Oceanography of Asian Marginal Seas

OCEANOGRAPHY OFASIAN MARGINAL SEAS FURTHER T/TL€SIN THlS S€R/€S Volumes 7-7, 15, 16, 18, 19,2 1 and 23 are out of pri...

Author: K. Takano

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OCEANOGRAPHY OFASIAN MARGINAL SEAS

FURTHER T/TL€SIN THlS S€R/€S Volumes 7-7, 15, 16, 18, 19,2 1 and 23 are out of print. 8 E. LlSlTZlN SEA-LEVEL __ . _ _ CHANGES _ . .. .. ._ _ _ 9 R.H. PARKER THE STUDY OF BENTHIC COMMUNITIES 10 J.C.J. NIHOUL (Editor) MODELLING OF MARINE SYSTEMS 11 01 MAMAYEV ANALYSIS OF WORLD TENIPERATURESAL~NITY OCEAN WATERS 12 E.J. FERGUSON WOODand R.E. JOHANNES TROPICAL MAF~~NE POLLUTION 13 E. STEEMANN NIELSEN MARINE PHOTOSYNTHESIS 14 N.G. JERLOV MARINE OPTICS . ... .. ... ._ _ . .. _ 17 R.A.GEYER 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 GRAVITY2 4 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 MODELLINGMARINE HYDRODYNAMICS 27 R.A. GEYER (Editor) MARINE ENVIRONMENTAL POLLUTION 28 J.C.J. NIHOUL (Editor) MARINE TURBULENCE 29 M. M. WALDICHUK, G.B. KULLENBERGand M.J. ORREN (Editors) MARINE POLLUTANT TRANSFER PROCESSES 3 0 A. VOlPlO (Editor) THE BALTIC SEA 31 E.K. DUURSMA and R. DAWSON (Editors) MARINE ORGANIC CHEMISTRY 32 J.C.J. NIHOUL (Editor) ECOHYDRODYNAMICS 33 R. HEKlNlAN PETROLOGYOF THE OCEAN FLOOR 34 J.C.J. NIHOUL (Editor) HYDRODYNAMICS OF SEMI-ENCLOSEDSEAS 35 B. JOHNS (Editor) PHYSICAL OCEANOGRAPHYOF COASTAL AND SHELF SEAS 36 J.C.J. NIHOUL (Editor) HYDRODYNAMICS OF THE EQUATORIALOCEAN 37 W. LANGERAAR SURVEYINGAND 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 4 1 H. KUNZEDORF (Editor) MARINE MINERAL EXPLORATION 42 J.C.J. NIHOUL (Editor) MARINE INTERFACES ECOHYDRODYNAMICS 43 P. LASSERRE and J.M. MARTIN (Editors) BIOGEOCHEMICAL PROCESSESAT THE LANDSEA BOUNDARY 44 I.P. MARTINI (Editor) CANAD!AN INLAND SEAS

45 J.C.J. NIHOU and B.M. JAMART (Editors) THREE-DiMENstoNkL MODELS OF MARINE AND ESTUARINE DYNAMICS 46 J.C.J. NIHOUL and B.M. JAMART (Editors) SMALL-SCALE TURBULENCE AND MIXING IN THE n r FA N

47 M.R. LANDRY and B.M. HICKEY (Editors) COASTAL OCEANOGRAPHY OF WASHINGTON AND OREGON 48 S.R. MASSEL HYDRODYNAM~CSOFCOASTALZONES 49 V.C. LAKHAN and A.S. TRENHAILE (Editors) APPLICATIONS IN COASTAL MODELING 50 J.C.J. NIHOUL and B.M. JAMART (Editors) MESOSCALE IN GEOPHYSICAL TURBULENCE SYNOPTICCOHERENT STRUCTURES 5 1 G.P. GLASBY (Editor) ANTARCTIC SECTOR OF THE PACIFIC 52 P.W. GLYNN (Editor) GLOBAL ECOLOGICALCONSEQUENCESOF THE 1982-83 EL NINO-SOUTHERN OSCILLATION 53 J. DERA (Editor) MARINE PHYSICS

Elsevier Oceanography Series, 54

OCEANOGRAPHY OF ASIAN MARGINAL SEAS Edited by

K. TAKANO Institute of Biological Sciences University of Tsukuba Tsukuba, lbaraki 305, Japan

ELSEVIER Amsterdam - Oxford - New York -Tokyo

1991

ELSEVIER SCIENCE PUBLISHERS B.V. Sara Burgerhartstraat 25 P.O. Box 21 1, 1000 AE Amsterdam, The Netherlands Distributors for the United Stares and Canada:

ELSEVIER SCIENCE PUBLISHING COMPANY INC 655, Avenue of the Americas New York, NY 10010, U S A .

L i b r a r y o f Congress C a t a l o g i n g - i n - P u b l i c a t i o n

Data

Oceanography o f A s i a n m a r g i n a l seas / e d i t e d b y K. Takano. p. cm. ( E l s e v i e r o c e a n o g r a p h y series ; 5 4 ) I n c l u d e s b i b l i o g r a p h i c a l r e f e r e n c e s and i n d e x . ISBN 0-444-88805-5 1 . O c e a n o g r a p h y - - N o r t h P a c i f i c Ocean--Congress. I. Takano. 11. S e r i e s . Kenro. 1929GC79 1 .024 1991 9 1-8897 551.46'55--dc20

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CIP

ISBN 0-444-88805-5

0 Elsevier Science Publishers B.V.. 199 1 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 publisher, Elsevier Science Publishers B.V./ Academic Publishing Division, P.O. Box 330, 1000 AH Amsterdam, The Netherlands. Special regulationsfor readers in the USA - 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 USA. All other copyright questions, including photocopying outside of the USA, should be referred to the publisher. No responsibility is assumed by the Publisher for any injury and/or damage to 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 material herein.

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V PREFACE The fifth JECSS Workshop was held in Kangnung, Korea on 18-22 September 1989. JECSS, an acronym for Japan and East China Seas Study, was initiated by Prof. lchiye of Texas A&M University in 198 1 as an international forum by, and for, those who have interest in oceanography, especially in physical oceanography, of the Japan Sea, East China Sea and adjacent waters. It is a forum not only for exchange of information, views and data, but also for discussion and planning of joint studies. The workshop was convened by Profs. T. lchiye (Texas A&M University) and K. Takano (University of Tsukuba), organized by Profs. K. Kim (Seoul National University), Byung Ho Choi (Sung Kyun Kwan University) and Dr. Heung-Jae Lie (Korea Ocean Research and Development Institute). It was associated with the Regional Committee for the W ESTRAC of the Intergovernmental Oceanographic Commission (IOC), supported by the Research Institute of Oceanography of Seoul National University, and sponsored by the Oceanological Society of Korea, the Korean Oceanographic Commission and the American Geophysical Union. The JECSS workshops have been held every t w o years. The numbers of participants are listed below.

1st June 198 1 2nd April 1983 3rd May 1985 4th September 1987 5thSeptember 1989

Number of participants China Japan Korea Philippines Thailand UK USA USSR 6 7 27 1 6 6 27 7 6 8 30 7 7 15 33 9 55 1 1 1 9 4 13 30

Compared t o the previous four JECSS workshops held at Tsukuba in Japan the fifth workshop is worthy of special mention: through much effort of the local organizers, colleagues could attend the workshop from the People's Republic of China and the USSR which have no diplomatic relation with Korea, and from the Philippines, Thailandpnd the UK for the first time. Participationfrom many countries on Asian marginal seas is of crucial importance to JECSS activity. Seventy-four papers were presented at the workshop. This Proceedings volume contains thirty-one papers includingfive papers from Chinese scientists who could not attend the meeting because of difficulties with formalities. We would like t o thank the anonymous reviewers of the submitted papers for their comments and suggestions. Authors and titles of the papers which were presented at the workshop but do not appear in this volume are listed on the next pages.

VI The Proceedings of the first to fourth workshops were published in a special issue of La Mer, 20 (1982), of la Socibtb franco-japonaise d'ocbanographie, Elsevier Oceanography Series, 39 (1984), Progress in Oceanography, 17 (1986), pp. 1-399, and Progress in Oceanography, 2 1 (1989), pp. 227-536. Editors of the Proceedings Kenzo Takano (Editor-in-chief, University of Tsukuba) Kuh Kim (Seoul National University) Ya Hsueh (florida State University) Hsien-Wen Li (National Taiwan Ocean University) Fagao Zhang (Institute of Oceanology, Academia Sinica) Kazuo Kawatate (Kyushu University)

VII List of papers which do not appear in this volume

Keynote address._-____________-__--____________________------------------T. lchiye On the statistical features and variation in the oceanographic conditions in K. Fushimi, Y. Konishi and S. Ebara the Japan Sea ._________________________________ Seasonal and interannual variability in temperature of the upper layer off the southern coast of Korea. __________________-_____________________----H.-J. Lie Some features of water structure and dynamics in the northern Japan G.I. Yurasov, LA. Zhabin, V.G. Yarichin and Y.I. Zuenko Sea. -__________________ The relation between expansion of the coastal cold water and periodic wind farcing.------------------------------------------------------------------ S. Han and J. Na Numerical experiment on the circulation in the Japan Sea - Branching of the Tsushima Current. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . J.-H. Yoon and C.-H. Hong Numerical experiment of thermohaline circulation in the Japan Sea. --- Y. lsoda Polychaete assemblages on the continental shelf and slope of the East Sea (southeasternSea of Japan). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . J.W. Choi and C.H. Koh Influence of the Tsushima Current on the occurrence of the cladoceran Podon schmackeri in the Sea of Japan. ____-_____________________________ S.-W. Kim Drift bottle experiments on the current in the adjacent sea of Korea _-______________________________________------------------Y.G. Kim and K.-S. Chu Characteristics and distribution of physical and chemical properties of the East Sea in 1988 and 1989. ................................. K . 4 . Kim and K. Kim Clay minerals of the fine-grained sediments on the Korean continental shelves. ........................................................ Y.A. Park and B.K. Kim Study on suspended matter in seawater in the southern Yellow Sea. ............................................................. Y.S. Qin and S.L. Zhao Sediment distribution, dispersal and budget in the Yellow sea. ---- H.J. Lee and S.K. Chough Review of results obtained with moored acoustic doppler current R.L. Gordon profilers,_______________________________________--------A preliminary study of the current in the Yellow Sea. -- F. Zhang, W. Wang and K. Yu The formation of bottom water and interannual variability of its temperature in the yellow Sea. ............................................................. S.Tawara Numerical study on the ocean surface winds induced by tropical storms and its application to the prediction of sea surface drift in the East China Sea. ________________________________________--------H. Choi, K.M. Jang and B.G. Lee Water mass exchange between the Kuroshio and the East China Sea. The upwelling formed at the confluence of the Kuroshio and the East China Sea S.F. Lin and C.-T. Liu The structure of the Kuroshio west of Kyushu.-- C.C. Chen, R.C. Beardsley and R. Limeburner Dynamics of the East China Sea. I: The Kuroshio connection. -------- S.-Y. Chao Scaling invariance in Lagrangian observations of meso-scale S. Nakamoto, J.P. Liu, A.N. lndest and A.D. Kirwan, Jr. eddies _________-______ _____________________________________r__------------------------

VIII

Direct measurements of the Kuroshio in the East China Sea. --------- S.Mizuno, K. Kawatate and T. Nagahama Numerical simulation of the Kuroshio path near Taiwan. ...................... F. Yin An overview on using spaceborne sensors to study the physical processes --- W.T. Liu of the ocean. ________________________________________------------------ M. Kubota New technique of cloud test for AVHRR. ___________-__--_----------------Satellite thermal imageries of the Taiwan Strait. ------- S.P. Cheng, C.-T. Liu and S.-C. Chen Comparison of coastal zone color scanner (CZCS) imagery of western boundary current fronts in the East China Sea and off southeastern U.S. ............................................................................. J.A. Yoder Investigation on the diffusion and dispersion of thermal plume by remote sensing and numerical modelling. ...................... J.Y. Chung and T.S. Lee Measurements of near-shore surface currents by H.F. radar. ---------- D. Prandle Profiling of wind-driven currents in the ocean mixed layer. T. Hosoyamada and A. Kaneko Interpretation of the tides in the East China Sea in terms of the normal K. Rikiishi modes. ........................................................................... Scattering of tidal waves around an island on the shelf sea. ------------- S.H. Lee Numerical analysis of water waves for Pusan harbor under the third development plan. ............................................................. J.W. Lee Numericalexperiments on the tide and surge in the Korea Strait.----- K.-A. Park and H.-S. An On the sea level variations of the Japan sea in relation t o a typhoon passage. __________________-_____________________----------C.-H. Hong and J.-H. Yoon Geomorphology and sediments dynamics near the Keum River estuary.--------------------------------------------------------B.H. Choi and H.R. Yo0 Tsunami in the Japan sea ......................................................... C.N. Go Storm surges on the Japan Sea coast of USSR. ------------------ A.V. Ravinovich Some characteristics of the storm surges caused by typhoons passing through the south sea of Korea. .............................. I.S. Oh and S.4. Kim Seasonal characteristics of water masses at the entrance of the Korea Strait. ________________________________________--------------S.-K. Byun and K.4 Chang Dispersion of sediment suspensions in turbulent fluids. ---------------- Y. Noh and H.J.S. Fernando Rossby-Kelvin instability. _____--___-_____________________________~--------S.Sakai On the abyssal circulation in the Philippine Sea. ------ A. Masuda, K. Uehara and K. Taira On the origin of a warm ENS0 event in the western Pacific -- Y. Masumoto and T. Yamagata Radiating instability of nonzoal ocean currents with application t o the Kuroshio Extension. J.-Y. Yin Circulation in the South China Sea and the Gulf of Thailand------------ A. Siripong Physical oceanography in the Philippines: status and prospects. ----- G. Jacinto (in order of presentation)

I________-___

IX CONTENTS

Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

v

Observational characteristics of internal temperature fluctuations in the mid-latitude North Pacific. Jae-Yul Yun, James M. Price and Lorenz Magaard . . . . . . . . . . . . . . . . . . . . . . . .

1

Tidal computation of the East China Sea, the Yellow Sea and the East Sea. Sok Kuk Kang, Sang-Ryong Lee and Ki-Dai Yum ........................... 25 Nonlinear Rossby waves in the inertial boundary current and their possible relation to the variability of the Kuroshio. Qinyu Liu and Zenghao Qin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

49

Laboratory experiments of periodically forced homogeneous flow in a rotating cylindrical container. Jungyul Na and Bongho Kim . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

63

Remote sensing for modelling of variation in primary production fields. Akira Harashima . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

75

Reflection of the oceanic fronts on the satellite radar images. L.M. Mitnik and V.B. Lobanov . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

85

SST structure of the polar front in the Japan Sea. Y. Isoda, S . Saitoh and M. Mihara . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

103

numerical experiment on the seasonal variation of the oceanic circulation in the Japan Sea. Yoshihiko Sekine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

113

On the Intermediate Water in the southeastern East Sea (Sea of Japan) C . H . Kim, H.-J. Lie and K.-S. Chu . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

129

Modern sedimentation of San'in district in the southern Japan Sea. K. Ikehara . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

143

Effects of winter cooling on subsurface hydrographic conditions o f f Korean coast i n the East (Japan) Sea. Young-Ho Seung and Soo-Yong Nam . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

163

A

An observation of sectional velocity structures and transport of the

Tsushima Current across the Korea Strait. A . Kaneko, S.-K. Byun, S.-D. Chang and M. Takahashi

. . . . . . . . . . . . . . . . . . 179

Measuring transports through straits. D. Prandle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

197

A cross-spectral analysis of small voltage variation in a submarine

cable between Hamada and Pusan with speed variation of the Tsushima Warm Current. Kazuo Kawatate, Akimasa Tashiro, Michiyoshi Ishibashi, Takashige Shinozaki, Tomoki Nagahama, Arata Kaneko, Shinjiro Mizuno, Jyun-ichi Kojima, Toshimi Aoki, Tatsuji Ishimoto, Byung Ho Choi, Kuh Kim, Tsunehiro Miita and Yasunori Ouchi . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

207

Outflows from straits. Takashi Ichiye . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

223

X Dispersal patterns of river-derived fine-grained sediments on the inner shelf. S.C. Park and K.S. Chu . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

231

Estimation of atmospheric variables and fluxes on the ocean surface around Korean Peninsula. .................... . . . . . . . . . . . . 241 In-Sik Kang and Maeng-Ki Kim . . . . . . . Identification of water masses in the Ye ow Sea and the Eas cluster analysis. K. Kim, K.-R. Kim, T.S. Rhee, H.K. Rho R. Limeburner and .................. R.C. Beardsley .......................

China Sea by

. . . . . . . . . . . . 253

Synoptic band wintertime heat exchanges in the Yellow Sea. Y. Hsueh and James H. Tinsman, 111 ....................................

269

Numerical prediction of the vertical thermal structure in the Bohai and Huanghai Seas - Two-dimensionalnumerical prediction model. Wang Zongshan, Xu Bochang, Zou Emei, Gong Bin and Li Fanhua . . . . . . . . . . . 277 Development of towed vehicle systems for acoustic doppler current profiler. W. Koterayama, A. Kaneko, M. Nakamura and T. Hori .....................

293

A study of the Kuroshio in the East China Sea and the currents east of

the Ryukyu Islands in 1988. Yaochu Yuan, Jilan Su and Ziqin Pan

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

305

Three numerical models o f the Guanhe estuary. Zhang Dongsheng and Zhang Changkuan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

321

Geostrophic current and associated vertical motion off northeastern Taiwan. Hsien-Wen Li . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

335

Water volume transport through the Taiwan Strait and the continental shelf of the East China Sea measured with current meters. Guohong Fang, Baoren Zhao and Yaochua Zhu .............................

345

A subsurface northward current o f f Mindanao identified by dynamic

calculation. D. Hu, M.C. Cui and T.D. Qu

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

359

The mechanism of explosive cyclogenesis over the sea east of China and Japan. Simei Xie, Dingying Wei, Chenglan Bao and Aoki Takashi . . . . . . . . . . . . . . . . 367 The effect o f swell on the growth of wind waves. Hisashi Mitsuyasu and Yoshikazu Yoshida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

381

The thermal effluent problems of three nuclear power plants in Taiwan. K.L. Fan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

393

A simple dynamic calculation of launching a line. Kazuo Kawatate, Takashige Shinozaki, Yoshio Hashimoto, Tomoki Nagahama, Hideo Ishii and Akimasa Tashiro . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 405 Free transverse vibration of an elastic circular cylinder in a fluid. Guanghuan Liang and Kazuo Kawatate . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

421

1

OBSERVATIONAL CHARACTERISTICS OF INTERNAL TEMPERATURE FLUCTUATIONS IN THE MID-LATITUDE NORTH PACIFIC 1 Jae-Yul YunlY2, James M.Price3 and Lorenz Magaard 1 Department of Oceanography, U n i v e r s i t y of H a w a i i , Honolulu, Hawaii 96822, U.S.A. 2 Department of Oceanography, Korea Naval Academy, J i n h a e , Kyungnam, Korea 3 Hawaii I n s t i t u t e of Geophysics, U n i v e r s i t y of H a w a i i , Honolulu, Hawaii, 96822, U . S . A .

ABSTRACT Using the TRANSPAC XBT data at a depth o f 300 m, the regional variability of energy, time scales, length scales, and phase propagation of internal temperature fluctuations in the mid-latitude North Pacific is examined. lhe results show that the regional variability in the eastern half of the basin is substantially different from that in the western half. In the eastern North Pacific, the energy of the internal temperature fluctuation is very low and fairly uniformly distributed. Time scales and meridional length scales are distributed over broad ranges, zonal length scales are relatively small, and the direction of phase propagation is almost due west. At the eastern boundary, the opposite tendency in time and length scale distribution holds. In contrast, the energy in the western North Pacific is high, particularly along the main axis of the Kuroshio Extension Current (KEC), and decreases toward the east. It also decreases toward the north and south. Time scales are small near the western boundary and increase eastward. Both zonal and meridional length scales are large near the western boundary and decrease eastward. Phase propagation along the KEC appeared to be eastward, while that in the outer regions north and south of the KEC seemed to be westward with poleward components to the north and south, respectively. INTRODUCTION Since the development of the TRANSPAC XBT program (White and Bernstein, 1979), the number of studies on the variability of energy, time scales, length scales and phase propagation of the internal temperature fluctuations bas substantially increased for the mid-latitude North Pacific. It is well known from many studies that the potential energy or the variance of internal temperature fluctuations is much higher in the western North Pacific, especially along the axis of the Kuroshio current system, than in the eastern North Pacific with a relatively sudden change at about 17O0\J (Roden, 1977; ItJhite, 1977; Wilson and Dugan, 1978; Kang and Hagaard, 1980; Bernstein and White, 1981; White, 1982; Harrison et al., 1983; Mizuno and White, 1984). Concerning the time scales the results of earlier studies suggest that there are predominant interannual fluctuations (White and Walker, 1974; Price and Magaard, 1980; White, 1983) and near-annual ones with periods ranging from

2

several months to two years (Bernstein and White, 1974; fiery and Magaard, 1976; Bernstein and White, 1981; White, 1982) in the North Pacific. This is also clearly shown in Nagaard's (1983) composite model spectrum for the area of 20-25'N, 175-130%. The most commonly observed wavelength of the internal temperature fluctuations both in the zonal and in the meridional direction is reported to be in a range of 400- 1000 km in the mid-latitude North Pacific (Roden, 1977; Wilson and Dugan, 1978; Kang and Plagaard, 1980; Harrison et al., 1981). The phase propagation in the mid-latitude North Pacific, according to the results of earlier studies (Kang and Magaard, 1980; Bernstein and White, 1981; Wiite, 1982), is basically westward with a speed of 1-4 cm/sec. However, Mizuno and White (1984), using the same data set as the present one, recently found that the phase propagation is eastward along the axis of the Kuroshio Extension Current (KEC). Although the basic features of the observational characteristics in the mid-latitude North Pacific seem to be known from earlier studies, the regional variabilities of the characteristics have yet to be studied rather systematically. The purpose of this paper is to investigate the geographical distribution of energy, time scales, length scales and phase propagation of the temperature fluctuations at a depth of 300 m in the mid-latitude North Pacific. In this paper we utilize several standard analysis methods such as spectrum analysis, autocorrelation analysis, time-longitude and time-latitude contours, and harmonic analysis without discussing the methodology in detail.

THE DATA In this paper we use seasonally averaged temperature data at 300 m depth, mostly from the TRANSPAC XBT program and partly from the Japanese Oceanographic Data Center for the data in the region near Japan. This data set has 0.5' latitude by 0.5' longitude grid coverage from the coast of Japan to the coast 5 ' N and extends, in time, from sunnner 1976 to of California between 30 and 4 spring 1980. In the course of the edition of this data set each grid point has been interpolated first by fitting a trend surface to the nearest eight surrounding observations and then selecting the value at the grid point from this surface. This data set was provided by Dr. Warren White of the Scripps Institution of Oceanography. It has been used by White (1982) in the eastern North Pacific (18O-12O0W) and by Mizuno and White (1984) in the western North For more detailed descriptions of the data set the Pacific (130°E-1700W). reader is referred to their papers.

MEAN AND STANDARD DEVIATION The mean temperature profiles at a depth of 300 m (Fig. 1) indicate that the

3

45

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LONGITUDE Fig. 1. Mean temperature at 300 m depth KEC appezrs to be bifurcate at about 152OE and to become two nonzonal branches,

the northeastward (NE) branch (along the isotherms of 6-8OC) and the southeastward (SE) branch (along the isotherms of ll-14OC). This bifurcaFion of the mean flow may be triggered by the Shatsky Rise as is inferred from its location near the point of the bifurcation. The profiles also show that the mean temperature gradients decrease from about 3°C/100 km at 15OoE to about loC/10O !a at 165'E along the SE branch and to about 0.3°C/100 km at 16OoW indicating the eastward decrease in geostrophic velocity. The standard deviation of the temperatures (Fig. 2) shows a drastic change at

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LONGITUDE Fig. 2. Standard deviation of the temperature at 300 m depth.

173% along the main axis of the KEC suggesting that the vigorous current-related fluctuations terminate there. West of 173OW the standard deviation is mostly larger than 0.5OC and especially large along the SE branch of the KEC. East of 173% the standard deviation is mostly less than 0.5OC with a fairly uniform distribution. Along the SE branch, the standard deviation shows a distribution of highs and lows with intervals of 300-750 km between the neighboring maxima. These intervals may be considered as wavelengths of the fully developed finite amplitude instability of the KEC. The shape of the l0C contour line of the standard deviation indicates that the energy sources are apparently located along the axis of the KFC and that the energy decreases outward (particularly toward the east, north and south) from the sources. In the far-fields away from the KEC the l0C contour line of the

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o

4

standard deviation indicates a radiation of energy from the east toward the west, probably as stable Rossby waves. If areas with values of the standard deviation less than 0.5OC are considered as areas of background fluctuations, the meridional length scale of decay from the current-related source region to the northern region of the background fluctuations in the far-fields is about 1000 km. TIlE SCALES Frequency spectrum analysis Frequency spectra have been estimated using the direct Fourier transform method and a spatial smoothing instead of a smoothing in the frequency domain, because of the small number of data (16 points) in time. The Fourier coefficients have been calculated from the time series at each grid point. These coefficients are then averaged over each of the six subareas (Fig. 3 ) ,

LONGITUDE Fig. 3. Distribution of subareas I-VI and contours indicating the numbers

of the missing data point in time. Inside the label 1 there is no missing data point in time over the area west of 17OoV, over the area east of 17OoW, and over the entire study area. The frequency spectra are then calculated from the spatially averaged coefficients. Therefore, a total of 9 spectra have been produced. In this analysis the grid points with gappy data have been excluded in determining the spectra. Thus, the actual area covered by this analysis is within the contour line labled 1 in Fig. 3 . The same estimation procedure is repeated after removing the linear trend in each time series to determine to what degree the linear trend affects the spectrum. Since the data are not independent from a grid point to another, the effective number of degree of freedom has to be calculated to assign the confidence interval for each spectrum. Using the formula given by Emery and Magaard (1976) the effective number of degree of freedom is calculated for subarea I which has the smallest number of grid points ( 3 5 4 ) and it is 100. Since all the other subareas contain more number of grid points than subarea I does, the number of degree of freedom of 100 is used safely to assign 95%

5

confidence interval for each spectrum

-32T

0

32

T

In Figs. 4 and 5 the broken lines

SlBARFA I NO. OF AVG. 354

.:

2

4 6 8 FREQUENCY (CYCLES/16 SEASONS)

0

4 6 8 FREQUENCY (CYCLES/16 SFASONS)

32 -

SUBAREA IV NO. OF AVG. 789

m z

.'

2

SUBAREA v NO. OF AVG. 584

0

4 6 8 FREQUENCY (CYCLES/16 SMONS)

h

E 1.

c3

::

z2

I:

\

o!

0

,-;

-. -___ .-9

-

I

, I

L-

f

N U 0 ,

\

v

\

0

__

2

32 ..

HVI

, '\ ,

SUBAREA I11 NO. OF AVG. 677

SUBAREA VI NO. OF AVG. 377

'

: ,

\-

2

4 6 8 FREQUENCY (CYCLES/16 SEASONS)

Fig. 4. Spatially smoothed power spectra in each subarea. Solid (broken) lines indicate the spectra without (with) the linear trends.

indicate the spatially smoothed spectra with the linear trends included and the solid lines indicate those with the linear trends removed. In subarea I (near the western boundary), a considerable amount of energy is found only in a frequency range of 0.75-1.25 cycles/year (cpy). In this region the removal of the linear trends makes no difference for the shape of the spectra indicating a lack of very low frequency fluctuations. In subareas I1 and 111 (the interior western North Pacific) the energy levels are high (the highest in subarea 11) at all frequencies. In the spectra, distinct peaks are shown at 1 cpy (15% of the total energy) and 1.5 cpy (9%) in subarea I1 and at 1

6

cpy (16%) and 1.75 cpy (6%) in subarea 111. When the linear trends are removed, both peaks in subarea I1 become more distinct (from 15% to 23% for 1 cpy and from 9% to 14% of the total energy for 1.5 cpy) while both peaks in subarea I11 show little change (from 16% to 21% for 1 cpy and remained at 6% for 1.75 cpy). The spectra illustrate that the linear trends account for a larger portion of the energy in subarea I1 than in subarea 111, while the energy at the lowest frequency after the removal of the linear trend is higher in subarea I11 (67%) than in subarea I1 (33%). In subareas I V and V (the interior eastern North Pacific) the energy level is very low at all frequencies compared to that in subareas I1 and 111, as is anticipated from the distribution of the standard deviation. The energy at 1 cpy is not as distinct as that in subareas I1 and 111. In these subareas the spectra look quite different depending on whether the linear trends are included or not. When the linear trends are included, the energy level decreases drastically from the lowest frequency (62% in subarea I1 and 58% in subarea 111) to the next lowest (14% in subarea I1 and 15% in subarea 111) and then slowly decreases toward the higher frequencies. When the linear trends are removed, the energy at 1 cpy (16%) in subarea I V and at 0.5 cpy (32%) and 1 cpy (15%) in subarea V become important. In subarea V I (near the eastern boundary) the linear trend and the lowest frequency account for a large portion of total Energy but the portion is not as large as that in subareas I V and V. In this subarea the energy level is relatively high at 1.5 cpy (10 and 11% respectively with and without the linear trends) and 1.75 cpy (5 and 9%) as well as at 1 cpy (6 and 8%). This indicates a large increase in very small time scale fluctuations toward the eastern boundary. Hence, it is very clear that the spectra in subareas IV and V are distinct from those in subareas I1 and IJT. Jt i s certain that long-term fluctuations ir! the eastern North Pacific are different from those in the western North Pacific with respect to relavant time scale and the amount of energy. We speculate that the high energy at very low frequency in subareas I V and V may indicate the existence of the same low frequency signal as shown in Magaard's (1983) model spectrum. The spectra also indicate a different distribution of the small time scales between the two rygions. The most dominant periods are annual (23 and 34% respectively with and without the linear trends) in the interior western North Pacific (Fig. 5a) and annual (5 and IS%), biannual (15 and 20%) and semiannual (6 and 11%) in the interior eastern North Pacific (Fig. 5b). The average spectra over the entire study area (Fig. 5c) show the annual period as being the most dominant (13 and 34%).

7

(a) Western North Pacific NO. of Average 1583

32 h

cn

T

Eastern North Pacific NO. of Average 1491

..

HVI

GZ

\

\

::

\

'

0

2 4 6 8 FREQUENCY (CYCLES/16 SFASoqTs)

0

2

4 6 F R J ? Q m C Y

8

(CSCLES/16 srmws (C)

btire study area NO. of Average 3074

\ \

, 1

0

0

2

4 6 FREQLJENCY

8

( CYCLES/ 16 SEASONS)

Fig. 5. Spatially smoothed power spectra in the western half, the eastern half, and the entire study area. Solid (broken) lines indicate the spectra without (with) the linear trends. Autocorrelation analysis In order to study the time scales by a different method, an autocorrelation analysis method has been applied. This method determines the first zero-crossing time lag (hereafter called ZXTL) at each grid point. In this analysis the time series have been used without any filtering and the only grid points that have more than 12 consecutive data points are included in the analysis. For the present data, autocorrelations beyond the ZXTL are found to

8

be drastically reduced at most of the grid points. Therefore, one can consider the ZXTL as decorrelation time scales. Table 1 displays the total number of estimated ZXTLs in each subarea and TABLE 1

Total number of estimated first zero-crossing time lags in each subarea and their distribution (in percentage of the total number) in 7 categories of time lags.

< 3 3-6 6-9 9-12 12-15 15-18 > 18

48 30 11 8 1 0 1

36 37 11 7 5 3 1

35 34 12 10 7 2 0

14 17 15 13 19 17 5

29 16 15 13 13 9

4

58

23 10 7 2 0 0

their distribution (in percentage of the total number) in 7 categories of time lags. The most commonly occuring ZXTL is less than three months. A comparable decorrelation time scale of about 2 months has been obtained by Bernstein and White (1981) from about 2 years of the TRANSPAC XBT data in the western North Pacific. The next most commonly occuring ZXTLs are 3 to 6 months. The ZXTLs of less than three months occur most commonly in subareas I and V I (near both boundaries) and it is especially so in subarea V I . The ZXTLs are generally more broadly distributed in the interior than in the boundary region. In the interior region the 2 x n s are distributed over a narrower range in the western than in the eastern North Pacific. As one proceeds from subarea I to subarea I V , the distribution of the ZXTLs becomes broader with a gradual increase in number of larger time lags. However, the small ZXTLs of less than 3 months and 3 to 6 months are still dominant in subareas I1 and 111. The tendency of the distribution of the ZXTLs broadening toward subarea I11 is indicated by the percentage of the total number of ZXTLs in the first two categories of time lag (less than 3 months and 3 to 6 months) decreasing from 78% in subarea I to 73% in subarea I1 and then to 69% in subarea 111. The ZXTLs of 6 months or larger at 155-170°E as shown in Fig. 6 may be a consequence of the Shatsky Rise, since the topography can scatter energy toward larger time scales (Rhines and Bretherton, 1974). The ZxTLs are distributed over the broadest range in subarea I V . From subarea I V to subarea V I the number of small ZXTLs increase substantially, although the distribution tends to be broad still in subarea V. The percentage of time lags in the first two categories increase from only 31% in subarea I V

9

LONGITUDE Fig. 6. Geographical distribution of the first zero-crossing time lags (in months) of the temperature at 300 m depth. to 45% in subarea V and then to 81% in subarea VI. The percentages of time lags larger than 9 months are 10, 16, 19, 54, 39, and 9% in sunareas I to V I , respectively. This again indicates a possible existence of long-term fluctuations in the eastern North Pacific as discussed in the frequency spectrum analysis. However, the confidence interval of the autocorrelation function will be much larger at these larger ZXTLs than at the smaller ones. LENGTH SCALES

Wavenumber spectrum analysis Wavenumber spectra are estimated from the space series of the temperature data using the standard Blackman-Turkey method. For the zonal wavenumber spectra in the western North Pacific, the space series covering the distance from near Japan to the date line are used after removal of the spatial linear trends. The average record length is about 45 degrees in the zonal direction; In order to obtain a composite zonal wavenumber the sampling rate is 0.5'. spectra at latitudes, 30, 31, 45, individual spectra are estimated for each Then for each latitude lo, 32 season at latitudes'1 and (1+0.5)'. individual spectra (2 spatial points x 16 time points) are averaged at each wavenumber. The effective number of degree of freedom for the composite spectrum is estimated using the formula given in Ehery and Magaard (1976) and is about 60. This effective number of degree of freedom is used to assign the 95% confidence interval to the composite spectra. The composite wavenumber spectra N in the western North at several selected latitudes (33, 36, 39 and 4' Pacific and 32, 35, 39 and 4' N in the eastern North Pacific) are shown in Figs. 7a and 7b. For the zonal wavenumber spectra the energy levels are high at 33-36'N (the maximum is at 35-36'N) at most wavenumbers and low at all wavenumbers at 42-45'N (the minimum is at 42'N) in the western North Pacific (Fig. 7a). In the eastern North Pacific the energy levels are nearly two orders of magnitude smaller than the maximum energy in the western North Pacific. In the eastern

....

10

I

--

Z m l wavenunber spectrum I W a t e m Pacific)

16* 7 10-4

I

. .

