1
Biogeochemical and hydrographic controls on chromophoric dissolved
organic matter distribution in the Pacific Ocean
Chantal M. Swana, David A. Siegela b, Norman B. Nelsona, Craig A. Carlsonc, Elora
Nasira
a
Institute for Computational Earth System Science, University of California, Santa
Barbara, USA
b
Department of Geography, University of California, Santa Barbara, USA
c
Department of Ecology, Evolution and Marine Biology, University of California, Santa
Barbara, USA
0
Corresponding author: Chantal M. Swan, email: swan@icess.ucsb.edu, Ph: +1 805 4501564, Fx: +1 805 893-2578
Abstract
Recent in situ observations of chromophoric dissolved organic material (CDOM) in the Pacific Ocean
reveal the biogeochemical controls on CDOM and indicate predictive potential for open-ocean CDOM in
diagnosing particulate organic matter (POM) remineralization rates within ocean basins. Relationships
between CDOM and concentrations of dissolved oxygen, nutrients and inorganic carbon in the
subthermocline waters of the Pacific reflect the relative influences of water mass ventilation and water
column oxidative remineralization. Apparent in situ oxygen utilization (AOU) accounts for 86% and 61%
of variance in CDOM abundance, respectively, in Antarctic Intermediate Water and North Pacific
Intermediate Water. In the deep waters of the Pacific below the zone of remineralization, AOU explains
26% of CDOM variability. The AOU-CDOM relationship results from competing biogeochemical and
advective processes within the ocean interior. Dissolved organic carbon (DOC) is not statistically linked to
the CDOM or AOU distributions, indicating that the majority of CDOM production occurs during the
remineralization of sinking POM and thus potentially provides key information about carbon export. Once
formed in the ocean interior, CDOM is relatively stable until it reaches the surface ocean where it is
destroyed by solar bleaching. Susceptibility to bleaching confers an additional tracer-like quality for
CDOM in water masses with active convection, such as mode waters that appear as subsurface CDOM
minima. In the surface ocean, atypically low CDOM abundance highlights a region of unusually extreme
oligotrophy: the subtropical South Pacific gyre. For these hyper-oligotrophic waters, the present CDOM
observations are consistent with analysis of in situ radiometric observations of light attenuation and
reflectance, demonstrating the accuracy of the CDOM spectrophotometric observations. Overall, we
illustrate how CDOM abundance in the ocean interior can potentially diagnose rates of thermohaline
overturning as they affect regional biogeochemistry and export. We further show that relative surface ocean
CDOM abundances are driven in large part by processes occurring in the deep layers of the ocean. This is
particularly significant for the interpretation of the global surface distribution of CDOM using satellite
remote sensing.
Keywords: CDOM, AOU, Pacific, water masses, hydrography, bio-optical
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1. Introduction
Chromophoric dissolved organic matter (CDOM) is the light-absorbing fraction
of the total oceanic dissolved organic matter pool, and as such is quantifiable in the
surface ocean from satellite ocean color data (Morel, 1988; Siegel et al. 2002). CDOM is
characterized by an absorption spectrum that is high across ultraviolet (UV) wavelengths
and decreases monotonically toward longer wavelengths (Bricaud et al., 1981). CDOM
abundance is typically reported as absorption coefficient (m-1) at a single wavelength
(325nm) as in the present study. In the open ocean, CDOM is locally produced as a
byproduct of microbial transformations of organic material, is destroyed by
photobleaching in the surface ocean, and is distributed throughout the deep ocean in a
way that reflects water mass subduction, residence time and seasonal convective events
that redistribute light-exposed surface CDOM (Vodacek et al., 1997; Nelson et al., 2004;
2007; Morel et al., 2007a; Yamashita and Tanoue 2008). The hydrographic and
biogeochemical tracer-like characteristics of CDOM in the Pacific Ocean are the focus of
the present paper.
The study of CDOM is timely as satellite estimates reveal that CDOM accounts
for nearly one-half of total non-water light absorption in the global surface ocean at
440nm, near the maximum of spectral light absorption by chlorophyll in phytoplankton.
Regional variation in CDOM is not yet accounted for in current satellite chlorophyllbased estimations of global phytoplankton biomass, and this oversight impacts global
biogeochemical predictions (Siegel et al., 2005b). Further, CDOM absorbs over 90% of
total UV radiation entering the euphotic layer, thereby serving a photoprotective role in
biological processes (Zepp, 2002). UV photochemistry of CDOM also contributes to the
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sea-surface flux of climate-relevant trace gases (DMS, CO, CO2, and COS) (Cutter et al.,
2004; Johannessen and Miller, 2001; Stubbins et al., 2006; Toole et al., 2006; Zafiriou et
al. 2008), and to the photochemical alteration of DOM bioavailability in the surface
ocean (Miller, W.L. et al., 2002).
The synoptic global satellite view of CDOM in the surface ocean suggests that
CDOM may be a tracer of upper ocean circulation (Siegel et al., 2005a), as it provides
information about the relative contributions of photobleaching, biological production of
CDOM, and vertical mixing inputs of CDOM on various time scales of ocean renewal
processes (Nelson et al., 2007; Siegel et al., 2005a). The 8-yr. mean satellite-based
surface CDOM distribution calculated using a semi-analytical inverse method based on
SeaWiFS ocean color data is shown in Fig. 1 for the Pacific sector of the global ocean
(GSM algorithm: Siegel et al., 2005a; 2005b). Basin-scale oceanographic features such as
the subtropical gyres and coastal and equatorial upwelling regions are noticeably
reflected by the surface CDOM distribution (Fig. 1). Surface abundances of CDOM
increase poleward in each hemisphere, consistent with the degree of photobleaching
expected with latitude given meridional gradients in solar irradiance and mixed-layer
depth. There is a hemispheric asymmetry in surface CDOM abundance, pronounced at
high latitudes and weighted toward the northern hemisphere (Siegel et al., 2002; 2005a)
that is examined within the present study.
Characteristics of CDOM in the ocean interior relating to the meridional
overturning circulation have recently been described for the North Atlantic basin (Nelson
et al., 2007; Nelson et al. in review). In the highly ventilated North Atlantic, abundance
and spectral characteristcs of CDOM within deep waters are traceable to their formation
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region, suggesting that satellite surface CDOM data could be applied to quantifying water
mass advection. Several studies that relate the fluorescent fraction of dissolved organic
matter in the Pacific to biogeochemical properties implicate CDOM as a potential means
for assessing rates of organic carbon turnover by way of its relationship to
remineralization indices such as apparent oxygen utilization (AOU) (Chen and Bada,
1992; Coble, 1996; Coble, 2008; Hayase and Shinozuka, 1995; Yamashita et al., 2006;
Yamashita and Tanoue, 2008). These results, expanded upon within this paper, indicate
the scope of potential applications of CDOM measurements to marine geochemistry.
The deep ocean distributions of CDOM have only begun to be mapped in situ
over the global ocean during the last several years (Nelson et al., 2007; Nelson et al. in
review). Absorption by open-ocean CDOM (remote from terrestrial influence) was
previously too low for instrument detection and thus regarded as relatively insignificant
in marine systems (Nelson and Siegel, 2002). The advent of long-path spectrometers have
changed this (Miller, R.L., et al., 2002). Rapid, routine detection of CDOM in the open
ocean with such instruments has provided insights into the stability of CDOM in the deep
ocean, and the capacity of CDOM to function as a biogeochemical proxy, even as its
cycling mechanisms appear to occur independently from those of dissolved organic
carbon (DOC).
Systematic observations of CDOM in the open-ocean Pacific were conducted
along with hydrographic measurements under the auspices of the U.S. CO2/CLIVAR
Repeat Hydrography Program. The present study incorporates measurements taken along
CLIVAR cruise tracks P16 and P2, shown as black lines superimposed on Fig. 1. Here
we present the full-depth, trans-Pacific profiles of CDOM in relation to the
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biogeochemical and water mass structure of the basin, discuss an intriguing coupling
between the surface ocean and subthermocline waters of the North Pacific that relate to
the observed hemispheric asymmetry in satellite-derived CDOM, and explore the tracerlike qualities of CDOM in the open ocean. Further, by examining the CDOM distribution
in the context of the global thermohaline circulation and ventilation processes of the
Pacific, we build an argument for the potential use of CDOM in diagnosing relative rates
of water mass advection versus carbon flux and remineralization among ocean basins.
