only the analytical incompleteness of the wet digestion method for DON, but a different view on the chemical nature and role of DON as the energy source for microorganisms living in seawater. However, Walsh (1989), using a new high temperature combustion method, could not obtain such a high DON concentrations in the North Pacific Ocean. Furthermore, Schlitzer (1989) reported that the nutrient and carbon cycle models using the d a t a of Sugimura and Suzuki (1988) were not feasible for the North Atlantic Ocean. In a recent paper, Williams and Druffel (1989) pointed out several problems in order to clarify the differences in the DON concentration and distribution pattern in the North Pacific Ocean. They claimed the necessity for the new analytical techniques of identifying the nitrogen containing conipounds and their origin and biochemical stabilities among others. Hence, in this paper, we have compared the two methods used for DON determination, the HTCO and persulfate oxidation methods, to evaluate the DON concentration present in seawater accurately. Furthermore, we have analyzed water samples collected from the subarctic and subtropical regions of the North Pacific Ocean. The vertical and horizontal distributions of DON are discussed in relation to the hydrographical parameters, and biological production and decomposition.

187 Table 1. Sampling date and depths and type of filters used in the determination of dissolved nitrogen compounds in seawater samples collected from the North Pacific Ocean. Cruise No.

Date

0887-18 OS88-23 0889-28 OS89-43 KH85- 2 KH86- 3 KT87-18 KT89-10

Jun.23, Jun.13-Aug. 6, Aug. 2-Aug. 4, Jun. 2-Jun.10, Apr. 4-May 7, May 6, Nov.20-Nov.24, Jul. 4-Jul. 7,

1987 1988 1989 1989 1985 1986 1987 1989

Sampling depth (m) 0-2000 0-2000 0-2000 0- 900 0-2000 0-3000 0-1500 0-3000

Filter used

Description

HA type GF/F GF/F GF/F HA type GF/F HA type GF/F

in laboratory on board for TDN and NO3 on board for NO3 in laboratory in laboratory in laboratory in laboratory on board for TDN

Table 2. Analytical conditions followed for the determination of dissolved total nitrogen by high temperature catalytic oxidation method for dissolved total nitrogen. Instrument Seawater volume injected Temperature distribution in furnace Catalyst Flow rate of carrier gas Final form of nitrogen Detection of nitrogen Oxidizing agent Absorbent solution volume Cell size Wavelength Time of measurement

2.

Sumigraph N-200 500 p1 630-690°C 3% Pt on A1203 or Manganese

300 ml/min. NO2 Spectrophotometric 5% KMn04 5ml 5% HzS04 5ml 5 ml 2 cm 545 nm 6 min. (After Maita and Yanada, 1990)

+

Materials and Methods

Seawater samples were collected from the surface to 2,000 m depth at 33 stations of the North Pacific Ocean (Fig. 1)during the cruises of R/V Hakuho Maru (KH852, KH86-3), R/V Tansei Maru (KT87-18, KT89-lo), T / V Oshoro Maru (0387-18, 0888-23, 0889-28) and T / V Hokusei Maru (H089-43). The seawater samples were filtered using pre-washed 45 mm membrane filters (Millipore HA) or 45 mm glass fiber filters (Whatman G F / F ) which were previously heat,ed for 3 h at 450°C.

188 The filtrates were analyzed either immediately on board or stored in polyethylene bottles below -20°C (Table 1). T h e inorganic nitrogen compounds (NOs--N, NOz--N, NH4+-N) were analyzed according to the method described by Stricklarid and Parsons (1972) using an Auto Analyzer (Technicon, Model 11). The total dissolved nitrogen concentration (TDN) was determined by High Temperature Catalytic Oxidation method (HTCO) using Sumigraph N-200 (Sumitomo Chemical Industries Co.). Table 2 shows the analytical conditions followed by Maita and Yanada (1990). The DON concentrations were calculated by subtracting the DIN concentration from the TDN concentration. Moreover, two methods in measuring the TDN concentrations in seawater, the HTCO method and the persulfate oxidation method (Sol6rzano and Sharp, 1980; Parsons et al., 1984), were compared using seawater samples. Molecular weight fractionation and determination of dissolved total amino acids (DTAA) were also carried out using seawater samples collected at one station (KH86-3). Seawater samples were filtered with pre-treated Whatman G F / F filters. The filtrates were stored in polyethylene bottles at room temperature after adding HgC12 (final concentration: l o p 4 M). Size fractionation was performed using three kinds of ultrafilters with molecular weight cut-offs as indicated: YM2, 1,000; YM10, 10,000; YM100, 100,000 (Amicon Co.). The filtrate from each fractionation was hydrolyzed in 6 N HC1 at 110°C for 24 h. Hydrolyzates were concentrated and purified by ion-exchange chromatography according to Tada and Maita (1988). The DTAA in the samples were determined by High Performance Liquid Chromatography (HPLC) using o-phthalaldehyde as the fluorogenic agent.

3. Results and Discussion 3.1.

Re-evaluation of analytical methods for total dissolved nitrogen

The total dissolved nitrogen (TDN) concentrations in natural seawater samples were determined by both high temperature catalytic oxidation and persulfate oxidation methods. The persulfate method used here was a modified one according t o D’Elia et al. (1977) or Koroleff (1983) with a few modification respect to pH and the concentration of persulfate (Sol6rzano and Sharp, 1980). In our analysis, two kinds of catalysts, platinum and manganese, were used for the determination of TDN by the HTCO method. Although the TDN values obtained using the platinum catalyst were slightly lower than those by manganese, the difference is not significant (Maita and Yanada. 1990). The analytical instruments, both Sumigraph N-200 and TN-7 (Yanagimoto Seisakusho Co.) gave the same TDN values (Maita and Yanada, 1990). The results from two station (KT87-18 and OS87-18) are shown in Fig. 2. It is evident that a highly linear relationship exists between the TDN values obtained by the two methods ( T = 0.98), although the HTCO method gives slightly higher value by about 3% on average. This shows that the TDN concentrations obtained by the HTCO and persulfate oxidation methods were not significantly different. In the reports of Suzuki et al. (1985) and Sugimura and Suzuki (1988), the TDN concentrations obtained by the HTCO method were significantly higher than those by persulfate method. T h e difference was assumed to be due to incomplete

189

t

Y=X

/

vl,,,li I

Y=o.97xto.53 r=0.98

n=42

, I , ,

0

10

20

30

TDN b y HTCO (

40

50

60

70

method, ,ug-at.N/1

After

M a i t a and Yanada.1990

)

Fig. 2. Comparison of the total dissolved nitrogen concentrations determined by the high temperature catalytic oxidation (HTCO) and persulfate oxidation method in seawater samples collected from the North Pacific Ocean. oxidation by the persulfate oxidation method of some unknown organic nitrogen compounds in natural seawater. To test this hypothesis, the recovery of nitrogen by the two methods was checked, using some standard organic nitrogen compounds and humic acids which were extracted from natural sediments. T h e recoveries by the two methods were found t o be almost. quantitat,ive for all the nitrogen compounds used (Maita and Yanada, 1990). Possible changes in the TDN concentration during storage were also checked by determining seawater samples kept in polyethylene bottles and stored frozen at -20°C for one month (Maita and Yanada, 1990). The results indicated t h a t the TDN concentrations in frozen samples were relatively lower than concentrations of seawater samples analyzed immediately on board. It is clear from the above that, despite of the difference in the analytical techniques, such as oxidation methods, type of catalysts and analytical instruments, the results obtained do not show any significant difference in the amount of TDN measured in the sample.

190 TDN, p ~ - a t . N / l 10

-

20

30

40

50

6

0

10

DIN. 1 1 g - a t . N I l 20 30 40

DON, y c - a t . N / l

50

-2

0

0

100

r c

500

1000

1500

2000 3000

Fig. 3. Vertical profiles of the total dissolved nitrogen (TDN), dissolved inorganic nitrogen (DIN) and dissolved organic nitrogen (DON) concentrations. The values represent the average concentration for the 33 stations in the North Pacific Ocean; bars represent the standard deviation. 3.2.

Vertical distribution of total dissolved n i t r o g e n a n d dissolved organic nitrogen

Figure 3 shows the vertical profiles of TDN, DIN and DON at 33 stations in the North Pacific Ocean. The TDN concentration increases with depth, from 16f4.6 pg-at./l in the surface waters to a maximum value of 48.9f2.6 pg-at./l at 1,200 m. The coefficient of variations of TDN concentrations in the upper 700 m are higher (22-39010) than those below 800 m (5511%). The vertical profile and variation for DIN almost coincide with those for TDN ( T = 0.99, n = 451). This implies that the major constituent of TDN is DIN. Moreover, since the nitrate nitrogen shows the same profile as TDN and a higher Concentration than nitrite nitrogen (0.23f0.17 pg-at./l) and ammonium nitrogen (1.5f0.6 pg-at./l), it seems that the TDN concentration is largely influenced by the nitrate nitrogen concentration. On the other hand, the DON concentration tends to decrease slightly with depth, from the surface (6.451.9 pg-at./l) t o 200 m (2.9f1.9 pg-at./l), and below 200 m the concentrations were low and almost constant with depth. The vertical variation of DON is quite small as compared t o that of DIN. The concentrations and the vertical profiles of DON obtained in this study for the North Pacific Ocean are very similar t o those reported for open oceans (Sharp, 1983), and disagree with the high concentration (about 40 pM) and the vertical profiles reported by Suzuki et al. (1985) and Sugimura and Suzuki (1988). The analytical conditions of our HTCO method were almost identical to those described in the above literatures, except. for t,he volume of sample injected into

191 D T A A con.,

0

50

nH 100

150

0

0

Percentage, X 20 4 0 60 80

100

100

1000 E L U

n

:2000 3000

4000 I

I

Fig. 4. Vertical profiles of the dissolved total amino acids (A) and the contribution (B, %) of the low molecular weight fraction (LMW; 1,00010,000); intermediate molecular weight fraction (IMW; 10,000-100,000); and the high molecular weight fraction (> 100,000). the analyzer (500 p1 for us and 200 pl for Suzuki et al.). Therefore, it is not clear and curious why they obtained such high DON values. Considering t h a t we have made a detailed examination of the methods used for DON as described earlier, and found no problem with it, there is no reason for us to believe that the high DON values reported by Suzuki et al. (1985) are real. Thus, we conclude, based on our data in the North pacific Ocean, that the DON concentrations in surface waters are only slightly higher than those in deeper waters but never exceed 10 pg-at.11. 3.3.

Molecular weight distribution of a constituent of dissolved organic nitrogen in seawater

Dissolved total amino acids (DTAA) in seawater were separated into three molecular weight fractions: low molecular weight (LMW) fraction, 1,000-10,000; intermediate molecular weight (IMW) fraction, 10,000-100,000; and high molecular weight (HMW) fraction, >100,000. Since amino acids are a major component of organic matter containing nitrogen in seawater, it may be possible to know the molecular weight distribution of DON in the North Pacific Ocean (KT86-3).

192 Figure 4 shows the vertical profile of the DTAA concentration in each molecular weight fraction and the contribution of each fraction t o the total DTAA. The concentration of LMW-DAA is found to be high in the surface waters, showing a decrease with depth to about 20 nM a t 2,000 m depth. The vertical profile of IMW-DAA is generally similar to that of LMW-DAA except t h a t the maximum concentration is observed at a deeper layer (1,000 m depth). In contrast, the HMW-DAA is low in the surface waters and gradually increases with depth. The sum of the LMW- and IMW-DAA (1,000-100,000) constitutes about 80% of the total dissolved amino acid from the surface to 1,000 m and about 30 to 40% below 1,000 m. Tada and Maita (1988) have shown t h a t the concentration of dissolved free amino acids (molecular weight is below 1,000) is also high in surface waters and decreases with depth. Assuming that amino acids or proteinous matter are the major constituents of DON, the DON is present as low molecular weight organic nitrogen in the upper layers and as high molecular weight organic nitrogen in deeper layers. Table 3. Total dissolved nitrogen (TDN), dissolved inorganic nitrogen (DIN) and dissolved organic nitorgen (DON) concentration (z&s.d.) at each layer in the subarctic, transition and subtropical regions of the North Pacific Ocean. Compounds TDN

DIN

DON

Layers

Subarctic Transition pg-at.N /1

Subtropic

(m) 0- 100 100- 200 200- 500 500-1000 1000-3000

20.9f6.1 (n=128) 41.3f7.5 (n= 66) 43.2f6.0 (n= 54) 47.1f3.1 (n= 49) 47.2f2.7 (n= 48)

16.3f4.5 (n=38) 21.1f4.0 (n=14) 31.2f8.9 (n=15) 42.0f7.4 (n=15) 51.0f.3.1 (n= 8)

11.0f3.6 15.5f5.0 22.2f7.8 34.4*8.7 43.4314.8

0- 100 100- 200 200- 500 500-1000 1000-3000

16.956.7 (n=128) 33.1f7.6 (n= 66) 40.6f6.1 (n= 54) 45.0f3.2 (n= 49) 44.0f6.9 (n= 48)

11.3f5.2 (n=38) 17.4f3.9 (n=14) 28.3f.9.3 (n=15) 40.1f6.2 (n=15) 46.6f0.7 (n= 8)

4.7f3.9 (n=44) 11.865.9 (n=18) 20.1f8.7 (n=17) 32.6f.8.5 (n=21) 39.6f4.6 (n=19)

(n=44) (n=18) (n=17) (n=21) (n=19)

0- 100 4.1f2.3 (n=128) 5.0f1.9 (n=38) 6.2f2.0 (n=44) 100- 200 2.551.8 (n= 66) 3.6f0.8 (n=14) 3.7f1.5 (n=l8) 200- 500 2.6f1.9 (n= 54) 2.2f1.8 (n=15) 2.1f1.7 (n=17) 500-1000 2.0f1.5 (n= 49) 1.9f2.7 (n=15) 1.3f1.8 (n=2l) 1000-3000 2.3k2.3 (n= 48) 4.4f3.5 (n= 8) 2.0f1.3 (n=19) Subarctic, transition and subtropical domains were classified in Fig. 5 ( s e e also the text in detail)

193

Fig. 5. Vertical and horizontal variations of the physical properties (temperature, salinity, density and dissolved oxygen) at the meridian (A), 155"E (B) and 50 to 52"N (C). 3.4.

Geochemical significance of dissolved organic nitrogen in the North Pacific Ocean

Relationship with water masses Figure 5 shows a diagrammatic representation of the physical properties of the water masses along either latitude or longitude. According to the definition by

194 Dodimead et al. (1963), along the meridian, the water mass north of 48"N belongs t o the subarctic region and the water mass south of 39"N belongs t o the subtropical region. The water mass between these two water masses is called the transition region (Anma et al., 1990). At the 155"E longitude, the water masses north of 42.5"N and south of 36.5"N belong to the subarctic and the subtropical regions, respectively. In the subtropical region, a strong pycnocline is formed between 200 t o 500 m depth and a sharp dissolved oxygen minimum (< 1 ml/l) is found between 1,200 to 1,500 m depth. The water mass present at > 1,000 m in the subtropical regions ascends t o about 200 ni in the subarctic region. In the subarctic region, a strong pycnocline is not observed at the corresponding level and a dissolved oxygen minimum spreads over a wide range of depths (200-1200 m). All the stations in Fig. 5C, except for that a t 134"W, belong to the subarctic water mass, which has a strong pycnocline between 30 to 100 m depth and a wide minimum dissolved oxygen concentration layer. Table 3 shows the average TDN, DIN and DON concentrations in each layer which has different physical characteristics of temperature, salinity, density and dissolved oxygen. For the surface layer of 0-100 m, both the TDN and DIN concentrations in the subarctic region were higher than those in the southern regions. Inversely, the DON concentrations in the subarctic surface water are slightly lower than those in the southern surface water. The same trend continues for the depths of 100-200 m. The concentrations found a t the subsurface layer in the subarctic are high as well as those found at deep layers in the subtropic. The high TDN and DIN values are likely due to the upwelling of the subtropical deep water to the surface in the subarctic region, considering that the same water mass, which has a density ( u t )of 27.0-27.1 and high N o s , comes up from the southern deep layer to the northern subsurface layer (Fig. 5). Although the TDN and DIN concentrations in the surface water vary from region to region, the DON concentrations below 200 m are invariant. It seems that the DON in the North Pacific Ocean shows only slight vertical variation but it does not seem t o be closely related with the water masses.

Relationshap with biological production The surface layer is the biologically active zone in open ocean waters. In this layer, phytoplankton take up not only DIN but also DON from the seawater (McCarthy, 1972; Schell, 1974; Antia et al., 1975; Sahlisten, 1987). Heterotrophic microorganisms (e.g., bacteria) also consume DON from seawater (Williams, 1975). On the other hand, it is known that both phytoplankton and zooplankton release relatively low molecular weight DON into the seawater (Gagosian and Lee, 1981; Small et al., 1983; Lancelot, 1984). Hence, the concentration and distribution of DON in the shallow waters may depend on the balance between biological uptake and release. Referring t o the distribution pattern of phytoplankton biomass in the North Pacific Ocean reported by Odate and Maita (1989), the biomass of phytoplankton is greater in the subarctic than in the subtropic. This may partially explain the

195 Table 4. Slope (m) of the regression line and the coefficient of correlation ( T ) between parameters: nitrate nitrogen (NO3), dissolved organic nitrogen (DON) and apparent oxygen utilization (AOU), a t each layer in the North Pacific Ocean.

Y-x

NO3 - AOU r n+ 0.84 253

m* 4.8

Upper layer (0-200 m) Deep layer 4.3 0.81 (200-2000 m) Total layer 4.7 0.94 * Units: pg-at.N/ml t Units: pg-at.N/pg-at.N -+ Number of samples

148 401

DON - AOU m* r n+ -0.42 0.39 253 0.43 -0.31

0.18

148

0.41

401

NO3 - DON mt r n+ -3.2 0.61 253 0.1

-3.2

0.04

148

0.52

401

observation that the DON concentrations in surface waters of the subarctic region are slightly lower than those in the subtropic, whereas both the TDN and DIN concentrations in the subarctic region were relatively higher than those in the subtropic. This is because that the higher phytoplankton activity in the subarctic may result in the slightly lower DON . Relationship w i t h biochemical decomposition

The free-living heterotrophic bacteria in seawater can transform DON t o DIN as nitrate nitrogen. Since DON is decomposed by bacteria, and if DON is a dominant source of DIN, the DON concentrations in seawater should show an close relationship with DIN and AOU. Table 4 shows the slope of the regression line and coefficient of correlation between NOS, DON and AOU in each layer. In the upper layer (0-200 m), a positive linear relationship exists between NO3 and AOU and the slope of the regression line is about 4.8 which is statistically significant. On the other hand, a significant correlation also exists between DON and AOU ( r = 0.39, P < 0.01). An inverse linear relationship is found between DON and NO3 in the upper water. The slope which is estimated to be about -3 indicates that about 30% of the regenerated NO3 result from the decomposition of DON. However, in waters below 200 m, the inverse relationship between DON and AOU was no longer found. This suggests that the DON compounds in deeper waters are composed of nitrogen compounds that are not easily degraded. Miyake et al. (1982 and 1985) have deduced that the DOC or DON in seawater must be strongly correlated in a stoichiometric sense with AOU. Their hypothesis was not supported by this study, however. The AOU used for decomposition of DON was only about 10% of the AOU required for regeneration of DIN as estimated based on the values of Table 4. Hence, it appears that a large portion of AOU resulted in the decomposition of particulate organic matter which has transported by sedimentation from the euphotic zone in situ (Suess, 1980) and/or by horizontal

196 advection (Asper, 1987).

Acknowledgment We would like t o t h a n k Prof. Emeritus A. H a t t o r i a n d Prof. I. Koike of the Ocean Research Institute of Tokyo University for their encouragements a n d useful discussion. T h i s work was funded by t h e Ministry of Education, Circulation subject No. 62610501 a n d 63610501. T h i s is also contribution No. 266 from t h e Research Institute of North Pacific Fisheries, Faculty of Fisheries, Hokkaido University.

References Anma, G., K. Masuda, G. Kobayashi, H. Yamaguchi, T . Meguro, S. Sasaki and K. Ohtani, 1990. Oceanographic structures and changes around transition domain along 180 longitude, during June 1979-1988. Bull. Fac. Fish. Hokka. Univer., 41, 73-88. Antia, N.J., B. R. Berland, D. J. Bonin and S. Y . Maesrini, 1975. Comparative evaluation of certain organic and inorganic sources of nitrogen for phototropic growth of marine microalgae. J. Mar. Biol. Assoc. U. K., 55, 519-539. Asper, V. L., 1987. Measuring the flux and sinking speed of marine snow aggregates. Deep-sea Res., 32, 1-17. D’Elia, C. F., P. A. Stedudler and N. Corwin, 1977. Determination of total nitrogen in aqueous samples using persulfate digestion. Limnol. Oceanogr., 22, 760-764. Dodimead, A. J., F. Favorite and T . Hirano, 1963. Salmon of the north Pacific Ocean - 2. Review of oceanography of the subarctic Pacific region. Bull. Int. North Pacific Fish. Comm., 13, 1-195. Gagosian, R. B., and C. Lee, 1981. Processes controlling the distribution of biogenic organic compounds in seawater. In: Marine Organic Chemistry, E. K. Duursma and R. Dawson (ed.), Elsevier Scientific Publishing Company, .41nsterdam-Oxford-New York, pp. 91-123. Koroleff, F., 1983. Total organic nitrogen. In: Methods of Sea Water Analysis, 2nd ed., K. Grasshoff, M. Ehrhard, and K. I-hemling (ed.), Verlag Cemie, Weinheim, pp. 162173. Lancelot, C., 1984. Extracellular release of small and large molecules by phytoplankton in the Southern Bight of the North Sea. Estuar. Coast. and Shelf Sci., 18, 65-77. McCarthy, J. J., 1972. The uptake of urea by natural populations of marine phytoplankton. Limnol. Oceanogr., 17, 738-748. Maita, Y., and M. Yanada, 1990. Vertical distribution of total dissolved nitrogen and dissolved organic nitrogen in seawater. Geochem. J., 24, 245-254. Miyake, Y . , T. Sagi and K. Saruhashi, 1982. The biogeochemical cycle of nitrogen and phosphorous in the ocean. Geochem. Res. Assoc. Sci. Report, 1-22. Miyake, Y . , K. Saruhashi and T . Sagi, 1985. On the dissolved carbon in sea water. Bull. Soc. Seawater Sci. Jpn., 38, 353-367. Odate, T., and Y. Maita, 1989. Regional variation in the size composition of phytoplankton communities in the Western North Pacific Ocean, Spring 1985. Biol. Oceanogr., 6, 65-77. Parsons, T . R., Y . Maita and C. M. Lalli, 1984. A Manual of Chemical and Biological Methods for Seawater Analysis. Pergamon Press, London, 173 pp. Sahlsten, E., 1987. Nitrogenous nutrition in the euphotic zone of the Central North Pacific Gyre. Mar. Biol., 96, 433-439. Schell, D. M., 1974. Uptake and regeneration of free amino acids in marine waters of southeast Alaska. Limnol. Oceanogr., 19, 260-270.

Schlitzer, R., 1989. Modeling the nutrient and carbon cycles of the North Atlantic. 2. New production, particle fluxes, COz gas exchange, and the role of organic nutrients. J. Geophys. Res., 94, 12,781--12,794. Sharp, J. H., 1983. The distributions of inorganic nitrogen and dissolved and particulate organic nitrogen in the sea. In: Nitrogen in the Marine Environment, E. J. Carpenter and D. G. Capone (ed.), Academic press, pp. 1--33. Small, F., S. W. Fowler, A. Moore and J. LaRosa, 1983. Dissolved and fecal pellet carbon and nitrogen release by zooplankt,on in tropical waters. Deep-sea Res., 30, 1199-1220. Sol6rzano, L., and J. H. Sharp, 1980. Determination of total dissolved nitrogen in natural waters. Limnol. Oceanogr., 25, 751-754. Strickland, J . D. H., and T. R. Parsons, 1972. A Practical Handbook of Seawater Analysis. Fish. Res. Bd. Canada, 167, 310 pp. Suess, E., 1980. Particulate organic carbon flux in the ocean surface productivity and oxygen utilization. Nature, 288, 260-263. Sugimura, Y., and Y. Suzuki, 1988. A high temperature catalytic oxidation method on non-volatile dissolved organic carbon in seawater by direct injection of liquid samples. Mar. Chem., 24, 105-131. Suzuki, Y., Y. Sugimura and T. Itoh, 1985. A catalytic oxidation method for the determination of total nitrogen dissolved in seawater. Mar. Chem., 16, 83-97. Tada, K., and T. Maita, 1988. Fluorometric determination using HPLC of dissolved amino acids in seawater. Bull. Fac. Fish. Hokka. Univer., 39, 151--159. Walsh, T. W., 1989. Total dissolved nitrogen in seawater: a new high-temperature combustion method and a comparison with photo-oxidation. Mar. Chem., 26, 295-311. Williams, P. J. LeB., 1975. Biological and chemical aspects of dissolved organic matter in seawater. In: Chemical Oceanography 2, 2nd ed., J. P. Riley and G. Skirrow (ed.), Academic Press, London, pp. 301-363. Williams, P. M., and E. R. M. Druffel, 1989. Dissolved organic matter in the ocean: comments on a controversy. Oceanography, 1, 14-17.

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Deep Ocean Circulation, Physicul and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

199

Determination of Some Oxyacid Elements and Manganese in Seawater and their Distributions in Some Unique Environments of the North Pacific Eiichiro NAKAYAMAT Yoshiki SOHRINt and Kenji ISSHIKIt

Abstract A column extraction method using macroporous resin impregnated with 7Dodecenyl-8-quinolinol was utilized for preconcentration of molybdenum and tungsten in seawater, followed by their simultaneous determination with a catalytic polarography in a special electrolytic solution. The concentration of tungsten in oceanic water was found to be 60 pM. Another column extraction method using a combination of 8-quinolinol and a macroporous resin was also applied to the speciation of chromium in seawater. Furthermore, an automated flow-system suitable for the onboard analysis of manganese in seawater was constructed. This system was proven to be useful for proving the hydrothermal activity in the sea floor.

1.

Introduction

In recent years, we have developed a variety of analytical techniques necessary for the accurate measurements of some trace elements in seawater using column extraction methods and electrochemical methods (Isshiki et al., 1989, Sohrin et al., 1987, 1989, Nakayama et al., 1989). Our research has been chiefly aimed at determining the concentrations of oxyacid elements such as chromium, iodine, molybdenum and tungsten of which behaviors in the ocean are considerably different from those of common transition metals, because of their presence in seawater in more than two valency states. The originally developed column extraction methods have been proven t o be very useful for the preconcentration of molybdenum, tungsten and chromium species. Two automated analytical methods were also successfully developed which enable us to determine trace elements such as iodine species and manganese on board ship in the contamination free closed systems (Nakayama et al., 1985 and 1989a, b). Using these methods a number of seawaters from various oceanic environments were analyzed. The noteworthy results obtained for manganese in the study of the hydrothermal activity in the east Manus Basin are presented here. Some results obtained for the oxyacid elements in the different oceanic regions are also discussed in this paper. *Research Center for Instrumental Analysis, Faculty of Science, Kyoto University, Kyoto 606, Japan +Institutefor Chemical Research, Kyoto University, Uji, Kyoto 611, J a p a n :Department of Applied Science, Kohochi Women’s University, Kohochi 780, J a p a n

200

Sampling and Analytical Methods

2.

Seawater samples for chromium, molybdenum and tungsten were collected by using Niskin PVC bottles with an inner spring of Teflon coated stainless steel. Contamination from the sampler is believed to be insignificant for the oxyacid elements (Tsubota et al., 1985). Seawater samples for manganese were also collected with Teflon coated Go-Flo samplers mounted on a CTD rosette. Various precautions against contamination were made in the sample treatment. For example, filtration of seawater samples and preconcentration by the column extraction were performed in a simple closed box covered with PVC seats. Evaporation of eluents after the column extraction were also carried out in a glass cabinet using a hot plate under continuous flow of filtered clean air. In addition, hydrochloric acid, nitric acid and aqueous ammonia of analytical-reagent grade were further purified by rapid isopiestic distillation and used for all the experiments.

2.1.

Column Extraction Using Macroporous Resins

In this method, trace metals are concentrated by combinations of organic ligands and macroporous resins (Amberlite XAD-4 resin and CHP-BOP gel) which are styrene-divinylbenzene copolymer. 7-Dodecenyl-8-quinolinol (DDQ) is an organic ligand which is capable of forming complexes with a variety of metals, chemically stable and strongly adsorbed on t o the resin. Since the DDQ loaded on XAD-4 resin is hardly.eluted at all with both the strong acidic and alkaline solution due t o the large hydrophobic vinyl substituent group, DDQ-impregnated resin (DDQ resin) can be used repeatedly like a chelating resin (Isshiki et al., 1987). DDQ resin is easily prepared by adding water (500 mL) gradually to a mixture of dry XAD-4 resin (ca. 5 g), acetone solution containing the required amount of DDQ (800 mg) and 20 mL of hydrochloric acid. The resulting DDQ resin is slurry-packed in a Teflon column (8 mm i.d., 40 mm length) fitted with porous Teflon filter (pore size, 1015 pm) and Teflon fittings. The DDQ column extraction method has advantages in that the rate of chelation with metal ions is relatively rapid as compared with the conventional Chelex-100 column extraction. The DDQ column extraction can be applied t o the preconcentration of almost all common trace metals in seawater except for cobalt and chromium of which trivalent state ions are inert in the ligand exchange. For cobalt and chromium, macroporous resins can be used as a n adsorbent for their chelate compounds when their metal-organic chelate were formed in the sample solution. We will briefly describe below the methods for simultaneous determination of molybdenum and tungsten and the fractional determination of Cr(II1) and Cr(V1) by using the column extraction method. Molybdenum and tungsten

A 500 mL filtered seawater sample adjusted to pH 5 is pumped through the DDQ column at a flow rate of 5 mL/min. Then the two elements are eluted with 2 M aqueous ammonia. The eluate is evaporated to dryness, and organic matter which might be present in the eluate is decomposed with nitric acid and hydrogen peroxide. The residue is dissolved with a 10 mL of supporting electrolyte solution containing 0.5 M potassium chlorate, 0.03 M sulfuric acid, 0.3 M benzilic acid and

20 1 * E (30

I40

IM

160

170

1110

110

I60

I50 ' W

Fig. 1 Sampling stations for molybdenum, tungsten, chromium and manganese in the North Pacific.

0

10

20

30

40

50

km

I

Fig. 2 Sampling stations in the eastern Manus Basin of Bismark Sea. 0.2 M 2-methyl-8-quinolinol. After deaeration of the resulting solution, the polarogram is recorded in the sampling-d.c. mode using a PAR 174 polarographic analyzer with a PAR 303 static mercury drop electrode. The potential is scanned from 0 t o -1.0 V at a rate of 10 mV/sec. Two catalytic current curves corresponding to the

202 two metals are obtained at -0.25 and -0.85 V, respectively. The detection limits (S/N = 3) of this method for molybdenum and tungsten are 20 and 3 pmol/L, respectively. The more details of this method will be found in Sohrin et al. (1987, 1989).

Chromium In the case of Cr(III), a 300 mL of filtered seawater sample is adjusted to pH 9.5 with an ammonia-ammonium chloride buffer solution. Then, a methanol solution of 8-quinolinol is added to the sample solution. T h e solution is heated in a microwave oven at 85°C and is passed through a CHP-20P column a t a flow rate of 5mL/min. After rinsing the column with dilute aqueous ammonia, the adsorbed complex is eluted with a methanol-hydrochloric acid mixture (100 1). Subsequently, the eluate is evaporated to dryness and the organic residue is decomposed with a mixture of nitric acid and hydrogen peroxide by heating on a hot plate. The residue was dissolved in 1 mL of 0.1 M nitric acid. For Cr(II1) plus Cr(V1) (i.e., total Cr), a 100 mL of seawater sample is acidified with 1 mL of 2 M hydrochloric acid and then to reduce Cr(V1) t o Cr(II1) 1 ml of 1 M hydroxylamine solution is added. After 1 hour standing, the solution is adjusted to pH 9.5 with aqueous ammonia. Subsequent treatments are the same as those for Cr(II1). The concentration of chromium is determined by graphite furnace atomic absorption spectrometry. The more details of this method is described in Isshiki et al. (1989).-

+

w

( X l O p m o l I L)

0

0.2

0.2 0.r 0.c 0.P

1

E

24

-

2

0

3

i :

n

4

€

B 6-

Fig. 3 Vertical profile of molybdenum and tungsten in the North Pacific Ocean (DE-2).

203 Chromium

(nmol/l)

L Fig. 4 Vertical profile of chromium species in the center of warm core ring off Sanriku, northern Japan. 2.2.

Shipboard Determination of Manganese

The new automated system for the determination of manganese on board ship developed by Nakayama et al. is based on the electrolytic concentration and chemiluminescence (CL) detection (Nakayama et al., 1989b). In this method, Mn(I1) in the seawater sample adjusted to pH 5.0 is electrochemically oxidized t o Mn(1V) oxide and quantitatively adsorbed on a flow-through glassy carbon electrode. The manganese is then eluted from the electrode with an acidic hydrogen peroxide solution (ca. pH 4.0) which reduces Mn(1V) oxide to Mn(I1) instantaneously. After passing the eluent through the aboxre stated DDQ column to remove the other interfering metal ions such as iron and copper, the eluent is mixed with an alkaline (0.24 M potassium carbonate) luminol solution and the mixture is introduced into the CL detection cell. The concentration of manganese is measured by the chemiluminescence intensities because Mn(I1) catalyzes the oxidation of luminol with hydrogen peroxide in proportion to its concentration. With this method, 0.3 to 20 nmol/L of manganese can be accurately determined from seawater samples of less than 10 mL within 10 min. per sample.

3.

Results and Discussion

Samples were collected from various environments of the Pacific Ocean, including the North Pacific between Japan and Hawaii (46'44'N, 162'22%; 44"40'N, 117"OO'W; 30°00'N, 159"50'W) for molybdenum and tungsten as shown in Fig. 1, DE-2, DE-4 and DE-7 respectively, warm core ring off Sanriku, Japan (around

204

L

Fig. 5 Detailed vertical profile of chromium species down to 500 m in the center of warm core ring off Sanriku, northern Japan. T h e d a t a indicated by crosses are from Figure 4. 40"05'N, 145"00'E, Fig. 1: WCR) for chromium and Okinawa Trough (around 27"35'N, 127"09'E, Fig. 1: CB-6), Sagami Bay (35"00'N, 139"14'E, Fig. 1: H-site) near Japan and East Manus Basin, in the Bismark Sea (around 3"41'S, 151"52'E, Fig. 2) for manganese. 3.1.

Tungsten in the North Pacific

As shown in Figure 3, both the dissolved molybdenum and tungsten show the conservative-type profile (Sohrin et al., 1987). The average concentration of tungsten was 53-GO pM (normalized to a salinity of 350/,) and that of molybdenum was 94-106 nM in all of the North Pacific waters investigated. The molar ratio of dissolved concentration in oceanic water is Mo : W = 1800 : 1 whereas the crustal abundance of molybdenum and tungsten is 1.5 and 1 ppm (Bowen, 1979), respectively and hence the molar ratio, Mo : W = 3 : 1. Despite of the similarity of their chemical properties, molybdenum is clearly enriched in oceanic water compared with tungsten. This may be due to the difference of the two elements during weathering in the solubility and/or in the removal process from the ocean. Since the molar ratio of dissolved concentration in river water is Mo : W = 33 : 1 (Bowen, 1979), the removal process from seawater must account for the remaining 55-fold enrichment of molybdenum in seawater. The concentration ratio of Mn : Mo is almost constant in various sediments and manganese nodules (Shimmield and Price, 1986), and the similar relationship is also found for tungsten (Sohrin et al., unpublished). The concentration factor defined as K D = [Metal concentration in mol kg-'Imanganese nodule / [Metal concentration in mol L-llseawater in natural

205 manganese nodule is 4 . 4 ~ 1 0for ~ molybdenum and 5 . 5 ~ 1 0for ~ tungsten. These values are approximately equal to the partition coefficients ([Metal concentration in mol kg-'],,lid phase/[hletal concentration in mol L-l]aqueous phase) obtained ~ for rnolybdenuni and by laboratory experiments using MnO;! which are 5 . 0 lo4 3 . 5 l~o 6 for tungsten (Sohrin et al., unpublished). Therefore, tungsten appears t o be removed about two order of magnitude faster than molybdenum by incorporation in marine manganese nodule and sediments. This is consistent with what we observed on the dissolved concentrations of molybdenum and tungsten in seawater.

3.2.

Chromium Species in the Warm Core Ring off Sanriku

The vertical profiles of chromium species were obtained in the center of the warm core ring off Sanriku northern Japan. The concentration of total-Cr increased with depth in the deep waters (Fig. 4). This trend is similar t o the distribution of chromium species generally observed in open oceans (Cranston and Murray, 1978; Nakayama et al., 1981). The increase of total-Cr in the deep waters suggests t h a t chromium is regenerated from sediments a t the sea-floor where Cr(II1) is oxidized to Cr(V1) presumably by the catalytic effect of manganese(1V) oxide (Nakayama et al., 1981). Figure 5 shows the detailed vertical profiles of chromium species in the shallow waters down t o 500 m at the center of warm core ring. There were two nitrite maxima at 30 and 80 m depths in this station. Two clear minima of totalM n CONC./pg/L

oo

01

02

P

I

I

03 1

0 4

I

0 5 I

1000a 0

1500-

L

Fig. 6 Vertical profile of manganese in the Okinawa Trough. Solid line, determined by shipboard analysis (symbols are indicated slightly different stations in the same marine area). Dashed line, determined by 8-quinolinol extraction-ICP AES method.

206 Cr were also noted corresponding to the nitrite maxima. Since the nitrite maxima usually couples with chlorophyll a maxima, the good correlation between nitrite maxima and minima of total-Cr (or Cr(V1)) suggests that the reduction of Cr(V1) into Cr(II1) occurs during the photosynthetic process. This reduction reaction must be rapid enough to follow the temporally variability of nit,rite maxima. 3.3.