IO-~ WAVENUMBER (CYCLES/KM)

IOP

WAVENUMBER ( CYCLES/KM

lo2

I

195% CONFIDENCE

-

10'

--

*e .3

W n u J

INTERVAL

176W 1 -

164W

2-0

142W

a '

5-

-

5:

10-1--

Meridonol wavenumpr spectrum (Western Pocific)

(d

)

(Eastern Pacific)

lo-*

lo2 10"

16" WAVENUMBER (CYCLES/KM)

10-

10-4

Iu3

10-2

WAVEWMBER (CYC L E S / K M l

Fig. 7. Zonal wavenumber spectra (a) at 33, 36, 39 and 44ON in N in the the western North Pacific and (b) at 32, 35, 39 and 4' eastern North Pacific. Meridional wavenumber sgectra (c) at 146, 159 and 176'E, and (d) at 176, 164 and 142 W. North Pacific there is very little difference in energy levels among different latitudes. The energy levels are slightly higher at 39% and slightly lower at 45ON in the eastern North Pacific. For the zonal wavenumber spectra the spectral slopes change at wavelengths of 600-750 !an in the western North Pacific and at wavelengths of 420-500 km in the

11

eastern North Pacific. The difference in the wavelengths at which the slope change occurs reflects the difference in dominant length scales. The slope changes are more gradual in the eastern than i n the western North Pacific. In the western North Pacific the spectral slopes at 35-45'N are -2.7 to -3.2 at wavelengths of 100-750 km and nearly flat at wavelengths larger than 750 km. However, those at 31-35'N are -2.7 to -3.0 at wavelengths of 100-750 km and -1 to -1.2 at wavelengths larger than 750 km indicating a broader distribution of length scales at 31-35'N than at 35-45'N. At 31-3Z0N there are also marginal peaks at a wavelength of about 750 km. The spectral slopes in the eastern North Pacific are -2.6 to -3.0 at wavelengths of 100-500 km. Hence, they are gentler than those in the western North Pacific. The gentler slopes in the eastern North Pacific indicate an increase in number of very small scale fluctuations. The slopes at wavelengths larger than 500 km are about -1 (slightly flatter at 31-34ON) in the eastern North'Pacific. The steeper slopes at this wavelength range indicate a more gradual change in length scale distribution. At 38-42'N a marginal peak is found at wavelengths 600-750 km. Composite meridional wavenumber spectra are estimated using the meridional Each composite spectrum is obtained by averaging space series covering . ' 5 1 spectra from 2.5' longitude (5 grid points) and 16 seasons. The effective number of degrees of freedom for these composite spectra is about 55. The spectra at several selected longitudes (146, 159, 176'E, 176, 164 and 142OW) are shown in Figs. 7c and 7d. In the meridional wavenumber spectra the highest energy is found at 138.5-146'E at all wavenumbers. The energy levels decrease quite monotonically toward the east at all wavenumbers. The various spectra in the eastern North Pacific have about the same shape and energy level at each wavenumber. The energy minima are found at 151.5, 149, 141.5 and 129OW. For these meridional spectra the spectral slopes change at a wavelength of

400 km. The spectral slopes are -2.5 to -3.4 over wavelengths of 100-400 lan with steeper slopes in the western than in the eastern North Pacific indicating more frequent occurence of small scale fluctuations in the latter region. The spectral slopes at wavelengths larger than 400 h are -0.4 to -1.0 with subareas I, IV, 11, 111, V, and VI in a sequence of descending magnitude of the slope. Autocorrelation analysis The autocorrelation analysis method is used to study the length scale distribution in both zonal and meridional directions in each subarea. In this method the zonal first zero-crossing space lags (hereafter called ZurSL) are computed in each subarea from the zonal space series of the temperature fluctuations. For the same reason as described in the autocorrelation analysis of the time series, the ZZXSLs can be interpreted as decorrelation length scales. Table 2 shows the total number of estimated ZZXSL in each subarea and

12

TABLE 2 Total number of estimated Z O M ~ first zero-crossing space lags in each subarea and their distribution (in percentage of the total number) in 9 categories of space lags. 150-170E 170E-170W 170-150W 150-130W 140-120W (446) (326) (438) (444) (4511

50-100 100-150 150-200 200-250 250-300 300-350 350-400

400-450 450-500

0 9 27 18 18 23 3 2 0

I1

I11

IV

V

vr

0 17 36 15 14 8 5

0 27 31 21 12 3

1 31 36 11 10

1 18 30

4

4

2 1

2 2 1

2 25 34 18 10 6 3 2 1

1

0

4

19 14 11 5 2

their distribution (in percentage of the total number ) in 9 categories of space lags. In this analysis only the space series that have more than 30 consecutive data points are included. The ZZXSLs are generally large and they are distributed over a broad range in subareas I and VI. In the interior North Pacific the ZZXSLs are generally small and they are distributed over a narrow range. The ZZXSLs are distributed over the broadest range in subarea VI, although the space lags of 150-200 km are still dominant (30%) among all the categories. The ZZXSLs which have the maximum percentage are the smallest and they are distributed over the narrowest range in subarea IV. The percentages of the total number of estimated ZZXSLs in the categories of 100-150 1071 and 150-200 km are 36, 53, 58, 67, 59, and 48% in subareas I-VI, respectively. As one proceeds away from subarea IV toward both boundaries, the number of large ZZXSLs increase gradually. The percentages of the ZZXSLs larger than 300 lan are 28, 1 7 , 10, 9, 15, and 19% in subareas I-VI, respectively. For the entire study area the most commonly occuring ZZXSLs are 150-200 lan. Meridional first zero-crossing space lag (hereafter called MZXSL) is estimated at each longitudinal grid point and for each season. In this analysis only the meridional space series that have more than 20 consecutive data points have been included. Table 3 shows the total number of estimated MZXSLs in each subarea and their distribution (in percentage of the total number) in 8 categories of space lags, In general the MZXSLs are smaller than the ZZXSLs. The MZXSLs are distributed over a narrow range in subareas I and VI but the most comonly occuring MZXSLs are different between the two subareas. They are 200-250 lan (46%) in subarea I while 100-150 lan (50%) in subarea VI. In the

13

TABLE 3

Total number of estimated meridional first zero-crossing space lags in each subarea and their distribution (in percentage of the total number) in 8 categories of space lags. 140-150E 150-17OE 170E-170W 170-150N 150-130W 130-120W (279) (625) (640) (640) (631) (164)

50-100 100-150 150-200 200-250 250-300 300-350 350-400

400-450

I

I1

I11

1 21 25 46 7 0 0 0

1 28 39 24 7 1 0 0

2 38 31 19

a

1 1 0

IV

V

VI

3

7 39 28 12 10 2 1 1

12

21 28 13 17 14

4 0

50 27 9 1 1 0 0

latter subarea the percentage at 50-100 km is also large with 12%. In the interior region they are distributed more broadly in subareas IV and V than in subareas I1 and 111. The broadest distribution is shown in subarea IV. A s one proceeds from subarea I to subarea 111, the most commonly occuring MZXSLs change from 250-300 km (46%) in subarea I, to 150-200 km (39%) in subarea 11, and then to 100-150 km (38%) in subarea 111. From subarea I11 to subarea IV the distribution of the PlZXSLs becomes broad quite suddenly with an increase in number of large space lags. In subarea V the distribution of the MZXSLs is over a narrower range than in subarea IV and the commonly occuring MZXSL is again 100-150 Ian (39%). The percentages of the MZXSLs in the first three categories (the space lags of 50-100, 100-150, and 150-200 !a) are 47, 68, 71, 52, 74, and 89% and those in the range larger than 250 !an are 7, 8, 10, 35, 14, and 2% in subareas I-VI, respectively. PHASE PROPAGATION Time-longitude time-latitude contours Time-longitude matrices at each full degree of latitude are contoured to examine the direction of zonal phase propagation. Figs. 8a-h exhibit the contours at each odd-numbered latitude. In the figures an upward slope of contour toward the west indicates westward phase propagation. The zonal phase propagation at 44-45'N (Fig. 8a for the contour at 45'N) does not show any 3 ' N (Fig. preferred direction as there is no consistent slope. However, at 4 8b) there is a tendency of upward slopes to the west intermittently over short time periods and locally over short distances. Westward phase propagation is relatively evident at 37-42'N except in the regions near both eastern and western boundaries (Figs. 8c-e at 41, 39 and 37'N, respectively). The westward phase propagation is most pronounced at 3 9 ' N with speeds of about 2

14

(a) 4 5 ' N

(b) 43'N

JXNGITUDE

(c) 41°N

(d) 3 9 ' N

Fig. 8. Time-longitude plots of the temperature (at 300 m depth) variability at each odd-numbered latitude. The contour interval is 0.5OC.

it becomes less evident. West of 160°E the direction of cm/sec. At 37-38'N phase Propagation is not obvious even at 37-45'N. At 34-38'N (Figs. 8e and f at 37 and 35'N, respectively) there appears to be eastward phase Propagation in the Western North Pacific in some seasons, and it Shows most clearly at 35-36ON. The eastward phase Speed is about 3 aii/sec for the first few years at 160°E-170%. Hmever, even in the region where phace Propagation appears to be eastward, there is some indication of westward phase Propagation at least for a short period of time.

15

LONGITUDE

Fig. 8. (continued) Pacific westward phase propagation appears to occur even at 35-36'N. Recalling that this latitudinal band is basically located along the southeastward branch of the KEC west of 165'E and located in the region between the two branches at 165-17goW, eastward phase speed indicates the disturbances being carried along the axis of the current (Hansen, 1970; lzizuno and White, 1984). At 30-33'N eastward phase propagation tends to disappear and westward phase propagation tends t o appear as one prodeeds southward (Fig. 8h at 31°N). The phase speed seems t o be higher toward the south, for example, about 4.5 cm/sec between 165OE and 168OW for the last one year

16

period at 31°N. The-latitude Rwtrices are contoured in Fig. 9 to examine the direction of

LATITUDE

LATITUDE

16 14

16 14

z l2

2 10 O

(d) 171°E

(c) 159OF,.

cn 6

8z

l2 lo r n 8

6 4 2

4 2 o ~ - z f a m o ~ - z f m m m m m u e e

LATITUDE

LATITUDE

LATITUDE

LATITUDE

Fig. 9. Time-latitude plots of the temperature (at 300 depth4 variability at several selected longitudes. The contour interval is 0.5 C.

meridional phase propagation. The slopes of the contours indicate phase propagation from the axis of the KEC to the north and south. The northward and southward phase propagation respectively in the regions north and south of the axis of the KEC are relatively evident at 141-171OE. A typical phase speed both to the north and south is approximately 1-2 cmlsec, although the estimates are often available over short period of time only.

17

Harmonic analysis A harmonic analysis method is employed to investigate phase propagation of the temperature fluctuations for the annual period. Percentages of the total variance explained by the annual harmonics are also obtained. In order to reduce the error bar the original half degree grid data have been averaged over each 1' latitude by lo longitude before applying the method. The percentages of the total variance explained by the annual harmonics (raw bandwidth of 10.7-13.7 months) are mostly 20-40% in the western North Pacific and 10-20% in the eastern North Pacific. This is an agreement with the results of the time scale analysis. The zonal phase propagation for the annual harmonics is shown in Figs. 10a-h. In the figures westward phase propagation is indicated by an upward

130

140

I50

160

170

180

170

160

150

140

130

120

LONGITUDE

Fig. 10. Phase of the annual harmonics along zgnes at (b) 43O, (cb 41°, (d) 39 , (e) labitudes og (a) 45' 37 , (f) 35 , ( 8 ) 33*, and (h) 31 N.

18

t

130

140

I50

160

170

180

170

160

1x1

140

130

120

LONGITUDE

Fig. 10. (continued) slope to the west. This analysis shows generally the same phase propagation characteristics as discussed in the time-longitude matrices. The westward phase speed at 43ON is roughly 2.1 cm/sec for longitudes between 180 and 168OW. At 3 9 ' N the westward phase propagation is very obvious at 150°E-1400W, and the phase speed between 167'E and 167OW is about 1.9 cnfsec. At 35 and 3 6 ' N eastward phase propagation occurs between 165'E and 170% with a speed of roughly 3.8 cm/sec (Fig. 10f at 35ON). As one proceeds away from 3 5 ' N toward the south, westward phase propagation becomes relatively apparent again. At 31-33ON the westward phase speeds are roughly 2.5-3.2 cm/sec. The meridional phases are shown in Fig. 11. It turns out to be possible to categorize areas with similar distributions of meridional phases while the

19

a 162-161 W

4 5 4 0 3 5 3 0

LAT I T UDE

LATITUDE

LATITUDE

-

Fig. 11. Phase of the annual harmonics along meridians at various longitudes. The range of longitudes within a box is selected to group similarly looking curves.

distribution changes wildly from an area to the next. The meridional phases for 154-165'E show a considerably persistent pattern in which the relative phase minima exist around 36'N and the phases increase from the region of the phase minima toward the north and south. The phase speed is approximately 2 . 3 cmlsec both to the north and south. This tendency of phase propagation to the far-fields from about 3 6 ' N seems to prevail only in the western North Pacific. In some cases there are two relative phase minima indicating that the energy sources are located along the two branches of the current. East of 165OE, meridional phase propagation often appears to be alternating northward and southward.

20 I

I

I

'

.'.

I

'

'

155-154w

.*

0

-I I

+

t

151-IsbW

- t

t

ii

w 40 35 30

45

LATITUDE

v

LATITUDE

LATITUDE

Fig. 11. (continued)

The distribution of the phase of the annual harmonics over the entire study 0 area is displayed in Fig. 12. In this figure only the 120 phase lines are

LONGITUDE Fig. 12. Phase of the annual harmonics in the study area. The shaded region indicates phase less than 120'.

contoured i n order to see a simple phase pattern. The figure shows fairly

21

regular intervals of the phase lines, except in the regions along the axes of the KEC. The figure appears to indicate that the phase propagates to the east along the axis of the KEC, to the northwest in the region north of the axis, and to the southwest in the region south of the axis in the western North Pacific. In the entire eastern North Pacific the phase propagates to the west. SUMMARY AND CONCLUSIONS

The energy levels are much higher in the western (the highest in subarea 11) than in the eastern North Pacific at all frequencies with a drastic chmge occuring at 1730W. The annual peak is highly distinct in the western North Pacific but in the eastern North Pacific it becomes distinct when the linear trends are removed. The linear trend accounts for a large portion of the total energy in the eastern North Pacific indicating a possible existence of strong interannual fluctuations. For the zonal wavenumber spectra the energy levels are highest at 35-36'N and lowest at 4Z0N at most wavenumbers in the western North Pacific. They are highest at 39'N and lowest at 45ON in the eastern North Pacific, but with only a small difference. For the meridional wavenumber spectra the energy levels are highest at 138.5-146'E and lowest at 151.5, 149, at all wavenumbers. 141.5 and 129'14 In subarea I (near the western boundary) the time scales are small and both zonal and meridional length scales are large in general. In this subarea the time and meridional length scales are distributed over a narrow range but the zonal length scales are distributed over a broad range. The dominant period, zonal wavelength and meridional wavelength are shorter than a year, ranges of 600-800 km and 1200-1400 km and a range of 800-1000 km, respectively. In subarea VI (near the eastern boundary) the time and meridional length scales are very small but the zonal length scales are generally large compared to those in subareas IV and V. In this subarea the time scales are distributed over a narrow range and the zonal length scales are distributed over a broad range as in subarea I. However, the meridional length scales are distributed over a narrow range unlike those in subarea I. The dominant period, zonal wavelength and meridional wavelength are shorter than a year, a range of 600-800 lan and a range of 400-600 h,respectively. Subareas I1 and I11 (the interior western North Pacific) are regions of gradual transition for the time series from a narrow range of distribution of predominantly short scales in subarea I to a broad range with an increase in number of large scales in subarea IV. The gradual transition also occurs for the zonal length scales from a broad distribution with relatively large scales in subarea I to a narrow distribution of predominantly small scales in subarea IV. The meridional length scales are distributed over a narrow range in the western North Pacific but the dominant length scales changes gradually from

22

large scales in subarea I to small scales in subarea 111. The dominant period, zonal wavelength and meridional wavelength are respectively near-annual, a range of 600-800 km, and ranges of 600-800km in subarea I1 and 400-600 !a in subarea 111.

In subareas IV and V (the interior eastern North Pacific) the time and meridional length scales are distributed over broad ranges with increase in number of large scales compared to those in subareas I1 and 111. However, the zonal length scales are distributed over a narrow range of predominantly small scales in these subareas. In subarea IV the distributions of the time and meridional length scales are the broadest while that of the zonal length scales are over the narrowest range (especially at 38-42'N) among all the subareas. In general, the zonal length scales are smaller than the meridional ones in this subarea. In these subareas there appears to be no single dominant time scale, but the near-annual, biannual period and interannual fluctuations may be similarly important. The dominant wavelengths are in a range of 400-800 km in both directions. From subarea IV toward subarea VI the number of small time and meridional length scales and the number of large zonal length scales tend to become large. Phase propagation is to the east along the axis of the KEC, to the northwest in the region north of the axis and to the southwest in the region south of the axis in the western North Pacific. It is to the west in the eastern North Pacific. The zonal phase speeds are about 3-4 cm/sec eastward along the axis of the KEC and 2-4 cmlsec westward in the region away from the KEC with a slight increase toward the south. The meridional phase speeds are about 2 cmlsec to the north and south in the western North Pacific. REFERENCES Bernstein, R., and W. White, 1974. Time and length scales of baroclinic eddies in the central North Pacific. J. Phys. Oceanogr., 4: 613-624. Bernstein, R., and W. White, 1981. Stationary and traveling mesoscale perturbations in the Kuroshio &tension current. J. Phys. Oceanogr., 11: 692-704. Emery, W. J., and L. kgaard, 1976. Baroclinic Rossby waves as inferred from temperature fluctuations in the eastern Pacific. J. Mar. Res., 34: 365-385. Hansen, D. V., 1970. Gulf Stream meanders between Cape Hatteras and the Grand Banks. Deep-sea Res., 17: 495-511. Harrison, D. E., W. J. Emery, J. P. Dugan, and B.-C. Li, 1983. ?,lid-latitude mesoscale temperature variability in six multiship XBT surveys. J . Phys. Oceanogr., 13: 648-662. Kang, Y. Q., and L. Magaard, 1980. Annual baroclinic Rossby waves in the central North Pacific. J. Phys. Oceanogr., 10: 1159-1167. Magaard, L., 1983. On the potential energy of baroclinic Rossby waves in the North Pacific. J. Phys. Oceanogr., 7: 41-49. Mizuno, K., and W. B. White, 1984. Annual and interannual variability in the Kuroshio current system. J. Phys. Oceanogr., 13: 1847-1867. Price, J. M., and L. Magaard, 1980. Rossby wave analysis of the baroclinic

23

potential energy in the upper 500 meters of the North Pacific. J. Mar. Res., 38: 249-264. Roden, G. I., 1977. On the long-wave disturbances of dynamic height in the North Pacific. J. Phys. Oceanogr., 7: 41-49. White, W. B., 1977. Secular variability in the baroclinic structure of the interior North Pacific from 1950-1970. J. Mar. Res., 35: 3, 587-607. White, W. B., 1982. Traveling wave-like mesoscale perturbations in the North Pacific. J. Phys. Oceanogr., 12: 231-243. White, W. B., 1983. Westward propagation of short term climatic anomalies in the western North Pacific Ocean from 1964-1974. J. Mar. Res., 41: 113-125. White, W. B., andL. Bernstein, 1979. Design of an oceanographic network in the mid-latitude North Pacific. J. Phys. Oceanogr., 9: 592-606. White, W. B., and A. E. Walker, 1974. Time and depth scales of anomalous subsurface temperature at Ocean Weather Station P, N, and V i n the North Pacific. J. Geophys. Res., 1979: 4517-4522. Wilson, W. S., and J. P. Dugan, 1978. Mesoscale thermal variability in the vicinity of the Kuroshio Extension. J. Phys. Oceanogr., 8: 537-540.

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25

TIDAL COMPUTATION O F THE EAST CHINA SEA. THE YELLOW SEA AND THE EAST SEA

Sok Kuh KANG. Sang-Ryong LEE* and Ki-Dai W M Coastal Engineering Laboratory, Korea Ocean Research and Development Institute *Department of Marine Science, Pusan National University, Korea

ABSTRACT The M. tidal phenomena in the entire surrounding seas of Korea Peninsula have been investigated under the condition of one model area, majorly to understand which factors govern the amphidromic system in each sub-area. In this study two-dimensional numerical model based upon an implicit scheme has been used. Due to the large depth differences of each sea in the modelled region the tide generating force is included as well as other inertia terms. The model region toward south and south-eastern boundaries is extended to Ryukyu Islands to utilize the measured data. The computed M, co-tidal and co-amplitude lines in the Yellow Sea and the East China Sea are in good agreement with those of existing tidal charts. The four amphidromic points known to exist in the Yellow Sea and the East China Sea as well as two amphidromic points in the East Sea (Japan Sea) are successfully reproduced. The results of computation also show that the tide generating forces play greatly different roles in the each sub-sea model region. The tidal amplitude due to the tide generating force explains several percent of the tide generating force in the Yellow Sea and the East China Sea. It is shown that the realistic amphidromic system in the East Sea can not be explained without the tide generating forces. It is inferred that the reflected wave at the head region of Tartary Bay is more reinforced by the tide generating force. Also the role of the driving force at the head of Tartary Bay for the amphidromic system is discussed. An analysis for tidal volume fluxes and volume transport is made for the five straits existing within model area to assess the possible contribution to tidal regime in the linked sub-model area. And the tidal volume fluxes through live straits in the model region are compared with those calculated by simple analytical method (Defant, 1961) and those calculated by Odamaki (l989b).

I. INTRODUCTION In this paper the M, tidal regimes of the East China Sea, the Yellow Sea and the East Sea (Japan Sea) are investigated under one model area by using a two-dimensional depth averaged numerical model based upon an implicit scheme. The model domain covers the whole region of the Yellow Sea and the East China Sea extending to the southeast as far as Ryukyu Islands, to the south as far as Taiwan Strait as well as covering the East Sea connected by the Korea Strait (see Figure I). The depths used for the model were obtained from SYNBAPS II data set from MIAS and

US chart No. 529. As shown Figure 2, the mean depth of each sea in the model area is very different from each other. It is also noted that there exist abrupt changes of depth at several regions. From the Ryukyu Islands, the ocean bottom rises rapidly from 2,000 m to 200 m and the depths of the entire north-western model area are less than 100m. The East Sea is a semi-closed sea connected with other seas by four straits : the Korea Strait between Korea and Japan, the Tsugaru Strait between Honshu and Hokkaido. the Soya Strait between Sakhalin and Hokkaido and the

26

55'N

50

40

30

@

East China Sea ,,

ge

.$

'8

@ Korea Strait @ Tsugaru Strait

0 Soya Strait 0 Mamiya Strait

South China Sea

20 120

t30

140OE

Figure. I. Map showing the geographical names.

21

Figure. 2. Bathymetric map of the modelled area. Numerals show depth in meters.

28

Mamiya Strait between Sakhalin and the continent. The most comprehensive work on the tides and the tidal currents in these boundary seas was done by Ogura (1933). who has edited the co-tidal M, and co-range (M.+S.) charts in these seas using the method of Proudman and Doodson (in Defant(l961)). based upon a considerable number of tidal measurements. And recently the co-tidal and co-range charts of semidiurnal and diurnal tides in the East China Sea, the Yellow Sea and the East Sea are redrawn with the new tide and tidal current data by Nishida (1980). The M2 tidal charts are shown in Figure 3(a) and (b) following Nishida. According to the tidal chart of MI tide there exist four amphidromic points in the Yellow Sea and the East China Sea and two amphidromic points in the East Sea. It is worth noting that the large tidal ranges up to 6 m along the west coast of Korea occur while those of the East Sea are less than 0.2 m in the central part of the sea. In this paper the tidal charts by Nishida are used for the comparison between model-generated tidal charts and existing ones because Nishida's charts have an explicit expression of co-amplitude chart of

while that

by Ogura (1933) is for (M2+S2) co-range chart. Recently Fang (1986) also edited the tidal charts for M. tide based upon the observed tide on coast and islands. Especially the differences from those of Ogura and Nishida appear in the locations of the amphidromic points occurring in Liatung and Pohai bays. As noted by Defant (1961). the tidal phenomenon of the entire East China Sea is almost exclusively conditioned by those water-masses which penetrate through the canals between the Ryukyu Islands within a half tidal period and then flow out again within the following half tidal period. Defant also described that the tides of this boundary seas (the East China Sea and the Yellow Sea) are essentially of Pacific origin as a result of qualitative analyses of water transport during 6 h of the M2 tide through the canals of the Ryukyu Islands (350 km'), the northern entrance of the Taiwan Strait (130 km') and the Korea Starit (20 km'). In this paper the qualitative assessment of water transport during 6 h of the M , tide through Taiwan Strait and Korea Strait as well as three other straits in the East Sea will be given based upon the computed results of the numerical model. Besides the works mentioned above there have been several analytical or numerical studies to investigate the tides of the Yellow Sea and the East China Sea or the East Sea. An (1977) developed a Yellow Sea numerical model with the open boundary being established on a straight line from southern tip of Korea to Shanghai of China. He gave some explanation for the high tidal ranges in Kydnggi Bay in terms of its shallowness and a resonant oscillation of 10 hour period. He also commented that the amplitude due to the tide generating potential is very small (less than about 3 percent) compared to that of the boundary condition in the Yellow Sea. An extensive study by numerical model for the tidal phenomena in the Yellow Sea and the East China Sea was carried out by Choi (1980). In his study he employed the modelling techniques based upon an explicit scheme. The model area in his study extends to the south as far as Nothern Taiwan (Formosa) and seaward as far as the edge of the continental shelf with the grid resolution l/5" in latitude and 1/4" in longitude. In his study general arnphidromic system in the Yellow Sea were reasonably reproduced while in the Liatung Bay it appears as deamphidromic pattern. Through the comparison of the cotnputed tidal constants with the observed ones of A4: tide he showed that the differences are approximately 10 percent in amplitude and loo in phase.

29

I

I

I

I

.55"N

.50

.40

30

120

Figure. 3. Tidal charts for the

130

140"E

tide in the seas around the Korea Peninsular (Nishida. 1980)

(a) Co-amplitude chardamplitude : cm).

30

j5"N

50

40

30

31

But as the open boundaries in the study are located at the edge of the continental shelf, he had to use the tidal charts edited by Ogura (1933) for the open boundary input. He also does not include the tide generating force in governing equations. This tidal model has been further refined to resolve the flow over the continental shelf in more detail with the resolution of l / l S O in latitude and 1/12’ in longitude in another two-dimensional study (Choi. 1988). Kang (1984) has investigated the co-oscillating tides in the Yellow Sea south of Shantung Peninsula analytically by a superposition of Kelvin and Poincare waves. He explained that the large tidal range of Kydnggi Bay is due to the modifications of Poincare and Kelvin waves by Ongjin Peninsular. He also showed that the asymmetry of amphidromic system of M, tides arises primarily due to the partial penetration of tidal energy through the opening at the bay head. He also expected that the tidal energy dissipation by bottom friction might contribute to the asymmetry of the amphidromic system. But for mathematical simplicity the frictional effect to investigate the asymmetry was excluded in his analyses.

For the tide of the entire East Sea a hydrodynamical theory has been given by Ogura on the basis of the canal theory. Recently Odamaki (1989a) re-edited co-tidal chart in the Korea Strait

on the basis of intensive observed data and the amphidromic point of the M, are similar to Ogura’s result while the amphidromic points of the K , and 0,tides are remarkably shifted toward the Korean coast compared to Ogura’s charts. Odamaki (1989b) also investigated the generation of tides in the East Sea with relation to the tidal volume fluxes at the attached straits, which are estimated with the observed tidal current data. In his study he separated the tides into the co-oscillating tides induced by tidal volume fluxes and the independent tide by tide-generating force using an one-dimensional tidal model. Through his study he estimated that the amplitude of the independent tide in the Tartary Bay is the same as that of the co-oscillating tide attributed to the Korea Strait. But due to the limitation of the one dimensional study the effect of tide generating force on the amphidromic system could not be realistically investigated and the contribution of the Mamiya Strait to the tidal regime of the East Sea was ignored by letting the strait closed. In this study the M, tide in the entire surrounding seas of Korea Peninsula is investigated to understand which factors determine the amphidromic system. To meet the requirement that nearly same accuracy should keep over the largely different depth diifference in model area, the two dimensional numerical model employing implicit scheme is used in this study. And as the model region is including both the shallow and deep water depths the tide generating force is introduced

to see what different effect it has on each sub-model area. The model region toward south and south-eastem boundaries is also extended to Ryukyu Islands for the model to utilize the measured data along the islands and to avoid possible error that might occur when using the edited tidal charts. Also the four straits surrounding the East Sea are set open to study possible contribution to the tidal regime in the East Sea.

2. A TWO-DIMENSIONAL TIDAL MODEL 2.1 Governing equations The governing equations are described under the spherical polar coordinate system to consider the curvature of the earth and the variation of Coriolis force with latitude and the variation of the generating forces with latitude as well as longitude. The usual depth averaged two-dimensional

32

equations of shallow water waves can be alternatively represented in terms of depth-integrated velocities as follows :

av gh -+ar R

where

a

-(q-(l+k2-hi)E)=

a@

x is longitude, Q latitude,

Vtana

-2osinQURh

ka V

d h’

W

(2)

h water depth, q surface elevation. R radius of the earth. kh=g/C’.

g gravity acceleration, C=Chezy coefficient with uniform value or (h+q)”’/n with n constant. The term ( I + k . . - h i ) E in (1) and (2) represents tide generating forces. The modification factor by earth tide, (1 +k>-h;) is set to 0.69 following Pingree and Griffths (1987). and k2 and h:’ are Love numbers. And E = H e X c o s 2 8 X c o s ( o r - P L + ~ ) ,

in which He, a P. L and S denote the amplitude of

the equilibrium tide for the M,, angular frequency, species number (2 for MI component), longitude. latitude and standard time. respectively. In this study the convection terms are ignored as the main purpose of this study is mainly to investigate the amphidromic system. The bottom friction term has been denoted as quadratic friction form following the general applicability of the quadratic friction law for semi-diurnal tide modelling(Pingree and Griffths, 1987). 2.2 Numerical method and computation i) Numerical method. The governing equations were solved using the linearized Abbott (1979) type scheme with four component equation of 2-dimensional equations which is an implicit scheme and has second order accuracy in space and time. In the computation the two equations that is. x-momentum and continuity equation are alogrithmically connected in a first stage in such a way as to advance U directly from n6t to ( n

+ 1)6t while advancing q only

from n6r to ( n + 1/2)6r. The

second stage connects y-momentum equation and continuity equation algorithmically to advance V directly from n6t to ( n + 1 ) 6 r while advancing q from (n+1/2)6r to ( n + 1 ) 6 t . In each stage the

equations are solved using the double sweep method. The more detailed description has been given by Abbott (1979). Due to the large differences of depth of each sea in the model area and the existence of a deep water depth in the East Sea the proper choice of time interval may be important and from this point this scheme is efficient in choosing the time interval (6r) than the explicit scheme which should satisfy the Courant-Friedrichs-Lewy condition. Letting mean depths of the East Sea and the Yellow Sea 2,500 m and 500 m. repectively, the Courant number ratio is about 9 for the chosen time interval and grid size. The characteristics of the scheme used in the region with basins of greatly different mean depths should show nearly same accuracy for the large different Courant number. The characteristics of Abbott type scheme has been thoroughly investigated by Sobey (19701, Abbott (1979) and Abbott et al. (1981).In Figure

33 4 the phase portrait of the scheme used in this study is shown Following Abbott (1979).

The amplification factor is 1 for all wave number and the celerity ratio is not unity. which is functions of the direction of propagation relative to grid line. This figure shows phase portraits in phase shifts as function of numbers of points per wave length when propagation occurs at angle

h=45" to the grid lines. x or y . Also it is shown that in order to use the capabilities of the scheme for working at the high Courant number, something in the range of 20-100 points per wave length is required for the main wave component. As shown in Figures 3(b) and 5 (grid system). the typical grid points per wave length in the Pohai Bay and the East Sea are about 80 and several hundreds, respectively. And the chosen time interval (6t) is about 175 s and the mean interval of grid space between h and U (or V) is about 7,000 m. The mean depth in the Pohai Bay and the East Sea is about 20 m and 2.500 m, respectively, so that the Courant number is 0.5 and 5.5, respectively. It is therefore expected that for such ranges of Courant number the scheme used in this study is expected to be sufficient enough to reproduce the accurate amplitude and phase of M , tide from the theoretical point of view. ii) compirfurion. The initial and boundary conditions required for the solution of equations (1) to (3) are described in what follows. At r=O. U(x. 8. t ) . V ( x , 8, I ) and

I& 8. f ) = O

are specified

at every points. At a land boundary, the component of the flow normal to the boundary is zero. The open boundaries consist of five sections ; section along the Ryukyu Islands, section along the Taiwan Strait and three sections in the East Sea, that is, the Tsugaru Strait. the Soya Strait and the Mamiya Strait. Along an open boundary, elevation is specified as a function of time. For the

M2 tidal constituent the open boundary values are given by q(x. 8, t ) = ~ ~ . & . e) c o d ~ ~ ~ 0)) g ~ d ~ ,

(4)

where gM2 is local phase lag referred to 135" E and HM2,OK are amplitude and angular frequency of M, component, respectively. The values of tidal constants (HM2and gM2) are directly obtained from measured data not from the tidal chart. By using the measured data as bourdary condition with some interpolatin we can avoid the possible error that could occur when we use tidal chart. But in this study the inner disturbances due to tide-generating force were not permitted to radiate through the open boundaries under the assumption that the effect is limitted around the boundary. The grid resolution is 1/8"(26x) in longitude and 1/6"(260) in latitude. the time interval is 174.6647s with which one M , period corresponds to 256 time steps. To check the convergence of numerical solution 13 cycle run was made. The grid system and the locations of 12 checking points for elevation and five strait sections for the compution of tidal volume fluxes are indicated in Figure 5. 3. RESULTS AND ANALYSIS

A series of numerical experiments were carried out to simulate the M , tidal distribution in order to reach reasonable agreemet with tidal charts newly edited from additional observations (Nishida, 1980). The various bottom friction parameters were used, that is, with the spatially uniform Chezy value or nonuniform Chezy value with respect to water depth at each grid point. At this stage the overall reproducation of M 2tidal distribution with uniform frictional factor is shown due to its relatively more realistic reproduction of tidal charts than using a nonuniform parameter. In Figure 6 the computed elevations at 12 grid points are shown to confirm the fact that

34

Figure. 4. Phase portrait for propagation along direction at 45" to grid lines for linearized Abbott type scheme(Abbott( 1979)).

Strait

Amplitude

Korea

5.34

Tsugaru

1.93

Soya

0.54

Tartary

0.09

Taiwan

2.64

Table. 2. Inflow and Outflow of volume transport in the model region during one M, cycle. (unit : x 10"hZ'h) Strait Korea

Inflow

outflow

Eastern channel

3.714

3.677

Western channel

3.922

3.936

Sum

7.633

7.608

2.875

2.638

Tsugaru Soya

0.891

0.684

TartaIy

0.140

0.122

Sum in the East Sea boundary

1 1.539

11.052

Taiwan

3.828

3.687

35

Figure. 5. The location of the checking points( 1-12) for numerical soultion’s convergence and sections

for computing the tidal volume fluxes and volume transport through five straits.

36 the computed results reach stationary state around 7th cycle after the start of calculation from the initial state and to see the different amplifications at various sub-sea points. The short period oscillations superposed on the major oscillation are noted at points I and 2 which are located in continental slope. These are thought to be due to the generation of short period component associated with the abrupt change of bottom topography rather than numerical instability. The large tidal amplitude at the point 3 of Kybnggi Bay occurs as expected. At the points 4. 5, 6. 7 located Seohan, Pohai, Liatung bays relatively different converged results are shown especially with the strong nonlinear characteristic occurring in grid point 6. The elevations at points 9, 10. 11. 12 in the East Sea are shown to converge around 7th cycle while in point 11 the nonlinear characteristic is shown. The results were analysed for M, constituent by the Fourier analysis. 3.1 Comparison of model-generated tidal charis with aisring tidal cham Figure 7(a) and 7(b) show the model-generated co-amplitude and co-tidal charts of the M. tide. The results have been simulated with inclusion of tide generating forces with friction parameter

kh=0.0025. The general patterns of the computed charts are in good agreement with the existing ones in Figure 3(a) and 3(b). The four amphidromic points in the Yellow Sea and the East China Sea as well as two amphidromic points in the East Sea have been reasonably reproduced. Also 8 hour co-tidal line as well as co-range line located nearly along the continental slope has been fairly well simulated without any serious oscillation. The large tidal ranges in the Kybnggi Bay west coast of Korea are also reasonably generated. The low tidal ranges in the almost entire area of the Eas Sea due to the deep water depth are also apparent in the model results. But some deviations of co-tidal hour in the bays of Liatung and Pohai are shown as well as a little northward movement of co-amplitude line of 10 cm in the Tartary Bay of the East Sea. For the locations of amphidromic points occurring in the bays of Liatung and Pohai the computed ones move more westward or northwestward, compared with those of Nishida (1980) or Ogura (1933), but appearing more consistent with the results by Fang (1986). Choi’s fine grid model study (1988) also show that the amphidromic point in Liatung Bay appears in deamphidromic pattern while the amphidromic point in Pohai Bay moves northwestward similarly in this study. From these observational and numerical results the more thorough analysis based upon measurement seems to be required to confirm where and how the amphidromic system exists. A little severe dissipation of tidal amplitude at Tartary Bay appearing as a little northward movement of co-amplitude line of 10 cm and slightly westward movement of amphidromic point in Korea Strait are thought to be partly due to the relatively high bottom dissipation by using uniform friction parameter over the whole model area irrespective of the difference in water depth between the seas. In fact, when considering the bottom friction factor as variable of depth (the form of (depth)”‘), the friction factor ratio is 2 to 1 for the depths 50 m and 250 m for the Yellow Sea and the East Sea, respectively. In fact, the response for the less friction factor is also studied and, as expected, such effects appear though not shown in this study. The ocean current, a branch of Kuroshio, in the East Sea is also expected to play a role to the movement of amphidromic system located at the Korea Strait through wave current interaction. 3.2 The effect of tide generating forces

37

- -2

0.911 1

If

0-

-1.11;

0-

,

,

,'

,

,

' ,

, '

,

0 1 2 3 4 5 6 7

!