FIGURE 1
2. Methods
2.1 Hydrographic measurements
The CLIVAR P2 field campaign was conducted during July – Aug. 2004 in the
North Pacific along 30N latitude, spanning longitudes 133E (Yokohama, Japan) to
117W (San Diego, California). The CLIVAR P16 transect was conducted in two field
campaigns, P16S (South Pacific, Jan. – Feb. 2005) and P16N (North Pacific, Feb. – Mar.
2006), achieving a latitudinal coverage from 71S (Antarctic circle) to 55N (Kodiak
Island, Alaska) along 150W longitude in the Pacific (Fig. 1). Station spacing was
approximately 55 km along the P2 and P16S transects, and approximately 110 km on the
P16N transect.
The P2 and P16 hydrographic data considered herein: dissolved inorganic carbon
(DIC), dissolved oxygen, dissolved nutrients (NO3, PO4), salinity and potential
temperature, were collected during full-depth 36-bottle CTD/rosette casts as part of the
core measurement suite of the CLIVAR Program following standard WOCE protocols
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(http://ushydro.ucsd.edu/). Dissolved organic carbon (DOC) data included in the present
study were collected from identical bottle samples as CDOM during the campaigns. DOC
samples were drawn from the rosette into HDPE bottles via in-line gravity filtration
through precombusted GF/F filters and stored frozen for later on-shore analysis. DOC
analysis was conducted via high temperature catalytic combustion on a modified
Shimadzu TOC-V analyzer according to established protocol (Carlson et al., 2004;
Farmer and Hansell, 2007). Apparent oxygen utilization (AOU), neutral density (n) and
potential vorticity variables used in the present analysis were derived from P2 and P16
hydrographic data using Ocean Data View version 3.2.2 software (Schlitzer, 2002).
2.2 CDOM absorption measurements
Our ability to conduct a relatively high-resolution survey of the oceanic CDOM
distribution has been possible due to the application of liquid capillary waveguides in
absorption spectroscopy (Miller, R.L. et al., 2002; Nelson et al., 2007). We utilized an
UltraPathTM single-beam instrument (World Precision Instruments, Sarasota, FL), which
consists of a UV-VIS light source (250-800nm), 194.3cm liquid waveguide cell and
TIDAS II spectrometer module. This system is capable of detecting a dynamic range of
light absorption values (especially low absorption values typical of the oligotrophic
ocean) at a relatively fast sample injection and analysis rate (< 2 min.). Instrument drift is
minor, yet routinely corrected for through a pure water reference scan immediately before
and after sample analyses. A limitation when using the UltraPathTM is that salinity of
seawater confers a higher index of refraction than does reference pure water (Byrne and
Kaltenbacher, 2001; Miller, R.L. et al., 2002). A higher index of refraction increases
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photon transmission through the waveguide cell to beyond that of pure water, which
creates a negative offset in the raw absorption signal. We have established an empirical
procedure to correct for this salinity effect following Nelson et al. (2007). Briefly, we
compiled absorption spectra on discrete dilutions of artificial seawater media (ASW;
Goldman and McCarthy, 1978) over a salinity range encompassing oceanic values (0 –
40ppt), and interpolation was used to determine corrections for intermediary salinity
levels (Nelson et al., 2007). The resultant ASW absorption spectrum corresponding to the
sample salinity is subtracted from the sample’s raw absorption spectrum as measured by
the UltraPathTM. Due to gradual, slight solarization of the waveguide and fiber optics of
the instrument during repeated usage, periodic recalibration using ASW media is
performed as described above to ensure the most accurate refractive index correction. All
data included in this study therefore reflect up-to-date instrumental correction factors.
CDOM absorption data were acquired from water samples that were drawn from
Niskin bottles into pre-combusted amber glass vials with Teflon-lined caps. Each sample
was vacuum-filtered through a 25mm 0.2m Nuclepore polycarbonate membrane preconditioned with ultrapure water on an all-glass filtration apparatus prior to spectroscopic
analysis. Samples from the P16 transect were collected, filtered and analyzed shipboard
within the same day. CDOM samples were collected and filtered shipboard in the same
manner along P2, but stored and shipped at 4C for spectroscopic analysis on-shore
(within 1 month). No appreciable change in CDOM absorption has been observed in prefiltered open-ocean samples that have been refrigerated at 4C in the dark for months to
up to one year (Swan et al. in review). All samples were equilibrated to room temperature
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prior to spectroscopic analysis to eliminate temperature effects, and reference water was
supplied by a Barnstead Nanopure Diamond UV/UF water purification system.
The estimated accuracy of our methods for quantifying CDOM absorption at
325nm is 0.0008 m-1 (assessed as the sum of RMS uncertainty of calibrations) and the
estimated precision at 325nm is 0.005 m-1 (assessed through the RMS difference between
replicate samples). Precision works out to be less than 4% of the mean observed CDOM
signal at 325 nm. CDOM absorption values measured at overlapping stations in the South
Pacific on the P16S and P16N campaigns were identical within this precision. This is
further support for methodological repeatability, particularly as the two campaigns
occurred more than one year apart.
The approximately exponential shape of the open-ocean CDOM spectrum
permitted estimation of the spectral slope parameter via a non-linear curve fit to the
wavelength range 300 – 650nm as described in Nelson et al. (2007). This parameter,
referred to as Snlf (nm-1) herein, is used to evaluate relative compositional changes in
CDOM, as high Snlf values are potentially indicative of long surface residence time and
photobleaching, while low Snlf values are potentially indicative of newly-formed CDOM
from microbial processes in the open ocean (Twardowski and Donaghay, 2002; Kitidis et
al., 2006). Confidence in the Snlf fitting method decreases when CDOM absorption levels
at 325nm fall below instrumental detection limits (0.005 m-1) resulting in geographic bias
in values of Snlf for some regions (e.g., the South Pacific gyre). For this reason, the
relative patterns in Snlf are described qualitatively as they support hypotheses related to
controls on the overall CDOM distribution in the present study.
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2.3 Radiometric and particle absorption measurements
Measurements of bulk in-water apparent optical properties (AOPs) in the euphotic
zone, specifically the spectral irradiance reflectance, R(), and the diffuse attenuation
coefficient of downwelling spectral irradiance, Kd() (m-1), were made during the P16
campaigns. A free-falling MicroPro IITM multispectral radiometer and mast-mounted
surface reference radiometer (Satlantic, Inc., Halifax, Nova Scotia) were used in tandem
to measure the underwater light field in the approximate upper 150m of the ocean. Daily
measurements of the spectral downwelling irradiance. Ed(),and spectral upwelling
radiance, Lu() at 11 wavelengths within 324 – 683nm were acquired around local solar
noon. Calculation of AOPs (and subsequently IOPs) from our direct measurements of Ed
and Lu are discussed in section 3.2.2.
In addition to the full-depth hydrographic CDOM measurements and radiometric
profiles conducted daily, we compiled particulate absorption (ap) spectra for surface
waters along P16. An uncontaminated ship surface-water intake was used to collect 4L
bottle samples from approximately 5m depth. Samples were immediately vacuum-filtered
onto 25mm Whatman GF/F filters, then flash frozen and stored in liquid nitrogen for later
on-shore spectroscopic analysis using the quantitative filter technique (Nelson et al.,
1998; Mitchell, 1990). The ap spectra (300 – 750nm) of the samples were analyzed using
a UV-2401PC spectrophotometer (Shimadzu Scientific, Inc., Columbia, MD) equipped
with an integrating sphere attachment and a reference GF/F filter saturated with ultrapure
water. Pathlength amplification effects were determined as described in Nelson et al.
(1998).
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Our determinations of particle absorption and of optical properties from
radiometric data required knowledge of chlorophyll-a (Chl) (mg/m3) concentration.
Filtered samples from the upper 250m of the water column along the P16 transect were
analyzed for Chl concentration following standard fluorometric protocol (BATS Methods
Manual: http://bats.bios.edu/methods/chapter14.pdf).