Manganese in the Okinawa Trough, Sagami Bay and Manus Basin

Figure 6 shows the vertical distribution of manganese in the Okinawa Trough in southern Japan obtained by the shipboard analysis (Nakayama et al., 1989). The profile reflects the hydrothermal activity, exhibiting a plume of highly concentrated manganese at a depth around 1300 m. The dashed line indicates the results obtained by Ishibashi et al. (1989) using inductively coupled plasma atomic emission spectroscopy after preconcentration with 8-quinolinol-solid phase extraction at the site 2 miles from our site. The profile also shows a maximum of manganese a t a depth of around 1300 m. Figure 7 shows the vertical profile of manganese in the Sagami Bay central Japan where a large seepage of methane was found due to the force of subductioninduced compaction (Sakai et al., 1987). Manganese was highly concentrated in the bottom water, especially in the case of samples collected during 1989 cruise. Although the bottom manganese can be derived from seepage like methane, the manganese concentration in the sediment pore water was found to be very low in this area (Masuzawa, private communication). Therefore, it is likely manganese in the bottom water is derived from other sources such as the shallow organic rich Mn CONC./ p g f L

0

r

0.5 I

1.0 I

// II

4.0 I

4.5 I

50C t!

\

H PI

w

a loo(

1501

Fig. 7 Vertical profile of manganese in the Sagami Bay. The data indicated by circles and triangles were obtained in February 1988 and in June 1989, respectively.

207

Fig. 8 Manganese anomaly in ppt in t,he eastern Manus Basin, Bismark Sea. sediments. Since the d a t a of 1989 were obtained just after a submarine eruption, it is also possible that the extraordinally high concentration of manganese was related to the seismic activity. Figure 8 exhibits east-west section of distribution of manganese in the eastern Manus Basin. Since the analysis was made based on unfiltered waters, the concentrations of manganese should be regarded as those of total adsorbable manganese which is the sum of dissolved manganese and active solid manganese oxide adsorbable onto the glassy carbon electrode. Extremely high enrichments of manganese exceeding 4000 ppt were found at a depth around 1750 m of station 49 and at a depth around 1650 ni of station 37. Relatively high anomalies were also seen at a depth around 1100 m of above two stations. These observation indicate t h a t the active hydrothermal sites are located in association with the manganese anomalies. This is also supported by aluminum anomaly found a t the station 49 and methane anomalies found at the stations 49 and 37.

References Bowen, H. J. M., 1979. Environmental Chemistry of the Elements. Academic Press, London. Cranston, R. E., and J. W. Murray, 1987. The determination of chromium species in natural waters. Anal. Chim. Acta, 99, 275-278.

208 Isshiki, K., F. Tsuji, T . Kuwamoto and E. Nakayama, 1987. T h e preconcentration of trace metals from seawater with 7-dodecenyl-8-quinolinol-impregnatedmacroporous resin. Anal. Chem., 59, 2491-2495. Isshiki, K., Y. Sohrin, H. Karatani and E. Nakayama, 1989. Preconcentration of chromium(II1) and chromium(V1) in seawater by complexation with quinolin-8-01 and adsorption on macroporous resin. Anal. Chim. Acta, 224, 55-64. Nakayama, E., H. Tokoro, T . Kuwamoto and T . Fujinaga, 1981. Dissolved state of chromium in seawater. Nature, 390, 768-770. Nakayama, E., T . Kimoto and S. Okazaki, 1985. T h e automatic determination of iodine species in natural waters by a new flow-through electrode system. Anal. Chem., 57, 1057-1060. Nakayama, E., T . Kimoto, K. Isshiki, Y. Sohrin and S. Okazaki, 1989a. Determination by using a and distribution of iodide- and total-iodine in the North Pacific Ocean new automated electrochemical method. Mar. Chem. 27, 105-116. Nakayama, E., K. Isshiki, Y Sohrin and H. Karatani, 1989b. Automated determination of manganese in seawater by electrolytic concentration and chemiluminescence detection. Anal. Chem., 61, 1392-1396. Sakai, H., T. Gamo, K. Endow, J. Ishibashi, T. Ishizuka, F. Yanagisawa, M. Kusakabe, T . Akagi, G. Igarashi and S. Ohta, 1987. Geocheniical study of the bathyal seep communities a t the Hatushima site, Sagami Bay Central Japan. Geochem. J., 21, 227-236. Shimmield, G. B., and N. B. Piice, 1986. T h e behavioi of molybdenum and manganese during early sediment diagenesis - offshore Baja, California. Mar. Chem., 19, 261280. Sohrin, Y., K. Isshiki, T . Kuwamoto and E. Nakayama, 1987. Tungsten in north Pacific waters. Mar. Chem., 22, 95-103. Sohrin, Y., E. Nakayama, K. Isshiki, S. Kihara and M. Matui, 1989. Simultaneous determination of tungsten and molybdenum in seawater by catalytic current polarography after preconcentration on a resin column. Anal. Chim. Acta, 218, 25-36. Tsubota, H., 1985. The distribution and behavior of heavy metals in the ocean. In “Ocean characteristics and their changes”, Chapter 4, K. Kajiwara (ed.), Koseisha, Tokyo, pp. 225-236 (in Japanese). ~

Chapter 4 Vertical Flux of Chemical Substances and their Behaviours in the Deep Oceans

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

21 1

Seasonal Variation of Lithogenic Flux in Japan Trench Continental Slope Measured by Sediment Trap Shinichiro NORIKI, Ryosuke SAITO, Chizuru SAITO and Shizuo TSUNOGAI*

Abstract

Settling particles were collected at the continental slope of Japan Trench with a time-series sediment trap. Total particulate fluxes varied seasonally and were 134-747 mg/m2day. Lithogenic fractions are about a half, by assuming that the fraction contains 8.2% of Al. Large A1 fluxes were observed at three periods, from September to November, February, and April and May. The Fe/A1, Mg/K and Opal/Al ratios of settling particles and surface sediments strongly suggest that the large A1 flux in February is caused by lateral transport of surface sediments occuring at depths less shallower than the trap site.

1. Introduction Lithogenic particles have been transported t o the ocean by two pathways. One is river input, the other is aeolian transport. Rex and Goldberg (1958) have pointed out t h a t contribution of lithogenic materials t o pelagic sediments through atmosphere is significant. Since then, the long-range aeolian transport of lithogenic particles t o the open ocean has been actively discussed (e.g., Duce e t al., 1980; Uematsu et al., 1983; Blank et al., 1985). On the other hand, it has been found t h a t A1 fluxes measured by sediment traps increase with depth in almost all water columns a t sediment t r a p stations (Brewer et al., 1980; Honjo et al., 1982; Tsunogai et al., 1982, 1990). Saito et al. (1992) have indicated t h a t the interpolated A1 flux at 2 km depth observed with sediment trap is larger than atmospheric A1 deposition on sea surface at each station in Pacific Ocean. This is due t o the fact t h a t lithogenic particles are horizontally and vertically transported from shallower area t o deeper ocean. Recently, horizontal transport of particles from the continental shelf zone t o the deep basin were reported (e.g., Honjo et al., 1982; Biscaye e t al., 1988; Narita et al., 1990b; Ramaswamy e t al., 1991). And, much attention has been focused on the lateral transport of lithogenic particles in the ocean (e.g., Walsh et al., 1988; Monaco et al., 1990a). *Department of Chemistry, Faculty of Fisheries, Hokkaido University, Hakodate 041, J A P A N

212 In this study, seasonal variations of particulate flux were measured by a timeseries sediment trap moored at the continental slope between Japan Island and Japan Trench in the western North Pacific. Some metals of settling particles were determined. In this paper, only the origin of lithogenic fraction in the settling particles collected in the sediment trap are discussed.

2.

Experimental

Mark VI time-series sediment trap (Honjo and Doherty, 1987) was deployed at 40°55'N, 144'13'E (Fig. 1). Water depth was 5300 m, and the t r a p was moored at 100 m above the sea floor. Vertical section from 43"N, 140"E to 40"N, 146"E is also illustrated in Fig. 1. The sediment t r a p was set up at the steep slope. The

Fig. 1. Location of sediment trap experiment.

213 settling particles were collected a t intervals of 20 days, from 10 September of 1988 to 28 May of 1989. The receiving cups were filled with neutral formaldehyde solution to prevent of organic matter decomposition. The settling particles collected in receiving cups were filtered through a preweighed Nuclepore filter of 0.6 mum. Its dry weight was determined by the method of Uematsu et al. (1978). The sediment samples were collected at Stations HB2 and SR74 (Fig. 1) with a box corer. The dried sample of about 10 mg for the settling particle or surface sediment was put into a decomposition vessel of Teflon and was completely dissolved with a mixture of H N 0 3 , HC104 and HF (Noriki et al., 1980). T h e silicate and metal contents were determined by a molybdenum yellow colorimetry and a flameless atomic absorption spectrometry, respectively, referring t o a certified standard material (Okamoto, 1982). A1 and Fe contents in the residue remaining after the 2% CH3COOH and then 1 M NH2OH in 2% CH3COOH treatments were determined. The content of biogenic silica, that is opal, was obtained from the difference in concentrations between total Si and aluminosilicate-Si (Taylor, 1964) in the sample. Lithogenic fraction was calculated by assuming t h a t the fraction contains 8.2% of Al. The results obtained are given in Table 1. Organic materials (Handa et al., 1991) and natural radionuclides (Narita et al., Table 1. Total particulate fluxes and Al, Fe, and Opal contents of particles. Period (20 days)

Tot a1

Total A1

S1 S2 s3 S4 S5 S6 S7 S8 S9 S10 S11 S12 S13

From 10 Sep. 1988 30 Sep. 20 Oct. 09 Nov. 29 Nov. 19 Dec. 08 Jan. 28 Jan. 17 Feb. 09 Mar. 29Mar. 18 Apr. 08 May

Residue

flux

To mg/m2day 30 Sep. 389 20 Oct. 345 09 Nov. 371 29 Nov. 176 19 Dec. 250 08 Jan. 1989 134 28 Jan. 191 17 Feb. 394 09 Mar. 291 29 Mar. 246 18 Apr. 142 08 May 443 28 May 747

Sediment, 0-1 cm HB2 41"50'N, 142"ll'E (1080 m) SR74 40"39'N, 143"41'E (2600 m)

Fe

A1

Fe

Opal

%

%

%

%

%

2.58 4.04 3.97 4.03 4.56 4.48 4.63 4.77 4.67 4.60 4.70 3.32 3.71

1.51 2.32 2.29 2.32 2.58 2.48 2.55 2.67 2.58 2.53 2.70 1.90 2.17

2.46 3.83 3.76 3.83 4.35 4.29 4.40 4.55 4.43 4.36 4.45 3.15 3.50

1.24 1.96 1.88 1.91 2.12 2.07 2.10 2.19 2.10 2.05 2.22 1.54 1.78

46.2 38.7 37.3 37.0 34.2 32.4 29.3 31.4 31.4 32.9 31.6 48.6 45.1

6.45 5.27

3.22 2.47

21.7 34.0

214 1990a) in these samples were analyzed.

3. Results and Discussion 3.1. Total and A1 fluxes Total particulate fluxes during September ( S l ) and November (S3) were nearly constant at 400 mg/m2day, and then decreased t o about 200 mg/m2 day. A large total particulate fluxes was observed a t February (S8), and then decreased in early April. In spring (Sl2 and S13), total particulate fluxes were again increased. Mean total flux was 115 g/m2year which was the same order in high productive western North Pacific (Noriki and Tsunogai, 1986). A1 concentrations of settling particles were 3-5% (Table 1). The source of A1 is exclusively lithogenic aluminosilicate. By assuming that the fraction contains 8.2% of A1 (Taylor, 1964), lithogenic fractions are about 30-60% (Table 2). A1 fluxes are illustrated in Fig. 2. Average A1 flux is 4.7 g/m2year and about 10 times those in open ocean (Saito et al., 1992). Large A1 fluxes as observed in this area have been found at a coastal slope and deep layers of a basin (Honjo et al., 1982; Biscaye et al., 1988; Monaco et al., 1990b). As shown in Fig. 2, it seems t h a t A1 flux is synchronized to the total particulate flux. Periods when large A1 fluxes were observed are September ( S l ) t o November (S3), February ( S 8 ) , April (S12) and May (S13). 3.2.

Source of A1

The Fe/A1 and Mg/K ratios in settling particles are useful to characterize the origin of lithogenic fraction (Saito et al., 1992). Fe/A1 and Mg/K ratios in settling particles were listed in Table 1. Fe/A1 and Mg/K were nearly constant at 0.55-0.59 and at 1.1-1.3, respectively. These constant values suggest t h a t the main origin of large lithogenic fluxes should be identical.

S1 52 53 SL 55 56 S 7 S8 S9 SlOSll 512513

S O 1988

Fig. 2.

N

D

J F 1989

M

A

M

Total and A1 fluxes observed with sediment trap.

215 Some investigations (e.g., Uematsu et al., 1983; Blank et al., 1985) indicated that lithogenic materials from the Asian continent have been carried by the longrange transport over the western North Pacific. At Mt. Yokotsu-Dake (1000 m high) near the tra p station (Fig. l),atmospheric deposition of A1 was 13 mg/m2day in spring, while it was 2 mg/m2day in other seasons (Suzuki, 1987). Annual A1 deposition was calculated to be 0.55 g/m2year. A1 flux observed with sediment trap was, on the other hand, 4.8 g/m2year, and was one order of magnitude larger than the atmospheric input. Biscaye et al. (1988) and Narita et al. (1990b) have found that an estimated atmospheric input was smaller t,han observed fluxes with sediment traps in continental area by using "'Pb. It is clear t h a t atmospheric flux Table 2.

Comparison of chemical composition Lithogenic*l

%

Fe/A1 Total Residue

Mg/K*2

Settling particle

s1 s2

s3 s4 s 5 S6 s 7 S8 s 9 s10

s11 s12 S13

31.5 49.3 48.4 49.1 55.6 54.6 56.5 58.2 57.0 5 6.1 57.3 40.5 45.2

0.59 0.57 0.58 0.58 0.56 0.55 0.55 0.56 0.55 0.55 0.57 0.57 0.58

0.50 0.51 0.50 0.50 0.49 0.48 0.48 0.48 0.47 0.47 0.50 0.49 0.51

1.29 1.07 1.13 1.20 1.10 1.13 1.17 1.04 1.06 1.10 1.07 1.17 1.08

Sediment HB2 SR74

0.50 0.47

1.17 0.83

Asian soil*3 35"08'N, 81"30'E 35'49"' 79"18'E 37"03'N, 77"Ol'E 37"38'N, 78"17'E 39"30'N. 76"03'E

0.41 0.58 0.48 0.38 0.47

0.35 0.71 0.62 0.79 0.71

*' *' *3

See text. Saito et al. (1992). Suzuki: Personal communication.

216

Mukawa River 1988 and 1989

J F M A M J J A S O N D

Fig. 3. Seasonal variation of suspended load from the Mukawa River. of lithogenic materials is insufficient to settling flux observed with sediment trap in coastal and continental areas. Fe/A1 and Mg/K ratios of Asian soils are listed in Table 2. Although Fe/A1 ratios of Asian soils were the same as those of settling particles, Mg/K ratio is significantly smaller than that of settling particles. Judging from quality and quantity, the large A1 fluxes observed with sediment tr a p in this area cannot be expected only by atmospheric input. The other source of lithogenic particles is river input. T h e big rivers draining into the study area are Mukawa River and Tokachi River among rivers shown in Fig. 1. In spring, suspended loads were striking large from rivers in northern Japan Island because of the thawing of snow. In Fig. 3, seasonal variations of average suspended load of 1988 and 1989 of Mukawa river were shown (Hokkaido Kaihatu Kyoku, 1988, 1989). Spring peak of suspended load coincides with a large A1 flux observed in Spring (S12 and S13). We suggest that one of origins of a large A1 flux is the river discharge in spring time. A1 flux in February (S8) was about twice of t h a t in early February (S’i). There is no factor to increase atniospheric input and/or river discharge in late January and early February. In this case, surface sediment may be a source of a large A1 flux. This is presumed from the fact that Honjo et al. (1982) have found horizontal movement of lithogenic matter in Panama Basin, and that lateral transport of sediment particles has been revealed in continental slope, for example, of the eastern United States (Biscaye et al., 1988) and Okinawa Trough(Narita et al., 1990b). Fe/A1 ratios of surface sediments were 0.50 and 0.47 (Table 2) a t sites HB-2 and

217

E . 7 200

(u

20 15

X'

=3

10 1: m

100

m

Q

0

Q

5 0 "

0

0 S1 S2 53 S4 S5 S6 57 S 8 S9 510 Sll S12 S13 8

S O 1988

N

Opal flux

Fig. 4.

D

J F 1989

-

,

M

A

M

OpallAl

Opal fluxes and Opal/Al ratios in settling particles.

SR-74 (Fig. l),respectively. These well coincide with t o Fe/A1 of settling particles, especially those in the residual fraction. Mg/K ratios of surface sediments were about 1, while Mg/K ratio of Asian soils were apparently less than 1. Results of lithogenic elements suggest t h a t the lithogenic fraction of settling particle is similar to the surface sediments of shallower area than the sediment t r a p point. In high latitudes, diatom blooms are generally observed in spring and fall, so biogenic opal fluxes were large in late April t o May (S12 and S13) and September t o early October (SlLS3). T h e Opal/Al ratios were over 10 in biological productive seasons. Opal flux was increased, like Al, in February ( S 8 ) , b u t the Opal/Al ratios of settling particles in the winter period were around 7. T h e Opal/Al ratio was not increased, while the opal flux was increased in early February (S8). These results indicate t h a t the peak of opal flux in winter (S8) is not derived from fresh biological activity. From Table 1, Opal/Al ratios in surface sediments a t HB-2 and SR-74 calculated t o be 3.4 and 6.5, respectively. Fe/Al, Mg/K and Opal/Al of settling particles and surface sediments suggest t h a t the large A1 flux in February (S8) is caused by the transport of surface sediments shallower than t r a p site. This is consistent with the results from organic coniponents (Handa et al., 1991) and natural radionuclides (Narita et al., 1990a).

acknowledgement We are most grateful t o Dr. K. Harada, National Institute for Resources and Environment, for his strong support throughout all phases of this study. We also thank Prof. N. Handa, Nagoya University, for his kind help and advice in the sampling. We wish t o thank t o the staff of our laboratory for their cooperation in the arrangements of sediment trap. We would like t o thank the captain, crew and scientific parties of Hokusei Maru of Hokkaido University for their assistance in the sediment t r a p experiment. We are also indebted t o the Japanese Ministry

218 of Education for t h e financial support.

References Biscaye, T. G., R. F. Anderson, and B. L. Deck, 1988. Fluxes of particles and constituents to the eastern United States continental slope and rise: SEEP-I. Cont. Shelf Res., 8, 855-904. Blank, M., M. Leinen, and J. M. Prospero, 1985. Major Asian aeolian inputs indicated by the mineralogy of aerosols and sediments in the western North Pacific. Nature, 314, 84-86. Brewer, P. G., Y. Nozaki, D. W. Spencer, and A. P. Fleer, 1980. Sediment trap experiments in the deep North Atlantic: isotopic and elemental fluxes. 3. Mar. Res., 38, 703-728. Duce, R. A., C. K. Unni, B. J. Ray, J. M. Prospero, and J . T. Merrill, 1980. Long-range atmospheric transport of soil dust from Asia to the tropical North Pacific: Temporal variability. Science, 209, 1522-1524. Handa, N., N. Harada, M. Itoh, and T. Nakatsuka, 1991. Origin of settling particle by 14C age; the case of Hidaka Basin. 1991 Annual Meeting of The Oceanographical Society of Japan. Hokkaido Kaihatsu Kyoku, 1988. Annual report of river quality. Hokkaido Kaihatsu Kyoku, 1989. Annual report of river quality. Honjo, S., and K. W. Doherty, 1987. Large aperture time-series sediment traps; design objective, construction and application. Deep-sea Res., 3 5 , 133-149. Honjo, S., S. J. Manganini, and L. J. Poppe, 1982. Sedimentation of lithogenic particles in the deep ocean. Mar. Geol., 50, 199-220. Monaco, A,, P. Biscaye, J. Soyer, R. Pocklington, and S. Heussner, 1990a. Particulate fluxes and ecosystem response on a continental margin: the 1985-1988 Mediterranean ECOMARGE experiment. Cont. Shelf Res., 10, 809-839. Monaco, A., T. Courp, S. Heussner, J. Carbonne, S. W. Fowler, and B. Deniaux, 1990b. Seasonality and composition of particulate fluxes during ECOMARGE-I, western Gulf of Lions. Cont. Shelf Res., 10, 959-978. Narita, H., K. Harada, and S. Tsunogai, 1990a. Natural radionuclides in marine sediments. 1990 Annual Meeting of The Geological Society of Japan. Narita, H., K. Harada, and S. Tsunogai, 1990b. Lateral transport of sediment particles in the Okinawa Trough determined by natural radionuclides. Geochem. J., 24, 207-216. Noriki, S., K. Nakanishi, T . Fukawa, M. Uematsu, Y. Uchida, and S. Tsunogai, 1980. Use of sealed teflon vessel for the decomposition followed by the determination of chemical constituents of various marine samples. Bull. Fac. Fish. Hokkaido Univ., 31, 354-361. Noriki, S., and S. Tsunogai, 1986. Particulate fluxes and major components of settling particles from sediment trap experiments in the Pacific Ocean. Deep-sea Res., 33, 903-912. Okamoto, K., 1982. The certification of Pond Sediment, In: Preparation, Analysis and Certification of POND SEDIMENT Certified Reference Material, K. Okamoto (ed.), Res. Rep. Natl. Inst. Environ. Stud., No. 38, pp. 81-104. Ramaswamy, V., R. R. Nair, S. Manganini, B. Haake, and V. Ittekkot, 1991. Lithogenic fluxes to the deep Arabian Sea measured by sediment traps. Deep-sea Res., 38, 169184. Rex, R. W., and E. D. Goldberg, 1958. Quartz contents of pelagic sediment of the Pacific Ocean. Tellus, 10, 153-159. Saito, C., S. Noriki, and S. Tsunogai, 1992. Particulate flux of Al, a component of land

219 origin, in the western North Pacific. Deep-sea Res., 39, 1315-1327. Suzuki, T., 1987. Study on the transport of soil dust from land to ocean over the atmosphere. Thesis for Doctor of Fisheries Science, Hokkaido University, 116 pp. Taylor, S. R., 1964. Abundance of chemical elements in continental crust. Geochim. Cosmochim. Acta, 28, 1273-1285. Tsunogai, S., S. Noriki, K . Harada, and K. Tate, 1990. Vertical-change index for the particulate transport of chemical and isotopic components in the ocean. Geochem. J., 24, 229-243. Tsunogai, S., M. Uematsu, S. Noriki, N. Tanaka, and M. Yamada, 1982. Sediment trap experiment in the western North Pacific: Undulation of settling particles. Geochem. J., 16, 129-147. Uematsu, M., M. Minagawa, H. Arita, and S. Tsunogai, 1978. Determination of dry weight of total suspended matter in seawater. Bull. Fac. Fish. Hokkaido Univ., 29, 164-172. Uematsu, M., R. A. Duce, J. M. Prospero, L. Chen, J . T. Merrill, and R. L. McDonald, 1983. Transport of mineral aerosol from Asia over the North Pacific Ocean. J . Geophys. Res., 88, 5343-5352. Walsh, J. J., P. E. Biscaye, and G. T . Csanady, 1988. The 1983-1984 Shelf Edge Exchange Processes (SEEP)-I experiment: hypotheses and highlights. Cont. Shelf Res., 8, 435456.

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved

221

Vertical Fluxes of Organic Materials in the Northern North West Pacific and Breid Bay, Antarctica, with Special Reference in the Effect of Phytoplankton Bloom Nobuhiko HANDA and Takeshi NAKATSUKA*

Abstract Mooring systems were deployed at the sediment trap sites in the northern North West Pacific (40°55.33’N, 144’12.75’E; water depth, 5,369ni) in September 1988 to May 1989 and in Breid Bay, Antarctica (70’11.536’s: 24’18.679’E; water depth, 300 m) in December 1985 to February 1986. Organic carbon and nitrogen fluxes were determined with the range of 4.44 to 24.6 mg C m-’ d-’ and 0.93 to 4.26 nig N m-’ d-l respectively, for the northern North West Pacific and 12.3 to 116 mg C m-2 d-’ and 1.79 to 15.4 mg N m-’ d-’ for Breid Bay, Antarctica, respectively. Sinking particles collected were analyzed for organic carbon and nitrogen, stable carbon and nitrogen isotopes and carbohyd-rates. High values of SI3C were found in the sinking particles collected in September to October 1983 and April t o May 1989 of the northern Fiorth West Pacific and in early January 1986 of Breid Bay, Antarctica, which were coincident with the time when t,he maxima of organic carbon and nitrogen fluxes were observed in both of the ocean areas. These facts indicated that the organic matter was derived from the phytoplankton in the logarithmic phase of their growth. Increase abundance of glucose and/or glucan was found in the water extractable carbohydrate of the sinking particles collected from Breid Bay, Antarctica ill early January 1989, indicating that phytoplankton cell of the late logarit,hmic to stationary phases had to be a source of the organic matter of the sinking particles.

1.

Introduction

Organic materials produced by phytoplankton photosynthesis in the euphotic layer have an extensive effect on marine organisms living in the surface through the deep ocean floor since these materials provide an energy source when they are transferred to the ocean floor (Hinga et al., 1979; Suess, 1980). Several sediment trap experiments have been made in various ocean areas to estimate vertical flux of organic materials of biological interest such as carbohydrates (Tanoue and Handa, 1980; Wefer et al., 1982), amino acids and protein (Lee and Cronin, 1982; Liebezeit, 1987) as well as lipid materials (Tanoue and Handa, 1980; Wekaliam et al., 1980; DeBaar et al., 1983). *Water Research Institute, Nagoya university, Chikiisa-ku, Nagoya 464-01 ,Japan

222 25"OOE

24"OOE 7O0O0S

7P00S Breid Bay

.

Sediment Trap Site

70'30's

c

.

Queen Maud Land

1

Asuka Base 24'08.17'E 70'31.34'5

24"OOE

70'30S

25OOO'E

180'

Fig. 1. Sediment trap sit,e in Breid Bay, Antarctica. Extremely heavy phytoplankton bloom reportedly occurs in the high latitude areas of the northern North West Pacific and Southern Ocean in accordance with seasonal variation of chlorophyll a (Fukuchi et al., 1985) and particle flux (Wefer et al., 1988). Despite the great importance of such phytoplankton bloom for the vertical flux of organic matter, few data set of the vertical fluxes of phytoplankton cellular organic constituents have been reported so far. We intend t o report here the change in the vertical flux of particulate organic matter collected by a time-series sediment trap system in the northern North West Pacific and Breid Bay of Antarctica with special emphasis on changes in carbon and nitrogen stable isotope composition and in monosaccharide composition during the phytoplankton bloom occurring in spring and autumn in the northern North West Pacific and in austral summer of the Antarctica.

2.

Materials and Methods

Time-series sediment trap systems were deployed in 110 m depth from 28 December 1985 to 13 February 1986 a t a station located at 70'11.536's and 24O18.679'E (water depth 300 m) in Breid Bay and in the water layer of 150 m above the bottom (water depth 5369 m) from 10 September 1988 to 28 May 1989 at a station located at 40'55.33" and 144'12.75'E in the northern North West Pacific. The sinking particles collected by the sediment traps were analyzed for organic carbon and nitrogen by means of dry combustion using a CHN analyzer after removal of carbonate materials by treatment with 0.6 M HC1 overnight. For the determination of the monosaccharide composition, the particles were treated with dilute sulfuric acid (0.5 M) at 100°C for 12 h to hydrolyze combined

223

45"N

40"

35"

30"

25"

Fig. 2. Sediment trap site in the northern North West Pacific. sugars to free monosaccharides, which were analyzed by gas chroinatography after conversion of the monosaccharides to alditol per acetates (Handa et al., 1992). Fractionation of the carbohydrates in the sinking particles was conducted by the continuous extraction of the carbohydrate with hot water followed by 5% NaOH. Resulting hot water soluble, 5% NaOH sobuble and residual fractions of the carbohydrates of the sinking particles were hydrolyzed to determine monosaccharide composition. Organic materials of the sinking particles were transferred t o the quartL tube. After complete removal of air, the tube was sealed and then combusted a t 850°C for 2 h to convert its contents t o carbon dioxide and gaseous nitrogen in the presence of cupric oxide and silver foil. These gases were introduced into a mass spectrometer (Model-GR) and ratios were determined as * 3 C / 1 L Cand 15N/14N(Wada and Hattori, 1976; Minagawa and Wada, 1984), which were represented as SI3C and 6I5N respectively. Peedee belemnite (PDB) and atmospheric nitogen were used as standard for the determination of 613C and S15N of the samples.

224 3.

Results

Vertical fluxes of organic carbon and nitrogen varied in the range of 12.3 t o 116 and 1.79 t o 15.4 mg m-’ d-’ respectively, with their maxima centered a t Sample No. 5 in Breid Bay, Antarctica (Fig. 1). Values ranging from 4.44 to 24.6 and from 0.93 t o 4.26 mg m-’ d-’ for the organic carbon and nitrogen fluxes respectively were determined in the northern North West Pacific with the maxima of the fluxes a t Sample Nos. 1, 8 and 13 which were collected in September of 1988, January to February and May of 1989 (Fig. 2). The 5I3C and 6I5N values of organic materials of the sinking particles from Breid Bay, Antarctica were in the range of -26.7 to -22.8./, and 1.8 t o 3.37/, respectively (Fig. 3). Relatively high values of 6I3C were obtained for Sample Nos. 1 through 6, whereas the values tended to decrease with time from Sample Nos. 8 to 13. Low values of 615N were obtained for Sample Nos.1 through 4 without any significant change over time. The values tended t o increase with time up to 3.30/, at Sample No. 12. In the northern North West Pacific, the 6I3C of organic materials of the sinking with the maxima at Sample Nos. 1 particles were in the ranges of -22.9 to -20.3 and 1 2 but not at Sample No. 8, indicating that the variability of 613C of the organic materials of the sinking particle are much influenced by the primary productivity in the surface water being a determinative factor controlling the vertical flux of organic carbon (Fig.4). S1’N of the sinking particle did not change significantly during the whole course of the sediment trap experiment. Monosaccharide composition of the sinking particle collected from Breid Bay, Antarctica were determined (Table 1). Glucose, galactose and mannose were dominant in the sinking particles, while fucose and ribose were minor monosaccharide constituents. Arabinose, xylose and rhaninose were a little less abundant in Sample Nos. 1 through 5, however arabinose as well as galactose became dominant in Sample Nos. 8 through 12, irrespective of the tendency towards marked decrease in the vertical flux of carbohydrate over time as has been previously mentioned.

yrn

Table. 1. Monosaccharide composition (%) of the sinking particles from Breid Bay.

Rhamnose Fucose Ribose Arabinose Xylose Mannose Galactose Glucose

1 10.5 6.0 5.1 9.2 7.2 14.2 16.6 31.1

2 9.4 6.4 6.0 9.9 9.1 13.5 19.1 26.2

3 9.9 7.0 5.8 13.8 11.1 14.3 18.9 19.5

4 8.7 5.8 5.1 11.7 9.6 12.9 17.2 29.7

Sample No. 5 6 7 8 10.7 9.9 9.9 7.4 7.2 7.6 6.9 6.8 5.1 5.1 4.9 4.2 9.9 13.6 13.0 20.7 7.5 10.9 12.9 12.5 10.9 16.7 14.3 13.5 17.1 20.6 17.5 22.7 31.7 30.0 21.0 11.8

9 1 0 1 1 1 2 6.3 5.9 4.8 5.8 6.2 5.0 4.9 5.1 4.1 3.8 3.3 3.8 20.1 27.1 22.0 23.7 14.3 16.5 15.1 13.5 14.8 9.5 11.8 10.0 24.5 21.3 21.4 24.5 9.6 10.5 9.5 11.7

225 6 x

2

5

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- 4

E

z C

$ 3 c ._

- 3

z

.-0

- 2

. % 2 P

0

1

- 1

0

- 0

30

25

130 - 25

20

- 20

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- 15

10

-

i Q

0, E

e 0

.-0

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P

10

0

5

- 5

1

0

S -1988-p-

0

N

D ( J

F

M

A

- 0

M

1989-

Fig. 3. Temporal variabilities of the organic carbon and nitrogen fluxes in the northern North West Pacific. The carbohydrates of the sinking particles were fractionated into three fractions t o assess the physiological state of the phytoplankton. Water extractable fraction, 5% NaOH extractable fraction and residual fraction accounted for 16.0 t o 37.5, 55.1 to 73.0 and 6.5 t o 13.7% of total carbohydrates in the sinking particles, respectively. However, the vertical flux of the water extractable carbohydrate fraction varied t o a great extent during the course of the sediment trap experiment relative t o those of the 5% NaOH extractable and residual carbohydrate fractions. Glucose was abundant only in the water extractable fraction of Sample Nos. 1 through 6, strongly suggesting that accumulation of reserved glucan with some

226

14

I

SampleNo.

1

2

3

4

5

6

7

8

9

10

11

12

Fig. 4. Temporal variabilities of the organic carbon and nitrogen fluxes in the Breid Bay, Antarctica. oligosaccharides consisting of glucose within the phytoplankton cells had occurred in these samples (Fig. 5). Rhamnose, ribose and mannose were moderately abundant in this fraction, in which fucose, arabinose and xylose were much less abundant. Carbohydrates in the 5% NaOH extractable fraction were a major component of the sinking particles; galactose, mannose and arabinose were major monosaccharide constituents throughout the samples of the sinking particles. Carbohydrates of the residual fraction represented only a minor carbohydrate flux ranging from 0.8 t o 2.9 mg C m-' d-l. They consisted mainly of glucose, galactose, mannose and xylose with some rhamnose and fucose, however practically no ribose and arabinose were detected.

4. 4.1.

Discussion Organic and nitrogen fluxes

Average values of organic carbon and nitrogen in Breid Bay which were calculated as 60.3 mg C m-2 d-l and 8.1 mg N m-' d - l , respectively, are comparable to those obtained at various sites in the Antarctic Ocean during the austral summer with the range of 27 to 103 mg C mP2 d-' and 17.1 mg N ni-' d-' for organic carbon and nitrogen fluxes respectively (Fukuchi and Sasaki, 1981; Wefer et al., 1982; Dunbar, 1984; Schnack, 1985; Bondungen, 1986; Gersonde and Wefer, 1987). However, these values ranged much higher than those obtained in the tropical and temperate oceanic areas (Honjo, 1980; Honjo et al., 1982). Such high vertical fluxes of organic carbon and nitrogen are considered to be mainly due to the outbreak of

227

Sample NO.

Fig. ., 5 . Vertical fluxes of monosaccharide constituents in the water extractable 5% NrtOH extractable ........... and residual car-

n m l

bohydrate fractions

I *

phytoplankton bloom often occurring in austral summer (Wefer et al., 1988). Organic carbon and nitrogen fluxes in the northern North West Pacific were 10.7 mg C m-' d-' and 2.23 mg N m-' d-' in average, of which values are much higher than those obtained in the northern North Pacific (47"55.1'N, 176"20.6'E at 5.25 km depth; Tanoue and Hanad, 1980), in the central North Pacific (15"21.1'N, 151'28.5'W a t 5.58 km depth; Honjo, 1980). This is because that our sediment trap site is deployed much closer to the coast, which results in enough supply of nutrients from the coastal areas. Seasonal variability of the organic carbon and nitrogen fluxes also was distinctive in this oceanic area. Extremely higher values of the organic carbon and nitrogen fluxes were obtained in September t o October

228

Fig. 5 . (continued) and in April t o May, which were coincident with time when phytoplankton bloom occurred in the sea areas around the northern part of Japan every autumn and spring (Imai, 1988). Higher values of the ratio of opaline Si t o A1 were obtained in the sinking particles collected in September to October and in April t o May relative t o the samples collected in other seasons of the year (Noriki, personal communication), indicating that an increase in the organic carbon and nitrogen fluxes in September to October and April t o May was obviously caused by the autumn and spring bloom of phytoplankton in the surface water. 4.2.

S13C and 6"N

The 13C and I5N contents of the sinking particles showed low values as compared t o those obtained in oceanic areas from tropical t o subarctic regions (Sackett et al., 1965; Saino and Hattori, 1980) because of low temperature and low light intensity (El-Sayed, 1990). Surface water temperature was often observed to range from -1.1 to -1.7"C during sediment trap deployment in Breid Bay, Antarctica (Fukuchi et al., 1988), and high pCOz of the surface water relative t o the atmospheric pCO2

229 11

Glucose

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(continued)

in the Antarctic Ocean (Inoue and Sugimura, 1985) gave rise t o low I 3 C content in the organic matter of the phytoplankton, in which S13C values ranged from -26 to -280/, (Sackett et al., 1965; Degens et al., 1968; Sweeney et al., 1978). A high rate of phytoplankton production of organic matter, however, results in shortage of cellular carbon dioxide due to the insufficient supply of carbon dioxide by diffusion across the phytoplankton cell membrane. Thus, phytoplankton must assimilate the carbon dioxide without enough discrimination between lighter carbon ("C) and heavier carbon ("C). This reaction process results in the formation of organic matter with high I3C content (O'Leary, 1981; Cifuentes et al., 1988). In light of these facts, it can be seen that high content of I3C in the organic matter of the sinking particles collected from the end of December 1985 t o the middle of January 1986 in Breid Bay, Antarctica was caused by the high phytoplankton growth rate in the euphotic layer which produced a phytoplankton bloom, resulting in an increase of the vertical flux of organic matter. However, the decrease in 13C content of the sinking particles collected after the middle of January clearly indicated that the phytoplankton growth rate had tended to decrease toward the end of the sediment trap experiment, and this resulted in a decrease in the vertical

230 flux of organic matter t o the underlying waters. T h e 6I5N of the sinking particles of Sample Nos. 1 through 4 measured in the range from 1.8 t o 1.9./, which were much lower values than for SI5N from nitrate dissolved in seawater (70/,, Wada et al., 1987), although the 615N of the sinking particles tended t o increase with time toward the end of the sediment trap experiment. Relatively low values of 615N, less than 2Ym for the organic matter of phytoplankton, have often been measured when nitrate is relatively abundant in the ambient waters but does not limit the active growth of phytoplankton, as observed in the boreal North Pacific (Wada, 1980). Considering, the fact t h a t nitrate is abundant in the euphotic layer due t o intensive vertical mixing of the surface by underlying waters during summer in the antarctic continental shelf area (Fukuchi and Sasaki, 1981) and in the Ross Sea (El-Sayed et al., 1983), the low values of 6I5N in sinking particles such as Sample Nos. 1 through 4, must be due t o large fractionation of nitrogen isotopes during the photosynthetic reaction of phytoplankton. T h e 15N content of the sinking particles of Sample Nos. 5 t o 12 tended to increase with time. This could be due t o enrichment of I5N in nitrate dissolved in the surface water caused by the preferential assimilation of I4N nitrate. High values of 613C was also found in the sinking particles collected from the northern North West Pacific in September t o October and April t o May, when the maxima of the organic carbon flux and the opaline Si/A1 ratio coincidently were observed in the sinking particles. From these facts, it can be seen t h a t phytoplankton bloom also occurred in the surface water of our sediment t r a p site in the northern North West Pacific in September t o October of 1988 and in April t o May of 1989. However, S15N in the sinking particles of this sediment t r a p site ranged from 4.73 t o 5.160/,, which were rather higher comparing t o those of Breid Bay, Antarctica. Considering t h a t nitrate concentration in the euphotic layers of this oceanic areas is reportedly low all year round, it is most likely t h a t a little shortage of nitrate ions comparing t o the photosynthetic requiremerit of phytoplankton caused elevated values of 615N of the sinking particles.

