UBC Faculty Research and Publications

Thorium-230 normalized particle flux and sediment focusing in the Panama Basin region during the last… Calvert, Stephen E.; Francois, Roger 2007

Your browser doesn't seem to have a PDF viewer, please download the PDF to view this item.

Item Metadata


52383-Francois_AGU_2006PA001357.pdf [ 2.97MB ]
JSON: 52383-1.0103619.json
JSON-LD: 52383-1.0103619-ld.json
RDF/XML (Pretty): 52383-1.0103619-rdf.xml
RDF/JSON: 52383-1.0103619-rdf.json
Turtle: 52383-1.0103619-turtle.txt
N-Triples: 52383-1.0103619-rdf-ntriples.txt
Original Record: 52383-1.0103619-source.json
Full Text

Full Text

PALEOCEANOGRAPHY, VOL. 22, PA2213, doi:10.1029/2006PA001357, 2007  Thorium-230 normalized particle flux and sediment focusing in the Panama Basin region during the last 30,000 years S. S. Kienast,1 M. Kienast,1 A. C. Mix,2 S. E. Calvert,3 and R. Franc¸ois3 Received 9 August 2006; revised 27 November 2006; accepted 5 December 2006; published 16 May 2007.  [1] Application of the 230Th normalization method to estimate sediment burial fluxes in six cores from the eastern equatorial Pacific (EEP) reveals that bulk sediment and organic carbon fluxes display a coherent regional pattern during the Holocene that is consistent with modern oceanographic conditions, in contrast with estimates of bulk mass accumulation rates (MARs) derived from core chronologies. Two nearby sites (less than 10 km apart), which have different MARs, show nearly identical 230Th-normalized bulk fluxes. Focusing factors derived from the 230Th data at the foot of the Carnegie Ridge in the Panama Basin are >2 in the Holocene, implying that lateral sediment addition is significant in this part of the basin. New geochemical data and existing literature provide evidence for a hydrothermal source of sediment in the southern part of the Panama Basin and for downslope transport from the top of the Carnegie Ridge. The compilation of core records suggests that sediment focusing is spatially and temporally variable in the EEP. During oxygen isotope stage 2 (OIS 2, from 13–27 ka BP), focusing appears even higher compared to the Holocene at most sites, similar to earlier findings in the eastern and central equatorial Pacific. The magnitude of the glacial increase in focusing factors, however, is strongly dependent on the accuracy of age models. We offer two possible explanations for the increase in glacial focusing compared to the Holocene. The first one is that the apparent increase in lateral sediment redistribution is partly or even largely an artifact of insufficient age control in the EEP, while the second explanation, which assumes that the observed increase is real, involves enhanced deep sea tidal current flow during periods of low sea level stand. Citation: Kienast, S. S., M. Kienast, A. C. Mix, S. E. Calvert, and R. Franc¸ois (2007), Thorium-230 normalized particle flux and sediment focusing in the Panama Basin region during the last 30,000 years, Paleoceanography, 22, PA2213, doi:10.1029/2006PA001357.  1. Introduction [2] Upwelling of subsurface waters in the eastern equatorial Pacific (EEP) results in elevated levels of primary production and particle export. Despite intensive research over the last several decades to reconstruct past variations in productivity and particle flux from surface waters to the seafloor, the interpretation of the sedimentary record in the EEP remains highly controversial (see Loubere [1999] and Paytan et al. [2004] for recent reviews). Linear sedimentation rates calculated from core chronology often indicate higher mass accumulation of sedimentary material during the Last Glacial Maximum (LGM), which has been taken to imply higher vertical fluxes of biogenic sediment and higher productivity during this time interval [e.g., Sarnthein and Winn, 1988; Paytan et al., 1996; Lyle et al., 2002]. In contrast, ‘‘constant flux’’ tracers such as 230Th and extraterrestrial 3He indicate that the vertical rain rates of partic-  1 Department of Oceanography, Dalhousie University, Halifax, Nova Scotia, Canada. 2 College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, USA. 3 Department of Earth and Ocean Sciences, University of British Columbia, Vancouver, Canada.  Copyright 2007 by the American Geophysical Union. 0883-8305/07/2006PA001357  ulate biogenic material changed little or even decreased during the LGM [Marcantonio et al., 1996, 2001; Paytan et al., 1996; Higgins et al., 1999, 2002; Loubere et al., 2004; Kienast et al., 2006; Bradtmiller et al., 2006]. A similar discrepancy exists between different reconstructions of past aeolian dust fluxes to the central equatorial Pacific. While calculations based on linear sedimentation rates yield no consistent relationship between dust fluxes and glacialinterglacial stages [Murray et al., 1995], recent 230Thnormalized data suggest that dust fluxes were systematically higher during cold periods [Anderson et al., 2006]. [3] Studies using constant flux tracers invoke lateral redistribution of sediment on the seafloor to explain changes in sediment mass accumulation rates. Redistribution of recently deposited material (i.e., syndepositional redistribution) from zones of sediment winnowing to zones to sediment focusing can significantly affect accumulation rates without impairing the stratigraphy or chronology of the sedimentary sequence and can thus be easily overlooked [Franc¸ois et al., 2004]. This process now appears much more common in the ocean than previously thought [e.g., Mollenhauer et al., 2003, 2005; Franc¸ois et al., 2004, and references therein; Mitchell and Lyle, 2005; Kienast et al., 2005], and failing to take it into account could obviously lead to significant errors if mass accumulation rates are used to evaluate past changes in the vertical rain rate of particles from overlying surface waters.  PA2213  1 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  Figure 1. Core sites used in this study. Bathymetry from Smith and Sandwell [1997], (http:// topex.ucsd.edu/marine_topo/mar_topo.html).  [4] Normalization to the activity of the radioisotope 230Th has been proposed as a means to more accurately determine the vertical particle flux originating from surface waters [Bacon, 1984]. This approach relies on the assumption that the flux of 230Th to the seafloor approximates its rate of production by radioactive decay of 234U. Since it was first proposed, 230Th normalization is increasingly being used and its underlying assumptions are being tested and refined (see Franc¸ois et al. [2004], Henderson and Anderson [2003], and Franc¸ois et al. [2007], for recent reviews). Model simulations [e.g., Henderson et al., 1999; Marchal et al., 2000] and sediment trap observations [e.g., Yu et al., 2001, Scholten et al., 2005] confirm that the flux of scavenged 230Th to the seafloor is within 30% of its rate of production over most areas of the ocean. Laboratory experiments reveal that while 230Th has different affinities for different particle types (clay, biogenic opal, manganese oxide, and carbonate), all particle types strongly scavenge 230 Th [Geibert and Usbeck, 2004]. [5] Using seismic data, Lyle et al. [2005] have recently questioned the validity of the interpretation of 230Thnormalized fluxes in the EEP, arguing that 230Th normalization grossly overestimates sediment redistribution. Here, we present new 230Th and compositional data from 4 cores from the Panama Basin region (Figure 1 and Table 1). We also revisit sites Y69-71 [Loubere et al., 2004; Lyle et al., 2005] and P7 [Pedersen et al., 1988; Yang et al., 1995] from the Panama Basin. Our results further validate the 230Th normalization method by showing that (1) 230Th-normalized fluxes of bulk sediment and organic carbon during the Holocene are regionally consistent with present-day sea surface conditions in the EEP, (2) two nearby sites with different sediment mass accumulation rates produce nearly  identical 230Th-normalized bulk fluxes, and (3) the difference in focusing factors derived from 230Th accumulation rates at these two sites is consistent with the difference in sediment accumulation rates estimated from seismic data [Lyle et al., 2005]. During the last glacial period (oxygen isotope stage 2, or OIS 2 from 13– 27 ka BP), sediment redistribution appears to have been more intense, and we submit two possible explanations for this observation.  2. Core Sites and Oceanographic Setting [6] Cores ME0005-24JC and Y69-71 were recovered within 10 km of each other from an area of abyssal hills just to the north of the Carnegie Ridge (Figure 1 and Table 1). ME0005-24JC is located 200 m deeper than core Y69-71 in an abyssal valley [Mix et al., 2003], whereas Y69-71 was recovered midslope on an adjacent abyssal hill to the north, farther away from the Carnegie Ridge [see Lyle et al., 2005, Figure 1]. Core ME0005-24JC served as a site survey core for ODP Site 1240, which was subsequently drilled at this location. Core P7 was raised several hundred kilometers to the northeast from an area of relatively low  Table 1. Cores Used in This Study Core ID  Water Depth, m  Longitude  Latitude  ME0005-24JC Y69-71P P7 ME0005-27JC TR163-31P TR163-19P  2941 2740 3085 2203 3209 2348  86°27.7880W 86°290W 83°59.180W 82°47.200W 83.95 °W 90°57.10W  0°01.3020N 0°060N 2°36.260N 1°51.2010S 3.60 °S 2°15.50N  2 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  relief. Core ME0005-27JC was raised on the southern flank of the Carnegie Ridge at 2203 m water depth (wd), from a bench that gently slopes to the south, approximately 20 km south of the main scarp of the Carnegie Ridge. ODP Site 1238 was subsequently drilled at this location. Further to the south, core TR163-31P was recovered from 3209 m wd. Finally, core TR163-19P was raised from 2348 m wd on the outer flank of the Cocos Ridge, which forms the western boundary of the Panama Basin. [7] The surface waters in the Panama Basin region are influenced by wind-driven upwelling of water from the equatorial undercurrent (EUC). Off Peru, nutrient-rich water originating from the deeper parts of EUC stimulates intensive production in the equatorial cold tongue as it recirculates into the south equatorial current [Toggweiler et al., 1991; Toggweiler and Carson, 1995; Dugdale et al., 2002]. Core sites TR163-31P and ME0005-27JC are affected most by these highly productive waters. Further northwest, cores ME0005-24JC and Y69-71 are situated within the equatorial divergence zone, whereas cores P7 and TR163-19P are located at the northern boundary of the equatorial upwelling zone. The cores span a gradient in biological productivity with higher values (200 gC/m2/year) observed at sites TR163-31P and ME0005-27JC close to the Peru upwelling and lower values (140 gC/m2/yr) at sites P7 and TR163-19P at the northern limit of the shallow equatorial upwelling zone [Antoine et al., 1996].  