2.4 Water mass definitions
This study espouses several of the major Pacific water mass layer definitions from
the literature in order to evaluate the CDOM distribution in relation to the vertical
structure and circulation dynamics of the Pacific Basin. The abbreviations, approximate
depths and ranges of neutral density anomaly (n, kg/m3) for these water mass layers are
summarized in Table 1. (As the vertical distribution of a water mass may change
dynamically with latitude and longitude, depth ranges in Table 1 should be interpreted
only as a rough guideline. Neutral density provides the more effective means of tracing
water mass distribution.) In brief, North Pacific Subtropical Mode Water (STMW) is
defined approximately by n layer 25.0 – 25.8 kg/m3 (Sonnerup et al., 1999). Subantarctic
Mode Water (SAMW) in the South Pacific is roughly delineated by a n pycnostad layer
of 27.0 – 27.2 kg/m3 (Sarmiento et al., 2003), and is formed in close association with
Antarctic Intermediate Water (AAIW) defined by n layer 27.0 – 27.8 kg/m3 (Sloyan and
Rintoul, 2001). We define North Pacific Deep Water (NPDW) by n layer 27.8 – 28.1
kg/m3, and group together Circumpolar Deep Water and Antarctic Bottom Water (CDW
+ AABW = n > 28.1 kg/m3) when referring to the abyssal waters of the Pacific
originating in the Southern Ocean (Ganachaud, 2003). Finally, the North Pacific
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Intermediate Water mass (NPIW), which was collectively well-sampled by P2 and P16
campaigns, is defined by n layer 26.0 – 27.4 kg/m3 (Ganachaud, 2003). As NPIW is
predominantly a zonally-oriented water mass in its formation (Miyao and Ishikawa, 2003;
Talley, 1997), the P2 hydrographic dataset was targeted at characterizing NPIW.
Therefore, relationships among hydrographic properties within NPIW are evaluated from
both zonal (P2) and meridional (P16) perspectives in the present study.
Water mass
Abbreviation
Depth range (m)
ReferenceR
n range (kg/m3)
North Pacific Intermediate Water
NPIW
26.0 – 27.4
400 – 900
Ganachaud 2003
North Pacific Subtropical Mode Water
STMW
~25.0 – 25.8
100 – 400
Sonnerup et al. 1999
North Pacific Deep Water
NPDW
27.8 – 28.1
1500 – 3000
Ganachaud 2003
Subantarctic Mode Water
SAMW
~27.0 - 27.2
400 – 700
Sarmiento et al. 2003
Antarctic Intermediate Water
AAIW
27.0 – 27.8
500 – 1500
Sloyan and Rintoul 2001
Circumpolar Deep Water and
CDW + AABW > 28.1
> 3000
Ganachaud 2003
Antarctic Bottom Water
Table 1. Water mass layer definitions relevant to this study and adopted from the literature as noted. Depth ranges are approximate as
water mass distribution varies dynamically with latitude and longitude during circulation.
3. Results and Discussion
3.1 CDOM distribution in the major water masses of the Pacific
The full ocean-depth distributions of CDOM absorption at 325nm, spectral slope
(Snlf), apparent oxygen utilization (AOU) and dissolved organic carbon (DOC) in the
Pacific as measured along P16 (150W) and P2 (30N) are shown in Figs. 2a-d and 3a-d,
respectively. The contour lines on each panel represent neutral density surfaces, with the
exception of Figs. 2c and 3c where contour lines of salinity overlay the respective P16
and P2 AOU profiles to outline the injection of intermediate waters, specifically the low
salinity tongue (Fig. 3c) that characterizes the distribution of NPIW (You et al., 2003).
The CDOM distribution across the meridional Pacific transect at 150W (Fig. 2a)
reveals several features of the general upper-ocean circulation that are consistent with
what is known about the light-driven component of CDOM cycling in the open ocean
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(Siegel et al., 2002). The subtropical gyres, centered at 30 latitude in each hemisphere,
appear as pools of relative CDOM minima (Fig. 2a). This is explained by extensive
photobleaching of CDOM due to the long residence time of surface waters in the
subtropical gyres and high solar insolation of the subtropics, which is corroborated by the
relatively high Snlf values observed within the gyres (Fig. 2b). The CDOM-depleted
surface waters of the subtropical gyres are gradually downwelled via Ekman pumping,
and this convergent circulation impedes input of elevated CDOM waters from below. The
South Pacific subtropical gyre has an even lower stock of CDOM and higher overall Snlf
values than its northern hemisphere counterpart (Fig. 2a). This is in part due to
outcropping isopycnals extending the “bowl” of the subtropical gyre in the South Pacific
to greater depths than in the North Pacific (Feely et al. 2004), and in part due to the
deeper penetration of intermediate waters in the South Pacific than in the North Pacific
that transport low-CDOM surface waters from formation regions to intermediate depths
(Sabine et al., 2004). An additional factor is the extreme oligotrophy that typifies the
South Pacific subtropical system, where chlorophyll values and water clarity are among
the lowest documented in oceanic waters (Morel et al., 2007a; 2007b; Tedetti et al.,
2007). We measured exceptionally low in situ CDOM absorption (ranging from near-zero
to 0.05 m-1 at 325nm) and the highest slope value (Snlf = 0.03 nm-1) in the South Pacific,
which we examine in greater detail in section 3.2.2.
FIGURE 2
The equatorial upwelling associated with the tropical current system in the Pacific
is also reflected in the CDOM distribution (Fig. 2a). The pattern of CDOM between 10S
and 10N follows the shoaling of neutral density surfaces at the equator, where there are
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vertical inputs of elevated subsurface CDOM into the surface waters (Simeon et al.
2003). The Snlf values in the equatorial distribution (Fig. 2b) are similar to those found in
the intermediate and deep water north of the equator, suggesting that the CDOM
upwelled at the equator may have origins associated with North Pacific Intermediate
Water (NPIW) southward flow. Renewal of upwelled waters at the equator is frequent
enough for elevated levels of CDOM to be sustained, as is confirmed by satellite
observations of CDOM (Fig. 1) and the relatively lower Snlf values that are sustained in
surface waters of the equatorial Pacific (Fig. 2b).
The degree of surface photobleaching of CDOM, subsurface production of
CDOM, and the residence time of a water mass are concomitant factors in setting the
magnitude and distribution of CDOM absorption and spectral characteristics in the ocean
interior. This is revealed when examining CDOM in the subthermocline water masses of
the Pacific. There is a well-defined signature of Antarctic Intermediate Water (AAIW)
and Subantarctic Mode Water (SAMW) formation in the meridional Pacific CDOM
distribution at the Polar Frontal Zone near 55 - 60S between approximately 800 –
1000m depth (Fig. 2a). AAIW and SAMW inject photobleached surface waters that are
low in CDOM and have higher Snlf values relative to the deeper waters beneath the front
(CDW and AABW) that form further south in the more productive Ross Sea (Sloyan and
Rintoul, 2001). SAMW is the precursor of AAIW, formed in all sectors of the Southern
Ocean north of the Polar Frontal zone through convection during austral winter. As such,
SAMW entrains subantarctic surface waters to approximately 600m depth (Sarmiento et
al., 2003; Schneider and Bravo, 2006). Its signal in the CDOM distribution is discernible
as a minimum relative to AAIW and the subsurface layer above it (Fig. 2a). There is also
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a relative maximum in the Snlf distribution that corresponds to the core of SAMW at
approximately 500m at 40S (Fig. 2b), which further indicates entrainment of
photobleached surface waters. This approximate location of SAMW as it corresponds to
the relative CDOM minimum and Snlf maximum is confirmed by a relative potential
vorticity minimum (not displayed) (Sloyan and Kamenkovich, 2007).
A characteristic feature in meridional sections of the North Pacific is the uniform
distribution of hydrographic properties below 3000m (Pickard and Emery, 1990). This is
manifest in the CDOM and Snlf distributions in the abyssal waters (>3000m) as well.