4.3.

Carbohydrates

T h e accumulation of water extractable carbohydrate are commonly found in the stationary phase of diatom cultured in the laboratory (Handa, 1969; Haug and Myklestad, 1976). Glucan with hemicellulose sometimes account for more than 40% of the total cellular carbohydrate of the diatom, resulting in a decrease of cell wall polysaccharides less than 55% of the total cellular carbohydrate. Much less abundance of the water extractable carbohydrate, however was found in the diatoms of the lag phase and/or logarithmic phase of their growth, which resulted in relatively higher abundance of cell wall polysaccharide soluble in dilute alkali. Taking these facts into account, it can be seen t h a t the great abundance of water extractable carbohydrate in the sinking particles in Breid Bay centered around Sample No.5 must be due t o the accumulation of reserved glucan and water-soluble hemicellulose within the diatom in the stationary phase (Fig. 4). T h e monosaccharide composition study sufficed t o support this explanation because glucose was

23 1 extremely abundant in then acid hydrolysate of the water extractable carbohydrate fraction of the sinking particles of Sample Nos. 1 through 7. The increased abundance of 5% NaOH extractable carbohydrate in the sinking particles of Sample Nos. 8 t o 12, however, appears to be due t o the diatoms between the lag and exponential phases. Glucose, galactose, mannose, xylose, arabinose, ribose, fucose and rhamnose were also detected in the water extractable carbohydrate fraction (Fig. 4). It is most likely t h a t these monosaccharides are the building blocks of water soluble heteropolvsacrharide (Mvklestarl et al. 1972. Sa-hlz+wT am4 HXK~..1_98F\\,IY&C&. has many structural similarities with heteropolysaccharide dissolved in seawater (Sakugawa and Handa, 1985). Thus, it is conceivable that water extractable carbohydrates other than reserved glucan also play an important role as source materials of the carbohydrates in dissolved organic matter. Galactose, mannose, xylose and glucose were often found in the diatom cell wall as major monosaccharide constituents together with some rhamnose and fucose, but very little ribose and arabinose have been found upon acid hydrolysis. However, the major monosaccharide composition changes depending on the diatoni species considered (Handa and Yanagi, 1969; Hecky et al., 1973; Haug and Myklestad, 1976). Still, acid hydrolysis revealed that galactose and arabinose are major monosaccharide components of the 5% NaOH extractable carbohydrate fraction together with glucose, mannose, xylose, ribose, fucose and rhamnose. Monosaccharide compositions of the 5% NaOH extractable carbohydrates from the sinking particles of Breid Bay appear to be much different from those of the diatoms cultured in the laboratory. The great abundance of galactose in this carbohydrate fraction, however, strongly suggests t h a t these monosaccharides must be a building block of the cell wall polysaccharide of the diatoms. Arabinose, which was scarcely found in the cell wall polysaccharide of diatom (Hecky et al., 1973; Haug and Myklestad 1976) is abundant in the 5% NaOH fraction of the sinking particles from Breid Bay (Fig. 4). Arabinose is generally abundant in higher plants, such as arabinoxylan in grasses and arabinoglucuronoxylan in conifers, both of which are soluble in 4 or 5% NaOH. Nevertheless, the arabinose found in the sinking particles could hardly be derived from these sources. Recently, arabinose as well as glucose and ribose, have been found in the zooplankton of hydrozoans (Arai et al., 1989). Thus, zooplankton may be as a source of arabinose for the sinking particles to some extent. Ribose and arabinose were not present in the residual carbohydrate fraction in all of the samples of the sinking particles from Breid Bay. This is because the ribose and arabinose of the sinking particles are completely extracted with water and 5% NaOH. Hecky et al. (1973) also reported t h a t almost all of the ribose was found in the cell content when diatom organic matter was separated into cell content and cell wall. These facts suggest t h a t most of the ribose found in the diatom cell is a component of low molecular weight compounds but not of biopolymers.

References Arai, M.N , J. A . Food, N. C. Whyte,

1989. Bichemical composition of fed and starved

232 Aequorea victoria (Murbach et Shearer, 1902) (Hydromedusa). J. Exp. Mar. Biol. Eco., 127, 289-299. Bondungen, B. V., 1986. Phytoplankton growth and Krill grazing during spring in the Bransfield Strait, Antarctica - Implication from sediment trap studies. Polar Biol., 6, 153-160. Cifuentes, L. A., J. H. Sharp, M. L. Fogel, 1988. Stable carbon and nitrogen isotopes biogeochemistry in the Delaware estuary. Limnol. Oceanogr., 33, 1102-1115. DeBaar, H. J., J. W. Farrtington, S. G. Wakeham, 1983. Vertical flux of fatty acids in the North Atlantic Ocean. J. Mar. Res., 41, 19-41. Degens, E. T., R. L. Guillard, J. A. Hellebust, 1968. Metabolic fractionation of carbon isotopes in marine plankton, I. Temperature and respiration experiments. Deep-sea Res., 15, 1-9. Danbar, R. B., 1984. Sediment tap experiment on the Antarctic continental margin. Ant. J. US., 19, 70-81. El-Sayed, S. Z., D. C. Biggs, Holm-Hansen, 1983. Phytoplankton standing crop, primary productivity, and near-surface nitrogen nutrient fields in the Ross Sea, Antarctica. Deep-sea Res., 30, 871-886. Fukuchi, M., H. Hattori, H. Sasaki, T . Hoshiai, 1988. A phytoplankton bloom and associated processes observed with a long term moored system in Antarctic water. Mar. Ecol. Prog. Ser., 45, 279-288. Fukuchi, M., H. Sasaki, 1981. Phytoplankton and zooplankton standing stocks and downward flux of particulate material around fast ice edge of Lutzow-Holm Bay, Antarctica. Mem. Natl. Inst. Polar Res., 34, 13-36. Fukuchi, M., A. Tanimura, H. Ohtsuka, T . Hoshiai, 1985. Report on the BIOMASSoriented research at Shyowa Station in 1982. Antarctic Record of Natl. Inst. Polar Res., 85, 102-117. Gersonde, R., G. Wefer, 1987. Sedimentation of biogenic siliceous particles in Antarctic waters (Atlantic section). Mar. Micropaleontology, 11, 311-332. Handa, N., 1969. Carbohydrate metabolism in the marine diatom Skeretonema costatum. Mar. Biol., 4, 208-214. Handa, N., and K. Yanagi, 1969. Studies on water extractable carbohydrates of the particulate matter from the northwest Pacific ocean. Mar. Biol., 4, 197-210. Haug, A., S. Myklestad, 1976. Polysaccharides of marine diatoms with special reference to Chaetoceros species. Mar. Biol., 34, 217-222. Hecky, R. E., K. Mopper, P. Kilham, E. T . Degens, 1973. The aminoacid and sugar composition of diatom cell-wall. Mar. Biol., 19, 323-331. Hinga, K. H., J . McN. Sieburth, G. R. Heath, 1979. The supply and use of organic material at the deep-sea floor. J. Mar. Res., 37, 557-579. Honjo, S., 1980. Material fluxes and sedimentation in the mesopelagic and bathpelagic zones. J. Mar. Res., 38, 53-97. Honjo, S., S. Manganini, J. Cole, 1982. Sedimentation of biogenic matter in the deep ocean. Deep-sea Res., 29, 609-625. Imai, M., 1988. Seasonal variation of chlorophyll a in the seas around Japan. Oceanol. Mag., 38, 23-32. Inoue, H., Y. Sugimura, 1985. Pcoz and 613C values in surface seawaters and atmosphere in southern oceans. Abstract in the Seventh Symposium on Polar Biology, Natl. Inst. Polar Res., Tokyo, p 2. Lee, C., C. Cronin, 1982. The vertical flux of particulate organic nitrogen in the sea: decomposition of amino acids in the Peru upwelling area and the equatorial Atlantic. J. Mar. Res., 40, 227-251.

23 3 Liebezeit, G., 1987. Early diagenesis of carbohydrates in the marine environment. I. Sediment trap experiments. In “Particulate Flux in the Ocean”, E. T. Degens, E. Izdale, S. Honjo (ed.), Selbstverlag des Geologisch-Paleontrogischen Institutes der Universitat Hamburg, pp. 279-299. Minagawa, M., E. Wada, 1984. Stepwise enrichment of 15N along food chains: Further evidence and the relation between “ N and animal age. Geochim. Cosmochim. Acta., 48, 1135-1140. Myklestad, S., A. Haug, 1972. Production of carbohydrates by the marine diatom Chaetoceros affinis var. willei (Gran) Hustedt. 11. Preliminary investigation of the extracellular polysaccharide. J. Exp. Mar. Ecol., 9, 137-144. O’Leary, M. H., 1981. Carbon isotope fractionation in plant. Phytochemistry, 20, 553567. Sackett, W. M., W. R. Eckelmann, M. L. Bender, W. H. Be, 1965. Temperature dependence of carbon isotope composition in marine plankton and sediments. Science, 148, 235-237. Saino, T., A. Hattori, 1980. ” N abundance in oceanic suspended particulate matter. Nature, Lond., 283, 752-754. Sakugawa, H., N. Handa, 1985. Isolation and characterization of dissolved and particulate polysaccharides in Mikawa Bay. Geochim. Cosmochim. Acta., 8, 185-196. Schnack, S. B., 1985. A note on the sedimentation of particulate matter in Antarctic waters during summer. Meeresforschungsergebnisse 30, 306-315. Suess, E., 1980. Particulate organic flux in the oceans-surface productivity and oxygen utilization. Nature, Lond., 288, 260-263. Sweeney, R. E., K. K. Liu, I. R. Kaplan, 1978. Ocean nitrogen isotopes and their use in determining the source of sedimentary nitrogen. New Zealand Dept. Sci. Ind. Res. Bull., 9-26. Tanoue, E., N. Handa, 1980. Vertical transport of organic materials in the northern North pacific as determined by sediment trap experiment. Part I. Fatty acid composition. J . Oceanogr. SOC.Japan, 36, 231-245. Wada, E., 1980. Nitrogen isotope fractionation and its significance in biogeochemical processes occurring in marine environment. In “Isotope Marine Chemistry”, E. D. Goldberg, Y. Horibe, K . Saruhashi (ed.), Uchida-Rokakuho, Tokyo, pp. 375-398. Wada, E., A. Hattori, 1976. Natural abundance of 15N in particulate organic matter in the North Pacific. Geochim. Cosmochim. Acta, 40, 249--251. Wakeham, S. G., J. W. Farrington, R. B. Gagosian, C. Lee, H. DeBaar, G. E. Nigelli, B. W. Tipp, S. 0. Smith, N. H. Few, 1980. Organic matter fluxes from sediment tarps in the equatorial Atlantic Ocean. Nature, Lond., 286, 798-800. Wefer, G., G. Fischer, D. Fuetterer, R. Gersonde, 1988. Seasonal particle flux in the Brabsfiel Strait, Antarctica. Deep-sea Res., 35, 891-898. Wefer, G., E. Suess, W. Balzer, G. Liebezeit, P. J. Muller, A. Ungerer, W. Zenk, 1982. Flux of biogenic components from sediment trap deployment in circumpolar waters of the Drake Passage. Nature, Lond., 299, 145-147.

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved

235

Organic Composition of Sinking Particles (JT-01) in Japan Trench as Revealed by Pyrolysis Gas Chromatography/Mass Spectrometry Ryoshi ISHIWATARI: Shuichi YAMAMOTO: Nobuhiko HANDAt and Yoshiyuki NOZAKIS

Abstract Organic matter in sinking particles from Japan Trench (JT-01) was analyzed by pyrolysis-gas chromatography/mass spectrometry with the following results: The sinking particles produced mainly aromatic hydrocarbons and nitrogen compounds with lesser amounts of phenols, furans, aliphatic hydrocarbons and aliphatic cyanides. Ratios of n-alkanes/n-alkenes and nalkanes/indole and the production of aromatic hydrocarbons and pyridines were used t o determine whether organic matter in particles is fresh or mature. It was concluded that organic matter in the sinking particles from Japan Trench is essentially composed of fresh and matured (i.e., resuspended) ones.

1.

Introduction

Organic matter in sinking particles in the ocean is composed of that produced in the surface/subsurface waters (autochthonous), that derived from lands (allochthonous), and probably resuspended one. Studying the composition, source and chemical changes of organic matter in sinking particles is important for understanding the organic carbon cycle in the ocean. Thus, many works have been done on this study (e.g., Crisp et al., 1979; Suess, 1980; Tanoue and Handa, 1980; Deuser and Ross, 1980; Wefer et al., 1982; Wakeham, 1982; Gagosian et al., 1982), and have revealed several aspects of carbon cycles in the ocean including (1) the presence of rapid transformation process of organic matter froni surface to deep water, (2) a big difference in the organic composition between sinking particles and bottom sediment, and (3) existence of land derived organic matter in both sinking particles and sediment. However, our knowledge on organic matter in sinking particles is yet insufficient t o draw a complete picture of geochemistry of organic matter in the ocean. Most conventional methods of organic compounds in geochemical samples involves extraction, separation by chromatography and determination, which are time *Tokyo Metropolitan University, Minami-Ohsawa, Hachioji, Tokyo 192-03, J a p a n t Soka University, Tangicho, Hachioji, Tokyo 192, J a p a n $Water Research Institute, Nagoya University, Frocho, Chikusa-ku, Nagoya 464, J a p a n §Ocean Research Institute, University of Tokyo, Minamidai, Nakano-ku, Tokyo 164, J a p a n

236 consuming and require a large amount of samples for one analysis. Therefore, using such a method is not necessarily suitable for organic analysis of sinking particles because the amount of samples is usually very limited. Pyrolysis-gas chromatography (Py-GC) and pyrolysis-gas chromatography/mass spectrometry (Py-GC-MS) have been often used for obtaining the general feature of organic matter in geochemical samples (e.g., Bracewell and Robertson, 1976; McHugh et al., 1976, 1978). This method is simple as compared with conventional methods of organic analysis. A milligram amount of sample is pyrolyzed a t high temperature (commonly 600-800°C) in the absence of oxygen. This leads to decomposition of organic matter into small fragmenh, which are analyzed by GC or GC-MS. The nature and quantity of these fragments reflect the nature and the quantity of the parent (initial) organic matter. Several authors have successfully applied this method for characterizing particulate organic matter in natural waters (Van de Meent et al., 1980a; Sigleo et al., 1982; Whelan et al., 1983; Yamamoto and Ishiwatari, 1987; Yamamoto et al., 1988). In this paper, we report the result of Py-GC/MS of sinking particles from Japan Trench.

2. 2.1.

Experimental Materials

Sinking particle samples (JT-01) were collected at the JT location (maximum water depth 9200 m; 34"10.4'N, 141'58.9'E) in Japan Trench. The sampling device was set a t depth of 8798 m from August 30, 1986, t o May 4, 1987, and in total 13 particle samples were taken. The detailed description was given by Handa (1989). At sampling, sinking particles were divided into two subsamples with different grain sizes by wet sieving using 1 mm sieve. The samples smaller than 1 mm were numbered as #1 to #13 and the 13 samples larger than 1 mm were combined into four groups and named as A to D, where sample A corresponds to #1-#3, Sample B to #4-#6, sample C to #7-#10 and sample D t o #11-#13, respectively. Those particle samples were freeze-dried before analysis. 2.2.

Methods

Pyrolysis-GC/MS analysis was conducted by using a pyrolyzer (Chemical Data Systems Pyroprobe Model 100) connected to a Varian Model 3400 GC (equipped with a fused silica capillary column wall-coated with DB-5 30 mx0.25 mm id.)Finnigan Model INCOS 50 quadrupole mass spectrometer. In a typical operation, 2-3 mg of a freeze-dried sample was loaded in a quartz tube ( 2 mm 4~ 15 mm), inserted into an interface preheated at 300°C and pyrolyzed at 750°C (actual temperature is probably 650°C for 20 sec (Fig. 1). The temperature of gas chromatograph was first maintained at 40°C for 5 min. and programmed from 40°C to 300°C a t a rate of 6°C min-l. Mass spectrometric data were acquired in the electron impact mode (70 eV), scanning from 50-650 mass units a t 1.5 s per decade. The pyrolysis products (pyrolysates) were quantified from the area of a characteristic mass fragment of a compound (given in Table 1) by comparison with the area m/z 71 of normal octadecane used as an external standard.

237 3. Results and Discussion Fig. 2 shows the representative gas chromatogram (RIC) of pyrolysates of J T (Japan Trench)-sinking particles. All the pyrolysis gas chroniatograms of sinking particles are given in Appendix. Many peaks are observed on the gas chromatogram. Table 1 lists the tentative identification of pyrolysates and their probable source organic matter. Aliphatic hydrocarbons, fatty acids/aliphatic cyanides, alkylbenzenes, phenols, furans and nitrogen compounds were detected in the pyrolysates. Table 2 gives quantitative results of pyrolysis products. As shown in av. Fig. 2 and Table 2, major products are nitrogen compounds (JT-#1-#13: 34%; JT-AND: av. 34%), aromatic hydrocarbons (JT-#1-#13: av. 43%; J T A-D: av. 30%), phenols (JT-#1-#13: av. 18%; JT-A-D: av. 15%) and furans (JT-#1-#13: av. 11%;JT-AND: av. 9%). Aliphatic hydrocarbons (JT-#1-#13: av. 2%; JT-AND: av. 3%) are minor and aliphatic cyanides ( J T - # 1 ~ # 1 3 : av. 0.5%; JT-A-D: av. 0.8%) were produced in small amounts. Major products and relation with their parent materials are summarized below. (1) Aliphatic Hydrocarbons: n-Alkanes and n-alkenes ranging from Clo to C27 were produced by pyrolysis. These compounds may be derived from lipid materials with long methylene chains. Isoprenoid hydrocarbons, such as pristane, phytane and pristenes (prist-1-ene and prist-2-ene), are considered to be derived from phytol side chain of chlorophyll pigments (Van de Meent et al., 1980b). (2) Fatty Acids/Aliphatic Cyanides: Fatty acids niay come directly from particles and/or from fatty acid esters by decomposition. Being different from several sinking particles reported (Yamamoto and Ishiwatari, 1987; Yamamoto et al., Coil e l e m e n t

: 750?2"C,75'C/msec,

2Osec

q u a r t z tube

Injection port

n

G C or GC-MS

Pyroprobe 100 S e r i e s Solid Pyrolyzer Chemical

D a t a S y s t e m s , Inc. Fig. 1. Pyrolysis apparatus.

Oven temp. LO'C 5min LO - 31 0°C 5" C/min Column f u s e d silica DB-5 0.25

mme x 30 m

23 8 1988), the samples did not produce fatty acids but produce fatty cyanides (n-Cl2n-Cls) on pyrolysis. Aliphatic cyanides are thought t o be derived from fatty acid amides. Therefore, the aliphatic cyanides are closely related t o fatty acids. T h e reason why the pyrolysis products of JT-sinking particles are different from those from other places is unknown at present. ( 3 ) Alkylbenzenes: Benzenes (Cl-Cq), indenes ( C O - C ~ )naphthalenes , (COC 4 ) , biphenyls (Co-C4) and antracenes/or phenanthrenes (Co--C3) were present in the pyrolysates. Toluene and xylene with short alkyl side chains are known t o be derived from phenylalanine of amino acids (Shulman and Simmonds, 1968; Tsuge and Matsubara, 1985). T h e source organic matter of the other aromatic hydrocarbons (benzenes with larger alkyl groups ( >C2), indenes, naphthalenes, biphenyls, antracenes and phenanthrenes) are not clear a t present. (4) Phenols: Phenols (Co-C,) were detected in the pyrolysates. These compounds are likely produced from both proteins (Simmonds e t al., 1969) and lignin (Bracewell and Robertson, 1976; Faix et al., 1987; Saiz-Jimenez et al., 1987). If lignin is present in a sample, methoxyphenols should be detected in a pyrolysate. However, these phenols were not found in the pyrolysates of sinking particles. This Table 1. Pyrolysis products and their possible precursors Abbre- Pyrolysis viation product aliphatic hydrocarbons n n-alkanes n= n-alkenes i-n isoprenoid alkanes i-n= isoprenoid alkenes other aliphatic compounds ACn f a t t v acids CNn aliphatic cyanides aromatic hydrocarbons BZn C,-benxenes INn C,-indenes N An C, - naphthalenes BIn C,-biphenyls APn C,-anthracenes APn C, -phenanthrenes uhenols PHn C,-phenols MPn

C,-methoxyphenols

identified range #1--#13 A-D

mass fragment for analysis

possible precursor lipid lipid phytol phytol

n=9-24 n=9-24 n=19, 20 n=19(1,2-enej

n=9-32 n=9-29 n=lg(l,Z-ene)

71 69 71 69

N. D. n=12-18

N . D. n=12-18

73 110

fatty acid fatty acid?

n=1-4 n=0-3 n=9-4 n =0-4 n=0-3 n=0-3

n=1-4 n=0-3 n=0-4 n=0-4 n=0-3 n=0-3

92+14n 116+14n 128+14n 154+14n 178+14n 178+ 14n

protein(n=1-2 j $7 ?

n=0--3

n=0-3

94+14n

N. D.

N . D.

124+14n

protein(n=0-2) lignin(n=O-q) lignin(n=O -3)

n=1-3 n=0-1 n=0-3

n=1-3 n=0-1 n=0-3

68+14n 96+14n 118+14n

carbohydrate carbohydrate

n=0-4 n=0-2 n=0-4 n=0-3

n=0-4 n=0-2 n=0-4 n=0-3

67+1411 79+14n 103+14n 117+14n

protein(n=0-2) protein+? protein+? protein(n=O-I)

n=19, 20

7 ? 0 0

furans FRn C,-furans FAn C,-furan aldehydes BFn C,-benzofurans N compounds PRn C ., -Dvrroles ." PDn C,-pyridines BCn benzene-C,-cyanides IDn C,-indoles

9

239 indicates that lignin is probably not present in the particles, and therefore, phenols may have been derived from proteins. ( 5 ) Furans and related compounds: Furans (Cz-C,), furaldehydes (Co-C,) and benzofurans (Co-C,) were detected in the pyrolysates. Furans and furaldehydes are produced from carbohydrates (Simmonds et al., 1969). Since benzofurans are produced from melanoidins (condensation product of amino acids with carbohydrates) and heated carbohydrates (Yamamoto and Ishiwatari, unpublished data). these compounds have been derived from carbohydrates and related materials. (6) Nitrogen compounds: Pyrroles (Co-C,) pyridines ( C O - C ~ )benzene alkylcyanides (Co-C,) and indoles (CO-C,) were produced from sinking particles on pyrolysis. These compounds are produced from proteins, amino acids and their derivatives. According t o Simmonds et al. (1969), Danielson and Rogers (1978), Bracewell and Robertson (1984), and Tsuge and Matsubara (1985), indoles, pyrroles and benzene alkyl cyanides are derived from tryptophane, glutamic acid (or proline) and phenylalanine, respectively. Pyridines are not produced from amino acids or protein, but are obtained from an altered protein such as melanoidins (Yamamoto and Ishiwatari, unpublished data). 3.1.

General feature of organic matter of sinking particles

Fig. 3 shows mass chromatograms at m/z 71 and ni/z 69 corresponding t o nalkanes and n-alkenes, respectively, in the pyrolysates. T h e abundance of aliphatic

T a b l e 2.

Py-GC results of sinking particles from J a p a n Trench

grain size > 1 m m A B C D av. #1 # 2 relative comDositionlOln of the total) ,. aliphatic HCs 2.8 2.8 2.6 2.6 2 .7 2.6 2.6 aliphatic cyanides 0.9 0.8 0.7 0.6 0.6 0.6 aromatic HCs 29.5 30.0 30.3 28.8 29.7 35.7 34.5 phenols 13.8 15.4 14.5 15.2 14.8 17.2 18.4 furans 10.5 7.1 9.2 8.3 8.8 10.2 10.2 N compounds 42.6 43.6 42.4 44.4 43.2 33.9 33.8 nitrogen compounds(% of the total nitrogen compounds) pyrroles 41.4 41.4 42.1 42.3 41.8 25.5 2 2 . 2 pyridines 38.5 36.7 36.7 34.9 36.7 53.6 56.1 benzene cyanides 9.5 9.7 9.5 9.2 9 .5 12.8 14.3 indoles 10.5 12.3 11.8 13.6 12.1 8.1 7.4

grain size #3

#4

#5

#6

<

#7

h,

P 0

1 mm #8

#9

#lo

#11

#12

#13

av.

2.7 0. 7 33.8 18.9 10.4 33.6

2.2 0.4 33.4 17.9 11.2 34.9

2.3 0.3 34.1 18.5 10.7 34.2

2.3 0.4 34.6 21.5 10.4 30.8

2.4 0.4 33.7 19.9 10.1 33.5

2.2 0.5 30.8 20.9 11.7 33.9

2.4

-I

n-alkanes H/L* n-alkenes H/L* A/Et A/I: E/I§

0.24 0.16 1.2 0.63 0.52

0.28 0.18 1.2 0.52 0.43

0.34 0.22 1.3 0.55 0.43

0.26 0.28 0.18 0.19 .2 1.2 1 0.44 0.54 0.36 0.44

* c2o-c27/c1o-c19 t n-alkanes(Cio-C27)/n-alkenes(C~0-Cn) 1 n-alkanes(Clo-C27)/indole 3 n- alkenes( Cio-C27)/indole

2.5 2.5 2.6 2.0 2.7 0.6 0.4 0.4 0.3 0 . 3 34.6 36.4 34.2 34.0 35.0 18.1 16.3 17.5 17.0 17.8 10.7 9.9 11.2 11.9 11.6 33.5 34.6 34.1 34.9 32.6

24.3 21.2 21.1 24.3 25.9 26.2 25.7 26.2 25.0 24.6 28.3 53.3 59.3 59.2 57.4 54.8 53.4 55.4 53.7 52.1 54.0 51.5 15.2 13.6 14.7 14.1 13.8 14.3 13.6 15.0 17.4 15.7 12.7 7.1 5.8 5.0 4.2 5.6 6.1 5.4 5.1 5.4 5.7 7.5

0.16 0.11 0.14 0.13 0.28 0.08 0.14 0.10 0.14 0.11 0.11 0.11 0.10 0.08 0.09 0.08 0.11 0.04 0.08 0.'05 0.10 0.09 0.07 0.08 1.6 1.7 1.3 1.3 1.7 1.5 1.6 1 . 5 1.6 1.4 1 . 5 2 . 2 1.2 1.1 1.1 1.3 2.4 1.6 1.8 1.4 1.1 1.6 1.7 1.5 1.1 0.94 0.78 0.70 0.80 0.88 1.1 0.99 1.0 1.1 0.91 0.94

0.5 34.2 18.4 10.8 33.7

24.7 54.9

14.4 6.0

0.13 0.13 0.09 0.0(3 1.3 1.5 .4 0.84 1 0.67 0.92

24 1 compounds with long chain length (>C,o) is very small. In addition, methoxyphenols which are a n indicator of lignins is not detected in the pyrolysates. These two facts indicate extremely small abundance of land-derived organic matter in JT-sinking particles. This is in accordance with t,he conclusion made by Handa (1989) from analysis of carbohydrates (i.e., Glucose/Arabinose ratio). Fig. 4 summarizes the relative abundance of pyrolysates. T h e analytical result of sinking particles from the Sagami Bay is also presented for comparison in Fig. 3. Their relative abundance is quite similar for all JT-#1-#13. T h e pyexcept rolysate composition of JT-A-D is also similar t o t h a t of JT-#1-#13 15

71 ("-alkane)

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242 for the following point. The abundance of aromatic hydrocarbons for the former (29-30%) is slightly smaller than that for the latter ( 3 1 ~ 3 6 %while ) the relative abundance of nitrogen compounds shows a reverse relation. It is known t h a t oxygen and nitrogen contents of sedimentary organic matter decrease on diagenesis. It is expected, therefore, that matured (altered) organic matter produces on pyrolysis a larger amount of aromatic compounds and less amounts of nitrogenous and/or oxygenated compounds than fresh (unaltered) organic matter. The fact that the pyrolysates from JT-A-D are more abundant in nitrogen compounds and less abundant in aromatic compounds than those from JT-#1-#13, indicates that

Fig. 4. Relative composition of pyrolysis products of sinking particles from Japan Trench and Sagami Bay.

243 the organic matt.er in the former is more fresh than that in the latter. Here, analytical result of sinking particles of Sagami Bay (SB) was compared with JT-ones. The sinking particles were collected a t the location of about 30 km distant from the coast. The water depth where the samples were taken, is 360 m and 560 m, respectively. Therefore, the organic matter in SB-sinking particles is expected to be more fresh than that in JT-ones. Both aromatic hydrocarbons and nitrogen compounds from JT-A-D and JT#1-#13 produces in larger amounts than those from SB ones. On the other hand, the yields of fatty acids, phenols and aliphatic hydrocarbons for JT-sinking parti-

Sagami Bay 1983 A (360m) 1983 B (560m) Japan Trench JT-01 l m + < A

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.

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.

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Fig. 5 . Relative composition of nitrogen compounds produced from sinking particles of Japan Trench and Sagami Bay.

244 cles are lower than those for SB-ones. According to our previous results (Yamamoto and Ishiwatari unpublished), fresh living organic matter (e.g., phytoplanktons) produces a larger amount of fatty acids, phenols, nitrogen compounds (pyrroles and indoles) and furans than sedimentary organic matter. It is also known t h a t kerogens from sedimentary rocks produces a larger amount of aromatic and aliphatic hydrocarbons, and less amount of oxygen compounds (fatty acids, phenols and furans) or nitrogen compounds than kerogens from recent sediments (Graas et al., 1981; Larter and Douglas, 1982; Yamamoto and Ishiwatari, 1986 and unpublished). If we ignore the yield of nitrogenous compounds, we can concluded t h a t the organic matter in JT-sinking particles is older (more matured) than that in SB-ones. As mentioned above, nitrogen compounds are produced from JT-sinking particles in a larger amount than from SB-ones (see Fig. 4). This would mean that organic matter in JT-sinking particles is younger than t h a t in SB-ones. This is inconsistent with the conclusion deduced from non-nitrogen pyrolysis products. So we examined in detail the nature of nitrogen compounds t o explain this inconsistency. Fig. 5 shows the percentage of pyrroles, indoles, pyridines and benzene cyanides in the total nitrogen compounds. As shown in Fig. 5 , JT-A-D produce indoles and pyrroles more than JT-#1-#13. The relative abundance of indoles and pyrroles for JT-A-D is closer t o that of SB-sinking particles than t h a t of JT#1-#13. This seems to indicate that organic matter of JT-A-D is more fresh than t h a t of JT-#1-#13. In contrast, JT-#1-#13 produce pyridines in a larger amount than JT-A-D and those from SB-sinking particles, indicating that organic matter in the former is more mature than that in the latter. It is known that matured organic matter tend to produce saturated hydrocarbons rather than unsaturated hydrocarbons on pyrolysis. According t o our unpublished data, the ratio of n-alkanes/n-alkenes (A/E) ratio is: 0.7-0.8 for young kerogens from lake sediments; 0.7-0.8 for organic matter in sinking particles from Sagami Bay; 0.7 for organic matter in a surface sediment of Sagami Bay and 1.1-1.4 for kerogens from marine sediments of Cretaceous age. The A / E ratios for JT-sinking particles are considerably high, being in the range of 1.2-2.2, as shown in Fig. 6. This result indicates t h a t matured organic matter is probably present in a considerable amount in JT-sinking particles. It may be concluded from the above results that organic matter in JT-sinking particles is composed of both fresh organic matter and matured one, which probably have come from bottom sediments by resuspension. 3.2.

Seasonal Variation of Organic Matter Composition of Sinking Particles

Fig. 6 shows the seasonal variation of A / E (n-alkanesln-alkenes) ratios of JTsinking particles. The A/E ratios are high for JT-#5-#7 (1.6-2.2) and J T #lo-#12 (1.5-1.7), whereas those for the other samples JT-#1-#4 (1.4-1.6), JT-#8-#9 (1.3) and JT-#13 (1.3) are low. JT-A-D give relatively low (1.2-1.3) A/E ratios, being similar each other. This means that matured organic matter is supplied to a larger amount in the periods of JT-#5-#7 (the middle of December-the beginning of January) and JT-#lo-#12 (the late of Febrary-the

245 beginning of April). Fig. 7 gives the seasonal variation of n-alkanes/indole (A/I) and nalkenes/indole (E/I) ratios. As described already, indole can be used as an indicator of fresh organic matter. Therefore, the A/I and E/I ratios may be another parameter t o test whether organic matter is mature or not. If the organic matter is mature, an A/I ratio should be high. If the organic matter is fresh, an A/I ratio should be low. An E / I ratio should similar between mature and fresh organic matter. The variations of A/I and E/I ratios are little for JT-A-D (0.44-0.63 and 0.36-0.52, respectively). This means t h a t the ratio of matured organic matter to fresh organic matter does not vary considerably for these samples. In the case of JT-#1-#13, on the contrary, the A/I ratio varies considerably (0.8-2.4), while the E/I ratio does not (0.7-1.1). In particular, the A/I ratio is high for JT-#5-#7 (1.6-2.4) and JT-#10-#12 (1.5-1.7). These results irriply t h a t the extent of contribution of matured organic matter to the total organic matter is higher for JT-#5-#7 and JT-#10-#12 than for the others. These results are consistent with those of A / E ratios. According t o Handa (1989), high values of organic carbon flux (0.7-3.2 g C mP2yr-') were obtained from the sediment trap experiments (JT-01 and J T 02) in Japan Trench. It seems difficult t o explain these high organic carbon flux value by the surface/or subsurface primary productivity because the primary productivity is thought t o be not high in this area (Handa, unpublished). Narita et al. (1990) reported that a sedimentation rates in Japan Trench and its western

1

S

2

0

3

4

N 1986

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6

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8

9

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hinUmber

1987

Month

Fig. 6.

Monthly variation of n-alkanes/n-alkenes (A/E) ratio.

246 slope are about 1000 times larger (140-160 cm/kyr) than t h a t in a deep sea area (0.23 cm/kyr) at its east site, and claimed t h a t resuspension and redeposition are actively taking place in Japan Trench area and this process have made a n apparent sedimentation rate high in this area. Therefore, the large organic carbon flux may be explained by the fact t h a t organic matter in the sinking particles is composed of fresh organic matter and resuspended organic matter.

4.

Conclusions

Organic matter in sinking particles from Japan Trench (JT-01) was analyzed by pyrolysis-capillary gas chromatography/mass spectrometry with the following results: (1) Long-chain ( > C z 0 ) n-alkanes, n-alkenes, aliphatic cyanides and methoxyphenols were absent or almost absent in the pyrolysis products, suggesting terrestrial organic matter does not considerably contribute t o total organic matter in sinking particles. (2) Sinking particles with smaller grain size (< 1 nim) produced a large amount of aromatic hydrocarbons and less amount of pyrroles and indoles than those with larger grain size (> 1 mm). This result indicates t h a t organic matter of the former particles is matured, t h a t is, altered t o a larger extent than the t h a t of the latter.

Fig. 7. Monthly variation of n-alkanes/indole (A/I) and nalkanes/indole (E/I) ratios.

247 (3) T h e ratios of n-alkanesln-alkenes a n d n-alkanes/indole a r e higher for t h e samples obtained from December t o J a n u a r y (JT-#5w#7) a n d t,hose from February t o April ( J T - # 1 0 ~ # 1 2 ) t h a n those for t h e other samples, indicating t h a t t h e percentage of “matured” organic m a t t e r in t h e t o t a l organic m a t t e r is large in these sinking particles.