3. Experimental Section 3.1. Analytical Methods [8] Total 230Th, 232Th and 238U activities were determined at the Woods Hole Oceanographic Institution (WHOI) and the University of British Columbia (UBC) by isotope dilution on an inductively coupled plasma mass spectrometer (ICP-MS) following total acid digestion of sediment samples equilibrated with 229Th and 236U spikes. Details and principles of the procedure can be found in the work of Choi et al. [2001] and Franc¸ois et al. [2004]. Total 230 Th data for core Y69-71 are from Loubere et al. [2004] and P. Loubere (personal communication, 2005) but multiplied by 0.85 to correct for a standardization error at the time of measurement. For core P7, excess and decay corrected 230Th data were taken directly from Yang et al. [1995]. [9] Organic carbon (Corg) contents of cores ME0005-24JC, ME0005-27JC, TR163-19P and TR163-31P were obtained at UBC by determining total C by combustion gas chromatography and subtracting carbonate carbon values determined by coulometry [Calvert et al., 1993]. Carbonate (CaCO3) values were obtained from the coulometric CO2 determinations assuming no other carbonate-bearing phase was present. Precisions were ±3% of the reported values. For core Y69-71, Corg and CaCO3 values are taken from Lyle et al. [2002]; they were determined at Oregon State University (OSU) by the acidification/wet oxidation technique described by Weliky et al. [1983] with modifications described by Lyle et al. [2002]. Organic carbon values of core P7 are taken from Pedersen et al. [1988].  PA2213  [10] Biogenic opal (Siopal) of core ME0005-24JC was determined at UBC by extraction of silica from 20 mg subsamples by a 2M Na2CO3 solution at 85°C for 5 hours. Dissolved silica concentrations in the extract were determined by molybdenum blue spectrophotometry [Mortlock and Froelich, 1989]. Precision was ±4.3% (1s). Biogenic opal of core Y69-71 is taken from Lyle et al. [2002] and was obtained by determining Na2CO3-soluble Si by atomic absorption spectrophotometry [Lyle et al., 2002, 1988]. [11] Total Iron, Al, Mg, Si, and Mn contents of core ME0005-24JC were determined by ICP-OES analysis of subsamples fused in LiBO2 and redissolved in dilute HNO3. Accuracy was controlled by analyzing a wide range of international rock standards, and precisions were better than ±2%. Nonbiogenic Si in this core was derived by subtracting opaline Si (determined by alkaline extraction) from total Si (determined by ICP-OES). [12] Radiocarbon dates are based on accelerator mass spectrometry (AMS) of hand-picked N. dutertrei specimens and bulk carbonate samples and were carried out at the National Ocean Science Accelerator Mass Spectrometry Facility (NOSAMS) using standard procedures. 3.2. Core Chronologies [13] The accuracy of both mass accumulation rates (MARs) and focusing factors is strongly dependent on the availability of reliable core chronologies. A significant part of this work is thus dedicated to the presentation and discussion of age models and their effect on focusing factors and MARs. 3.2.1. Y69-71 [14] A new chronology for this core is presented here which is based on correlating the new high-resolution benthic d 18O record (not shown) to core MD95-2042 off Portugal, which has been matched to the GRIP ice core record [Shackleton et al., 2000]. This tuning is confirmed by 8 14C dates obtained on N. dutertrei and bulk carbonate (Table 2 and Figure 2). Radiocarbon ages were calibrated to calendar years using the software CALIB 5.0.2 (Stuiver and Reimer [1993]; deltaR = 167 ± 106 years), except for the oldest 14C date, which was calibrated using the radiocarbon calibration of Fairbanks et al. [2005]. The new age model results in age assignments which are up to $2000 years older before 22 ka BP compared to the age model used in Loubere et al. [2004], which was based on the previous low-resolution d 18O record and its correlation to SPECMAP (Figure 2). The differences in age models and their effects on focusing factors and MARs are discussed in section 4.3. 3.2.2. ME0005-24JC [15] The age model for core ME0005-24JC is based on correlating its carbonate record to that of core Y69-71 [see Lyle et al. 2005, Figure 2]. The resulting age model agrees well with 3 reservoir corrected (deltaR = 167 ± 106 years) and calibrated 14C dates on N. dutertrei during the deglaciation, but gives a significantly older age assignment than a radiocarbon-based age point during the glacial period (Figure 2 and Table 2). Whereas we use the carbonate based stratigraphy for the main part of this contribution, we also discuss MARs and focusing factors that result from a radiocarbon based age model (section 4.3). Note that the  3 of 19  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  PA2213  Table 2. Radiocarbon Data of Cores ME0005-24JC, -27JC and Y69-71P Core Depth, cm  14  C Age, years Calendar Age,a years Laboratory Number  N. N. N. N.  dutertrei dutertrei dutertrei dutertrei  81 230 291 351  9040 ± 40 13,900 ± 35 15,700 ± 85 18,050 ± 65  9560 ± 130 15,820 ± 230 18,440 ± 270 20,630 ± 190  ME0005-24JC OS-33897 OS-33192 OS-33898 OS-33899  N. N. N. N.  dutertrei dutertrei dutertrei dutertrei  41 91 120 160  7430 ± 40 13,650 ± 65 17,200 ± 70 22,200 ± 95  7730 ± 110 15,470 ± 220 19,750 ± 70 26,180 ± 110b  ME0005-27JC OS-33900 OS-33193 OS-33194 OS-33195  11 – 18 25 – 30 50 85 120 165 180 260  3240 ± 25 6110 ± 30 9760 ± 40 12,130 ± 180 14,700 ± 65 17,170 ± 350 19,150 ± 70 28,000 ± 190  2870 ± 120 6370 ± 120 10,210 ± 150 13,360 ± 200 16,830 ± 270 19,670 ± 510 21,950 ± 370 32,350 ± 170b  Y69-71P OS-45680 OS-45682 OS-33894 (45) OS-33191 (45) OS-33895 OS-33896  N. dutertrei N. dutertrei N. dutertrei Bulk CaCO3 N. dutertrei Bulk CaCO3 N. dutertrei N. dutertrei  Reference this this this this  study study study study  this this this this  study study study study  G. Mollenhauer and T. Eglinton, unpublished data, 2006 G. Mollenhauer and T. Eglinton, unpublished data, 2006 Clark et al. [2004] Clark et al. [2004] Clark et al. [2004] Clark et al. [2004] Clark et al. [2004] this study  a Calibrated to calendar ages with the software CALIB 5.0.2. [Stuiver and Reimer, 1993] using a reservoir correction of deltaR = 167 ± 106 yrs and the calibration data set of Hughen et al. [2004]. Calendar ages are rounded to the nearest 10 years. b Reservoir corrected (À467 years) and calibrated to calendar ages using Fairbanks et al. [2005].  radiocarbon age model requires linear extrapolation from the oldest 14C age fix point (at $20 ka) to the base of stage 2 at 27 ka. For the core top in the radiocarbon model, we adopted the age assignment derived from the correlated age model. 3.2.3. ME0005-27JC [16] The age model for core ME0005-27JC is based on 4 radiocarbon dates on N. dutertrei (Table 2), which were corrected for reservoir effects (deltaR = 167 ± 106 years) and then calibrated to calendar ages using CALIB 5.0.1, except for the oldest age, which was calibrated using the radiocarbon calibration compilation of Fairbanks et al. [2005]. Preliminary benthic d 18O data (not shown) confirm this stratigraphy. 3.2.4. P7 [17] The age model of core P7 is taken from Pedersen et al. [1988] and Yang et al. [1995] and is based on 8 radiocarbon dates. 3.2.5. TR163-19P [18] The age model for TR163-19P is based on planktonic d 18O [Lea et al., 2000] augmented by 2 radiocarbon measurements from the deglacial section of the core [Spero and Lea, 2002]. 3.2.6. TR163-31P [19] An age model for core TR163-31P based on a benthic d 18O record and its correlation to SPECMAP was presented by Martin et al. [2002]. More recently, however, Skinner and Shackleton [2005] suggested that the postglacial benthic d18O rise in the EEP lagged the global glacioeustatic d18O component by $2 ka because of local hydrographic effects in the deep Pacific. Skinner and Shackleton [2005] base their findings on tight radiocarbon control on Kasten core TR163-31B, the companion core of TR163-31P used in this study. We thus present MARs and  focusing factors based on the original stratigraphy [Martin et al., 2002] but also investigate the ‘‘Skinner effect’’ by subtracting 2 ka from core depths corresponding to the benthic d18O rise in TR163-31P.  Figure 2. Depth-age relationships of core Y69-71 based on the new stratigraphy used here (thick dashed line) as well as the previously used age model (thin dashed line, Loubere et al. [2004]) and of core ME0005-24JC (thick solid line). Open symbols are calibrated 14C dates (Table 2).  4 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  Figure 3. (a) Downcore 230Th-normalized fluxes and (b) mass accumulation rates (MARs) of bulk sediment over the last 30 ka for all cores used in this study. MARs are averaged over the time intervals from the core top to 9.5, 9.5– 13.4, 13.4– 21, and 21– 27 ka. Note different Y axes.  3.3. Mass Accumulation Rates [20] Bulk accumulation rates (MARbulk) are derived from the product of linear sedimentation rates (cm/ka) and dry bulk densities (g/cm3) averaged over chosen core sections (see 3.5). The dry bulk density of cores ME0005-24JC, -27JC, TR163-19P and À31P was estimated from the carbonate content using the relationship r = 1/(3.6 – 0.0279  %CaCO 3 ) published by Snoeckx and Rea [1994]. Dry bulk density data for core Y69-71 are taken from Lyle et al. [2002], and data for core P7 are taken from Yang et al. [1995]. Mass accumulation rates of Corg and carbonate are derived as MARX = MAR  (X)/100, where X is weight% Corg or CaCO3 averaged over the corresponding core sections. 