CDOM absorption averages 0.12 m-1 in abyssal waters of Pacific (Fig. 2a), and only a
very slight gradient in CDOM and Snlf is observed between the southern and northern end
of the transect as compared to the meridional CDOM gradient in overlying intermediate
and deep waters. The water column above 3000m reveals hydrographic patterns that are
consistent with the degree of oxidative remineralization as water masses age along the
global oceanic conveyor belt (Feely et al., 2004). Indeed the most prominent feature of
the meridional Pacific CDOM distribution is the pronounced gradient of increasing
CDOM in subthermocline waters (Fig. 2a) and decreasing spectral slope (Fig. 2b) from
the South Pacific into the North Pacific (Fig. 2a), and the strong association of the
CDOM gradient with that of AOU (Fig. 2c). This relationship suggests that
biogeochemical processes are the dominant driver of CDOM dynamics in the Pacific
basin, and is consistent with the current paradigm of CDOM formation through microbial
organic matter remineralization in the open-ocean, which is supported by the gradient of
decreasing Snlf (Nelson et al., 2004; 2007; Yamashita and Tanoue, 2008). The DOC
distribution does not share similar correspondence with CDOM and AOU in the Pacific
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(Fig. 2d). This implies several things: 1) in the Pacific, DOC cycling processes are
decoupled from those of CDOM, as was shown for the Atlantic (Nelson et al., 2007), 2)
the contribution of DOC remineralization to AOU (and likely CDOM) is small (<10% at
depth; Aristegui et al. 2002), and 3) organic matter decomposition resulting in CDOM
formation on multi-decadal time scales likely involves the sinking flux of particulate
organic carbon. This seems plausible given that the region of the water column where
AOU and CDOM appear to have the strongest relation (Figs. 2a-c) is the remineralization
zone over which organic particles are solubilized and decomposed (Azam, 1998; Feely et
al., 2004).
The distribution of CDOM and Snlf in the upper 1000m of the zonal Pacific along
30N (P2 section) is shown in Figs. 3a and 3b, respectively. CDOM increases with depth
across the basin, with a slight offset of elevated concentrations at shallower depths near
the North American coast that is presumably due to upwelling (Fig. 3a). Snlf values are
highest in surface waters and in the western section of the P2 transect, and lowest
between 500 – 1000m (Fig. 3b). The gradient of Snlf values with depth along P2 is again
indicative of CDOM generation through organic matter decomposition over the depth
horizon of remineralization (1000m) in the North Pacific.
Between approximately 160°E and 160°W in the depth range 500 – 1000m, there
is very high CDOM abundance (0.3 m-1) and associated low Snlf values (0.012 nm-1)
relative to the waters to the east or west of this area along P2 (Fig. 3a) that indicate a
localized source of CDOM. Only the upper 1000m of properties along P2 are shown here
as CDOM absorption data below 1000m is very sparse; however, these few CDOM
observations suggest that this feature of high (~0.3 m-1) CDOM absorption (observable at
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longer observation wavelengths as well) extends to roughly 2500m. This feature is
uncoupled with the AOU distribution, implying a non-microbial source of CDOM. It is
possible that the Hawaiian Ridge – Emperor Seamount Chain is influencing the CDOM
distribution as observed along P2 as its longitudinal range intersects with the area of
observed elevated CDOM abundance (Clague and Dalrymple, 1987). Boyle et al. (2005),
for example, attributed Fe enrichment observed within intermediate depths around
Station ALOHA to the nearby Loihi Seamount submarine hydrothermal vents (18°55’N,
156°15’W) that erupt between 1000 – 1500m depth. The most active areas within the
Emporer seamount system are located just a few degrees north of the P2 transect at
approximately 170E. Hydrothermal activity can cause resuspension of organic matter
into the water column from surrounding sedimentary material at the vent site (Levesque
et al., 2005). Either through direct transport, or microbial processing of this material in
the water column, proximate hydrothermal activity in the subtropical North Pacific could
be the source of elevated CDOM observed along P2; however, as noted it is difficult to
validate this hypothesis with the current zonal dataset.
In the western portion of the P2 section, between 140°E – 175°E, a minimum in
the subsurface CDOM distribution is found between 100 – 500m that corresponds to the
location of North Pacific Subtropical Mode Water (STMW) (Fig. 3a). STMW is a
pycnostad (much like SAMW, but of smaller scale) occurring between approximately
100 – 400m in the water column, that is formed as a result of successive convective
overturn events (Oka and Suga, 2003; Sonnerup et al., 1999). Subtropical surface waters
that are low in CDOM and have high Snlf values due to subtropical irradiation are
entrained between the seasonal and the permanent thermocline during STMW formation.
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The Snlf distribution along the P2 transect is consistent this formation mechanism as there
is a bolus of high Snlf (0.29 nm-1) observed at ca. 300m at 150°E within STMW (Fig. 3b).
Nelson et al. (2007) show a similar pattern for CDOM in the North Atlantic subtropical
mode water. In both the subtropical North Atlantic and Pacific basins, the CDOM
minima associated with the subtropical mode waters are corroborated by local potential
vorticity minima (not displayed) (Nelson et al., 2007; Talley, 1988).
FIGURE 3
The overlain salinity contours on the P2 AOU distribution in Fig. 3c illustrate the
low salinity tongue (~34.25 on the practical salinity scale) that characterizes North
Pacific Intermediate Water (NPIW) (You et al., 2003). NPIW is the main ventilation
pathway of the North Pacific; however, the water mass does not have a clear signature in
the CDOM distribution as the other ventilated water masses in the Pacific do (SAMW,
AAIW, STMW) (Fig. 3a). This is likely caused by the unique formation mechanism of
NPIW. NPIW has a surface source water region in the far western Pacific, but
predominantly acquires its geochemical properties through slow interior mixing at
intermediate depths with southward-flowing, relatively fresh Oyashio current waters and
northward-flowing warmer, more saline Kuroshio current waters (Miyao and Ishikawa,
2003; Talley, 1997; You et al., 2003). This admixture is transported along the northern
rim of the subtropical gyre and spreads laterally eastward once under the weaker
influence of the eastern boundary current. The relatively unventilated nature of NPIW is
why CFC invasion has failed to distinctly trace its evolution (Sonnerup et al., 1999).
There is a vertical gradient in CDOM and Snlf within NPIW (Fig. 3a-b) that roughly
coincides with that of AOU (Fig. 3c), albeit not as robustly as observed across the
18
meridional Pacific transect. NPIW therefore represents a case where ventilation and
advection compete with remineralization processes in establishing the CDOM patterns
observed. A quantitative evaluation of these statements and biogeochemical relationships
with CDOM are given in section 3.3.
It should be noted that absorption by CDOM in the visible region of the spectrum
(e.g., 443nm) closely followed the distributional patterns of CDOM absorption at 325nm
as discussed throughout the paper and presented in Figs. 2 and 3. However, as absorption
in the visible is much less than CDOM absorption in the UV, hydrographic features and
gradients are not as easily resolvable in a depth contour of absorption at 443nm, and are
therefore not displayed herein.
3.2 Patterns of CDOM in surface waters of the Pacific
3.2.1 Hemispheric asymmetry in satellite-derived and spectroscopy-derived CDOM
The distribution of CDOM in the surface layer (<15m) of the Pacific along
150°W, as both measured in situ during P16 and from satellite retrievals using the GSM
algorithm (Siegel et al., 2005a), are plotted together in Fig. 4. Absorption coefficient
(m-1) at 325nm is represented. As absorption coefficients of CDOM at 443nm were below
practical detection limits in many surface samples along P16 and P2, values of CDOM at
325nm (Fig. 4) were extrapolated from the GSM algorithm retrieval of CDOM at 443nm
using the GSM CDOM spectral slope value of 0.0206 m-1, a mean value for the open
ocean as determined from global field data (Siegel et al., 2002). The match-up in the
Pacific between spectroscopy-derived CDOM and satellite-derived (GSM) CDOM values
(Fig. 4) is good (r2 = 0.72, n = 59, CDOMGSM = 1.743*CDOMin situ – 0.007), as well as
19
consistent with previous assessments of GSM performance using global datasets (Siegel
et al., 2005a, 2005b; Nelson et al., 2007). The slope of this relationship differs from
unity, which is likely due to both the use of the global average spectral slope to calculate
absorption at 325nm as determined by satellite, and due to the small (< 15%) contribution
from nonphytoplankton particulate matter in satellite-derived estimates of CDOM
absorption coefficient (Siegel et al. 2002; Nelson et al., 1998).
FIGURE 4
The hemispheric asymmetry observed in the satellite-derived global surface
distribution of CDOM (Siegel et al., 2005a) is corroborated by our field observations.