, ' , : , ' 8 9 1 0 1 1 1213

',

-1.li1, , , , , ., ,. , , , 0 1 2 3 4 5 6 7 8 9 1 I

,

,

,_I

111213

3-

3.0-

0-

-3.0-

0-

0-

0-

-

'

'

-

0--

-0.5]',

, '

,

:

-0.5';.

' ,

, ' ,

' I

I

'

4

I ' ' -

Figure.6.Temporal variation of the computed surface elevation in meters at 12 points for 13 cycles

of M> tiddtime unit in tidal period).

I

38

Figure. 7. Model-generated tidal charts with tide generating force included. (a) co-amplitude chart (amplitude : em).

39

Figure.7(b). Co-tidal chart referred to 135" E.

40

Figure 8(a) and 8(b) show the computed co-tidal co-amplitude lines without tide generating forces with an uniform friction parameter (kb=0.0025). The difference of co-tidal lines between Figure 7(b) and 8(b) is obvious in the results of the East Sea where without tide-generating force the amphidromic point in the Tartary Bay disappears while the amphidromic point just above the Korea Strait moves unrealistically toward northwest direction compared with the results including tide generating force and with existing tidal chart (Nishida (1980)). In case of co-amplitude lines (Figure 7(a) and S(a)) any significant differences in the Yellow Sea and the East China Sea are not noticeable probably due to minor changes while some differences at the Tartary Bay of the East Sea are shown. To see the effect of tide generating forces upon the tidal dynamics of the modelled area the pure contribution of independent tide due to tide generating force has been computed by analysis of the differences of elevations with tide generating force and without it have been compared. Figure 9(a) and (b) show the model generated co-amplitude and co-tidal lines due to the independent tide by tide generating force. As shown in Figure 9(b), the independent tide in the Yellow Sea

shows the pattern that the equilibrium tidal wave propagates from south and south-east direction to nothern direction. The amphidromic points by independent tide also appear at the nearly same

locations in the existing tidal chart except the amphidromic point located in the Gulf of Pohai. In the East Sea the independent tide propagates from central area both to north-eastern direction and south-westem directions. Figure 9(a) shows the computed co-amplitude lines by the independent tide in the model area. In the Yellow Sea and the East China Sea the large amplitudes up to 11 cm due to the tide generating force appear in the west coast of Korea as in the co-oscillating tide, but in the Tartary Bay of the East Sea the amplitudes due to the independent tide increase gradually and reach up to about 25 cm, which is almost the same order of magnitude of tide flowing into the East Sea through Korea Strait, as also discussed by Odamaki (1989b). The independent tide in the distribution of amplitudes accounts for about 1-4% in the Yellow Sea and the East China Sea while in the East Sea it explains more than 10% over the nearly whole area. and in the Tartary Bay up to about several tens of percent. Through these analysis the independent tide has been shown to be very important over the whole area in the East Sea. Especially it is worth noting that the amphidromic point located in the Tartary Bay does not appear without the independent tide even though the bay head is open. Therefore it could be inferred that the reflected Kelvin wave at the head area of the Tartary Bay is more reinforced by tide generating force than the reflected wave formed only by the interaction

of incoming wave through Korea Strait with the waves imposed on the open boundaries of the Mamiya Strait.

3.3 Tidal volume fluxes through jive straits The tidal volume fluxes through five straits (see Figure 5 for the location) in modelled region have been computed and the time variations of tidal volume fluxes through each strait are plotted in Figure 10, which shows that the volume fluxes passing through the Taiwan, the Korea, the Mamiya straits reach stationary state during 7 or 8th cycle while in the Tsugaru and the Soya straits they reach stationary state beyond loth cycle. The amplitude of tidal volume fluxes and the volume transports during one M, tidal cycle have been listed in Table 1 and Table 2. The amplitude of

41

Figure. 8. Model-generated tidal charts with tide generating force excluded. (a) co-amplitude chart (amplitude : cm).

42

i.iz [I

43

Figure. 9. The effect of tide-generating force. (a) co-amplitude chart (amplitude cm).

44

Figure.9Cb). Co-tidal chart referred to 13.5" E.

45

- '"1

T

1.7-

0-

3.5

0

-3.5 5.9

0

-6.1

0

1

2

3

4

5

6

7

8

9

1 1 1 1 2 1 3

4.7

0

-3.3

I

Figure. 10. Convergence of tidal volume fluxes in Sv throutgh five straits (Taiwan, Korea, Tsugaru, Soya, Tartary Straits in order of bottom to top. Time unit in tidal period).

46

tidal volume flux at the Taiwan Strait is about 2.64 Sv (=lo6

rnys) and the volume transport

during half period of M , is about 38 krn], but the ratio of each strait to total volume flux are similar to those estimated by Odamaki. The inflow volume transport into the East Sea for about half M , tidal cycle through the Korea, the Mamiya straits are 76.3 km’, 1.4 krn’ while those through the Soya and the Tsugaru straits are 8.9 krn’and 28.8 km’. So it could be said that the co-oscillating tide through the Korea Strait is dominant for the M.. tide, as noted by Defant (1961). The contribution by the Tsugaru Strait also seems to be influential considering the amount of volume transport. The amounts of inflow and outflow through the Korea Strait have been computed to be almost the same as each other. In contrast to the Taiwan Strait the volume transport through the Korea Strait is about

5 times as large as that computed by Defant(l961). One of the reason of this difference is expected to be partly due to the possible difference of the location of section where the tidal amplitudes change rapidly from place to place for computing volume transport. 4. CONCLUSION AND DISCUSSIONS

Through this study it has been shown that the tidal phenomena in the entire seas surrounding Korea Peninsular could be investigated under one oscillating system of the whole seas using a two-dimensional depth averaged numerical model. In this study southern part of open boundaries were extended to Ryukyu Islands for the boundary value to be obtained from the observed data. Due to the large mean depth difference between each sea the tide generating force has been included in the computation. To overcome the problem related to time interval limitation by CFL condition in case of using the explicit method the implicit scheme has been successfully employed in this study. The computed

M2

co-amplitude lines in the Yellow Sea and the East China Sea are in good

agreement with those of existing tidal charts. The four amphidromic points known to exist in the Yellow Sea and the East China Sea as well as two amphidromic points in the East Sea are successfully reproduced with inclusion of tide generating forces. Even along the continental slope where rapid depth changes occur the computed results have been founded to be in good agreement with the observed results. The results of computation also show that tide generating forces play greatly different roles in the each sub-sea of the model region. The tidal amplitude due to the tide generating force in the Yellow Sea and East China Sea explains several percent of the real observed amplitude while the tidal regime in the East Sea is greatly influenced by the tide generating force. It might be inferred that the reflected waves at the head area of the Tartary Bay is rather reinforced by the tide generating force. In the Tartary Bay of the East Sea the experiment letting the head as open boundaries was found to give more reasonable results for the proper simulation of the amphidromic system than letting the head closed, even though the tidal volume fluxes through the Mamiya Strait is very small compared with those passing through other straits, as shown in the section 3.3. In case a bay with its head opening is too narrow and too deep for Poincare modes to propagate semi-diurnal energy, Hendershott and Speranza (1971) showed that the co-oscillating tide must be primarily a superposition of oppositely travelling waves and that an asymmetry of amphidromic system in such a bay results from partial absorption of incident power flux at the bay head. In case of the

41 Tartary Bay, in contrast to the Yellow Sea investigated by Kang (1984). it has a step in depth along the bay from 1,000 m to 100 m and the amphidromic point is located on the continental slope, which step as well as bottom friction make it difficult to investigate the amphidromic system analytically. The critical period of the first Poincare wave, 7; is as follows (see Kang (1984) ).

where

g, H, rn, B are Coriolis parameter. gravity acceleration, depth, cross channel mode number

of Poincare wave and bay width, respectively. For depth of 100 m, 500 m. 1,000 rn and width of 150 km the critical periods are 2.6, 1.2, 0.8 hr, respectively. So, even though considering the depth slope, it could be said that the co-oscillanting tides in the Tartary Bay consist of Kelvin waves. The possible asymmetry in the geometry condition like the Tartary Bay is first expected to occur due to a partial absorption of incident power flux at the bay head (Hendershott and Speranza (1971)) and the energy dissipation by bottom friction (Rienecker and Teubner (1980)). To check

the role of head opening another experiment has been made with the head closed. The computed results even with tide generating force showed that amphidromic points within realistic range of bottom friction factor. seems to move inside the land. When considering that the head opening has the effect of letting amphidromic point move westward in the nothern hemisphere it is thought that the driving force at the head of the Tartary Bay plays some role in the amphidromic system as well as the tide generating force, as discussed in section 3.2. Forth- bottom friction the quadratic form was found to give general agreements with observed ones even th,

some evidence of a little excessive dissipation appears in the results of the East

Sea area, wh,,,,

was also shown through simple analysis. Further studies for more accurate

reproduction of tidal amplitude and phase seem to be needed with relation to bottom friction. In this study the nonlinear phenomena in the modelled area have not been investigated as well as not including the study on the tidal regime of diurnal component. There topics will be considered in the next study associated with a series of studies for investigating the tidal regime in these seas. ACKNOWLEDGEMENTS The authors would like to thank Drs. H. J. Lie (Korea Ocean Research and Development Inst.), Y. H. Seung (Inha Univ.) and Y. Q. Kang (Pusan Fishery Univ.) for their reading manuscripts and valuable comments. We appreciate Dr. B. H. Choi for review and helpful comments. Thanks are also due to Mr. W. D. Paik for drawing figures. The work was partially funded by the Ministry of Science and Technology and the computations by Cray2s was supported by System Engineering Research Institute, KIST.

REFERENCES Abbott, M.B.. 1979. Computational Hydraulics : Element of the Theory of Free-Surface Flows, Pitman, London.

48 Abbott, M.B., A. McCoWan. and 1. R. Warren, 1981. Numerical modelling of free-surface flows and coastal waters, Ed. H.B. Fisher, Academic Press, New York. An, H.S. 1977. A numerical experiment of the M Ztide in the Yellow Sea. Jour. of the Oceanogra. SOC.of Japan, 33 : 103-110. Defant, A. 1961. Physical Oceanography, Vol. 2. Choi, B.H. 1980. A tidal model of the Yellow Sea and the Eastern China Sea. KORDI Rep. 8002, Korea Ocean Research and Development Institute, Seoul 72 pp. Choi. B.H. 1988. A fine grid two-dimensional M, tidal model of the East China Sea. Jour. of Korean Assoc. of Hydrol. Sc., 21(2) : 183-192. Fang, G. 1986. Tides and tidal currents in the marginal seas adjacent to China. Presented at tenth international symposium on earth tide. Madrid. Hendershott, M.C. and A. Speranza. 1971. Co-oscillating tides in long, narrow bays ; the Taylor problem revisited. Deep-sea Res., 18 : 959-980. Kang, Y.Q. 1984. An analytical model of tidal waves in the Yellow Sea. Jour. Mar. Res., 42 : 473483. Nishida, H. 1980. Improved tidal charts for the western part of the north Pacific Ocean, Rep. of Hydrogra. Res., No. 15. Odamaki, M. 1989a. Tides and tidal currents in the Tsusima Strait. Jour. of the Oceanogra. SOC. of Japan, 45 : 65-82. Odamaki, M. 1989b. Co-oscillating and independent tides of the Japan Sea. Jour. of the Oceanogra. SOC.of Japan, 45 : 217-232. Ogura. S. 1933. The tides in the Seas adjacent to Japan., Hydrogr. Bull. Dep. Imp. Jap. Navy. 7. 1-189. Pingree, R.D. and D.K. Griffiths. 1987. Tidal friction for semidiurnal tides. Continental Shelf Res., 7(10)

1181-1209.

Rienecker, M.M. and M.D. Teubner. 1980. A note on frictional effects in Taylor's problem. Jour. Mar. Res., 38 : 183-191. Sobey, RJ. 1970. Finite difference schemes compared for wave-deformation characteristics ets., Tech. Memor. No. 32, US Army Corps of Engineers, Coastal Eng. Res. Center, Washington, D.C.

49

NONLINEAR ROSSBY WAVES IN THE INERTIAL BOUNDARY CURRENT THEIR POSSIBLE RELATION TO THE VARIABILITY OF THE KUROSHIO

AND

Qinyu Liu and Zenghao Qin Institute of Physical Oceanography, Ocean University of Qingdao, Qingdao 266003 (China)

ABSTRACT Based o n a qualitative theory of ordinary differential equations, the stability characteristics of nonlinear barotropic Rossby waves propagat ing in the western inertial boundary current of the ocean is discussed with geostrophic momentum approximation. A criterion is obtained about the occurrence of the bifurcation in the inertial boundry current, which may be applied to the variability of the Kuroshio. INTRODUCTION In recent decades, remarkable advances and achievements have been made in the study of the large scale ocean waves and their instabilities,especially of the linear barotropic and baroclinic Rossby waves. Some phenomena closely related to the large scale ocean waves have been clarified. Pedlosky( 1979) has systemati-

cally studied the linear Rossby wave instability under various circumstances. Kang et a1 ( 1 9 8 2 ) have studied it in a non-zonal

.

shearing ocean current. However, very few authors have dealt with the nonlinear Rossby waves and their instabilities in the oceans. This is because not only the nonlinear problem is mathematically difficult to deal with but also the nonlinear terms can be neglected in most cases in large-scale ocean motions. Within the inertial boundary current zone in the western boundary of the ocean, e.g. ,the Kuroshio and Gulf Stream, the current velocity is usually one order of magnitude greater than that in other regions, and in some processes nonlinear terms in the momentum equations can not be omitted. Because the western boundary region is mostly over the continental shelf and the sea-floor topography is quite complicated , the interactions between the currents and those between waves can not lected. The inertial boundary current results from the

be negcombin-

ation of the westward-propagating Rossby waves obstructed by the western boundary and the beta-ef f ect ( Pedlosky ,1979). Hence, the study of the nonlinear Rossby waves in this region is helpful to

50

understand the dynamical mechanism of inertial bouridary currents such as the Kuroshio and Gulf Stream. Liu et a l . (1983, 1987) have studied nonlinear waves both in the atmosphere and ocean using stability theory of ordinary differential equations. However, their studies ignore the bottom topography, basic current and frictional effects. The present authors ( Liu and Qin,1990) have studied the topographically trapped waves and the nonlinear waves in the tropical ocean and atmosphere, which are, however, not applicable for the nonlinear waves in the inertial boundary currents. An attempt is made to study nonlinear Rossby waves in inertial boundary current and to explain to some degree Kuroshio variability.

the the

2 NONLINEAR BAROTROPIC ROSSBY WAVES

In order to generalize the discussion of the inertial boundary current, the local Cartesian coordinates, x, y , z are chosen : the y-axis coincides with the straight coastline and x-axis lies perpendicular to it with positive direction shore and z-axis

pointing

goes upward pointing to the local

off-

zenith. The

7

( x , y , t) is the free surface elevation above the equilibrium level. Assuming the bottom friction to be linear for the velocity component with bottom resistance coefficient r and denoting the satisfying constant basic inertial boundary current by V plane

oxy coincides with undisturbed sea surface and

-

-V = ( g / f )

2/ax > 0

, we get the nonlinear equations for barotropic

motion of an inviscid, homogeneous and incom-pressible overlying the uneven sea-bed z = -h(x,y) :

water

where u and v are the depth-averaged components of the horizontal velocity in the x- and y-directions, respectively.

51 T1,.

p.~r~iiieCar f-

Cr,i-icJl i

is

s~lssnmed to

f = f,

be

+&x

+

By

y,

is the Coriolis parameter and (3 pSin+ , p,=pcos+ 9 f, is its northward gradient at a prescribed latitute and+is the counterclockwise angle of the x-axis from the east. The other symbols in E q s (1)-(3) are usual. The geostrophic momentum approximation (Hoskins,l375) is made to filter out the inertia-gravity waves which is inherent in the model and simultaneously unchdnged, i e

to

retain

the

nonlinear

effect

. .

s o that E q s . (1) vorticity equation

and ( 2 ) can

The term ( a v h x -au/ay)(auI;,x

be

combined into the following

+ a v / a y ) is

negligibly

small in

comparison with term f, ( I J u h x +2v/ay), so that it is omitted.

For Rossby wave of wave numbers k and 1 directions and angular frequency &). , solutions ( 4 ) are of form

in s- and of Eqs. ( 3 )

-

yand

where 8 =kx+ly-cJt , , and the parameter V is assumed to be independent on8. Eqs. ( 3 ) and ( 4 ) are then transformed to

+

1v

+

1G) kZ

I

kU + 1V) = 0

+

1’ ) H ” ’ +

(kk -6,l)H’

+

r(k”+ l*)H”

52 (-0

+

+ 1T)H' + [(kU + lV)(H

Here,

the

superscript

ti

+ r ) ] ='

denotes

.

the

respect to 8 Integrating Eq. (7) and constant to be zero, we have

(kU + 1V) = H(O- lF)/(H

+

h

0

(7)

differentiation putting

the

with

integral

+?

Inserting Eq. (8) into ( G ) results in

where

The zero solution or equilibrium points of the ordinary differential equation (9) are H=H'=H''=O, which is equivalent to the average inertial boundary current with the corresponding uneven sea surface? and without perturbed u , v and 1. F(H, H', H") is an analytical function if the conditions W 17, and H f -(h h o l d good. Eq.(9) can be replaced by its equivalent set of equations:

*

dH/dS

+T)

= H'

dH'/d8 = H"

dH/ds

= P(H) H"

+ Q(H) H'

Consequently, there are some equilibrium points (H, , O , O ) in the phase plane in which H, depends upon the values of P and Q. The zero solution must be the equilibrium point no matter what the

53 values of P and Q will be. Expanding Eq. (9) in Taylor's series near the point ( 0 , 0, 0) leads to

+

H' ' ' = Q , H ' + Q z H "

= ?-/(a-

,Q,

- f:/[g(k%

+

lT),

Q3HH'

+ Q+HH"

+

......

Q , = ($,k - p * l ) / [ ( U - lv)(kr

'1 )(h

+T

(11))

+7)]

+ l+)(h

$1

Eq.(ll) is a nonlinear ordinary differential equation, which is difficult to solve. If the nonlinear terms higher than the third order and the friction terms are neglected, E q . (11) can be replaced by the equation : = Q,H'

H"'

+ Q3 H H '

(12)

which is a KdV equation provided that the horizontal variations of ti and areyneglected.Its solution is an elliptical cosinoidal function (Liu, 1983) and its frequency is related not only to the wave length, ocean depth, basic current orientation, but also to the wave amplitude. the

and

coast

When the friction is taken into account, the derivation of solution would be quite difficult. Now let's use the

stability theory of the ordinary differential equation explore the characteristics of the solution of Eq. (9). 3 STABILITY

to

AND BIFUREATION

Eq.(9) can be approximated by a nonlinear equation

+

+ Q,HH'

+ Q4HH"

H"'

= Q,H'

The

stability in the vicinity of the zero solution of Eq. (13)

QaH"

(13)

54

within the phase space (H,

H I ,

H")

can

be studied

from

the

stability theory of the ordinary differential equation. First of all, let's discuss the stability of the linearized Eq. ( 1 3 )

= Q,H'

H"'

+

Q,H"

(14)

Its corresponding characteristic equation is

,(-

Ah

+ B

where A = Q,, B =

= O

-Q, ,

Set:

From the characteristic root of Eq.(15) and the expressions for Q1 and Q2, the stability behaviour i n the vicinity of the zero solution of Eq (14) can be determined as shown in Table 1. Table 1 Stability behaviour of equilibrium point (zero solution) stability behaviour of equilibium point r M M>1/4 unstable focal point M=1/4 unstable degenerate nodal point O < M < 1 / 4 unstable nodal point r> M>1/4 stable focal point M=1/4 stable degenerate nodal point O < M < 1 / 4 stable nodal point unstable saddle point Mh$ centre r= M 10m/s). a n t i c y c l o n e 3 d i d n o t remain c o n s t a n t .

The

r a d a r image t o n e i n s i d e t h e

B r i g h t n e s s is lower i n t h e n o r t h e r n

(warmer) p a r t of t h e eddy. B r i g h t n e s s is enhanced i n t h e s o u t h e r n p a r t where t h e v o r t e x movement c e n t r e

K a s

l i k e l y l o c a t e d and t h e narrow stream of Oyashio

waters was observed. However, as a whole, on t h e r a d a r image w a r m eddy waters have a l i g h t t o n e and t h e c o l d s u b a r c t i c o n e s a d a r k one. C o r r e l a t i o n between t h e SST and t h e b a c k s c a t t e r e d r a d a r s i g n a l l e v e l Oo is v i s i b l e a t periphery of

A3.

Thus,

t h e narrow stream

(about

10 miles) o f

Oyashio waters 8 , involved i n t h e v o r t e x movement and i t s b r a n c h e s 9 a p p e a r d a r k e r on t h e r a d a r

The w a r m area 10 east of

image t h a n t h e eddy.

these

branches is as l i g h t as t h e eddy. The d a r k s t r e a k 11 d e n o t i n g t h e minimum roughness zone a d j o i n s t h e eddy from t h e n o r t h . The r a t h e r w a r m zone 1 2 w i t h l i g h t tone

extends

farther

t o the

north.

The c o l d e s t Oyashio waters are

l o c a t e d h e r e . On t h e I R image t h e zone 12 s t a n d s o u t i n d i s t i n c t l y , which may be e x p l a i n e d by t h e ueak thermal c o n t r a s t . The s p i r a l v o r t e x f o r m a t i o n 5 developed by waters of t h e Tsugaru C u r r e n t 13

is t r a c e d n o r t h of

t h e a n t i c > - c l o n i c eddy 3 .

The width of

changes from about 5 miles i n i t s t a i l t o 25-30

miles i n t h e r o t a t i o n a l

c e n t r e . The c e n t r a l p a r t of t h e s p i r a l 5 with t = 9-10°C h i g h e r v a l u e s of

Qo

on t h e r a d a r image.

t h e warm s p i r a l

is c h a r a c t e r i z e d b y

Water t e m p e r a t u r e and r a d a r

image

b r i g h t n e s s are d e c r e a s e d i n t h e narrow p o r t i o n of t h e s p i r a l . The b o u n d a r i e s of t h e t h e r m a l and r a d a r c o n t r a s t s c o r r e l a t e well.

The Oyashio c o l d water zone with a low Oo l e v e l s t a n d s o u t east o f t h e Tsugaru S t r a i t .

The warm area 6

formed by

the

anticyclonic

eddy of

the

91

Kuroshio is not visible on the radar image, because it is located at the far edge of the swath where only the significant changes of

(ro

can be registered

(Mitnik and Victorov, 1990 1 . In spring 1987 the anticyclone A3 and other heterogeneities in the SST field were registered on several radar images. During the radar survey on April 1 (Fig. 5(a)) a baric field with low gradients was located above eddy 3. Wind speed measured almost simultaneously with satellite observations was 2-3 m/s in the vicinty of the Honshu, and south of the eddy at 0600 Z. The weak wind zones 1 are reflected by a dark gray tone. A warm front north of Hokkaido is seen in the surface weather map. A strong (15 m/s) westerly was noted in the frontal rear and resulted in

U0

increase (area 2 ) . On the I R image obtained h;-

NOAA-10 satellite 3.5 hours after radar one the larger part of the eddy was

shielded by clouds, which makes it difficult to get detailed correlation.The position and form of the contrast zone 5 coincide with the bank of cumuli clouds and are distinguished well on the I R image. Alternate light and dark belts about 16-18 km apart are noted in almost all areas of the anticyclone to the southwest of the boundary 5 . Variations in sea surface wind caused by lee waves (atmospheric internal gravity waves) resulting from interaction of the eastward air flow over mountains on Honshu are responsible for such modulation of the backscattered radar signal level (Mitnik and Victorov, 1990). These waves are expressed in the cloud field also. Quasiperiodic brightness variations are visible in the strong wind area 6 where the western part of the anticyclonic eddy is located (this eddy designated by 6 in Fig.J(a)). The eddy can be recognized on the I R image. However, its indication on the radar image is practically impossible due to strong wind, in spite of the fact that it is situated at the near-edge of the radar swath. During radar measurements on May 2 (Fig. 5(b)) wind speed did not excced I m/s in the greater part of the area under investigation, as in the previous two cases. The nearest I R image obtained two days before radar one permits to interpret the brigtness variations. I n the anticyclone area 3 the sea silrface is rougher than in surrounding colder waters, except for area 1 to the west, where the brightness increase is associated with 'the wind action. Compared to image on April 1 the eddy A3 was shifted by about 100 km to the north. This estimation is supported by I R images analysis. The arced line 2 fits the boundary of waters with t = 3-5OC (area 4 ) and 10-12'C (area 5). The higher radar contrast features the eastern boundary I; of the first meander of the Kuroshio. The tiuroshio waters 7 have temperatures of 19-20°C. It is 2 - 4 O C higher than in adjacent waters 8 and 9. The brightness

heterogeneities in area 7 where the SST variations from IR data are small can be attributed to variations in the wind speed and direction. In spring 1987 eight radar images east of Japan were obtained. The eddy A.3

92

Fig. 5. Cosmos-1766 SLR images of an area east of Japan. ( a ) At 0605 GMT on A p r i l 1 , 1987. ( b i A t 1330 GWT on May 2, 1987.

93

area is t r a c e d w i t h d i f f e r e n t d e g r e e of c o n t r a s t on s i x images. On two images d a t e d A p r i l 12 and 24 t h e eddy is n o t r e v e a l e d because o f s t r o n g wind (10-15 m/s a c c o r d i n g t o t h e s y n o p t i c maps).

2 . 2 . S p r i n g 1988 In

April-May

1988

fifteen

radar

images

of

areas

east

of

Japan

were

o b t a i n e d . Most o f them showed d e f i n i t e e l e m e n t s o f t h e f r o n t a l zone. However, those

elements

i n f l u e n c e of

were

on two

absent

images.

SST f i e l d was n o t v i s i b l e ,

passage w i t h v e r y s t r o n g winds.

In

one

case

because of

(April

l l ) , the

t h e a t m o s p h e r i c front,

In t h e o t h e r case (May 211, no t e m p e r a t u r e

c o n t r a s t s was n o t r e v e a l e d i n t h e roughness f i e l d , because of calm weather with winds r a n g i n g from 0 t o 3-5 m / s .

In May hydrometeorological measurements

were c a r r i e d o u t from a weather s h i p "Ocean"

estimate

to

the

SST,

air

t e m p e r a t u r e and wind speed v a r i a t i o n s i n t h e areas observed on r a d a r images. On t h e r a d a r images i n 1988 t h e o c e a n i c f r o n t s were observed n o t o n l y i n t h e r e g i o n o f t h e f i r s t Kuroshio meander,

b u t also t o t h e east o f i t , . Two

a n t i c y c l o n i c e d d i e s of t h e second branch o f t h e tiuroshio (147-151°) A5 and $16 s t a n d o u t s h a r p l y on I R image i n A p r i l 28 ( F i g . 6 ) . T h e i r t e m p e r a t u r e \-as

about 13-14 and 15-16OC, r e s p e c t i v e l y . The SST c o n t r a s t on t h e eddy b o u n d a r i e s w a s about 4OC.

waters.

Both e d d i e s had a high r a d a r c o n t r a s t a g a i n s t t h e f r o n t a l zone

I n t h e v i c i n t y of e d d i e s t h e t h e r m a l f r o n t s t r u c t u r e c o i n c i d e s w e l l

w i t h u0 f i e l d v a r i a t i o n s .

Because of c l o u d s , t h e r a d a r and IR images ( P i g . 6 ) can n o t make c l e a r whether

t h e s o u t h e r n boundary

of

t h e eddy .A6 i n

t h e SST f i e l d

f i t s the

c o n t r a s t zone 1 p o s i t i o n . The c l o u d i n e s s c o n t i n u e d t o c o v e r t h i s area in t h r s u c c e e d i n g days. However, t h e d i s t i n c t i v e o u t l i n e s of eddy 1 s o u t h e r n boundary and Kuroshio meander s t e p w i s e boundary 2 were observed c l e a r l y a g a i n on t h e r a d a r image on May 1. Over t h r e e last d a y s t h e e d d i e s ' kept,

while

their

contrast

to

the

background

c o n f i g u r a t i o n s were

substantially

decreased.

Apparently, t h e h y d r o l o g i c a l c h a r a c t e r i s t i c s of t h e e d d i e s c o u l d n o t change g r e a t l y f o r such a s h o r t p e r i o d , as shown by I R images. A p o s s i b l e r e a s o n of c o n t r a s t d e c r e a s e is t h e W weaking down t o 0-3 m / s

by f o r m a t i o n o f a v a s t

low-gradient b a r i c area by May 1 o v e r t h e r e g i o n c o n s i d e r e d . I n c o n t r a s t , wind s p e e d s of 2-7m/s

were observed on A p r i l 28.

The h i g h e s t r a d a r c o n t r a s t s on r a d a r images were a s s o c i a t e d w i t h i n t r u s i o n of

t h e Kuroshio

Kuroshio stretched

waters

waters t o t h e Oyashio area.

In the

(18-19OC)

A4

t o the north

occupied

(Fig.

the

eddy

area

middle of and

.April tlie

formed

6 ( a ) ) . Then a w a r m streamer o f

meandor

16-l'ieC

H ~ S

s e p a r a t e d from t h e eddy A 4 , approached t h e eddy A3 boundary and began t o enteii n t o i t s area c u r v i n g i n c l o c k w i s e d i r e c t i o n ( F i g . 7 ) .

Waters of t h e warm streamer 1 and t h e eddy A 4 2 were c h a r a c t e r i z e d by n v e r y high l e v e l b a c k s c a t t e r e d r a d a r s i g n a l s .

The lowest v a l u e s of

go

h'ere

94

Fig. 6. Satellite images of the Kuroshio frontal zone on A p r i l 28, 1988 ( a ) NOAA-9 IR image at 0553 GWT.

.

95

(b) Cosmos-1766 SLR image at 1525 GMT.

96

Fig. 7 . Satellite images of the huroshio first meander, ( a ) hOAA-10 I R image at 2 2 2 i GWT on Yaj 8. ( b ) Cosmos-1766 SLH image at 1435 GMT on Ma5 7 , 1988.

97 found i n t h e Oyashio waters 3 qoing around t h e e d d i e s .A3 and .A1 from t h e e a s t . The n o r t h e r n boundary of

t h e warm streamer was n o t

remarkable because

it

c o i n c i d e d with t h e wind a c t i o n area 4 w h i l e t h e eddy .44 boundary was i n high c o n t r a s t with t h e s u r r o u n d i n g waters. The h o r i z o n t a l d i s t r i b u t i o n of SST a c r o s s t h e eddy .4-4 eastern boundary on May 8

is shown i n Fig.

8.

The

SST changes a t

the

eddy b o u n d a r i e s

are

c h a r a c t e r i z e d by an a b r u p t jump by about l2OC o v e r a d i s t a n c e of o n l y 2-3 lcm. That area c o r r e s p o n d s t o a zone of s h a r p b r i g h t n e s s c o n t r a s t on t h e r a d a r image. A s h a r p f r o n t on t h e edge of warm streamer occupied eddy A3 remained in t h e succeeding d a y s and was v i s u a l l y observed from t h e weather s h i p " Ocean" on May 17-22. During 5 days of low windy weather d i f f e r e n c e s i n t h e s u r f a c e s t a t e between l i g h t e r Oyashio and d a r k e r waters of warm streamer were marked well from t h e s h i p w i t h i n a d i s t a n c e of 1-3 miles.

16 14

12 10 OC

8

6 4

0

20

40

60

80

km

Fig. 8 . Sea s u r f a c e t e m p e r a t u r e ( s o l i d l i n e )

and a i r t e m p e r a t u r e (clashed l i n e ) distributions across the eastern edge of warm eddy A 4 o b t a i n e d by weather s h i p "Ocean" a t 10-14 GMT on May 8. Roughness

differences

of

the

Kuroshio

and

Oyashio

surface

waters

are

confirmed by v i s i b l e images i n s u n g l i t t e r a r e a . One o f such images o b t a i n e d on May 2 is given t o g e t h e r w i t h synchronous L R image i n Fig. 9. T h e r e g i o n of t h e f i r s t Kuroshio meander w i t h e d d i e s A3 and A 4

is i n sun p l i t t e r a r e a .

Higher b r i g h t n e s s of c o l d waters 1 on t h e v i s i b l e image is e x p l a i n e d by t h e f a c t t h a t t h e sea s u r f a c e is smoother. Agreement between SST ( I R image) and roughness ( v i s i b l e image) f i e l d s is high enough:

in the brightness f i e l d th?

SST f i n e s t r u c t u r e i n warm water i n t r u s i o n 2 and areas s u r r o u n d i n g t h e f r o n t a l

zone are o u t l i n e d .

86

99 3 . DISCUSSION

Analysis of radar and IR images shows that the spectrum of ripples is correlated with the SST. The following problems arise. By which mechanisms is this relation accomplished and which hydrometeorological and other procesxes work for the indication of temperature inhomogeneities by radar technique? The influence of the SST variations on spectral density of capillarL-gravity waves may depend on change of the sea water characteristics and stratification of the atmosphere in the air-sea boundary layer caused by these variations. Two parameters, the surface tension coefficient a and a kinematic viscosity U , which are functions of temperature, govern the spectrum of short surface waves of scales of centimeters. The dependence of 0' on the SST obtained from processing of scatterometer data by Seasat satellite was explained by decrease of a with temperature increase (Woiceshyn et al., 1986). However, the change of a did not exceed 1.6% at a temperature difference of 68OC, which brings about no marked change in the spectrum for waves of h=2-4cm. The dependence of the kinematic viscosity on water temperature is much stronger: V = 1.84 cSt at t = O°C, = 1.36 cSt at t = 10°C and u = 1.56 cSt IJ

at t = 2OoC (at a salinity of 35 ppt). As a result, the increase of water temperature leads to decreasing the minimum wind speed Wo(A) at which waves of wavelength A begin to form. Wo values for edges of the radar swath are different since wavelength A , for which resonance scattering condition holds true, depends on incidence angle 8. At the near edge where 8 21° and A = 4.1 cm, Wo is 2 . 7 m/s at t = O°C, 2.4 m/s at t = 10°C and 2.1 m/s at t = 20°C. At the far edge (8 = 46O, A = 2.1 cm) W o decreases from 3 . 8 m/s to 2.9 m / s with variation of t from O°C to 20°C (Donelan and Pierson, 1987). For example, the dark tone of areas 2 (Fig.J(b)) and 1 (Fig.5(a)) possibly results from W < No. If the wind speed fluctuates, then the SST variation must affect the value of v0 for W > 4-5 m/s also. Actually, if there is an abrupt change in local qusts,patches of enhanced intense of ripples arise. The ripple amplitudes will 2 depend on the SST because the damping decrement is given by 7(t)=8n2U(t)/A and the wave amplitude is damped with time t in proportion to exp(-yt). At 5 =

10 s, the wave amplitude with wavelength A = 3 cm (resonance scattering condition for this wave is fulfilled at 0 = 30°) at t=lO°C will be about 1.5 times larger than that at t =O°C. The above estimation refers to the pure sea water without surfactant film. However, in the Kuroshio confluence zone under consideration, water masses are different not only in temperature, but also in chemical and biological properties. In particular, the plankton concentration may be one order hisher in cold Oyashio waters than in warm frontal waters and in the Kuroshio saters (Kun, 1969). The difference in plankton concentration is also revealed bv remote sensing, as shown by satellite images obtained by multichannel scanning

100 device from Meteor (Ginsburg and Fedorov, 1986) and

CZCS

and Emery, 1988). The surface-active films are a product of

the vital

from Piimbus-7 (l'liomas functions and

decomposition of plankters. These films damp ef fectivly short capillarygravity waves (Ermakov, 1987; Ermakov et al., 1987). Damping mechanism of films is not thoroughly understood. In the open ocean, three possible processes may come into play: change of energy inflow from wind, change of nonlinear interactions between waves, and direct influence of the surface f i l m - damping. The damping of waves may be associated with decrease of surface tension of water. Then, when the surfactant concentration exceeds a v a l u e Gmin' the elasticity of film P = - G & ( G ) / a G and surface wave dampinz decrement increases by almost one order of magnitude within a narrow transition area of width O.lGmin, and then decrease slowly (Monin and Krasitskii, 1985; Ermaliov, 1 9 8 7 ) . The viscosity increase in waters with films is an additional factor which increases 1 (Carlson,1987). There are a number of indirect mechanisms increasing Uoof warmer waters.The main one is formation of unstable stratification of the atmosphere above them which increases the wind stress and therefore the surface roughness increase. For example, Fig. 8 shows SST and air temperature profiles across the eddj- .A4 boundary.