3.4. The 230Th-Normalized Fluxes and Focusing Factors (y) [21] 230Th normalization relies on the approximation that the flux of scavenged 230Th reaching the seafloor is equal to its known production rate by 234U decay in the overlying water [Bacon, 1984; Suman and Bacon, 1989; Franc¸ois et al., 1990, 2004]. Sediment trap [Yu et al., 2001; Scholten et al., 2005] and modeling [Henderson et al., 1999; Marchal et al., 2000] studies indicate that deviations from this assumption do not exceed 30% over most oceanic regions. Consequently, there is a simple inverse relationship between the vertical flux of particulate material (Fv) and its scavenged 230Th activity which can be used to calculate preserved vertical fluxes from the activity of scavenged 230Th in the bottom sediment. Scavenged, or initial excess 230Th (xs230Th(0)), was derived by correcting the total 230Th activity for (1) contributions from detrital and authigenic sources using the activities of 232Th and 238U in the same samples and following principles summarized  by Franc¸ois et al. [2004]; and (2) radioactive decay since the time of deposition using the chronologies described above. Estimates of total vertical flux (prFV – bulk) are derived from prFV – bulk = b  z/ex230Th(0), where b is the production of 230Th from 234U decay in the water column (0.0267 dpm mÀ3 yÀ1) and z is the water depth (m). Preserved vertical fluxes of sedimentary constituents such as Corg (prFV – Corg), carbonate (prFV –CaCO3), and opal (prFV – opal) are derived from prFVi = prFV – bulk  i/100, where i is weight percent Corg, carbonate, or opal. [22] In settings were particles settle through the water column and accumulate on the seafloor without lateral redistribution, the inventory of scavenged 230Th in the sediment should match the 230Th production in the overlying water column integrated over the time of sediment accumulation. 230Th inventories in excess of production, on the other hand, are evidence of sediment focusing, while 230Th inventories lower than predicted from production are indicative of sediment loss, or winnowing, from the core site. Focusing factors (y : ) are thus calculated as the decay corrected sedimentary accumulation rate of scavenged 230Th divided by its rate of production in the overlying water column integrated over the time interval represented by the core depth interval [Suman and Bacon, 1989]. Zr2 Y ¼ r1  À230  Á Thxs;0 rr dr  bzðt1 À t2 Þ  where r1 is sediment depth (cm), t1 is the corresponding age (ka), and rr is mean dry bulk density (g/cm3). The xs230Th0  5 of 19  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  PA2213  Figure 4. (a) Downcore 230Th-normalized fluxes and (b) MARs of organic carbon over the last 30 ka. MARs are averaged over the time intervals from the core top to 9.5, 9.5– 13.4, 13.4 – 21, and 21– 27 ka. Note different Y axes. and rr are averaged over the depth interval z1 to z2. Note that we recalculated y values for core Y69-71 here on the basis of 230Th values from Loubere et al. [2004], multiplied by 0.85 (see section 3.1) and using the new chronology described previously (see section 3.2). 3.5. Time Intervals Used for MAR and Focusing Calculations [23] We calculated MARs and focusing factors for OIS 1 (core top to 13 ka BP) and for OIS 2 (13 – 27 ka B.P). Note that for all cores, 0 cm core depth does not equal zero years. Although the definition of shorter time intervals may not necessarily be justified, particularly in cores TR163-31, TR163-19 and ME0005-27, which have a lower number of age control points, we nevertheless obtained MARs and focusing factors for the core top to 9.5, 9.5 to 13.4, 13.4 to 21, and 21 to 27 ka BP time intervals in order to facilitate comparison with the results of Loubere et al. [2004].  4. Results 4.1. The 230Th-Normalized Fluxes [24] Th-normalized fluxes of bulk sediment in the Panama Basin region range from lower values during the Holocene to higher values during OIS 2 (Figure 3a). As expected for nearby sites, cores ME0005-24JC and Y69-71 show very similar Th-normalized bulk fluxes. Th-normalized fluxes of organic carbon range from low values during the Holocene to maximal values during the early deglaciation (Figure 4a). Note that the gradient between cores sites displayed by bulk and organic carbon fluxes is maintained throughout the last 30 ka. However, Th-normalized organic carbon fluxes differ by approximately 30% between nearby sites Y69-71 and ME0005-24JC (Figure 4a; see discussion in section 5.7).  4.2. Mass Accumulation Rates (MARs) and Sediment Focusing (y) [25] Traditional MARs are consistently higher than Thnormalized fluxes (Figures 3b and 4b). MARs of bulk sediment and Corg range from lower values during the Holocene to higher values during OIS 2. Both Corg and bulk sediment MARs display large and temporally variable gradients between core sites. During the 9.5– 13.4 ka time interval, for example, bulk-MARs at site ME0005-24JC are almost 3 times higher than at neighboring site Y69-71. [26] As expected, focusing factors reflect the same temporal and regional trends displayed by bulk MARs (Figure 5) when calculated over the same time intervals. Similar to bulk MAR, the focusing factor is maximal (y = 7.9) at site ME0005-24 during the 9.5– 13.4 ka time interval, implying that nearly 90% of the total sediment cannot be accounted for by vertical particle rain. This brief event is not evident in the other records. Holocene focusing factors are >2 at the foot of the Carnegie Ridge (sites ME0005-24JC and Y69-71) and Cocos Ridge (TR163-19P) indicating that more than 50% of the bulk sediment at these sites has been advected laterally during this period (Figures 5a and 5b). At all sites except ME0005-27 focusing factors are higher during OIS 2 (but see section 4.3). During OIS 2, y is 5.4 at ME0005-24JC, the highest measured for this period (Figure 5a). 4.3. Effect of Chronology on Focusing Factors [27] By definition, focusing factors and MARs are highly dependent on the age control of the sedimentary sequence. Figure 6a compares focusing factors and MARs for core Y69-71 calculated on the basis of the new stratigraphy used here and on the basis of the stratigraphy used previously [Loubere et al., 2004]. The shorter (in core depth) glacial interval based on the new stratigraphy significantly  6 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  Figure 5. Focusing factors for (a) the time interval of OIS 1 (core top –13 ka BP) and OIS 2 (13 –27 ka BP) and (b) the time intervals from the core top to 9.5, 9.5– 13.4, 13.4 to 21, and 21– 27 ka.  decreases the MAR and focusing factor obtained for the interval from 21– 27 ka BP. Acknowledging that the stage 2/3 boundary is difficult to define in d18O stratigraphies, we examine the effect of a ±1.5 ka age uncertainty at the OIS 2/3 transition (see error bars in Figure 6a). The lower focusing factors (4.4 –2.6) and MARs obtained in this way for the 21– 27 ka interval are not very different from those during the Holocene (y = 2.2). Similarly, if we were to reject the tuned age model altogether and rely on the  radiocarbon date (at $32 calendar ka BP) to infer the core depth corresponding to the stage 2/3 boundary at 27 ka, focusing in the 21– 27 ka time interval would decrease to 3.3, still higher than during the Holocene but significantly lower than previously estimated [y = 8.2, Loubere et al., 2004]. [28] On the basis of tight radiocarbon age constraints, Skinner and Shackleton [2005] recently suggested that the postglacial benthic d 18O rise in the EEP lagged the benthic d18O rise in the deep NE Atlantic by $4 ka and the global  Figure 6a. Comparison of bulk MARs (left) and focusing factors (right) of core Y69-71 calculated on the basis of the age model used here (solid line) and the previous age model (gray dashed line). Vertical lines indicate the error bars resulting from a ±1.5 ka uncertainty at the stage 2/3 boundary. 7 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  Figure 6b. Comparison of bulk MARs (left) and focusing factors (right) of core TR163-31P calculated using the stratigraphy of Martin et al. [2002] (gray dashed lines), which is largely based on d18O correlation to SPECMAP and based on adjusting this chronology by À2 ka during the deglaciation to account for the potential lag of deep Pacific d18O versus the global glacioeustatic component suggested by Skinner and Shackleton [2005] (solid lines). Vertical lines indicate the error bars resulting from a ±1.5 ka uncertainty at the stage 2/3 boundary.  glacioeustatic d 18O component by $2 ka because of local hydrographic effects in the deep Pacific. If we adjust the chronology of TR163-31P for this effect by subtracting 2 ka from core depths corresponding to the benthic d 18O rise,  less sediment is assigned to the glacial interval, making glacial MARs lower, which in turn translates into smaller glacial focusing factors (Figure 6b). Allowing an age uncertainty of ±1.5 ka at the OIS 2/3 transition could further  Figure 6c. Comparison of bulk MARs (left) and focusing factors (right) of core ME0004-24JC calculated on the basis of the stratigraphy used here (solid lines) and on the basis of a radiocarbon age model which requires extrapolation from the oldest radiocarbon date at $20.6 ka to the OIS 2/3 boundary (dashed lines). Vertical lines indicate the error bars resulting from a ±1.5 ka uncertainty at the stage 2/3 boundary. 8 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  reduce focusing factors and MARs (error bars in Figure 6b). Nevertheless, in the case of TR163-31P, the focusing factor and MAR during OIS 2 remain higher (y = 3– 3.6) than during OIS 1 (y = 2) (Figure 6b). [29] The age model of ME0005-24JC used in this study is based on correlation of its down core carbonate profile to that of Y69-71 (see section 3.2). Taking the same ±1.5 ka age uncertainty at the OIS 2/3 into account illustrates that focusing factors and MARs during OIS 2 might not be significantly different from those observed during the 0 – 9.5 kyr time interval (Figure 6c). If, on the other hand, we were to reject the tuned age model and rely exclusively on the radiocarbon dates, we would have to linearly extrapolate sedimentation rates from the 14C age at 20 ka to the base of OIS 2 at 27 ka. Doing so increases the sediment interval corresponding to the 21– 27 ka BP time interval, which in turn increases the corresponding MARs and focusing factors (y = 7.9 – 12.6) significantly over those observed in the 0 – 9.5 ka interval (y = 4.3, Figure 6c). Note also that the maxima in focusing and MAR between 9.5 and 13.4 ka BP in the carbonate based age model (Figures 3b and 5b) decrease in comparison with glacial time intervals when using the radiocarbon base age model (Figure 6c), once again illustrating the delicate dependence of focusing factors and MARs on core chronology. 