Mean CDOM absorption values across surface waters of the North Pacific exceeds that of
the South Pacific by nearly 0.03 m-1 (Fig. 4). One hypothesis for this observed
phenomenon is that solar irradiance dose in the surface ocean (a function of mixed layer
depth (MLD) and solar irradiance climatology) is higher in the southern than northern
hemisphere, leading to greater surface bleaching on multi-annual time scales. However,
climatologies of MLD in the respective hemispheres revealed that MLD is, on average,
shallower in the North Pacific during summer months (de Boyer Montégut et al., 2004).
This would lead to higher integrated irradiance dose over the North Pacific than South
Pacific, particularly given that solar fluxes are comparable among hemispheres in the
Pacific (Wenying Su, pers. comm., 2008). These considerations notwithstanding, it could
be argued that CDOM in the South Pacific is more susceptible to solar bleaching than it is
in the North Pacific; however, experimental determinations of apparent quantum yield for
photobleaching of CDOM in the open-ocean Pacific do not support this argument (Swan
20
et al., in review.) nor the hypothesis that hemispheric asymmetry in CDOM is lightdriven.
Given the elevated levels of CDOM observed in the intermediate and deep waters
of the North Pacific (Fig. 2a), and the slow thermohaline upwelling that occurs in the
subarctic Pacific (Ganachaud, 2003; Qiu and Huang, 1995), we pose that the
bathypelagic supply of CDOM in the North Pacific exerts control on the surface
expression of CDOM in this region on multi-annual time scales. Deep water inputs of
CDOM to the surface layer may cause the observed hemispheric asymmetry in CDOM in
the Pacific observed from the satellite data. Bulk thermohaline upwelling also occurs in
the Pacific sector of the Southern Ocean; however, deep waters of the Antarctic are
younger and lower in CDOM abundance than North Pacific deep waters. Therefore,
surface inputs of CDOM by deep water upwelling in the Southern Ocean do not balance
the high latitude surface distribution of CDOM (Fig. 1) in the northern hemisphere.
Compared to other ocean basins, the subarctic Pacific also exhibits higher surface
nutrient availability (Schlitzer, 2004), which further emphasizes the influence of deep
water outcropping on surface water concentrations in the North Pacific.
Terrestrial DOM is also not likely to be driving the observed hemispheric
asymmetry in surface layer CDOM. Riverine inputs have been shown to contribute only a
small fraction (0.7–2.4%) of the total DOM in the ocean, and the oceanic residence time
(20–130yr) of terrigenous DOM is significantly less than that of marine DOM (Opsahl
and Benner, 1997). Such findings were based on assessments of oceanic concentrations
of lignins and humics (found only in terrigenous plant material) that are chromophoric in
nature (Nelson and Siegel, 2002). Any potential CDOM contributions from major
21
Alaskan river inputs into the subarctic region sampled on P16 would likely appear as
small-scale lenses of very elevated surface CDOM in the neritic zone of the northern P16
transect, much like that of the Orinoco River plume observed in meridional sections of
the North Atlantic (Nelson et al., 2007). As such a feature is not visible in Fig. 2a, it is
unlikely that terrigenous inputs to the Gulf of Alaska drive the observed hemispheric
asymmetry in the open-ocean CDOM distribution (see also Yamashita and Tanoue,
2008). If terrestrial inputs had substantial impact on global surface CDOM abundance, it
would be accordingly expected that the Atlantic have disproportionately higher CDOM in
surface waters than the Pacific, as there is overall 3.6 times greater riverine discharge
entering the North Atlantic (Opsahl and Benner, 1997; Del Vecchio and Subramaniam,
2004); but this is not observed either (Siegel et al., 2002; 2005a). The large-scale spatial
patterns in satellite-observed global CDOM are overall oceanic in their distribution
(Siegel et al., 2005a).
3.2.2. In situ CDOM vs. radiometrically-derived CDOM in the subtropical South Pacific
Values of CDOM absorption in the South Pacific subtropical gyre were extremely
low, occasionally falling below our instrumental limit of detection (0.005 m-1). The
subtropical South Pacific is a notoriously hyper-oligotrophic region of the world ocean
where surface waters have been observed to be more optically clear than ultrapure water
reference standards (Morel et al., 2007a; 2007b; Tedetti et al., 2007). We sought to
compare radiometry-based estimates of CDOM with spectroscopically-measured CDOM
values along the P16S transect in an effort to reconcile the near-zero absorption values
measured spectrophotometrically. Radiometric in-water light measurements and
22
particulate absorption (ap) spectra measured daily along P16S were used to independently
determine in situ CDOM absorption following the inversion scheme of Morel et al.
(2007a).
A mast-mounted and underwater profiling radiometer concurrently measured
fluxes of the spectral downwelling irradiance, Ed(), and spectral upwelling radiance,
Lu(), over wavelength () range 324 – 683nm in the upper 150m along P16. Lu() was
converted to the spectral upwelling irradiance, Eu(), using factor Q(), which is modeled
as a function of environmental conditions of zenith sun angle, s, and chlorophyll-a
concentration, [Chl], at each wavelength (look-up table supplied by B. Gentili, 2008).
Spectral irradiance reflectance, R(), was then computed as the ratio Eu()/Ed(). The
diffuse attenuation coefficient for downwelling spectral irradiance, Kd() (m-1), was
computed from linear regression of the natural logarithm of Ed() over the top 20m
against depth. R() and Kd() were used in calculating total in-water absorption, atot(),
in the following manner (adopted from Morel et al. 2007a):
atot() = 0.962 Kd() d(s, [Chl], ) {1 – R() / f’(s, [Chl], )
(1)
where d and f’ are dimensionless coefficients that are also a function of s, [Chl] and 
(see ftp://oceane.obs-vlfr.fr/pub/morel/).
Values of atot represent the sum of absorption due to particulates (ap), CDOM
(acdom) and pure water (aw) at each wavelength:
atot() = ap() + acdom() + aw()
(2)
As mentioned in section 2.3, ap() was measured in situ along P16. To define a range for
aw(), we adopted the data set of aw() values (300 – 500nm) reconstructed by Morel et
al. (2007a and references therein) using estimates from prior studies. It was then possible
23
to infer acdom by inversion of equations 1 and 2. We refer to this inverse determination of
CDOM as a’cdom and compare it to the spectroscopic determination of CDOM from bottle
samples, acdom.
We selected four stations along P16S for the assessment (see Fig. 1, yellow
triangles). Two of these stations, at a respective 28°S and 33°S, fall within the lowchlorophyll (< 0.03 mg/m3) waters of the subtropical South Pacific gyre. The other two
stations were located outside of the hyper-oligotrophic region near the subtropical front at
latitudes 43°S and 45°S, and were selected for comparison. Fig. 5 demonstrates the
contrast in light attenuation, as represented by Kd(), among the four stations. Light
penetration at 325nm is nearly three times as great at the stations within the subtropical
gyre (Fig. 5, solid lines) than at the stations outside of the gyre (Fig. 5, dashed lines). The
Kd() values in the UV domain, and the Kd() minima observed at 410nm for the two
subtropical gyre sites, are in excellent agreement with those reported by Morel et al.
(2007a) for the eastern South Pacific gyre waters near Easter Island, a region
hypothesized as the clearest natural waters in the entire global ocean (Morel et al.,
2007b).
FIGURE 5
The results of the inversion are displayed in Fig. 6a – d, where individual panels
show the component absorption spectra for each of the four stations analyzed. The mean
deviation between a’cdom and acdom among all stations and wavelengths is 0.005 m-1 ±
0.001 m-1, with a bias toward higher absorption by a’cdom. The one exception to this
occurs at 28°S, where the two estimations, a’cdom and acdom, are a near match (Fig. 6a).
The mean deviation between a’cdom and acdom is not great enough, given our analytical
24
uncertainty (0.005 m-1), to apply as a correction for reference water impurity. The lowest
literature estimates for pure water absorption are limited to empirical values from
experimental laboratory attempts at purifying water (Morel et al., 2007a). Values of a’cdom
and acdom at 28°S for  > 375nm remain slightly negative, suggesting that even the aw
value supplied by Morel et al. (2007a) overestimates absorption by truly pure water (Fig.