References Bracewell, J. M., and G. W. Robertson, 1976. A pyrolysis-gas Chromatography method for discrimination of soil humus types. J. Soil Sci., 27, 196--205. Bracewell, J. M., and G. W. Robertson, 1984. Quantitative comparison of the nitrogencontaining pyrolysis products and amino acid composition of soil humic acids. J. Anal. Appl. pyrolysis, 6, 19-29. Crisp, T., S. Brenner, M. I. Venkatesan, E. Ruth, and I. R. Kaplan, 1979. Organic chemical characterization of sediment trap particles from San Nicolas, Santa Barbara, Santa Monica and San Pedro Basins, California. Geochim. Cosmochim. Acta, 43, 1791-1801. Danielson, N. D., and L. B. Rogers, 1978. Determination of tryptophane in proteins by pyrolysisgas chromatography. Anal. Chem., 50, 1680-1683. Deuser, W. G., and E. H. Ross, 1980. Seasonal change in flux of organic carbon to the deep Sargasso Sea. Nature, 283, 364-365. Faix, O., D. Meie, and I. Grobe, 1987. Studies on isolated lignins and lignins in woody materials by pyrolysis-gas chromatography-mass spectrometry and off-line pyrolysisgas chromatography with flame ionization detection. J. Anal. Appl. Pyrolysis, 11, 403-416 Gagosian, R. B., S. 0. Smith, and G. E. Nigrelli, 1980. Vertical transport of steroid alcohols and ketones measured in a sediment trap experiment in the Equatorial Atlantic Ocean. Geochim. Cosmochim. Acta, 46, 1163-1172. Grass, G. V., J. W. De Leeuw, P. A. Schenck, and J . Haverkamp, 1981. Kerogen of Toarcian shales of Paris basin. A study of its maturation by flash pyrolysis techniques. Geochim. Cosmochim. Acta, 45, 2465-2474. Handa, N., 1989. Symposium: Sediment trap of Japan Trench. Kaiyo, 197-203 (in Japanese). Larter, S. R., and A. G. Douglas, 1982. Pyrolysis methods in organic geochemistry: An overview. J. Anal. Appl. Pyrolysis, 4, 1-19. McHugh, D. J., J. D. Saxby, and J. W. Tardif, 1976. Pyrolysis-hydrogenation-gas chromatography of carbonaceous material from Australian sediments, Part I: Some Australian coals. Chem. Geol., 17, 243-259. McHugh, D. J., J. D. Saxby, and J. W. Tardif, 1978. Pyrolysis-hydrogenation-gas chromatography of carbonaceous material from Australian sediments, Part 11: Kerogens from some Australian coals. Chem. Geol., 21, 1-14. Narita, H., K. Harada, and S. Tsunogai, 1990. Natural radio nuclei in marine sediments. Abstract of Nihon Kaiyo Gakkai in 1990 (in Japanese). Nozaki, Y., 1989. Symposium: Sediment trap of Japan Trench. Kaiyo, 187-191 (in Japanese). Saiz-Jimenez, C., J. J. Boon, J. I. Hedges, J. K. C. Hessels, and J. W. De Leeuw, 1987. Chemical characterization of recent and buried woods by analytical pyrolysis. Comparison of pyrolysis data with 13C NMR and wet chemical data. J. Anal. Appl. Pyrolysis, 11, 437-450. Shulman, G. P., and P. G. Simmonds, 1968. Thermal decomposition of aromatic and

248 heteroaromatic amino acids. Chem. Comm., 1040-1042. Sigleo, A. C., T. C. Hoering, and G. R. Helz, 1982. Cornposition of estuarine colloidal material: organic components. Geochim. Cosmochim. Acta., 46, 1619-1626. Simmonds, P. G., G. P. Shulman and C. H. Stembrige, 1969. Organic analysis by gas chromatography-mass spectroscopy. A candidate experiment for the biological exploration of mars. J. Chrom. Sci., 7, 36-41. Suess, E., 1980. Particulate organic carbon flux in the oceans-surface productivity and oxygen utilization. Nature, 288, 260-263. Tanoue. E., and N. Handa, 1980. Vertical transport of organic materials in the Northern North pacific as determined by sediment trap experiment. Part 1. Fatty acid composition. J. Oceanogr. SOC.Japan, 36, 231-245. Tsuge, S., and H. Matsubara, 1985. High resolution pyrolysis-gas chromatography of proteins and related materials. J. Anal. Appl. Pyrolysis, 8, 49-64. Van de Meent D., J. W. De Leeuw, and P. A. Schenck, 1980a. Cheniical characterization of non-volatile organics in suspended matter and sediments of the river Rhine delta. J. Anal. Appl. Pyrolysis, 2, 249-263.. Van de Meent D., J. W. De Leeuw ,and P. A. Schenck, 1980b. Origin of unsaturated isoprenoid hydrocarbons in pyrolysates of suspended matter and surface sediments. In: Adv. Org. Geochem. 1979, A. G. Douglas and J. R. Maxwell (ed.), Pergamon Press, pp. 469-474. Wakeham, S. G., 1982. Organic matter from a sediment trap experiment in the Equatorial North Atlantic: Wax esters, steryl esters, triacylglycerols, and alkyldiacylglycerols. Geochim. Cosmochim. Acta, 46, 2239-2257. Wefer, G., E. Suess, W. Balze, G. Liebezeit, P. J. Muller, C. A. Ungerer, and W. Zenk, 1982. Fluxes of biogenic components from sediment trap development in circumpolar waters of the Drake Passage. Nature, 299, 145-147. Whelan, J. K., M. G. Fitzgerald and M. Tarafa, 1983. Analyses of organic particulates from Boston Harbor by thermal distillation-pyrolysis. Environ. Sci. Technol., 17, 292298. Yamamoto, S., R. Ishiwatari, and P. R. Philp, 1986. Pyrolysis gas chromatography-mass spectrometry of insoluble organic matter (kerogen) from a recent lacustrine sediment. Chikyukagaku (Geochemistry), 20, 39-50 (in Japanese). Yamamoto, S., and R. Ishiwatari, 1987. Pyrolysis gas-chromatography of sediment trap samples from Sagami Bay. Bulletin of Toin-gakuen Technical College, 2, 25-32 (in Japanese). Yarnamoto, S., R. Ishiwatari and PIT. Handa, 1988. Relationship of products by Py-GC to source organic material. Res. Org. Geochem., 6, 91-95 (in Japanese).

249

JT-01 R l (‘8618130

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JT-01 1 5 (‘86/11/14

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Appendix. Pyrograms of sinking particles from Japan Trench.

250

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Appendix.

(continued)

-

1/29]

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25 1

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I

Appendix.

(continued)

i

-I

Chapter 5 Modelling of Subsurface Water Circulation and Associated Dynamical Processes

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved

255

Numerical Modelling of the Philippine Sea Masahisa KUBOTA'

Abstract Abyssal circulation is investigated using a linear reduced gravity model. First, effects of the location of an inflow on the abyssal circulation is studied using an idealized simple model. It is shown that a stagnation point disappears when an inflow is located north of the latitude of a stagnation point expected from the Stommel-Arons theory in the northern hemisphere. Second, the abyssal circulation in the Philippine Sea is studied using a model with realistic coastal and topographic geometries. It is assumed that the bottom water is supplied from the main western North Pacific into the Philippine Sea through the Yap-Mariana Junction as suggested by observations. Numerical experiments are carried out for models with and without bottom topography. In results for the model without bottom topography the strong abyssal boundary current is formed along the western and northern boundaries, and a stagnation point, not formed. Results from the experiment including a topographic effect show that abyssal circulation in the eastern Philippine Sea (the Shikoku and West Mariana Basins) is more active than that in the western Philippine Sea (the Philippine Basin). It is encouraging to see that a modelled circulation pattern compared well to observational results obtained by mooring current meters and hydrographic data in spite of the simplicity of the model.

1.

Introduction

Dynamical framework about abyssal circulation was given by Stommel and Arons (1960) many years ago. Since they focused on global abyssal circulation, the model was highly schematic and did not include detailed features in each ocean, for example the Philippine Sea. Our understanding of abyssal circulation has been slowly developed because there were many difficulties in measurements of abyssal currents. It took a long time to test and break through the Stommel-Arons dynamics. However recent developments of various kinds of observational techniques lead to do such things. Especially we can detect absolute velocities by long-term direct current measurements using moored current meters. Since differences in temperature and salinity at great depths are quite small, it may be very difficult to determine a reference level and obtain exact geostrophic current. Thus, direct current measurements are very important in investigating the abyssal circulation. The Philippine Sea, which is located in the western part of the North Pacific, mainly consists of the Shikoku, the west Mariana and the Philippine Basins . *School of Marine Science a n d Technology, Tokai University, Shimizu, Shizuoka, 424 J A P A N

256 The geometry of the Philippine Basin is shown in Fig. 1 where areas with depths shallower than 3500 m are shaded and the thick shaded area represents lands. Though the abyssal Philippine Sea is connected to the main North Pacific Basin only through a few narrow passages, recent observational results from the direct current and hydrographic measurements indicate that the remarkable circulation exists in the Shikoku Basin (Fukasawa et al., 1986, 1991) and its abyssal water is supplied through the Yap-Mariana Junction (Uehara and Taira, 1990). However, observational d a t a are still far short at great depths in the Philippine Sea to understand the whole abyssal circulation field. Thus, it is important and necessary to supplement the observation by modelling studies of the abyssal circulation. Moreover we can obtain useful information from model results for designing observation systems.

L O N G I TUDE Fig. 1. Map of the Philippine Sea. 1. Nankai Trough; 2. Ryukyu Trough; 3. Bashi Channel; 4. Yap Trench; 5. Palau Trench; 6. West Caroline Basin; 7. Yap-Mariana Junction

257 In modeling the Philippine Sea, our basic idea is to develop as a simple model as possible. There exist several lacks in the Stornmel-Arons model in order to investigate the abyssal circulation in the Philippine Sea. Kuo (1974,1978) investigated the topographic effect on the deep circulation and pointcd out that the resulting deep flow is quite different from the circulation for a flat bottom. His results suggest that it is crucial to include the bottom topography in the numerical model in order to reproduce the deep circulation pattern of the Philippine Sea. Besides this, parameterization of the vertical velocity may be most important in modeling the abyssal circulation. A uniform upwelling was assumed in the Stommel-Arons model. But this assumption is effective only for the interior region and may be not so realistic in the real ocean. Kawase (1987) successfully developed the deep ocean circulation model which determines upwelling internally by including a Newtonian damping term in the continuity equation. Kubota and Ono (1992) and Masuda and Uehara (1992) investigated the abyssal circulation in the Philippine Sea using basically the same as his model. Masuda and Uehara (1992) indicated t h a t the Newtonian-damping formula can be easily derived from the equations for a continuous stratified ocean and proved the parameterization t o be quite useful. Lastly it is an important feature of the abyssal circulation in the Philippine Sea to be forced by a zonal inflow through the Yap-Mariana Junction. A similar situation can be found in the Central Indian Basin where the lower deep water derives from the deep western boundary current in the West Australian Basin by overflow across the Ninetyeast Ridge. Warren (1982) studied the circulation pattern in this basin by the Stommel Arons dynamics. Also Ono and Kubota (1990) studied dependence on several conditions, e.g., the location of an inflow region, and the resulting circulation pattern. It is the purpose of this paper to construct an abyssal circulation image in the Philippine Sea by mainly using results from Ono and Kubota (1990) and Kubota and Ono (1992). Results of simple mechanistic models are presented in Section 3 . Numerical modeling of the Philippine Sea and the results are shown in Section 4. Discussions and summary are given in Section 5 .

2.

Simple Mechanistic Model

A linear reduced-gravity transport model is used in this study. Recently this type of model has been successfully adopted for the research of the upper tropical ocean dynamics (e.g., Kubota and O’Brien, 1988). This remarkable success depends on the special feature about the ocean structure in the tropical ocean, for example the existence of the shallow and sharp thermocline. Assuming that the thickness of the lower layer is infinite, the dynamics associated with the first baroclinic mode can be simply described by the shallow water equation about the upper layer thickness. This simplicity may be one of important factor of the model. It should be noted t h a t the Stommel-Arons model, which has been considered as a prototype of the abyssal circulation model, represents only the lower layer of the two-layer model. Thus, the similarity between the two models, the tropical reduced-gravity model and the Stommel-Arons model, points out the possibility of an application of the reduced gravity model t o the study of the abyssal

258

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Circulation patterns driven by an inflow (On0 and Kubota,

circulation. The basic equations on an equatorial beta-plane in a non-dimensional form are

U and V are the eastward and the northward components of non-dimensional volume flux; is the nondimensional displacement of the interface from the rest Py) is the nondiposition and H M is the non-dimensional depth. f M (= f o

+

259 mensional Coriolis parameter; OG = g'h/pXL3; O A = A H / , B L ~O,R = X/PL are non-dimensional numbers. AH (= lo6 cm2/sec) is the horizontal eddy viscosity coefficient; X is the Newtonian damping coefficient; g' is the reduced gravity; h is the interfacial displacement; p is the gradient of the Coriolis parameter and L is the horizontal characteristic scale. The nondimensional parameter, O R is the ratio of the damping time to the inertial period, O A the ratio of the Munk boundary layer t o the basin width, and OG the ratio of the damping scale to the transit time of the long Rossby wave across the basin, respectively. The gravity wave speed is represented by ( H M O G O R ) ' / ~ . Fig. 2 shows interfacial displacement and velocity fields at a steady state for the various inflow latitudes. The model basin covers from the equator t o 30"N. The source in the Fig. 2(a) is located on the equator along the western boundary. This corresponds t o the case of the North Pacific Ocean in the Stommel-Arons model. On the other hand the source is located on the eastern boundary in the Fig. 2(b)(d). A stagnation point in the basin is expected to exist a t y = 0.5 from the Stommel-Arons dynamics. The course of the flow from the source region is based on the conservation of potential vorticity. Therefore the abyssal water entering from the eastern boundary flows westward in a zonal jet along the contour line of ambient potential vorticity (Pedlosky, 1979). However a weak northward flow in the whole interior region can be found in order t o compensate vortex stretching associated with the upwelling. The circulation pattern shown in Fig. 2(b) is quite similar t o the southern part of the schematic pattern given by Warren (1982). There does not exist a stagnation point in Fig. 2(c) and (d), and the western boundary current flows southward everywhere. Since Warren (1982) pointed out the similar characteristics by using the Stommel-Arons model, this feature is independent of effects of the Newtonian damping. It should be noted t h a t the western boundary current always flows southward when the source is located around or north of the latitude of the stagnation point. Also the location of the stagnation point is unvaried when it exists.

3.

Numerical Modelling of the Philippine Sea

The model used here is almost same as that in the previous section. However dimensional equations are used in this section for the purpose of simulating the abyssal circulation pattern in the Philippine Sea. The basic equations are

For the present study, we adopt g' = 0.196 cm/s2, X = 1/1000 days-' and AH = lo6 cm2/sec. The Newtonian damping is associated with the vertical diffusion and can easily be derived from the equations for a continuously stratified ocean (Masuda and Uehara, 1992). However the coefficient value is not so definite. The

260 Newtonian damping has the effect to freeze the flow a t the point of set up of the eastern boundary current because the long Rossby waves are damped (Kawase, 1987; Ono and Kubota, 1990). The Stomniel-Arons model can be considered t o be the limit of negligible decay of long Rossby waves. The choice of t,he Newtonian damping value is crucial for the circulation pattern if we adopt the large value for a Newtonian damping coefficient. However such a situation may be not so realistic. Thus, in order to avoid such a situation the decay time of 1000 days is used in the present study considering t h a t the traveling time of the long Rossby wave to cross the abyssal Philippine Sea is about 500 days at the latitude of the source. We also carried out a case with the damping time of 3300 days as an additional experiment. However the final flow pattern (not shown here) was not so different from the pattern for the case of the damping time of 1000 days. Only the Philippine Sea deeper than 3500 m is considered, and hence the model sea surface represents the interface at 3500 m in the real ocean. Since we are mainly

T=

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Fig. 3. Displacements and velocity fields in the flat model after 10000 days. The contour interval is 1 m.

26 1

T= 8000 ( D A Y S 1 !..........I. ............................................

m-

................. . . . . . . . . . . . . . . .&) . . . ........................... ......................

........................................... I

125

I

I

130 135 LONG I TUOE [ E l

I

140

Fig. 4. As in Fig. 3 except in the model with bottom topography after 8000 days. interested in the bottom water circulation, we selected 3500 m as the interface. However, this value of 3500 m is fairly arbitrary. We use 10 km as a grid size with the Arakawa C grid (Mesinger and Arakawa, 1976) t o reduce computer core storage. Thus, we can include topographic effects by using a very fine grid size. The equations are integrated in time using a forward-backward scheme. We concentrate ourselves on a steady state in the present study. From estimation of the traveling time of the baroclinic Rossby wave t o cross the abyssal Philippine Sea, we may realize that the model should be integrated longer than 4000 days at least in order to get the steady state. Our model is forced by the supply of the abyssal water through the Yap-Mariana Junction as indicated by the observations (Uehara and Taira, 1990). The transport flowing into the Philippine Sea through the Yap-Mariana Junction is assumed t o be 0.3 Sv. The transport value is not definite and should be made clear by observation in the future.

262 4.

Results

First, a case of flat bottom is studied as a preliminary experiment. A result after 10000 days is shown in Fig. 3 . Solid lines indicate contour lines of the displacements of the interface. Velocity vectors are represented by arrows. Larger displacements are found in the Philippine Basin and especially t h a t around the Bashi Channel is the largest. The general features of the velocity field are t.he weak north-eastward flow and the strong western and northern boundary flows as a complement. The stagnation point in the present model is expected at 12.8"N from the StommelArons dynamics, while the inflow region extends from 10"N to 13"N. Therefore, there does not exist a stagnation point in the present model as shown in Fig. 2(c). Fig. 4 shows the result of the model with bottom topography after 8000 days. An extremely complicated pattern compared with Fig. 3 is seen, but an anticlockwise circulation over the whole area is common. We can find the remarkable northern and western boundary currents in Fig. 3, while such strong currents do not continuously but intermittently exist in Fig. 4. As a whole the circulation in the eastern part (the Shikoku and the West Mariana Basins) is more active than that in the western part (the Philippine Basin). But it should be noted that the strong flow can be exceptionally found east of the Philippine Islands even in the Philippine Basin. The inflow through the Yap-Mariana Junction splits into two main branches. One flows along the eastern boundary and the other flows north-westward along the Kyushu-Palau Ridge which divides the Philippine Sea into the eastern and western parts. The first branch is considered to be the bottom-controlled current due t o the topographic ,!?-effect, while the second branch is the western boundary

Fig. 5 . Trajectories of markers released at the inflow region.

263 current due to the planetary P-effect. The former turns to the west in the Shikoku Basin and joins the latter around the Ryukyu Trench a t 130"E, 25"N. The latter involves the anticlockwise recirculation flow in the Philippine Basin on the way to the north-western boundary. Such a circulation pattern can be supported by the trajectory of markers put on the source (Fig. 5 ) . Displacements of the interface are quite large along the eastern and north-eastern boundaries. This result suggest,s that the upward motion there is active. We can easily convert the value of the displacement into t h a t of the upward velocity. The largest value of the upward velocity is 1 . 6 ~ 1 cm/s 0 ~ ~and the average value is 1 . 0 ~ 1 0 -cm/s ~ in the Philippine Sea.

5.

Comparison with Observational Results

High quality CTD d a t a and Co-operative Study of the Kuroshio and Adjacent Regions (CSK) standardized chemical tracer d a t a in the western North Pacific are analyzed by Kaneko (1984). Fig. 6 shows dynamic topography on the 2000 d b surface referred t o the 4000 d b surface. The pressure map shows a rise of the isopycnal along the western boundary of the Philippine Sea, which suggests northward flow when the deeper surface of no motion is assumed. Also the circulation in this figure is clockwise while the model results in Figs. 3 and 4 give the circulation to be anticlockwise. An interesting feature about the structure of the deep mean current is pointed out from results of the direct current measurement in the Shikoku Basin

Fig. 6. Maps of geopotential anomaly (dyn*m). 2,000 d b surface relative t o the 4,000 d b surface. Area with depths less than 2,000 m are shaded, and the light line is the 4,000 m contour (Kaneko, 1984).

264 (Fukasawa et al., 1986). They showed t h a t the mean current increases with depth and that the vertical extent of the mean current is estimated t o be about 2000 m above the sea bottom. This leads t o the idea that the 2000 d b surface t o be no motion level, rather than the 4000 d b surface. Thus Fig. 6 can be interpreted as the dynamic topography on the 4000 d b surface referred to the 2000 d b surface. Consequently Fig. 6 shows an anticlockwise circulation on the 4000 d b surface which is consistent with the model results. Techniques for the mooring current meters have been developed since the 1960s. Though this technique has the disadvantage t h a t measurements are made a t one location only, it is very useful because mooring current meters can provide absolute velocity of various depths. Direct current measurements have been recently carried out in the Philippine Basin (e.g., Fukasawa et al., 1986). Here we compare those results with the velocity field given by the present model results. The extended maps are given in Fig. 8 together with observational results shown by the thick arrow. The location of the extended maps is indicated by the square in the Fig. 7. Fig. 8(a) represents the south West Mariana Basin. Yoshioka et al. (1988) made a direct current measurement from July 1985 to July 1986 in the West Mariana Basin. They demonstrated that the abyssal current flows southward throughout the observed period and the average speed (0.8 cm/sec) was not so large. They continue the direct current measurement a t the same location after July 1987. The average speed after July 1987 was also weak, less than 1 cm/sec, but the current direction during the first year was west by south and during the second year was north-east (Yoshioka, private communication). On the other hand the current at

Fig. 7.

Locations of the extended maps.

265

:++ 1.0

135

LONG1 T U D E

(El

140

Fig. 8. Displacements and velocity fields in the extended region indicated in Figure 7. A thick arrow indicates observed current speed. A thin arrow indicates model current speed. the same location in the present model is also so weak and flows northward. The feature that the current at this location is very weak is in coinmon between the observation and the model result. However the current direction is not necessarily the same between them because the observed current direction was not stable. Also direct current measurements were carried out in the Bashi Channel (Fig. Sb), the Ryukyu Trough (Fig. 8c) and the Nankai Trough (Fig. 8d). Observational results strongly suggest the existence of a steady westward abyssal flow along the northern and the northwestern periphery of the Shikoku Basin (Fukasawa et al., 1986). Moreover there exists an eastward abyssal flow indicated by the arrows of SB1 and SB3 off there (Kawabe, private communication). The model current field demonstrates similar feature as the direct current measurement does concerning not only the current direction but also the relative amplitude. For example both results of the current measurement and the model in the Bashi Channel shows a southward and very weak flow. On the other hand those in the Ryukyu Trough shows a southwestward and moderate flow. It should be noted that these remarkable features are well reproduced by the model current field.

266

5% 1.0

125 130 LONG1 TUDE ( E l Fig. 8.

(continued)

Sudo (1986) described time series of spatial distributions of potential temperature and dissolved oxygen concentration below 3,800 m in the Shikoku Basin during 1975-1983. He pointed out t h a t the bottom water mass enters the Shikoku Basin from the south a t early stage and then slowly spreads clockwise in the northern area at least north of 31"N. The northward inflow of the bottom water into the Shikoku Basin can be clearly seen in the present model results (Fig. 4). Moreover even the clockwise circulation in the north Shikoku Basin is also indicated by the extended map as shown in Fig. 8(d).

6.

Abyssal Oxyty Distribution

Fig. 9 shows property distributions on the isopycnal surface U Q = 27.77, which lies at about 3,000 m (Kaneko, 1984). We can easily find a maximum region in the north-western Shikoku Basin and a minimum region in the south-eastern westMariana Basin in the oxyty map. The latter suggests a deep inflow through the Yap-Mariana Junction since the water of low oxyty is found in the North Pacific. This low-oxyty region distributes along the western boundary and crosses the West Mariana Basin as shown by the contour line of 3.2 ml/l. Velocity fields of the model results are surprisingly consistent with this feature of the low-oxyty. On

267

Fig. 8. (continued) the other hand the former cannot be explained from the viewpoint of a horizontal circulation because there does not exist high oxyty on the same isopycnal surface around there. Thus, Kaneko (1984) concluded that the high oxyty water is formed through vertical mixing of abyssal water. The model result also indicates the fairly strong upwelling in the north-western Shikoku Basin. However strong upwelling regions are not necessarily confined there and also exists along western and northern boundaries in the model results. Kaneko (1984) suggested from water mass analysis that abyssal water below 3000 m in the Philippine Sea may come from the main North Pacific Basin. The oxyty at the depth of 3000 m in the western North Pacific is lower than that in the Philippine Sea as shown in Fig. 9, while that near bottom in the western North Pacific is expected to be higher because there is no mechanism creating deep sea water in the Philippine Sea. Thus the lower oxyty water flows into the Philippine Sea at the depth of 3000 m as shown in Fig. 9 while the high oxyty water does at the deeper ocean. The high oxyty water gradually upwells and mixes with the upper low oxyty water as spreading horizontally by the current expected from the model results . Consequently the observational results as shown in Fig. 9 may be consistent with the current fields of the present model. However it is impossible to explain the high oxyty distribution indicated by the contour of

268

0 3

c H

t

<

J

135

s++

LONG I TUOE

(El

1.0

Fig. 8. (continued)

3.3 ml/l in the southern Philippine Sea by the current velocity fields of the present model. The high oxyty water may feed by the inflow south of the Philippine Sea except t h a t through the Yap-Mariana Junction. Whether it is the case or not will be clarified by the future observations. Ono (1990) investigated the abyssal oxyty distribution in the Philippine Sea using the horizontal advection-diffusion equation. T h e model current velocity shown in Fig. 4 is adopted as the advection velocity in his model. Vertical diffusion is neglected in his model. Since he considered the oxyty distribution only on the isopycnal surface go = 27.7, the low oxyty water (3.5 ml/l) from the main Pacific flows into the Philippine Sea which is filled with the high oxyty water (4.0 ml/l). Fig. 10 shows the oxyty distribution after 32850 days. The relatively high oxyty regions can be found in the Shikoku Basin and south of the Philippine basin. Such features are also identified in Fig. 9. This similarity suggests a possibility t h a t the high oxyty water in both regions can be explained without vertical motion. However Fig. 10 presents not the steady state but the transient state. The final oxyty distribution will be uniform because the decay or the vertical advection is not included in his model. The understanding of the mechanism of both high oxyty regions is a topic of future research.

269 7.

Summary and Discussions

An abyssal circulation model of the Philippine Sea has been developed. T h e present model was created primarily for constructing an image of the abyssal circulation of the Philippine Sea. T h e circulation produced by the model is also very useful for the observational array design as direct current measurements. T h e present model is based on a linear, shallow water equations representing only one baroclinic mode. Such a model has been successful t o study the equatorial dynamics. T h e success of the equatorial model strongly depends on a special hydrographic feature of an equatorial ocean such as the sharp and shallow thermocline. Though there does not exist such remarkable thermocline in the abyssal ocean, the results from the present model are encouraging. T h e current fields compare t o the past observational results including not only direct current measurements but also hydrographic data. The model current fields surprisingly reproduce the results of direct current measurements though the model is very simple. This remarkable similarity between them confirm the model current fields t o be qualitatively realistic. This success of the present model may be associated with several intrinsic features of the abyssal circulation. Since heat and momentum exchanges between ocean and atmosphere through the sea surface are essential for the surface circulation, it is impossible t o understand a mechanism of the surface circulation without knowledge about sea surface flux. Also the motion itself in the surface layer is very active, and this makes the surface circulation fairly complicated. O n the other hand the abyssal circulation is basically side-boundary forcing because the abyssal 10'

30'

20'

12O'E

130'

150'

160'

170'

Fig. 9. Oxyty distributions on the a0 = 27.77 surface. Area with depths less than 3,000 m are shaded (Kaneko, 1984).

270

Fig. 10. The distribution of the oxyty after 32850 days, calculated with the model velocity field shown as Fig. 4. water is formed in only polar regions. Also the abyssal flow is generally not so strong and complicated as the surface current. Moreover it should be noted that bottom topography directly affects the abyssal circulation, especially the bottom circulation. These features of the abyssal circulation suggests the simplicity of the abyssal circulation. The success of the simple reduced-gravity model in the present study is due t o the simplicity of the abyssal circulation. It may be more important t o have recognized the simplicity of the abyssal circulation than to have well simulated the observational features. The pattern of the abyssal circulation in the Philippine Sea given in the present study is tentative. There still exists several disagreements between the model and observational results. First, the current speed in the model is very weak compared with that by direct current measurements. This difference is caused by resolution in the not only horizontal but also vertical directions. Because a layer below 3500 m is considered to be only one layer in this study, currents are considered to represent average values below 3500 m. Thus, it appears t o be expected that model currents are so weak compared with the result of direct current measurements. In order to

27 1 compare quantitatively between them, more sophisticated model must be required. Second it is impossible t o explain the high oxyty distribution in the southern Philippine Sea (Kaneko,1984). T h e high oxyty region suggests the possibility t h a t abyssal water flows into the southern Philippine Sea from the West Caroline Basin. The West Caroline Basin is almost disconnected with the main Pacific except the Yap Trench. Thus, we may consider t h a t the southward abyssal water originated from the western North Pacific passes through the Yap and Palau Trenches and flows northward into the southern Philippine Sea. It is necessary t o extend the model domain in order t o confirm the above abyssal circulation. Finally the ability of the linear reduced-gravity model t o well produce a n abyssal circulation pattern is demonstrated in this study. T h e results presented here are encouraging for extending the model over the whole Pacific. Results from the whole Pacific Ocean model will be presented in the near future.

Acknowledgement This research has been supported by Priority Area Programme, Dynamics of the Deep Circulation of the Ministry of the Education of Japan. T h e authors wish t o thank Drs. S. Horibe, N. Suginohara and M. Fukasawa for their many useful comments. We would like t o thank K. Ono and A. Ishii for their assistance during the study.

References Fukasawa, M., T. Teramoto and K. Taira, 1986. Abyssal current along the northern periphery of Shikoku Basin. J. Oceanogr. SOC.Japan, 42, 459-472. Fukasawa, M., T. Teramoto, K. Taira, and M. Kawabe, 1991. Hydrographic structure in association with deep boundary current in the North of the Shikoku Basin. Submitted to J. Oceanogr. SOC.Japan. Kaneko, I., 1984. Structure of Mid-depth Water in the Philippine Sea. Doctor’s thesis, University of Tokyo, 97 pp. Kawase, M., 1987. Establishment of deep ocean circulation driven by deep-water production. J. Phys. Oceanogr., 17, 2294-2317. Kubota, M., and J. J. O’Brien, 1988. Variability of the upper tropical Pacific ocean model. J. Geophys. Res., 93, 13930-13940. Kubota, M., and K. Ono, 1992. Abyssal circulation model of the Philippine Sea. Deep-sea Res., 39, 1439-1452. Kuo, H. H., 1974. The effect of bottom topography on the stationary planetary flow on a sphere. Deep-sea Res., 21, 933-945. Kuo, H. H., 1978. Topographic effect of the deep circulation and the abyssal oxygen distribution. J. Phys. Oceanogr., 8, 428-436. Masuda, A., and K. Uehara, 1992. A reduced-gravity model of the abyssal circulation with Newtonian cooling and horizontal diffusion. Deep-sea Res., 39, 1453-1479. Mesinger, F., and A. Arakawa, 1976. Numerical Methods Used in Atmospheric Models. GARP Publ. Ser. 17, Vol. 1, World Meteorol. Organ., Geneva, 64 pp. Ono, K., 1990. Abyssal Circulation in the Philippine Sea. Master’s thesis, Tokai University, 135 pp. (in Japanese) Ono, K., and M. Kubota, 1990. Stagnation point in the abyssal circulation. J. Fac. of Mar. Sci. and Tech. Tokai Univ., 31, 117-132. Pedlosky, J., 1979. Geophysical Fluid Dynamics. Springer-Verlag, New York, 624 pp.

272 Stommel, H., and A. B. Arons, 1960. On the abyssal circulation of the world ocean. An idealized model of the circulation pattern and amplitude in ocean basins. Deep-sea Res., 6, 217-233. Sudo, H., 1986. Deep water property variations below about 4,000 m in the Shikoku Basin. La mer, 24, 21-32. Uehara, K., and K. Taira, 1990. Deep hydrographic structure along 12'N and 13'N in the Philippine Sea. J. Oceanogr. SOC.Japan, 46, 167-176. Warren, B. A,, 1982. The deep water of the Central Indian Basin. J. Mar. Res., 4O(Supplement), 823-860. Yoshioka, N., M. Endoh and H. Ishizaki, 1988. Observation of the abyssal current in the West Mariana Basin. J. Oceanogr. SOC. Japan, 44, 33-39.

Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved

273

The Occurrence of Vacillation in a Model Ocean Driven by Wind, Heat and Salinity Flux Kensuke TAKEUCHI* and Yuji KASHINOt

Abstract

A simple, idealized model ocean, which roughly models the north Pacific Ocean, is driven by wind stress, heat and salinity flux through the sea surface. Different formulations are used for the heat and salinity flux, Newtonian damping type for heat flux versus fixed flux for salinity. The model starts vacillating with a period of about 23 years, even though the boundary conditions remain constant. It takes place in the northern portion of the model ocean, where stratified and deep convected state appear alternatively. Dependence of this vacillation on several parameters are examined.

1.

Introduction

Driving mechanism of ocean circulation has been studied by many researchers for a long time. However, except for purely wind-driven cases, driving mechanisms are poorly understood. Especially, combined effect of wind, heat and salinity forcing is hard t o understand, as the process become highly non-linear. So far, numerical modelling seems to be the only possible weapon t o attack the problem. Recently, number of numerical experiments are performed in attempts t o throw light on this driving mechanism. In those experiments, curious behaviors of models are found when so called ‘mixed boundary condition’ is used. In the past, newtonian damping type boundary conditions are used usually for both heat and salinity fluxes through the sea surface. This boundary condition is useful for simulation type modelling, as it tend t o restore model t o observed climatology. However, from physical point of view, there is no rationale t o use this type of boundary condition t o salinity flux. For heat flux, it is reasonable t o use this type of formulation, as there are feed-back mechanism, for example, if sea surface temperature become too high, outgoing heat flux increases and tend t o cool it down. However, this kind of feedback mechanism does not exist for salinity flux. Even if sea surface salinity become too high, it does not reduce the evaporation nor enhances rainfall if the salinity remain in realistic range. For this reason, researchers start using fixed salinity flux at the sea surface in their models (Bryan, 1986a,b; Manabe and Stouffer, 1988; Marotzke e t al., 1988; Weaver and Sarachik, 1990). This mean t h a t different type of boundary conditions are used for *Department of Geophysics, Faculty of Science, Hokkaido University, Sapporo, 060 J a p a n + J a p a n Marine Science and Technology Center Natushima-cho, Yokosuka-city, 237, J a p a n

274 heat and salinity, as damping type formulation is used for heat flux, so it is called ‘mixed boundary condition.’ In most of these experiments, damping type of flux formulations are used to reach steady states, and then the flux at the steady states are fixed t o continue the calculations. Even though the same flux are given at the sea surface, the model oceans become unstable in various ways. They are categorized into two types of instability. In the first type, which occurs only in model oceans extends t o both hemisphere symmetrically, one of the meridional gyre circulations becomes dominant and pushes away the other beyond the equator and finally occupies all the domain, so that formation of deep water takes place only in one of the hemispheres. The other type is termed as ‘polar halocline catastrophe’by Bryan (1986a), in which overturning circulation a t high latitudes reduces significantly. This ‘catastrophe’ commonly happens in those models using mixed boundary condition, but their appearances differs from model t o model. In some model, this weak convection state lasts to the end of the calculation, while in some other model it vacillates between week and strong overturning states irregularly. The present model experiments also use the mixed boundary condition. The purpose of the present experiments are to get better understanding about the circulation of the North Pacific Ocean. Especially we focused on intermediate depth circulation, the circulation of cold and fresh water formed at high latitudes of the North Pacific Ocean. Accordingly, the present model is not designed to simulate deep convection and formation of deep water. In that mean, the present model

pk

A-

6 4 0 0 km

Fig. 1. Schematic diagram of the model ocean.

275 differ from others. Still the ‘polar halocline catastrophe’ also occurs in the present models, but somewhat in a different way. In the present models, rather regular oscillation between two states takes place. In the present paper, the oscillation is described and analyzed. Also, the dependence of this oscillation on parameters like horizontal diffusivity, model horizontal resolution and strength of wind driven circulation is examined.

2.

Model

Fig. 1 shows the schematic view of the model ocean, which roughly models the North Pacific Ocean north of 10’N. For simplicity, the curvature of the earth surface is ignored, except that beta-plane approximation is used for the Coriolis parameter. The governing equat,ions are ones used commonly in oceanography, with hydrostatic, rigid lid and Boussinesq approximations. The coefficients for the eddy viscosity and the eddy diffusivity are set constant, and are 1.5 and 1.0 for vertical direction, l.0x109 and 1 . 5 ~ 1 0for ~ horizontal direction (all in CGS unit), respectively. We employ the formula proposed by Mamaev (1964) for the calculation of sea water density (Eq. 1). p = 1028.14 - 0.0735T - 0.0469T2

+ (0.802 - O.O02T)(S - 35.0)

(1)

where p , T and S are density, temperature and salinity of sea water, respectively. No heat and salinity fluxes are allowed through any boundary except the sea surface. The northern and southern walls are slippery, while non-slip condition is assumed at the eastern and western walls, as well as ocean bottom. The model ocean is driven by momentum, heat and salinity fluxes through the ocean surface. Wind stress has only zonal component and vary only in the meridional direction. For heat flux, Haney type formulation (Haney, 1971) is used in the present study, H = r(Ts - Ta) (2) where H is heat flux into the ocean, T, is model sea surface temperature and y is the coefficient, which is set t o 70 (cal/cm-’/day) in the present model. T, is equilibrium temperature and is given as a function of latitude only, as shown in Fig. 2. As discussed in the Introduction of this paper, salinity flux is given directly as a function of latitude in this model, rather than using damping type formula. The distributions of zonal wind stress, the equilibrium temperature and salinity flux are shown in Fig. 2. At the initial state, the ocean is at rest and temperature and salinity are uniform horizontally. The way that the North Pacific Ocean is modeled in the present study, the model is not designed to simulate the formation of deep water. The numerical scheme is similar to t h a t explained in Bryan (1969). T h e model ocean has 10 levels and the vertical resolution is 25 m near the sea surface. The model ocean is spun up using coarse grid model with horizontal mesh size 320 km for 500 years. The integration continued for additional 150 years using finer grid model with mesh size 160 km, which is to be mentioned as the basic model hereafter.

276 I .o

v)

"

0.0

v1

- I .o

Fig. 2. Zonal distributions of boundary conditions used in this model. Solid line: Equilibrium temperature ( T a in Eq. 2). Dash line: Wind stress in dyn/cm (positive eastward). Dot dash line: Precipitation minus evaporation in one year (cm).

01.5"

t1J:OA 281.5"

'2S.O"

bO0.U"

I '00

1

0

100

200

300

I

500

I 600

TEAR

Fig. 3. Time evolution of averaged salinity in each of top 6 layers of the model. At year 500, when finer mesh model takes over the integratAon, the model ocean starts vacillating.

3.