4.4. Geochemistry [30] The down core record of ME0005-24JC exhibits a pronounced maximum in Fe/Al and Mg/Al ratios between 13.5 and 10 cal ka (Figures 7b and 7c), and Sinonbiogenic/Al and Mn/Al ratios are also elevated during this period (Table 3). This metal enrichment is also observed in the same time interval in nearby core Y69-71 (S. E. Calvert, unpublished data, 2006), but is absent in all the other cores considered here. The period of metal enrichment corresponds to the time of maximal y (Figures 5b and 7d) and MAR values (Figure 3b) in the carbonate based age model and a minimum in %CaCO3 content at site ME0005-24JC (Figure 7a). A more modest % CaCO3 minimum is also apparent at site Y69-71 at this time (Figure 7a).  5. Discussion 5.1. The 230Th-Normalized Fluxes and Focusing Factors in the Equatorial Pacific [31] The advantages of Th normalization are its relative insensitivity to stratigraphic errors and the possibility of a high-resolution, point-by-point determination of paleofluxes [Franc¸ois et al., 2004]. Earlier studies have also documented a much higher degree of consistency in the sedimentary record of adjacent cores when using 230 Th normalization versus MARs [e.g., Anderson et al., 2006; Loubere et al., 2004; Franc¸ois et al., 2004; Thomson et al., 1999; Frank et al., 2000, 1999; Ru¨hlemann et al., 1996]. The same is clearly true for this study, where Th-normalized fluxes display a much more coherent pattern between core sites than MARs. (Figures 3 and 4). Both Th-normalized fluxes and MARs follow the trend expected from modern surface productivity (Figure 8), although absolute MARs are much higher than Th-normalized fluxes and there is a  PA2213  large difference between MARs at closely adjacent sites Y69-71 and ME0005-24JC. The difference in MAR values between these two sites is even larger prior to the Holocene (Figures 3b). This observation suggests that MARs, at least at site ME0005-24JC, are not solely driven by vertical flux. [32] Using chirp subbottom and seismic reflection profiles over sites ME0005-24JC and Y69-71, Lyle et al. [2005] noted a consistent bias of sediment deposition to the deeper basin over the entire depositional period ($2.5 Ma yrs), with site Y69-71 collecting sediment at a rate 30% slower than the deeper site ME0005-24JC. On the basis of the depth of a major seismic horizon which has a mean depth of 15.7 m in the subbottom profile, but is found at 16.7 m at site Y69-71 and 24.3 m depth at sites ME0005-24JC, Lyle et al. [2005] estimated focusing factors of $1.1 (Y69-71) and $1.5 (ME0005-24JC), assuming that sediment focusing is supplied by local sediment redistribution within the area acoustically surveyed. The relative difference between these estimates (1.1/1.5 = 0.7) is consistent with the difference in 230 Th-based focusing factors at both sites averaged over the last 27 ka (3.1 for Y69-71 versus 5.3 for ME0005-24JC; 3.1/5.3 = 0.6). Thus on this regional scale at least, there is good agreement between seismic and geochemical evidence for sediment focusing. Both methods thus show that, despite similar vertical fluxes at both sites (due to their proximity), there is more lateral input of sediment at the deeper site. The chief difference between the seismic and geochemical approach is that 230Th based estimates of focusing are much higher and suggest that both sites received an additional lateral input of sediment from outside the region acoustically surveyed. 5.2. Sources of Lateral Advection to Sites ME005-24JC and Y69-71 [33] We see two potential sources of laterally advected sediment to sites ME0005-24JC and Y69-71. The first one is general downslope transport of sediment from the surrounding Carnegie and Cocos Ridges into the southern boundary region of the Panama Basin. The second one is the episodic advection of hydrothermally influenced material, presumably from the Galapagos Spreading Center (see Figure 1 for location). 5.2.1. Downslope Transport [34] On the basis of seismic reflection profiles from the early 1970s, it has been known for some time that the sediment cover to the acoustic basement on the Carnegie and Cocos Ridges is thin and even partially absent [Van Andel et al., 1971]. Seismic imaging during ODP Leg 202 confirmed these observations for the eastern Carnegie Ridge [Mix et al., 2003]. Textural and compositional analyses revealed that the Carnegie and Cocos Ridges are indeed characterized by lag deposits mainly composed of coarse carbonate (intact planktonic foraminiferal shells), sandsized ash particles and hydraulically equivalent volcanic minerals, whereas fine carbonate and opal are transported into the Panama Basin by downslope transport [Van Andel, 1973; Moore et al., 1973; Kowsmann, 1973]. Substantial sediment transport was also deduced from observations of transverse dunes and barchans in an erosional valley excavated into the northern flank of the Carnegie Ridge  9 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  Figure 7. (a) Carbonate weight percent in core ME0005-24JC (solid gray line) and core Y69-71 (dashed gray line; data from Lyle et al. [2002]) and Th-normalized carbonate flux in core ME0005-24JC (solid black line) compared with (b) Fe, (c) Mg, and (d) focusing factors. Black lines indicate 230Thnormalized fluxes of Fe and Mg and gray lines indicate metal/Al ratios.  [Lonsdale and Malfait, 1974]. The valley itself is a several hundred meter deep depression cut into pelagic sediments. The dunes in the valley point northwestward, indicating down-valley sediment transport from the saddle in the central Carnegie Ridge into the Panama Basin. They are composed of sand-sized, broken (70%) and intact tests of Quaternary foraminifera [Lonsdale and Malfait, 1974]. [35] Although deep sea tidal currents are neither unidirectional nor continuous, they are often strong enough to entrain pelagic sediments and remove fine particles from the  seafloor (>10 cm/sec, Beaulieu [2002],) especially in areas of high relief and where flow is channeled [e.g., Lonsdale et al., 1972; Cacchione et al., 2002; Shanmugam, 2003]. Such resuspension would automatically lead to downslope transport and preferential sediment accumulation in troughs such as the abyssal valley where sites Y69-71 and ME0005-24JC are located. Indeed, tidal currents reaching 15– 20 cm/sec were observed in an erosional valley on the southern side of the Carnegie Ridge (Spiess et al. [1973], as cited by Lonsdale and Malfait [1974]), and barchans and related  10 of 19  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  PA2213  Table 3. Summary of Major and Minor Element Data of Hydrothermal Deposits From the Vicinity of the Galapagos Rift in the Panama Basin Compared With Average Continental Crust and Basalt  Fe/Al Mg/Al Si/Al Mn/Al  Upper Continental Crusta  Average Basaltb  Hydrothermal Depositsc  ME0005-24JC During Eventd  ME0005-24JC Before/After Evente  0.44 0.17 3.8 0.0074  1.3 0.54 2.8 0.0211  43 ± 8.7 2.9 ± 0.6 44.7 ± 4.3 0.99 ± 0.64  4.2 2.9 5.9f 0.63  0.75 0.5 2.5f 0.15  a  Taylor and McClennan [1985]. Faure [1991]. Ratio of average metal to average Al concentration in well-crystallized nontronites dredged from hydrothermal sediment mounds south of the Galapagos Spreading Center (samples N1 – N5 [Corliss et al., 1978]). d Maximal values during the %CaCO3 minimum event. e Average values before and after the %CaCO3 minimum event. f Opaline Si was subtracted from total Si. b c  dune forms generally indicate velocities of 40– 100 cm/sec in restricted passages [Masson et al., 2004, and references therein]. 5.2.2. Hydrothermal Input [36] An additional source of laterally advected sediment could be the hydrothermal mound deposits located in an area of high heat flow south of the Galapagos Spreading Centre [Lonsdale, 1977; Corliss et al., 1978], which is approximately 50 nautical miles to the north of sites ME0005-24JC and Y69-71 (Figure 1). The bulk chemical composition of ME0005-24JC and Y69-71 is significantly  different between 13.5 to 10 ka BP. During this time period, Fe/Al, Mg/Al, Mn/Al, and Si/Al ratios are elevated over values commonly observed in upper crust and basalt (Figures 7b, 7c, and Table 3). The hydrothermal deposits found south of the Galapagos Spreading Center are enriched in Fe-rich smectite (nontronite), which is characterized by extreme Al depletion (0.01– 1.12 weight%) signifying the presence of Fe in the octahedral as well as the tetrahedral sites of this mineral phase. The hydrothermal deposits are also highly enriched in Mn and Si, and somewhat enriched in Mg compared to upper crust or basalt. They are thought  Figure 8. Th-normalized fluxes and MARs of organic carbon (averaged from the age of the core top to 9.5 ka BP) versus modern primary production at the core sites [Antoine et al., 1996]. Both proxies display the regional gradient expected on the basis of modern conditions. However, MARs are larger and show a significant difference between sites ME0005-24JC and Y69-71. Site P7 is omitted because it only has one data point for the last 9.5 ka. 11 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  Figure 9. Focusing factors (Y) in the equatorial Pacific during (a) the Holocene, (b) the last glacial period, and (c) the ratio of glacial to Holocene focusing factors in the eastern equatorial Pacific. Data and references are reported in Table 4.  