6a). We therefore conclude that oceanic waters within the South Pacific subtropical gyre
may have lower absorption than the lowest literature estimates of pure water in the  >
375nm range. We interpret the extraordinary optical clarity of these waters as a
consequence of the extremely low biological activity in the gyre, supported by the low
Chl observations (Morel et al. 2007a; 2007b), and likely due to iron-limitation
(Behrenfeld and Kolber, 1999). Moreover, the large-scale downwelling circulation of the
gyre limits the input of elevated CDOM from subsurface waters, and the few dissolved
absorbing substances that are present in surface waters of the gyre are only further
bleached by the intense solar exposures at such latitudes.
FIGURE 6
3.3 Hydrography of CDOM in intermediate and deep water masses of the Pacific
We initially observe a strong linear relationship (r2 = 0.82, n = 1161) between
AOU and CDOM in the subthermocline waters of the Pacific (from 300m - bottom) (Fig.
7a), which contrasts with the weak relationships of DOC vs. CDOM (r2 = 0.03, n = 1146)
in the Pacific (Fig. 7b) and AOU vs. CDOM (r2 = 0.14, n = 1094) in the Atlantic (Fig.
7c). The correlation between AOU and CDOM in the Pacific is slightly weaker once
corrected for isopycnal mixing within the water masses (discussed in section 3.3.1), but
25
overall is still considerably higher than that observed within the Atlantic. There is a
significant negative correlation between AOU and the spectral slope parameter, Snlf (nm1
), (r2 = 0.52, n = 1152) over the Pacific (300m – bottom) as displayed in Fig. 7b. The
observation of low spectral slopes in association with high AOU is consistent with
production of new CDOM via remineralization processes in the Pacific. The correlation
coefficients between Snlf and AOU are smaller than that of CDOM and AOU because the
observed values of Snlf are close to zero and inhabit a small range.
Association of AOU with the optical properties of dissolved organic matter
(specifically the fluorescent fraction, FDOM, which is closely related to CDOM) have
been recently reported by Yamashita and Tanoue (2007; 2008) for the Pacific and
Southern Oceans, and linked to the oxidative remineralization of organic material with
water mass aging. However, a strong AOU-CDOM trend is not found in the Atlantic
(Fig. 7c), even as water masses age (Nelson et al., 2007; Nelson et al., in review), which
challenges the hypothesis that biogenic formation of CDOM occurs irrespective of
oceanic regime (Yamashita and Tanoue, 2008). Formation and advection of North
Atlantic Deep Water (NADW) is large-scale and rapid enough such that NADW transport
offsets any increases in CDOM from local biology (Nelson et al., 2007; Nelson et al. in
review), thus the AOU-CDOM correlation is much weaker in the North Atlantic interior
(Fig. 7c). There is no convective source of deep water in the North Pacific as there is for
the NADW due to the narrow Bering Strait and shallow Bering Sea impeding the inflow
of cold bottom water from the Arctic region (Pickard and Emery, 1990; Qiu and Huang,
1995). North Pacific Deep Water (NPDW) instead forms as an amalgamation of
southward-returning upwelled bottom waters (including CDW, AABW and the oldest
26
NADW) and vertical mixing with overlying intermediate waters (Fukasawa et al., 2004;
Ganachaud, 2003). The sluggish circulation throughout the deep parts of the Pacific
allows for CDOM accumulation over centennial time scales. As a result of these
considerations, and following mixing corrections (outlined in section 3.3.1), we
hypothesize that CDOM abundance and distribution can predict the relative strengths of
ventilation and advection versus local bioremineralization in ocean basins.
FIGURE 7
3.3.1 Isopycnal mixing in Pacific water masses
We have shown a strong relationship between AOU and CDOM in the Pacific,
which we attribute to biogeochemical drivers (Fig. 7a). In order for the interpretations of
the CDOM-AOU relationship initially observed across the Pacific between 300 – 5500m
(Fig. 7a) to be valid, it was necessary to break down the Pacific basin into water mass
layers as outlined in section 2.4 to determine which geographic and vertical sections of
the ocean are the dominant contributors to the observed strong trends of CDOM with
AOU. Having determined the water masses in which the strongest biogeochemical trends
with CDOM are observed, the potential influence of conservative processes (i.e.,
isopycnal mixing) on the observed covariance in hydrographic properties must be
considered (Deutsch et al., 2001; Hansell et al., 2004). We conducted a binary mixing
analysis following the methods of Hansell et al. (2004) for the water mass layers (Table
2) in which statistically significant biogeochemical relationships of CDOM were
observed.
27
The make-up of any given parcel of water within a water mass has a conservative
component due to physical mixing of preformed end-members that is independent of, and
thus not attributable to, any biogeochemical or diagenetic alteration (Anderson and
Sarmiento, 1994). We evaluated the potential effect of isopycnal mixing by subtracting
the preformed concentrations of hydrographic properties of interest from the observed
values in a water mass, and regressing the residual values against one another (Deutsch et
al., 2001; Hansell et al., 2004). We conducted a binary mixing analysis following the
methods of Hansell et al. (2004) for the water mass layers defined in Table 1 in which
statistically significant biogeochemical relationships of CDOM were observed. Clear
north and south end-member relationships were observed in the deep waters of the North
Pacific (NPDW, CDW and AABW), which were evaluated collectively as ‘Pacific Deep
Water’ (PDW). While it is clear from Fig. 2a that PDW spans the entirety of the Pacific
as sampled along 150W, only very weak CDOM-AOU correlation coefficients were
observed in the South Pacific (r2 < 0.10, n = 359), indicating this region was at best a
minor contributor to the overall basin-scale CDOM-AOU relationship observed (Fig. 7a).
This is likely due to the relatively homogeneous AOU and CDOM distributions in the
South Pacific sector of PDW.
The intermediate water of the North Pacific (NPIW) had distinct north and south
end-members as the water mass forms from the mixing of subarctic and subtropical
waters. We thus were able to account for end-member mixing in NPIW over its
latitudinal range (0 – 40N) as sampled on P16. No east-west end-member relationships
were present in NPIW, negating the need to correct for mixing when assessing
biogeochemical relationships in the zonal distribution of NPIW as sampled on P2. The
28
same outcome was found with AAIW, for which no north-south end-members could be
discerned among the hydrographic variables of interest. Lack of discernable endmembers within a water mass (such as within AAIW) indicated cases where local nonconservative (biogeochemical) processes, and not physical mixing, dominated the
hydrographic relationships to such a degree that mixing-correction for the regression
variables was neither a valid nor necessary step.
The correction method is illustrated as follows using the North Pacific section of
PDW as an example. To calculate the fractional contributions of northern and southern
end-members of each water mass, potential temperature, Tpot (°C), (a conservative tracer)
was plotted against the neutral density (γn) range of the water mass in question (Fig. 8,
PDW example). Visual inspection determined the approximate latitudinal range
encompassing the respective northern and southern envelopes of end-member data points.
Following this, the linear correlation between Tpot and γn for each envelope had to meet
the criterion of r2 ≥95% in order for the regression line to be used in modeling the
northern, fn, and southern, fs, fractional contributions of the water mass (Deutsch et al.
2001; Hansell et al. 2004). In the case of PDW, 48 – 55°N and 0 – 8°N encompassed the
northern and southern components, respectively. (Fig. 8, annotated black regression
lines).
FIGURE 8
The regression of Tpot with n for each envelope becomes the linear model that
determines the northern and southern end-members of potential temperature (Tn and Ts)
for all γn within the water mass range. These model values were applied in calculating fn
and fs in conjunction with the observed potential temperature values (Tobserved) as follows:
29
fs = (Tn – Tobserved) / (Tn – Ts)
(3)
fn = 1 – fs
(4)
The preformed value can be determined for any variable of interest, C, within a
water mass. Cpreformed is calculated using fn and fs along with the northern, Cn, and
southern, Cs, end-members of that variable of interest determined by regression of C with
γn. The following equations were used to calculate Cpreformed in this manner, and
ultimately the residual concentration, C, in which preformed values are subtracted from
the originally observed values of the property of interest.