The horizontal current shear is another possible factor affecting the level of scattered radar signals. The water speed in synoptic eddies, streamers and the Kuroshio major branch is 1-2m/s. The abrupt drop of the speed down to 0.1-0.3 m/s

occurs several miles apart from the flow boundary. Similar inhomogeneities of current.s at wind speeds of 2 to 7 m/s cause a striking (tT,o

or three orders of magnitude) change of spectral intensity of the capillary-

.

gravity waves due to nonlinear effects (Van Gastel, 1987). Wind speed variations with respect to moving water surface should be also taken into account.Such cross-stream variations may reach 20-50% at wind speed of ,4-6m/s. Such an effect is observed in an eddy, where wind and current directions are the same on one side of the eddy and the opposite on the other side. Probably, inhomogeneities of the brightness field in the anticyclone A3 region (Fig. 3 ) are associated with this. 4. CONCLUDING REMARKS

Similarity of radar reflectivity and temperature fields can be used to study the oceanic frontal zones. Radar information is particularly useful in the regions covered by clouds. The observation of the oceanographic phenomena on radar images is, however, limited by the rather narrow range of the surface wind speed variability (from 2 to 8-10 m/s). To explane the relationship between no and the SST, one should take into account the dependence of water

101 iiscosit y

on

the

temperature, the close connection between

surfactant

concentration and SST field, and the variation in the stability of the air flow above the sea surface caused by changes in SST. The determination of the relative contribution of each of the / above mechanisms in radar contrast formation is a complicated problem 'and requires special satellite and subsatellite

measurements.

The

effects

of

water

temperature,

film

concentration, current speed variation and other parameters on roughness spectrum change with wave length A (Donelan and Pierson, 1987; Ermakov et al., 1 9 8 7 ) . Therefore, radar sensing of the sea surface at different wavelengths A

is of importance to solving this problem and to estimating potential power of

radar techniques in the detection and study of oceanographic phenomena. ACKNOWLEDGEMENTS We

wish

to

thank

N.V.

Bulatov

(Pacific Institute of

Fishery

and

Oceanography) for helpful discussion and reviewers for their comments and suggestions. REFERENCES Carlson, D.J., 1987. Viscosity of sea-surface slicks. Nature, 329: 823-825. Donelan, # . A . and Pierson, W.J., 1987. Radar scattering and equilibrium ranges in wind-generated waves with application to scatterometry. J. Geophys. Res., 92: 4971-5029. Ermakov, S . h . , 1987. Film slicks on sea surface. I n : A . V . Gaponov-Grekov and S . A . Khristianovich (Editors), Nethods of Hydrophysical Research. Waves and Vortices. Inst. Pricl. Phys., Gorky, pp. 259-277 (in Russian). Ermakov, S.A., Zuikova, A.M. and Salashin, S.G., 1987. Transformation of short wind wave spectra in film slicks. Izv. Acad. Sci. USSR, Atmos. Oceanic Phys,, Engl. Transl. 23: 707-71.5. Ginsburg, A . I . and Fedorov, K.N., 1986. Near-surface water circulation in the subarctic frontal zone from satellite data. Issled. Zemli iz Cosmosa, Xo. 1 : 8-13 (in Russian). Kun, M.S., 1969. Seasonal changes in mesoplankton constitution and distribution in the liuroshio (observat,ions of 1965-1966). Izv. TINRO, 68: 93-110 (in Russian). Mitnik, L.M. and Victorov, S.V. (Editors), 1990. Radar Sensing o f the Earth Surface from Space. Gidrometeoizdat, Leningrad, 200 pp. (in Russian). Monin, A.S. and firasitskii, V.P., 1985. Phenomena on the Ocean Surface. Gidrometeoizdat, Leningrad, 3 i 5 pp. (in Russian). Thomas, .-\.C. and Emery, W. J., 1988. Relationships between near-surface plankton concentrations, hydrography, and satellite-measured sea surface temperature. J. Geophys. Res., 93: 11733-15748. Van Gastel, i i . , 1987. Imaging by N band radar of subsurface features: nonlinear phenomenon. J. Gephys. Res., 92: 11857-11865.


103

SST STRUCTURE OF THE POLAR FRONT IN THE JAPAN SEA Y. ISODA' ,S. SAITOH2 and M. MIHARA3 lDepartment of Ocean Engineering, Ehime University, Bunkyo 3, Matsuyama 790

(J 2Research Institute, Japan Weather Association, Kojimachi, Chiyoda-ku, Tokyo I02 (Japan)

Kaiji Center Bldg.,5,4-chome,

3ECOH; Enviromental Consultant for Ocean and Human, Co., Ltd., Minamisenju, 59-7, 1-chome, Arakawa-ku, Tokyo 116 (Japan)

ABSTRACT Horizontal s t r u c t u r e and s e a s o n a l v a r i a b i l i t y of t h e P o l a r f r o n t in t h e Japan Sea a r e examined by analyzing NOAA 9 infrared images (IR) and hydrographic d a t a in 1987. T h e Polar f r o n t o r i e n t e d in t h e w e s t - e a s t d i r e c t i o n s p r e a d s f r o m t h e Korean c o a s t t o t h e Tsugaru S t r a i t . Its n o r t h e r n e n d c o r r e s p o n d s t o t h e l o c a t i o n of a s y n o p t i c w a r m eddy t r a p p e d o v e r t h e Y a m a t o Rise. Besides, t h e Polar front has a different structure between western and eastern parts of the Yamato Rise. The western Polar front spreads over the Korean coast from 39"N to 42"N. I t s s t r u c t u r e is not so d i s t i n c t t h a t i t d i s a p p e a r s in a n IR i m a g e in s u m m e r by t h e d e v e l o p m e n t of s t r a t i f i c a t i o n d u e t o a s u r f a c e h e a t i n g . On t h e o t h e r hand, t h e e a s t e r n f r o n t h a s a r e l a t i v e l y r e m a r k a b l e SST ( S e a S u r f a c e T e m p e r a t u r e ) g r a d i e n t in all seasons. INTKODUCTION T h e Tsushima C u r r e n t s y s t e m i s a m a j o r hydrographic f e a t u r e in t h e J a p a n Sea. Warm and s a l i n e w a t e r of t h i s c u r r e n t e n t e r s through t h e Tsushima S t r a i t

-

and flows northward above a homogeneous water of low temperature (0" low salinity (34.0-

l ° C ) and

34.1 psu) which is called the Proper Water in the Japan Sea.

Then, most of t h e w a r m w a t e r flow o u t through t h e Tsugaru S t r a i t . It i s well known that this warm water can not spread to the northern end of the Japan Sea d u e t o t h e f o r m a t i o n of a t h e r m a l f r o n t in t h e c e n t r a l region. This t h e r m a l f r o n t h a s b e e n c a l l e d " P o l a r f r o n t ' ' a n d is o n e of t h e i m p o r t a n t p h y s i c a l phenomenon f o r understanding t h e h o r i z o n t a l c i r c u l a t i o n p a t t e r n s in t h e J a p a n Sea. However, l i t t l e i s known a b o u t t h e o b s e r v a t i o n a l c h a r a c t e r i s t i c s of t h e Polar front in the Japan Sea. This lack in our knowledge mainly results from the l i m i t a t i o n in t i m e and s p a c e resolution of shipboard o b s e r v a t i o n d u e t o t h e large horizontal scale of the Polar front, about 1000 km. Recently, some numerical experiments, which were done by basically inflow-outflow model, were carried o u t in o r d e r t o s u p p l e m e n t t h e knowledge a b o u t d y n a m i c a l s t r u c t u r e of t h e

104 Tsushima C u r r e n t s y s t e m (for e x a m p l e , Yoon, 1982a,b; K a w a b e , l 9 8 2 a a n d Sekine,1986). However, t h e P o l a r f r o n t w a s n o t r e p r o d u c e d in t h e s e models and i t s f o r m a t i o n mechanism c o u l d not b e clarified. Thus, w e c a n n o t y e t say t h a t t h e f u n d a m e n t a l s t r u c t u r e of t h e P o l a r f r o n t in t h e J a p a n Sea h a s been understood fairly well. The Polar front should be interpreted as the boundary between s u b a r c t i c and s u b t r o p i c a l w a t e r s , i.e.

t h e P r o p e r W a t e r in t h e J a p a n S e a a n d

Tsushima Current water. A further work will be important t o name a more approp r i a t e terminology f o r t h i s front. But w e shall u s e "Polar f r o n t " through t h i s paper

for convenience.

S a t e l l i t e i n f r a r e d (IR) i m a g e s a r e useful to e x a m i n i n g t h e h o r i z o n t a l SST s t r u c t u r e of

t h e P o l a r f r o n t . P r e v i o u s s t u d i e s (Toba et a1.,1984;

Saitoh,1988 and S u g i m u r a et al.,1989)

Isoda and

i n d i c a t e a usefulness of IR i m a g e s f o r

i n v e s t i g a t i n g t h e v a r i a b i l i t y of t h e Tsushima C u r r e n t . A f t e r a s e a r c h of a l l images published in the JMA-NOAA d a t a catalog(1987), it was found that 105 partial cloud free i m a g e s a r e a v a i l a b l e i n t h e J a p a n S e a f o r 1987. In t h e p r e s e n t s t u d y , w e discuss t h e h o r i z o n t a l s t r u c t u r e and s e a s o n a l v a r i a t i o n of t h e Pol.ar f r o n t on t h e basis of t h e r e s u l t s f r o m analysis of IR i m a g e s and h y d r o g r a p h i c d a t a in the Japan Sea for 1987. Although the present study is a descriptive one, i t will give useful i n f o r m a t i o n a b o u t f u t u r e i n v e s t i g a t i o n on t h e g e n e s i s of t h e P o l a r front. PROCESSING OF THE SATELLITE IR IMAGES The d a t a utilized in t h e present analysis consist of 105 IR images collected by NOAA 9 AVHRR from January to December, 1987 as shown in Table 1. Our study

area is shown by the thick line

in Fig. 1. W e chose the AVHRR channel 4 d a t a

f o r observing SST. T h e s e s e l e c t e d IR i m a g e s a r e p a r t i a l l y a f f e c t e d by clouds. Isoda and Saitoh(1988) developed a composite method removing the cloud d a t a by

TABLE 1 Temporal distribution of NOAA 9 AVHRR images selected in this study.

JAN FEB

MAR

APR HAY JUN JUL AUG SEP 0 CT NOV DEC

8 6

9 9 9 1

6 9

9 11 12 10

105

105

B :Korean Plateau C:Tsushirna S t r . D:Tsugaru Str.

0

200

400km

141O E

129OE

Fig. 1. The b o t t o m topography in t h e J a p a n Sea. Thick outlined region i s t h e area covered by the NOAA 9 AVHRR imagery utilized in this study.

divided into appropriate grids according to t h e spatial scale for the phenomena of concern. Although i n f o r m a t i o n of small s c a l e e d d i e s w i t h

s h o r t l i f e t i m e is

crashed by this method, the synoptic SST structure can b e represented by selecting an appropriate grid size and composite period. The major hydrographic features in the southern interior region of t h e Japan

Sea a r e

meandering paths of t h e Tsushima Current or synoptic scale eddies, of

which scales a r e a few hundred km (e.g. Kawabe,l982b). Figure 2 shows the traj e c t o r i e s of w a r m and cold e d d i e s in 1987. T h e i r positions show c o n s i d e r a b l e t e m p o r a l v a r i a t i o n s ; f o r e x a m p l e , t h e n o r t h e a s t w a r d s h i f t of a w a r m eddy lying

to t h e west of the Sad0 Island and the counter-clockwise rotation of a warm eddy between the Noto Peninsula and the Oki Islands. However, their moving distance in one month is shorter than several tens of

kilometers. This may indicate that

t h e s y n o p t i c eddy s t r u c t u r e c a n be also r e p r e s e n t e d by t h e monthly c o m p o s i t e

106

(a)Warm eddy 40"N

3S0N 1 3b0E

13b0E

1 UboE

( b ) C o l d eddy 40'N

35'N

1 2 3 4 5 6 7 8 9 101112

(month) Fig. 2. T h e t r a j e c t o r i e s of ( a ) w a r m and ( b ) c o l d e d d i e s in 1987( a f t e r the intantaneous ocean map of JMSA; 1987). Number indicates t h e month.

Fig. 3. Temporal change in the spatial mean SST (open circle) and spatial standard deviation (thick line) of t h e composite IR images in 1987.

images. Therefore, w e chose 0.25"mesh

(approximately 27X 20km) as the spatially

averaged

grid size and one month as the composite period.

Then, w e made up

12

monthly composite IR images in 1987 with a method by Isoda and Saitoh(1988). The right hand side numbers in Table 1 denote the composite numbers in each month. MONTHLY COMPOSITE IR IMAGES F i g u r e 3 shows t h e s p a t i a l m e a n and s t a n d a r d d e v i a t i o n of SST in e a c h monthly composite IR image. The seasonal variation of spatial mean SST is characterized by the minimum (about 5°C) in January through April, and the maximum (about 17°C) in July through September. Such a seasonal variation of SST may b e affected by both the heat exchange through the sea surface and the volume transport of warm waters through the Tsushima Strait. Although the volume transport increases in summer, spatial standard deviation of SST tends to become minimum. This may b e b e c a u s e a s t r o n g s t r a t i f i c a t i o n i s developed at a b o u t 5 0 m d e p t h d u e to a s u r f a c e h e a t i n g ( Minami et a1.,1987).

shallower than

On t h e o t h e r

107 hand, o c e a n i c s t r u c t u r e in w i n t e r m a y e a s i l y a p p e a r in a n SST d u e t o t h e vert i c a l mixing of a s u r f a c e cooling. F i g u r e 4 shows t h e c o m p o s i t e IR i m a g e s f o r even months in 1987. Warm waters extend over the southern part of the Japan Sea in all seasons. These warm waters seem to converge at the south of the Tsugaru S t r a i t ( t h e l o c a t i o n of Tsugaru S t r a i t is d e n o t e d by t h e black a r r o w in Fig, 4 ) and afterward t o flow northward along t h e Japanese coast.

of t h e J a p a n Sea, a s t r o n g SST g r a d i e n t region, i.e. the east-west

In the interior region

a t h e r m a l f r o n t , runs in

direction with significant meandering, particularly

from

winter

to

spring. N o t e t h a t t h i s f r o n t is n o t c o n n e c t e d t o t h e s o u t h e r n e n t r a n c e of warm w a t e r and in d i r e c t c o n t a c t w i t h t h e K o r e a n c o a s t .

We h e r e a f t e r c a l l t h i s

thermal front in t h e central region of t h e Japan Sea, the Polar front. Meanders of the Polar front from winter t o spring correspond well t o the bottom topography around the Yamato Rise (Figs. 1 and 4). That is, the existence of a synoptic

eddy just above the Yamato Rise can be inferred from such a mean-

dering. Moreover, t h e s h a r p n e s s of t h e P o l a r f r o n t is d i f f e r e n t b e t w e e n t h e eastern and western parts of t h e Yamato Rise. That is, the western front struct u r e is w e a k e r t h a n t h e e a s t e r n one. T h e w e s t e r n P o l a r f r o n t s p r e a d s o v e r t h e K o r e a n c o a s t and i s

divided i n t o t w o p a r t s around 3 9 " N and 4 2 " N . T h e P o l a r

front around 39"N tends t o become sharp only near the Korean coast. Then, this western Polar front cannot be traced in an IR image in t h e heating season from June t o August and another thermal front is formed from the Tsushima Strait to the Yamato Rise in August. The SST structure of the Polar front in t h e western region begins to appear again in October, when a sea surface cooling begins. On t h e o t h e r hand, t h e e a s t e r n P o l a r f r o n t f r o m t h e Y a m a t o R i s e t o t h e Tsugaru S t r a i t h a s a r e l a t i v e l y s t r o n g SST g r a d i e n t . Especially, t h e P o l a r f r o n t around t h e Tsugaru S t r a i t is confined t o t h e J a p a n e s e c o a s t a l region and c a n b e d e t e c t e d in all s e a s o n s as m a i n t a i n i n g t h e s h a r p t h e r m a l f r o n t . T h e s e r e s u l t s may indicate that the Polar front

i s influenced by a s y n o p t i c s c a l e s t r u c t u r e above

t h e Y a m a t o R i s e and h a s a d i f f e r e n t SST s t r u c t u r e b e t w e e n i t s w e s t e r n and e a s t e r n parts. Hence, w e shall i n v e s t i g a t e t h e r e l a t i o n s h i p b e t w e e n t h e P o l a r f r o n t and t h e b o t t o m topography around t h e Y a m a t o Rise. F i g u r e 5 s h o w s SST d i s t r i b u t i o n s in e a c h monthly IR i m a g e s and t h e s e c t i o n of b o t t o m topography along 132"E, 134"E and 136"E (thick broken lines in Fig.

1 ) . Seasonal variations

of SST common t o the three lines a r e as follows. The increase of SST from March

to August is uniform everywhere and the Polar front loose its clearness. On t h e o t h e r hand, t h e d e c r e a s e of SST from S e p t e m b e r t o F e b r a u a r y m a k e s t h e P o l a r front clear again. Such SST variations show that SST is mainly influenced by the heat exchange through the sea surface. That is, oceanic structure will b e masked

or weakened by the development of a surface stratification in summer. Therefore, we cannot judge that the western Polar front does not exist in summer from only SST d a t a . T h e s h a r p n e s s of t h e P o l a r f r o n t along 134"E i s s t r o n g e r t h a n t h o s e

Fig. 4. T h e c o m p o s i t e IR i m a g e s for e v e n m o n t h s i n 1987. 1x1 d e n o t e s t h e g r i d lacked by t h e removal of cloud data. The black arrow shows th e location of t h e T s u g ar u S t r a i t .

109 ( a ) 132OE L i n e (West o f the Yamato Rise) 36oN

U

16

t

'2

ln ln

38'N

40"N

42ON

40°N

20

36ON

38'N

40°N

42ON

4 4 0 ~

3G0N

38'N

40"N

42-N

44'N

36"N

38'N

90'N

42-N

44"N

20

U

16

t-

12

ln ln

8

8

4

4

( b ) 134OE L i n e ( O v e r t h e Yamato Rise) 3G0N

3a0N

40°N

42ON

44ON

20

20

U

16

U

16

I-

'2

k

'2

ln

ln

m

m

8

8

4

4 -0

-0

552 2

5 5 2 w

4

4

"N

( c ) 136OE L i n e ( E a s t of the Yamato Rise) ON

20

W

16

c

'2

m

u)

8 4 - 0

552 a w

36ON

38'N

40'N

42'N

4

44'N

Fig. 5. SST p r o f i l e s in e a c h c o m p o s i t e IR i m a g e ( u p p e r ) and t h e s e c t i o n o f b o t t o m topography ( l o w e r ) along 1 3 2 " E ( a ) , 134"E(b) and 1 3 6 " E ( c ) lines shown in Fig. 1.

a l o n g o t h e r l i n e s . B e s i d e s , i t s l o c a t i o n Ts a l s o s t a b l e i n t i m e a b o v e t h e Thus, i t i s s u g g e s t e d t h a t a s t r o n g o c e a n i c

n o r t h e r n end of t h e Y a m a t o Rise.

structure trapped above the Yamato Rise exists in all seasons.

110

HORIZONTAL DISTRIBUTION OF TEMPERATURE AT 100 M DEPTH Japanese hydrographic d a t a a r e quantitatively insufficient t o cover t h e whole J a p a n Sea, and w e i n v e s t i g a t e t h e o c e a n i c c o n d i t i o n s only around t h e Yamato Rise. Figure 6 shows the horizontal distributions of temperature at 100 m depth in 1987 a f t e r the "ocean instantaneous map" compiled by JMSA(l987). The

2000 m isobath is indicated by a thick broken line. The cold water region occupies both sides' basins of the Yamato Rise, i.e. Yamato Basin and Tsushima Basin. The warm water region is stably confined above the Yamato Rise in all seasons. A warm eddy is formed above the western part of t h e Y a m a t o Rise, although i t s h o r i z o n t a l t e m p e r a t u r e g r a d i e n t i s n o t a l w a y s large. This w a r m eddy l o c a t i o n i s c o n s i s t e n t w i t h o n e i n f e r r e d

f r o m t h e SST

meandering of t h e P o l a r f r o n t in an IR image. Thus, i t is c o n f i r m e d t h a t t h e s p a t i a l p a t t e r n of w a r m and cold w a t e r regions well c o r r e s p o n d s t o t h e b o t t o m topography. This relationship is

a l s o f o u n d i n o t h e r r e p o r t s b a s e d on t h e

s a l i n e w a t e r d i s t r i b u t i o n s (Tanioka ,1962 a n d Ogawa, 1974). Moreover, Kolpack (1982) suggested t h a t t h e position of t h e P o l a r f r o n t in t h e J a p a n S e a is c o n t r o l l e d by t h e b o t t o m topography. T h e s e r e p o r t s support t h e p r e s e n t view t h a t t h e s y n o p t i c eddies in t h e J a p a n Sea are a f f e c t e d by t h e b o t t o m topography, especially over the Yamato Rise. CONCLUSION AND DISCUSSION The Polar front oriented in the east-west

direction is formed in the central

region of t h e J a p a n Sea. S o m e f u n d a m e n t a l p r o p e r t i e s of t h i s Polar f r o n t are understood through t h e p r e s e n t s t u d y using t h e monthly c o m p o s i t e IR i m a g e s in 1987. W e point out the following three points which are considered t o be import a n t subjects. ( 1 ) T h e n o r t h e r n e n d o f t h e P o l a r f r o n t c o r r e s p o n d s t o t h e l o c a t i o n of a

synoptic warm eddy trapped over the western part of the Yamato Rise. This warm eddy s t a b l y e x i s t s in a l l seasons. T h e r e f o r e , t h e P o l a r f r o n t d e p i c t s significant meandering around the Yamato Rise. (2)

T h e s h a r p n e s s of t h e P o l a r f r o n t is d i f f e r e n t b e t w e e n in t h e e a s t e r n p a r t

and w e s t e r n p a r t of t h e Y a m a t o Rise. T h e e a s t e r n f r o n t is s t r o n g e r and m o r e s t a b l e t h a n t h e w e s t e r n o n e in a l l seasons, p a r t i c u l a r l y around t h e e x i t of t h e Tsugaru S t r a i t .

( 3 ) The e a s t e r n e n d of t h e P o l a r f r o n t c o n n e c t s to t h e Tsugaru S t r a i t or along the Japanese coast. Such a feature shows that warm waters of Tsushima Current flow o u t through t h e Tsugaru S t r a i t or Soya S t r a i t . On t h e o t h e r hand, its western end is in direct contact with t h e Korean coast form 39"N to 42"N. I t is known that the East Korean Warm Current flows northward from the Tsushima Strait along t h e K o r e a n Peninsula. N e v e r t h e l e s s , a s i g n i f i c a n t n o r t h w a r d flow c a n n o t be inferred from t h e composite IR images. In a cloud free image off the Korean

f Y)

i.

f 0

z

T-0 Y)

i i

In

f

z

0 ir

z

0 In T-0

Y

F i g . 1. S t a t i o n s o c c u p i e d i n January, 1986 ( a ) and J u l y , 1986 (b). Symbols i n t h e b o t t o m panel r e p r e s e n t c r u i s e s and o r g a n i z a t i o n s l i s t e d i n T a b l e 1. I s o b a t h s i n meters.

111

112

c o a s t ( n o t shown), s o m e s y n o p t i c e d d i e s w i t h t h e f o r k - l i k e s p l i t t i n g of w a r m w a t e r a r e o f t e n o b s e r v e d ; h o w e v e r , i t s e e m s t h a t t h e b e h a v i o r s of e d d i e s a r e complicated. These fact suggest t h a t a synoptic warm eddy over t h e Y a m a t o Rise plays an i m p o r t a n t r o l e in t h e f o r m a t i o n of t h e P o l a r f r o n t . T h e r e f o r e , m o r e d e t a i l e d o b s e r v a t i o n , f o c u s i n g on t h e d y n a m i c s t r u c t u r e of t h i s w a r m e d d y should b e u n d e r t a k e n in t h e f e a t u r e . In a d d i t i o n , i n t e n s i v e w o r k s a r e n e e d e d to i n v e s t i g a t e t h e dynamical s t r u c t u r e of t h e E a s t Korean Warm Current. ACKNOWLEDGMENTS The authors express their sincere thanks to Dr. T. Yanagi of Ehime University for his discussion and encouragement during t h i s study. T h e d a t a analysis was carried out on a FACOM M 360 A P of Ehime University. REFERENCES Japan Weather Association, 1987. NOAA d a t a catalog. Japan Maritime Safety Agency, 1987. Ocean instantaneous map. Isoda,Y. a n d S.Saitoh, 1988. V a r i a b i l i t y of t h e s e a s u r f a c e t e m p e r a t u r e obtained by t h e statistical analysis of AVHRR imagery. - A case study of t h e south Japan Sea J.Oceanogr.Soc.Japan, 44, 52-59. Kawai,H., 1974, T r a n s i t i o n of c u r r e n t i m a g e s in t h e J a p a n S e a . In: T h e Tsushima Warm Current - Ocean s t r u c t u r e and fishery. ed.by Fishery SOC. Japan,Koseisha-Kouseikaku, pp.7-26. (in Japanese) Kawabe,M., 1982a. Branching of t h e Tsushima Current in t h e Japan Sea, P a r t 2. Numerical experiment. J.Oceanogr.Soc. Japan, 38, 183-192. Kawabe,M., 1982b. Branching of t h e Tsushima Current in t h e Japan Sea, P a r t 1. D a t a a n a l y s i s . J . O c e a n o g r . S o c . J a p a n , 38, 95-107. Kolpack,R.L., 1982. T e m p e r a t u r e and salinity changes in t h e Tsushima Current. Lar mer, 20, 199-209. M i n a m i , H , Y , H a s h i m o t o , Y . K o n i s h i a n d H.Daimon, 1987. S t a t i s t i c a l f e a t u r e s of t h e oceanographic condition in t h e Japan Sea. Umi t o Sora, 4, 163-175. (in J a p a n e s e ) Ogawa,Y., 1974. The relation between t h e high saline w a t e r in t h e Japan Sea and Tsushima Current. Bull. Japan Soc.Fish.Oceanogr., 24, 1-12. (in Japanese) Sekine,Y., 1986. Wind-driven c i r c u l a t i o n in t h e J a p a n S e a a n d i t s i n f l u e n c e on the branching of t h e Tsushima Current. Progress Oceanography, 17, 297-312. Sugimura,T.,S.Tanaka and Y.Hatakeyama, 1984. Surface t e m p e r a t u r e and current vectors in t h e sea of Japan from NOAA-7/AVHRR data. pp.133-147. In: R e m o t e sensing of shelf sea hydrodynamics, ed.by Jacques C. J.Nihoul,Elsevior, 38. T a n i o k a K., 1962. A r e v i e w of sea c o n d i t i o n s in t h e J a p a n S e a ( 2 ) - o n t h e cold, warm and saline w a t e r regions Umi to Sora, 38, 115-128. (in Japanese) Toba,Y.,H.Kawamura,F.Yamashita and K.Hanawa, 1984. S t r u c t u r e of horizontal t u r b u l e n c e in t h e J a p a n Sea. pp.317-332. In: O c e a n H y d r o d y n a m i c s of J a p a n a n d E a s t C h i n a S e a . ed.by T . I c h i y e , E l s e v i o r , 39. Yoon,J.H., 1982a. Numerical experiment on t h e circulation in t h e Japan Sea. P a r t 1. F o r m a t i o n of t h e E a s t K o r e a n w a r m C u r r e n t . J . O c e a n o g r . S o c . J a p a n , 38, 43-51. Yoon,J,H., 1982b. Numerical experiment on t h e circulation in t h e Japan Sea. P a r t 2. Formation of t h e nearshore branch of t h e Tsushima Current. J. Oceanogr.Soc.Japan, 38, 119-124.

-.

-.

113

A NUMERICAL EXPERIMENT ON THE S E A S O N A L VARIATION O F T H E O C E A N I C CIRCULATION IN THE JAPAN S E A

YOSHIHIKO S E K I N E Institute o f Oceanography. Faculty of Bioresources. Mie University. 1 5 1 5 Kanihaeachou. Tsu Hie 514 (Japan)

ABSTRACT The seasonal variation o f t h e g e n e r a l c i r c u l a t i o n in the J a p a n S e a is studied numerically with special reference t o the branching o f t h e Tsushima Current. A two-layer model with bottom and coastal topograpgv o f the J a p a n Sea is used. T h e seasonal c h a n g e in t h e inflow corresponding t o t h e T s u s h i m a Current and two o u t f l o w s corresponding t o t h e Tsugaru and S o y a Currents are given a s open boundary conditions. The wind-driven circulation dominates in winter. whereas t h e circulation in summer is d u e to the in-and o u t f l o w . In winter, a strong anti-cyclonic circulation i s formed in t h e deep Japan Basin by the intensified wind s t r e s s w i t h negative curl. In summer. the intensified inflow flows along the Japanese Coast and generates an anti-cyclonic circulation in t h e Japan Basin. The circulation in s p r i n g and autumn is transitional from the winter t o s u m m e r r e g i m e and from t h e s u m m e r t o winter regime. respectively. O n the whole. the Tsushima current h a s three dominant branches in the .Japan Sea. t h e coastal branch along t h e J a p a n e s e main Islans. t h e o f f s h o r e central branch a n d -the western boundary branch, of w h i c h remarkable seasonal variations depend on t h e v a r i a t i o n s of t h e in- and o u t f l o w and wind stress. INTRODUCTION A major feature of the current system in t h e Japan S e a is an inflow o f the Tsushima Current

t h r o u n h the Tshushima

into three branches

(Suda et a1.,1933;

Strait.

The T s u s h i m a Current s p l i t s

Uda.1934.36):

one branch flows along

the Japanese coast. the second branch flows in a central renion and t h e third branch runs along t h e Korean Coast

(

Fig.

I).

However,

Horiyasu

(1972a.b)

considered this flow pattern as a large m e a n d e r path rather than t h e branching into t h e three barnches,

because the Tsushima

Current

s h o w s a c l e a r large

meander at the s o u t h e r n Japan Sea. Because of t h e interannual and s e a s o n a l variation o f the inflow, the details of the Tsushima Current in t h e J a p a n S e a have not been clarified yet. So far. s o m e theoretical s t u d i e s on the Tsushima Current were carried out

(Table I ) .

The flow along the Japanese Coast i s d u e to t h e topographic guiding

effect along t h e geostrophic c o n t o u r (f/h). where f is t h e C o r i o l i s parameter and h is t h e water depth (Yoon.1982~). which is hereafter referred t o a s the topographic

branch

(TPB).

However,

the

direct

current

measurements

by

Matsuvama et al. (1986) showed t h e baroclinic structure of the TPB. w h i c h is not always controlled by the barotropic flow character along f/h. The s e c o n d dominant flow is along the K o r e a n coast. and i s identlfied with the K o r e a Warm

114

P a c i f i c Ocean

145'E

135O

125O

Fig. 1 . Schematic representatlon of three representative branches of the Tsushina Current flowinn into the Japan Sea. L e t t e r s TPB. O C P and WBCB mean the topographic branch alonn the Japanese Coast. the offshore central branch and the western boundary current branch correspondinn t o t h e Korean Warm Current. respectively.

TABLE 1 Model characteristics of the Japan S e a so far developed. Author

Density Coastal Bottom stratification topography topoqraphy

Wind

Heat*

Inflow

Yoon (1982 a)

Yes

Rectangular

Flat

Constant

Yoon (1982 b )

Yes

Realistic

Flat

S e a s o n a l S e a s o n a l Stationary variation variation

Yoon (1982 c)

No

Rectangular Simplified

No

No

Stationary

Kawabe(l982b)

Yes

Rectangular Simplified

No

No

Variable

Sekine(19861

Yes

Realistic

Realistic

No

Stationary

Present study

Yes

Realistic

Realistic

No

Seasonal variation

Present study

Yes

Realistic

Realistic

No

Seasonal variation

Seasonal variation

No Seasonal v a r 1 at i on

Constant

Stationary

* Constant, S e a s o n a l variation and N o mean s t a t i o n a r y t h e r m a l condition. seasonally varying thermal boundary condition and n o c o n s i d e r a t i o n on t h e t h e r m a l condition, respectively.

115 Current. I t is formed b y the planetary B effect

Yoon.1982a.b

(

as the western boundary current branch (WBCB).

referred to

) ?

'The last dominant flow in an

offshore central region (OCB) has not been clarified well b y the historical observation.

Kawabe(l982b)

showed

that

this

branch

is

by

caused

the

propagation of the lowest two modes of upper shelf waves which are generated by the increase in the inflow of the Tsushima Current in summer. On the other

hand.

Sekine(1986) pointed out that because the wind stress is very weak in

summer. this branch is formed by the westward shift of the kinetic energy of the TPB carried by planetary Rossby waves. Kawabe(l982a) showed from historical data analysis that the TPB along the Japanese coast exists at least from sprinn t o summer. the O C B in summer only and the WBCB throughout the year.

Sekine (1986) demonstrated

by

a numerical

model that the observed branching of the Tsushima Current except for summer season is well predominates

simulated by

in winter

the seasonally varying wind

under

the strong

wind

stress:

stress, whereas

the WBCB

the TPB

is

intensified in spring to summer because of the positive wind stress curl over the Japan Sea. which intensifies cyclonic circulations such as the TPB, but weakens anti-cyclonic circulations such as t h e WBCB. However. Sekine (1986) failed to simulate the WBCB in summer, which is observed throughout the year. This may be due to a constant in-and outflow given at the boundary (see, Table

I

).

Figure 2

displays the observed current velocity at the Tsushima Strait

but

the

inflow

by

A weak inflow with small vertical shear is seen in winter.

Inoue et a1.(1985).

is

intensified

in

summer

and

the vertical

shear becomes

prominent in autumn. Sekine(1989) pointed out that because the surface mixed laver penetrates t o the bottom in the Tsushima Strait from January t o April. the inflow is blocked by the sill toponraphy of the strait. which is suxgested by

the barotropic

isopleth

o f depth

( (

n o density stratification e.g..

Pedloskg.

1979).