to form by deposition of Fe, Mn and Si from hydrothermal fluids, which circulate through the basaltic crust and overlying sediments in this region [Corliss et al., 1978]. The geochemistry of core ME0005-24JC thus points to the influence of allochthonous hydrothermal (nontronitic) material between 13.5 – 10 ka BP, which has diluted the carbonate fraction, thereby producing the marked %CaCO3 minimum in this core (Figure 7a). The addition of laterally sourced sediment during this brief time interval is consistent with the high focusing (Figure 7d) and MAR (Figure 3b) values based on the correlated age model at that time. It is also consistent with the slight decrease in Th-normalized CaCO3 fluxes during this time period (Figure 7a), as Th  normalization is thought to somewhat underestimate the vertical particle flux of individual sediment components in settings where a deep nepheloid layer advects material with a composition that is different from that found at the final deposition site [Franc¸ois et al., 2004]. Nearby core Y69-71 also displays a drop in %CaCO3 at this time (Figure 7a), as well as enrichments in Fe, Si and Mg (not shown) indicating that this site was also affected by hydrothermal input, albeit to a lesser extent. [37] These two sources of sediment, the Carnegie Ridge and the Galapagos Spreading Center lie outside the subbottom profile presented by Lyle et al. [2005], suggesting that their acoustically surveyed section might not represent a  12 of 19  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  PA2213  Table 4. Focusing Factors (Y) in the Central and Eastern Equatorial Pacific Core id  Latitude  Longitude  Depth, m  Y Holocene  TT013 MC112 TT013 MC34 TT013 MC27 TT013 MC97 TT013 MC69 TT013 MC19 61 BX1 MANOP B25 TT013-PC18 TT013-PC72 MANOP B18 54 BX1 53 BX1 52 BX1 ODP-849B 50 BX1 RC13-110 ODP-846B ME0005-24JC Y69-71 P7 Y71-6-12  5.00N 5.00S 3.00S 2.00N 0.00N 2.00S 0.91N 1.05N 2.00S 0.00N 1.06N 1.01N 0.98N 1.01N 0.18N 1.03N 0.10N 3.10S 0.02N 0.10N 2.60N 16.45S  140.00W 140.00W 140.00W 140.00W 140.00W 140.00W 140.04W 139.02W 139.00W 139.00W 138.80W 133.70W 124.63W 118.20W 110.52W 104.41W 95.65W 90.82W 86.46W 86.48W 83.79W 77.57W  4560 4391 4540 4440 4475 4328 4401 4354 4298 4281 4212 4516 3857 3851 3473 3231 3307 2941 2740 3085 2734  1.0 1.1 1.1 1.1 1.9 1.9 1.4 0.7 2.8 2.3 0.9 1.5 1.8 1.2 2.1 1.8 2.0 2.1 4.9 2.5 1.2 1.4  1.9 0.8 0.7 2.3 0.6 2.0 3.0 1.8 3.0 2.3 1.6 3.7 5.5 3.6 1.9 0.9  1.4 1.1 0.3 1.0 0.7 1.3 1.7 1.5 1.4 1.3 0.8 1.8 1.1 1.4 1.6 0.6  ME0005-27JC VNT01-8PC  1.85S 0.03N  82.79W 110.48W  2203 3791  1.8 1.6  1.7 2.3  0.9 1.4  TR163-19 TR163-31 V21-40  2.26N 3.60S 5.52S  90.95W 83.95W 106.76W  2348 3209 3128  2.2 1.6 3.5  2.5 3.4 3.2  1.1 2.1 0.9  V19-28 V19-29 KLH 093 KLH 068 RC13-140  2.37S 3.58S 1.50N 1.23N 2.87N  84.65W 83.93W 102.06W 101.61W 87.75W  2720 3157 3260 2870 2246  2.0 2.4 2.1 1.3 2.9  3.5 2.4 2.5 1.6 3.6  1.8 1.0 1.2 1.2 1.2  RC11-238  1.52S  85.82W  2573  3.0  2.0  0.7  V19-30  3.38S  83.52W  3091  3.2  4.4  1.3  RC13-114  1.65S  103.63W  3436  3.8  2.2  0.6  closed system with respect to sedimentation. This implies that their claim that 230Th normalization overestimates sediment focusing is unwarranted. 5.3. Local Differences in the Degree of Focusing in the Panama Basin and Regional Sediment Mass Balance [38] The compilation of data from the Panama Basin in Figure 5 indicates that focusing factors are highly variable in time and space. High values are restricted locally to the area of the seafloor at the foot of the Carnegie Ridge (ME0004-24JC and Y69-71) and the Cocos Ridge (TR163-19P). They are less pronounced on the slope of the Carnegie Ridge (ME0005-27JC) and do not extend far into the Panama Basin (P7). This is consistent with the view that sediment focusing and winnowing are controlled by bathymetry and current flow on a local scale [e.g., Turnewitsch et al., 2004; Loubere et al., 2004; Scholten et al., 1994]. [39] On the basis of earlier findings of very high glacial focusing factors at site Y69-71 (y = 8.2; Loubere et al.  Y Glacial  Y Gl/Y Hol  Reference Berelson et al. [1997] Berelson et al. [1997] Berelson et al. [1997] Berelson et al. [1997] Berelson et al. [1997] Berelson et al. [1997] Yang et al. [1990] Marcantonio et al. [2001] Marcantonio et al. [1996] Marcantonio et al. [1995] Marcantonio et al. [2001] Yang et al. [1990] Yang et al. [1990] Yang et al. [1990] Pichat et al. [2004] Yang et al. [1990] Loubere et al. [2004] Loubere et al. [2004] this study this study this study S. S. Kienast et al., unpublished data, 2006 this study S. S. Kienast et al., unpublished data, 2006 this study this study S. S. Kienast et al., unpublished data, 2006 Lao et al. [1992] Lao et al. [1992] Frank et al. [1994] Frank et al. [1994] L. I. Bradtmiller et al., unpublished data, 2006 L. I. Bradtmiller et al., unpublished data, 2006 L. I. Bradtmiller et al., unpublished data, 2006 L. I. Bradtmiller et al., unpublished data, 2006  [2004], Lyle et al. [2005] disputed the validity of 230Thnormalized fluxes by arguing that such high sediment focusing extrapolated over the entire region where they have identified a glacial MAR maximum (a $ 6° square area) would necessitate a source region barren of sediment covering a 17° square area, which would be easily noticed but has not been observed in the equatorial Pacific [Lyle et al., 2005]. Our data show, however, in light of the spatial and temporal heterogeneity of focusing factors reported here, that extrapolating a focusing factor estimated from one site to the entire EEP cannot be justified. In addition, the original glacial focusing factors reported at site Y69-71 need to be revised down significantly (see section 4.3 and below), further decreasing the need for an inordinately large source region to account for the glacial increase in focusing. 5.4. Focusing Factors in the Central and Eastern Equatorial Pacific [40] Focusing factors reported thus far for the central and eastern equatorial Pacific region are summarized in Figure 9  13 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  and Table 4. At most sites, focusing factors are greater than one during the Holocene (Figure 9a). This is not surprising given that high-sedimentation sites are preferentially chosen for coring operations because of their expected high temporal resolution. The compilation also reveals that focusing factors are generally higher in the eastern equatorial Pacific, particularly in the vicinity of the Panama Basin. This observation is consistent with the view that topographic relief should favor downslope sediment transport and focusing. Note, however, that a more modest, but still prevalent focusing effect is also found at sites between 95° and 140°W and 0° and 2°N, far away from major topographic features. Another remarkable observation is the relative increase in focusing during the last glacial period apparent at many sites, especially those near topographic features (Figure 9c). 5.5. Possible Mechanisms for Higher Glacial Focusing Factors [41] We see two principal ways to explain the higher glacial focusing factors observed in the EEP. One possibility is that the glacial increase in focusing in this EEP is only apparent and partly or largely results from inaccurate age control. Focusing factors are essentially the accumulation rate of scavenged 230Th divided by its rate of production, and thus the largest source of uncertainty in estimates of focusing comes from uncertainties in age models and MARs. Very high focusing factors (y = 8.2) were reported for the 23– 27 ka time interval of core Y69-71 on the basis of an earlier chronology [Loubere et al., 2004]. Using the revised chronology now available for this core (see section 3.2), MARs and focusing factors are reduced here by more than a factor of 2 (y = 3.6) for this time interval. Allowing for a ±1.5 ka age uncertainty at the stage 2/3 boundary could either increase or further reduce focusing and MARs in this time interval (y = 4.4– 2.6), with the reduced values being much closer to Holocene ones (y = 2.2; Figure 6a). In core ME0005-24JC, which displays the highest overall focusing values and MARs observed in this study, a ±1.5 ka age uncertainty could similarly increase or reduce the difference between Holocene and glacial focusing and MARs (Figure 6c). In core TR163-31, the same error would still leave glacial focusing factors slightly higher than Holocene ones (Figure 6b). It is obvious that focusing factors and MARs are highly sensitive to the age model, and it is possible that the apparent increase during the glacial period observed here might be partly due to inaccuracies in age control. Many of the records compiled in Figure 9 are dated using d 18O stratigraphies, in which the base of glacial stage OIS 2 at 27 ka BP is notoriously difficult to identify. Additional uncertainty has recently been implied during the deglaciation and at the beginning of OIS 2. Skinner and Shackleton [2005] suggested that the postglacial benthic d 18O rise in the EEP lagged the benthic d18O rise in the deep NE Atlantic by $4 ka and the global glacioeustatic d 18O component by $2 ka because of local hydrographic effects in the deep Pacific. If true, this effect would have led to an overestimation of the thickness of the glacial sediment intervals in previous studies based on d 18O stratigraphies in the Pacific, rendering glacial focusing factors and MARs  PA2213  erroneously high. Previously published Holocene focusing factors and MARs might also not be free of uncertainties. Piston cores often overpenetrate the sediment during coring operations, which leads to a real age of the top most core sample that is unknown and non zero. Assigning an age of zero years to the core top, however, results in a longer time interval for the Holocene core interval. This results in MARs and focusing factors that are erroneously low, thereby further enhancing the apparent glacial-interglacial contrast in focusing factors and MARs. Analyzing multicores with undisturbed sediment surfaces together with piston/gravity cores, such as done in the Southern Ocean [Frank et al., 2000], would allow evaluation of this potential bias. [42] If, on the other hand, we assume that the increase in glacial focusing is real and not solely driven by uncertainties in age control, we need a mechanism that can explain this general increase in sediment redistribution during glacial periods. One possibility could be related to an increase in tidal dissipation during glacial stages. Wunsch [2003] suggested that tidally induced ocean mixing should have increased during the LGM because of an increase in deep ocean tidal dissipation when shallow energy sinks were removed by lower sea level stands. Recent modeling studies indeed suggest that globally averaged barotropic tidal kinetic energies and deep ocean dissipation were significantly greater for LGM bathymetries [Egbert et al., 2004; Arbic et al., 2004]. Since there is significant dissipation of tidal energy over rough topography in the open ocean [Egbert and Ray, 2000, 2001], this effect could have been particularly strong in the Panama Basin area in view of the rugged local seafloor topography. The overall increase in tidal dissipation would have caused an enhancement in abyssal tidal currents and is likely to have resulted in greater entrainment and redistribution of sediment particles. The tidally induced forcing of sediment redistribution proposed here is time variant and thereby addresses one of the requirements of Paytan et al. [2004] for a climatically driven physical mechanism to explain glacial-interglacial changes in sediment focusing. 