Cpreformed = (fn * Cn) + (fs * Cs)
(5)
C = Cobserved – Cpreformed
(6)
Relationships among the residuals of hydrographic properties are then assessed
independently of isopycnal mixing effects. Fig. 9 shows for the case of PDW that the
initially strong correlation observed between AOU and CDOM (Fig. 9a, r2 = 0.83, n =
249) is considerably weakened when the variable concentrations are corrected for mixing
effects (Fig. 9b, r2 = 0.26, n = 249). This indicates that lateral isopycnal mixing plays a
strong role in influencing the distribution and coincident patterns of both CDOM and
AOU observed in the deepest waters of the North Pacific. This result is somewhat
expected given that the approximate depth horizon of PDW (2000 – 5500m) lies below
the remineralization zone of the water column. Very limited biological activity (and thus
oxygen utilization) that contributes to CDOM accumulation takes place below 2000m;
thus oxidative bioremineralization in fact accounts for only one-fourth of the variance in
CDOM within PDW.
FIGURE 9
30
The negative residuals of CDOM and AOU in Fig. 9b resulting from the endmember analysis suggest a loss of material relative to preformed abundances. In the
absence of considerable net production of CDOM through microbial activity in the water
column, it is possible that there is slight degradation of CDOM over long periods of time
in the abyssal layers of the ocean. Diagenesis of CDOM spectral quality is hypothesized
to occur over time scales of North Atlantic Deep Water circulation in the deep ocean
(Nelson et al., 2007). The potential adsorption of refractory dissolved substances onto
particles in the deep ocean as hypothesized by Druffel et al. (1998) may also explain
CDOM degradation. On the other hand, oxygen evolution in the deep ocean as implied by
the slightly negative AOU residuals, is most likely an artifact resulting from the
undersaturation of oxygen in the surface ocean in regions of deep water formation as
confirmed by Ito et al. (2004).
3.3.2 Mixing-corrected basin-scale biogeochemical relationships of CDOM
We can now make some final assessments regarding the strength of
biogeochemical influence on CDOM abundance and distribution within the Pacific water
masses. Table 2 displays the least-squares linear regression values among CDOM, AOU,
DOC, NO3 and DIC in the Pacific water masses after mixing corrections were applied
where necessary. (Table 2 excludes data points within the 0 – 100m surface layer where
solar bleaching directly influences CDOM abundance.) Across the meridional extent of
NPIW (at 150°W), mixing had a relatively modest influence on the biogeochemical
relationships with CDOM. A strong correspondence between AOU and CDOM in NPIW
(r2 = 0.61, n = 176) remains after correcting for mixing. In the zonal direction across the
31
North Pacific (along 30°N), where isopycnal mixing was not a factor in NPIW, there is a
weaker correlation between CDOM and AOU (r2 = 0.29, n = 155).
Water mass
Lat. range
Lon. range
AOU vs. CDOM
DOC vs. CDOM
NO3 vs. CDOM
DIC vs. CDOM
NPIWcorrected
0 – 40N
150W
r2 = 0.61, n = 176
r2 = 0.00, n = 174
r2 = 0.49, n = 176
r2 = 0.52, n = 171
NPIW
30N
r2 = 0.29, n = 155
r2 = 0.17, n = 129
r2 = 0.30, n = 156
r2 = 0.30, n = 155
PDWcorrected
(NPDW+CDW+
AABW)
AAIW
0 – 55N
133E –
145W
150W
r2 = 0.26, n = 249
N/A
r2 = 0.01, n = 250
r2 = 0.41, n = 251
60S – 15°N
150W
r2 = 0.86, n = 297
r2 = 0.06, n = 288
r2 = 0.75, n = 298
r2 = 0.77, n = 292
Table 2. Least-squares regression statistics among hydrographic variables in Pacific water mass layers. r2 values in bold are
significant at the 95% confidence interval. NPIW (at 30N between 133E – 145W) and AAIW did not have specifiable endmembers in the parameters of interest, therefore no mixing correction was made. N/A = No regression due to unspecifiable endmembers of DOC in PDW.
The overall strongest linear relationships between CDOM and remineralization
indices (AOU, NO3, DIC) throughout the Pacific basin are observed within AAIW and
NPIW (along 150°W). It is expected that the ventilated nature of intermediate waters
would prevent observation of CDOM accumulation from subsurface microbial processes;
however, the approximate depth horizon of intermediate waters (300 – 1500m) overlaps
the remineralization zone that spans the mesopelagic. The strength of the AOU-CDOM
relationship in intermediate waters therefore provides information on the competing rate
processes of advection and remineralization. For example, AAIW is a source of renewal
for the subthermocline waters of the South Pacific, however its formation and transport is
notably slow in the Pacific sector of the subantarctic (England et al., 1993). The rate of
local remineralization in AAIW must significantly exceed the advection rate given the
remarkably tight correspondence of CDOM with AOU (r2 = 0.86, n = 297). The CDOMAOU relation is slightly weaker (r2 = 0.61, n = 176) in NPIW (along 150°W) than in
AAIW, suggesting advection competes to a greater extent with CDOM production within
NPIW than AAIW. When considering NPIW over its east-west distribution (along 30°N),
32
which aligns with the zonal orientation of the water mass, the competitive role of
advection in NPIW is emphasized by the fact that remineralization indices explain
narrowly one-third of CDOM variability (Table 2).
CDOM does not substantially covary with free-standing DOC in any of the water
masses assessed (Table 2). This indicates that microbial utilization of labile and semilabile DOC that is released during particulate matter solubilization is the likely source of
CDOM production in the interior of the open ocean. The particulate flux of organic
matter appears to provide the essential initial substrate for the process of microbial
CDOM formation. This is further supported by experimental evidence revealing that
microbial utilization of free-standing labile DOC by itself can not account for CDOM
accumulation in the water column, as microbes simultaneously produce and consume
CDOM during this pathway (Nelson et al., 2004). Furthermore, CDOM comprises a
small percentage of the recalcitrant DOC pool (Nelson et al., 2002), and CDOM
dynamics appear decoupled from bulk DOC dynamics in the deep ocean. This trend is
corroborated by data from the North Atlantic (Nelson et al., 2007) and confirms that
CDOM and DOC stocks are not regulated by the same mechanisms in the open ocean.
The relationships of CDOM with NO3 and DIC closely follow the AOU-CDOM
relationships as further support for a remineralization-driven CDOM signal in the open
ocean (Table 2). The slightly weaker regressions of CDOM with these variables may be
explained by potential influence of nitrification/denitrification and carbonate
precipitation/dissolution processes on deep ocean concentrations of NO3 and DIC,
respectively (Deutsch et al., 2001; Li and Peng, 2002).
33
The use of graphical inspection for identifying water mass end-members imparts a
subjective element to isopycnal mixing analyses (Li and Peng, 2002). Given this, other
attempts at evaluating mixing influence in the water masses discussed might vary
quantitatively, but not qualitatively, from the present study. Nevertheless, our assessment
provides a valuable general indication of where the relative biogeochemical and physical
drivers dominate with respect to the CDOM distribution in the Pacific basin.
On a final note, it is unlikely that the oxygen-related CDOM distribution in the
Pacific is due to abiotic processes (e.g., alteration of spectral properties due to changes in
redox state of the dissolved substances). As collection and storage of CDOM samples are
conducted in the presence of oxygen, our measurement protocols would inhibit
observation of such a phenomenon (and thereby the CDOM-AOU relationship) in the
deep ocean. Furthermore, copious experimental data show that net production of CDOM
occurs through biotic means (Yamashita and Tanoue, 2004; Nelson et al., 2004;
Steinberg et al., 2004), while there is no empirical evidence to date for abiotic oxygenrelated changes in CDOM absorption in the non-sunlit layers of the open ocean.
4. Conclusion
We conclude that the patterns in subsurface CDOM abundance in the Pacific
basin are attributed to the precipitation of particulate organic material in the water
column as it is modulated by subduction and advection. Azam (1998) proposed that
phases of organic matter in the water column (i.e., particulate, dissolved) are more
accurately conceptualized as a continuum that includes classifications of submicrometer
particles, gel-like matrices and colloids that generate polymeric “hotspots” of bacterial
34
activity. In view of this, it is particularly fascinating that an organically-formed optical
property, CDOM, emerges as a stable tracer of temporally integrated microbial activity in
a chemically and physically heterogeneous pool of organic material.