Vacillation

After we switched t o the finer grid model, the model ocean start vacillating despite the boundary conditions being held constant. Fig. 3 shows evolutions of averaged salinity in each top 6 layers of the model. The period of the vacillation is

277 about 23 years, and it reaches as deep as 6th layers (600 m). Fig. 4 shows the horizontal distribution of salinity in the surface layer at several phase of one cycle. At the beginning of the cycle, a t year 627 (Fig. 3a), there is a tongue of a low salinity water extending from the north-western corner of the model ocean t o the east. It gradually collapse and at year 638 (Fig. 4e), it disappears nearly completely. Then it reappears at the north-western corner, extends t o the east again and by year 648, it returns to similar state a t the beginning. During the cycle, the southern half of the model ocean remains nearly steady. Density, temperature and salinity distributions in the zonal plane along 50.4"N a t the two extreme phase of vacillation are shown in Fig. 5. At year 628, when

SllL 1 N I T I

627.0TERRS

2S.OH

SOL I N 1 T I

630.5YEflRS

28.0~

(d)

(C) S A L I NI T Y

631.5YERRS

SOL IN1 T I

2S.OH

50

SO"

vo

YO

632.0YERnS

28.m

Y

3

30

J

30s

20"

20

10

IOH

"

OlllAHEt lnnl

0

I600

3200 OISIWiiE

UBOO

6"DO

IKnl

Fig. 4. Horizontal salinity distribution in t,he surface layer, at years (a) 627, (b) 630.5, (c) 631.5, (d) 632, (e) 638, ( f ) 641, (g) 645 and (h) 648, respectively. The contour interval is 0.1 psu.

278 the fresh water tongue is a t full development, there is a mass of warm and saline water in subsurface layers, forming sharp transition layer between t h a t water mass and fresh and cold surface layer. At year 633, the upper 600 m of the ocean is well mixed vertically, and water in the surface layer becomes colder and more saline.

4.

Mechanism of Vacillation

To inspect the mechanism of the vacillation, the cycle of sea water properties in the upper 2 layers at the location where amplitude of vacillation is largest, are displayed on T-S diagrams (Fig. 6). After strong vertical mixing due to the initiation convection (indicated by arrows in the figures), sea water become cooler and fresher by cooling and freshwater flux through the sea surface. At the beginning, cooling effect is stronger and den-

SRLINI TT

638.OYERRS

28.011

Fig. 4.

SRLI N l l Y

(continued)

6'11.01ERRS

2S.On

279

0

200

g

'100

x c 600

800

loo0 I

0

t

16W

TEHPE?+W&E

(b

,

UBOO

3200

mmix

6VOO

inmi

628.0

YERRS L R T 1 1U O E = 50. Y N

0

I

1600

SAL IN177

(C)

I

xao DIST'IIE

VBOO

I GVOO

lKHl

628.0

TERRS LRTITUDE=50.4N

,000 0

I

1600

1 3200

01SlnuTE

UBOO

I-

6'100

,*XI

Fig. 5. Density (a,d), temperature (b,e) and salinity (c,f) in the zonal plane along 50.4"N at year 628 (a,b,c) and 633 (d,e,f), respectively. sity of sea water increases. However, as the temperature decreases and approaches the equilibrium temperature, cooling reduces while salinity flux remain the same. Eventually, the effect of salinity flux become larger, and density in the surface

280 POTENTIRL DENSITY

6 3 3 . 0 YEflRS LRTITUDE=SO.~N

DtSlRNCE

TEMPERRTURE

(el

LRTllUOE=50.4N

Q

I

I

I

1600

UB 0

3200

oIsin'jcE

(f)

IKMl

6 3 3 . 0 iERFi5

SRL 1 N I T Y

6'100

IKHI

6 3 3 . 0 YEflRS LnTITUflE=SO.4N

I 3200 '1800 6'100

00

L l l S l n l i C L Inn,

Fig. 5 .

(continued)

layer start decreasing. Then, convection stops and strat,ification start to develop. In the surface layer, both temperature and salinity keep decreasing, while subsurface water becomes warmer and more saline as it is cut off from the effect of surface boundary condition and fluxes from lower latitudes become dominant.

28 1

(b )

(a) TRAJECTORY

627.0 To 650~0yEARS

(1920KH. 50.4N.

25.0H13

5hLlNITT (PER H I L L € I

?PhJECTORY

F R O M 6 2 7 . 0 T O 650.0YEhRS

(1920KM. 5 0 . 4 H .

ShLlliITT

87.5M)

(PER H l L L E l

Fig. 6. Trajectories of water property in T-S diagram, for the top two layers at 50.4"N and 1920 kni from the western boundary. It is more difficult to understand the process in the opposite phase. After the surface water reaches the lowest density state in the cycle, suddenly salinity start increasing. Upon examining the heat and salinity budget, it is found that horizontal diffusion is dominant in this stage. Although temperature also increases a t the same time, the effect is not strong enough to compensate the effect of salinity. Thus, density difference between the surface and the subsurface water reduces and finally vanishes to start overtnrning again. Once a convection occurs at one location, it induce a sharp horizontal gradient in temperature and salinity between the convected water column and stratified water colunin adjacent to it. That increases horizontal salinity and heat flux by diffusion into the stratified water column, causing another overturn there. Thus stratified region reduces and the whole system goes back to the initial phase.

5.

Dependence on Parameters

In this section, the dependence of the vacillation on horizontal diffusivity coefficient, grid size and wind driven circulation is examined. For this purpose, several additional models with different parameters are carried out.

5.1.

Horizontal diffusivity coefficient

Two models, in each of which horizontal diffusivity coefficient is twice as big as, or a half of t h a t of the basic model, are tested. The high coefficient model uses the year 600 of the basic model as the initial condition when the cold and fresh water tongue is at maximum development. The model goes back to the convective mode, and stays there for the rest of experiment, without showing any sign of restarting the vacillation again. In the low coefficient model, the vacillation is stronger but irregular even near the end of this 100 years of run. The average period of the oscillation is about 24

years, which is similar to t h a t of the basic model.

5.2.

Grid size

As is mentioned in the earlier section, in the coarse grid model used for the spin up, vacillation does not occur. However, in t h a t coarse grid model, a larger horizontal diffusivity coefficient ( 2 . 0 l~o 7 ) is used along with larger horizontal viscosity coefficient (4.0x109), thus it should not be concluded t h a t this is a effect of grid size only. In a model with a half grid size model (80 km), vacillation is similar to that in the model with smaller horizontal diffusivity coefficient. However, again, the horizontal diffusivity is smaller ( 0 . 5 l~o 7 ) ,and therefore it is difficult to distinguish the effect of grid size.

5.3.

Wind driven circulation

In the basic model, the speed of extension of the cold and fresh water tongue is similar to that of the surface current. So, it is speculated that the period depend on the speed of the surface current. To examine this hypothesis, the model wind stress is enhanced by a factor 1.5. However, the result is completely different from expectation. After repeating two cycle in the first 70 years, vacillation suddenly stops a t the state when the cold and fresh water tongue develops half way t o its maximum state. In the first two cycle, the period is about 35 years.

6.

Discussion and Conclusion

In the previous sections, it is shown that vacillation can occur in models with constant forcing. The present models are different from the niodcls in the previous works mentioned in the introduction of the present paper in many details. The most essential difference is that no deep water is formed in models in the present work. Also, mixed boundary condition is used from the initial conditions. The details of the ’polar halocline catastrophe’ are also different from those in other models. In the present models, it starts at the north western corner of the model ocean and extends to south west, in contrast to the model in Bryan (1986a) in which it is the area near the eastern boundary which trigger the catastrophe. Also the oscillation is much more regular in the present models and the period found in the present experiments is not similar to any of others. However, one essential thing is common. T h a t is, in models using mixed boundary condition, overturning convection in higher latitudes is not stable. As far as the authors know, ‘polar halocline catastrophe’ is found in all the numerical models using mixed boundary condition. This indicate that, even though their appearance differ from model t o model, the essence of the phenomena is the same, and it is intrinsic to thermohaline circulation with mixed boundary condition. The question is, whether it is happening in the real ocean? So far, any evidence of this kind of phenomenon is not found. The suggestion by Weaver and Sarachik (1990) of the relation between decadal variability and oscillation of the same time scale occurring in their model recently attracts a lot of attention. However, we found the period depends on models and sensitivity to some parameters. It seems

283 too early to conclude that this relationship has substance. More investigations about the mechanism of this type of instability are needed to be able to say it, can happen in the real ocean.

References Bryan, F., 1986a. Maintenance and variability of the thermohaline circulation. Ph. D. thesis, Princeton University, Princeton, New Jersey, 254 pp. Bryan, F., 1986b. High-latitude salinity effects and interhemispheric thermohaline circulations. Nature, 323, 301-304. Bryan, K., 1969. A numerical method for the study of circulation of the world ocean. J. Com. Phys., 4, 347--376. Haney, R. L., 1971. Surface thermal boundary condition for ocean circulation models. J. Phys. Oceanogr., 1, 241-248. Mamaev, 0. I., 1964. A simplified relationship between density, temperature and salinity of sea water. Bull. Akad. Sci., USSR Geophys. Ser., 2, 180-181. Manabe, S., and R. J. Stouffer, 1988. Two stable equilibria of a coupled ocean-atmosphere model. J. Climate, 1, 841-866. Marotzke, J., P. Welander, and J. Willebrand, 1988. Inst,ability and multiple steady states in a meridional-plane model of the thermohaline circulation. Tellus, 40A, 162-172. Weaver, A. J., and E. S.Sarachik, 1990. The role of mixed boundary conditions in numerical models of the ocean’s climate. submitted to J. Phys. Oceanogr.

This Page Intentionally Left Blank

Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved

Modelling of Western Pacific Abyssal Circulation - Preliminary Experiment Nobuo SUGINOHARAT Shigeaki AOKIt and Minoru XAIiATA*

Abstract

As a first step t o model the western Pacific abyssal circulation, buoyancydriven circulation in a basin which has a simple geometry hut retains characteristic features of the western Pacific is studied using a multi-level numerical model. The circulation is forced by cooling inside the ocean at the southwest corner of the basin and uniform heating through the sea surface. It is clearly demonstrated that the bottom and deep waters take different paths from the southern ocean to the Philippine Basin. The bot,toni water flows northward along the western boiindary crossing the equator, and at the northern end of the Marshall Islands it turns westward t o flow to the opening of the Philippine Basin. On the other hand) the deep water at mid-depths flows directly into the Philippine Basin along the coasts. Numerical modelling studies which the present study is based upon are reviewed.

1.

Introduction

T h e ocean general circulation may be divided irit,o wind-driven a n d therniohaline circulations in terms of t h e forcing agencies: although superposition is not possible due t o nonlinearity of t h e system. In general, t h e wind-driven circulatiori is considered t o dominate in t h e surface layer above t h e main thermocline while t h e thermohaline circulation dominates in t h e deep layer. However, t h e interface between t h e circulations is not well underst.ood. In t h e present study, we will t r y t o understand t h e abyssal circulation in t h e Pacific as t,he thermohaline circulation. T h e progress on t h e development of a theory for the thermohaline circulation had been very slow until numerical modelling emerged into t,his study by t h e aid of development of high speed computers. Bryan (1987) demonstrated strong dependence of t h e thermohalirie circulation on t h e effect of vertical diffusivity in a general circulation model, i.e., with increasing diffusivity t h e thermocline deepens a n d t h e circulation becomes stronger. Colin de Verdiilre (1988) studied t h e thermohaline circulation within t h e context of planetary geostrophy, and demonstrated detailed dependence of t h e diffusivity although convincing explanation for its physics was not given. Suginohara a n d Fukasawa (1988) attempted t o interpret 'Center for Climate System Research, University of Tokyo. M r g u r o - k u , Tokyo 153, Japan National Research Institute for Pollution and Resources, 16-3, Onogawa, Tsukuba-shi, Ibaragi 305, J a p a n

286 and improve the far-reaching idea of the deep circulation of Stomniel and Aroiis (1960), and clearly demonstrated how the thermocline and associated deep circulation are formed. In the set-up processes, Kelvin and Rossby wave-type density currents play an essential role in establishing the thermal structure. Similarly dynamics is also understood by a simple inverted reduced gravity model of Kawase (1987) which assumes pre-existence of stratification. Also they showed that set-up in the equatorial region occurs in a different way from off the equator, i.e., higher vertical mode motions dominate forming a set of alternating zonal jets (stacked jets) along the equator. Suginohara and Aoki (1991) studied nature of the therniohalinr circulation, and gave a convincing explanation why the thermohaline circulation depends on vertical diffusivity. Further, they gave a possible explanation for the generation of the stacked jets along the equator which is an important ingredient of the thermohaline circulation. All of the above studies showed that most of the ocean basin tends t o be occupied by the coldest (heaviest) water formed at the sea surface. This feature is common t o the horizontal convection in a non-rotating tank (Rossby, 1965, and Beardsley and Festa, 1972), because the Kelvin wave-type density current assisted by Rossby wave-type density current plays an essential role in establishing the stratification as discussed in Suginohara and Fukasawa (1988). For the Pacific abyssal circulation, Stornmel and Arons (1960) proposed the circulation pattern. Since then many tried to simulate the observed distributions of tracers such as salinity, dissolved oxygen and 14C using the Stommel and Arons’ pattern of the deep circulation (Kuo and Veronis, 1973 and Fiadeiro and Craig, 1978). However, as Fiadeiro (1982) concluded, the Stonimel and Arons’ pattern needs to be confined to the deeper part of the deep ocean in order to reproduce the distributions of the observed tracers. i n fact, the existence of the benthic front (Craig et al., 1972) and also the 14C distribution along 180’ line of GEOSECS strongly indicate this feature. In addition, meridioiial velocity fields obtained by various inversion techniques (Wunsch et aZ., 1983) using the SCORPIO expedition d a t a (Stommel e t al., 1973) show that the northward flowing western boundary current is confined to the deeper part of the deep ocean. A possible cause of this may be due to effects of the existence of Drake Passage which blocks the deeper part of the abyssal circulation and leads t o introduction of only the bottom water as was demonstrated by Cox (1989). Suginohara, Aoki and Obata (1992) examined an abyssal circulation driven by cooling inside the ocean at the southwest corner of their basin. The reference density for cooling is taken to be vertically distributed, following a basic experiment for a non-rotating case by Aoki, Ishikawa and Suginohara (1992). They clearly demonstrated t h a t the Stommel and Arons’ pattern is confined t o the deeper part of the deep ocean, and even well below the thermocline the circulation reverses, which is just the abyssal circulation pattern proposed by Fiadeiro (1982). Nakata, Aoki and Suginohara (1992) studied effects of the continental slope along the western boundary on the circulation obtained by Suginohara, Aoki and Obata (1992). In the Philippine Basin, Fukasawa et al. (1986) made direct current measurements and found t h a t the deep flows are in the opposite direction to the surface Kuroshio at several stations south of Japan. Kubota and Ono (1992) reproduced

287 a circulation well compared with t h a t observed by Fukasawa rt al. (1986) in an inverted reduced gravity model. In their model the inflow into the Philippine Basin is prescribed at the Yap-Mariana Junction where the inflow is really taking place as shown by Uehara and Taira (1990) in the detailed CTD measurements. In this paper, following the modelling studies discussed above, we study how

1500km

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1000 2000 3000 4000 5000

0

1000 2000 3000 4000 5000

, 1 3 0

0 m 1000

2000 3000 4000 -5000-4000-3000-20001000

km

0

m 1000 2000 3000 4000 50004OOCF3000-2000-1000

km

Fig. 1. Meridional circulation, zonal flow and density fields along the meridional section at the center of the basin (a) and zonal circulation, meridional flow and density fields along the zonal sections a t the distance 2000 km south of the equator and the equator (b) for the case of the uniformly distributed reference density after Suginohara, Aoki and O b a t a (1992). For the zonal flow, the contour interval is 0.2 cm s-' and the shaded areas indicate the westward flow. For the meridional flow, the contour interval is 0.4 cm s-l and the shaded areas indicate the southward flow. For the density, the contour interval is 0.2 for ct less than 27.8 and 0.005 for ut greater than 27.9. Numerals on contours are ut - 20.0.

288 [h)

-2000km

Okm

... 1000.y:.-..*.. ....................................... ..................... .......

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Fig. 1. (continued)

the bottom and/or deep water is introduced into the Philippine Basin and how the deep water returns to the southern ocean using a multi-level numerical model. The model basin has a simple geometry but retains characteristic features of the western Pacific. This study is a first step to model the Pacific abyssal circulation, and we hope this will help to construct a more realistic model for the whole Pacific circulation. This paper is organized as follows. In section 2, several papers which the present study is based upon are reviewed as they are not published yet. We describe a model used in this study in section 3 and then discuss results in section 4. In section 5 , discussion and concluding remarks are given.

2.

Description of Results of Previous Works Which the Present Study is Based upon

Suginohara, Aoki and Obata (1992) studied an abyssal circulation driven by cold water formation inside the ocean a t the southwest corner of the basin and uniform heating through the sea surface. Their model ocean is a rectangular basin on a P-plane which extends 5000 km from the equator in the north and south directions, respectively and 3000 km in the east-west direction. For the cold water formation, they considered two cases; one is a case where the reference density is uniform in the vertical direction and the other is a case where the reference density is vertically distributed. Figs. l a and b show the velocity and density fields along the meridional

289 section at the center of the basin (a) and those along the zonal sections ( b ) for the case where the reference density is vertically uniform. Also plotted in Figs. 2a and b are those for the case where the reference density is vertically distributed. I t is clearly seen in Fig. 2 t h a t for the case where the reference density is vertically distributed the Stomniel and Arons' pattern for the deep circulation is confined t o the deepest part of the deep ocean, and consequently tlic reversed circulation fornis even well below the thermocline. Being associated with the predominance of higher vertical mode motions, the western boundary current has characteristics which are well compared with the observation (Warren, 1976, and Reid, 1986) as shown in the zonal section at 2000 km south of the equator, i.e., the lower northward flow doesn't accompany a n offshore countercurrent, while t h a t in Fig. l b docs. -41~0 the boundary currents accompany intensive vertical motions, and of course, those with higher vertical modes dominate for the case where the reference density is vertically I

.........

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Fig. 2. As in Fig. 1 but for the case of the reference density with vertical profile.

290 Okm

-2000km

0 m 1000

m 1000

0

2000

2000

Fig. 2.

(continued)

distributed. These features may be explained by the work of Masuda and Uehara (1992), where characteristics of the boundary layer are examined using a reduced gravity model with horizontal diffusion and viscosity. They classify the dynamics of the boundary layer into viscous and diffusive regimes. In both regimes, lateral diffusion dominates to determine the distribution of vertical velocity. T h e criterion dividing the two regimes is given by

where L R denotes the radius of deformation, K H coefficient, of horizontal diffusion and A H t h a t of horizontal friction. When the left-hand side is larger than the right-hand side, the boundary layer is in the viscous regime. For this regime, the boundary current accompanies the offshore countercurrent, and an intense vertical motion in a layer adjacent to the coast and an opposite vertical motion on the offshore side take place. The width is close t o that of the Munk layer, O ( ( A H / P ) + )In . contrast, when the left-hand side is smaller than the right-hand side, the boundary layer is in the diffusive regime. This is the layer which Warren (1976) obtained, i.e., the boundary current accompanies no offshore countercurrent and intensive vertical motions occur as in the viscous regime. For the case where the reference density is vertically distributed, the higher vertical mode motions dominate, i.e., the radius of deformation is smaller and hence as shown in Fig. Zb, the Warren type boundary layer forms.

29 1 A set of alternating zonal jets along the equator forms in both cases. For the case where the reference density is vertically distributed, number of jets increases. This may be so because predominance of the higher vertical mode motions in the western boundary in the southern hemisphere leads to occurrence of the higher mode structure of the stacked jets as discussed in Suginohara and Aoki (1991). Nakata, Aoki and Suginohara (1992) studied effects of a continental slope along 10"s

4000

0

- 4

4000

10

20

30

0

10

20

30

4000

Fig. 3. Zonal circulation, meridional flow and density fields along the zonal sections a t 10"s and the equator for the case of the flat bottom after Nakata, Aoki and Suginohara (1992). For the meridional flow, the contour interval is 1.0 cm s-l and the shaded areas indicate the southward flow. For the density, the contour interval is 0.1 for ct less than 27.9 and 0.01 for at greater than 27.9. Numerals on contours are at - 20.0.

292 the western boundary on the circulation obtained by Suginohara, Aoki and Obata (1992), using a rectangular ocean basin on a spherical earth which extends 30 degrees from the equator in the north and south directions, respectively and 30 degrees in the west-east direction. Figs. 3 and 4 show the velocity and density fields along the zonal sections for a case without the slope (Fig. 3 ) and with the slope (Fig. 4). It is clearly seen that effects of the slope is to enhance the upper northward flowing western boundary current (compare Fig. 3b with Fig. 4b). The enhanced upper northward flowing boundary current is accompanied by the offshore countercurrent, which leads to formation of an anticyclonic circulation a t mid-

4000

40001'

, , , , , I

Fig. 4. As in Fig. 3 but for the case of the continental slope along the western boundary.

293 depths over the slope. I t seems t h a t the density and velocity fields in the interior are not affected by the existence of the slope. Vertical motions associated with the boundary current retain those for the flat bottom case (see Figs. 3 and 4). T h e stacked jets along the equator are formed in both cases.

3.

Description of the Model

T h e ocean basin has a simple geometry but retains characteristic features of the western Pacific as shown in Fig. 5. The width of the basin is taken t o be very small compared with the Pacific, and the southern boundary is located at 20"s. On the western side of the basin, coasts are composed of, froin the sout,h, the Tonga Islands, Solomon Islands, New Guinea, Philippine Islands and Japan. Along the coasts, continental slope is considered. S o r t h of New Guinea and t>lieSolomon Islands, we consider a relatively shallow depth region. T h e Philippine Basin has the Izu-Ogasawara-Mariana Ridges its its eastern boundary and is connected t o the Pacific deep basin through the Yap-Mariana Junction. North of the Tonga Islands, the Marshall Islands are considered as a ridge which extends t o 15"N. T h e model used in this study is the multi-level numerical model described in Bryan (1969) and adopted t o modern vectorizing computers by Cox (1984). T h e continuous equations and a detailed description of the finite difference formulation may be found in the references given above. But, for the continuous equations only the equation of density under the assumption t h a t density and temperature have a linear relation is considered. For the finite difference formulation, revision is made on the vertical advection term in the equation of density, i.e., a weighted upcurrent

/

Fig. 5 .

Schematic view of the model ocean basin.

294 scheme is adopted so as not t o induce false increase or decrease in density at every grid point. This avoids the occurrence of unrealistic features around the equator as discussed in Suginohara, Aoki and Fukasawa (1991). The cold (heavy) water formation is made inside the ocean at the southwest corner in the southern hemisphere by adding a mass source term to the equation of density following the idea of Suginohara, Aoki and Obata (1992),

where C, is the specific heat of constant volume, p the sea water density and p c ( s ) the reference density which is vertically distributed as shown in Fig. 6. The coefficient, p(X, 4) is a function of longitude, X and latitude, 4 as shown in Fig. 7. It has a maximum value, l o p 2 cal("C cm2day)-'cm-l around the southwest corner of the basin. In addition to the cold water formation inside the ocean, density flux is imposed through the southern boundary,

0-

1000

2000 DEPTH(A1)

k--

3000

10

20

IT,

4000

a

LON

5000. 24

25

26

27

28

at Fig. 6. Vertical distribution of the reference density, p c ( 2).

Fig. 7. Horizontal distribution of the coefficient, p(X, 4).The contour interval is cal(OC cIn2day)-'cm-'. The maximum value lies around the southwest corner and is cal("C crn2day)-'cm-'.

295

Model layer 1 2 3

4 5 6 0

I

8 9 10 11 12 13 14 15 16 17 18 19 20

Table 1. Vertical grid size in meters 300 300 300 300 300 300 300 300 300 300 300 300 300 300 300 250 150 100 50 50

The coefficient, y,(X) has a maximum value, 3 ~ 1 cal("C 0 ~ cni2day)-l near the southwest corner, and the longitudinal dependence of it is the same as t h a t of p(A, 4).At the sea surface, density flux is imposed YS

-(Ps

Po

c,

-

P)

(4)

where p s is the reference density and is taken to be 24.0 in ut units, and ys is constant, 150 cal("C cm2day)-'. Both velocity components and the normal gradient of density are zero a t the lateral boundaries except the latter a t the southern boundary. Wind stresses are not taken into account, and at the bottom a quadratic drag law is considered. Horizontally, the resolution is constant throughout the grid at half a degree in the longitudinal and latit,udinal directions. Vertically, there are 20 levels as shown in Table 1. The mixing coefficients have been set to K H = l o 7 cm2sP1 for horizontal diffusion, AH = 8x107 cm2s-' for horizontal friction, K v = 1 cm2s-' for vertical diffusion and Av = 1 cm2s-l for vertical friction. As the initial condition a weakly stratified ocean is considered to suppress density currents with large amplitude as discussed in Suginohara and Fukasawa (1988). To accelerate a calculation and to reach a steady state efficiently, Bryan's accelerating method (Bryan, 1984) is adopted except in the initial transient stage. A calculation is carried out until a thermally and dynamically steady balance is est ablished.

296

Fig. 8. Horizontal velocity fields (a) and horizontal distribution of density (b). The contour interval for the density is 0.01 in C T ~units but 0.1 for the depth 450 m.

4. Results Calculation was made for more than 2000 years to entirely cover the vertical diffusion time scale. At the end of the calculation, the net density flux is almost zero and no trends can be seen in density and velocity fields. The horizontal velocity fields (a) and horizontal distribution of density (b) are shown in Figs. Sa and b. Also plotted in Figs. 9a and b are the velocity and density fields along the meridional sections (a) and those along the zonal sections (b). As is clearly seen in the meridional sections, the stratification is almost uniform in the whole ocean basin. The thermocline is formed around the depths, 300 m. The flow pattern a t the depth of 4950 m in Fig. Sa represents the circulation in the lowest part (see also Figs. 9a and b). In the cold water formation region, an anticyclonic circulation forms and the westward flow feeds the northward flowing western boundary current which crosses the equator and feeds the interior eastward flow with the poleward component in the northern hemisphere. When the boundary current reaches the northern end of the Marshall Islands, the current turns its direction toward the west and flows to the opening of the Philippine

297

Fig. 8. (continued) Basin and strong upwellings occurs there (see Fig. 9a). In the northern part of the northern hemisphere a cyclonic circulation which extends as south as 20"N forms. At the depth 4350 m where the Philippine Basin exists, the bottom water tends to flow into the basin. In the Philippine Basin, there is the westward flowing boundary current along the northern boundary, and the interior flow is eastward with the northward component. At this depth, the western boundary current flows in the same direction as below and the interior flow becomes weak. At depth, 3750 m, which is just below the bottom of the relatively shallow region north of New Guinea, the western boundary current becomes weak and that around the equator reverses its direction, and the interior flow reverses its direction except in the northern part. In the Philippine Basin, the flow pattrrn is the same as below being fed from the opening. At the depth 3150 m where the relatively shallow region north of New Guinea exists, the western boundary current reverses its direction all along the Tonga Islands and Marshall Islands. As seen in the zonal section a t 1"N, the upwelled water flows into this region along the equator and feeds the circulation there. The flow pattern at depth, 1950 m, represents the circulation a t mid-depths. The western boundary current off the Tonga Islands flows northward feeding the westward current along the New Guinea coast which feeds the flows in

298 7" 0-

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Fig. 9. Meridional circulation, zonal flow and density fields along the meridional sections a t 7", 11" and 35" (a) and zonal circulation, meridional flow and density fields along the zonal sections a t lo's, 1"N, 10"N, 20"N and 30"N. For the zonal flow, the contour interval is 0.2 cni s-' and the shaded areas indicate the westward flow. For the meridional flows, the contour interval is 1.0 cm s-' for the solid lines and 0.25 cm s-' for the dashed lines, and the shaded areas indicate the southward flow. For the density, the contour interval is 0.1 in ct units for the solid lines and 0.02 for the dashed lines. Numerals on contours are ut - 20.0.

35" ~~

205

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(continued)

the Philippine Sea. On the eastern side of the Marshall Islands, the western boundary current forms and flows southward being fed by the eastward flows from the western boundary via the Izu-Ogasawara-Mariana Ridges. It is noted t h a t there formed a circulation north of New Guinea and off the Tonga Islands. At depth, 450 m, the western boundary current reverses its direction again and the interior flow is westward with the equatorward component. In the northern part, there is an anticyclonic circulatiorl which extends as south as 20"N which is the central latitude in the northern hemisphere. Together with the cyclonic circulation discussed above, it indicates t h a t the interior upwelling is almost uniformly distributed and the flow is in geostrophic balance (Kawase, 1987).

300

T h e western boundary current a t 10"s in Fig. 9b has resemblance t o t h a t in Fig. 4b. There are two northward flowing currents. T h e upper one accompanies the countercurrent but the lower one does not; the upper one is in the viscous regime and the lower one in the diffusive regime of Masuda and Uehara (1992) as discussed in Nakata, Aoki and Suginohara (1992). As seen in other zonal sections, some boundary currents accompany countercurrents being associated with predominance of the gravest baroclinic mode motions. A set of alternating zonal jets (stacked jets) along the equator forms even off New Guinea. T h e jets on the eastern side of the basin have the characteristics similar t o those in Suginohara, Aoki and O b a t a (1992) (compare Fig. 2 with Fig. 9).

301 20"N

Fig. 9.

(continued)

T h e eastward jet a t the bottom levels accompanies marked upwelling a t t he equator and downwelling on both sides of it (see Fig. 9b). For the stacked jets on the western side, th e bottom level flow is westward, which is fed by upwelled water of th e westward flow on the eastern side. In general, it seems t hat the jets are disconnected at t he edge of the ridge as seen in the zonal section a t 1"N. It should be remarked th at the westward jet a t the bottom level on the western side feeds the circulation over t he relatively shallow region north of New Guinea (Fig. 8a). T h e density balance for the advection type representation of the density equation in the interior of the deep eastern basin and t he boundary current regions is the same as in Suginohara and Aoki (1991). T h a t is, for most of the interior, dom-

3 02 30"N

::I Fig. 9.

$

_....___._._.----__._.___..---.--..

(continued)

inant balance is between the vertical diffusion and vertical advection, and at the bottom levels the horizontal advection plays an important role. The importance of the horizontal advection at the bottom levels can be understood when comparison is made between the flow pattern and density distribution in Figs. 8a and b. Each of the terms in the density equation for the Philippine Basin is plotted in Fig. 10. As seen in the balance, a t all depths the horizontal diffusion plays an important role. This may be understood by characteristics of the boundary layer controlled by the effects of the lateral diffusion (Masuda and Uehara, 1992). Vorticity balance is the same as in Suginohara and Aoki (1991). In the interior, even for the Philippine Basin, the balance is linear, i.e., P u = f d w l d z . Along the

303 coasts and the equator, the lateral friction becomes important.

5.

Discussion and Conclusive Remarks

As the first step to model the western Pacific abyssal circulation, buoyancydriven circulation in a basin which has a simple geometry but retains characteristic features of the western Pacific has been studied using a multi-level numerical model. The circulation is forced by cooling inside the ocean and uniform heating through the sea surface. For the deep and bot,tom water formation inside the ocean, we

Fig. 10. Terms is the density equation along the meridional section a t 7" at the depths of 3150 m and 4150 m. HA and HD are for the horizontal advection and diffusion term; and VA and VD for the vertical advection and diffusion term. The unit of the originate is s - l c t .

3 04

Fig. 11. Dissolved oxygen distributions in the bottom layer deeper than 4000 m (a) after Mantyla and Reid (1983) and at mid-depths between 2500 and 2600 in (b) after Reid (1981). Units are nil l-'.

followed the idea of Suginohara, Aoki and Obata (1992) where the reference density is vertically distributed. In their model it is clearly demonstrated that the Stomme1 and Arons' deep circulation pattern for the Pacific is confined to the deeper part of the deep ocean. Only their flow pattern can explain the observed tracer distributions such as 14C and the benthic front as pointed out by Fiadeiro (1982). Major attention is paid on how the bottom and deep waters are introduced into the Philippine Basin from the southern ocean. It is clearly demonstrated that the bottom and deep waters take different paths. T h e bottom water flows northward along the western boundary crossing the equator, and at the northern end of the Marshall Islands it turns westward and flows to the opening of the Philippine Basin. On the other hand, the deep water at middepths flows directly into the Philippine Basin along the coast. Between the two boundary currents there is a southward flowing current. It should be remarked t h a t a circulation north of New Guinea and off the Tonga Islands forms at mid-depths. It seems t h a t the interior circulation on the eastern side of the basin has nothing to do with this circulation. Figs. l l a and b show the observed dissolved oxygen distributions in the bottom layer deeper than 4000 in (Mantyla and Reid, 1983) and at mid-depths, 2500 through 2600 m (Reid, 1981). The distributions indicate t h a t the bottom water is introduced into the Philippine Basin via the northern end of the Marshall Islands, while the deep water at mid-depths is introduced directly along the coasts of the Solomon Islands and S e w Guinea. The circulation obtained in the present model suggests that this occurs, thoiigh the Solomon Sea is not included. But Wyrtki (1961) showed no inflow from the Solomon Sea into the Philippine Basin at mid-depths. We carried out many case-studies for tracer distributions, changing parameters such as coefficients of vertical and horizontal diffusion. The results, as expected

305 from the obtained flow patterns in t.he present model, show that the calculated tracer distributions may be well compared with those in Figs. l l a and b. It may be said t h a t the waters below the thermocline in the Philippine Basin come from the southern ocean taking two paths. A set of alternating zonal jets along the equator forms even on the western side of the basin where the relatively shallow depth region exists. The jets on the deep eastern side may be compared with that obtained in Suginohara, Aoki and Obata (1992), but detailed structure is slightly different as the cold water formation is made closer to the equator in the present model. The stacked jets on the eastern and western sides of the basin are seemingly not connected. We also calculated a couple of cases changing the vertical profile of the reference density for the cold water formation. In either case, two paths to the Philippine Basin exist for the deep and bottom wat,ers. A n interesting result may be t h a t when both of the bottom and deep waters are very homogeneous, there is no stacked jets on the western side of basin. In conclusion we believe that the present study will serve as the first step t o further model the whole Pacific abyssal circulation.

Acknowledgment We would like to express our heartfelt thanks to Prof. Toshihiko Teramoto for his leadership on promoting studies of the deep circulation. Thanks are extended t o Dr. Masao Fukasawa, Prof. Yutaka Nagata and Dr. Kensuke Nakajima for pleasant discussions and Ms. Emi Tabata for typing the manuscript.

References Aoki, S., I. Ishikawa, and N. Suginohara, 1990. Convection induced by cooling at one side wall in two-dimensional noii-rotating fluid. J . Oceanogr. (submitted). Beardsley, R. C., and J . F. Festa, 1972. A numerical model of convection driven by a surface stress and non-uniform horizontal heating. J . Phys. Oceanogr., 2, 444-455. Bryan, F., 1987. Parameter sensitivity of primitive equation general circulation models. J. Phys. Oceanogr., 17, 970-985. Bryan, K., 1969. A numerical method for the study of ocean circulation. J. Comput. Phys., 4, 347-376. Bryan, K., 1984. Accelerating the convergence to equilibrium of ocean-climate models. J. Phys. Oceanogr., 14, 666-673. Colin de Verdikre, A,, 1988. Buoyancy driven planetary flows. J. Mar. Res., 46, 215-265. Cox, M. D., 1984. A Primitive Equation, 3-dimensional Model of the Ocean. GFDL Ocean Group Tech. Rep., No. 1, 143 pp. Cox, M. D., 1989. An idealized model of the world ocean. Part I: The global-scale water masses. J. Phys. Oceanogr., 19, 1730-1752. Craig, H, Y . Chung, and M. E. Fiadeiro, 1972. A benthic front in the South Pacific. Earth Planet. Sci. Lett., 16: 50--65. Fiadeiro, M. E., 1982. Three-dimensional modeling of tracers in the deep Pacific Ocean: 11. Radiocarbon and the circulation. J. Mar. Res., 40, 537-550. Fiadeiro, M. E., and H. Craig, 1978. Three-dimensional modeling of tracers in the deep Pacific Ocean: I. Salinity and oxygen. J . Mar. Res., 36, 323-355. Fukasawa, M., T. Teramoto, and K. Taira, 1986. *4byssal current along the northern periphery of Shikoku Basin. J . Oceanogr. Soc. Japan, 42, 459-472.