5.6. Paleoproductivity Reconstructions in the EEP From Th-Normalized Fluxes of Organic Carbon [ 43 ] Despite the difference between MARs and 230 Th-normalized fluxes in the Panama Basin, both proxies show overall higher Corg fluxes in OIS 2. These findings are in qualitative agreement with previous studies pointing to higher export production during OIS 2 [Pedersen, 1983, 1988; Farrell et al., 1995; Yang et al., 1995; Herguera, 2000; Lyle et al., 2002]. However, 230Th normalized Corg fluxes are maximal during the deglaciation ($18– 13 ka BP) rather than the full glacial per se, suggesting enhanced export production or organic matter preservation at that time, which is consistent with a recent study identifying the time interval from 20– 15 ka BP as the period of most intense upwelling in the EEP over the last 40 ka [Martinez et al., 2006]. [44] On the other hand, higher glacial 230Th-normalized Corg fluxes are at odds with lower estimates of paleoexport production during OIS 2 in the Panama Basin based on a  14 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  Figure 10. 230Th-normalized organic carbon flux compared to paleoproductivity estimates based on benthic foraminifera [Loubere et al., 2003] at site Y69-71.  benthic foraminiferal transfer function [Loubere et al., 2003]. Both methods have been applied at site Y69-71, and lead to opposite results (Figure 10). This discrepancy is perplexing, and unresolved at the time of writing. Benthic foraminiferal assemblages are known to record the relative changes in the flux of labile organic carbon to the seabed [e.g., Kuhnt et al., 1999; Loubere and Fariduddin, 1999; Wollenburg and Kuhnt, 2000], as does the 230Th normalized organic matter burial. Further work will be needed to resolve this discrepancy. 5.7. Difference in 230Th-Normalized Organic Carbon and Opal Fluxes Between Sites Y69-71 and ME0005-24JC [45] 230Th-normalized organic carbon fluxes are approximately 30% higher at deeper site ME0005-24JC compared to Y69-71 from 27 to 9 ka BP (Figures 4a and 11a). A similar offset between both sites is observed in 230Thnormalized opal fluxes (Figure 11b). Sedimentation rates (Figure 2) and focusing factors (Figure 5b) are also significantly higher at ME0005-24JC than at site Y69-71 in the time interval from 27 to 9 ka BP. Thorium normalized fluxes of bulk sediment on the other hand which are dominated by carbonate are very similar at both sites throughout the last 30 ka (Figure 3a). The higher 230Th normalized biogenic fluxes at site ME0005-24JC compared to Y69-71 are related to higher weight percentages of these constituents at the former site, which translate directly into higher 230Th-normalized flux estimates (see equations in section 3.4). At least for organic carbon, analytical differences between ME0004-24JC and Y69-71 (see section 3.1) cannot fully account for the intersite difference, as an offset  remains even when using organic carbon data produced by the same laboratory (Figure 11a). We recognize three non exclusive possibilities to explain the difference in weight percentages and hence 230Th normalized fluxes of organic carbon and opal at these sites: [46] 1. Higher sedimentation rates at ME0005-24JC could have led to faster burial and better preservation of organic carbon and biogenic opal at this site compared to Y69-71 [sensu Mu¨ller and Suess, 1979]. [47] 2. A subtle increase in carbonate dissolution with depth would lead to increased weight percentages of organic carbon, opal as well as terrigenous material at site ME000524JC, which is 200 meters deeper than site Y69-71. Indeed, %opal and %terrigenous material (not shown) are slightly higher in core ME0005-24JC than in core Y69-71, where as %carbonate is slightly higher in core Y69-71 (Figure 7a). [48] 3. Preferential removal and advection of organic matter and opal concentrated in the fine sediment fraction to site ME0005-24JC offers a third explanation. In this scenario, the preferential advection of a minor sediment constituent such as organic carbon or opal would not significantly affect the 230Th-normalized flux of bulk sediment and major sediment constituents, but would increase the %Corg and % opal at the receiving site, resulting in higher 230Th-normalized values of these constituents (see section 3.4). [49] These possibilities need to be investigated further, particularly in light of the discrepancy with paleoproductivity reconstructions based on benthic foraminiferal transfer functions (Figure 10). It is possible that this discrepancy might be related, in part, to a possible preference of benthic foraminifera for very fresh organic matter delivered directly  15 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  PA2213  Hence we suspect that while preferential advection of organic carbon (and opal) under conditions of elevated focusing is possible, it is not the dominant control on the Th-normalized organic carbon and opal fluxes in core ME0005-24JC.  6. Summary and Conclusions  Figure 11. Th-normalized fluxes of organic carbon (a) and biogenic opal (b) at sites ME0005-24JC and Y69-71. Note that the difference in organic carbon fluxes between both sites largely remains when using Corg data produced by the same laboratory (OSU) for the flux calculations. Percent organic carbon data of ME0005-24JC used for Th-normalized flux calculations are from this study (red solid line), from Lyle et al. [2005], red dashed line), and from Lyle et al. [2002] for core Y69-71 (blue line). Percent biogenic opal data used for Th-normalized flux calculations are from this study (ME0005-24, red line) and Lyle et al. [2002], Y69-71, blue line). from the sea surface versus laterally advected organic matter which could be more refractory. However, we note that the pattern of 230Th-normalized Corg flux at ME0005-24JC strongly resembles that at other sites in the Panama Basin, including sites ME0005-27JC and P7, which are only marginally affected by focusing (Figures 4a, 5a, and 5b).  [50] The use of Th normalization as a tool for reconstructing past particle fluxes from the sea surface to the seafloor is still controversial [Lyle et al., 2005; Franc¸ois et al., 2007]. This study shows that 230Th-normalized fluxes of bulk sediment and organic carbon display a regionally coherent pattern in the EEP, which is consistent with present-day surface oceanographic conditions. Nearby sites show almost identical 230Th-normalized bulk fluxes and a 30% difference in 230Th-normalized Corg fluxes. Traditional MARs of bulk sediment and Corg, on the other hand, are significantly higher, and display large site-to-site variability. Thus the results of this study further validate 230Th normalization. [51] Holocene focusing factors are 2 – 4 at the foot of the Carnegie Ridge, suggesting that lateral sediment transport to this region is significant. This observation is consistent with earlier observations of lag deposits and bare rock on top of the Carnegie Ridge and of an erosional valley on the basinfacing slope of the ridge. The relative difference between focusing factors at two nearby sites the Panama Basin estimated by 230Th normalization and using seismic profiles [Lyle et al., 2005] is comparable. [52] Sites that are affected by focusing in the Holocene show even higher focusing factors during OIS 2, similar to previous findings in the central and eastern EEP. Examination of the effect of chronological uncertainties at the OIS 2/3 and 1/2 boundaries in our records shows that the magnitude of the glacial increase in focusing factors and MARs is very sensitive to chronological uncertainties. We thus offer two possible explanations for the glacial increase in focusing and MARs. The first one is that the apparent increase in lateral sediment redistribution is partly or even largely an artifact of insufficient age control in the EEP. The second explanation, which assumes that the observed increase is real, invokes enhanced deep sea tidal current flow during periods of low sea level stand [Wunsch, 2003]. Only improved stratigraphic control will help to distinguish between these two scenarios. [53] The controversy revolving around 230Th normalization was partly triggered by very high estimates of glacial focusing (y = 8.2) at site Y69-71 in an earlier study [Loubere et al., 2004]. Extrapolating from this site to the entire EEP, Lyle et al. [2005] assumed the need for a sediment source covering a 17° square area to account for the high focusing in the Panama Basin during the LGM, which is not observed in seismic surveys. Using improved chronological control at site Y69-71, we show here that glacial focusing values have to be revised down considerably (y = 4.4– 2.6), although lateral sediment input is still considerable. In addition, the large spatial and temporal heterogeneity of focusing factors observed in the EEP and elsewhere does not warrant extrapolation from one site to the entire region. Thus there appears to be no requirement  16 of 19  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  for an inordinately large area of seafloor devoid of sediment to account for the mass of laterally supplied sediment during the glacial. [54] Over the course of the last 30 ka, both 230Thnormalized fluxes and MARs of Corg show a general increase in OIS 2 compared to the Holocene. Th-normalized Corg fluxes are maximal in the deglaciation, similar to Thnormalized opal fluxes. This could indicate enhanced export production at that time, in conflict with results inferred on the basis of a benthic foraminifera transport function in the Panama Basin [Loubere et al., 2003]. [55] Although 230Th-normalized bulk fluxes are nearly identical at two nearby sites, there is a $30% difference in 230Th-normalized fluxes of organic carbon and opal between these sites. This apparent difference in 230Thnormalized fluxes is caused by a difference in weight% of these constituents at the two sites. The difference in weight%, in turn, could be caused by either changes in sedimentation rate or carbonate dissolution with depth. Alternatively, preferential advection of minor sediment  PA2213  constituents associated with fine-grained particles such as opal and organic matter might have increased the concentration of these constituents and their 230Th-normalized fluxes at the deeper site. These observations may potentially hold clues to help explain the difference between results of paleoproductivity reconstructions based on benthic transfer functions and other proxies.  [56] Acknowledgments. We gratefully acknowledge official reviews by Robert Anderson, Mitchell Lyle, Martin Frank, and one anonymous reviewer which helped to improve and clarify the manuscript. We also wish to acknowledge fruitful discussions with Zanna Chase, Gesine Mollenhauer and Tim Eglinton. Gesine Mollenhauer, Tim Eglinton, and Louisa Bradtmiller shared unpublished 14C data. Maureen Soon, Alan Fleer and Susan Brown-Leger provided invaluable technical assistance. Core material was provided by the Core Repositories of Oregon State University (supported by NSF grant OCE97-12024) and the University of Rhode Island (supported by NSF grant OCE-9102410). This research was supported by grants from the Canadian Foundation for Climate and Atmospheric Sciences (CFCAS), the Natural Sciences and Engineering Research Council (NSERC) Canada, and NSF.  References Anderson, R. F., M. Q. Fleischer, and Y. Lao (2006), Glacial-interglacial variability in the delivery of dust to the central equatorial Pacific, Earth Planet. Sci. Lett., 242, 406 – 414. Antoine, D., J. M. Andre, and A. Morel (1996), Oceanic primary production: 2. Estimation at global scale from satellite (coastal zone color scanner) chlorophyll, Global Biogeochem. Cycles, 10, 57 – 69. Arbic, B. K., D. R. MacAyeal, J. X. Mitrovica, and G. A. Milne (2004), Ocean tides and Heinrich events, Nature, 432, 460. Bacon, M. P. (1984), Glacial to interglacial changes in carbonate and clay sedimentation in the Atlantic Ocean estimated from 230 Th measurements, Isot. Geosci., 2, 97 – 111. Beaulieu, S. E. (2002), Accumulation and fate of phytodetritus on the seafloor, Oceanogr. Mar. Biol., 40, 171 – 232. Berelson, W. M., et al. (1997), Biogenic budgets of particle rain, benthic remineralization and sediment accumulation in the equatorial Pacific, Deep Sea Res., Part II, 44, 2251 – 2282. Bradtmiller, L. I., R. F. Anderson, M. Q. Fleisher, and L. H. Burckle (2006), Diatom productivity in the equatorial Pacific Ocean from the last glacial period to the present: A test of the silicic acid leakage hypothesis, Paleoceanography, 21, PA4201, doi:10.1029/ 2006PA001282. Cacchione, D. A., L. F. Pratson, and A. S. Ogston (2002), The shaping of continental slopes by internal tides, Science, 296, 724 – 727. Calvert, S. E., T. F. Pedersen, and R. C. Thunell (1993), Geochemistry of surface sediments of the Sulu and South China Seas, Mar. Geol., 114, 207 – 231. Choi, M. S., R. Franc¸ois, K. Sims, M. P. Bacon, S. Brown-Leger, A. P. Fleer, L. Ball, D. Schneider, and S. Pichat (2001), Rapid determination of 230Th and 231Pa in seawater by desolvated micro-nebulization inductively coupled plasma magnetic sector mass spectrometry, Mar. Chem., 76, 99 – 112. Clark, P. U., A. M. McCabe, A. C. Mix, and A. J. Weaver (2004), Rapid rise of sea level 19000 years ago and its global implications, Science,  304, 1141 – 1144. Corliss, J. B., M. Lyle, J. Dymond, and K. Crane (1978), The chemistry of hydrothermal mounds near the Galapagos Rift, Earth Planet. Sci. Lett., 40, 12 – 24. Dugdale, R. C., A. G. Wischmeyer, F. P. Wilkerson, R. T. Barber, F. Chai, M.-S. Jiang, and T.-H. Peng (2002), Meridional asymmetry of source nutrients to the equatorial Pacific upwelling ecosystem and its potential impact on ocean-atmospheric CO2 flux: A data and modeling approach, Deep Sea Res., Part II, 49, 2513 – 2531. Egbert, G. D., and R. D. Ray (2000), Significant tidal dissipation in the deep ocean inferred from satellite altimeter data, Nature, 405, 775 – 778. Egbert, G. D., and R. D. Ray (2001), Estimates of M2 tidal energy dissipation from TOPEX/ Poseidon altimeter data, Geophys. Res. Lett., 106, 22,475 – 22,502. Egbert, G. D., R. D. Ray, and B. G. Bills (2004), Numerical modeling of the global semidiurnal tide in the present day and in the Last Glacial Maximum, J. Geophys. Res., 109, C03003, doi:10.1029/2003JC001973. Fairbanks, R. G., R. A. Mortlock, T.-C. Chiu, L. Cao, A. Kaplan, T. P. Guilderson, T. W. Fairbanks, A. L. Bloom, P. M. Grootes, and M.-J. Nadeau (2005), Radiocarbon calibration curve spanning 0 – 50000 years BP based on paired 230Th/234U/238U and 14C dates on pristine corals, Quat. Sci. Rev., 24, 1781 – 1796. (Available at http://radiocarbon.ldeo. columbia.edu/research/radcarbcal.htm) Farrell, J. W., T. F. Pedersen, S. E. Calvert, and B. Nielsen (1995), Glacial-interglacial changes in nutrient utilization in the equatorial Pacific Ocean, Nature, 377, 514 – 517. Faure, G. (1991), Principles and Applications of Geochemistry, 2nd ed., 600 pp., Prentice-Hall, Upper Saddle River, N. J. Franc¸ois, R., M. P. Bacon, and D. O. Suman (1990), Thorium 230 profiling in deep sea sediments: High-resolution records of flux and dissolution of carbonate in the equatorial Atlantic during the last 24,000 years, Paleo-  17 of 19  ceanography, 5, 761 – 787. Franc¸ois, R., M. Frank, M. M. R. van der Loeff, and M. P. Bacon (2004), 23OTh normalization: An essential tool for interpreting sedimentary fluxes during the late Quaternary, Paleoceanography, 19, PA1018, doi:10.1029/2003PA000939. Franc¸ois, R., et al. (2007), Comment on ‘‘Do geochemical estimates of sediment focusing pass the sediment test in the equatorial Pacific’’ by M. Lyle et al., Paleoceanography, 22, PA1216, doi:10.1029/2005PA001235. Frank, M., J.-D. Eckhardt, A. Eisenhauer, P. W. Kubik, B. Dittrich-Hannen, M. Segl, and A. Mangini (1994), Beryllium 10, thorium 230, and protactinium 231 in Galapagos microplate sediments: Implications of hydrothermal activity and paleoproductivity changes during the last 100,000 years, Paleoceanography, 9, 559 – 578. Frank, M., R. Gersonde, and A. Mangini (2000), Similar glacial and interglacial export bioproductivity in the Atlantic sector of the Southern Ocean: Multiproxy evidence and implications for glacial atmospheric CO2, Paleoceanography, 15, 642 – 658. Frank, M., R. Gersonde, and A. Mangini (1999), Sediment redistribution, 230Thex-normalization and implications for the reconstruction of particle flux and export paleoproductivity, in Use of Proxies in Paleoceanography: Examples from the South Atlantic, edited by G. Fischer and G. Wefer, pp. 409 – 426, Springer, New York. Geibert, W., and R. Usbeck (2004), Adsorption of thorium and protactinium onto different particle types: Experimental findings, Geochim. Cosmochim. Acta, 68, 1489 – 1501. Henderson, G. M., and F. Anderson (2003), The U-series toolbox for paleoceanography, Rev. Mineral. Geochem., 52, 493 – 531. Henderson, G. M., C. Heinze, R. F. Anderson, and A. M. E. Winguth (1999), Global distribution of the 230Th flux to ocean sediments constrained by GCM modeling, Deep Sea Res., Part I, 46, 1861 – 1893. Herguera, J. C. (2000), Last glacial paleoproduc-  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  tivity patterns in the eastern equatorial Pacific: Benthic foraminifera records, Mar. Micropaleontol., 40, 259 – 275. Higgins, S. M., W. Broecker, R. F. Anderson, D. C. McCorkle, and D. Timothy (1999), Enhanced sedimentation along the equator in the western Pacific, Geophys. Res. Lett., 26, 3489 – 3492. Higgins, S. M., R. F. Anderson, F. Marcantonio, P. Schlosser, and M. Stute (2002), Sediment focusing creates 100-ka cycles in interplanetary dust accumulation on the Ontong Java Plateau, Earth Planet. Sci. Lett., 203, 383 – 397. Hughen, K. A., et al. (2004), Marine04 Marine radiocarbon age calibration, 26 – 0 ka BP, Radiocarbon, 46, 1059 – 1086. Kienast, M., M. J. Higginson, G. Mollenhauer, T. I. Eglinton, M. T. Chen, and S. E. Calvert (2005), On the sedimentological origin of down-core variations of bulk sedimentary nitrogen isotope ratios, Paleoceanography, 20, PA2009, doi:10.1029/2004PA001081. Kienast, S. S., M. Kienast, S. Jaccard, S. E. Calvert, and R. Franc¸ois (2006), Testing the silica leakage hypothesis with sedimentary opal records from the eastern equatorial Pacific over the last 150 kyrs, Geophys. Res. Lett., 33, L15607, doi:10.1029/2006GL026651. Kowsmann, R. O. (1973), Coarse components in surface sediments of the Panama Basin, eastern equatorial Pacific, J. Geol., 81, 473 – 494. Kuhnt, W., S. Hess, and Z. Jian (1999), Quantitative composition of benthic foraminiferal assemblages as a proxy indicator for organic carbon flux rates in the South China Sea, Mar. Geol., 156, 123 – 157. Lao, Y. L., R. F. Anderson, and W. S. Broecker (1992), Boundary scavenging and deep-sea sediment dating: Constraints from excess 230Th and 231Pa, Paleoceanography, 7, 783 – 798. Lea, D. W., D. K. Pak, and H. J. Spero (2000), Climate impact of late Quaternary equatorial Pacific sea surface temperatures variations, Science, 289, 1719 – 1723. Lonsdale, P. (1977), Deep-tow observations at the mounds abyssal hydrothermal field, Galapagos Rift, Earth Planet. Sci. Lett., 36, 92 – 110. Lonsdale, P., W. R. Normark, and W. A. Newman (1972), Sedimentation and erosion on Horizon Guyot, Geol. Soc. Am. Bull., 83, 289 – 316. Lonsdale, P. F., and B. T. Malfait (1974), Abyssal dunes of foraminiferal sand on the Carnegie Ridge, Geol. Soc. Am. Bull., 85, 1697 – 1712. Loubere, P. (1999), A multiproxy reconstruction of biological productivity and oceanography in the eastern Pacific for the past 30,000 years, Mar. Micropaleontol., 37, 173 – 198. Loubere, P., and M. Fariduddin (1999), Benthic foraminifera and the flux of organic carbon to the seabed, in Modern Foraminifera, edited by B. S. Gupta, pp.181 – 200, Springer, New York. Loubere, P., M. Fariduddin, and R. W. Murray (2003), Patterns of export production in the eastern equatorial Pacific over the past 130,000 years, Paleoceanography, 18(2), 1028, doi:10.1029/2001PA000658. Loubere, P., F. Mekik, R. Franc¸ois, and S. Pichat (2004), Export fluxes of calcite in the eastern equatorial Pacific from the Last Glacial Maximum to the present, Paleoceanography, 19, PA2018, doi:10.1029/2003PA000986. Lyle, M., D. M. Murray, B. P. Finney, J. Dymond, J. M. Robbins, and K. Brooksforce (1988), The record of late Pleistocene biogenic sedimenta-  tion in the eastern tropical Pacific Ocean, Paleoceanography, 3, 39 – 59. Lyle, M., A. Mix, and N. Pisias (2002), Patterns of CaCO3 deposition in the eastern tropical Pacific Ocean for the last 150 kyr: Evidence for a southeast Pacific depositional spike during marine isotope stage (MIS) 2, Paleoceanog r a p h y, 1 7 ( 2 ) , 1 0 1 3 , d o i : 1 0 . 