The strongest association observed between CDOM and heterotrophic oxygen
utilization in the Pacific occurs within intermediate water mass layers. This association is
governed by both the meridional gradient in ventilation age, and the vertical gradient in
remineralization in the intermediate layer as it coincides with the zone (100 – 1000m) of
high oxygen utilization rates in the water column (Feely et al., 2004). Organic carbon
remineralization rates at intermediate depths in the Pacific are observed to be highest in
regions characterized by high fluxes of calcium carbonate and biogenic silica (ballast
minerals) such as the subarctic (Feely et al., 2004). This is consistent with our hypothesis
that subsurface CDOM production, and correspondingly the AOU-CDOM relationship, is
primarily moderated by the downward flux of particles, and thereby closely linked to
carbon export.
Antarctic intermediate waters are of paramount importance to the oceanic sink for
anthropogenic CO2, whose greatest uncertainty is at intermediate depths (Sabine et al.,
2004; Sloyan and Rintoul, 2001; Talley, 1997). We have shown that low surface
abundances of CDOM entrained by intermediate and mode waters, as well as high
CDOM abundances resulting from export production at intermediate depths, each provide
potential information regarding the relative residence time and fate of organic carbon
harnessed in the intermediate layers of the global ocean.
Overall, we have established that strong end-member mixing relationships explain
very little variability in our parameters of interest in the Pacific Ocean except in the
35
oldest deep waters below the zone of remineralization. In these layers, isopycnal mixing
has considerable influence on the hydrographic relationships and there is minimal
apparent in situ oxygen consumption. The implication of this finding is that precautions
must be taken when assuming CDOM production rates in abyssal waters from simple
AOU-derived respiration rates as previous studies have attempted to do (Yamashita and
Tanoue, 2008). The covariance of AOU and CDOM distributions in abyssal waters
appears to be primarily regulated by physical processes and not microbial processes.
CDOM in the mesopelagic ocean holds potential as a predictive tool for
diagnosing the relative strengths of biogeochemistry and upper ocean renewal processes
not only within an ocean basin (e.g., NPIW versus AAIW), but among ocean basins as
well (e.g., North Pacific versus North Atlantic). Recent surveys of CDOM in the Indian
Ocean, a basin characterized by a similar meridional oxygen gradient as in the Pacific
due to lack of ventilation north of the equator, show that AOU and CDOM also have very
close correspondence (Nelson et al., in review). Schlitzer (2004) described an inverse
method for modeling biogeochemical rate constants and physical transports from watercolumn concentration data (specifically nutrient, carbon and oxygen data). We observe
trends in CDOM abundance in the open ocean that consistently reflect the basin-scale
interrelationships between physical dynamics and biogeochemistry. Moreover, the seasurface distribution of CDOM captures examples of coupling between epipelagic and
deep water masses. There is clear potential for water column CDOM data acquired via
the CO2/CLIVAR Repeat Hydrography Program (as well as the accompanying satellite
record of global surface CDOM data) to be integrated into a biogeochemical and 3-D
ocean circulation model, such as described by Schlitzer (2004), to predict flow fields and
36
carbon export fluxes in the world oceans. We have constrained areas of the ocean where
CDOM has a near-stoichiometric relationship with O2 utilization, and can qualitatively
account for where departures from this relationship are directly influenced by ventilation
and flow strength. It is conceivable that the future multi-year dataset of surface and
thermocline/deep ocean CDOM could be parameterized within a hindcast model, as has
been demonstrated using temporal variability in O2 (Deutsch et al., 2006), to even
estimate climate-related physical and biogeochemical changes in the upper ocean.
Acknowledgements
We acknowledge the support of NSF Chemical Oceanography (OCE-0241614
and OCE-0648541) and NASA Ocean Biology and Biogeochemistry to N. Nelson, D.
Siegel and C. Carlson, and NASA Earth System Science Fellowship Program to C. Swan.
We thank the CO2/CLIVAR Repeat Hydrography Program, chief scientists Jim Swift,
Chris Sabine, Richard Feely, and Paul Robbins, as well as captains and crew of the R/Vs
Revelle, Melville and Thompson. Dennis Hansell’s group (UMiami) conducted DOC
analysis for the P16N and the western section of P2. We thank Stacy Brown (UMiami)
for CDOM collection on P2, Dave Menzies, Stuart Goldberg, Meredith Meyers and Elisa
Wallner (UCSB) for field assistance on P16S and P16N, and Kirk Ireson for analytical
assistance.
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FIGURE CAPTIONS:
Fig. 1. SeaWiFS eight-year mean composite of absorption by CDOM and detrital
particulates at 443nm (‘acdm443’) in the Pacific as determined by the GSM algorithm
(Siegel et al. 2005a, 2005b). Black lines = CLIVAR transects P2 (zonal 30N) and P16
43
(meridional 150W). Yellow triangles = P16 stations at which inverse determinations of
apparent optical properties using radiometry data have been conducted for this study.
Fig. 2. A – C. Full-depth (5500m) hydrographic distributions along 150°W in the Pacific
from 71°S to 57°N (P16 transect). (A) Chromophoric dissolved organic matter (CDOM)
with overlain neutral density (γn) isopycnals. Annotations indicate approximate locations
of the water masses discussed. (B) Spectral slope parameter, Snlf, with overlain neutral
density (γn) isopycnals. (C) Apparent oxygen utilization (AOU) with overlain contour
lines of salinity.(D) Dissolved organic carbon (DOC) with overlain γn isopycnals.
Fig. 3. A – C. Upper ocean (1000m) hydrographic distributions along 30°N in the North
Pacific from 133°E to 117°W (P2 transect). (A) CDOM with overlain γn isopycnals.
Annotations represent approximate locations of NPIW and North Pacific STMW. (B) Snlf
with overlain γn isopycnals. (C) AOU with overlain contour lines of salinity. (D) DOC
with overlain γn isopycnals.
Fig. 4. Distribution of CDOM in surface water samples (ca. 5m) collected and analyzed
shipboard (▲) along 150°W (P16), and absorption by CDOM and detrital particulates at
325nm (*) as extrapolated from GSM retrieval of CDOM absorption at 443nm using
merged SeaWiFS and MODIS Aqua data (Siegel et al. 2005a, 2005b). (Regression
between satellite-derived (GSM) and in situ CDOM is r2 = 0.72, n = 59, CDOMGSM =
1.743*CDOMin situ – 0.007.) Open-ocean detrital absorption is minor, thus satellitederived values are primarily representative of CDOM (Siegel et al. 2002).
Fig. 5. Diffuse attenuation coefficient, kd (m-1), of spectral downwelling irradiance for
wavelength range 325 – 555nm computed over 0 – 20m using radiometric profile data
from four station locations along 150°W (P16).
Fig. 6. A – D. Component absorption (m-1) in surface waters (ca. 5m) at each of four
stations (P16) used in radiometric inversion analysis. Solid black lines represent the
observed (spectroscopic) in situ CDOM absorption, acdom (▲) and the resultant inverselycalculated (radiometric) CDOM absorption, a’cdom (●). Dotted lines represent the pure
water absorption values, aw (x), adopted from Morel et al. (2007a), in situ absorption by
algal particulates, ap (◊), as determined using the quantitative filter technique, and total
absorption, atot (□).
Fig. 7. Scatter plots of (A) AOU vs. CDOM (r2 = 0.82, n = 1161) (B) AOU vs. Snlf (r2 =
0.52, n = 1152) and (C) DOC vs. CDOM (r2 = 0.03, n = 1136) in Pacific waters (300 –
5500m) from 60°S to 55°N (P16). (D) Scatter plot of AOU vs. CDOM (r2 = 0.14, n =
1094) in Atlantic waters (300 – 5500m) within geographic limits 60S – 60N by 20W –
69W (CLIVAR transects A20, A22 and A16). Data within surface waters (z < 100m) are
omitted.
Fig. 8. Scatter plot of potential temperature (Tpot) vs. γn for PDW (γn = >27.8 kg/m3).
Black lines are the regressions used in the present analysis for determining the northern
44
(r2 = 0.99, n = 67) and southern (r2 = 0.99, n = 50) end-member components of PDW.
Data within surface waters (z < 100m) are omitted.
Fig. 9. Scatter plots of AOU vs. CDOM in PDW (A) before mixing effects are considered
(r2 = 0.83, n = 249), and (B) after isopycnal mixing effects are removed (r2 = 0.26, n =
249). Data within surface waters (z < 100m) are omitted.