Kawase, M., 1987. Establishment of mass-driven abyssal circulation. J. Phys. Oceanogr., 17, 2294-2317. Kubota, M., and K. Ono, 1992. Abyssal circulation model of t,he Philippine Sea. Deep-sea Res., 39, 1439-1452. Kuo, H.-H., and G. Veronis, 1973. T h e use of oxygen as a test for a n abyssal circulation model. Deep-sea Res., 20, 871-888. Mantyla, A. W . , and J. L. Reid, 1983. Abyssal characteristics of world ocean waters. Deep-sea Res., 30, 805-833. Masuda, A , , and K. Uehara, 1992. -4 reduced-gravity model of t h e abyssal circulation with Newtonian cooling and horizontal diffusion. Deep-sea R.es., 39, 1453-1479. Nakata, M., S. Aoki, and 13. Suginohara, 1991. Effects of a continental slope alongh the western boundary on t h e abyssal circulation. J. Oceanogr., 48, 193-219 (submitted) Reid, J. L., 1981. O n the mid-depth circulation of t h e world ocean. In: Evolution of Physical Oceanography, B. A. Warren and C . Wunsch (ed.), MIT Press, Cambridge, 632 pp. Reid, J. L., 1986. O n t h e total geostrophic circulation of the South Pacific Ocean: Flow Pattern, tracers and transports. Progress in Oceanogr., 16, 1-61. Rossby, H. T., 1962. O n t h e thermal convection driven by non-uniform heating below: An experimental study. Deep-sea Res., 12, 9--16. Stommel, H., and A .B. Arons, 1960. On the abyssal circulation of the world ocean 11. An idealized model of the circulation pattern and amplitude in oceanic basins. Deep-sea Res., 6, 217-233. Stommel, H., E. D. Stroup, J. L. Reid, and B. A. Warren, 1973. Transpacific hydrographic sections a t Lats. 43’s and 28’s: t h e SCORPIO Expedition I. Preface. Deep-sea Res., 20, 1-7. Suginohara, N . , and S. Aoki, 1991. Buoyancy-driven circulation as horizontal convection on /?-plane. J. Mar. Res., 49, 295-320. Suginohara, N., S. Aoki, and M. Fukasawa, 1991. Comments on “On t h e importance of vertical resolution in certain oceanic general circulation models”. J. Phys. Oceanogr., 21, 1699-1701. Suginohara, N., S. Aoki, and A. Obata, 1992. Modelling of double structure of Pacific abyssal circulation. J. Mar. Res. (to be submitted). Suginohara, N., and M. Fukasawa, 1988. Set-up of deep circulation in multi-level numerical model. J. Oceanogr. Soc. Japan. 44, 315-336. Uehara, K., and K. Taira, 1990. Deep hydrographic structure along 12’N and 13’N in t h e Philippine Sea. J. Oceanogr. Soc. Japan, 46, 167-176. Warren, B. A , , 1976. Structure of deep western boundary currents. Deep-sea Res., 23, 129-142. Wunsch, C., D. Hu, and B. Grant, 1983. Mass, heat, salt and nutrient fluxes in the sout,h Pacific Ocean. J . Phys. Oceanogr., 13, 725 -753. Wyrtki, K., 1961. T h e flow of water into the deep sea basins of t h e western South Ocean. Aust. J. Mar. Freshw. Res., 1 2 , 1--16. ~

~

Deep Ocean Circulation, Physical and Chemicul Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved

3 07

Diagnostic Approaches on Deep Ocean Circulation Norihisa IMASATOY Hideki NAGASHIMA! Hidetaka TAKEOKA: and Shinzou FUJIO'

Abstract Three different approaches on the diagnostic study of deep ocean circulation are discussed. Firstly, the time scale of the circulation and the age of the water are studied by using a reservoir model. The results demonstrate that the multi-layered structure of the deep layer is essential to produce the middepth age maximum which is found in the North Pacific Ocean. Secondly, the diagnostic method to obtain reasonable flow field from limited numbers of direct current observations is proposed. The nietliod is applied to the deep layer in the Philippine Sea, and the obtained flow field appears very reasonable. Lastly, by using climatological data of density field, the current field in the deep Pacific Ocean is obtained by following the diagnostic method given by Sarmiento and Bryan (1982), and the exchanges of water masses among sub-basins in the deep Pacific are discussed by Lagrangian tracking method of water particles. Generally speaking, the results are consistent with those obtained by various investigators.

This chapter consists of three different approaches on the diagnostic study of deep ocean circulation. In the first section, the time scale of the circulation and the age of the water are discussed by using a reservoir model based on the work by H. Takeoka (1991). The diagnostic method t o obtain reasonable flow field from limited numbers of direct current observations is given in the second section based on the work by H. Nagashima (present study). In the last section, by using climatological data of density field, the current field in the deep Pacific Ocean is obtained by following the diagnostic method given by Sarmiento and Bryan (1982), and the exchanges of water masses among sub-basins in the deep Pacific are discussed by Lagrangian tracking method of water particles. The study in the last section is based on the work by N. Imasato and S. Fujio (present study).

* D e p t . of Geophysics, Faculty of Science, Kyoto University, Sakyo-ku, Kyoto 606, J a p a n . + I n s t i t u t e of Physical a n d Chemical Research, Wako-shi, Saitanra 351-01, J a p a n . ZDept. of Ocean Engineering, Faculty of Engineering, Ehime University, Matsuyama-shi, Ehime 790, J a p a n .

308 1. 1.1.

Time Scale of the Deep Circulation and Age Distribution of the Water - A Reservoir Model Introduction

Relating to the abyssal circulation of the world ocean, an age of the water (an elapsed time since the water left the sea surface) and its spatial distribution are often discussed (e.g., Tsunogai, 1981). However, as the water would be considerably modified before it reaches the present position due t o mixing processes, the definition of the age is not straightforward, and sometimes its meaning is not clear when it is defined conventionally. The several time scales such as water age, residence time, transit time and turn-over time are widely used in the coastal oceanography in order t o describe the behavior of materials injected into a coastal sea or bay which is assumed as a reservoir (Bolin and Rodhe, 1973; Zimmerman, 1976; Takeoka, 1984). In this section, we investigate the time scale of the deep water circulation and the distribution of the water age in the ocean by using analogous method. We consider a dissolved or suspended material which moves with water particles. Here, we consider two basic time scales which represent the characteristics of the reservoir; an average age of the material r, which is obtained by averaging an age of material (a time elapsed since the material entered the reservoir) in a reservoir and an average transit time rt which is obtained by averaging the age of material which is leaving the reservoir. The average age r, characterize the time scale necessary for an initially empty reservoir to be filled with the material of the present amount, or the time scale for the whole material in the reservoir to be carried away after the supply of the material is stopped (Takeoka, 1984). S o , ra is a measure of response time of the reservoir to a time-varying input. The average transit time rt is the same as the turn-over time which is a ratio of the total amount of the material in the reservoir to the flux of the material into or out of the reservoir. The value of r, relative t o t h a t of rt is an important factor to represent the nature of the reservoir. For example, if the inlet and the outlet are located on the

source

(a@

sink

@: :s

source

Fig. 1-1. If the sink of the material is located on the opposite side of the source as in the case (a), the average transit time rt for the lake is larger than the average age 7,. If the sink is located near the source as in the case ( b ) , the average transit time rt is smaller than t,he average age ra.

309 opposite side of a lake as shown in Fig. 1-la, ra would be smaller than rt. On the contrary, if the inlet and the outlet are located close to each other as shown in Fig. 1-lb, r, would be larger than ~t (Takeoka, 1984). Since rt is constant for the fixed values of the capacity of the lake and of the river discharge by its definition, r, would be considered as a essential parameter of the renewal rate of the lake ~ be water. So, the difference r, - rt or the normalized difference ( 7 , - T ~ ) / Twould a good measure of the renewal speed of the lake or the reservoir. As seen in the example shown in Fig. 1-1, the locations of the source and the sink of the material are essential factors in the renewal problem of the reservoir. For some material in the ocean, it may be supplied through sea surface and be lost by depositing on sea floor, then the situation would be similar to the case in Fig. 1-la. In the present study, however, we shall deal with the case that the source and sink of the material are located on sea surface. Namely, the situation is rather similar to the case in Fig. 1-lb.

1.2.

Model

For simplicity, we consider a two-dimensional ocean with the horizontal extent

L and the depth H as shown in Fig. 1-2. The ocean consists of the deep layer with thickness h = 3H/4 and the surface layer with thickness H - h = H/4. The material is supplied a t the uppermost few grids (20x20 grids are set in the ocean) in the surface layer as an initial condition, and the change of the distribution of material concentration is calculated for several prescribed flow fields. The formation of the deep water occurs in the limited areas in the northern

source = s i n k t & t & t & t & t l t &

Fig. 1-2. Configuration of the vertically two-dimensional modeled ocean with the depth of H and the horizontal extent of L (left panel). We consider two layer system, and the thickness of the deep layer is h. As shown in arrows at x = h, we set a narrow sinking region to the right and a wide upwelling region to the left. As to the vertical velocity field in the deep layer of the upwelling region, three prescribed profiles (Case 1 through Case 3) as shown in the right panel are used in the present calculations.

310 part of the North Atlantic and in the Weddell Sea, and the deep water is believed t o return gradually t o the surface layer of the world oceans (Stommel and Arons, 1960). So, the sinking is assumed to occur in the right-hand narrow region of the width of L/10, and the upwelling in the other wide region. In the case of the Pacific Ocean where no deep water formation occurs and the deep water is supplied from the southern boundary, the left side in the modeled ocean may be regarded as the north. The vertical velocity distribution at the interface is given as shown in Fig. 12, and is assumed t o be horizontally uniform in the upwelling region and in the sinking region, respectively. Uniform shrinking of the water column is assumed in the surface layer, but three different vertical velocity profiles in the deep layer, Case 1 through Case 3 as shown in the right figure of Fig. 1-2, are examined in order t o check the effect of multi-layered structure as discussed by Fiadeiro (1982). Then, the horizontal velocity field is obtained from the continuity equation. The velocity fields in Case 1 through Case 3 are shown in Figs. 1-3a through 3c, respectively. The governing equation of the material concentration C is dC dC dC -+u-+w-=K,,+Kz-@at ax dz

d2C

d2C

dx

(1.1)

where t is the time, x and z the horizontal and vertical coordinates, u and w the horizontal and vertical velocity components and K , and h7,the horizontal and vertical diffusivities. The equation may be written in normalized form as: dC dC dC -+u-+w-=--+--

dT

dX

dZ

1 d2C P,, 8x2

1 d2C

P,, dZ2

(1.2)

where t = LhT/lw(h) = ridT, x = L X , z = h Z , P,, = ( L 2 / h F Z ) / r t d P,, , = (h2/K,)/rtd,U = h u / l w ( h ) and W = Lw/lw(h), and 1 is the width of the upwelling region (= 9L/10) and w ( h ) the upwelling velocity at the interface ( z = h ) in the upwelling region. r t d is the turn-over time of the water of the deep layer. P,, and P,, are the ratios of the horizontal and vertical diffusion time scales to the turn-over time, respectively. The larger values of Pp, and P,, indicate that the advection is more dominant in the process of the material transport. The material is supplied in the uppermost few grids at t = 0. The values of r, and rt depend on the thickness of the layer where the material is initially supplied. However, the age distribution pattern in the ocean is not affected by the thickness (Takeoka, 1991), and only the case in which the material is supplied in the uppermost one grid level will be discussed here. The distribution of the material concentration and its change are numerically calculated by solving Eq. 1.2. As the boundary condition, the concentration a t the sea surface is assumed to be 0. This means that the material once leaving the ocean through 5ea surface never returns. It is assumed that no flux of the material across the other solid boundaries, and that gradient of the concentration is 0 on these boundaries.

1.3.

Results

Not only the average age ra and the average transit time rt for the overall deep layer but also the local average age for each grid point are calculated from the time

311 ....................--,, ........... .................. -

-

-

L

-

.,*,,,,,*.**<-.***

................ , , , , , , , , , . . . . . . . . . p v .. .. .. .. .. .. .. .. . . .......,” .................. _ _ _ _ _ _ _ _ _ - _ - - - - -+:?.

, . , , , , , , , r r r r r r r r r

.................... .................. ..................

.................... \ t t 7

\ . . , , . . . \ \ V V

.................. .................. .. .. .. .. .. .. .. .. .. .. .. .. ............ .................. .. .. .. .. .. .. .. .. .. .. .. .. .. .. .. .. .. .. ..................

..................

.................... (a) case 1

9

,

,,

.... (bl case 2

! sg L !!#fq; (cl case 3

Fig. 1-3. Velocity dist,ributions corresponding t o Case l(a), Case 2(b) and Case 3(c).

.. .. .. . .. . . . . . . . . . . . .

.. .. .. .. .. ..... . . . . .. .. .. .. . .. .. .. .. .. .. .. .. . .. .. \.

I: : : : 1 : : ( a ) case 1

:la!?:

: : : : :A: : :I

[ b ) case 2

[ c ) case 3

Fig. 1-4. Calculated distributions of the local average age for Case l ( a ) , Case 2(b) aiid Case 3(c). change of the concentration distribution by using the iiiethod given by Takeoka (1984). The distributions of the local average age for Case 1 through 3 are shown in Figs. 1-4a through 4c, respectively. Here, we take L = lo4 km, h = 3 kin and Ttd = 1,000 years (or w ( h ) = lo-’ crn/s). The results are 51lOWn for P,, = P,, = 5 or for K, = 6x106 cm2/s and K z = 0.6 cin2/s. The averagc ages r, aiid the average transit times rt for the overall deep layer for these three cases are givrn iii Table 1-1. Each value of r, in Table 1 is almost one order of magnitude larger than the value of rt (the same for three cases), and shows that the deep layer of tlie ocean has a similar characteristics t o a reservoir in Fig. 1-lb. T~ in Case 1 is sigiiificantly larger than those in Cases 2 and 3. This is attributable to the difference in tlie flow patterns in Fig. 1-3. The downwelling in the sinking region in Case 1 does not penetrate near to the bottom in coniparison with that in Case 2 or Case 3 , and the flow near the bottom is very weak. Thus, the water in the almost stagnant region at the left bottom corner of the ocean has the large local average age as seen in Fig. 1-4a, and T~ for all-over deep layer becomes large. It is noted t h a t the renewal of the material is accelerated in the cases of multi-layered structure. The considerable difference between a single layered case (Case 1) and a multi-

3 12 Table 1-1. Comparison between the average transit time rt and the average age r,. case rt

r,

1 0.22 1.55

2 0.22 1.40

3 0.22 1.34

layered case (Case 2 or Case 3) can be also seen in the distribution of the local average age shown in Fig. 1-4. In Case 1, the age increase monotonously downward and leftward, and the distribution does not show the maximum age in mid depths as seen in the actual age distribution in the Pacific Ocean (Tsunogai, 1981). Several calculations are conducted for Case 1 by varying the values of P,, and P,,, but the obtained nature of the age distribution is the same as in Fig. 1-la, and the maximum age never appears in mid depths. T h e depth of the maximum age is changeable with the values of P,, and P,,. The calculations for multi-layered ocean (Case 2 and Case 3 ) are also done for various values of P,, and P,,. It is shown that the maximum age appears in mid depths for the larger P,, and P,,, but that it appears at the bottom for smaller Pez and P,,, say about 1 or less. This means t h a t , in order t o create the mid-depth age maximum, the diffusion time scale needs to be larger than the advection time scale or t h a t advection is more effective than diffusion in the niaterial transport in the real ocean. By using a simple reservoir model, we demonstrated t h a t the multi-layered structure of the deep layer is essential to produce the mid-depth age maximum. As pointed out by Fiadeiro (1982). the multi-layered structure in the deep layer should be taken into account in order to understand the deep water circulation.

2.

Estimation of the Flow Field form Sparse Current Data Variational Method

-A

2.1.

Introduction

Because the current data based on direct current incaburements are very few, the conventional dyiiamical calculation is widely used to obtain current fields in the ocean. In order to solve the inherent reference-level problem in the dynaniical calculation, several methods have been proposed recently; e.g., the ,B spiral method by Stonimel and Schott (1977), the inverse method by Wunsch (1978) and the advanced inverse method by Killworth (1986). The robust diagnostic model (Sarmiento and Bryan, 1982), which we shall apply t o the Pacific Ocean in the next section, is also one of the attempts t o obtain a realistic current field in the ocean from the hydrographical data. However, with the recent progress of current measuring techniques, number of the observational d a t a on deep ocean currents are increasing rapidly, so it might be possible to get a reasonable current field by interpolating the observed d a t a in near future if we develop a suitable objective analysis method. Sasaki (1958) introduced

313 a variational method to wind field analysis, and Dickerson (1978) developed this method to apply the air pollution problems of San Francisco Bay Area. We shall develop this method to be applicable to current field in the deep ocean, by introducing two physical constraints in which the effects of bottom topography is taken into account. Then, the method will be applied to the deep layer of the Philippine Sea where several direct current measurements have bceri carried out by Japanese oceanographers (see Chapter 2 in this book). 2.2.

Variational m e t h o d a n d physical constraints

Take x and y are the horizontal coordinates, z the vertical coordinate, and u , u and w the P ,y- and a-components of the velocity. The variational functional E is expressed as

where is the Lagrangian multiplier. ah and a , are the Gauss precision inotluli and a: = ;gL2, where g : and cz are the error variances of defined by a: = ?pi2 the horizontal and vertical velocity components. ( ( L O , Z'O, W O ) is the observed flow field or the initially guessed fields based on the observed current data. T h e constraint function F ( x ,y, z ) is chosen from a physical requirement as discussed below, and should be zero if the estimated flow field is consistent with t h a t requirement. By introducing this term with a suitable X which minimize the value of E , the estimated flow field will be modified so as to fit the physical requirement. The Sverdrup constraant Firstly, we introduce the constraint from the Sverdriip relation: pu = f o

dw

.dz

Similar constraint is used by Provest and Salmon (1986) in their flow field analysis in the Labrador Sea. By integrating Eq. 2.2 from the bottom ( z = - h ) to z = - H , we have

where W H and WB are the vertical velocity component at the top of the deep layer under consideration and at the bottom, and u g and U B are the horizontal velocity components a t the bottom, respectively. For simplicity, here we assiime t h a t the horizontal velocity in the deep layer does not depend on z , and replace U B and U B by u and u , respectively. Then we have

3 14 h in the second term in Eq. 2.4 can be replaced by h - H if the top of the layer is horizontal. If we use h’ = h - H and if we use Eq. 2.4 as a constraint function, Eq. 2.1 becomes

’

where we drop in h’ for simplicity. Taking the variation for each variable and setting it t o be zero t o minimize E : u = uo

dh +(1/24). A. dX

dh = WO - (1/2@:). A . (/?h/fo - -) dY

and

R R E Q D E E P E R THAN 3000M

Fig. 2-1. Bottom topography of the Philippine Sea. T h e portion shallower than 3000 m is shaded.

315 In the present system, the best value of X is determined froln Eqs. 2.4, 2.6, 2.7, and 2.8 for each grid point: A=

+

24(-'U.o . d h / d z vo(Ph/fo - dh/dy) - w,) ( d h / d z ) 2 ( P h / f o - d h / d y ) Z y2

+

+

(2.9)

By substituting this X into Eqs. 2.6 through 2.8, the flow field adjusted for the Sverdrup constraint is given:

where (2.13)

2cn15 ---3 I30

I40

150L

Fig. 2-2. Available current velocity d a t a in the Philippine Sea. The scale of current speed is shown above the figure. The velocities a t stas. 1 through 14 are based on direct current measuren1ent.s. The velocity a t sta. 15 is deduced from the property distributions given by Kaneko, 1984) and that at sta. 16 is assumed from the estimation of the inflow through the Palau Trench (Taira, personal communication).

316 and (2.14)

Muss conservatron coastrarnt The flow field estimated by the above procedure does not iieed t o satisfy the mass conservation. In ordcr t o assure the mass conservatio~i,we may modify it by introducing the second constraint: d(hlL)/dZ

+d(hv)/dy +

1UH

=0

(2.15)

Taking the variations of Eq. 2.1 under this constraint, we have (2.16)

(2.17)

Fig. 2-3. Flow field given as a n initial guess, which is given by interpolation and/or extrapolation from the d a t a shown in Fig. 2-1, by using the scheme 2-22.

317 and the equation t o determine the best value of X is h . V2X

+ 2 ( d h / d ~ d. X / d +~ d h / d y . d X / 3 y ) - y 2 / h . X +

where

3

+

3

0

0

1

(2.19)

+

(2.20) = Zct.;/h. ( z ( h ~ o )- ( h v o ) ZUH~) 3h If we introduce a new variable 4 (fAh), Eq. 2.17 can be converted into the simple Helmholtz equation: 0

V2d - { ( l / h ) V %

+ (y"P)}q!J +

0

=0

(2.21)

Eqs. 2.16 through 2.18 give the flow field adjusted for the mass conservation. After this second constraint is applied, the satisfaction for the Sverdrup relation may not be the best, and so the final estimation might be obtained by repeating the above two procedures iteratively. However, it appears to give a good estiniate with no such iteration.

2.3.

Application to the Philippine Sea

In order to apply the method to the estimation of the flow field in the ocean, we need to choose reasonable values of the parameters in the above equations. The

Fig. 2-4. Flow pattern obtained by the variational method with the Sverdrup constraint, from the field in Fig. 2-3.

318 horizontal and vertical Gauss precision moduli ah and a , of the observed field would be determined from the observation. In the present study, we selected these values arbitrary as ah = 1 and a , = 10". w e use the small value of y = in order t o exaggerate the effect of the bottom topography. The depth of the top of the deep layer is chosen to be 3,000 m. T h e bottom topography deeper than 3,000 m of the Philippine Sea is shown in Fig. 2-1. The positions of the available direct current observations are shown with the mean current vectors in Fig. 2-2. We ignore the difference in depth of the observation if the observation was made below 3,000 m depth as we assuiiie that the current velocity is vertically uniform in the present analysis. For the grid point ( A , ] ) where no observation is available, the initial assumption of the current velocity [ ( u o )( ~V ~O ,) ~ is~ ]given by interpolation and/or extrapolation from the observed velocities by using the following scheme:

n ]is , the current velocity at the n-th observation point, and where [ ( Z L ~ (~u o~b s)) ~ ( T ~ is ) the ~ ~ distance between t h a t observation point and the grid ( i , j ) under consideration. The initial assumption of the current field obtained is shown in Fig. 2-3.

Fig. 2-5. Flow pattern obtained by the variational method with the mass conservation from the field in Fig. 2-4.

3 19 The estimated flow field obtained by the variational method with the Sverdrup constraint only is shown in Fig. 2-4. The estimated velocity tends to be affected by the bottom topography in the vicinity of the estimated point, and so the velocity field shows considerable spatial variability. Secondly, we applied the variational method with the mass conservation t o the current field given in Fig. 2-4. The result is shown in Fig. 2-5. The second constraint has the effect to smooth the flow field as seen in the figure. The strong currents seen a t several grid points may be erroneous as these come from the shallowness at those points. As we assume the vertically uniform flow, the mass conservation constraint may produce the locally exaggerated current velocity above shallows. Although the flow pattern in Fig. 2-5 is deduced from only 9 observed current data, it appears well to reproduce the counterclockwise circulation in the Philippine Sea. The northward flow along the western edge of the Izu-Ogasawara Ridge is supplied its water both from the inflow from the Palau-Yap Junction and from the northeast current at the southern margin of the circulation. This flow nature appears very reasonable too. We believe that the variational method proposed here would give a powerful method to get a reasonable flow field if the number of the direct current observations is increased in the future.

3. Diagnostic Calculation for Water-Mass Movement in the Deep Pacific 3.1.

Introduction

The deep circulation has been studied with a core-layer method from the distribution of water-masses characterized by potential temperature, salinity and dissolved oxygen. The core-layer method has successfully revealed the locations of formation regions and the spreading routes of deep water (Mantyla and Reid, 1983). However, it could show only a qualitative flow pattern, and it is not easy t o trace water-masses in the deep Pacific, because no deep water is formed in the deep Pacific and the tracer concentrations are rather uniform. We diagnose the deep circulation in the Pacific by using the robust diagnostic model of Sarmiento and Bryan (1982) with a small revision. The quantitative velocity field diagnosed by the model enables us t o investigate water-mass movements with the Euler-Lagrangian method. Since the deep water is thought to be supplied t o the Pacific by the Antarctic Circumpolar Current, our objective is to show how the Circumpolar water enters the deep Pacific and where it extends.

3.2.

Model and procedure of the calculation

The model uses the spherical coordinates of X (longitude), 4 (latitude) and t (height). The hydrostatic balance is assumed for the vertical motion, whereas the conservation equation of momentum, including inertial and viscous terms, is used for the horizontal motion. The eddy viscosities, A H and A v , are taken as 8x10' and 1 cm2sP1, respectively. The equation of continuity is also used. The equation

320 for potential temperature 8 or salinity S is given as Eq. 3.1,

dA dt

-+--

+-

1 d(uA) 1 d(vAcos4) a c o s 4 dX acos4

d2A +-d(wA) =A7~VZ24+K~~+y(A*-A) dz d,

(3.1) where the variable A denotes 6' or S . The eddy diffusivities, K H and KV, are taken as 1x107 and 1 em's-', respectively. The last term in Eq. 3.1 is called y term, which is introduced by Sarmiento and Bryan (1982) to prevent calculated values from deviating greatly from observed potential temperature 0" or salinity S * . If there is an observed density which significantly deviates from a local advectivediffusive balance, the density is smoothed by the model t o satisfy the balance to some extent. The degree of modification is represented by y. For a large y, the model is restricted by the d a t a and approaches a purely diagnostic model. We prefer a large y because we intend to diagnose a velocity field from hydrographic data, not to predict it. T h e boundary condition for momentum is a non-slip condition a t the lateral walls and a slippery condition a t the bottom. The sea surface is assumed t o be a rigid-rid boundary and the surface momentum flux is imposed. The boundary condition for potential temperature and salinity is a no-flux condition both at the lateral walls and at the bottom. The surface fluxes of heat and salt are given by Eq. 3.2,

a

K,-(8, dz

S ) = p(BS+- 8, ss*- S )

(3.2)

Parameter X corresponds to a time scale for potential temperature and salinity t o be restored to observed surface values 6'; and S;. We choose 20 days as this restoring time scale. Since the potential temperature and salinity are constrained internally by the observed values, the influence of the surface condition is small. Since our purpose is t o obtained the steady circulation, we use annual-means for all observed values. The observed potential temperature and salinity are taken from Levitus (1982), and the sea surface stress from Hellerinan and Rosenstein (1983). The model covers the whole Pacific Basin (Fig. 3-1). The open boundary conditions for the Antarctic Circumpolar Current are as follows: The total transport is specified as 130 Sv ( = 1 3 0 ~ 1 0~~1 ~n ~ s - according l) t o Whitworth et al. (1983). The normal velocity gradient and the tangential velocity are imposed t o be zero along the open boundaries (shown as dotted meridional lines along 125"E and 45"W in Fig. 3-1). The potential temperature and salinity are fixed to observed values within the buffer regions (shaded area in Fig. 3-1). The model domain is divided into water columns of 2"latitude x 2"longitude. Each column is also divided vertically into 15 boxes. The lowermost box of each column is permitted t o change its thickness individually, so that the model represents the bottom topography smoothly. Calculation starts from the rest state. The model reaches an almost steady state after only a few years. However, we continue integration for 5000 days (13.7 years) to make sure of steadiness of the model. The parameter y plays a very important role in the robust diagnostic model. In a purely diagnostic case, an extraordinary vertical circulation appears near the equator, while the tracer distribution (Mantyla and Reid, 1983) does not infer

321 such a strong current. As the horizontal motion is quasi-geostrophic, even the small error in density field produce erroneous flow field near the equator where the Coriolis parameter vanishes. To prevent this effect, we assunie that the parameter y is proportional to the Coriolis parameter, i.e., y = yo/sir1 41. After carrying out a series of calculations, the most suitable value of yo is decided as yo = 1 / 50 day-', which gives 1 / 100 day-' of y at latitude 30". The calculated distribution of the zonally averaged meridional transport is shown in Fig. 3-2. As seen in this figure, smooth transition of the flow field is produced across the equator, and the difference between calculated and observed densities is reasonably sinall for whole region. In order to follow the water mass movement, many labeled particles are tracked by using the obtained Eulerian velocity field. A particle is regarded as a water parcel which holds a specific characteristic of tracers; hence, a group of them is thought t o be a water-mass. This technique called as the Euler-Lagrangian method has succeeded in studies of the tidal exchange through a narrow strait (Imasato et

50"

30"

IO'N

10's

30'5

50'5

70'5 120'E

150'E

180'

150'W

IZO'W

90 'W

60'W

Fig. 3-1. Geometry and bottom topography of the modeled Pacific Ocean. The depth contour interval is 1000 m. The portion shallower than 3,000 m is shown by stippled area. The thick lines represent the coast lines in the model. The shaded regions in the Antarctic Circumpolar Current region are the buffer domain where the potential temperature and salinity are fixed to observed values, and the normal velocity gradient and the tangential velocity are set to be zero at the open boundary of these regions along 125"E and 45"W.

322 M E R l D l O N A L TRANSPORT DAYZONALLY INTEGRATED

5000

0 I00

200

300 4 00

-

E

5DO 500

1500

I I-

4 2500 n 3500

4500

5500 TO'S

50's

10'5

30's

10"

JO'N

50"

Fig. 3-2. Obtained distribution of the zonally averaged meridional transport. The contour interval is 4 Sv. al., 1980) and of the water exchange between the continental shelf and the Kuroshio

Current (Awaji et al., 1990). If a three-dimensional velocity vector U = ( u , u , w ) is given at any point, the position of a particle, X ( t ) = (ucosq5 x A(t), u x + ( t ) ,z ( t ) ) is obtained by solving the initial value problem:

d X-u

at

(3.3)

In the present study, particles are initially deployed at an interval of 1" in the horizontal direction and 250 m in the vertical direction. Hence, each particle occupies a volume of 2 . 5 ~ 1 0 ~cm3 ' (= 100 km x 100 km x 250 m). Since the particle is thought to keep roughly this volume during its movement, the volume of the water-mass can be estimated by counting the number of particles. Eq. 3.3 is integrated in time with the fourth-order Runge-Kutta scheme. A simple linear interpolation is performed for the three-dimensional velocity field to find U at any point. Whenever a particle approaches a boundary (e.g., within a fourth of a horizontal box size), the velocity component normal to the boundary is imposed to be zero. This treatment prevents the particle from sticking t o the boundary. Since particle motion is reversible, it is possible t o integrate Eq. 3.3 backward in time to obtain its position in the past. Thus, we can calculate where a particle comes from by tracking it backward, as well as where it goes by tracking it forward.

323

Fig. 3-3. Obtained horizontal velocity field at 4000 m. The portion shallower than 4,000 m are shown with hatched area.

3.3.

The diagnosed flow pattern

A deep overturning cell appears in the deep layer below 1,000 m depth in the South Pacific (Fig. 3-2), and its circulating water is 9.3 Sv. This feature agrees with direct current measurements by Warren and Voorhis (1970) and Reid and Lonsdale (1974) who detected a southward mid-depth flow in addition to a northward deep flow. It also agrees with the meridional cell structure of Wunsch et al. (1983) who estimated its transport t o be 12 Sv by the Inverse Method. The present model reveals that this cell does not extend t o the North Pacific. We also emphasize that the Equatorial Undercurrent, in addition to the South and North Equatorial currents, are well reproduced. Fig. 3-3 shows the horizontal velocity field at 4000 m. An anticyclonic circulation appears around the East Pacific Rise, which is contrary to the result of Reid (1986). Fig. 3-4 shows the distribution of the diagnosed velocity on the vertical section along 43"S, which agrees well with the result of Wunsch et al. (1983). Especially, both of them show deep flows in an anticyclonic sense over the East Pacific Rise. Lupton and Craig (1981) inferred a westward flow at mid-depth near 15"S, 115"W over the East Pacific Rise. These results support t h a t the flow corresponds to the anticyclonic circulation. A swift deep current east of New Zealand agrees with descriptions of deep flows by Warren (1973) and Warren and Voorhis (1970). The present model estimates its maximum velocity to be 5 cm s-l and its northward transport near 43"s t o be

3 24 NORTHWARD V E L O C I T Y DAY-5000 ALONG L A T I T U D E 43.0's 0

I00 200 300

400

-

-E

$88 1500

tI

n

w 2500 c1

3500

4500

C.I.

5500 120'E

ISO'E

180'

15O"W

I20'W

1 . 0 cm 90'Y

5

60'W

Fig. 3-4. Meridional velocity across the vertical section at 43"s. The contour interval is 1 cm s-' for the upper portion (shallower than 500 m) and 0.5 cm s-' for the lower port,ion (deeper than 500 m). The stippled area represents the southward flow. roughly 10 Sv. The diagnosed deep current splits into two parts at the Samoan Passage: One flows through the Samoan Passage and the other flows through the passage to the east of the Manihiki Plateau. However, Reid and Lonsdale (1974) stated that this deep current passes through the Samoan Passage and flows into the Central Pacific Basin. This disagreement is probably attributed to the fact that the Samoan Passage is not resolved well in the present coarse-grid model. The flows in the Northwest Pacific Basin is very weak (Fig. 3-3) as similar to the result of Roemmich and McCallister (1989). In the Northwest Pacific Basin, the deep water flows northward along the Izu-Ogasawara Ridge and the Kuril Islands as the deep western boundary current. After reaching Emperor Seamounts, the deep water flows southward along its western side. This anticyclonic flow pattern is consistent with that given by the Inverse Method (Roemniich and McCallister, 1989), but differs from that inferred by the core-layer method (Kenyon, 1983) which shows that the northward current flows along the western side of the Emperor Seamounts. The deep water flows from the Northwest Pacific Basin into the Philippine Basin through the gap a t 10"N of the Izu-Ogasawara Ridge (Moriyasu, 1974; Mantyla and Reid, 1983) is well produced in the present model. 3.4.

Water-mass movement in the deep Pacific

Fig. 3-5 shows the particle trajectories projected on the horizontal plane at 4000 m depth. The initial positions of the particles are initially shown by small

325 dots, and each particle is tracked for 5 years. Generally, the particles in the North Pacific move slowly, and the particles released in the Northwest and Northeast Pacific Basins tend to circulate inside the basins and scarcely flow away. On the other hand, the particles in the South Pacific move far from their initial positions. The water particles a t 4,000 m depth in the Southwest Pacific Basin (the domain enclosed thick full and dashed lines in Fig. 3-6) are tracked backward in order to know where these particles come from. The particles are tracked for 60 years, and those which originate from outside of the basin are picked. Their horizontal positions are shown in Fig. 3-6 (data points in the basin domain). D a t a points outside the basin domain indicate the positions of the particles 60 years before; as the particles come from various basins and various depths, the identification of the original basins which these particles come from is shown with four symbols in Fig. 3-6a, and the identification of the original depth ranges is shown in Fig. 3-6b in the similar way. Fig. 3-6 shows that four water-masses enter the deep layer of the South-west Pacific Basin. The first water-mass comes from the South Indian Basin through the gap south of New Zealand and this water is a mixture of the Circumpolar Deep Water and the Antarctic Bottom Water (this water-mass is referred to CDAB hereafter). The route of this water-mass has been widely accepted (Gordon, 1972). T h e PARTICLE TRAJECTORY ~ " " " " " "

I20.E

I 5 0 'E

180'

"

"

"

'

150'W

,

I20'X

90'W

60'W

Fig. 3-5. Particle trajectories projected on the horizontal plane at 4000 m depth. The particles are initially deployed on this plane, at the positions marked by small dots. Duration of tracking is 20 years. T h e contour of the 4,000 m isobath is shown in the figure.

326 second water-mass originates from the upper Antarctic circumpolar Current (this water-mass is referred to UACC hereafter) and comes from the deep layer of the Southeast Pacific Basin. A part of the upper water (500-1000 m) of the Antarctic Circumpolar Current is caught into the clockwise circulation in the Ross Sea and becomes dense to sinks off Antarctica (arrow in Fig. 3-6b). On the other hand, the Antarctic Circumpolar Water of 1500 m level flows away to the Drake Passage around the circulation in the Ross Sea, although the route is not demonstrated in any figure. The Ross Sea Shelf Water mixes with the Circumpolar Deep Water, INFLOW P A R T I C L E

YEAR-

-60

YEAR-

-60

30's

SO'S

70's I40'E

I 7O'E

160'W

I 3 0 'W

IOO'W

INFLOW P A R T I C L E

70'W

30'5

SO'S

T O ' S -I

140'E

I70'E

160'W

130'W

IOO'W

Fig. 3-6. Distribution of the particles at 4000 m depth in the Southwest Pacific Basin (enclosed by a solid line), which were located outside the basin 60 years before. The positions of the particles 60 years before are also shown: The origins of the particles is shown in (a): * from UACC, A from CDAB, x from MFITA and o from UNE. The depth of the origins are shown in (b): o from 0-1750 m layer, A from 1750-2750 m layer, x from 2750-3750 m layer and * from 3750-bottom layer.

TO'W

327 OUTFLOW

YEAR-

PARTICLE

50

IO'N

10's

30's

50's

70's IZO'E

150'E

180'

I50'W

I20.W

90'W

60'W

Fig. 3-7. Distribution of the particles a t 4000 m depth in the Southwest Pacific Basin (enclosed by a solid line), the locations of which will be outside the basin 50 years after. The positions of the particles 50 years after are also shown. The identification t o which depths go is shown: o from 0-1750 m layer, A from 1750-2750 m layer, x from 2750-3750 m layer and * from 3750-bottom layer. and this mixed water moves to the Southwest Pacific Basin through the Southeast Pacific Basin. Mantyla and Reid (1983), on the other hand, stated t h a t the deep water is stagnant in the Southeast Pacific Basin. The water flows westward along the northern side of the East Pacific Rise for about 40 years and reaches the southwestern corner of the Southwest Pacific Basin to meet the water-mass CDAB from the South Indian Basin. Both waters flow together northward as the deep western boundary current. The third water-mass (MFITA) originates from the mid-layer of the Fiji and Tasmania Basins and the forth water-mass (UNE) originates from the upper layer of the equatorial and northeast Pacific and comes through the western and the northeastern boundary of the South Pacific Basin. In order t o know how the particles flow out from the Southwest Pacific Basin, the water particles at 4,000 m depth in the basin are tracked forward for 50 years. The result is shown in Fig. 3-7 in the similar way in Fig. 3-6b. A few water particles upwell and flow out of the basin t o be captured by the Antarctic Circumpolar Current. Most of water particles flow into the Central Pacific Basin. It takes only 20 years for the fastest particles to move from the east of New Zealand t o the Central Pacific Basin. The particles east of the Tonga-Kermadec Ridge also ride on the deep western boundary current and flow away from the basin. A few particles from the Southwest Pacific Basin cross the equator in the western part of the Central Pacific Basin and most of them flow eastward. The distribution of particles seems t o be consistent with that of the tracers in the Pacific (e.g., Mantyla and Reid, 1983).

328 DAY-

B A S I N BOUNDARY

IZO'E

150'E

180'

150'W

120'W

90 'W

B A S I N BOUNDARY

IZO'E

150'E

SO'W

DAY-

180'

150'1y

I20'W

90 'W

0

SO'W

Fig. 3-8. Water exchange (in Sv) among the basins in the 1875-2875 m layer (short arrow) and in the 3875-bottom layer (long arrow), which are obtained from Net Eulerian transport (a), and Lagrangian transport evaluated by counting the numbers of water particles exchanged (b).

0

329 3.5.