1 0 2 9 / 2000PA000538. Lyle, M., N. Mitchell, N. Pisias, A. Mix, J. I. Martinez, and A. Paytan (2005), Do geochemical estimates of sediment focusing pass the sediment test in the equatorial Pacific?, Paleoceanography, 20, PA1005, doi:10.1029/ 2004PA001019. Marcantonio, F., N. Kumar, M. Stute, R. F. Anderson, M. A. Seidl, P. Schlosser, and A. Mix (1995), Express Letter: A comparative study of accumulation rates derived by He and Th isotope analysis of marine sediments, Earth Planet. Sci. Lett., 133, 549 – 555. Marcantonio, F., R. F. Anderson, M. Stute, N. Kumar, P. Schlosser, and A. Mix (1996), Extraterrestrial 3He as a tracer of marine sediment transport and accumulation, Nature, 383, 705 – 707. Marcantonio, F., R. F. Anderson, S. Higgins, M. Stute, P. Schlosser, and P. Kubik (2001), Sediment focusing in the central equatorial Pacific Ocean, Paleoceanography, 16, 260 – 267. Marchal, O., R. Franc¸ois, T. F. Stocker, and F. Joos (2000), Ocean thermohaline circulation and sedimentary 231Pa/230Th ratio, Paleoceanography, 15, 625 – 641. Martin, P. A., D. W. Lea, Y. Rosenthal, N. J. Shackleton, M. Sarnthein, and T. Papenfuss (2002), Quaternary deep sea temperatures derived from benthic foraminiferal Mg/Ca, Earth Planet. Sci. Lett., 198, 193 – 209. Martinez, I., D. Rincon, Y. Yokoyama, and T. Barrows (2006), Foraminifera and coccolithophorid assemblage changes in the Panama Basin during the last deglaciation: Response to sea-surface productivity induced by a transient climate change, Palaeogeogr. Palaeoclimatol. Palaeoecol., 234, 114 – 126. Masson, D. G., R. B. Wynn, and B. J. Bett (2004), Sedimentary environment of the Faroe-Shetland and Faroe Bank channels, north-east Atlantic, and the use of bedforms as indicators of bottom current velocity in the deep ocean, Sedimentology, 51, 1207 – 1241. Mitchell, N. C., and M. W. Lyle (2005), Patchy deposits of Cenozoic pelagic sediments in the central Pacific, Geology, 33, 49 – 52. Mix, A., R. Tiedemann, P. Blum, and the Shipboard Scientific Party (2003), Proceedings of the Ocean Drilling Program, Initial Reports, vol. 202, Ocean Drill. Proj., College Station, Tex. (Available at http://www-odp.tamu.edu/ publications/202_IR/202ir.htm) Mollenhauer, G., T. I. Eglinton, N. Ohkouchi, R. R. Schneider, P. J. Mueller, P. M. Grootes, and J. Rullkoetter (2003), Asynchronous alkenone and foraminifera records from the Benguela upwelling system, Geochim. Cosmochim. Acta, 67, 2157 – 2171. Mollenhauer, G., M. Kienast, F. Lamy, H. Meggers, R. R. Schneider, J. M. Hays, and T. I. Eglinton (2005), An evaluation of 14C age relationships between co-occuring foraminifera, alkenones, and total organic carbon in continental margin sediments, Paleoceanography, 20, PA1016, doi:10.1029/2004PA001103. Moore, T. C., Jr., G. R. Heath, and R. O. Kowsmann (1973), Biogenic sediments of the Panama Basin, J. Geol., 81, 458– 472.  18 of 19  PA2213  Mortlock, R. A., and P. N. Froelich (1989), A simple method for the rapid determination of biogenic opal in pelagic marine sediments, Deep Sea Res., Part A, 36, 1415 – 1426. Mu¨ller, P. J., and E. Suess (1979), Productivity, sedimentation rate, and sedimentary organic matter in the oceans: I. Organic carbon preservation, Deep Sea Res., Part A, 26, 1347 – 1362. Murray, R. W., M. Leinen, D. W. Murray, A. C. Mix, and C. W. Knowlton (1995), Terrigenous Fe input and biogenic sedimentation in the glacial and interglacial equatorial Pacific Ocean, Global Biogeochem. Cycles, 9, 667 – 684. Paytan, A., M. Kastner, and F. P. Chavez (1996), Glacial to interglacial fluctuations in productivity in the equatorial Pacific as indicated by marine barite, Science, 274, 1355 – 1357. Paytan, A., M. Lyle, A. Mix, and Z. Chase (2004), Climatically driven changes in oceanic processes throughout the equatorial Pacific, Paleoceanography, 19, PA4017, doi:10.1029/ 2004PA001024. Pedersen, T. F. (1983), Increased productivity in the eastern equatorial Pacific during the Last Glacial Maximum (19000 to 14000 yr B. P.), Geology, 11, 16 – 19. Pedersen, T. F., M. Pickering, J. S. Vogel, J. N. Southon, and D. E. Nelson (1988), The response of benthic foraminifera to productivity cycles in the eastern equatorial Pacific: Faunal and geochemical constraints on glacial bottom water oxygen levels, Paleoceanography, 3, 157 – 168. Pichat, S., K. W. W. Sims, R. Franc¸ois, J. F. McManus, S. B. Leger, and F. Albarede (2004), Lower export production during glacial periods in the equatorial Pacific derived from (231Pa/230Th)xs,0 measurements in deepsea sediments, Paleoceanography, 19, PA4023, doi:10.1029/2003PA000994. Ru¨hlemann, C., M. Frank, W. Hale, A. Mangini, S. Mulitza, P. J. Mu¨ller, and G. Wefer (1996), Late Quaternary productivity changes in the western equatorial Atlantic: Evidence from 230 Th normalized carbonate and organic carbon accumulation rates, Mar. Geol., 135, 127 – 152. Sarnthein, M., and K. Winn (1988), Global variations of surface ocean productivity in low and mid latitudes: Influence on CO2 reservoirs of the deep ocean and atmosphere during the last 21,000 years, Paleoceanography, 3, 361 – 399. Scholten, J. C., R. Botz, H. Paetsch, and P. Stoffers (1994), 230Th(ex) flux into Norwegian-Greenland Sea sediments: Evidence for lateral sediment transport during the past 300,000 years, Earth Planet. Sci. Lett., 121, 111 – 124. Scholten, J. C., et al. (2005), Radionuclide fluxes in the Arabian Sea: The role of particle composition, Earth Planet. Sci. Lett., 230, 319 – 337. Shackleton, N. J., I. A. Hall, and E. Vincent (2000), Phase relationships between millennial-scale events 64,000 – 24,000 years ago, Paleoceanography, 15, 565 – 569. Shanmugam, G. (2003), Deep-marine tidal bottom currents and their reworked sands in modern and ancient submarine canyons, Mar. Pet. Geol., 20, 471 – 491. Skinner, L. C., and N. J. Shackleton (2005), An Atlantic lead over Pacific deep-water change across Termination I: Implications for the application of the marine isotope stage stratigraphy, Quat. Sci. Rev., 24, 571 – 580. Smith, W., and D. Sandwell (1997), Global sea  PA2213  KIENAST ET AL.: PARTICLE FLUX IN THE PANAMA BASIN  floor topography from satellite altimetry and ship depth soundings, Science, 277, 1957 – 1962. Snoeckx, H., and D. K. Rea (1994), Dry bulk density and CaCO3 relationships in upper Quaternary sediments of the eastern equatorial Pacific, Mar. Geol., 120, 327 – 333. Spero, H. J., and D. W. Lea (2002), The cause of carbon isotope minimum events at glacial terminations, Science, 296, 522 – 525. Spiess, F. N., B. T. Malfait, P. F. Lonsdale, and R. C. Tyce (1973), Shipboard cruise report on Leg 5 of Expedition SOUTHTOW, Ref. 73-1, 9 pp., Scripps Inst. of Oceanogr., Univ. of Calif., San Diego, La Jolla. Stuiver, M., and P. J. Reimer (1993), Extended 14 C data base and revised CALIB 3.0 14C age calibration program, Radiocarbon, 35, 215 – 230. (Available at http://radiocarbon.pa.qub. ac.uk/calib/) Suman, D. O., and M. P. Bacon (1989), Variations in Holocene sedimentation in the North American Basin determined from 230Th measurements, Deep Sea Res., Part A, 36, 869 – 878. Taylor, S. R., and S. M. McClennan (1985), The Continental Crust: Its Composition and Evolution, 312 pp., Blackwell Sci., Malden, Mass. Thomson, J., S. Nixon, C. P. Summerhayes, J. Schoenfeld, R. Zahn, and P. Grootes (1999), Implications for sedimentation changes on the Iberian margin over the last two glacial/ interglacial transitions from (230Thexcess)0 systematics, Earth Planet. Sci. Lett., 165, 255 – 270.  Toggweiler, J., K. Dixon, and W. Broecker (1991), The Peru upwelling and the ventilation of the South Pacific thermocline, J. Geophys. Res., 96, 20,467 – 20,497. Toggweiler, J. R., and S. Carson (1995), What are upwelling systems contributing to the ocean’s carbon and nutrient budgets?, in Upwelling in the Ocean, Modern Processes and Ancient Records, edited by C. P. Summerhayes et al., pp. 337 – 361, John Wiley, Hoboken, N. J. Turnewitsch, R., J.-L. Reyss, D. C. Chapman, J. Thomson, and R. S. Lampitt (2004), Evidence for a sedimentary fingerprint of an asymmetric flow field surrounding a short seamount, Earth Planet. Sci. Lett., 222, 1023 – 1036. Van Andel, T. H. (1973), Texture and dispersal of sediments in the Panama Basin, J. Geol., 81, 434 – 457. Van Andel, T. H., G. R. Heath, B. Malfait, D. F. Heinrichs, and J. I. Ewing (1971), Tectonics of the Panama Basin, eastern equatorial Pacific, Geol. Soc. Am. Bull., 82, 1489 – 1508. Weliky, K., E. Suess, C. A. Ungerer, P. J. Muller, and K. Fischer (1983), Problems with accurate carbon measurements in marine sediments and particulate matter in seawater: A new approach, Limnol. Oceanogr., 28, 1252 – 1259. Wollenburg, J. E., and W. Kuhnt (2000), The response of benthic foraminifera to carbon flux and primary production in the Arctic Ocean, Mar. Micropaleontol., 40, 189 – 231. Wunsch, C. (2003), Determining paleoceano-  19 of 19  PA2213  graphic circulations, with emphasis on the Last Glacial Maximum, Quat. Sci. Rev., 22, 371 – 385. Yang, Y.-L., H. Elderfield, and M. Ivanovich (1990), Glacial to Holocene changes in carbonate and clay sedimentation in the equatorial Pacific Ocean estimated from Thorium 230 profiles, Paleoceanography, 5, 789 – 809. Yang, Y.-L., H. Elderfield, T. Pedersen, and M. Ivanovich (1995), Geochemical record of the Panama basin during the Last Glacial Maximum carbon event shows that the glacial ocean was not suboxic, Geology, 23, 1115 – 1118. Yu, E.-F., R. Franc¸ois, M. P. Bacon, and A. P. Fleer (2001), Fluxes of 230Th and 231Pa to the deep sea: Implications for the interpretation of excess 230Th and 231Pa/230Th profiles in sediments, Earth Planet. Sci. Lett., 191, 219 – 230.  ÀÀÀÀÀÀÀÀÀÀÀÀÀÀÀÀÀÀÀÀÀÀÀÀÀ  S. E. Calvert and R. Franc¸ois, Earth and Ocean Sciences, University of British Columbia, Vancouver, BC, V6T 1Z4 Canada. M. Kienast and S. S. Kienast, Department of Oceanography, Dalhousie University, 1355 Oxford Street, Halifax, Nova Scotia, B3H 4J1 Canada. (stephanie.kienast@dal.ca) A. C. Mix, College of Oceanic and Atmospheric Sciences, COAS Administration Building 104, Oregon State University, Corvallis, OR 97331-5503, USA.  


Citation Scheme:


Citations by CSL (citeproc-js)

Usage Statistics



Customize your widget with the following options, then copy and paste the code below into the HTML of your page to embed this item in your website.
                            <div id="ubcOpenCollectionsWidgetDisplay">
                            <script id="ubcOpenCollectionsWidget"
                            async >
IIIF logo Our image viewer uses the IIIF 2.0 standard. To load this item in other compatible viewers, use this url:


Related Items