)

flow character along the

In contrast

to this.

density

stratification develops in late summer to autumn and induces a strong inflow

WlNTER SPRING

SUMMER AUTUMN

WINTER SPRING

SUMMER

AUTUMN

2. Seasonal change in the current velocity a t t w o stations in the Tsushima Current (after: fnoue et a1.,1985). Locations of the two stations are shown by two solid marks in the map.

Frg.

116 with large vertical shear. Although a constant inflow is g i v e n in t h e m o d e l of Sekine (1986). there is a

possibility

Current

has

that an

t h e seasonal

important

role

variation in

the

in

the

circulation

inflow of of

the

the

Japan

Tsushima Sea

with

reference t o its hranchinn. F o r t h i s reason. t h e e f f e c t s of s e a s o n a l change in the in- and

outflow

tonether with

the effects of seasonal change

stress a r e examined in t h e present study (Table 1 ) .

in wind

The numerical model

*ill

be described in the next section. R e s u l t s are presented in s e c t i o n 3. Summary and discussion will be g i v e n in s e c t i o n 4 .

2 NUMERICAL MODEL Figure 3

shows t h e schematic r e p r e s e n t a t ~ o n o f the model ocean. Coastal

and bottom toponraphy are modelled in a simplified form. A two-laver model is used.

The

basic

equations

by

hydrostatic

balance.

13-plane.

rigid

lid

and

Boussinesq approximations are the s a m e as in S e k i n e (1986). T h e coefficient o f horizontal eddy viscosity is 10' c m z s 'and the reduced n r a v i t y 2.77cms details of t h e basic equations, s e e Sekine. wind

stress

stress

data

Sakurai(l982).

is obtained by

Japan

by

linear

interpolation

Yeteorological

for

'(

1986). The annual c y c l e o f t h e

Agency

of the monthly mean (1972)

and

wind

Kutsuwada

and

which is displayed i n Fig. 4. In winter. a s t r o n g northwesterly

from the Siberian H i g h prevails. while the wind s t r e s s is very weak in summer. The annual change o v e r the whole d o m a i n is s h o w n

i n Fig.

4(b).

The average of

wind s t r e s s curl is n e n a t i v e in winter. but positive in A p r i l t o July.

f" FIR. 3. ( a ) Domain o f t h e model s h o w n is stippled. Three in-and o u t f l o w are

of the model o c e a n (hatched renion), (b) bottom topography by isobaths ( m ) . Contour interval is 500m and s h a l l o w a r e a thick lines on t h e coastal boundary s h o w t h e renion where imposed.

117

NOV.

b

F i n . 4. (a) Monthly m e a n wind s t r e s s data f o r 1958 - 7 5 ( a f t e r J a p a n M e t e o r o i o n i c a l AnenCv.1972; K u s t u w a d a and Sakurai.1982). No v e c t o r s less t h a n 0.05 d y n e c m a r e plotted. (b) M e a n x-compenent of monthlv mean wind s t r e s s ( t x ) . mean v-component ( C Y ) and wind s t r e s s curl o v e r the w h o l e m o d e l d o m a i n (for the d i r e c t i o n s x-and v c o o r d i n a t e .see Ein.3).

-

118 The

inflow velocity

U,sin(xv/L). the inflow.

in the upper and

lower layers is assumed

to be

where U o is the maximum inflow velocity and L is the width of Similar velocity profiles are assumed at the two outflow

boundaries, The annual change in the inflow is shown in Fig. 5 in terms of U , for both

layers.

The velocity

and

total volume

transport

represent

the

observed features shown in Fig. 2. I t is also assumed that in all seasons. 652 ( 3 5 % ) of the inflow volume flow out through the Tsugaru (Soya) Strait. In order to see the individual influences of the in- and outflow and wind stress on the current system in the Japan Sea. numerical experiments are performed in three cases. The first case is driven by the in- and outflow and no wind stress is imposed. The second case is driven by both the in- and outflow and wind stress. The third case is the same as the second case except for the inflow condltion. Since the seasonal variation in the inflow is more remarkable in the western part of the Tsushima Strait (Kawabe. 1982a). the seasonal change in the inflow is confined to the western two thirds of the boundary , while the stationary inflow with the ainimus velocity of January is niven at the eastern one third. F o r these three cases. the time integration of the basic equations starts by the wind stress on 1 January. The initial flow is given alonn the southern boundary froa the inflow renion to the two outflow boundaries.

1

$ 2 Y

F I X . 5.

Seasonal channe in the inflow given in the model.fa) Total volume transport. (b) maximum velocity ( U n ) of the inflow with the sinusoidal distribution in the horizontal direction across the inflow boundary: u = U sin(nylL). where I, is the width of the inflow.

3 RESULTS A time integration was carried out for 10 years for the three cases. As was shown by Sekine ( 1 9 8 6 ) . periodic annual changes in kinetic enrrny and flow

patterns are clearly seen in the numerical solution. Therefore. the solution for the last one year is examined in each case.

119 The annual velocity field in t h e first case is displayed in Fig. 6. Except in late spring, the upper layer inflow h a s a tendency t o f l o w along t h e n o r t h e r n boundary, which c o r r e s p o n d s t o t h e WBCB. However. in summer, t h e u p p e r laver inflow flowing along t h e s o u t h e r n coast a l s o dominates, which corresponds t o t h e TPB. From late s u m m e r t o autumn. a f l o w i s formed in t h e offshore central region by the westward shift o f t h e k i n e t i c energy o f t h e intensified TPB. Although this f l o w may be different from t h e O C B d e n o t e d by Kawabe(l982a,b), i t is possibly formed in a n area west o f t h e T P B by t h i s process. In late autumn. i t is carried further westward and absorbed i n the WBCB. From autumn t o winter, most o f t h e u p p e r layer i n f l o w feed t h e W B C B and the TPB disappears in t h e u p p e r layer in November. Because of the predominance o f bottom intensifeid m o d e (Rhines. 1970; Suginohara, 1982). the velocity field in t h e lower layer h a s a s t r o n g tendency t o f l o w along the geostrophic contour. which is approximated by t h e isobath in a local area where c h a n g e in t h e Coriolis parameter

is relatively small.

Because there e x i s t s a relatively wide continental slope in o f f s h o r e o f t h e Japanese Coast, a n inflow in t h e lower l a y e r is forced t o f l o w along t h e s o u t h e r n boundary in all the season. F r o m s u m m e r t o autumn, a prominent anticyclonic circulation is formed in t h e n o r t h e r n Japan B a s i n with a d e e p and relatively flat bottom (see. Fig. 3b). T h i s circulation includes the n o r t h e r n boundary current and t h e westward f l o w along an isobath north off Y a m a t o Ridge. In particular. t h e l a t t e r westward f l o w along t h e bottom s l o p e may be connected with a s h a r p front observed here. Although t.he main g e n e r a t i o n p r o c e s s of t h e observed front may be due t o o t h e r processes. an important role o f t h e continental s l o p e off Y a m a t o R i d g e i s suggested from this experiment. The annual variation in total volume transport function in t h e u p p e r and lower layer is s h o w n in Fig. 7. To the south of t h e o u t f l o w corresponding t o the T s u g a r u Strait, most o f t h e TPB turn westward and f l o w along a n isobath, which generates a n anti-cyclonic circulation in the n o r t h e r n h a l f o f t h e basin. T h i s anti-cyclonic

circulation is m o r e remarkable from late s p r i n g t o

summer, but decays in autumn. It is c l e a r that t h i s circulation is formed by the increase in i n f l o w transport. At the i n f l o w region. t h e W B C B is w e a k in spring, but intensified gradually in late s u m m e r t o autumn. T h i s is c a u s e d by t h e development of t h e ocean r e s p o n s e to t h e strong inflow and the k i n e t i c energy s h i f t s t o t h e western boundary r e n i o n with lapse o f time. I f the stationary in- and o u t f l o w is Imposed, t h e W B C B is not c l e a r i n s u m m e r (Sekine, 1 9 8 6 ) . Therefore. t h e formation o f t h e WBCB in s u m m e r is conclusively due to the strong inflow transport that g i v e s t h e WBCB t h e k i n e t i c energy. T h e seasonal c h a n g e o f the velocity field in t h e second c a s e driven by both the in-and o u t f l o w and wind s t r e s s is characterized by t h e formation of a s t r o n h anti-cyclonic circulation in winter (Fig. 8 ) . T h i s c i r c u l a t i o n is generared by the strong wind s t r e s s in w i n t e r with the n e g a t i v e curl (see. Fig. 4). However, a s t h e wind s t r e s s is w e a k e n e d in spring, the center o f the

F i n . 6 V c l o c i t v f i e l d of t h e u p p e r (UI.VI) and l o w e r (U2,VZ) l a y e r i n t h e c a s e of no wind s t r e s s . V e c o t o r s a r e p l o t t e d a t e v e r y o t h e r n r i d p o i n t . No v e c t o r s s m a l l e r t h a n 0 . 0 5 cms ' a r e p l o t t e d .

anti-cyclonic

c i r c u l a t i o n moves westward and d e c a y s i n l a t e s p r i n g . The TPB

and o f f s h o r e c e n t r a l f l o w a r e g e n e r a t e d i n summer. Because t h e wind s t r e s s i s weak i n t h e l a t t e r h a l f of t h e y e a r . t h e v e l o c i t v f i e l d i s s i m i l a r t o t h a t i n

the

first

case

d i a g r a m s of wind-driven

in

circulation

due t o t h e in-and responce

to

this

period.

the velocitv vectors

these

which (

Fig.

is d i s p l a y e d 9).

more c l e a r l y

bv s t i c k

I t is thus concluded t h a t

the

is dominant i n w i n t e r and t h e c i r c u l a t i o n i s m a i n l y

o u t f l o w i n summer. F u r t h e r m o r e . t h e p r o c e s s o f t h e o c e a n i c two e v e n t s

is

achieved

: the

circulation

in

spring

is

121

2TRANSPURT MAY

Fin. 7. Seasonal change in he calculated volume transport function. C o n t o u r interval is 0.2 S v (isv = 1 0 ' 2 c m 3 s - ' ) and r e g i o n s with neRative transport function

(

cvclonic circulat on

characterized

bv

spin-down

)

of

are stippled. the

anti-circulation

circulation

formed

by

strong wind s t r e s s in winter. O n t h e o t h e r hand. t h e main D r o c e s s in a u t u m n is a westward shift o f t h e kinetic energy g i v e n by the enhanced T P B in summer. The difference

in these

flow patterns

from those of S e k i n e

(1986) w i t h a

constant inflow o f 2 Sv is small in winter but significant in s u m a e r t o a u t u m n when an inflow l a r q e r than 2 S v is specified in the present study. T h e volume transport function in t h e s e c o n d case and t h e net wind-driven circulation which

is obtained by subtracting

first case from that former

half

year.

characterized

by

in the s e c o n d case.

the the

seasonal formation

change of

the

the transport

in

the

total

winter

to

southwestward

spring.

counter

a

flow

northeastward in

the

volume

anti-cyclonic

northern basin in winter and its gradual spin-down From

function i n the

a r e displayed in Fig.

flow

southern

10. In the

transport

circulation

in

is the

in t h e following period. in

part

the of

TPB the

contacts

a

anti-cyclonic

ciruclation, which attenuates t h e TPB. T h e p r o c e s s d e p e n d s o n t h e r e l a t i v e intensity of the in- and o u t f l o w and t h e anti-cyclonic

circulation d r i v e n by

t h e wind stress. B e c a u s e t h e s i m u l a t e d volume transport of t h e anti-cyclonic circulation e x c e e d s 4 S v in the present model, larger than an observed i n f l o w (

2 S v at most) in w i n t e r (Inoue et al..

1985).

t h e TPB b e c o m e s weak

results is common t o t h e case with the larger inflow of 2 S v Hence,

the TPB

is expected

t o be weakened

in winter.

(

Sekine.

In spring.

. This 1986).

cvclonic

e d d i e s are formed i n the contact region. In t h e n o r t h e r n reqion. the formation

of t h e anti-circulation

is s l o w in this c a s e in comparison with the first

ZZI F i g . 8 . Same a s F i g . stress. case.

which

cyclonic

is

6 b u t f o r t h e s e c o n d c a s e w i t h s e a s o n a l c h a n g e i n wind

clearly

circulatlon

seen

in

i s weakened

the by

transport

function

the positive

I n May.

'The a n t i -

wind s t r e s s c u r l

in late

s p r i n 8 t o e a r l y summer. which g e n e r a t e s t h e c y c l o n i c c i r c u l a t i o n i n t h e J a p a n Basin

(Fig.

lob).

I n l a t e summer t o autumn, b e c a u s e t h e wind s t r e s s is v e r y

weak, no r e m a r k a b l e d i f f e r e n c e i s d e t e c t e d between t h e two c a s e s . F i g u r e 1 1 shows t h e r e s u l t s i n t h e t h i r d c a s e w i t h t h e e n h a n c e d s e a s o n a l variation

of

inflow

at

the

western

part

of

the

boundary.

The

difference

between t h e s e c o n d and t h i r d c a s e i s r e l a t i v e l y r e m a r k a b l e f r o m summer t o l a t e autumn and i n mid w i n t e r ( F i g . I l b ) . I n summer, t h e f l o w i n t h e c e n t r a l r e g i o n

is more i n t e n s i f i e d t h a n i n t h e s e c o n d c a s e . However. t h i s i n t e n s i f i e d f l o w is s h i f t e d westward more q u i c k l y and a b s o r b e d

i n t h e WBCB i n l a t e autumn.

The

u p p e r l a y e r v e l o c i t y f i e l d i n November i n d i c a t e s t h a t t h e WBCB i n autumn is

Fig. 9. Stick diagram o f the upper and lower mean velocities across (a) near the inflow boundary.(b) central region. Case with n o wind stress (upper panel) and case with seasonal change in wind stress(1ower panel). The thick lines show the sections where the mean velocities a r e calculated. caused by the westward shift of the kinetic energy of the TPB formed by the intensified inflow in summer. The difference between the t w o cases in winter is caused by the combined effect o f continuing westward shift o f the kinetic energy o f the TPB and the strong wind stress in winter. Figure 12 implies that the westward shift of the kinetic energy is accomplished in January. After this, the difference is relatively small until the increase in the inflow begins t o be intensified. Two cvclonic circulations are combined near the inflow and outflow regions. which is clearly shown in the volume transport function (Fig. 13b). One cyclonic circulation is caused by intensified s h e a r of the western part o f t h e inflow. while the other is caused by the enhanced flow in t h e central region. On the whole, there exists n o significant difference in the total flow pattern between the second and third cases. However. the seasonal variation o f the inflow confined into the western two thirds of the boundary yields fast oceanic response t o the increased inflow in sunmer, and the central offshore flow and the WBCB are formed faster than in t h e second case. 4 SUMMARY AND D I S C U S S I O N

The seasonal variation of the circulation in the Japan Sea has been examined numerically. A two-layer model is used and t h e observed in- and outflow is imposed through the boundaries. The main results are summarized a s follows: ( 1 ) In winter, a strong anti-cycvclonic circulation is formed in the northern Jaoan Basin by the intensified wind stress with negative curl. Because the inflow through the Tsushima Strait is weak, the TPB is suppressed by a counter current generated by the anti-cyclonic circulation. These results agree with those of Sekine ( 1 9 8 6 ) with a stationary inflow o f 2Sv.

124

b

F i n . 10. (a) S a m e as Fin. 7 but for the s e c o n d c a s e with s e a s o n a l c h a n n e in w ~ n d stress. (b) T r a n s p o r t of t h e wind-driven c i r c u l a t l o n o b t a i n e d b y s u b t r a c t i n n the t r a n s p o r t function of the first c a s e ( Fin. 7 ) from that of the s e c o n d c a s e s h o w n in (a).

125

I

F i g . 1 1 . i a ) S e a s o n a l change i n t h e upper l a v e r v e l o c i t y i n t h e t h i r d c a s e w i t h t h e s e a s o n a l channe c o n f i n e d t o t h e w e s t e r n p a r t o f t h e i n f l o w . ( b ) D i f f e r e n c e o f t h e v e l o c i t v v e c t o r s . ( t h i r d c a s e ) minus ( s e c o n d c a s e ) .

126

Fin. 12. (a) Same as Fin. 7 but for the third case.(b) D i f f e r e n c e of transport function , lthird case1 minus lsecond casel. (2)

In spring t o summer.

the increased inflow f l o w s along t h e Japanese Coast

and forms the TPB. But most of t h e increased inflow does not flow out directly through t h e Tsugaru Strait. summer.

t h e WBCB

but

forms an anti-cvclonic circulation.

In late

is formed g r a d u a l l y by the westward shift o f t h e k i n e t i c

energy of the TPB.

T h e s e r e s u l t s are not simulated

in Sekine(1986),

which

suggests that t h e intensified inflow in s u m m e r h a s an important role in t h e circulation in t h e J a p a n Sea. (3)

In

autumn,

the

flow

pattern

is chracterized

by

the

process

toward

accomplishinn the oceanic response t o t h e increased inflow in summer. In early autumn, formed a r e the central f l o w to the west o f the TPB in the u p p e r layer and a westward current alonn t h e n o r t h w e s t e r n bottom s l o p e of t h e Y a m a t o Ridne. In late autumun. t h e TPB and o f f s h o r e central flow are absorbed in t h e WBCB. T h i s flow pattern is maintained u n t i l the strong anti-cyclonic circulation is formed bv strong wind s t r e s s in January. ( 4 ) O n t h e whole. the total flow pattern in the J a p a n S e a is influenced by the seasonal c h a n g e in wind stress and in- and outflow. 'The Tsushima Current h a s three p o s s i b i l e branches in t h e s o u t h e r n Japan Sea: t h e first branch along t h e Japanese Coast formed by the topographic guiding effect o f t h e continental slope. t h e s e c o n d branch t o the west of t h e first branch formed by westward shift o f t h e k i n e t i c energv of the first branch. and t h e third branch along the Korean Coast formed by t h e planetary B effect (westward intensification). (5) T h e seasonal variation in t h e current s y s t e m characterized by t h e wind-driven circulation

of

the

in winter and

by

Japan

Sea

is

the enhanced

increased in-and o u t f l o w in summer. The spring and autumn flow patterns a r e o f transitional character froa w i n t e r t o s u m m e r regime and from s u m m e r t o winter regime. respectively. A next problem is to investigate the role of thermohaline circulation. However. modeling o f t h e thermohaline c i c u l a t i o n is very difficult b e c a u s e of its s e n s i t i v e dependence on t h e value of t h e coefficient of vertical eddy diffusivitv (e.R.. F.Brvan. 1987): if a larner (smaller) value of the coefficient

is assumed.

the thermohaline

(wind-driven) ciculation c o n t r o l l s

the current svstem. Another problem is that t h e frontal structure o f t h e current system c a n not be properly handled with t h e layer model. The circulation in Japan Sea will be more realistically simulated by use o f level mode Is. ACKOWLEDGMENTS The author would like t o express h i s s i n c e r e thanks t o P r o f e s s o r K. T a k a n o o f Tsukuba University for his critical readinn o f the Hanuscipt. Thanks are extended t o Dr. M. Matsuvaaa of T o k y o Universitv o f Fisheries. Dr. J . H . Yoon and Dr. M. Kawabe of University of Tokyo for t h e i r valuable comment and discussions. REFERENCES Brvan.F. 1987. P a r a m e t e r sensitlvitv o f primitive equatlon o c e a n trenerai circulation aodels. J. Phys. Oceanozr.. 17: 970-985. 1noue.N.. Miita. T. and Tawara. S . 1985. Tsushima Strait 11 Physlcs. In H. Kunishi et at. (Editors), Coastal oceanonraphy of Japanese Islands, Tokai University Press. 914-933 (in Japanese). Japan MeteorolonicaI Anencv 1972. flarine meteorological study o f the Japan Sea. Technical Report of the Japan ~ e t e o r o ~ o a i c a Anency. l 80: 1-161. Kawabe. M. 1982a. Branching of the Tsushima Current in t h e J a p a n Sea. P a r t I Data analysis. J. Oceanonr. SOC. Japan. 38: 95-107. Kawabe. M. 1982b. Branchinn of the Tsushima Current in t h e Japan Sea. P a r t I 1

128 Numerical experiment. J . Oceanogr. SOC. Japan. 38: 183-192. Kutsuwada. K. and Sakurai. K. 1982. Climatolonical maps of wind s t r e s s field over t h e north P a c i f i c Ocean. Oceanogr. Mag., 32: 25-46. Matsuvama. M.. Nazumi. T. and Takahata. T. 1986. S o m e characteristic velocitv fields near the Tailma Coast. Bull. Coast. Oceanonr.. 23: 129-138 (in Japanese). Morivasu. S. 1972a. T h e Tsushima Current. In. H. Stommel and K. Y o s h i d a (Editors). Kuroshio - Its physical aspects. Univ. Tokvo Press. Tokyo. 353369. Morivasu. S. 1972b. Hydrography of the Japan Sea. Marine S c i e n c e s ( K a i v o Kagaku. 4. 171-177 (in Japanese w i t h English abstract). Pedloskv. J. 1979. Geophvslcal fluid dynamics. S p r i n x e r Verlan. 6 2 4 PP. Rhines. P.3. 1970. Edge-bottom and Rossbv wave in a rotaing s t r a t i f i e d fluid. Geophis. Fluid. Dyn.. 1 : 273-302. Sekine. Y . 1986. Wind-driven circulation in the J a p a n Sea and its influence on the branching o f t h e Tsushima Current. Prog. Oceanogr.. 17: 297-313. Sekine. Y . 1989. O n the seasonal c h a n n e in t h e in-and o u t f l o w of the J a p a n Sea. Prog. Oceanogr.. 21: 269-279. Suda. K. and Hidaka. H. 1932. The r e s u l t s o f t h e oceanographic observation on board R.M.S. Svnmpu Maru in t h e southern part of the J a p a n Sea in t h e s u m m e r of 1929. Part I. J . Oceanogr. Imper. Mar. Obs.. 3: 291-375 (in Japanese). Suginihara. N. 1981. Ouasi-qeostrophic waves in a stratified o c e a n with bottom toponraphv. J. Phvs. Oceanogr.. 1 1 : 107-115. Uda. M . 1934. The results of s i m u l t a n e o u s oceanographic investinations in t h e Japan S e a and its adjacent w a t e r s in May and June.1932. J. Imp. Fish. Exp. St.. 5: 57-190 ( in Japanese with English abstract). Uda. M . 1936. R e s u l t s of s i m u l t a n e o u s oceanographic investigations in t h e Japan S e a and its adjacent waters during O c t o b e r and November. 1933. J . Imp. Fish. Exp. St.. 7: 91-151 ( in Japanese with English abstract). Yoon. J. H. 1981a. Numerical experiment on the circulation i n t h e Japan S e a Part I Formation of east Korean warm current. J. Oceanogr. SOC. Japan. 38: 43-5 1 Yoon. J. H . 1981b. Numerical experiment on t h e circulation i n t h e Japan S e a Part I1 Influence of seasonal variation in atmospheric c o n d i t i o n on t h e Tsushima Current. J . Oceanogr. SOC. Japan. 38: 125-130. Yoon. J. H. 1981~. Numerical experiment on the circulation in t h e J a p a n S e a Part 1 1 1 Mechanism o f nearshore branch of t h e Tsushima Current .I. Oceanogr. SOC. Japan. 38: 125-130.

.

129

ON THE INTERMEDIATE WATER IN THE SOUTHWESTERN EAST SEA (SEA OF JAPAN)

C.H. KIM'. H.-J. LIE' and K . 4 . CHU2 'Korea Ocean Research & Development Institute, P.O.Box 29, Ansan, Seoul 425-600, Korea 'Hydrographic Office. P.O.Box 56, Incheon 400-600. Korea

ABSTRACT The spatial and temporal variability of the East Sea Intermediate Water (ESIW) in the East Sea is examined by incorporating the CTD data, taken over three different surveys in August 1986. December 1987 and May 1988, with the historical hydrographic data obtained bimonthly. The ESIW is clearly identified by the salinity minimum and the dissolved oxygen maximum off the mid-east coast of Korea and in Korea Strait in August 1986. A core of the salinity minimum water appeared at the coastal region in December 1987, offshore in May 1988, and at both region in August 1986. In the historical bottle data location of salinity minimum core is not so clear as CTD data. Temporal variation in the properties of the ESIW suggests that this water appears interannually. 1. INTRODUCTION Water masses in the East Sea (Sea of Japan) have been discussed by various authors (Akagawa, 1954 ; Kajiura et al., 1958 ; Gong and Park. 1969 ; Miyajaki and Abe, 1960). Moriyasu (1972) has classified four water masses according to Kajiura et al. (1958) : Surface and Intermediate Waters, the Proper Water. and the fourth water mass. Along the east coast of Korea a cold water exists forming a strong front with a warm water camed by the Tsushima Warm Current flowing north. The cold water is known to show year-toyear variation in the width and strength (Gong and Son, 1982 ; Gong and Lie, 1984). Kim and Kim (1983) have observed that the cold water of 2-5 ? is characterized by high concentration

of dissolved oxygen(> 6.5 mlh) and low salinity(< 34.00 %o). They have suggested that this water is brought to the southwestern East Sea by the North Korean Cold Current, which has been believed to flow southward along the northern Korean coast (Uda, 1934). This cold water is a distinct water mass, which can he identified by the salinity minimum and the dissolved oxygen maximum in the vertical sections (Kim and Chung, 1984). The cold water mass off Korea is also found to have the same characteristics as those of the fourth water mass that has been observed at the polar front in the middle of the East Sea (Kajiura et al.. 1958 ; Moriyasu, 1972). This fact suggests that the cold water is widely distributed in the East Sea. For that reason Kim and Chung (1984) have proposed to call the cold water as the East Sea Intermediate Water (ESIW). Compared with the Tsushima Warm Current Water (TWCW) and the Proper Water, the ESIW takes a small portion of the total water volume in the East Sea (Yasui et al., 1967). However, it may play an important role in the circulation of the East Sea. Unfortunately, little is known about the spatial and temporal variability of the ESIW ; Kim and Kim (1983) concerned mainly on the existence and origin of the ESlW ; Kim and Chung (1984) focused on the characteristics of the ESIW. In this paper we have examined the spatial and temporal variations of the ESIW by analyzing

130

I

1

z

I

1

m

Line 105

4

;4

47

'I

*:I

.ULLEUNGC 45 44

42

e

43

m

KOREA 31

30 29

28 26

b m

27

'2 I

16

l5

1

Aug. 5-15,ISE (St.1 - 3 . 5 2 ) z Dec.12. 1987 May 28-29,19[ (St.3- St. I l l A Current Me asurernenl

.I b ~ 13

129"

130'

%i

m

13PE

Fig. 1. Locations of hydrographic stations occupied in August 1986, December 1987 and May 1988. The stations between Mukho and Ulleung Island are the same with line 105 which is occupied bimonthly by the Fisheries Research and Development Agency of Korea.

131 both CTD data, observed by the Korea Ocean Research & Development Institute (KORDI). and historical hydrographic data collected by the Fisheries Research and Development Agency (FRDA) of Korea.

2 SPATIAL DISTRIBUTION O F THE INTERMEDIATE WATER 2.1 August 1986 CTD casts were made over three different periods in the southwestern East Sea August 1986, December 1987 and May 1988 (Fig. I). Vertical profiles of temperature, salinity and dissolved oxygen along the 36" 30"

line (stations 35-30 in Fig. 1) show that isotherms are almost parallel

to the sea surface and a strong seasonal thermocline is formed at the depth of 10-30 m (Fig. 2a).

The upper surface water has high temperature (16-25 "c) and low salinity (32.40-34.00 %o). This water is originated from the Tsushima Warm Current Surface Water (TWCSW) (Lim and Chang, 1969). The Tsushima Warm Current Middle Water (TWCMW) below the upper surface water appears at about 50 ni with two separated cores of high salinity (> 34.40%0

and low concentration of

dissolved oxygen (< 5.6 ml/l). A permanent thermocline is found at 70-120 m underneath the TWCMW. Temperature below 150 m is lower than 2

e , so that

the Proper Water seems to occupy the whole water column

under the TWCMW. However, the salinity profile (Fig. 2b) distinguishes a salinity minimum layer (< 34.05

%o)

from the Proper Water (< 2 2 . 34.08-34.09 %a). The water of low salinity corresponds

to the ESIW mentioned by Kim and Chung (1984). The location of the salinity minimum layer

of the ESIW coincides with that of the dissolved oxygen maximum layer. The thickness of the salinity-minimum and dissolved oxygen-maximum layer (< 34.05 %o,

>

7.5 mlA) is about 100-130

m in the vertical section. Two cores of the salinity minimum and dissolved oxygen maximum appear : one in the coastal region and the other in the offshore region. The offshore core is more distinct in the salinity and dissolved oxygen character. The salinity minimum layer of the ESIW is also observed clearly in other sections. Fig. 3a shows the depth where the lowest salinity is observed in the salinity minimum layer and Fig. 3b is the spatial distribution of the lowest salinity at the depths corresponding to Fig. 3a. A core of saIinity minimum, less than 33.98 %o, appears at 130 m near Mukho and at 200 m depth west of Ulleung Island. In the south of 37" OO'N a salinity minimum core of 34.00-34.02

%a,

which

is slightly higher than at northern section, appears at depths of 130-140 m near the coast. Cores of salinity minimum also exist at the stations east of 130" 30' E at 120 m depth of station 15.

155 m depth of station 27 and 160 m depth of station 30. The ESIW extends to coastal region along the east coast of Korea and to the northern end of Korea Strait. As the hathymetry becomes shallow toward Korea Strait, the ESIW appears near the bottom. Salinity minimum layer does not appear outside the boundary shown as a broken line in Fig. 3. In the southwestern area of Ulleung Island the TWCMW occupies from 20 m almost to the maximum sampling depth of 250 m. Therefore, no salinity minimum layer is found within the sampling depth in the area. 2.2 December 1987 and May 1988

It is worthwhile to examine the spatial distribution of water masses in other seasons. Fig.

132 U

50

100

-E

__--

I50 -?\

200

I

250

I3534

-E

33

32

20c

I 250 fi 300

w

n 400

500

\

!

)O"i

401 I-

On (rni/i)

\

(Aüg. 940,1986). 50,

O_ioOKm

-_

-

31

30

Iw

Fig. 3. (a) Depth distribution of the lowest salinity in the salinity minimum layer in August 1986 and (b) its spatial distribution at the depths corresponding to (a). Isobath of 200 m is indicated. 133

134 4 is the vertical sections of temperature and salinity along the 37” 33”

(line 105 in Fig. 1) surveyed

in December 1987 and May 1988. The TWCSW composed of well-mixed water (14-16 “c, 33.9034.00 % o ) is lying from the surface to a depth of 100 m (Fig. 4a and 4b). Around the permanent thermocline there exists the saline TWCMW (8-14 “c, 34.20-34.40 % o ) . Though the dissolved oxygen was not measured, the ESlW (0.7-4 ‘c , 33.96-34.02

%a)

can be readily identified by the existence

of salinity minimum layer which is found between the TWCMW and the Proper Water of homogeneous salinity (34.02-34.03 %o).

The salinity minimum layer deepens from the coast to

offshore. in May 1988 the ESiW(1-4 “c,

J X + v HJY h

JT JX

-A

J T , ah

--)- Jx

Jx

rv'T,I JT,

+(---)-I Jy

Jy

ah Jy

281

Double integration of Eq. (12)with respect to I,first from 0 to q, then from 0 to 1, yields: 1 -J T , aT, JT -JT JT JT ( s - K r ) [ y + U ~ + V ~ - - A ~ V ’ T ~ I + KJt, [ +- u’ A-J X+ v L - AJy JY -T , - T n ah ah ah 1 2 A , aT, a T ~ah 2K7 -{ (- + U- + V- - A ,.V h)l--I- OX JX H-h Jl JX Jy

=[(ax

+

T ~ - T ah~ 2 [(GI ( H - h)’

JT J T ~ ah +(2 --)-1-2A JY JY JY I

-4 0

~

-

T s - T T , ah a h - 1 ) e 7 d ~ i A r[(GI ( H - h)’

u =(Q,

where

vlT 1

- i Q d q ) / Q,,

ah +(G) 11+[2K,-a, 1

+ (JhG z)I-= V Q r

=o

Q = (w”+ R). Simultanous equations (Eqs. (13) and (14)) to-

0

gether with Eq. (1 I ) give: JT,, JT, -+ UJt

ax

aT + YH - A,v’T,, ay ‘ h

=C

,

z -C

+ C,A.=

T - T ahi N), ( H - h)’

aha +(GI 1

QS-Qh

l

h

(15)

and

C, = ( 2 K Y - K,K,C, = [2a,K+,

2K

1

- K ) / (KrK,) 2 7

rc-,

a K

+ 4K

I

J ( q- 116 ,dq] / (K 7 K T ) 0

C, = ( K , v - K ~ ) / ( K , K , )

c, = (-21 K 1 - KO and A = 0 when x < 0. U . = ( T /~ pw)”*the friction velocity of sea water. In order to determine u, Eqs. (13) and (14) are transformed into: d[l - K , + K , n + K , B

+ (2Kr -n,)a]=

1

(21)

Eqs. (21) and (22) give: 1 u = [- - KY(1- n - 2 8 2

I

201)- (a1 + 4 j ( q - l)Ordq)al/ [I - K,(I - n -8 0

- 2a)-

a,a].

By use of historical temperature profile data and meteorological data, and corresponding u, v,

(24)

283

Qs and Qh calculated from Eqs. (5)-(7), (IS), and (191, y can be determined with Eq. (24). The average value of y is 0.06 in the warming period and 0.53 in the cooling period in the Bohai and Huanghai Seas. Then the numerical prediction model is constructed by using Eqs. (5) to (7),(ll), (IS) and (16). At the closed boundary, V, = 0, Ts =TH,h = H, Q, = Qh = QH = 0. At the open boundary from the Changjiang river mouth to Chejudo island, au / aX = 0, aV / ay = 0. The sea level ( at the closed boundary is specified by: n

c=

c ~ ( x , Y ) , C o 4 U -l t

(25)

R(X,Y),I

I- I

where i = 1,2, ... ... ,n denotes the serial number of tidal constituent, Hiand giare the harmonic constants of ith tidal constituent, ci is angular velocity. The initial conditions are given by:

u=v=O, C=O, h=h, Ts=T,. TH=Tho, whent=t,,. 3 NUMERICAL MODEL The parallel of 36 N is taken as x axis (positive eastward), and the meridian of 120 E as y axis (positive northward), the mean sea level a8 z = 0 (positive upward). 3.1 DIFFERENCESCHEME In order to get a finite difference analog of the prediction model in differential form, we chose Platman's alternative grid, and use alternating direction implicit method (ADI) (Leendertse, 1967; Zhang et al., 1983, 1988). The features of this method are: (1) the calculation is simple because an implicit differential form is alternatively used in one direction when estimating the variables in x and y directions. (2) an implicit and an explicit differential form are alternatively used in x and y directions that make the calculation stable and quick convergent. In the meantime a hydrodynamic numerical method with temporal forward difference and special central difference (HN)is used to disctetize the Eqs. (It), (15) and (16) describing the vertical temperature structure. We evaluate ( ,h. T , That a grid point (i. 2, u, w,at a grid point (i+l / 2, 3, v, w,at a grid point (i, jtl / 2) and H at a grid point (i+l / 2, j+l / 2). O

0 1 0 1 0 - + - + O

-

J

.

I

+

1 - 7 0

1

-

t

j-1

-

i 4

O -

l

G

t

-

0 1 0 1 0 -

f

0

1

-

0

+ '

Fig. 3. Definition for computed grid points and variables.+-(.