Water exchanges among the basins in the Pacific

By referring the 4,000 m isopleth of the Pacific, we defined twelve basins as shown in Fig. 3-8, and estimated the water exchanges among these basins. We used two methods: One is t o calculate Eulerian net volume flux across the boundary between each two basins, and the other to count the net number of particles flowing across the boundary for a given time interval (100 years in this study). When we consider the specified layer of the boundary between basins A and B, the number of particles t h a t move into the layer of basin A from various layers of basin B and that move into that layer of basin B from various layers of basin A is calculated, and the difference of these two numbers is assumed t o be a net particle flow. Then the volume flux is obtained by multiplying the net number flux by 2 . 5 ~ 1 0km3. ~ The estimation of the latter method may be called as Lagrangian volume flux. T h e water exchanges among the basins estimated from the above two methods are shown in Figs. 3-8a and b for two layers from 1875 m to 2875 m with thick and short arrows and from 3875 m to the bottom with thin and long arrows. Numerals attached t o these arrows indicate volume flux in the unit of lo6 m3s-l. The bottom water of the South Pacific moves t o the North Pacific through basin C E (See Fig. 3-8 for the positions of the basins). T h e mid-depth layer (17502785 m) crosses the equator of the eastern part of the Pacific and moves t o basin PH (the Philippine Basin). Figure 3-8a of Eulerian transport shows that the bottom layer water moves from basins FI and TA to basin SE via SW, and t h a t this transport is attributed to the ACC. On the other hand, Lagrangian transport flows from basin SE to basins FI and from TA via SW. The Lagrangian transport from basin SI to S W and SE is large while the Eulerian transport is small. Moreover, in the mid-depth layer, the Eulerian and the Lagrangian transports from the basin S W t o SE is in opposite directions. The Antarctic Circumpolar Current flows from the basin S W to SE, but the water exchange occurs from SE to SW across the boundary as shown in Fig. 3-6. This result shows that the Eulerian treatment does not succeed in evaluating the water exchange between the two basins. Our result shows that the theoretical Euler-Lagrangian tracking method is most useful in evaluating the water-mass exchange and material transport between deepbasins and in examining the exchange process. However this method does include some kind of ambiguity.

References Awaji, T., K . Akitomo and N. Imasato, 1991. Numerical study of shelf water motion driven by the Kuroshio: Barotropic model. J. Phys. Oceanogr., 21(1), 11-27. Bolin, B., and Rodhe, 1973. A note on the concepts of age distribution and transit time in natural reservoirs. Tellus, 25, 58-63. Dickerson, M. H., 1978. MASCON-A mass consistent atmospheric flux model for regions with complex terrain. J. Appl. Meteor., 17, 241-253. Fiadeiro, M. E., 1982. Three-dimensional modeling of tracers in the deep Pacific Ocean. 11. Radiocarbon and the circulation. J. Marine Res., 40, 537-550. Gordon, A. L., 1972. Spreading of Antarctic Bottom Waters 11. In: Studies in Physical Oceanography, vol. 2 , A. L. Gordon (ed.), Gordon and Breach Science Publishers, pp. 1-17.

Hellerman, S., and M. Rosenstein, 1983. Normal monthly wind stress over the world ocean with error estimates. J. Phys. Oceanogr., 13, 1093-1104. Imasato, N., T . Awaji and H. Kunishi, 1980. Tidal exchange through Naruto, Akashi and Kitan Straits. J. Oceanogr. SOC.Japan, 36, 151-162. Kaneko, I., 1984. Structure of Mid-depth Waters in the Philippine Sea. Dr. thesis, Univ. of Tokyo, 97 pp. Kenyon, K. E., 1983. Sections along 35'N in the Pacific. Deep-sea Res., 30, 349-369. Killworth, P. D., 1986. A Bernoulli inverse method for determining the ocean circulation. J. Phys. Oceanogr., 16, 2031-2051. Levitus, S., 1982. Climatological Atlas of the World Ocean. NOAA Prof. Pap., No. 13, U S . Government Printing Office, Washington, D.C., 173 pp. Lupton, J. E., and H. Craig, 1981. A major helium-3 source at 15's on the East Pacific Rise. Science, 214, 13-18. Mantyla, A. W., and J. L. Reid, 1983. Abyssal characteristics of the World Ocean waters. Deep-sea Res., 30, 805-833. Its Physical Moriyasu, S., 1972. Deep waters in the western North Pacific. In: Kuroshio Aspects, H. Stommel and K. Yoshida (ed.), Univ. of Tokyo Press, pp. 387-408. Provost, C., and R. Salmon, 1986. A variational method for inverting hydrographic data. J. Marine Res., 44, 1-34. Reid, J. L., 1986. On the total geostrophic circulation of the South Pacific ocean: Flow patterns, tracers and transports. Prog. Oceanogr., 16, 1-61. Reid, J. L., and P. F. Lonsdale, 1974. On the flow of water through the Samoan Passage. J. Phys. Oceanogr., 4, 58-73. Roemmich, D., and T. McCallister, 1989. Large scale circulation of the North Pacific Ocean. Prog. Oceanogr, 22, 171-204. Sarmiento, J. L., and K. Bryan, 1982. An ocean transport model for the North Atlantic. J. Geophys. Res., 87, 394-408. Sasaki, Y., 1958. An objective analysis based on the variational method. J. Meteor. SOC. Japan, 36, 72-79. Stommel, H., and A. B. Arons, 1960. On the abyssal circulation of the world ocean. I. Stationary planetary flow patterns on a sphere. Deep-sea Res., 6, 140-154. Stommel, H., and F. Schott, 1977. The beta spiral and the determination of the absolute velocity field from hydrographic station data. Deep-sea Res., 24, 325-329. Takeoka, H., 1984. Fundamental concepts of exchange and transport time scales in a coastal sea. Continental Shelf Res., 3, 311-326. Takeoka, H., 1991. Effects of the deep ocean circulation on the reservoir characteristics of the ocean. J. Oceanogr. SOC. Japan, 47, 27-32. Tsunogai, S., 1981. A method for chronology of the Pacific and Atlantic deep water and its application. Geochemical J . , 15, 70-76. Warren, B. A , , and A. D. Voorhis, 1970. Velocity measurements in the deep western boundary current of the South Pacific, Nature, 228, 849-850. Warren, B. A,, 1973. Transpacific hydrographic sections at Lats. 43's and 28's: the SCORPIO expedition; I1 Deep Water. Deep-sea Res., 20, 9-38. Whitworth, T., 111, W. D. Nowlin, Jr., and S. J. Worley, 1983. The net transport of the Antarctic Circumpolar Current through Drake Passage. J. Phys. Oceanogr., 1 2 , 960-971. Wunsch, C., 1978. The North Atlantic general circulation west of 50'W determined by inverse methods. Rev. Geophys. Space. Phys., 16 (4), 583-620. Wunsch, C., D. Hu, and B. Grant, 1983. Mass, heat, salt and nutrients fluxes in the South Pacific Ocean. J. Phys. Oceanogr., 13, 725-753. ~

33 1 Zimmerman, J. T. F., 1976. Mixing and flushing of tidal embayments in the Western Dutch Wadden Sea. Part I, Distribution of salinity and calculation of mixing time scale. Netherland J. of Sea Res., 10, 149-191.

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved

333

Laboratory Experiments of Dense Water Descending on Continental Slope Yutaka NAGATAT Ryuji KIMURA; Hiroyuki HONJI? Yasuhiro YAMAZAKI: Kazuhiro KAWAGUCHI" and Tokuzo HOSOYAMADAS

Abstract The behavior of dense water descending on sloping bottom in a rotation system is investigated experimentally. The flow pattern can be classified into several regimes, two dimensional or three dimensional or steady flow or wavyeddy flow. Its parameter dependence is discussed. The descending speed of the dense water front is measured and compared with theoretical ones based on a simple steady-state model. The agreement is fairly good, and indicates that the flow in the bottom Ekman layer is essential for downward transport of the dense water.

1. Introduction Deep and bottom waters of the ocean are formed in shelf regions of two limited areas: the seas off Greenland in the North Atlantic Ocean and around Antarctica, especially in the Weddell Sea. The dense water generated by severe winter cooling through sea surface and by ejection of brine water due to ice formation flows down on continental slope (e.g., Carmack, 1973; Carmack and Foster, 1975; Foldvik et al., 1985a,b,c), and spreads along the ocean bottom into world oceans. Smith (1975) and Killworth (1973, 1977) investigated the behavior of the descending water on slope, and pointed out t h a t the entrainment effect of surrounding water is essential t o produce the necessary volume of the deep water and t h a t the frictional force plays an important role in the rotation system. Killworth (1977) also indicated that the pressure dependence of the thermal expansion coefficient is needed t o be considered in order that enough volume of the dense water can reach the deep bottom layer in the Weddell Sea. However, details of the descending and spreading mechanisms of dense water still seems be unclarified. The purpose of this chapter is to review the recent results of laboratory experiments conducted at both Kyushu University and University of Tokyo, and t o clarify the basic mechanisms of descending and spreading dense water. T h e results of the simple models in our rotating tanks may not be applicable directly to the *Dept. of Earth and Planetary Physics, Faculty of Science, University of Tokyo, Bunkyo-ku, Tokyo 113, J a p a n . 'Ocean Research Institute, University of Tokyo, Nakano-ku, Tokyo 164, J a p a n XKyushu University Graduate School of Engineering Sciences, Kasuga-si, Fukuoka 816, J a p a n § P o r t and Harbour Institute, Yokosuka-si, Kanagawa 239, J a p a n

334 formation mechanism of deep and bottom waters in the real ocean. We believe, however, that identification of several flow regimes, the dependence of the advancing velocity of dense water front on various experimental parameters and so on will help to produce better understanding for the formation and circulation of the deep and bottom waters.

2. 2.1.

Structure of Dense Water Descending on Slope Experimental procedure

The experimental apparatus used in this section is shown schematically in Figs. l a and l b . A Plexiglas tank of 7 5 x 3 0 15 ~ cm is placed obliquely on a turntable so t h a t one of the side walls of the tank constitutes a sloping bottom with declining angle of 34". Firstly, the tank is filled with fresh t a p water u p to the upper end of the sloping bottom (Fig. la). Dense water is formed by a cooling tube which is placed horizontally 6.5 cm apart from the upper end of the sloping bottom. This tube consists of a brass tube of 73 cm in length and 1.0 cm outside diameter, inside of which cooling water is forced to circulate so as t o keep the tube a t constant temperature. Cooling is started after the water in the tank attains solid-state rotation. Soon after the cooling water has begun to circulate, the dense water is formed around the tube and flows down on the sloping bottom. The angular velocity (f / 2 : f is the Coriolis parameter) of anticlockwise rotation of the table is varied from 0 t o 1.1 s-1. T h e electrolytic precipitation method (Honji, 1981) is employed for flow visualization: when a low D.C. voltage is applied, white dye is produced from the surface of the cooling tube, which is coated with solder. Streaks of the dye are observed from top by a camera as shown in Fig. la. In addition, temperature distribution in the descending water along a line 6.5 cm apart from the cooling tube is mea-

bJ

Water tank

3,

Camera Thermocouple

Thermocouple

Brass tube

Turn table

( a )

( b )

Fig. 1 Schematic view of the apparatus used in the experiments in Section 2 : (a) side view and (b) top view. Numerals in the figure give scales in cm.

335

Fig. 2 Examples of top view of dye streaks in the descending water for various parameters: (a) p* = 00 and T a = 0 (no rotation), (b) p* = 5.36 and Ta = 2 . 0 7 lo3, ~ (c) p* = 1.64 and Ta = 6 . 9 4 lo3, ~ (d) p * = 0.770 and T a = 1.23x104, and (e) p* = 0.502 and T a = 2 . 3 7 ~ 1 0 ~The . horizontal bar seen near the top margin in each picture is the cooling tube. sured every 60 s with an accuracy of 0.1"C by an array of 9 thermocouples which are placed 0.7 cm above the sloping bottom surface (see Figs. la and l b ) . The thickness of the descending water layer is quite thin (order of several mm). As the motion of the environmental water is so small, our attention is paid only to the

336

24

23 h

0 0 v

6

22

21

(b)

Fig. 3 Horizontal temperature distribution and its temporal variation in the descending dense water at the linez 6.5 cm apart. from the cooling tube: (a) for the case of p* = 5.02 and T a = 2 . 7 7 ~ 1 0and ~ (h) for the case p* = 1.41 and T a = 9 . 0 7 ~ 1 0 ~X . in the abscissa indicates the distance measured from left-hand side wall of the tank in cm. The profiles are given for elapse times of 1, 3 , 5 and 7 min., respectively. descending water. 2.2.

Results and discussion

The descending flow pattern attains almost a steady state after a critical nondimensional elapsed time: v t / D 2 > 7 where v is the kinematic viscosity coefficient,

337

k-

lo

0 .I

Two-dimensional

I

I

I

I 1 1 1 1

I

1

I

I

I

I

L

t the time measured from the start of dense water formation, and D the diameter of the cooling tube. In our experiment, for u t / D 2 > 14 the flow is affected by the dense water accumulating near the bottom corner of the tank; so the observation was made for 7 < v t / D 2 < 14. The flow pattern is determined by two nondimensional parameters: the non-dimensional density differences p* = g ’ / ( f D ) , and the Taylor number T a = f 2 D 4 / u 2 ,where g’ = g A p / p is the reduced gravity, g the gravitational acceleration, p the density of the environmental water and A p the density difference between the dense water just near tlie cooling tube and tlie environmental water. Examples of photographed flow patterns are shown in Fig. 2 for various values of these parameters. In the case of no rotation (Fig. 2a), the dense water flows down straight in the slope direction, and the dye streaks are almost parallel from top t o bottom. When the rotation is imposed, the streaks are deflected toward the rights looking downstream. The dense water flows almost horizontally at the middle of the sloping bottom and forms the front of the dense water. The water which flows along this front is accumulated a t the left end of the slope, and then descends along the left edge of the tank. For smaller values of Ta or larger values of p* (Fig. 2b), the streaks are almost parallel except near the front, and the flow may be regarded as a kind of two-dimensional plane flow though some wavy undulations of the streaks are seen in the left half of the figure. The temperature distribution along the line 6.5 cm apart from the cooling tube and its temporal variation are shown in Fig. 3a for the similar case of two-dimensional plane flow regime. The front of descending dense flow passed over this line by t = 3 min., and no significant change occurs after that.

338 Camera

Buffer

t Main Tank

Subtank

Slope

Rotator

Fig. 5 Schematic view of the experimental apparatus. The dyed salt water is pumped u p from a sub-tank and fed through flow regulator t o the injector set on the top surface of the truncated cone slope. All apparatus shown are fixed on rotating table.

As the value of T a is increased and t h a t of p* is decreased, a regular wavy pattern is generated in the descending water region as shown in Fig. 2c. The crest lines of the wavy pattern are perpendicular to basic flow direction. For larger values of T a and smaller values of p * , the wavy pattern is broken into several eddies as shown in Fig. 2d, but the shapes and their spacing are still regular. We may call this situation as a regular wave-eddy regime. For much larger Ta or much smaller p * , such regular eddies collapse into irregular turbulent flows as shown in Fig. 2e (irregular turbulent regime). An example of the temperature distribution and its temporal variations for regular wavy regime are shown in Fig. 3b. The wavy structure is clearly seen also in the temperature distributions after t = 3 min., and such a larger wave amplitude cannot be generated from the irregularities in cooling tube or in slope surface. The flow regimes obtained in 10 experimental runs are plotted in p*-Ta plane (Fig. 4). Though the coverage of the d a t a is very limited due t o technical reasons, this would give some idea what kind of flow regime is realized for given parameters. If we t.z,ke the thickness of formed dense water D as of order of 100 m, f as 1 . 4 ~ 1 0 -s-'~ (corresponds to 75"S), A p / p as lop3, and v as 10 cm2/s, p* and T a become 5x103 and 2x106, respectively. Our experimental range is too limited t o predict the structure of the descending dense water in the real ocean. But we feel t h a t three dimensional structures may be found in the real ocean.

3. 3.1.

Descending Velocity of the Dense Water Front Experimental procedure

A cylindrical tank of 80 cm in diameter and of 30 cm deep is set on a rotating table. An axisymmetric slope of circular truncated cone is put on the bottom as shown schematically in Fig. 5. The declination of the slope is fixed as 6/35

339

2(

15

lo(

50

0

1 10

20

30

40

so

t

(S)

Fig. 6 An example of the temporal change of the descending dense water front: f = 0.90 s-', A p / p = ~ . O X ~ O - The ~ . distance IC is measured downwards from the top of the slope, and the time t from the start of dense water ejection. Vertical bars attached to d a t a point indicate standard error. throughout this experiment: the height of the truncated cone is 6 cm, the diameter of the top flat surface is 26 cm and that of the bottom surface is 76 cm. Firstly, t a p water is filled so that the water depth a t the top surface is 16 cm. A Lucite disk lid is installed in order to spin up the air above the water surface together with water in the tank. After the water gets to solid-state rotation, dense salt water is fed from the margin of the top surface of the truncated cone at the fixed flow rate of 4.65 cm3/s. The salt water is dyed, and the advancing front of the dyed water is observed visually. From this the descending velocity is determined from the time needed t o move over a fixed distance. The dyed water is also photographed from top by a camera mounted on the frame set on the rotating table. The later d a t a is mainly used for the case of smaller values of Coriolis parameter in which the front is a p t t o be wavy or not symmetrical. Though the entrainment of the surrounding water into the descending dense water looks to be very small visually, the density

340

Fig. 7 An example of photographed, dyed descending dense water: f = 0, A p l p = 1 . 0 ~ 1 0 -and ~ t = 86 s.

Fig. 8 An example of photographed, dyed descending dense water: f = 1.59 s-', A p / p = 2 . 0 ~ 1 and 0 ~ t~= 126 s. of the water near the front appears t o decrease slightly a4 it moves downward. Also due to geometrical elongation of the frontal line, the advancing speed may decrease as the front moves downwards. However, as shown in an example in Fig. 6, the speed is almost constant from the top to the bottom of the slope. As a result it is sufficient to discuss only the averaged descending speed here. In order to prevent thermal convection, the fresh water and salt water have been kept in temperaturecontrolled laboratory for more than 24 hours before each experimental run. The density differences between the feeded salt water and the environmental

341

Fig. 9 An example of photographed, dyed descending dense water: ~ . O X ~ Oand - ~ t = 124 s.

f = 1.3 s-', A p / p =

fresh water A p / p is varied from 0.001 to 0.020, and the Coriolis parameter f is varied from 0.00 to 2.67 s-'. The thickness of the dense water layer just near the feeder is about 0.5 cm, so the minimum value of p* and the maximum value of Ta defined in Section 2 are 0.26 and 0 . 4 7 l~o 4 , respectively. This flow would be either in the two-dimensional flow regime or in the regular wavy-eddy flow regime. The examples of the photographed descending dense water and its front are shown in Figs. 7 through 9. In the cases of no-rotation or very weak rotation, we needed to use a well adjusted feeder and a carefully manufactured slope in order to keep the front axisymmetric. Even so, some irregular frontal shape is generated as shown in Fig. 7. We read the position of the front in 16 directions and the average position is used for determination of the front speed. In the case of rotation, descending water may be divided into two regions: a very thin frontal region and a relatively thick region backwards. For larger rotating rates, a regular array of eddies as shown in Fig. 9 is generated in the backward region. wliile no structure is usually seen for smaller rotation rates even in the backward region as shown in Fig. 8. The boundary of these two regions is very hard t o identify either visually or photographically. Also, identification between the two-dimensional flow regime and the regular wavy-eddy regime is difficult because the backward region may be formed only inside of the feeder for some cases and also because the formed eddies are somewhat time-dependent and may not be well developed as our experimental time span is too short. The descending speed, however, does not appear t o depend on whether the eddies are formed in the backward regions or not. So we shall omit discussing about the eddies in the backward region in the discussions t o follow. Some of the obtained descending speeds of the front uf are plotted in Fig. 10 against Coriolis parameter f for various values of density difference Ap. It is clear t h a t the descending speed decreases with increase of f and with decrease of Ap.

342 0.50

I

I

I

I

I

I

dd= 1 0 dd= 3 + dd= 5 0 dd=lO X dd=20 A

0.45

0.40

x

-

A

-

0.35 E

0

0.30 d

E

a, a, (4

[I)

0.25

-

X X

u

X

X

x

_ ,

x

X

-

XX X X

0.20

4-

0.15

-

0

+

0

+

0

+

+

+

-

0

+

+

i

0.10

-

0.05

0.00

-

I

I

0.0

0.5

I

I

I

I

1.0 1.5 2.0 2.5 Coriolis parameter [l/s]

3.0

Fig. 10 Plots showing the dependence of the descending speed of the front on Coriolis parameter f and on density difference (dd). Different symbols are used for different dd’s as shown in top-right corner of the figure (in g/cm3). 3.2.

Governing non-dimensional parameters

As it is difficult to find the governing non-dimensional parameters only from our limited experiment, we shall discuss a simple two-layer model as shown in Fig. 11. We consider two dimensional steady flow behind the front assuming that its downstream component is a measure of the descending front speed. Taking the x-axis down-slope and the y-axis horizontally on the slope and z-axis upward orthogonal to the slope surface and denoting 1-and y-velocity components with u

343 and v, we have linearized steady-state equation of motion for V = u d 2 V / d t 2 - i ( f / v ) V + g’sin O/v = 0 d2V/dz2 - i ( f / v ) V = 0

for h 2 z 2 0 for z > h

+ iu: (1)

where h is the thickness of the lower dense water layer, g‘ the reduced gravity, 0 the declination of the slope. The boundary conditions imposed on the bottom surface and on the interface are

v=o V and d V / d z are continuous

at z = O at z = h

(2)

for h 2 z 2 O forz>h

(3)

For the case of no rotation (f = 0), we have u = (g’sinB/v)z(2h - z ) / 2 , v = O u = (g’sin O / v ) h 2 / 2 , v=O

The constant downward velocity in the upper layer may be unrealistic as such a steady state condition would never be achieved in the tank. Here, we ignore the flow in the upper layer assuming that only the flow just over the interface satisfies the above boundary conditions. Then, we have the downward transport q of the dense water in the lower layer per unit width as q = g‘h3 sin O / ( 3 v )

(4)

The thickness of the lower layer h is usually too small t o be measured in our experiment. For the case of no rotation, however, q can be estimated from flow rate of the feeded dense water divided by length of the front: if we consider the circle at the mid point of the slope, the length of the front is 27rx25.5 cm, and so q is 0.0290 cm2/s as the flow rate is fixed as 4.65 cm3/s throughout our experiment. We shall denote this estimated transport with qo, and use it as a reference transport. Then the thickness of the lower layer ho and the mean descending speed uo for f = 0 is given as follow: ho = 3vqo/(g’sin 0) (5) uo = g’hi

sin 0/(3v)

(6)

Hereafter, ho and uo will be used for reference values. For the case of rotation, the downward velocity component within t.he bottom Ekman layer would be essential. We may choose the square of the ratio holb, where b = (2v)3/f is the thickness of the Ekman layer, as a measure of the effect of rotation: f* = = f/fo (7)

All of our experimental results are plotted by using non-dimensional downward speed of the front u i = U ~ / U Oand non-dimensional Coriolis parameter f * = f / j o in Fig. 12. The d a t a points for f = 0 is nearly 1, indicating that the theoretical

344

i Fig. 11 A simplified two-layer flow model used in theoretical considerations. A steady state is assumed and the thickness of the lower dense water layer h is assumed t o be constant. 1.2

I

I

I

I

dd= 1 0 dd= 3 + dd= 5 0 dd=lO X dd=20 A max u max Q - -

1.0 d B

a

vl

B

a

0.8

A B

b

6

A

0.6

4

m

0

.d B

C

0.4

.A E

2

- _-- - -- - - _ _

0.2

0.0 0.0

0.5 1.0 1.5 Non-Dimensional Coriolis Parameter

2.0

2.5

Fig. 1 2 Plots showing the dependence of non-dimensional downward speed u f / u o of the dense water front on non-dimensional Coriolis parameter f*. The solid and dashed lines indicate the theoretical guesses obtained from the simple model in Section 3.3. guess in the above discussion gives a good estimate of the descending speed of the dense water front for f = 0. It should be noted that all of the experimental data

345 are well aligned on a single curve within our experimental range. This suggests that the mean downward velocity in the lower layer in the simple model may give a good measure of the descending velocity of the dense water front for wide range off*.

Theoretical guesses on descending speed of the dense water front

3.3.

The solutions of Equation (1) under the boundary conditions (2) for the case of rotation are

where co = g’sinB/f

and

k = (l+i)(f/2v)i = (l+i)/b The volume transport within the lower layer Q is given by

Q

=Soh V d z = -ico[h + (1/2K){-3 - exp(-2kh) + 4exp(-kh)}] I I

5

10

Fig. 13 Plots showing the dependence of the non-linear downward transport q* on the non-dimensional thickness of the lower layer h/b. The conditions of the maximum q* and the maximum non-dimensional mean downward velocity are indicated with “MAX Q” and “MAX U” in the figure.

346 If we use non-dimensional volume transport Q* = Q/(cob),Q* is a function of lch or h/b only. The non-dimensional downward transport q* (the real part of &*) of the dense water is shown in Fig. 13 as a function of h/b. q* has the maximum value ( M 0.794) at h/b = T (“MAX Q” in Fig. 13). This means that only a part of the fed dense water can flow down on slope if we feed dense water at the rate higher than this critical value, and t h a t the remaining volume of the fed dense water will be trapped in upper part of the slope near the feeder. Thus, the thickness of the lower layer cannot be determined by the flow rate of the fed dense water as for the case of no-rotation, and we need to set some criterion t o determine the layer thickness h for the case of rotation. It may be reasonable to assume t h a t the flow occurs so as to achieve the maximum value of downward transport under the given condition. If the given q* is smaller than the critical value “MAX Q”, we can easily find the corresponding h/b value from the curve in Fig. 13. (Plural solutions for h/b can be obtained for q* values a little less than the critical value. Here, we shall select the smallest h/b for simplicity.) If the given q* is larger than the critical value, we assume t h a t the realized transport remains at just the critical value. By using the thickness of the lower layer determined by this procedure, we can calculate the non-dimensional mean downward velocity u* in the lower layer as the function of the non-dimensional Coriolis parameter f * defined in the previous subsection. The obtained function is shown by the dashed curve in Fig. 12. The downward and horizontal velocity profiles for the critical point “MAX Q” are shown in Fig. 14a. The maximum of q* occurs when the lower layer coincides with the layer of positive downward velocity of the Ekman spiral. This condition, however, does not give the maximum mean downward velocity in the lower layer. The maximum mean velocity occurs at the point “MAX U” in Fig. 13 ( h / bM -1.776 and q* M 0.609). The velocity profiles for this condition are given in Fig. 14b. Another reasonable criterion to determine the layer thickness h of the dense water is that the flow occurs so as to achieve the maximum value of mean downward velocity under the given condition. This gives another guess on the dependence of the non-dimensional mean downward velocity in the lower layer on the nondimensional Coriolis parameter. The result is given by the solid curve in Fig. 12. The agreement between the experimental results and the theoretical guesses is fairly good, especially for the latter case of “MAX U” assumption. 3.4.

Discussion

The normalization factor u, in (6) of the descending speed of dense water used in Fig. 12 is derived from the consideration of the non-rotational case, but our experimental results suggest that this normalization is also effective for larger nondimensional Coriolis parameter f* (namely, for larger f and smaller Ap). Th‘is may result from the fact t h a t the boundary condition at the interface (Equation 2), which would be realistic for the rotational case but not be for the non-rotational case, gives a good estimate of the descending speed of the front even for the nonrotational case. Then, the theory using a simple model predicts the smooth tran-

347

-1.0

1.o

0

MAX

Q

-1.0

0

1.0

MAX U

Fig. 14 Vertical profiles of the downward and horizontal velocity components for “MAX Q” (a) and for “MAX U” (b). T h e vertical coordinate is normalized by the thickness of the lower layer and each velocity component by co = g’f-l sin8.

Fig. 15 The extent of the fed dense water, when a barrier in the radiative direction is set t o prevent geostrophic horizontal flow: f = 1.57 s-l, and t = 43.8 s. Aplp = sition from non-rotational case to rotational case as seen in the theoretical curve in Fig. 12. (If we confine our attention only to rotational case, co would be a more reasonable reference velocity.) The agreement between the experimental results and the theoretical guess is especially good for f’ < 1. However, the experimental values are generally a little smaller than the theoretical guess. This may suggest the existence of some discontinuity or sharp transition between non-rotational and rotational cases. In

348 the later stage of the experiments, we made our best, efforts for the determination of the descending speed for smaller f * values by u s i q a carefully designed feeder and a accurate specially-manufactured smooth bottom slope. However, we cannot find any critical value of f * which separates rotational and non-rotational regime. For f * > 1, the experimental values of the speed are significantly larger than the upper theoretical curve (“MAX U” curve). The difference tends t o increase with increase of f *: the experimental speed looks to reach a fixed value asymptotically, while the theoretical curve continuously decrease with increasing f*. This discrepancy appears t o start near the bending point of “MAX Q” curve. For f* values larger than this bending point, only a part of fed dense water flows down the slope and the remaining part will be accumulated near the feeder. The trapped water would circulate clockwise in the upper part of the slope and stay there under the state of geostrophic balance. The outer boundary of the geostrophic domain, however: would move downwards as the volume of the trapped water increases with time. This movement may affect and increase the descending speed of the dense

Fig. 16 Temporal variation of the front line: the experimental conditions are same as for Fig. 15. Numerals attached to contours indicate the time t in s.

349 water front under consideration. Though it is hard to observe the position of the outer boundary of the geostrophic domain and its movement in our experiment, the discrepancy between experimental and theoretical speeds for f * >> 1 might be explained by this mechanism.

4.

Concluding Remark

Two series of rotating tank experiments concerning the behavior of the descending dense water on sloping bottom are reviewed. The experimental ranges in the present study are too limited and some of the important parameters such as the layer thickness of the descending dense water are not successfully measured, so the results may not be applicable directly to the real ocean. Besides, the downward transport in the real ocean is not necessary to occur within the bottom Ekman layer as discussed here as bottom topography in the real ocean has usually complicated three-dimensional configuration. In the experiment using a flat slope in Section 2, the dense water easily flows down along the lefthand edge of the slope, where the dense water is accumulated by the horizontal flow. The behavior of the dense water, when a barrier in the radiative direction is set t o prevent geostrophic horizontal flow in the upper part of the slope, are shown in Figs. 15 and 16. Foster and Carmack (1976) showed that the dense water on the continental slope of the Weddell Sea extends offshorewards along the east coast of the Antarctic Peninsula. The topographic effect by which the generated dense water is accumulated in some spot would be much more important. However, the results obtained in this study such as an identification of the flow regime and its parameter dependence would give an idea of the behavior of the dense water in the real ocean. Also, the parameter dependence of the descending speed of the dense water front obtained experimentally which can be explained by using a simple two-layer flow model would give a useful suggestion on the understanding of the effect of friction on the dense water behavior.

Acknowledgement The authors wish to express thanks for Prof. T. Teramoto, Kanagawa University for his suggestions and encouragements. The work was supported by the Priority Area Programme “Dynamics of the Deep Ocean Circulation” defrayed by the Ministry of Education, Science and Culture.

References Carmack E. C., 1973. Silicate and potential temperature in the deep and bottom water of the western Weddell Sea. Deep-sea Res., 20, 937-932. Carmack E. C., and T. D. Foster, 1975. On the flow of water out of the Weddell Sea. Deep-sea Res., 22, 711-724. Foldvik, A., T. Gammelsred and T. Tmresen, 1985a. Hydrographic observations from the Weddell Sea during the Norwegian Antarctic Research Expedition 1978/79. Polar Res., 3 n.s., 177-193. Foldvik, A . , T. Gammelsrod and T. Torresen, 198513. Physical oceanography studies in the Weddell Sea during the Norwegian Antarctic Research Expedition 1978/79. Polar Res., 3 n.s., 195-207.

350 Foldvik, A , , T . Gammelsrcbd and T. T@rresen, 1985c. Oceanographic conditions on the Weddell Sea during the German Antarctic Research Expedition 1979/80. Polar Res., 3 n s . , 209-226. Foster, T. D., E. C. Carmack, 1976. Frontal zone mixing and Antarctic bottom water formation in the southern Weddell Sea. Deep-sea Res., 23, 301-317. Honji, H., 1981. Streaked streaming around an oscillating circular cylinder. J. Fluid Mech., 107, 509-520. Killworth, P. D., 1973. A two dimensional model for the forniation of Antarctic Bottom Water. Deep-sea Res., 20, 941-971. Killworth, P. D., 1977. Mixing on the Weddell Sea continental slope. Deep-sea Res., 24, 427-448. Smith, P. C., 1973. A streamtube model for bottom boundary currents in the ocean. Deep-sea Res., 22, 853-873.

Chapter 6 Development of Acoustic Technology for Ocean Measurement

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Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

353

Target Parameter Estimation by Linear-Period Modulated Signal Wataru MITSUHASHI and Hitoshi MOCHIZUKI'

Abstract A basic technique for acoustic measurement on moving targets is described. Some interested features of a linear-period modulated ( L P M ) signal are introduced and a new method using the LPM signal for simultaneous estimation of range and velocity of the target is proposed.

1.

Introduction

Acoustic measurement of target range and velocity along the lines of sight of transducers is fundamental t o the investigation of flow structures in underwater environment. An integration of motion information, such as the range of a moving target and its velocity vector observed with an array of sound transducers, will allow the flow structures t o be reconstructed three-dimensionally. In this article, we will introduce a basic technique t o estimate the range and velocity of moving targets along the line of sight of a single transducer. In general, echoes reflected from submerged sound-scattering particles are difficult to be detected because of their lower energy as compared with ambient underwater noises. To gain a large SNR (the ratio of signal energy to noise power) on the receiver output, the emitted signal energy must be widely dispersed both in time and in frequency: this can be realized by a code modulation with pseudo random sequences or by a certain type of frequency modulation. Inverse operations of these signal manipulating processes enable us to compress the echo waveform within a limited time interval and t o form a higher peak amplitude than the noise level. However, sound scattering particles move with ocean currents, and accordingly the Doppler effect on the observed echo waveform must be considered carefully; it cannot be ignored in acoustic measurements because of a relative slow speed of sound propagation. The Doppler effect is theoretically formulated as temporal expansion or contraction of the echo waveform (Kelly and Wishner, 1965). Hence, the frequency characteristics of echoes reflected from moving particles will be different from that of the emitted signal. For this reason, the pulse compressing process does not effectively work if the signal is designed unsuitable for motion measurement. Thus, for precise estimation of the range and velocity of moving particles, it is highly important to design a signal to be robust against the Doppler effect - this type of signal property is referred to the Doppler tolerance. 'University of Electro-Communications, Chofugaoka, Chofu, Tokyo 182, J A P A N

354

Ar

b)

Ar

Fig. 1. Ambiguity diagram of (a) LPM signal and (b) LFM signal in the case of upward chirping. Except for the direction of frequency sweep, both signals are designed identical in the form of envelope, frequency bandwidth and time interval in which these signals exist. A linear frequency modulated (LFM) signal, known as the Chirp signal, has been developed for increasing the average power capability in peak power limited RADAR systems (Cook, 1960). As a direct consequence of the capability, the LFM technique allows higher target range resolution t o be realized than can be realized with conventional pulse RADAR systems (Klauder et al., 1960). It is well known however t h a t the LFM signal has a n ambiguous property related t o error coupling in simultaneous estimation of range and velocity. On the other hand, a linear period modulated (LPM) signal, the period of this signal increases linearly with time, has been developed as one of the Doppler tolerant signals (Kroszczyriski, 1969). T h e ambiguity function of the LPM signal has a longer spread ridge along the velocity error axis with approximately flat amplitude. T h e use of the LPM signal will consequently enable us to realize unbiased estimation of target range, even though the target moves with a high velocity. However, the velocity estimate is still ambiguous with a single L P M signal. Although Altes has proposed a double pulse structure of the LPM signal t o resolve the velocity ambiguity (Altes and Skinner, 1977), we will adopt a single L P M signal because of its simplicity in the design of signal form. Fig. 1 illustrates ambiguity functions of two kinds of signals in connection with the Doppler tolerance.

2. 2.1.

Background Echo reflected from a moving target

Emitting a signal u ( t ) as a location sound toward a point target moving with constant velocity u,, we can observe a returning echo reflected from the target:

e ( t ) = 6. u ( s , . [t - 7 , ] )

(1)

355 where ro is the propagation delay time between the sound emitter and the moving target, and so denotes temporal expansion or contraction due t o the Doppler effect is necessary in order (Kelly and Wishner, 1965). In the above expression, to account for the fact t h a t the signal energy does not change under the Doppler transformation. However, we can disregard the term J.0 because of its slight influence on parameter estimation. In Eq. ( l ) ,we have assumed that the signal component a t the time origin, u(O), reflects on the target just at r0/2,and at the time, the distance from the emitter t o the target is exactly rO1. In this situation, the following relationships hold:

Jso

2r0 ro = -

and

C

c - vo so = c v,

+

where c is the velocity of sound propagation. Thus the measurement of the range and the velocity of a moving target is reduced to the problem t o estimate two parameters which correspond t o a propagation delay (ro)and a time scale factor (SO).

2.2.

A signal form suitable for target parameter estimation

Equation (1) shows t h a t the echo waveform will expand or contract along the time axis with the scale factor so. As one of the invariant transformation for time scale distortion of this sort, we can utilize the Mellin transform which is defined for a function u ( t ) by

Converting the variable t into an exponential function e x p x , we can rewrite the above equation in a more convenient form:

l, +m

uM

(O) =

u ( e z )exp(-jRz)dx

(4)

That is, the Mellin transform is equivalent to the Fourier transform following after the logarithmic transform of the abscissa. Expansion or contraction of the time axis is replaced with a time shift by the logarithmic transform, and finally a phase rotation is obtained as a result of the Fourier transform. To maximize the SNR on the Mellin transform U,(C2), the signal u ( e Z )after the logarithmic transform of the variable t should be expressed by a sinusoidal function as G(.)

= q x ) . ejbz,

x = logt, t

>0

(5)

Rewriting the signal G(x) with the variable t , we derive a linear-period modulated (LPM) signal: u ( t ) = a ( t ) .e j . b ' o g ( ' ) , t > O (6) 'This assumption does not necessarily mean t h a t t h e signal u ( t ) must have a n energy a t t h e time origin.