0

h, Ts,TH; --u, w,;

I-v,~,;

0-H

284 3.2 DIFFERENCE FORM Variables U’, V’, Ckareobtained from Eqs. ( 5 ) to (7) at k A t intervals. The appendix gives the implicit differential form for C and u, and explicit differential form for v from k A t to (k+1/ 2)A t; the implicit differential form for and v, and explicit differential form for u from (k+l / 2) A t to (k+l)A t; and the direrentialform ofEqs. (ll), (15) and (16)from k A t to (k+l)At. 4 APPLICATION TO TIDAL PREDICTION The wind speed w and air temperature T,,are known. The grid spacing is 20 km. The time step is 900 seconds. The basic parameters are specificed as in Table 1. TABLE 1 The relevant parameters in the model (cgs)

I

I

PI

I

I .229 x lo-)

I

CD

c,

0.938

K,

a

0.25 x lo-)

K,

g

I

980

I

I

(0.8719+0.000704W~0)Xlo-’

I

10’ 1 / ME’’‘ ,M = 0.016-0.018

I

* W,, means the wind speed at 10 m above the sea surface. In the forecasting procedure, first of all, we calculate the current field from the given wind field, then put the current field and the given air- temperature field into the governing equations to predict h, Tsand TH.Computation was made on IBM-4381 computer. 5 THE RESULTS AND DISCUSSION OF TRIAL PREDICTION Because of the limited observations in this area, we only take the wind, air temperature and sea surface temperature data during FGGE as known variables to forecast the vertical temperature structure. The period of validity is about 4 days. The surface temperature (Ts), thickness of the upper homogeneous layer (h), bottom temperature (TI(), meridional distribution of temperature along 123.5 ’ E and vertical temperature profiles on July 8, 1979 are presented in Figs. 4 to 8. These figures show that the sea surface temperature in the study area is 22-23C in most areas except that the SST in Haizhou Bay is higher than 25C and there is a cold eddy north of Cheng Shantou (Fig. 4). There are four areas where the thickness of the upper homogeneous layer is great (Fig. 5): the first one is located in the central Bohai Sea (greater than 5 m), the second one lies in the north ofthe northern Huanghai Sea Cold Water (greater than 15 m), the third one is at about 34.5 N, 123 ’ E (greater than 10 m) and the last one in the Huanghai Trough area has the greatest thickness (20 m-25 m). The thickness ofthe upper homogeneous layer reaches 10 m off the coast of northern Jiangsu province due to strong mixing and is about 5 m in the other areas. The temperature profile from the trial prediction (Fig. 8a) clearly shows the upper homogeneous layer, thermocline and deep layer. In order to verify the reliability of the trial prediction, we roughly compare the calculated results with observed data obtained from a simultaneous hydrographic survey (Fig. 8b). Comparison shows that: ( I ) the temperature profile of trial prediction coincides with the observation north of 36 N, but there is a discrepancy between them in the deep layer south of 36 ‘ N , which might be due to a large meridional distance and asynchronous observational data between stations. (2) the thermocline obtained from the trial prediction is thicker than the observed one and the predicted temperature is vertically homogeneous in the thermocline, wbicb may be due to the assumption that the current velocity is vertically homogeneous that induces the strong mixing between thermocline and its lower boundary. So three dimensional numerical prediction method is needed to put forward in the future.

285

30 *

Fig. 4. Sea surface temperature ( C ) from the trial prediction (July 8,1979 in the Bohar and Huanghai Seas.

Fig. 5. Thickness (m)of the upper homogeneous layer from the trial prediction (July 8, 1979) in the Bohai and Huanghai Seas.

30

Fig. 6 . Bottom temperature ( C , July 8,1979) in the Bohai and Huanghai Seas.

Fig. 7. The vertical temperature profile from trial prediction(Tp) and observation (T,) at some grid points in the Bohai and Huanghai Seas. a-j are the station numbers shown as in Table 2.

286

Fig. 8. Trial-predicted (a) and observed (b) rneridional temperature (C)at 123.5 ' E. Table 2 Characteristics of predicted (T,) and observed (T,) at 10 grid points in the Bohai Hunghai Seas

ACKNOWLEDGEMENTS The authors wish to thank Dr. Takano for his valuable suggestions and thanks go to Ms.Zhang Hongnuan for drawing the figures. This work was supported by the National Planing Committee, PRC.

287 Appendix: The implicit differential from For C and u ,and explicit differential from for v from k A t to &+1/ 2)At for Eqs . (5)-(7) are as follows:

where

where

288 the implicit differential form for ( and v, and explicit differential form for u from (k+l / 2)At to (k+l)Atare as follows :

where

where

289

vn

Velocity comonents F,T,Y, and (26)-(31) denote the average values at the calculated points, and d, is two times average depth of the point (i, j ). The differential forms of Eqs. (11),(15) and (16) from K A t to ( k + l ) A t are as follows :

290

REFERENCES Halpern, D., 1974, Observation of the deepening of the wind-mixed layer in the northeast Pacific Ocean. J. of Phys. Oceanogr., 4,454--466. Japan Meteorological Agency, 1985, The results of marine meteorological and oceanographical observations. No. 78. Kalazkii, V. E., 1978, Simulation of vertical thermal structure for the Ocean active layer. Leningrad (in Russian), 1978. Kalakzkii, V. E. and E. S. Nesterov. 1980, Numerical prediction of the ocean thermal structure with the influence of atmospheric process in weather scale. Tr. GMC SSSR (in Russian), No. 229; 37-44. Kitaigorodskii, S. A. and V. Z. Miropolskii, 1970, On the theory of the ocean active layer. LEV. Akad. Nauk SSSR, Ser. F A 0 (in Russian), No. 6,177--188. Kitaigorodskii, S. A. 1977, Dynamics of seasonal thermocline. Okeanologiya (in Russian), No. 4,6-34. Kraus, E. R. and J. S. Turner, 1975, A one-dimensional model of the seasonal thermocline II : the general theory and its consequences. Tellus, No. 19,98--106 Kraus, E. R., 1977, Modelling and prediction of the upper layers of the ocean. Pergmon Press. 1977. Leendertse, J. J., 1967, Aspects of a computional for long period water-wave propagation. Memorandum, RM--5394--PR, Rand Corporation, May, 1967. Malkki, P. and R. Tamsalu, 1985, Physical features of the Baltic sea. Finnish Marine Res., No. 252, 50--67. Murakami, M. et al., 1985, A numerical simulation of the distribution of water temperature and salinity in the Seto Inland Sea. J. of Oceanogr. SOC.of Japan, No.41,213--224. Nesterov, E. S. 1978, Numerical prediction for the thermal features of the upper layer in the north Atlantic Ocean. Tr. GMC SSSR (in Russian) ,No. 200,22-29. Nesterov, E. S., 1986, Response of upper ocean to the temperate zone cyclone. Tr. GMC SSSR (in Russian), No. 281.24-34. Ocean Group of the Science and Technology Commission, PRC, 1964, Comprehensive marine investigation report. No. 2 and 5 Omsteds, A. et al., 1983, Measured and numerically-simulated autumn cooling in the Bay of Bothnia. Tellus, No. 35a, 231-240. Resnanskii, U.D. and E. V. Trosnikov, 1980, Parametrization in the ocean active layer developed from the simulating zonal atmospheric circulation. Tr. GMC SSSR (in Russian), No. 229, 18--31. Resnanskii, U. D.. 1983, The influence of current on the evolution of features in the ocean active layer. Tr. GMC SSSR (in Russian), No.282,23--33. Resnanskii, U. D., 1986, Numerical experiment in the ocean active layer considering the term of space variation. Tr. GMC SSSR (in Russian) No. 291,3--14. Wang Zongshan, 1983, Asimplifed method of computing heat budget at the 8ea surface. Mar. S. Bull. (in Chinese), No., 2,22-25.

291

Wang Zongshan and Zou Emei, 1986, A parametric modle for thermal structure features of the ocean upper layer. Acta Oceanol. sinica, No. 5,16--21. Wang Zongshan, Xu Bochang, Zou Emei, Yang Keqi and Li Fanhua, 1990, A study on the numerical prediction model for the vertical thermal structure in the Bohai and huanghai seas ( I )---One-dimensional numerical prediction model. Acta Oceanol. Sinica (in Chinese), 14(2), in Press. Xu Bochang et al., 1983, Simulation of vertical distribution of sea-water density (ul) in the shallows. Mar. S. Bull. (in Chinese), No. 1,9-12. Xu Bochang et al., 1984, Simulation of vertical distribution of water temperature in the shallows. Mar. s. Bull. (in Chinese), No.3,1-6. Zbang Yangting and Wang Yijiao, 1983, Simulation of wind field and numerical computation of storm surge in the Bohai sea. Acta Oceanol. Sinica (in Chinese), 5(3), 261--272. Zhang Yangting and Wang Yijiao, 1988, The properties and the numerical modelling of the Typhoon surge of Huanghai sea. Preceedings of third science meeting of Chinese Oceanol. and Limnol. SOC. (in Chinese),57--63.


293

DEVELOPMENT OF TOWED VEHICLE SYSTEMS FOR ACOUSTIC DOPPLER CURRENT PROFILER

W. KOTERAYAMA, A. KANEKO, M. NAKAMURA and T. HORI Research Institute for Applied Mechanics,Kyushu University, Kasuga 816 (Japan)

ABSTRACT Two different types of towed vehicle, EIKO and DRAKE have been developed to house an ADCP (acoustic Doppler current profiler). EIKO is a simple and stable vehicle without any mechanism for the motion control. The depth and roll of DRAKE are controlled by the main wings and horizontal taiI wings. The contours of EIKO and DRAKE are designed so as to increase the hydrodynamic damping force and t o decrease the unstable roll moment acting on them. Structures of the two vehicles and results of on-site experiments for confirming their performances are described. The experiments showed that EIKO and DRAKE with ADCP are suitable for measurements of the ocean current. 1. INTRODUCTION

The acoustic Doppler current profiler (ADCP) measures the detailed vertical profile of ocean currents. Its accuracy is assured when mounted on the sea bottom (Lhermitte, 1982) or a stable platform. Some oceanographers have attempted to use a shipboard ADCP (Joice et al., 1982) which enables to collect spatially high dense data for drawing the detailed map of the velocity distribution in an ocean current. They reported that motions of the ship due to surface waves, the cavitation noise generated by the ship’s propeller and bubbles entrained under the hull had serious adverse effects on the ADCP data quality. These effects may be reduced by using the underwater vehicle. The underwater vehicle which could completely exclude these negative elements would be a non-tethered free-swimming one, but the speed and the radius of operation of the free-swimming vehicle are restricted owing to a difficulty in the power supply. Since we wanted to collect ocean data over wide areas of the sea we selected a towed vehicle system because of its high speed mobility. Two towed vehicles, EIKO and DRAKE were developed. EIKO is simple and light and has no motion control system

.

I t s maximum towing speed is 8 knots and its maximum

working depth is 10m. EIKO fulfills its function when used behind a small ship. DRAKE is developed for high speed towing of over 10 knots and its maximum towing speed is basically unlimited; the set point of submerged depth can be varied from 0 to 300m. The main wings and tail wings are controlled automatically to keep a desired depth without roll. T h e towing point is determined on the basis of theoretical analyses so that the trim and pitch angles are minimized under any conditions of the submerged depth and towing speed.

294

Cable Winch

w

Towed Vehicle

/

OUXE

-= Acoustic Doppler z Current P r o f i l e r -

uI

Control System

Fig.1. Schematic diagram of towed vehicle-ADCP system I n this report, the structures of EIKO and DRAKE and the results of on-site experiments for confirming their performances are described. 2. CONCEPT OF THE TOWED VEHICLE-ADCP SYSTEM

Figure 1 is a schematic diagram of the towed vehicle-ADCP system for ocean measurements. A CTD-sensor and an ADCP are housed in the vehicle. We chose a product of R D Instruments (MODEL RD-DR0150) as the current profiler. The length is about 1.8m and the weight in the air is 80kg. When mounted on the sea bottom or on a stable platform, the ADCP produces accurate ocean measurements (Lhermitte, 1982). When housed in a vehicle which oscillates by external forces, the correct measurement of ADCP's attitudes is essential because the backscattered acoustic signals are transformed into three dimensional velocity relative t o the ADCP in the instrument on the basis of the measured data of attitudes and direction. Pendulum type tilt meters are used to measure the attitudes. By this type of tilt meter the distinction between the surge acceleration and the pitch or the sway acceleration and the roll is basically impossible. Such accelerations are induced by tension variations of the towing cable caused by the motions of the towing ship or shedding vortices from members of the towed vehicle. Therefore the accelerations of motions of the vehicle must be minimized. We have developed towed vehicle systems for the purposes o f ;

-

measuring ocean currents accurately even under severe sea states,

-

collecting data a t a wider range of the depth than the ADCP normally covers, reducing the influence of cavitation noise generated by the ship's propeller,

*

eliminating interference caused by bubbles entrained under the ship's hull and

-

being able to use ADCP with any research vessel. All purposes but the first can be achieved simply by utilizing a cable of proper length

295

to tow the vehicle. To successfully achieve the first purpose, the vehicle's hull shape and control system were designed to minimize the oscillations of the towed vehicle in water. In addition, mother ship speeds should be measured accurately to calculate the velocity relative to the Earth by subtracting the vehicle's forward speed, of which the time average can be considered to be equal to that of mother ship speed, from the velocity relative t o the towed vehicle. The ADCP can measure its own velocity relative to the Earth by bottom tracking in shallow water, but in deep water other navigational methods must be used. The Loran-C system was adopted in these experiments. We use currently the Global Positioning System, which is much more accurate than the Loran-C. 3. STRUCTURE OF EIKO EIKO is pictured in Fig.%;the hull is made of fiber reinforced plastic (FRP). I t s principal features are shown in TABLE 1. TABLE 1 Features of EIKO

1Om L=%m.W=0.78m.H=O.55m 160kg (with ADCP)

Operating depth Dimensions Weight in air Weight in water Towing speed Instrumentation ~

~~

L V e r t i c a l Stabilizer

0

-

0 8 knots ADCP. CTD sensor

\w

Horizontal Stabilizer 1

Fig.% Figure of EIKO.

A.D.C.P.

296

The submerged operating depth (10m) is chosen considering that the draft of the largest observation ship in Japan is not deeper than 7m. It is enough to avoid the effects of the mother ship on measurements. The lighter EIKO is, the easier it handles, but some bumping of EIKO against the mother ship is inevitable in deployment and retrieval. Therefore, EIKO should be of sufficient strength to withstand this without damage. Frames of EIKO are counted as longitudinal strength members and also expected to act as a shock absorber for the ADCP. While the deployment and retrieval are easy with excess buoyancy, the excess buoyancy would cause the coupling motion of the surge and heave, so that it is made neutral in water. In the body of EIKO, the buoyant material, polyurethane foam, was filled. This has now been replaced by high pressure syntactic foam because EIKO sometimes goes down deeper than the designed depth during deployment and retrieval. The lower part of the frontal area of EIKO is designed so that it does not obstruct acoustic transmission of ADCP. A depressor is set a t its tail end for balancing the heavy head of the ADCP. It makes the trim of EIKO minimum by its downward lift force. The downward lift force can be controlled by changing the angle of the depressor. A horizontal stabilizer is set to increase hydrodynamic damping force for the pitching motion and minimize the trim angle. Vertical stabilizers are expected to increase the hydrodynamic damping forces for the roll and sway motion. The F R P body with the buoyant material protects the ADCP against the impulsive force induced by the bumping of EIKO against the ship. The body shape is streamlined and serves to diminish the drag force acting on the vehicle.

Fig.3 Forces acting on EIKO The dimension and angle

p

of the depressor and other properties are determined con-

sidering the static balance of the forces acting on EIKO as shown in Fig.3, in which To is the towing cable tension, L t and Dt are the lift and drag force acting on the horizontal stabilizer, La and D, are the lift and drag force acting on the depressor, W is the weight of the ADCP in water, B is the excess buoyancy of EIKO body, and X and Y are the x- and

297

y-components of the drag forces acting on EIKO with the ADCP. At first, the static balance of the moment of Lt,D , , L , , D, , W, B, X, and Y at the towing point T are considered. These values except W and B are functions of the towing speed Urn and the trim angle 9. In addition to that, D, and L, depend also on the set angle of the depressor. When Urn and

p are given, we obtain all these forces and the trim angle 9 through an iterative procedure. Next from the balance of the force at the towing point TI the value of the tension To and the angle

80

of the towing cable at the towing point T are calculated. The profile of the

towing cable and submerged depth of the vehicle are determined by substituting To and

80

into the inverse catenary theory. Until the depth reaches to the desired one, the calculation

is repeated by changing the value of p. The detail of the calculation method is shown by Hori et al., (1988).

4. STRUCTURE OF DRAKE Figure 4 and Table 2 show DRAKE and its principal feature. The hull is made of fiber reinforced plastic (FRP). The shape is streamline contour with its deep vertical contour designed to increase the damping- and added mass-force, thereby to reduce high frequency roll which is hard to control with the horizontal tail wings. The main wings have a symmetrical aerofoil profile. The lower part housing the ADCP and CTD sensor is designed so as to not obstruct the acoustic beam of the ADCP or the flow inlet for the CTD sensor. Table 2 Principal features of DRAKE and the towing system DRAKE Operating depth Dimensions Weight in air Towing velocity Depth control Roll control Instrumentation TOWING CABLE Length Diameter Breaking tension Conductors

-

0 300m L=2m, W=2m1H=1.5m 360kg 5 12knots By main wing By horizontal tail wing ADCP, CTD sensor 800m 12.9mm 9 tons Power conductors 1 pair Signal conductors 10

~

CABLE WINCH Weight Dimensions Maximum reeling tension Maximum stopping tension

2060kg 1.8m x 1.8m x 2.0m 1.5 tons 2.3 tons

The towing cable is a double-armored one with ten signal conductors and a pair of electric conductors for power supply to the sensors. The cable dynamics is very important

298

Fig.4 Picture of DRAKE.

Fig.5 Concept of calculation method for DRAKE system.

299

for the DRAKE system because the towing cable is very long and heavy. We developed a three-dimensional lumped mass method for cable dynamics and six-degree-freedom equation for the motions of the DRAKE (Koterayama et al., 1988), of which the concept is shown in Fig.5. In the lumped mass method the cable is modeled as N-discrete masses interconected by springs. All forces such as the drag force, the added mass force, the inertia force, the buoyancy and weight acting on the cable are considered to be concentrated load on each mass, then the motions of the cable can be represented by simultaneous differential equations. The equations for the motions of DRAKE are similar to those for the airplane. The motions of the towing vessel, cable and towed vehicle are represented by a set of simultaneous equations. The hydrodynamic coefficients used in the calculation were obtained from model experiments and theoretical estimations. The body and control system were designed by numerical simulations with this calculation scheme. Koterayama et al.( 1988) described details of the calculation method. The perspective (Fig.6) shows that the impeller set at the tail end is coupled to a hydraulic pump. This provides power for the actuation of the main- and horizontal tail- wings. The idea of using a stream-driven impeller to generate hydraulic power was introduced in the development of the Batfish by the Bedford Institute of Oceanography (Dessureault, 1976).

Pressure vessel

Potent lo-meter f o r main wing angle Horizontal

Acoustic D O current prof1 er

Fig.6 Perspective of DRAKE

300

5. ON-SITE EXPERIMENTS A N D DISCUSSION On-site experiments (Fig.7) of EIKO and DRAKE were separately carried out to confirm the performances of EIKO and DRAKE and to show the capability of t h e towed vehicleADCP system in ocean measurements. The maximum towing speeds tested were 8 knots for EIKO and 1 2 knots for DRAKE; the latter speed was the maximum possible by the mother ship and the maximum towing speed of DRAKE is basically unlimited. When the towing speed of EIKO exceeded 8 knots, the angle of the heel became greater than 30 degree which is the maximum allowable tilt angle of the ADCP. Figure 8 shows the static characteristics of DRAKE obtained from on-site experiments (Koterayama et

al.,

1990). Circles indicate the experimental results of main wing angle and

the solid line is the calculated result. The squares and broken line are the results of the trim of DRAKE. The double circles and one-dash chain line are the attack angle of the main wings relative to the uniform flow. The triangles and two-dash chain line are the tension a t the towing winch. This figure indicates the accuracy of theoretical estimations for the static ploblem. The motions of EIKO and DRAKE were found to be much less than those of the mother ship. A comparison between the power spectra of motions of the mother ship and DRAKE (Fig.9) shows that the heave, roll and pitch of DRAKE are much less than those of the mother ship. We did not measure the surge, sway or yaw, but the theoretical analysis (Koterayama et al., 1988) suggests that the sway and yaw are much less than those of the mother ship; the surge of DRAKE is also less though greater than the sway or yaw. T h e theoretical analysis also suggests that the ratios of motions of DRAKE and the ship decrease with the increase of incident wave height. Figure 10 compares d a t a by the EIKO-ADCP system with data by a mooring system. Both are in fairly good agreement. Figure 11 shows the velocity distribution of the Kuroshio taken by an ADCP mounted on DRAKE. I t is presumed that the disturbances were caused by small islands. There was no loss of measured data, which is often caused by the noise generated by the ship’s propeller and bubbles entrained under the ship’s hull when the ADCP is set on a surface ship. Velocity distribution in a n ocean current is usually obtained from geostrophic calculation using the measured vertical profile of water temperature and salinity or direct measurement using the mooring system, but with these methods it. is very difficult to obtain such detailed data as are shown in Fig. 11. The mooring system has an advantage over the ADCP carried by a vehicle or ship from the point of view of continuous measurement (Takematsu et al., 1986) while the ADCP boarded on the ship or vehicle is excellent in the point of spatial continuity. These two systems of measurement thus complement each other, and combined use of the two should become popular in ocean measurements. The system of ship-boarded ADCP has weak points on the data quality because the motions of the ship due to surface waves, the cabitation noise generated by ship’s propeller and bubbles entrained under the

301

EIKO (a)

DRAKE (b)

Fig.7 Views of on-site experiments of EIKO (a) and DRAKE (b).

I 0

50

100

150

200

Z(m)

Fig.8 Static characteristics of DRAKE (towing speed is 6 knots).

lo

250

302

1‘

HEAVE J

DRAKE

(EXPI-.DRAKE (CAL.) ----

0

1.0

30 2D w (rad Isec)

PlTCH

Fig.9 Power spectra of motions of the mother ship and DRAKE (towing speed is 11 knots).

0

50 V(cm/s)

I”’“’ -ADCP 0 ACM

100’

0

Fig.10 Comparison of data obtained by the EIKO-ADCP system and mooGng system with Aanderaa current meters on the current speed and direction at 34 04’N, 129’32’E. ( water depth is about loom, towing speed is 6 knot).

303

hull have serious adverse effects. T h e results of on-site experiments shown in Figs.10 and 11 indicated that the well-designed towed vehicle can exclude these negative elements, but more deeper studies by means of numericaI simulations and on-site experiments (Kaneko et al., 1990) are needed for the quantitative evaluation of the accuracy of measurements by an ADCP-towed vehicle system.

Acknowlegement This work is a part of the Ocean Research Project of the Research Institute for Applied Mechanics financed by the Ministry of Education, Science and Culture, Japan. The authors thank members of the Ocean Research Group of R.I.A.M.

v - -

-.

N --,--.,-

a . EO

5

'-

i:

i c -

-

--

Fig.11 Velocity distribution of the Kuroshio obtained by the DRAKE-ADCP system 28'57'N, 130'15'E (C9). at 30 5'N, 131'3'E ( C l ) (towing speed is 9 knot, submerged depth of DRAKE is 50m, December, 1988).

304

6. REFERENCES

Dessureault, J.G., 1976. -Batfish- A depth controllable towed body for collecting oceanographic data. Ocean Engineering, Vo1.3: 99-111. Hori, T., Nakamura, M., Koterayama, W., Honji, H., and Takahashi, M., 1988. A design of the towed vehicle system for the acoustic Doppler current profiler. Transactions of the West-Japan Society of Naval Architects, No.76: 97-112. Joyce, T.M., Bitterman ,J.R. and Prada K.E., 1982. Shipboard acoustic profiling of upper ocean currents. Deep-sea Research, No.29: 903-913. Kaneko, A., Koterayama, W., Honji, A., Mizuno, S., Kawatate, K and Gordon R.L., 1990. Cross-streem survey of the upper 400m of the Kuroshio by an ADCP on a towed fish. Deep-sea Research, Vo1.37, No.5: 875-889. Koterayama, W., Kyozuka, Y., Nakamura, M., Ohkusu, M. and Kashiwagi, M., 1988. Motions of a depth controllable towed vehicle. Proc. of the 7th International Conference on Offshore Mechanics and Arctic Engineerin 423-430. Koterayama, W., Nakamura, M., Kyozuka, Y., Kastiwagi, M. and Ohkusu, M., 1990. Depth and roll controllable towed vehicle DRAKE for ocean measurements. Proc. of the First Pacific/Asia Offshore Mechanics Symposium: 257-264. Lhermitte, R., 1982. Doppler sonar observation of tidal flow. Journal of Geophysical Research, No.88: 725-735. Takematsu, M., Kawatate, K., Koterayama, W., Suhara, T., and Mitsuyasu, H., 1986. Moored instrument observations in the Kuroshio south of Kyushu. Journal of the Oceanographycal Society of Japan, Vo1.42: 201-211.

305

A S T U D Y O F THE KUROSHIO I N T H E E A S T C H I N A SEA A N D T H E C U R R E N T S E A S T OF T H E R Y U K Y U I S L A N D S I N 1988 YAOCHU Y U A N , J I L A N SU AND ZIPIN PAN S e c o n d I n s t i t u t e of O c e a n o g r a p h y , S t a t e O c e a n i c A d m i n i s t r a t i o n , P . 0. B o x 1207, Hangzhou 310012, C h i n a

ABSTRACT The inverse method is used to c o m p u t e t h e Kuroshio in t h e East C h i n a S e a with h y d r o g r a p h i c d a t a c o l l e c t e d during early s u m m e r and a u t u m n 1 9 8 8 and t h e c u r r e n t s east of t h e R y u k y u I s l a n d s vith d a t a during early s u m m e r 1 9 8 8 only. I n t h e East C h i n a Sea t h e K u r o s h i o C o u n t e r c u r r e n t is located f u r t h e r t o the vest and its t r a n s p o r t is larger d u r i n g a u t u m n 1 9 8 8 t h a n i n latc s u m m e r 1 9 8 7 a n d early s u m m e r 1988. T h e K u r o s h i o water a l s o intruded o n t o t h e s h e l f vith a r e d u c e d width d u r i n g a u t u m n 1988. T h e t r a n s p o r t of t h e K u r o s h i o i n t h e East C h i n a S e a is 23.4 and 24.3~10' m3/s, r e s p e c t i v e l y , during early s u m m e r and a u t u m n 1988. The s e a s o n a l c h a n g e of t h e K u r o s h i o a x i s at t h e s u r f a c e d o e s not c o r r e l a t e vith t h e s e a s o n a l c h a n g e of t h e p o s i t i o n of the c e n t e r l i n e of t h e K u r o s h i o width. T h e r e is a c y c l o n i c gyre A part o f o n t h e s h e l f north o f T a i v a n d u r i n g all the cruises. t h e Taiwan Warm C u r r e n t in t h e s u r v e y area h a s a t e n d e n c y to c o n v e r g e t o t h e s h e l f break. C u r r e n t s east of t h e R y u k y u I s l a n d s flow northward o v e r t h e R y u k y u T r e n c h i n early s u m m e r 1988. of which t h e t r a n s p o r t is 26.9~10' m3/s. T h e c o r e of t h i s f l o w is between 4 0 0 and 8 0 0 m depths.

1 INTRODUCTION T h e r e have been many s t u d i e s on t h e c u r r e n t s t r u c t u r e and volume t r a n s p o r t of t h e K u r o s h i o in t h e East C h i n a S e a ( G u a n , 1982. 1988: Nishizawa et al., 1 9 8 2 ; S a i k i , 1 9 8 2 ; Yuan and S u , 1 9 8 8 ; Yuan, Endoh and Ishizaki, h e r e a f t e r r e f e r r e d t o a s YEI, 1990). Dynautic c o m p u t a t i o n s g i v e t h e o v e r a l l mean v o l u m e t r a n s p o r t t h r o u g h s e c t i o n G(PN) a s 21.3~10' m5/s for t h e y e a r s between 1955 and 1978 (Guan, 1982, 1988) or 19.7~10' m3/s if d a t a f r o m 1954 t o 1 9 8 0 a r e used (Nishizawa et al., 1982). G u a n (1982, 1988) and Nishizawa et al. (1982) used t h e 7 0 0 dbar as t h e r e f e r e n c e level. However, a r e c e n t work b y Y E 1 (1990) using !.hc inverse method s h o v e d that t h e r e may still b e a s i g n i f i c a n t c u r r e n t at 700 m depth. T h e r e f o r e , t h e Kuroshio t r a n s p o r t t h r o u g h t h e East C h i n a S e a may have been u n d e r e s t i m a t e d b y p r e v i o u s a u t h o r s using d y n a m i c c o m p u t a t i o n methods. S t u d i e s on t h e c u r r e n t s east o f t h e Ryukyu I s l a n d s a r e

306

r e l a t i v e l y few. K o n a g a et al. (1980) a n d Y u a n a n d Su ( 1 9 8 8 ) p o i n t e d o u t t h a t t h e r e is a n o r t h w a r d c u r r e n t e a s t of t h e R y u k y u I s l a n d s a n d t h i s may be p a r t l y r e s p o n s i b l e f o r s t r e n g t h e n i n g t h e Kuroshio south of Japan. YE1 (1990) found that the northward currents east of Ryukyu Islands flowed over the Ryukyu Trench d u r i n g Sep.-Oct., 1987. In t h i s s t u d y t h e i n v e r s e m e t h o d is u s e d t o c o m p u t e t h e Kuroshio C u r r e n t in t h e E a s t C h i n a S e a , u s i n g h y d r o g r a p h i c d a t a c o l l e c t e d d u r i n g t v o c r u i s e s in 1 9 8 8 (Hay t o J u n e a n d O c t o b e r t o November). T h e c u r r e n t s e a s t of R y u k y u I s l a n d s a r e a l s o c o m p u t e d with early summer data only. The current structure and transport of the Kuroshio and the currents east of the Ryukvu I s l a n d s will be d i s c u s s e d .

2 NUWERICAL CALCULATIONS S i n c e t h e i n v e r s e m e t h o d w a s d e s c r i b e d in d e t a i l in p r e v i o u s s t u d i e s ( Y u a n , S u a n d P a n , 1 9 9 0 ; YEI, 1 9 9 0 1 , we w i l l n o t r e p e a t t h e d e s c r i p t i o n of t h e m e t h o d here. T h e b a t h y m e t r y , h y d r o g r a p h i c s e c t i o n s a n d s t a t i o n s , a n d c o m p u t a t i o n b o x e s a r e s h o w n in F i g . 1. T w o b o x e s a r e s e t u p f o r t h e r e g i o n in t h e E a s t C h i n a S e a w h i l e e a s t of t h e R y u k y u I s l a n d s o n l y o n e box is s e t U P . Because the w a t e r d e p t h s a r e s h a l l o w w i t h i n b o x e s 1 a n d 2, we d i v i d e t h e b o u n d a r y s e c t i o n s o f t h e s e b o x e s into t w o layers. T h e v a l u e 6t.. o f t h e i s o p y c n a t level is t a k e n t o b e 25. The water depths in 122' 1?#7 ' ,

126

128'

,

130'

132'

/5

2P

2r

2J

F i g . 1 . (a) D e p t h (in m ) . h y d r o g r a p h i c s t a t i o n s a n d computation b o x e s d u r i n g e a r l y s u m m e r o f 1 9 8 8 (Hay a n d June. 1988).

307 120

1.

Fig.

boxes

122

126

124.

(b) D e p t h (in n), h y d r o g r a p h i c s t a t i o n s a n d c o m p u t a t i o n d u r i n g a u t u m n o f 1988 ( O c t o b e r a n d N o v e m b e r . 1 9 8 8 ) .

'

-.r

25----

LJ

27

30 -

1000h

E

.

7 27 1000 -rr" 2000:

v

c, 3000: n

6

4000 1

5000;

Fig.

b

2. I s o p y c n a l l e v e l s a l o n g (a) s e c t i o n

Fa

a n d f b f s e c t i o n Fa.

box 3 a r e d e e p , so t h a t we d i v i d e e a c h b o u n d a r y s e c t i o n o f b o x 3 into f i v e l a y e r s a c c o r d i n g a s t h e o + . , v a l u e s o f f o u r i s o p y c n a l l e v e l s 25, 27, 3 0 a n d 35. F i g u r e 2 s h o v s t h e i s o p y c n a l l e v e l s a l o n g s e c t i o n s F a a n d Fo, r e s p e c t i v e l y . T h e d e p t h s o f t h e o % . * = 2 5 , 2 7 , 30 a n d 3 5 l e v e l s lie b e t v e e n 1 2 5 t o 200 I, 300 t o

3 6 0 m , 7 0 0 to 7 3 0

a n d around 1 6 4 0 m , respectively.

q

TABLE 1

T h e b a l a n c e o f t h e t r a n s p o r t in box 3 d u r i n g early s u a m e r o f 1 9 8 8 -~

layer

Ai

AI.D

_ _ _ ~

1st

2nd

3r d

4th

5th

-2.03

-3.01

-7.61

9.73

1.46

x10-3

Xio-3

x10-"

x10-"

x I o -3

-1.09

-0.68

1.50

-2.88

-2.14

total -1.46

1 0 - :3 -8.29

Note b a l a n c e o f t h e transport into t h e i-th layer by inverse method (positive into t h e layer). A l m D :balance of t h e t r a n s p o r t into t h e i-th layer by d y n a m i c method (positive into t h e layer). (units: 10'' m3/s).

Ai:

T a b l e 1 c o m p a r e s t h e b a l a n c e o f t h e transport f o r each layer o f b o x 3 by the inverse method with t h o s e by dynamic method. The r e f e r e n c e level o f no motion is assumed a t t h e bottom f o r t h e d y n a m i c method. T h e d y n a m i c method y i e l d s large unbalanced t r a n s p o r t for each layer. T h e total unbalanced t r a n s p o r t a m o u n t s t o 8 . 3 ~ 1 0 " m3/s, which is t o o large t o b e ignored. This indicates t h a t c o n s e r v a t i o n o f mass is n o t s a t i s f i e d by d y n a m i c m e t h o d , when t h e r e f e r e n c e level o f no motion is a s s u m e d a t t h e bottom.

3 THE K U R O S H I O IN THE EAST C H I N A S E A In t h i s s e c t i o n we f i r s t d i s c u s s t h e K u r o s h i o Current s t r u c t u r e and its transport in early s u m m e r a n d a u t u m n o f 1 9 8 8 , and t h e n c o m p a r e t h e s e r e s u l t s with r e s u l t s from a n e a r l i e r cruise. 3 . 1 Early summer (May-June) o f 1 9 8 8 In early s u m m e r o f 1 9 8 8 t h e K u r o s h i o h a s only

one current In Fig. 3 and following figures; c o r e a t s e c t i o n S Z (Fig. 3 ) . t h e c o m p u t a t i o n p o i n t s a r e located a t t h e mid-points between neighboring hydrographic stations. T h e c o r e is located a t point S z - 6 w h e r e t h e water d e p t h is a b o u t 1 2 8 0 a . T h e maximum velocity is f o u n d a t t h e 5 0 a level vith a magnitude o f 9 7 . 3

309

cm/s. T h e velocities in the upper 150 in o f this c o r e a r e all I t s c u r r e n t s p e e d at thr? 600, 7 0 0 and 8 0 0 greater than 8 0 c m / s . m level is 1 7 . 3 , 7 . 9 and 1 . 6 cm/s, respectively. In t h e deep

layer there is a countercurrent. T h i s c o u n t e r c u r r e n t r e a c h e s t h e m level at point S z - 7 whose maximum s p e e d ( 6 . 6 CIR/S) i s at t h e 800 in level o f point S2-7. The transport o f the K u r o s h i o through s e c t i o n S Z in early s u m m e r is 23.4~10' m 3 / s and its width is about 170 km. 600

Computation Points

2

3

3 . Velocity distribution at (a) section S 2 and ( b ) s e c t i o n (positive value d e n o t e s northward flow). (units: c m / s ) .