356

h

- -- -- -w1

Instantaneous

w = b/t i \frequency

Fig. 2. A linear-period modulated signal and its related parameters. T h e center frequency and the frequency bandwidth are determined by the position of the envelope a ( t ) on the time axis. To localize the signal energy both in the R and in the x domain, we have expressed G(x) by a Gaussian function. Hence the signal envelope a ( t ) can be expressed with a log-normal function which is similar t o those expressions published so far (Altes and Skinner, 1977; Zakharia e t al., 1987). As will be shown lately, the parameter b determines the measurable range of target velocity. A simple example of the LPM signal and related parameters are shown in Fig. 2.

3. 3.1.

Linear-Period Modulated Signal Invariant property of instantaneous frequency of LPM signal

Since the envelope a ( t ) is a slowly varying time function in contrast with the phase blog(t), the temporal differentiation of the signal phase in Eq. (6) will yield the instantaneous frequency w : w=

dblog(t) - -b at t

(7)

We should take notice to this equation t h a t the instantaneous frequency w decrease as time passes; this corresponds t o the property of mammalian echo-locating sounds. I t has been pointed out t h a t the LPM signal has a typical feature of the location sounds produced by bats and cetaceans (Altes and Titlebaum, 1970). If the Doppler effect on the envelope a ( t ) is assumed t o be negligible, substitution of Eq. (6) into Eq. (1) leads t o e(t) = 6 . a ( s , . [t - To]). ej”og(so.[t-~ol)

ejb’og

SO

. u ( t - 7,)

(8)

357 The temporal expansion or contraction caused by the Doppler effect is converted to a phase rotation blogs, which originates in the signal design based on the Mellin transform. If the velocity of target motion is much slower than the sound propagation speed, then the following approximation holds: Iogs, = -2u,/c

(9)

Since the instantaneous frequency of Eq. (8) is given by

the frequency characteristic of the echo e ( t ) is equivalent to that of u ( t ) shown in Eq. ( 7 ) except for the delay time 7,. As a result, correlation between the emitted signal and the returning echo will have a maximum amplitude just at the time r,, even though the target moves with a high velocity. 3.2.

Coherent correlation detection of LPM signal

The LPM signal u ( t ) shown in Eq. (6) can be written in the form:

u ( t ) = a ( t ) c o s ( b l o g t ) + j ~ a ( t ) s i n ( b l o g t=) u 7 . ( t ) + j . u i ( t )

(11)

Since a real signal is emitted actually in SONAR problems, we must rewrite Eq. (8) in the following equation by selecting only the real part:

e ( t ) =cos(blogs,).u,(t-r,)

-sin(blogs,).u;(t-r,)

(12)

This equation shows that two parameters t o be estimated (the propagation delay T~ and the time scale factor s o ) are contained separately in the echo signal in terms of time shift and amplitude modulation. A schematic diagram of a coherent correlation detection system for target parameter estimation is illustrated in Fig. 3; the system output can be written in the form: e ( t ) . u*(t - r)dt = P ( r )+ j Q ( r ) (13)

/

where * denotes complex conjugation. The right hand side of Eq. (13) can be expressed in terms of correlation functions:

p ( 7 ) = cos(b1og so) . &?(r - 7 , ) - sin(b1og s,) . 4i7.(r- r,) Q ( r ) = - cos(b1og s o ) . b ; ( r- r,) sin(b1og s o ) . 4 i i ( r - 7 , )

+

(14) (15)

where, for example, d P T ( ~denotes ) the auto-correlation function of u r ( t ) , and also denotes the cross-correlation function between u r ( t )and u ; ( t ) . These correlation functions have the following relations:

358 Choose largest

Echo

Correlation detection

Q(T)

by 4 t )

tan-' Q I P

Fig. 3. A schematic configuration of coherent correlation detector. Range and velocity of a target are measured at the time where the correlator has the largest amplitude. where fc is the mean frequency of u ( t ) ,and A(T) is the envelope common t o all those correlation functions. Substitution of Eq. (16) and (17) into Eq. (14) and (15), and use of Eq. (9), yields the expressions:

P ( 7 ) = A(' - 7 , ) . C O S ( ~ T ~ , [ T- rO]- 2bv,/c)

(18)

Q(T) = A(T - T

(19)

~ ) sin(27rfc[r . - T,]

- 2bu,/c)

Since the phase measurement is practicable in general within the limits of f ~we, define the maximum measurable velocity as follows (Altes and Skinner, 1977):

V,,,

= crr/2b

(20)

Thus the LPM signal parameter b is given by

As a result, we can finally obtain the outFut representation of the correlation detector: J P 2 ( ~ ) Q2(7) = A ( r - T,)

+

3.3.

Velocity ambiguity

Eq. (22) shows t h a t by detecting the maximum peak of the output envelope, we can obtain an estimate of the propagation delay time, 6, without any a priori knowledge on the velocity uo. Hence the target range can be calculated by the relation T , = c 6 / 2 accordingly. However, Eq. (23) indicates t h a t the delay error AT yields directly the velocity error Au = -ZV,,, f c . AT. T h e property of the coherent correlation det,ector with a single LPM signal is summarized in the followings:

359 The error in velocity estimation depends heavily on that in delay estimation. On the contrary, the error in delay estimation is fully independent t o that in velocity estimation. Consequently, elimination of the error term, -2VmaxfcAr in Eq. ( 2 3 ) )will allow the ambiguous property in velocity estimation t o be cleared. However, this term cannot be eliminated as long as we use a single LPM signal, because both the cross correlation functions $ r i ( ~ ) and # i r ( r )in Eq. (17) do not become zero. These correlation functions will become zero and then the error term can be eliminated if a LPM signal is designed to be symmetrical with respect. t o the reversal of the time axis (Altes and Skinner, 1977). However we must use a single LPM signal for the convenience of signal design. In the following section we will propose a simple method to avoid the velocity ambiguity.

4. Simultaneous Estimation of Range and Velocity In Eq. ( 2 3 ) )Vmaxis an inherent parameter of the LPM signal, both v, and r, are the unknown to be estimated and 7~ is a constant, so that we can control only the variable f c . If the bandwidth of the detection system could be divided into an upper and a lower frequency bands, the following equations should hold in each of the frequency bands:

where f U and f i are the center frequency of the upper and the lower band respectively. Thus we can obtain the delay and velocity estimates by solving the above equations simultaneously.

4.1.

A bank of constant-Q filters

It is difficult in general t o divide a correlation detector into different frequency parts. Thus we have adopted a filter bank method in simultaneous estimation of range and velocity. It is widely known that frequency analysis is coarsely performed in the peripheral level of the auditory system and the logarithmic frequency axis is developed along the basilar membrane in the cochlea. If hair cells, which detect acoustical vibration, are distributed uniformly along the basilar membrane, the frequency analyzing function of the auditory mechanism can be modeled by a bank of constant-Q filters. For the LPM signal, a constant-Q filters can easily be implemented by the method shown in Fig. 4. The relation of the center frequency between the adjacent filters is given by e-2?r/b wk+l = wk '

3 60 3

3

u,.(t) = a ( t ) . cos(6. log t ) Fig. 4. A design of center frequency distribution for a bank of constant-Q filters in accordance with the phase characteristic of the emitted LPM signal

output

Echo e(t)

Fig. 5 .

Bandpass filters

Delay lines

“Delay and Sum” process with a bank of constant-Q filters.

where the index k denotes the filter number. Fig. 4 also shows t h a t the filters are arranged a t equal intervals along the logarithmic frequency axis. In order to realize an equivalent function to the coherent correlation detection, we must remove the non-linearity in spectral phase components of the received echo. This process can be realized by delaying each filter output with different temporal amount according t o the center frequency of the filter and by summing all of the outputs. The amount of time delay for the filter with center frequency wk is defined by d k = do - b/wk (27)

361 where do must be introduced t o satisfy the physical realizability of the filter system. Fig. 5 illustrates a schematic diagram of a bank of conatant-Q filters followed by time-delay devices. A quadratic implementation of the filters provides a n equivalent t o a coherent correlation detection system. We can obtain the equations (24) and ( 2 5 ) by dividing the bank equally into two parts along the frequency axis.

Fig. 6 Doppler tolerant property. Energy centroid of each echo waveform shifts with time according t o the target motion. T h e outputs processed by the proposed system are sharply compressed in time at the position which exactly corresponds t o the propagation delay, even though the target moves with a high velocity.

362 5.

Numerical Experiments

To verify the Doppler tolerant property of the LPM signal, we carried out numerical simulations on the assumption that targets move with high velocities. Fig. 6 illustrates simulated echoes compared with delayed and summed outputs. We can find clearly in this figure that the energy centroids of echo waveforms move along the time axis according t o temporal expansion (Fig. 6a) or temporal

.. I

.

I

time

,

.

,

,

,

I-.

I .

+

,

, 1

,

.

, I

,

,

~-L

,

,

,

,

I.

Fig. 7. Robustness against environmental noises. Three targets A, B and C are arranged stationary at equal distances along the line of sight with those intensities of 0.75, 1.00, and 0.50, respectively. The proposed system composed of a bank of constant-Q filters can detect clearly these three targets even from the echo buried in strong noise.

363

Simulated target velocity(mls) Fig. 8. Results of velocity estimation. Velocity estimates are plotted. = l m / s and SNR=34dB. Since the measurable extent of velocity is an estimated value beyond the limit is converted limited within ztV,,,, t o the value modulo V,,,. A proper design of ‘Q’ value of filters will cancel the systematic deviation shown in this figure.

V,,,

contraction (Fig. 6b,c) caused by the target motion. It can also be seen t h a t the outputs are sharply compressed at the time which exactly corresponds t o the propagation delay, even though the target moves with relative high velocity. T h e Doppler tolerant property of the LPM signal is consequently well confirmed. Fig. 7 shows that the filter bank work effectively even under a serious noisy environment. Finally we have estimated the range and velocity of a moving target according to the simultaneous solution of equations (24) and (25). Results of velocity estimation are shown in Fig. 8. Although biased estimates are observed in this figure, we can eliminate the biased amount by designing properly the ‘Q’ value of filters.

6.

Conclusion

We have developed a method for simultaneous estimation of the range and velocity of moving targets. The method emits a single LPM signal and processes the observed echo through a bank of constant-Q filters. Numerical experiments show that the bank of constant-Q filters can work effectively as an equivalent of a correlation detector. Doppler tolerant property of the LPM signal is confirmed. By dividing the bank of filters equally into two portions and solving their phase equations simultaneously, we can estimate the range and velocity of moving targets. Validity of this approach is also confirmed by the numerical simulation.

References Altes, R. A,, and D. P. Skinner, 1977. Sonar-velocity resolution with a Linear-Period Modulated pulses. J. Acoust. SOC.Am., 61, 4, 1019-1030. Altes, R. A,, and E. L. Titlebaum, 1970. Bat Signal as Optimally Doppler Tolerant Waveforms. J. Acoust. SOC.Am., 48, 4,1014-1020. Cook, C. E., 1960. Pulse Compression - Key t o More Efficient Radar Transmission. Proc. IRE, 48, 3, 310-316. Kelly, E. J., and D. P. Wishner, 1965. Matched Filter Theory for High-Velocity, Accelerating Targets. IEEE Trans. Military Electronics, MIL-9, 1, 56-69. Klauder, J. R., A. C. Price, S. Darlington, and W .J. Albersheim, 1960. T h e Theory and Design of Chirp Radars. BSTJ, 39, 4, 745-808. Kroszczynski, J. J., 1969. Pulse Compression by Means of Linear-Period Modulation. Proc. IEEE, 57, 1260-1266. Zakharia, M., P. Arzelies, M. E. Bouhier, J. P. Corgiatti, and B. FouchC, 1987. Improvement of Underwater Acoustic Localization by Coherent Processing. In: Progress in Underwater Acoustics, H. M. Merklinger (ed.), Plenum, New York, pp. 717-725.

Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

365

A Low Frequency Underwater Sound Source and Control Technique of Transducer Directivity Hiroyuki HACHIYA* a n d Motoyoshi OKUJIMAi

Abstract

In large scale monitoring techniques of the ocean such as ocean acoustic tomography, a wide band low frequency sound source and directivity control technique under high hydro-pressure in deep seas are required. In this paper, a new type of construction of a low frequency sound source and a new directivity control technique are proposed. In new type of a low frequency sound source, the cylindrical transducer is put between two metal circular plates. This source has wide band resonance and high reliability at high power operation. In new directivity control technique, a high-sound-speed material called syntactic foam is attached to a cylindrical transducer element to realize a transducer with desired directivity under high hydro-pressure. The source characteristics of resonance frequency and quality factor Q are examined using a theoretical model and experimental results are compared with theoretical results. Numerically designed results for the omnidirectional transducer using finite element analysis considering an infinite space agree well with measured directivity patterns. This low frequency sound source and directivity control technique are found to be valid in deep seas.

1.

Introduction

I n large-scale monitoring techniques of t h e ocean such as ocean acoustic tomography, a n underwater sound source which can transmit wide band low frequency sound with low transmission loss a n d highly accurate travel time measurement is required. T h e ability t o operate under high hydro-pressure is necessary when t h e source is moored in deep water of several thousand meters. A low frequency sound source having a resonant t u b e with a n effective length X/4 was proposed a n d used in a n acoustic tomography experiment (Worcester et al. 1985). B u t this sound source has high Q resonance, so four tube resonators were used t o broaden t h e bandwidth in t h e experiment. In addition, t h e reliability of adhesion between piezoelectric ceramics a n d metal of t h e flexural disc transducer used in this source is generally low during high power operation. In this paper, a new type of construction of a low frequency sound source in which a cylindrical transducer is p u t between two metal circular plates is presented. 'Precision and Intelligence Laboratory, Tokyo Institute of Technology, Nagatsuta, Midori-ku, Yokohama 227, Japan Faculty of Engineering, Toin University of Yokohama, Kurogane-cho, Midoriku, Yokohama 227, Japan

366 The space between the two plates operates as a resonator in the shape of a reel. This source has wide band resonance and high reliability when operated at peak power. The characteristics of this source are examined using theoretical analysis and results of measurement are discussed. On the other hand, a cylindrical transducer is frequently used in apparatus employed in deep seas such as an acoustic releaser. This cylindrical shell transducer has low sensitivity at certain angles from the axis of the cylinder. To realize a transducer with desired directivity under high hydro-pressure, we propose a new type of construction in which a high-sound-speed material called syntactic foam is attached to the inside or the outside of a cylindrical transducer shell element. T h e sound speed of the syntactic foam is higher than that of the water, so it is possible t o control the phase of the incident wave using this difference in sound speed between syntactic foam and water. In this paper, the geometrical arrangement and shape of syntactic foam is designed to realize an omnidirectional transducer. For ultrasonic transducers employed in deep seas, omnidirectivity is often required. A directivity pattern is calculated by changing the shape and arrangement of the syntactic foam attached t o a cylindrical transducer element. Measured directivity patterns of the transducer built according t o the numerical design are compared with calculated patterns.

2. 2.1.

Low Frequency Sound Source Principle of the low frequency sound source

Figure 1 shows the principle of the new type of low frequency sound source. T h e cylindrical transducer with outside diameter ro and height 2b is put between two metal circular plates with outside diameter 2a. The space between t h e two metal plates operates as a resonator in a reel shape and the resonance of the lowest mode in the radial direction is used. This reel shape resonator is equivalent to

P

2

To the p o i n t of observa t i 0 7

a Fig. 1. Principle of the low frequency sound source with an acoustic resonator in a real shape.

Fig. 2. Configuration for theoret,ical analysis.

367

ka

Fig. 3. Radiation impedance a t transducer surface 0.1).

r./o=

0. 5

IL

,o0.00

0. 1

0. 3

0. 2

0. 4

0. 5

01 0. 0

. 100

0. I

b/o

Fig. 4. Relation between k a at resonance and b / a ( T / a = 0.1).

(bla = 0.1, T / a =

0. 2

0. 3

0. 4

0. 5

b/a

Fig. 5. Relation between Q at resonance and b / u ( T / a = 0.1).

many resonant tubes in a horn shape (Report of a working group on low-frequency sound sources, 1983), so it is considered t o have a wide band frequency response. It is also very reliable a t high power operation and under high hydro-pressure due t o its simple construction. 2.2.

Analysis of the low frequency sound source

The characteristics of the low frequency sound source with the reel shape resonator are examined by theoretical analysis. Figure 2 shows an analytical model of the resonance characteristics. Two semi-infinite cylindrical baffles are separated by a distance of 2b; the space between the two baffles operates as a resonator. A

368

cyl indrical transducer

z T

Fig. 6. Scheme of low frequency sound source. cylindrical transducer with outside diameter 2r0 is located in the center of this space and vibrates uniformly. Acoustic radiation from the inside of the cylindrical transducer and walls of metal plates is neglected since it is considered t o be smaller than the acoustic radiation from the outside of the cylindrical transducer a t the resonance. Changing resonator diameter to 2 a , height 2b, and the outside diameter of the cylindrical transducer to 27-0, specific acoustic impedance a t the surface of the cylindrical transducer is calculated. When the surface of the transducer with radius ro vibrates with uniform vibration velocity U O , specific acoustic impedance Ztd at the transducer surface is obtained using the specific radiation impedance 2, of the open end calculated by analysis of the zonal vibration of a cylinder of infinite length (Nimura and Watanabe, 1953). Figure 3 shows the specific acoustic impedance at the surface of the cylindrical transducer as a function of ka in the case of b / a = 0.1 and ro/a = 0.1. There is a resonance point at k a = 2 and Q of the resonance is about 5.7. Figure 4 shows the relation between ratio of height and radius of resonator, b i n , and k a a t resonance of the specific acoustic impedance on the cylindrical transducer surface. T h e ratio between radius of resonator and outside diameter of transducer, T O / U , is 0.1. T h e product of wave number and resonator radius, ka, a t resonance decreases slightly with greater bla. Figure 5 shows the relation between b / a and Q of resonance of the specific acoustic impedance on the cylindrical transducer surface in the case of r o / a = 0.1. The quality factor Q of resonance decreases rapidly with greater b / a .

2.3.

Experiments and results

A low frequency sound source was fabricated aimed at a resonance frequency of 4 kHz and Q = 5 . The resonator is made of two iron disks with a thickness of 6 mm and is 282 mm in outside diameter (Fig. 6), so b / a is 0.106 and ro is 0.135. The PZT cylindrical transducer element used in these experiments is 33.4 mm in inside diameter, 38 mm in outside diameter and 30 mm in height. T h e cylindrical transducer is held between two metal circular disks with a rubber sheet by tightening the clamping bolts. Resonance frequency of the cylindrical transducer element

3 69

fdB

1

re

fiPa/V-ml

W i t h resonator

F r B qu e n cy Fig. 7. Frequency response of transmitting sensitivity.

0"

90

O

Fig. 8. Basic configuration of cylindrical transducer with syntactic foam. Syntactic foam is attached to t,he inside of the element. is 24 kHz. In measurements, a burst wave was used as a transmitting signal t o separate direct and reflected waves. The measured transmitting voltage response of the low frequency source is shown by the solid line in Fig. 7. Cylindrical transducer element sensitivity is also shown by the dotted line. Transmitting sensitivity

370

90"

I

i

Fig. 9. Basic configuration of cylindrical transducer with syntactic foam. Syntactic foam is attached to the outside of the element. increased by about 20 d B at resonance frequency 3.3 kHz with attachment of the reel shape resonator. The quality factor Q of resonance is 5.4.

3. Control Technique of Transducer Directivity 3.1.

Transducer construction

Figures 8 and 9 show the basic configuration of the cylindrical transducer with syntactic foam (Okujima et al, 1982). Syntactic foam is made of solidified hollow micro-glass balls using epoxy resin; this provides buoyancy under high static pressure. Since the acoustic impedance of the syntactic foam is nearly equal to that of the water, acoustic waves are transmitted without loss through the foamwater boundary. The 2500 m/s sound speed of syntactic foam is higher than the 1500 m/s of water. Using this difference in sound speed, the phase of the acoustic wave can be controlled. Syntactic foam is attached to the inside (Fig. 8) or the outside (Fig. 9) of the cylindrical transducer shell element.

3.2.

Design of omnidirectional transducer

In the design of a transducer with omnidirectivity, the syntactic foam is usually shaped into a cylindrical shell and is attached to the outside of the cylindrical transducer element as shown in Fig. 9. The directivity pattern was calculated by changing the length d of the syntactic foam projecting from the cylindrical transducer element end and the thickness t of the syntactic foam shell. We also calculated the directivity pattern of the transducer with syntactic foam shown in Fig. 8. In this configuration, however, there is little choice of the arrangement of syntactic foam and the directivity pattern was not as close t o the omnidirectional pattern. The directivity pattern is calculated by the finite element analysis considering an infinite space (Hachiya et al. 1985). An infinite space can be adopted in the finite element analysis in terms of the radi-

371 ation admittance. The radiation admittance at the boundary between the infinite space and finite element region is obtained using a spherical spreading procedure (Hunt et al., 1974). It is assumed that the sound field is axisymmetrical. Using the spherical harmonic expansion, the pressure a t a given point in axisymmetrical acoustic field can be easily obtained from boundary pressure. The PZT cylindrical transducer shell element used in these calculations and experiments is 46 mm in inside diameter, 62 nim in outside diameter and 30 mm in height. It is assumed that the inside and outside surfaces vibrate with the same amplitude and 180 degrees different phase. A contour map of the difference between the maximum and the minimum sensitivity in directivity is shown in Fig. 10. Frequency is 20 kHz. The thickness t of the syntactic foam shell and the projecting length d changes from 0 to 40 mm. A contour interval is 1 dB. Since the phase of acoustic wave radiated from outside of the cylindrical transducer element changes with the attachment of the syntactic foam cylindrical shell, the directivity pattern is changed by the projecting length d and the thickness t. The small value of this difference means that a directivity pattern is near the omnidirectional pattern. The directivity pattern a t t = 20 mni, d = 30 mm is closest to the omnidirectional pattern, as shown in Fig. 10. The difference between the maximum and the minimum sensitivity in directivity is about 0.4 dB. Figure 11 shows the calculated sound field near the cylindrical transducer element without syntactic foam. A contour interval is 5 dB. There is a low sensitivity region near the axis. This sound field results from interference between the sound waves radiated from the outside and the inside of the transducer element. The calculated sound field of the cylindrical t r a n s d x e r with syntactic foam at d = 30 mni and t = 20 mm is shown in Fig. 12. As is clear from this figure, the sound wave is radiated from the transducer in all directions with nearly equal intensity. Frequency : 2lkHz Cmm3

CmmJ

lhichess

t

Fig. 10. Contour map of sensitivity difference in directivity.

372 3.3. Experiments and results The cylindrical transducer with syntactic foam which was designed numerically was fabricated. The directivity pattern of this transducer was measured as a function of the angle of deviation from the axis over the limited range from -90 degrees to 90 degrees. Figure 13 shows the measured directivity pattern of the cylindrical transducer element without syntactic foam. This element has low sensitivity in the region near the axis and the difference between the maximum and minimum sensitivity in directivity is about 10 dB. Figure 14 shows the measured directivity pattern of the cylindrical transducer with syntactic foam a t d = 30 mm and t = 20 nim. As shown, the pattern is changed by attaching the syntactic foam cylindrical shell. The difference between the maximum and minimum sensitivity in directivity is about 3 dB which is considered nearly omnidirectional. Good agreement between the designed and measured patterns indicates t h a t numerical design for the shape and arrangement of syntactic foam is valid for realization of the desired directivity.

4.

Conclusion

A new type of construction of a low frequency sound source in which the cylindrical transducer is put between two metal circular plates has been presented. This source has wide band resonance and high reliability at, high power operation. The characteristics of the source, resonance frequency and quality factor Q, are examined using a theoretical model. Transmitting sensitivity of the low frequency sound source fabricated increases about 20 dB at resonance with the attachment of a reel shape resonator. It is clear that this structure can be useful for realization of a wide band low frequency projector for use in deep seas.

11. Sound field near the Fig. cylindrical transducer without syntactic foam. A contour interval is 5 dB.

Fig. 12. Sound field near the cylindrical transducer with syntactic foam. A contour interval is 5 dB.

373 r?

Lf

97f

0

-10

-20

-30

IdSJ

Fig. 13. Measured directivity pattern of cylindrical transducer element.

-& 0

gd -10

-20

-30

fd8J

Fig. 14. Measured directivity pattern of cylindrical transducer with syntactic foam ( t = 20 mm, d = 30 mni).

To realize a transducer with desired directivity under high hydro-pressure, we proposed a new control technique of transducer directivity in which a high-soundspeed material called syntactic foam attached t o a cylindrical transducer element. Optimum shape and arrangement of the syntactic foam for the omnidirectional transducer was obtained by numerical calculation using finite element analysis considering a n infinite space. T h e difference between the maximum and minimum sensitivities of the transducer made as a trial is 3 dB which almost agrees with the predicted value. Consequently, it is found t h a t this control technique of transducer directivity is useful, and the design of a transducer with desired directivity for use under high hydro-pressure in deep seas is proven t o be possible. References Worcester, P. F., R. C. Spindel and B.M.Howe, 1985. IEEE J. Oceanic Eng. OE-10, 123. Report of a working group on low-frequency sound sources, 1983. Marine Acoustic Society of Japan, 41 (in Japanese). Nimura, T., and Y. Watanabe, 1953. Rep. of Res. Ins. of Elect. Commu., 5(3), Tohoku University. Okujima, M., S. Ohtuski and H. Hachiya, 1982. Jpn. J .Appl. Phys., 22 (Suppl. 22-2), 39. Hachiya, H., S. Ohtuski and M. Okujima, 1986. Jpn. J. Appl. Phys., 25 (Suppl. 25-1), 88. Hunt, J T., M. R. Knittel and D. Barach, 1974. J. Acoustic. Sor. Am., 55, 269.

This Page Intentionally Left Blank

Deep Ocean Circulation, Physical and Chemical Aspects edited by T. Teramoto 0 1993 Elsevier Science Publishers B.V. All rights reserved.

375

Development of Multipaths Inverted Echosounder Tomoyoshi TAKEUCHI* and Keisuke TAIRAi

Abstract

A multipaths inverted echosounder was developed and tested for measuring vertical average currents by accurately measuring time differences for sound wave to travel in two opposite directions along a path in the ocean. The acoustic paths “sing around” vertical triangle with a base about 4 km long, and the time differences determine the horizontal component of velocity averaged through water column. The field tests were conducted by using two

multipaths inverted echo sounders (MIES). Each MIES provides the function of transponder and records all data. The results of the field tests are presented and discussed.

1.

Introduction

The so-called “sing around method” of flow measurement is based on the travel time difference of reciprocal acoustic waves between two points. Rossby (1975) discussed the principles and experimented in the laboratory. Motooka et al. (1985) also examined the method by laboratory experiments. On the other hand, the reciprocal transmission experiments aiming at the tomographic observation of midocean vortex was made by Worcester et al. (1985) in deploying two moorings 300 km apart in the Northwest Atlantic Ocean west of Bermuda. The same experiments were made by DeFerrari et al. (1985) in Florida Straits. Furthermore, Chaplin et al. (1986) has been investigating the measurement of horizontal currents averaged through the whole water column using two IES’s moored near the bottom. We have developed new inverted echo sounders which are named as multipaths inverted echo sounders by Professor Taira. We conducted an experiment of reciprocal transmission of acoustic waves around a triangle defined by one path directly between two of them moored at about 4 km horizontal spacing and about 200 m above the ocean bottom and another reflecting off the sea surface halfway between them. Here under described are the results of the experiment. By making the measurement in two orthogonal directions, the vectorial currents in the horizontal direction may be measured. Since the depth variation of the main thermocline can be measured with MIES a t the same time, the system may be very effective to monitor the transportation phenomena in the ocean. *University of Electro-Communications, 1-5-1,Chofu, Tokyo 182, J a p a n +Ocean Research Institute, University of Tokyo, 1-15-1, Minamidai, Nakano-ku, Tokyo 164, Japan

376 P

/t

/ \.

il'

CURRENT

..

L

Fig. 1 Schematic diagram for measuring mean currents in the ocean by using MIES's.

2.

The Theory

In Fig. 1, the acoustic waves emitted from MIESl transmit via the following four paths. Sound paths: I) MIESl - MIESZ, MIESZ - MIESl 2) MIESl - P (surface) - MIESZ, MIES2 - P (surface) MIESl 3 ) MIESl - bottom - MIES2, MIES2 bottom MIESl ~

~

~

Let the sound speed be c, and the reciprocal transmission time 6 t differences along a path is provided as

by setting a component of the flow velocity along the path which sound travels as u s , where u,/c << 1 as approximated. The approximation involved here is very good, since u,/c < lop3. Since the acoustic velocity varies in accordance with water temperature, pressure and salinity, the sound velocity can he provided as

c=co+6c

(2)

by setting the variation of the velocity as where c, is the mean sound velocity. Since 6c / c , < 0.003, equation (1) might be approximated as

377 Assuming that the vertical component of flow velocity in the ocean can be ignored and considering just the horizontal component of the velocity, we put the horizontal component in question as u ( z ) . In Fig. 1, if each time when the sound emitted from MIESl has traveled along the path of 1) and the path of 2) is respectively represented as t 2 or t 3 , they are provided as

ta = L+ + M - + N t 3 = L- + M + + N +

(4)

L , M , and N are respectively time for sound to travel along each side of the vertical triangle, whereas and - denote the opposite directions. From these equations, reciprocal travel time differences are provided as

+

6t = t3

- t2 =

(M+ - M - )

+ ( N , - N - ) - ( L , - L)

(5)

From equation ( 3 ) ,we obtains the following equations.

M+-M-

=

cz /”

--

u ( z )cot 0 d z

-0-6D

If the acoustic ray paths are almost straight lines, the following equation are valid. L‘

2 cot 0 = =

+ + 6L’

D

D+6D L‘ + 6L‘ cot.4 = 6D

Therefore.

is obtained. By putting mean currents ?i and G,

378

CURRENl

10‘

m

Fig. 2 Expected travel time difference as a function of horizontal distance between two MIES’s. Equation (8) is obtained from equation (7) and above equations.

Thus, it is known that the horizontal average currents can be measured, from equation (S), with the bottom mean flow as a reference. In equation ( 8 ) , L can be obtained from the travel time via the path l),where 4 might, be calculated by using L and the difference of depths between MIESl and MIES2. To measure the bottom flow, the synchronism of the clock of MIESl and MIES2 is required to be taken.

3.

Mooring Distance between MIESl and MIES2

When the average velocity in the horizontal direction is measured from equation (8) utilizing two units of MIES, the mooring distance should be a t first determined. Fig. 2 shows the relationship between the reciprocal acoustic travel time differences and the mooring distance of MIES’s taking the current as parameter when q5 is nearly equal 0. From the Figure, it is comprehended that the transmission time differences S t are 0.28 ms - 2.8 ms when L = 3 km and 0.4 ms - 4 ms when L = 5 km. Furthermore, they are 0.78 ms - 8 ms when L = 10 km, that is to say, the differences become large in accordance with the mooring distance between the individual units and the time measurement can be easily done. However, it should be kept in mind that the curve of acoustic ray paths will influence the measurement. Fig. 3 shows an example of the acoustic ray paths calculated from the profile of the water temperature measured with XBT. T h e acoustic waves emitted

("C)

Fig. 3 Sound speed profile (left, c) and ray paths (right) calculated water temperature profile (left, T) measured with XBT. X1 and X2 are points of moored MIES's. from the depth of 2800 m in the water arrive at sea surface and reflect off, and reach individually the position of the horizontal distance of approximately 4 km, and 10 km horizontal spacing in accordance with the emitting angles of 56' and 30" t o horizontal. These indicate that the acoustic ray paths are almost straight in the horizontal distance of 10 km and the effects of the curve of the acoustic ray paths can be ignored. When L is taken less than 10 k m , the acoustic transmission distance along the path 2) or 3) is less than 14 km. Thus, S/N required might be obtained by signal processing when the frequency of acoustic wave is 10 kHz. Time measurement accuracy should be less than 0.1 ms.

4. Acoustic Transceiver A source and receiver are integrated into a mechanical package t o form transceiver shown Fig. 4. Both are processed and controlled by a single CMOS V40 microprocessor. It operates on 1 MHz time base and is equipped with RS232C interface so that the transceiver and personal computer can be interconnected for d a t a communication. The transmitted signal was chosen to have the following characteristics: carrier frequency f o = 10 kHz digit length = 4 cycles of 10 kHz = 0.4 ms sequence length = 127 digits = 50.8 ms transmission length = 1 sequence period = 50.8 ms The transmitted signal is generated by reading out from PROM in which the phasecoded signal is stored as can be produced a square wave at frequency 10 kHz. The phase-coded digital signal is then amplified by a power amplifier, filtered by

380 cosw

,t

RAM

PROM

+ 128k bytes sinw

t

I

. . ,

128k bytes

x CPU

Reset(1 minute)

.

v10

I

PROM 64 k bytes

Receiver(on/off) RS-23ZC

Fig. 4 The block diagram of transmitter and receiver of MIES.

1 6 C ms

Fig. 5 Arrival peaks of correlation processing results of received acoustic signals (Sl: direct reached peak, S2: surface reflected peak, S3: surface and bottom reflected peak, S4: more surface and bottom reflected peak). a combination of broad band pass filter and impedance-matching network, and applied to the ring ceramic element.. Energy is supplied by lithium batteries. The receiver is microprocessor-controlled unit that digitize, process, and record the signal on EPROM. The projector transducer is used for acoustic signal reception also. A single CMOS V40 microprocessor performs both control and signalprocessing functions in the receiver. Signal acquisition is initiated by the processor at the source transmit time t o preset time for the maximal range. T h e incoming signal is amplified and filtered using a band-pass filter centered a t 10 kHz. The filtered signal is then shifted t o base band using a sinusoidal wave complex demodulator. The resulting quadrature components of the signal are digitized by followed

381 A/D converters per component. Each sample is the result of a 0.5 ms integration, giving a frequency response closely matched t o t h a t of the transmitted signal. A 0.2 ms sample interval is achieved to give two sample per digit. Since there are two quadrature components and 127 digits, a single period of the signal yields 254 complex samples. T h e processor squares each quadrature component and sums their results. The arrival time is obtained as the abscissa of resulting peak value. The transceiver responds after preset interval from that arrival time. The arrival time in each peak value and the some adjacent amplitudes are stored on 1 Mbits EPROM for later analysis.

5.

Preliminary Experiments

In the Zenisu Sea Area (30"N, 140"E) of 2.3 km water depth, acoustic transmission experiments for measuring mean currents were conducted by mooring MIES's spacing approximately 4 km about 800 m above the ocean bottom. Fig. 5 shows the records of waveforms detected, of which the acoustic waves emitted from MIESl traveled t o MIES2. We observe the four arrival peaks. The first peak is direct transmission sound. The second peak is the sound reflected from the surface and the third peak is the sound from the surface and the bottom. The last peak is sound which repeated several reflections from the surface and the bottom. The travel time differences will give the mean currents in the ocean from equation (8). It is explained that these have sufficient SNR. Reciprocal travel times along the multipaths ( l ) ,(2) are shown in Table 1. Each time base of IESl and MIES2 is synchronized at the time before deploying the MIES's and time is checked using a quartz oscillator. First column in the table shows the travel time along the direct path and second column shows the travel time along the path which the sound wave arrives at the sea surface, reflects off and reaches in each IESl to IES2 and its reverse transmission. The travel times in the first column shows the influence of clock drifts. Stl and St:! are travel time differences between the first and the second Table 1. Transmission times and reciprocal transmission time differences between two moored MIES's.

MIESl+MIES2 2731.0 3429.0 2719.5 3418.5 2714.5 3413.0 2723.0 3421.5 2715.0 3413.5 2724.0 3422.5 2697.5 3397.0 2687.5 3386.5 2684.5 3383.5

6t2

698.0 699.0 698.0 698.5 698.5 698.0 699.5 699.0 699.5

St, S t z - 6 t 3 (ms) MIES2+MIES1 2869.5 3567.0 697.5 0.5 2877.5 3572.0 697.0 4.0 696.5 2.0 2880.0 3576.5 2897.0 3591.5 694.5 4.0 2897.0 3594.0 697.0 1.5 696.0 2.0 2915.5 3611.5 2893.5 3592.5 699.0 0.5 2892.5 3592.5 700.0 -1.0 698.5 1.0 2900.0 3599.0 mean value 0.93

382 column respectively. T h e differences between these 6 t l and St:! are equal t o S t of equation (9). The mean transmission time along the direct path is 2801 ms. From this, L can be calculated as L = 4167 m since the water depth of each MIES is 1700 m and sound speed is 1488 m/s there. The mean value of reciprocal transmission time differences 6 t is 0.93 ms except anomalous value 4 s. Therefore, the mean currents is 26 cm/s directing from South West to North East from equation (8), since flows a t 800 m above the ocean bottom is almost 0 according to Anderra current meter which is moored along with MIES.

6.

Conclusion

An acoustic method was discussed for measuring average currents by accurately measuring time differences for sound wave t o travel in the opposite directions along a path in the ocean. Reciprocal acoustic transmission experiments were conducted by using two multipaths inverted echo sounders with the function of the slave transponder, and it was confirmed that the method might be effective for measuring average currents in the ocean. As the results, the reverberation by the sea surface is found t o be considerably strong. This will be one of the major sources of error. The sources of error such as scatters, sound speed variations, and mooring motion should be further investigated.

Acknowledgements This experiment has been conducted during the KT-91-4th research voyage of Tansei-maru belonging t o Ocean Research Institute, the University of Tokyo. The work was supported by the Ministry of Education (scientific research funds, Priority program, Dynamics of the Deep Ocean Circulation).

References Rossby, T., 1975. An oceanic vorticity meter. J. Marine Res, 32(2), 212-222. Motooka et al., 1985. Measurement of vortex by acoustic waves. In: The Ocean Charac-

teristics and their Changes, K . Kajiura (ed.), Koseisha Koseikaku (in Japanese). Worcester, P. F., et al., 1985. Reciprocal acoustic transmissions: Instrumentation for mesoscale monitoring of ocean currents. IEEE J. Oceanic Eng. OE-10(2), 123-136. DeFerrari, H. A,, et al., 1985. Acoustic reciprocal transmission experiments, Florida Straits. JASA 79(2), 299-315. Chaplin, G., et al., 1986. An acoustic ocean transport meter. IEEE Proc. Oceans, 426-429. Watts, D. R., 1977. Measuring dynamic heights with inverted echo sounders: Results from MODE. J. Phys. Oceanogr., 7(5), 345-358.