Fig. S4

The hydrographic s t a t i o n s a t s e c L i o n S h d o c s not reach d e e p waters. F i g u r e 3 ( b ) s h o w s that the c o m p u t e d K u r o s h i o velocities in t h e upper 1 0 0 m o f point S + - 4 arc' all greater than 80 cm/:;. T h e r e i s a c y c l o n i c c o l d g y r e to the west o f t h e K u r o s h i o o n the s h e l f north o f Taiwan (Fig. 4 ) . A c y c l o n i c g y r e i s often found here ( S u and P a n , 1 9 8 7 ; Yuan and Su, 1 9 8 8 ; Y E I , 1 9 9 0 ) . East o f this v o i d gyre and w e s t o f tho K u r o s h i o t h e r e is a northward c u r r e n t over t h e s h e l f which s e e m s t o be the o f f s h o r e branch o f the T a i w a n Warm C u r r e n t ITWCOB) (Su and P a n , 1 9 5 7 ; Yuan

310

et al., 1987) with a transport of about 0.11~10' m3/s. This may be a n underestimate, because the d i s t a n c e s between the s h e l f s t a t i o n s are too large to r e s o l v e t h e horizontal density gradient i n our and because t h e data noise is not completely removed inverse method t o make a reliable estiaate o f the small transport. I n fact. t h e inverse method is better s u i t e d for t h e computation of t h e velocity field i n a deep ocean than i n a shallow ocean.

F i g . 4. Distribution o f the transport in t h e computational region during early s u m a e r , 1988. (units: 10' n 3 / s ) .

3.2 Autumn (October-November) of 1988 In autuan 1 9 8 8 the Kuroshio again h a s only one current c o r e at section S Z . Its c o r e is located at point S Z - 7 . a s during The K u r o s h i o is s t r o n g e r than t h e early suaaer c r u i s e of 1988. during t h e early s u m m e r cruise. The velocities i n the upper 2 5 0 m at t h e c o r e a r e greater t h a n 100 cm/s and t h e maximum velocity Even i n deep layers, its of 117.3 cm/s is at t h e 100 m level. Velocities a r e still quite strong. For exaaple, at Point s z - 7 its speed is 82.4, 27.6 and 1 0 cm/s a t the 400, 8 0 0 and 1 0 0 0 level, respectively. The width o f t h e Kuroshio is about 110 km, much narrower t h a n during other c r u i s e s (September-October o f 1987 and H a y - J u n e of 1988). The transport of t h e Kuroshio through s e c t i o n S Z is 2 4 . 3 ~ 1 0 " a 3 / s (Fig. 8 ) . T h e r e is a region of strong c u r r e n t s near t h e s h e l f break around Point S Z - 4 . The maximum velocity in t h i s region is 43.9 cm/s, while the maximum

311

L

V e l o c i t y d i s t r i b u t i o n a t (a) s e c t i o n S 2 a n d (b) s e c t i o n the autumn c r u i s e (positive value d e n o t e s northward f l o v ) . (units: c d s ) . Fig.

5.

SS d u r i n g

312

velocity

in

this region during the early summer cruise is

23.9

CM/S.

1

122

124'

126

12s

1

C 3 1'

29

27

25-

Fig. 6 .

during

Distribution of the transport in the computaitonal region autumn 1 9 8 8 . (units: 10' m 3 / s ) .

There is a strong countercurrent at points S2-8 and s 2 - 9 . Its maximum speed is 4 2 . 1 cm/s at the 7 5 n level of point S 2 - 8 (Fig. 5a). The speed is 3 2 . 4 , 17.8 and 6 . 9 c n / s at 1 0 0 0 , 1 2 0 0 and 1500 m level of point S2-8, respectively (Fig. 5a). The width o f this countercurrent is about 55 km and its transport i s about 8.5~10' mS/s. A countercurrent is also found here during the September- October cruise of 1 9 8 7 , but its position was further to the east and both its width and volume transport were smaller (YEI. 1 9 9 0 ) . At section S a (Fig. 5b) the core o f the Kuroshio is located near the shelf break at point S S - 1 2 . Its velocities i n the upper 1 5 0 D of the core are e l l greater than 80 C ~ S .The maximum velocity o f 1 3 1 . 5 cm/s is at the 5 0 level. The Kuroshio axis seems t o move graduaily eastward below 250 a . This characteristic was also found in previous studies (Yuan and S u , 1 9 8 8 : Y E I , 1990).

Like the early summer cruise result there is also a cyclonic cold gyre to the west of the Kuroshio on the shelf north of Taiwan (Fig. 6 ) . The position of this gyre is almost the same a s that during early summer of 1 9 8 8 . The northeastward current which intrudes onto the shelf has a transport of about 0 . 1 7 ~ 1 0m~5 / s .

313

pointed out before, this value may be an underestimate. West of the gyre there is another northeastward current vhich seess to originate from the Taiwan Strait. The total transport of both northeastward currents into this computational region is about 0 . 6 3 ~ 1 0ms/s. ~ It is worthy to note that this value is not the total transport of the Taiwan Warm Current (TWC). because the computational region does not include the coastal region. In previous studies (Yuan and S u . 1 9 8 8 ; YEI. 1 9 9 0 ; Yuan, Su and N i , 1 9 9 0 ) , a part of the TVC had a tendency to converge to the shelf break and joins the Kuroshio, which is similar to the present resu 1 t As

.

3.3 Comparison of results from the three cruises We shall discuss the computed results from the three cruises, namely, September-October of 1 9 8 7 (YEI, 1 9 9 0 ) , early summer and autumn of 1 9 8 8 . Figure 7 shows the sketch of the Kuroshio width and the position of the main Kuroshio axis for the three cruise. The folloving main features are found: 1) The intrusion of the Kuroshio water onto the shelf during the autunn of 1 9 8 8 is close to the winter feature (Su and Pan, 1987). This means that the position of the Kuroshio is further into the shelf and its width is narrower during the autumn of 1 9 8 8 than its respective positions and widths in both late summer 1 9 8 7 and early summer 1 9 8 8 . I n addition, the countercurrent’s position i s further to the vest and its transport is greater during autumn of 1 9 8 8 than its respective positions and volume transports in both late summer 1 9 8 7 and early summer 1 9 8 8 (Fig. 7a). The Kuroshio undergoes a slight meander near 26O30’N (Fig. 71, which may be related to the change of orientation of the shelf-break line a s seen from the position of the 200 m isobath near 26’ 3 0 ” (Fig. 7 ) . Finally, the transport of the Kuroshio in the East China Sea at the section north o f Taiwan is 2 5 . 8 , 23.4 and 2 4 . 3 ~ 1 0 ”ms/s during the late summer 1 9 8 7 , early summer and autumn 1 9 8 8 , respectively. 2) The Kuroshio Countercurrent’s position is found to be furthest to the west and its transport i s the largest ( 8 . 5 ~ 1 0m~3 1 s ) during the autumn 1 9 8 8 cruise among the three cruises. 3) In all the three cruises, the position of the main axis of the Kuroshio (the position of the maximum surface speed) is found ismediatelr east of the 2 0 0 II isobath (Fig. 7b). At section S 2 the depth a t this position is 2 1 0 , 1260 and 1260m

314

during late summer 1 9 8 7 , early summer 1 9 8 8 and autumn 1 9 8 8 , respectively. During late summer 1 9 8 7 shelf-intrusion of the Kuroshio Surface Water was weak s o that the Kuroshio Subsurface Vater was lifted to the shelf (Su and P a n , 1 9 8 7 ) . This results i n large density gradient and hence the surface velocity of the Kuroshio i s larger near the shelf break ( ~ 2 0 0 m isobath) than at other positions. It is worthy t o note that the seasonal change of the centerline position of the Kuroshio width (Fig. 7a) does not correlate with the seasonal change o f its main axis position (Fig. 7b).

Fig.

7. The sketch of (a) the Kuroshio width and (b) the position

o f the main Kuroshio axis. Sep.-Oct. of 1 9 8 7 : ( 1 ) - - - , Yay -June of 1 9 8 8 : ( 2 ) - . - . - , and 0ct.-Nov. of 1 9 8 8 : ( 3 )

-.

4) During all the three cruises, there i s a cyclonic cold gyre west of the Kuroshio on the shelf north of Taiwan. Its position in late summer 1 9 8 7 is further to the southeast than during the two 1 9 8 8 cruises. This result is in agreement with the fact that in summer 1 9 8 7 the extent of the Kuroshio water intrusion onto the shelf i s less than during the two 1 9 8 8 cruises. 4

THE

CURRENTS EAST OF THE R Y U K Y U ISLANDS DURING EARLY

SUYYER

1988

This

section

presents

the current

structure

and

volume

315

transport at sections

Fa

and

Fa.

4 . 1 Section Fa There is

a large t o p o g r a p h i c variation along section F a T h e water d e p t h a t t h e west end o f s e c t i o n F a is s h a l l o w (about 1 2 0 0 m), but a t t h e m i d d l e o f t h e s e c t i o n t h e r e is a trench (the Ryukyu Trench) 5 8 2 0 m deep. East of the Ryukyu I s l a n d s t h e r e is a weak northward c u r r e n t o v e r t h e Ryukyu Trench. I t s maximum s u r f a c e s p e e d is a b o u t 1 7 . 8 cm/s and its s p e e d increases slightly with depth. r e a c h i n g t h e maximum velocity o f 3 8 . 2 C D / S a t t h e 400 m level of point F a - 1 . B e l o w t h e 400 m level its Max m u x speed s h i f t s gradually e a s t w a r d t o point P a - 2 . T h e c o r e o f his flow lies between .to0 and 800 m depths. This f e a t u r e was a s o found in a p r e v i o u s s t u d y (YEI, 1990). (Fig.

8).

Cornputatim Points

2000 ' h

E "

3000 -

4000.

5000.

5800 :

Fig.

8. Velocity distribution a t

section

Fa

1 9 8 8 (positive value d e n o t e s northward flow).

4.2 Section

d u r i n g early s u m m e r (units: cm/s).

Fe

T h e t o p o g r a p h i c variation is much larger a l o n g s e c t i o n F a than along section F a . T h e topography o f the western h a l f o f s e c t i o n F p r e s e m b l e s t h e topography o f s e c t i o n Fa. In the eastern h a l f o f section FS t h e r e is a r i d g e a n d a deep basin.

316

West o f the ridge t h e r e is a northward c u r r e n t , a n d east o f t h e r i d g e t h e r e is a s o u t h w a r d c u r r e n t in t h e upper 700 m layer (Fig. 9). Most o f t h e northward f l o w c o m e t h r o u g h s e c t i o n F a , and the r e m a i n d e r c o m e s f r o m t h e T o k a r a S t r a i t (Fig. 4 ) . The surface velocities o f t h e northward c u r r e n t vary f r o m 9 t o 3 0 cm/s. T h e velocity increases gradually with d e p t h a n d r e a c h e s t h e maximum o f 3 2 . 1 cm/s a t 6 0 0 m level of point Fa - 3 . At p o i n t F a - 3 i t s velocity is 1 8 . 2 , 1 0 . 9 , 5 . 8 a n d 1 . 7 cm/s a t 1 0 0 0 , 1 2 0 0 , 1500 a n d 2 0 0 0 m levels, respectively (Fig. 9 ) . Canpltaticm

Points

Fig. 9 . Velocity d i s - i b u ion a t s e c ion F e d u r i n g e a r l y s u m m e r 1 9 8 8 (positive v a l u e d e n o t e s northward flow). (units: c m / s ) . Near t h e r i d g e t h e r e is a n a n t i c y c l o n i c g y r e between t h e northward a n d s o u t h w a r d c u r r e n t s ( F i g s . 4 a n d 9). T h e maximum speed o f the southward current, 2 5 . 9 cm/s. i s a t t h e s u r f a c e o f point Fo-6. In Lhe southward c u r r e n t t h e r e i s a s m a l l c o r e of the northward c u r r e n t between 6 5 0 t o 1 3 0 0 water d e p t h s a t point F a - 6 , but its strength is rather weak with t h e maximum velocity o f a b o u t 4 . 1 cm/s. Below t h e s o u t h w a r d c u r r e n t the f l o w is again northward. I t s s p e e d is only a b o u t 0 . 1 cm/s o r less. 4 . 3 T h e v o l u m e transport F i g u r e 4 s h o w s that the n o r t h w a r d c u r r e n t e a s t o f t h e R y u k y u Islands f l o w s through t h e western part o f s e c t i o n F a a n d its Most of t h i s c u r r e n t c o n t i n u e s t o transport is 2 6 . 8 8 ~ 1 0m~S / s . f l o w northward through s e c t i o n F a with a transport o f 2 4 . 5 5 ~ 1 0 ~

317

T h e r e m a i n i n g 2.33~10' m3/s o f this c u r r e n t t u r n s e a s t w a r d f r o m t h e e a s t e r n part o f s e c t i o n F a and m a k e s a n a n t i c y c l o n i c meander t o f l o w southward through s e c t i o n FZ (Fig. 4 ) . T h e r e is a l s o another weaker northward c u r r e n t c o m i n g t h r o u g h s e c t i o n s F I and F O near t h e Tokara S t r a i t . I t s t r a n s p o r t is a b o u t 3 . 7 3 ~ 1 0 " m 3 Is. The totals o f t h e n o r t h w a r d a n d s o u t h w a r d t r a n s p o r t s at ~ respectively s e c t i o n FO a r e 2 8 . 2 8 ~ 1 0 " m3/s a n d 8 . 8 0 ~ 1 0 m3/s, (Fig. 4 ) . In t h e n e x t place, t h e t r a n s p o r t s o f t h e northward c u r r e n t t h r o u g h s e c t i o n F S in t h e f i r s t t o f i f t h layers a r e 2 . 1 6 , 4.18, 8.68, 1 0 . 2 9 and 1 . 5 7 ~ 1 0 " ms/s, r e s p e c t i v e l y , w h i l e those through s e c t i o n F O in t h e f i r s t t o f i f t h layers a r e 2 . 8 2 , 5.11, 9 . 4 1 , 9 . 1 4 and 1 . 7 9 ~ 1 0 " m3/s, respectively. T h e transport o f t h e s o u t h w a r d c u r r e n t through section F Z in t h e f i r s t t o fifth l a y e r s a r e 2.69, 1 . 7 8 . 1 . 6 7 , 2 . 3 4 and 2 . 6 6 ~ 1 0 ' m 5 / s , respectively. T h e hydrographic s t r u c t u r e in t h e upper water laver s h o w s a temperature d e c r e a s e s f r o m t h e west t o t h e e a s t (Fig. 10a) and d o e s not r e f l e c t a northward current. However, t h e d i s t r i b u t i o n o f t e m p e r a t u r e below the 4 0 0 m level (Fig. lob) d i f f e r s f r o m t h a t in t h e upper layer, reflecting t h e e x i s t e n c e o f t h e n o r t h w a r d current. T h u s , t h e measurement o f hydrographic s t r u c t u r e in the d e e p layer is important t o understanding t h e c u r r e n t f e a t u r e e a s t o f t h e Ryukyu Islands. m3/s.

1 0 . T e m p e r a t u r e d i s t r i b u t i o n a t (a) 2 0 0 m a n d ( b ) 5 0 0 levels e a s t o f t h e Ryukvu Islands in May-June, 1 9 8 8 .

Fig.

m

5 SUMWARY

Based o n hydrographic data f r o m both the early s u m m e r a u t u a n 1 9 8 8 cruises. t h e K u r o s h i o in t h e East C h i n a S e a a n d

and the

318

c u r r e n t s east o f t h e Ryukyu I s l a n d s a r e s t u d i e d by the inverse method. It is found that: 1 ) T h e c u r r e n t pattern in t h e East C h i n a S e a d u r i n g 0ct.Nov. 1 9 8 8 is c l o s e t o t h e winter pattern. T h e f e a t u r e s are: ( 1 ) T h e positions of t h e main K u r o s h i o C u r r e n t a n d c o u n t e r c u r r e n t a r e both f u r t h e r t o t h e w e s t t h a n d u r i n g o t h e r c r u i s e s ; ( 2 ) T h e c o u n t e r c u r r e n t is s t r o n g e r a n d t h e width o f t h e K u r o s h i o i s narrower during 0ct.-Nov. 1 9 8 8 than d u r i n g other c r u i s e s . 2) T h e s e a s o n a l c h a n g e o f t h e K u r o s h i o a x i s (the position o f t h e maxiaum s u r f a c e current) d o e s n o t c o r r e l a t e with t h e seasonal c h a n g e of t h e position o f t h e c e n t e r l i n e o f t h e K u r o s h i o C u r r e n t width. 31 T h e t r a n s p o r t o f t h e K u r o s h i o in t h e E a s t C h i n a S e a a t 2 3 . 4 a n d 2 4 . 3 ~ 1 0 ' m3/s, t h e s e c t i o n north o f T a i w a n is 2 5 . 8 , r e s p e c t i v e l y , d u r i n g t h e late s u m m e r 1 9 8 7 , early s u m m e r a n d a u t u m n 1 9 8 8 . D u r i n g all t h e s e t h r e e c r u i s e s a c y c l o n i c cold g y r e is present vest o f t h e K u r o s h i o o n t h e s h e l f north o f Taiwan. 4) During a u t u m n 1 9 8 8 t h e t r a n s p o r t o f t h e T a i w a n Warm C u r r e n t (TWC) f l o w i n g into our computational r e g i o n is a b o u t 0.63~10' m3/s. which is n o t t h e total t r a n s p o r t o f t h e TWC because the computational r e g i o n e x c l u d e s t h e coastal region. 5) In t h e s u r v e y area t h e c u r r e n t e a s t o f t h e Ryukyu I s l a n d s f l o w s n o r t h w a r d o v e r t h e Ryukyu Trench d u r i n g early is s u m m e r 1 9 8 8 . o f which t h e t r a n s p o r t a t s e c t i o n s F a a n d F e a b o u t 2 6 . 8 8 and 2 8 . 2 8 ~ 1 0 " . ' I s , respectively. T h i s n o r t h w a r d I t s c o r e lies c u r r e n t is not s t r o n g t h r o u g h o u t t h e depths. between 4 0 0 and 8 0 0 II.

6 REFERENCES Guan B i n g x i a n , 1 9 8 2 . Analysis o f t h e v a r i a t i o n s o f volume t r a n s p o r t o f K u r o s h i o in t h e E a s t China Sea. In: P r o c e e d i n g s o t h e J a p a n - C h i n a O c e a n S t u d y Symposium. Oct., 1 9 8 1 , S h i n i z u PP. 1 1 8 - 1 3 7 .

Guan B i n g x i a n , 1 9 8 8 . Major f e a t u r e a n d variability o f t h e K u r o s h in t h e East C h i n a Sea. Chin. J . O c e a n o l . Limnol. 6 ( 1 ) : 3 5 - 4 8 K o n a g a , S . . N i s h i y a m a , K., I s h i z a k i , H . a n d Hanazava. Y., 1 9 8 0 . G e o s t r o p h i c C u r r e n t S o u t h e a s t o f Yakushima Island. La H e r , 18: 1-16.

Nishizava. J., Kamihira, E.. Komura. K., K u m a b e , R. a n d H i y a z a k i , M . , 1 9 8 2 . Estimation o f t h e K u r o s h i o mass t r a n s p o r t flowing o u t o f t h e East C h i n a S e a t o t h e North Pacific. La Mer, 2 0 : 37-40.

Saiki. M., 1 9 8 2 . Relation between t h e g e o s t r o p h i c f l u x o f t h e K u r o s h i o in t h e Eastern C h i n a S e a a n d its large-Meanders in s o u t h o f Japan. T h e Oceanographical M a g a z i n e , 3 2 ( 1 - 2 ) : 11-18.

0

319

Su J i l a n and P a n

Y u s i u . 1987. O n t h e s h e l f c i r c u l a t i o n north of Taiwan. Acta O c e a n o l o g i c a S i n i c a , 6 ( S U P P . I): 1-20. Yuan Yaochu. S u Jilan a n d Xia S o n g y u n , 1987. Three dimensional diagnostic c a l c u l a t i o n of c i r c u l a t i o n over t h e East C h i n a Sea Shelf. Acta O c e a n o l o g i c a S i n i c a , 6 ( S U P P . I): 36-50. Yuan Yaochu and Su J i l a n , 1988. T h e c a l c u l a t i o n of K u r o s h i o Current S t r u c t u r e i n t h e East C h i n a S e a - Early S u m m e r 1986. P r o g r e s s in O c e a n o g r a p h y , 21: 3 4 3 - 3 6 1 . Yuan Yaochu, Endoh M a s a h i r o and Ishizaki H i r o s h i , 1990. T h e Study of t h e K u r o s h i o in t h e East C h i n a S e a and C u r r e n t s East of Ryukyu Islands. I n : P r o c e e d i n g s of J a p a n C h i n a J o i n t Symposium o f t h e C o o p e r a t i v e S t u d y on t h e K u r o s h i o , S c i e n c e and Technology Agency. Japan & S O A , China, PP. 39-57. Y u a n Yaochu, Su J i l a n a n d Pan Z i q i n , 1990. C a l c u l a t i o n o f the K u r o s h i o Current S o u t h of J a p a n During Dec., 1987-Jan., 1988. I n : P r o c e e d i n g of t h e I n v e s t i g a t i o n of K u r o s h i o ( 1 1 ) ( i n press). Yuan Yaochu, S u Jilnn and N i J u f e n , 1990. A P r o g n o s t i c Model o f the Winter C i r c u l a t i o n in t h e East C h i n a Sea. In: P r o c e e d i n g s of t h e I n v e s t i g a t i o n of K u r o s h i o (11) (in press).


321

THREE NUMERICAL MODELS OF GUANHE ESTUARY ZHANG DONGSHENG AND ZHANG CHANGKUAN Hohai University Nanjing, China

ABSTRACT Three numerical models of Guanhe estuary have been set up for different practical motives. The first model, dimension combined model, is to study the circulation pattern and to evaluate the influence of the river discharge on the current pattern in the region arround the entrance. An explicit characteristic difference scheme is used for the inner part while a kind of triangular element method is used for the outer part. The second one is to simulate current field, sediment transport and the corresponding bed variation. This second model is discreted by AD1 scheme in which the technique of non-uniform grids and movable boundary is introduced to improve the accuracy of the simulation in the vicinity of the river mouth. The third model has been developed with curvilinear orthogonal coordinate system to improve the computed current field close to the physical boundary. The current fields obtained by the three models are similar and show their own advantages in representing the nature. 1 INTRODUCTION The Guanhe estuary, with an irregular geometry and a seabed topography complicated by shoals at the entrance, is one of important estuaries entering the East China Sea, Jiangsu provence(Fig.1). Tides enter the estuary up to the top, 75km from the entrance. Freshwater enters the entrance through four sluice gates during flood season and mixes with sea water in the estuary. The numerical simulation of Guanhe estuary is of scientific significance on the studies about the interaction of sea water and the river discharge and about the formation of the entrance bar. The work so far carried out is for the purpose of predicting beach response and shoal change due to natural forces and constructed works. Three numerical models of Gaunhe estuary have been set up for different practical motives. The first model, an one-two dimension combined model, is to study the circulation pattern, and to evaluate the influence of the river discharge on the current pattern in the region arround the entrance. In this model an explicit characteristic difference method is used for the inner part and a kind of triangular element method with three nodes is used for the outer part. The model is successful in simulating tidal current fields, especially in demonstrating the influence of Guanhe river discharge on the current pattern, but the current structure in the vicinity of the entrance is not accurate enough to represent the influence of complicated topography near the entrance bar. The second model is set up to simulate the tidal current, suspended sediment fields and corresponding bed variation. The modelled domain includes the outer part and an 11km long inner part. This model is discreted by AD1 scheme. To improve the accura-

322

cy of the simulation in the vicinity of the entrance, a technique of non-uniform grids and movable boundary is introduced. The third model has been developed with curvilinear orthogonal coordinate system for the purpose of fitting the boundary better and considering the effect of the river bend on the current pattern in the vicinity of the entrance. Based on the solution of a set of Laplace equations, a curvilinear orthogonal grid system is generated by fixing a part of boundary points and moving the others. AD1 scheme has been extended for the transformed shallow water equations in curvilinear coordinate system.

2 MODELS 2.1 Basic equations 2.1.1 Basic eauations with Cartesian coordinate svstem (i) Equation for one dimensional calculation ( for the inner part of the first model )

au

au

- +u-

+g-

aH ax

ulul

+ g z

=o

C,R

where H is the total depth of flow (H=D+h, D: water depth with respect to the chart datum, h water surface elevation above the chart datum); Q is section discharge; u is section mean velocity; b is water surface width; R is hydraulic radius; g is gravitational acceleration and C,is Chezy coefficient. (ii) Equations for two dimensional calculation ( for the outer part of the first model and the whole domain in the second model)

+ 2( H U ) -a( ~ v )= o at ax aY au au au ah - +u+v-.fi+gat ax ay ax

g u v +--c: =o R

av av av ah -+u-+v-++fi+R-+-at JX ay ax

g vv c: R = o

+

where u and v are depth-averaged velocity components in x and y horizontal coordinate directions, respectively; f is Coriolis parameter; f is velocity modulus u v );and the others are same as in the equation (1). (iii) Equation of suspended sediment transport

(v=7 +

a0+a0 +-a ( H v C ) - - ( H D J

at

ax

ay

ax

- ) - - ( HJD aC =ax ay

aC -)-ao(fiS, yay

-yC)=O

(3)

323

where, C is depth-averaged sediment concentration, u is fall velocity of sediment particles, S,is capacity carrying sand by flow, u is settling probability of sand particles, D, and D, are dispersion coefficients along x and y directions, respectively, p and y are specified by the follows: 1

'={O

when u , v 2 u c when u,v < u c

1 "{O

when u , v 2 u r when u , v c u r

where uf is velocity for suspending sand particles and uc is threshold velocity. 2.1.2 Basic equations with curvilinear orthogonal coordinate system (i) Equations of unsteady flow in shallow water ( for the third model )

(ii) Basic Equations for generating curvilinear orthogonal coordinate

where 4 and q are variables in curvilinear orthogonal coordinate system and

and the corresponding boundary conditions are:

324

x =X )

(5,v = d , 1

Y=Y3

(5,s=d1)

x =x4

(4:,rl= d , )

Y =Y ,

(5,1=

on the boundary of q = d ,

on the boundary of q = d ,

d2)

2.2 Numerical models 2.2.1 One--two dimension combined model (model 1) In this model an explicit characteristic difference method is used for a 75km long inner part, which is divided into 30 segments about 2km long on the average. According to the Courant condition, the time step adopted is 150 sec. A kind of trianglar element method with three nodes is used for the outer part which has a size of about 60 * 40km. The shape of element is designed to fit the topographic feature of Guanhe estuary (Figs.1 and 2 ), and the size of elements is smaller for nearshore area than for offshore. The time step is 300 sec. Tidal levels are employed as boundary conditions at the open seaward boundaries. The measured water level variations are provided by the gauge stations located at Liu Wei and Zhong Shan, as shown in Fig. 1. The water level variations at the rest part of the open boundaries are obtained througth linear interpolation considering some damping seaward (Zhang Dongsheng et al. 1987). 2.2.2 Two dimensional model of suspended sediment transport (model 2) The modelled domain including a river section is about 60km * 25km and the upward boundary of the river section is located at Chen Gang, l l k m from the entrance (Fig.]). The simulated area is discreted by a grid of 1200m * 1200m. For the sake of closely fitting the complicated topography near the entrance, where the river width is 800-1000m only, finner grid with spacing of 200m is adopted (Fig.3 ). The time step used is 150 sec for current field computation and 450 sec for the sediment transport computation. Parameters such as S . , a,etc. are important for the sediment transport calculation. The capacity carrying sand by flow, S , ,is obtained by the following formula (Liu Jiaju, 1980) :

where p,is the denisty of sand, V,is the combined velocity of tidal and wind induced currents and V,is orbital velocity of wave motion.

325

N

China

712 Fig. 1 Map of Guanhe estuary showing the location of survey area and boundaries of simulation domains

Fig.2 Simulation area divided into triangular elements of model 1 (Dash lines stand for bottom topography)

Fig.3 Simulation domain of model 2 resolved by AD1 scheme

326

The fall velocity, a,is calculated by (Zhang Ruijin, 1989) : a=--1 P , - P R-d 2 24 P , Y where d is the particle size of sands; y is the kinematic viscosity of fluid. The threshold velocity, uc, and the velocity for suspending sand particles, uf, are calculated by (Qian Ning et al. 1983) : H

u C= (-)

d

0.14

PI -p -71.0+ H (17.6d+6.05 x 10 p, d O.” ~

f

1

where p ,p ,,are densities of water and sand, respectively. iment transport is:

a0+a0 +-a(HvC) = O at

ax

aY

The boundary condition for sed-

(in thecaseofoutflow)

or

c=c,

(in the case of inflow)

where C , is measured sediment concentration at Gl,G3,G4 and G5 (refer to Fig.3 for their locations). 2.2.3 Two dimensional numerical model for unsteady flow with curvilinear orthogonal coordinate (model 3) In consideration for the influence of river bend on the current pattern in the vicinity of the entrance, the modelled domain is limited to a smaller region. That is 15km * IOkm. The upstream boundary is the same as that in the model 2. The grid has been formed according to Eq.(5) (Lau P C M, et al. 1979, Middlecoff J F, et al. 1979) and the spacing varies from 160m to 250m. Fig.4 shows the grid in physical domain and F i g 3 shows the grid in the calculating domain. The time step used is 100 sec. 3 MODEL VERIFICATION AND SIMULATED RESULTS To collect the field data of tidal currents and water levels, four field measurements in Guanhe estuary were carried out in 1980,1981 and 1985 (two measurements), respectively. Parts of the field data obtained from the measurements have been employed to verify the above three models. Fin.6 is the comparison of tidal currents between the measured and the calculated bv the

32i

Fig.4 Grid of model 3 in physical domain

328

v e l o c i t y vectors a t GGI (1985. z 19)

Fig.6 Comparison between field data and calculated results water level at Yan Wei

_____

measured

-C a l c u l a t e d by , -2.0

.

x

mdel 2

calculated by model 3

L current v e l o c i t y

current velocity a t GGI

u (m/s 1

-0.8

u b/S) 0.8 -

-0.8-

at

u (m/s 1

1

u (n/s)

-0.8

-

Fig.7 Comparison between field data and calculated results

GG2

329

model 1 at points G G l and GG2 showing a good agreement between measured and calculated both in the magnitude of velocity and in the current direction. Fig.7 is the comparison of water levels at Yan Wei and tidal current velocities at G G l and GG2 between field data and numerical results. In Fig.7 dot lines, solid lines and cross marks stand for observed values, calculated values by the model 2 and the model 3, respectively. The comparison shows that the simulated values, either the water levels or tidal current velocities, as well as the velocity phases agree well with the measured. Fig.8 gives measured tidal current vectors during a complete tidal period at different observation stations. Figs.9 and 10 demonstrate simulated tidal current vectors in Guanhe estuary during a complete tidal period obtained from the model 1 and the model 2, respectively. Figs.] 1 and 12 are tidal current vectors at flood tide and ebb tide obtained from the model 3. From the comparisons it can be concluded that either the current speeds and the directions at GGI and GG2 or the general tidal current patterns in the whole estuary simulated by the three models are similar with the field observations. Fig.6 shows that the tidal current is semidiurnal, counter clockwise rotatory, and the calculated tidal ellipses are essentially the same as that obtained from field observations. Comparing the Figs.9, 10, 1 I, and 12 with the Fig.8, it can be found that the dominant current directions at flood and ebb tides on the western side of the estuary are SE, as well as NW and SSE and WNW on the eastern side, respectively. The Fig.9 indicates that there is a small area near the entrance within which the tidal velocity is particularly smaller than that of adjacent area. The magnitude of velocity is about 0.2-0.5m / s, and at the position of shoals it reduces to 0. Im / s. Fig.9 shows the boundary of the small velocity area in dash line. It would be worth pointing out that the water depth in the small velocity area is so shallow that observation boats can not reach this area. The existence of this small velocity area interprets the formation of the entrance bar, at least from the viewpoint of hydrodynamic influence. The computation results suggest that there is only a little difference in the current magnitude and direction between flood and dry seasons. The mean current velocities of flood and ebb tides during flood season are 0.01-0.05m / s larger than those during the dry season, and the region affected by the river discharge is limited to a small area near the entrance. This is because of the small river discharge released from four sluice gates, only 200m3/ s on the average during the flood season. Therefore, it can be concluded that river discharge variation under normal condition makes less contribution towards the current pattern of the outer part of the estuary. Suspended sediment calculation has given acceptable results showing good agreement with field data. Fig.] 3 gives the comparison between measured sediment concentration and computed at two points G2 and G6(refer to Fig.3 for their positions). Figs.14 and 15 are the distrbutions of suspended sediment concentration in Guanhe estuary during the flood tide and the ebb tide, respectively, obtained from the model 2. This two figures show that a reasonable distribution of suspended sediment concentration has been achieved from the numerical model. The value of sediment concentration is 1.Okg / m30n the average in the area around the entrance and decreases gradually seaward. During the period of flood tide, the suspended sedi-

330

/

0

10 20 30 40

50km 0.6m/s Y

Fig.8 Measured tidal current vectors within a complete tidal period

2.0 m.? u

Fig.10 Tidal current vectors within a complete tidal period simulated by the mode 2

331

Fig.11 Tidal current vectors at flood tide by the model 3

Fig.12 Tidal current vectors at ebb tide by the model 3

332

Fig.13 Comparison between the measured sediment concentration and simulated by the model 2

c

-

Zhong Shan

Fig.15 Distribution of sediment concentration at ebb tide ' ) obtained from the model 2(numbers in 0.1kg / m

333

ment is pushed into the shore while pulled off during ebb tide. The range of moving of suspended sediments during a complete tide period is about 1Okm. 4 DISCUSSION

The verifications between numerical results from the three models 1 to 3 and field measured data indicate that all the three models are suitable for tidal current simulation of Guanhe estuary, and the model 2 can give an acceptable estimate and a reasonable distribution of the sediment concentration. For different practical motives, by judging computation cost, accuracy requirement, etc., we can choose one of them. If we are more interested in the influence of the river discharge on the current pattern in the estuary area, then the model 1 is preferable. If we want to know more about the effect of complicated topography near the entrance bar on the current structure in the estuary region, the model 2 is better. If we are more concerning the fitting of physical boundaries, especially in the case of treating problems involving coastal engineering in which the velocities close to the coast or constructed works become more interesting, certainly the model 3 should be on the top of list of candidates. ACKNOWLEDGMENTS The authors would like to thank Mr. Zheng Xiaoping, Mr. Jiang Qin and Mr. Liu Jinfang for writting computer programs and for performing computations. This work was supported by the National Foundation for Natural Scientific Research. REFERENCES Lan P C M, 1979. Curvilinear finite difference method for biharmonic equation. International Journal for Numerical Method in Eng. Vol. 14, No.6, PP. 7Y 1-812. Lin Jiaju, 1980. Siltation calculation and prediction of the approach channel of Lianyun Harbour. Journal of Water Conservancy and Water Transpotation. No. 4. (in Chinese) Middlecoff J F and Thomas P D, 1979. Direct control of the grid point distribution in meshes generated by elliptic equation. AIAA Computation Fluid Dynamics Conference, July. Qian Ning and Wan Zhaohui ,1983. Mechanics of Sands. Science press (in Chinese) Zhang Dongsheng, Xie Jingzhan and Zheng Xiaoping, 1987. The numerical simulation of hydraulic factors of Guanhe Estuary. Proceedings of Coastal and Port Eng. in Developing Countries, Vol.11, PP.2215-2224. Zhang Ruijin, 1989. Dynamics of river sand. Water Conservency and Hydro power Press (in Chinese)


335

GEOSTROPHIC CURRENT NORTHEASTERN TAIWAN

AND

ASSOCIATED

VERTICAL

MOTION

OFF

Hsien-Wen Li Department of Oceanography, National Taiwan Ocean University, Kcelung 20224, Taiwan, CHINA

ABSTRACT The geostrophic velocity changes its direction with increasing depth by thermal wind relationship. The rate of rotation with depth is related to the vertical velocity by Hide's theorem. A method for calculating the three-dimensional absolute velocity by means of Hide's theorem and beta vorticity dynamics is developed. It is applied to hydrographic data in October 1987 and May 1988. The Kurosliio northeast of Taiwan flows nearly along the bottom contour. There exists a small cold core off northeastern Taiwan. The position of the Kuroshio's axis changes little in this period. The vertical velocity is mostly upward in the central part and west side of the Kuroshio and mostly downward east of the Kuroshio. Its lo-' m s-'. typical magnitude is N

INTRODUCTION In a rotating fluid, the Taylor-Proudman theorem implies that the horizontal velocity of a rotating homogeneous fluid is independent of the direction parallel to the rotation axis, and thus the motion is completely twodimensional. However, the velocity field of a non-homogeneous fluid is three-dimensional. Hide (1971) established a general theoretical relationship between the thre4imensional velocity field and density field for geostrophic flow of a non-homogeneous fluid. He found an equation for the rate of change with respect to the rotation axis of the direction of the horizontal flow. Schott and Stommel (1978) also derived an equation for the rate of turn of geostrophic velocity with height in the main thermocline, which they called the beta spiral by adding the beta vorticity constraint. In the atmosphere strong The atmosphere and the oceans are non-homogeneous. vertical motion occurs in the neighborhood of a front (e.g., Holton, 1972). In regions of cold or warm advection, the geostrophic wind backs or veers with height. Although the fronts are of the nature of highly ageostrophic pheno~nena,they are necessary concomitants of highly geostrophic flow (Hide, 1971). Thus there exists a relationship between the vertical motion and the turning of geostrophic velocity with height or depth near the fronts in the atmosphere and the oceans. The Kuroshio is an oceanic front in a sense that the temperature difference is remarkable across it. In this work a version of Hide's theorem is derived to depict the relationship between the rate of rotation of the horizontal velocity and the vertical distributions of temperature and salinity.

336

THEORY AND METHOD FOR DETERMINATION OF THE ABSOLUTE VELOCIT‘IES The thermal wind equation for the ocean is (e.g., Gill, 1982)

where x, y and z are Cartesian coordinates with x to t,he east, y to the north and z upward. The velocity components in x, y and z direction are denoted by u, v and w, respectively, f is the Coriolis parameter, g the gravity acceleration, T the temperature, cy the thermal expansion coefficient, S the salinity and y the expansion coefficient for salinity. We have steady-state equation for temperature and salinity,

where

lit and X,

are molecular kinematic diffusivities for temperature and salt, respectively.

In this study we consider only the large-scale geostrophic current far from the boundary layers, thus the eddy diffusivities due to turbulent smallscale flow may be neglected. We retain the molecular diffusivities according to physical nature of diffusion due to the existence of gradients of temperature and salinity. If we combine (1) multiplied by u with (2) multiplied by v, we get

Eqs.(B) and (4) are substit,ut,ed into ( 5 ) . Thus,

The velocity components can be expressed in polar form, u = V C O S ~, v = V sine where V and 0 are the magnitude and direction of the geostrophic velocity, respectively. Then (6) becomes

337

This is an equatioii for the rate of turn of t,he geostrophic current with depth. Data in October 1987 and May 1988 are analyzed. The orders of magnitude of the calculated expansion coefficients for temperature and salinity ale,

-

3.18 (at the surface)

x

10-4((0C:)-i

y = 7.36 (at the surface) 7.63 (at, 10001n)

x

10-4(psu)-’

N = 1.29 (at 1000m)

N

Typical values for the diffusivities of temperatiire and salinity (e.g., Gill, 1982) are,

= 1.5 x

m2 s-l

In the upper ocean where thermocline and halocline prevail, the vertical gradients of ternperat ure and salinity are niuch larger than their horizontal gradients, respectively. Thus a2T and the threedimensional Laplacians of temperature and salinity can be replaced by

6%

, respectively. Therefore, ( 7 ) can be approximated by

The numerical values of the first, derivatives of temperature and salinity with c1ept.h in the upper ocean are larger than their second derivatives, respect,ively. Thus if the nutnerical value of vertical velocity is larger than the values of diffusivities as shown above, i.e., if

O(w) > 10-7

111 s-1

then (8) can be approximated by

Eq.(9) gives the rate of change with depth of the direction of geostrophic velocity in relation to the vertical gradients of temperature and salinity, which is a version of Hide’s theorem. In the main thermocline, where

338

Eq.(9) ca,n be further simplified as follows,

E q . ( l l ) can be applied to the upper layer down to about 300 meters deep (upper thermocline) of the oceans in the low and middle latitudes, where the vertical gradient of temperature is stronger than the salinity gradient to satisfy (10). Thus (11) can be applied to the East China Sea. is positive except i n deep waters. The Coriolis parameter f is positive The gradient in the northern hemisphere, and the thermal expansion coefficient seawater. Therefore,

(I

is positive i n normal

w p 0 if aa85 0

Thus the upwelling or downwelling can be determined by the direction of rotation of geostrophic velocity with increasing depth in the upper thermocline through right-hand rule in the northern hemisphere. Schott and Stommel (1978) added the following beta vorticity constraint to the above mentioned current spiral and called it beta spiral. A linearized steady vorticity equation is / h - f z i3w =O

We will use (9) and (12) to determined the best-fit level of no motion by least3quares method. In beta spiral method different density surfaces at difdetermined from an observed densit,y field. The method present,ed here is quite simple, where no densit,y surface determination is required. We have N levels of temperature and salinity data for each station. The deepest, level is designated by N. The relative velocities between different levels can be determined by dynamical method. The procedure for determination of a best-fit level of no motion is as follows. 1) The deepest level N is assumed to be a reference level, where u, v and w are zero. The relat,ive horizontal velocities in t.he whole water column are calculated, which are substituted into (9). The vertical velocities in the whole water column a,re determined. 2) The velocities v and w at many depths determined at the first step, are substit,uted into (12). The right-hand side would vanish at all the depths if the assumed reference level was a real level of no motion. Practically it will not vanish. Deviations from zero (errors) are denoted by DNi , where the first subscript N refers to the reference level, and the second subscript i (i = 1,....,N-1) a level where D

Ni

is

339

calculated. The squares of the errors are summed up,

3) The above steps (1) and (2) are iterated by assuming different reference levels from the deepest level to the surface through a water column. We then have

N

1 Dnil 2

En=

n = l , .....,N

.

i=1

if,

The best-fit level of no motion is chosen in such a way that En becomes minimum.

27*00

26.00

I

1

+ f

+

+

+

+

25.00

24.00 121:oo

122:oo

123:OO

+

Fig. 1. Bottom topography off northeastern Taiwan. Symbols denote the the station of Cruise 124 (8 to 11 October 1987) and 157 ( 13 to 17 May 1988) of R/V Oc,ean Researcher 1.

340

HYDROGKAPHY Figure 1 shows the bott,oni topography of the sea northeast of Taiwan with CTD stations. The temperature and salinity distributions at 50m depth measured by CTD of R/V Ocean Researcher 1 on cruise 124 from 8 to 11 October 1987 are shown in dashed lines in Figs. 2 and 3, respectively. The region of maximum temperature gradient ma.y mark the front of the Kuroshio. There exists obviously a cold core off northeastern Taiwan. The cold core extends from the sea surface to about lOOm depth, where the water depth is approximately 200m. The dashed lines in Figs. 4 and 5 show the temperat>urea.nd salinity distributions at 50111 depth measured by CTD on cruise 157 from 13 to 17 May 1988. The temperat.ure gradient is apparently steeper in May 1988 than in October 1987. This indicates the Kuroshio is st’ronger in May 1988 than i n October 1987. The cold core off northea.stern Taiwan exisrs still there. The motion described by (9) and (12) is basically a large-scale geostrophic: flow. For eliminating the local effects on the temperature a.nd salinity distributions which are mostly due to the tidal motion a smoothing operation is done basically by a 5-point scheme, (Shuman, 1957) for inner stations, where TI, T,, T, and T4 are data at the stations nearest to the station of To. Data at outer stations are smoothed by -

To = JT,

+ 4 ( Ti + T2 + T,

)

Data at the corner stations are smoothed by

Isolines of temperature and salinity by smoothed data are also shown in Figs. 2-5. RESULTS By means of the method described above the velocities at 50m depth are calculated for the cruises 124 and 157, as shown in Figs. 6 and 7. The calculated best-fit levels of no motion at different stations in these two cruises are different, as shown in Fig. 8. The Kuroshio off nortlieastern Taiwan changes little its direction in October and May. It is stronger in spring (May, 1988) than that in autumn (October, 1987). There is a cyclonic eddy accompanied by the cold core northeast of Taiwan. The eddy is also stronger in May than in October. These figures show that upwelling exists in the central part and west side of the Kuroshio whereas downwelling on its east side. The calculated velocities are typically 10P m s-l. N

341

121

122

123

122

123

26

25

24 121

Fig. 2. Temperature distribution (-smoothed, October 1987.

--unsmoothed) at 50m depth from 8 to 11

-

121 I

24121

122

I

I

123

I

122

Fig. 3. Salinity distribution (-smoothed, 1987.

I

123

1

24

--unsmoothed) at 50m depth from 8 to 11 October

342

121 I

122 I

123 I

I

26

26

25

25

I

24 121

I

I

122

123

Fig. 4. Temperature distribution (-smoothed, May 1988.

121

122 I

24

--unsmoothed) at 50m depth from 13 to 17

123

I

I -

I

I

I

26

26

25

25

24

Fig. 5 . Salinity distribution (+moothed, 1988.

--unsmoothed) at 50m depth from 13 to 17 May

343

220

CRUISE:

124

129

1 M/SEC

Fig. 6. The calculated velocity distribution at a depth of 50m from 8 to 11 October 1987. Symbols 0 denote upwelling, whereas symbols @ downwelling.

Fig. 7. The calculated velocity distribution at a depth of 50m from 13 t,o 17 May 1988. Symbols @ denote upwelling, whereas symbols @ downwelling.

344

26.00

25.00

a

24.00 121.oo

122.00

123.00

Fig. 8. The number in pa.rentheses ( ) and brackets < > refers to the best-fit level of no motion at various stations on cruises 124 and 157, respectively. CONCLUSION The method for determination of the level of no motion presented here is quite simple. The magnitude of the calculat.ed vertical velocities in and near the Iiuroshio off northea.stern Taiwan is typically m s-l, which may be larger than that in region far from the frontal area,. On the whole, the central part and west side of the Kuroshio is an upwelling area and its east side is a dowiiwelling area. N

ACKNOWLEDGEMENTS The author is grateful t o Mr. Yao-Tsai Lo and Mr. Ming-Jeng Chang for preparing diagrams and typing the manuscript. Special thanks are due to the crew and working team of R / V Ocean Researcher 1 for their assistance during t.he hydrographic survey. The referee's valuable comments are appreciated. This study is sponsored by the National Science Council of the Republic of China through grant NSC 784209-MO19-01. REFERENCES Gill, A.E., 1982. Atmosphere-Ocean Dynamics. Academic Press, 662 pp. Hide, R., 1971. On geostrophic motion of a non-homogeneous fluid. J. Fluid Mech., 49, 4: 45-75 1. Holt,on, J.R., 1972. An introduction to Dynamic Meteorology. Academic Press, 319 pp. Schott, F. and Stommel, H., 1978. Beta spiral and absolute velocities in different oceans. Deep-sea Rcs., 25, 961-1010. Shuman, F.G., 1957. Numerical methods in weather prediction: 11. smoothing and filtering. Monthly Weather Rev., 85(11): 357-361.

345

WATER VOLUME TRANSPORT THROUGH THE TAIWAN STRAIT AND THE CONTINENTAL SKELF OF THE EAST CHINA SEA MEASURED WITH CURRENT METERS Guohong FANG, Baoren ZHAO and Yaohua M U Institue of Oceanology,Academia Sinica,7 Nanhai Road,Qingdao,Shandong 266071 ,China

ABSTRACT Historical current meter data from 24-hour anchored ships above the continental shelf of the East China Sea are analysed. The total (vertically integrated) flow averaged for each 2. 2" latitude X 2. 2" longitude square is clearly northeastward for all four seasons. The annual mean of the volume transport through the offshore area west of the 150 m isobath is estimated to be about 2 Sv. The anchored ship current meter data in the Taiwan Strait are analysd in combination with moored current meter data by Chuang (1985,1986) and Wang and Chern (1988) for a rough estimate of the volume transport through the Strait. The annual mean is also about 2 Sv. The volume transport through the Taiwan Strait and the shelf of the East China Sea is comparable, in magnitude, with the transport through the Tsushima Strait (Korea Strait), which is about 2. 8 Sv (Miita and Ogawa, 1984,1985). These results suggest the existence of an extended current system including the Tsushima current system on the shoreside of the Kurcshio, which extends from the Taiwan Strait to the Tsugaru Strait. This current system is called "Taiwan-Tsushima-Tsugaru Warm Current System (T-T-T WCS)". The dynamic height of the sea surface relative to 1000 db in the northeastern South China Sea is higher than that in the area east of the Tsugaru and Soya Strait by about 0. 7m. This sea-level difference is believed to be the main driving force of the T -T-T WCS. INTRODUCTION For a period of time it was believed that the current in the Taiwan Strait and the continental shelf of the East China Sea was driven by monsoon and thus went to the north in summer and to the south in winter. However, in the late 1950's and early 1960's direct current measurements showed that the current often flowed to the north even in winter, especially in the lower layers. Guan (1984, 1986) called it "Against-the-Wind

CurrentN or ffCounter-Wind Current", and

Fang and Zhao (1988) called it "Upwind Current". Up to date, however, no quantitative estimate of the volume transprt has been given. Since the flow in this area is mainly barotropic (Chuang, 1985, 1986; Fang and Zhao, 1988), and the conventional dynamic calculations p r e vide no reasonable results, it is thus necessary to make use of direct current measurement data. Our calculation will show that the mean volume transport through the Taiwan Strait and the shelf

of the East China Sea is comparable with that of the Tsushima Current, we shall pro-

existence

of a current system on the shoreside of the Kurcshio,known as the Taiwan- Tsushima-Tsugaru Warn Current System (T-T-T

WCS).

The driving mechanism of both the Taiwan Warm Current and the Tsushima Warm Current

has been topics of many papers. In this paper we shall give an explanation of the main forcing mechanism for the T-T-T

WCS, which was actually first proposed by Fang and Zhao (1988).

346

VOLUME TRANSPORT THROUGH THE TAIWAN STRAIT

Chuang (1985,1986) has made moored current measurements in the near-shore area west of Taiwan. Recently a moored current measurement lasting over one year has been reported (Wang and Chern, 1988). On the contrary, in the vast west part of the Strait no moored current mea-

.

current measurements have been made from an-

. The stations are shown in Fig. 1.

The measurements at stations Nos. 1 to 9 were

surement has been conduced A few 24-hour chored ship

.

made in wintertime Station 9 can be used to represent the channel between the Penghu Islands and Taiwan. The measurement was made by Chuang (1985,1986) with a moored current meter put at a level 20 m above the seabed. The mean northward component of the current was 24. 2 cm/s during the period of April 3 to May 5,1983 and 18.0 cm/s from March 11 to May 25,1985. The width of the passage is about 35 km and the mean depth about 55 m. Thus the volume transport

.

.

here is about 0. 4 Sv The current data at stations 2 to 8 are quoted from Fu et al (1989) and Guan (1980). These data are used to calculate the volume transport through the western Strait from the mainland to the Penghu Islands. The result is 0.6 Sv as shown in Table 1. The total volume transport in wintertime is thus 1.0 Sv.

TABLE 1 Calculation of volume transport through the Taiwan Strait 1. Winter (Dec. -May) Channel

Section

Width

Mean depth

(km)

(m) 45

P

250

Q

190

Mainland-Penghu Is.

Channel

Mainland-Taiwan

of N-wmp. b/s)

a.5 '25 }

Penghu Is. -Taiwan

2. Summer(June-Nov.

Mean speed

45 55

35

21. 1

Transport (SV)

0.6

0. 4

)

Section

R

3. Annual mean transport-2

Mean speed of

Width

Mean depth

(km)

(m>

(cm/s)

145

60

36

normal comp.

Transport (SV

)

3.1

Sv.

The current records in the southern part of the strait are inadequate to determine the summer volume transport passing through the section a c r w the Strait. There are a few stations, however,

347

4 2"

1

25

24'

23

22

Fig. 1.

Measured depth-mean current vectors in the Taiwan Stuait.

in the northern Strait as shown in Fig. 1. Station 13 is a mooring station maintained by Chuang (1985) and Wang and Chern (1988). Chuang obtained a mean current speed 28.5 cm/s during April 3 to May 5, 1983 at a level 14 m above the seabed. Wang and Chern gave a progressive current diagram for the period of July 19 to August 8 , 1984 at a depth 15 m below the sea surface. From the diagram a mean current speed of 55 cm/s is obtained. It is likely that the average of above two results gives an acceptable estimate for the vertically averaged current speed in summer. The current data at stations 10 to 12 are based on Wen et d. (1988) and Guan (1980). These records show that the current in the northern Taiwan Strait in summer is rather stable. The volum transport is calculated to be 3. 1 Sv. Thus the annual mean volume transport through the Taiwan Strait is about 2 Sv. VOLUME TRANSFORT OVER THE SHELF OF THE EAST CHINA SEA

There are 566 24-hour

anchored ship current measurements available in the East China Sea.

Most of these measurements were made with Ekman current meters. In addition to near-bottom measurements at 2 to 5 m above the seabed, measurements were done in most cases at depths of

348 120"

122"

126"

124"

130"

.. ..

0

0

0

.

0

0

.

0

'0

30" -

0

O 0 O b 0

0

0

0

do

28" -

-

0 sl

-

I

Fig. 2.

Position of anchored ship stations in the shelf part of the East China Sea.

5 , 10, 15, 20, 25, 30, 40, 50, 75, 100, and150mbelowtheseasurface, iftheselevelsare above the near -bottom layer. Figure 2 shows the measuring sites with each dot representing a current station. The seasonal distribution of the number of measurements is listed in Table 2 by days and by stations. It can be Seen that more measurements are carried out in spring and summer than in autumn and winter, but the difference is not very large. TABLE 2 Seasonal distribution of the number of measurements ~

Season

Days

Stations

spring

170

a5

Summer

193

98

Autumn

ioa

65

Winter

95

5a

Total

566

138

349 118" I

120"

122" I

124'

126"

I

I

128"

130"

34'

a-

32

'0

B

30

28

26

24

Fig. 3. Transport density in the shelf part of the East China Sea averaged for whole year. A , B and C are sections chosen to calculate stream function (see Fig. 8).

For each station and each measuring layer, we first make harmonic analysis of tidal current to get residual current (Fang, 1981). This approach is superior to 24-hour influence of the lunar tides Mz and

0 1 can

averaging in that the

be removed. Then for each 24-hour observation at a

site we calculate the transport density Q (The quantity Q has many names, such as transprt ,total transport, or current amount. In this paper we use transport density) by the following formula:

where K is the number of measuring layer ,zt and Vt the depth and residual current vector at the k -th layer, and h the water depth. The obtained 566 vectors of Q are grouped and averaged for

350

34'

32'

30

/ // //

28

D 6

4

Fig. 4.

Transport density in the shelf part of the East China Sea averaged for spring.

each 2. 2" latitudex 2. 2" longitude square centered at the points 1"apart. The vectors of the averaged transprt density are plotted in Fig. 3. Figure 4 to 7 show the averaged transport density for four seasons. It can be seen that all vectors of annual mean transport density are directed northa t w a r d except located at the northeast corner of the study area, where the current is likely to have an anticlockwise deflection. As for the seasonal means, the overall tendency is the same as that of the annual mean. It is especially noteworthy that the volume transport in winter is also northeastward regardless of the southward blowing wind. To give a quantitative estimate for the volume transport over the continental shelf of the East China Sea we choose three sections in the area as shown in Fig. 3. At each Sectton we take several

calculation points and calculate the transport density at these points by interpolation. Then the components normal to the section, Qu, are used to calculate the value of stream function defined by

35 1 118O

122"

124"

126"

128'

130"

As in Fig. 4 but for summer.

Fig. 5.

where

120"

5

is directed to the southeast along the section originating from the coast. The results are

shown in Fig. 8. From these results we obtain the distribution of stream function over the shelf of the East China Sea as plotted in Fig. 9. Figure 9 shows that the isopleth of 2 Sv is located slightly outside of the isobath of 100 m. Considering that in the above calculation more data are taken in summer and spring than in winter and autumn, and that most winter data were obtained under relatively calm weather condition, the transport in Fig. 9 should be slightly overestimated. Thus we can figure that the volume transport through the shelf part of the East Cbina Sea west 150 m isobath is about 2 Sv.

DISCUSSION;TAIWAN-TSUSHIMA-TSUGARU MAIN DRIVING FORCE

WARM CURRENT SYSTEM AND ITS

In the above sections we obtained the volume transport through the Taiwan Strait and the shelf

of the East China Sea. Miita and Ogawa (1984)calulated the volume tranport through the Tsushi-

ma Strait (Korea Strait) also from direct current measurements. The averaged volume transport is

352

34

32

30

//

28

d4 0 d

26

24'

Fig. 6.

As in Fig. 4 but for autumn.

1.8 Sv through the western channel and 1.66 Sv through the eastern channel. But most of the data used in their calculation were taken in summer (62%) and spring (21%). Soon later Miita (Coastal Oceanography Research Committee, The Oceanograplucal Society of Japan,

1985) got

an annual mean transport, 2.77 Sv from the same data with seasonal correction. On the other hand, recent studies indicated that the Tsushima Current did not directly come from the Kuroshio

near Kyushu (Rikiish and Ichiye, 1986). Fang and Zhao (1988) proposed that a current might run parallel to the Kuroshio from the northeastern part of the South China Sea to the area east of Hokkaido passing the East China Sea and Japan Sea by the Taiwan, Tsushima , Tsugaru and Soya Strait. In the present study we have obtained the transport through the Taiwan Strait and the area west of the 150 m isobath in the East China Sea to be about 2 Sv. The remaining transport (0.8 Sv) of the Tsushima Current should also be from the Kuroshio but is likely not through the Taiwan

Strait. We p r o p here to call this current the Taiwan-Tsushima-Tsugaru tem, abbreviated to T-T-T

WCS,which Ichiye (1989) called it T-T-T-S

rate the Soya Strait. A schematic representation of the T-T-T

Warm Current SysWCS to incorpo-

WCS is given in Fig. 10.

353

118"

120"

122"

124"

126"

128"

130" I

I

34' /

/

32

/

@. d J f

Q

P'

0

30

28'

26'

24'

Fig. 7.

As in Fig. 5 but for winter.

The Tsushima Current in the Japan Sea is seperated from the Kuroshio by the Japanese Islands.

But the current from the Taiwan Strait to the Tsushima Strait keeps in direct contact with the Kurahio. Thus strong water exchange may occur between these two currents. The driving mechanism of the Tsushima Current and the current in the Taiwan Strait has attracted many oceanographers' attention. Chuang (1985) assumes that the current in the Taiwan Strait is driven by the sea surface slope and wind. After a regression analysis he concludes that there should exist a sea surface slope from the south down to the north with a slope varying from 1.5 X lo-lto 2. 33 X lo-'.

This sea surface

slope was further verified based on land leveling data by Fang and Zhao (1988), who point out that the sea surface height along the southeast coast of China is higher in the southwest than in the northeast with a slope of 2.62 X lop7. The forces from the mean wind stress, atmospheric pressure and water density difference have effects opposite to the sea surface slope, but they altogether only amount to about one third of the force caused by the sea surface slope. The sea surface slope is certainly the main factor driving the northeastward current off the southeast coast of China. Ho-

354

Fig. 8.

Year-mean value of stream function along section A,B and C.

wever , there is still a question I what causes the sea surface slope? It is quite evident that the current in the Tsugaru Strait is also driven by sea surface slope along the Strait. There are several tidal stations near the Tsugaru Strait (Nakano and Yamada, 1975). Iwasaki and Hachinohe are located on the coast of northern Honshu with the former by the Japan

Sea and latter by the Pacific Ocean. The observed mean sea level at Iwasaki is 15 cm higher than at Hachinohe. Because both of them are on the southern side of the Tsugaru Current, this sea level difference can be regarded as the sea surface drop along the Strait. The mean sea level at Hakcdate on the northern side of the Tsugaru Current is lower by 23 cm than the mean sea level at Iwasaki and Hachinohe. However, this sea level difference only reflects the geostrophic balance in the crass-strait direction and has nothing to do with the driving mechanism. Along the northern coast of Hokkaido there is also a sea surface drop of 1 4 cm from Wakkanai to Abashiri. This indicates that the current in the Soya Strait is also driven by sea surface slope. But there is still a question : why the sea level on the Japan Sea side is higher than on the Pacific Ocean side? Many oceanographers have also attributed the Tsushima Current to the difference between the sea surface heights at the Tsushima and Tsugaru Straits. Minato and Kimura (1980) first consid-

ered the sea level variation along the western boundary of the ocean on the basis of Stommel's (1948) wind-driven

ocean circulation model. Sekine (1988) improved Minato and Kimura's

model by introducing the effect of density stratification. However, as pointed out by Ichiye (1984), the argument based on barotropic effect has two essential difficulties. First, the sea level

355

34

32'

30

2a

26'

24'

Fig. 9.

Year-mean stream function (in Sv=106m3s-').

difference derived from the wind-driven ocean circulation model is too small. The annual mean of the sea level drop from 30"N to 42"N is only 8 cm according to the calculation by Ichiye, which is even inadequate to account for an observed sea level drop along the Tsugaru or Soya Strait. Sec-

ond, the seasonal variation in the sea level difference is almost opposite in phase to the observed

seasonal variation in both observed sea level difference and volume transport of the Tsushima Current, and even negative in autumn (Ichiye, 1984, Sekine, 1988). Fang and Zhao (1988) noticed a large difference of the sea level along the western boundary of the ocean resulting from the thermohaline effect. Figure 11 shows the sea surface dynamic height relative to 1000 db in the northwestern Pacific Ocean, redrawn from Wyrtki (1975). This figure indicates that the sea surface height in the northeastern South China Sea is about 70 cm higher than that east of the Tsugaru/Soya Straits provided that the 1000 db pressure surface mincides with the geopotential surface. This assumption should be basically true because the current in deep layers such as down to 1000 m is generally much weaker than the geostrophic current at the

sea surface. Our diagnostic numerical model of the northwestern Pacific circulation reveals even

356

120"

130"

140"

150"

40

30

Fig. 10.

Schematic representation of the Taiwan-Tsushima-Tsugaru

Warm Current System.

a larger sea level difference along the west boundary. A sea level difference of 70 cm or so can produce a sea surface slope of order of 2X

This along-stream slope is comparable to that

observed or derived in the Taiwan Strait (Fang and Zhao, 1988; Chuang,1985) and is thus large enough to drive the T-T-T coast of China (--3Ocm)

WCS. And furthermore, the sea surface drop along the southeastern and along the Tsugaru Strait (-15

c m ) m be reasonably included in

this sea level difference. With this sea level difference as a basis, the seasonal variation of the transport of the T-TT WCS can be more easily explained. Because the steric height in the deep ocean varies little from w o n to season, in summer the southerly monsoon and the wax of vertical stratification of sea water enhance this current, while in winter the northerly monsoon and the wane of vertical stratification weaken it.

357

Fig. 11. The dynamic height of the sea surface of the northwestern Pacific relative to 1000 db (in dynamic cm). Redrawn from Wyrtki (1975). ACKNOWLEDGEMENTS The authors are especially grateful to Prof. Takashi Ichiye for his interest and valuable comments. Thanks are extended to Prof. Zilang Fu of Amoy University for making the current data in the Taiwan Strait available to the authors. The authors also thank the reviewers and the editor for their corrections and improvements to this paper.

REFERENCES

, 1985. Dynamics of subtidal flow in the Taiwan Strait. Journal of the 0ceanographical Society of Japan. 41 :65-72. Chuang, W. -S. , 1986. A note on the driving mechanisms of current in the Taiwan Strait. Journal of the Oceanographical Society of Japan. 42: 355-361. Coastal Oceanography Research Committee, the Oceanogcaphical Society of Japan, 1985. Coastal Oceanography of Japanese Islands (in Japanese). TOWUniversity press. 1106 pp. Fang , G. , 1981. Quasi-harmomic constituent method for analysis and prediction of tides, IU A practical procedure for analysing tidal stream and tidal elevations (in Chinese with English abstract). Studia Marina Sinica. 18: 19-39. Fang , G. and B. Zhao, 1988. A note on the main forcing of the northeastward flowing current off the southeast China coast. Progress in Oceanography. 21: 363-372. Chuang, W. - S .

.

358

Fu, Z. ,J. Hu and G. Yu, 1989. Winter volume transport in the Taiwan Strait (Manuscript, in Chinese).

Guan .B. .1980. Oceanic current (residual current)in the Minnan (Southern Fujian) fishing &o&d (Manuscript,in Chinese). Guan ,B. , 1984. Major features of the shallow water hydrography in the East China and H u m hai Sea. In : T. Ichiye (Editor), Ocean Hydrodynamics of the Japan and East China Elsevier, Amsterdam, pp. 1-14. Guan ,B. 1986. Evidence for a counter-wind current in winter off the southeast coast of chiM.Chinese Journal of Ooeanology and Limnology. 4: 319-332. Ichiye, T. , 1984. Some problems of circulation and hydrography of the Japan Sea and the Tsushima Current. In :T. Ichiye (Editor) ,Ocean Hydrodynamics of the Japan and East China Sea. Elsevier, Amsterdam, pp. 15-54. Ichiye , T. , 1989. J E C S news No. 1: On JECSS V. Journal of the Oceanographical Society of Japan. 45: A79-82. Minato ,S. and R. Kimura, 1980. Volume transport of the western boundary current penetrating into a marginal sea. Journal of the oceanographical Society of Japan. 36: 185- 195. Nakano, M. and S. Yamada, 1975. On the mean sea levels at various locations along the coasts of Japan. Journal of Oceanographical Society of Japan. 31 : 71 -84. Rikiishi, K. and T. Ichiye, 1986. Tidal fluctuation of the surface currents of the Kuroshio in the East China Sea. Progress in Oceanography. 17: 193-214. Sekine, Y . , 1988. On the seasonal variation in in-and outflow volume transport of the Japan Sea. Progress in Oceanography. 21 :269-279. Wang ,J. and C. -S. Chern, 1988. On the Kuroshio branch in the Taiwan Strait during wintertime. Progress in Oceanography. 21 : 469-491. Wen,X. ,L. Huang,H. Lian and H. Li,1988. Hydrographic features of the middle and northern part of the Taiwan Strait (in Chinese). In: Report of Marine Comprehensive Survey of the Middle and Northern Taiwan Strait. Science Press, Beijing, pp. 138-188. Wyrtki, K. , 1975. Fluctuations of the dynamic topography in the Pacific Ocean. Journal of Physical Oceanography. 5: 450-459.

-.

,

359

A SUBSURFACE NORTHWARD CURRENT OFF MINDANAO IDENTIFIED BY DYNAMIC CALCULAT ION D.X.HU, M.C.CU1, T.D.QU and Y . X . L I Institute of Oceanology, Academia Sinica, 7 Nanhai Road, Qingdao, Shandong, China ABSTRACT On the basis of CTD data along a section perpendicular to Mindanao by RIV Science I in October 1987 to 1989. a subsurface northward flow is clearly identified by dynamic calculation, which is named the Mindanao Undercurrent (MUC). The MUC is of a double-core feature. The first core is located about 80 km away from the coast of Mindanao in 1987 and 1989 and close to the coast in 1988, and centered at a depth of about 500 m with a width of about 50 k m , where the velocity is greater than 5 cmls. The maximum velocity of the first core can be over 20 cmls and the transport ranges from 3 to 9XlO'm'Is. The second is generally located about 200 km away from the coast, and is centered at a depth of about 370 in with a width of about 100 km. The maximum velocity can be over 30 cmls and the transport ranges from 5 to 16X10'm'/s. As a whole. the MUC transport ranges from 8 to 2 2 X lO'm'/s for the three years.

1 INTRODUCT IOh' The western

boundary current (WBC) in the north Pacific (east

is distinctly

Philippines)

different

from WBCs

in

the

other

of

the

oceans,

especially in terms of branching. There are two branches of the WBC east of the

Philippines-the

Kuroshio and the Mindanao Current, which should

significant effect upon the climate in East Asia

and

upon

the

have

formation

and evolution of the warm pool in the western equatorial Pacific. So there have been many studies on them ( Cui and Hu, 1989; Guan, 1986, 1989a and 1989b3 Hu. 1989; Hu and Cui, 1989; Lukas, 1988; Masuzawa, 1969; Nitani, 1970 and 1972; Wyrtki, 1956 and 1961

).

However, most of those studies are

concerned with the northward or southward surface currents only and few deal with return flows (undercurrents). The present study focuses on the subsurface northward current off Mindanao,based on geostrophic calculation.

2 DATA AND METHOD Every

October

Oceanology In

the

1987 to 1989, the Academia

Sinica

Institute

of

(ASIO) carried out a cruise in the WBC area with R/V Science I ,

which provided (Fig.1)

from

us

a number of CTD data by Neil Brown CTD or Sea-Bird CTD.

following a CTD section along 7" 3 0 ' N perpendicular

to

is geostrophically analyzed by dynamic calculation. All

Mindanao the

CTD

casts only reached down to 1500 m . For the choice of the level of no motion

360

20"-

15"-

10" -

Fig.1 Study area in the Philippine Sea and 7" 3 0 ' N section (thick with station dots)

line

the profiles of AD, the difference of dynamic depth between two adjacent stations, are plotted in Fig.2. According to Defant (1965). the level of no motion was chosen as 1200 m like Nitani (19701, 1300 m and 1500 m , respectively. The MUC transports for these levels of no motion are listed i n Table 1 .

TABLE I Volume transport (XlO'm'/s) year level of no motion 1200 m 1300 m 1500 m

of the MUC for three levels of no motion

1987

1988

1989

14.2 15.6 18 .O

7.2

23.4 23.4

7.4 8.6

22.9

361

-10.00

10.oo

0

c

E v

dyn cm -13 0 :.OO " "

'

.

"

'

'

"

0

7.0

.-

c

6 v

1500

I

I

I

I

dyn cm -11 .oo

0

10.00

dyn cm

Fig.2. Vertical distribution of dynamic depth difference between stations in 1987Ca). 1988(b) and 1989(c) It motion

is apparent from Tab.1 that different choice of the

level

results in little difference of volume transport of the

1987, 1988, 1989, respectively);

adjacent

MUC

of

no (for

i.e., the maximum deviation from the mean

.

is 2.1XlOCm'/s ( in 1987 ) There is significant interannual variability, which is not caused by the choice of the level of no motion and is one of the MUC's features. In addition, d(AD)/dz is calculated for 1200, 1300 and

362

1500 m at all station pairs for the three years. I f we take the 3-year average of d(AD)/dz at each depth of 1200, 1300 and 1500 m , respectively, the minimum will take place at 1500 m , which is 1.39XlO.' (dyn.crn) / m . 5o.a depth of 1500 m was chosen as the level of no motion for the following

calculation. 3 RESULT5 The calculated geostrophic velocities are shown in Figs.3a,b,c, relative 0"E

c

E

c E

v

u

c CI a

c c, a

c)

0

1

13"[ 1500

,) {

,[

,

200

1,

I

13001 ,

127"

128"

I

J

a, 1984 400 distance (km) ,

1500 0

I

8

0

I

b I

distance 129"

1988 I

200

130"

(

kin)

1 1"E

1500 200

400

distance ( km)

Fig.3. Geostrophic velocity section relative to northward velocity over 5 cm/s)

1500

db

(shaded

is

363

to 1500 db.

It

is evident

from

generally consists of two cores.

that

Fig.3

The first core

the

Mindanao Undercurrent

is

about 80 km away from

the coast in genera1,with an exception of being close to the coast in 1988. It is about 50 km wide and about 700 m thick, centered at about 500 m depth with maximum velocity of 21 cmls and maximum transport of 9.2XlO'm'Is with average of about 6XlO'm'Is. The second one is about 200 km away from the Mindanao coast, about 100 km wide, about 900 m thick if it is defined by speeds greater than 5 cmls. It is centered around 370 m in depth with maximum velocity over 30 cmls and maximum transport over 16X 108m'/s. The characteristics of the Mindanao Undercurrent are summarized in Table n

L.

TABLE 2 Characteristics of the Mindanao Undercurrent, characteristics

max. vel. (cmls) and depth ( m )

transport (10'm'ls)

width(km) -

thickness (m) -*

21

17

1987 465

379

10

20

1988

724

258

20

31

1989

362

482

17

23

517

373

average

n

-

corel core2 corel core2 corel core2 corel core2

year

distance from coast (km) corel core2

9.2

8.8

75

80

920 1050 (230- (170- 110 1150) 1220)5*

230

3.2

5.4

50

80

600 530 (600- (1701200) 700)

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