The effect of surface irradiance on the

Transcription

Deep-Sea Research I 63 (2012) 52–64
Contents lists available at SciVerse ScienceDirect
Deep-Sea Research I
journal homepage: www.elsevier.com/locate/dsri
The effect of surface irradiance on the absorption spectrum of chromophoric
dissolved organic matter in the global ocean
Chantal M. Swan a,n, Norman B. Nelson a, David A. Siegel a,b, Tihomir S. Kostadinov a
a
b
Earth Research Institute, University of California, Santa Barbara, USA
Department of Geography, University of California, Santa Barbara, USA
a r t i c l e i n f o
a b s t r a c t
Article history:
Received 10 October 2011
Received in revised form
13 January 2012
Accepted 23 January 2012
Available online 1 February 2012
The cycling pathways of chromophoric dissolved organic matter (CDOM) within marine systems must
be constrained to better assess the impact of CDOM on surface ocean photochemistry and remote
sensing of ocean color. Photobleaching, the loss of absorption by CDOM due to light exposure, is the
primary sink for marine CDOM. Herein the susceptibility of CDOM to photobleaching by sea surfacelevel solar radiation was examined in 15 samples collected from wide-ranging open ocean regimes.
Samples from the Pacific, Atlantic, Indian and Southern Oceans were irradiated over several days with
full-spectrum light under a solar simulator at in situ temperature in order to measure photobleaching
rate and derive an empirical matrix, esurf (m 1 mEin 1), which quantifies the effect of surface irradiance
on the spectral absorption of CDOM. Irradiation responses among the ocean samples were similar
within the ultraviolet (UV) region of the spectrum spanning 300–360 nm, generally exhibiting a
decrease in the CDOM absorption coefficient (m 1) and concomitant increase in the CDOM spectral
slope parameter, S (nm 1). However, an unexpected irradiation-induced increase in CDOM absorption
between approximately 360 and 500 nm was observed for samples from high-nutrient low-chlorophyll
(HNLC) environments. This finding was linked to the presence of dissolved nitrate and may explain
discrepancies in action spectra for dimethylsulfide (DMS) photobleaching observed between the
Equatorial Pacific and Subtropical North Atlantic Oceans. The nitrate-to-phosphate ratio explained
27–70% of observed variability in esurf at observation wavelengths of 330–440 nm, while the initial
spectral slope of the samples explained up to 52% of variability in esurf at observation wavelengths of
310–330 nm. These results suggest that the biogeochemical and solar exposure history of the water
column, each of which influence the chemical character and thus the spectral quality of CDOM and its
photoreactivity, are the main factors regulating the susceptibility of CDOM to photodegradation in the
surface ocean. The esurf parameter reported herein may be applied to remote sensing retrievals of CDOM
to estimate photobleaching at the surface on regional to global scales.
& 2012 Elsevier Ltd. All rights reserved.
Keywords:
Marine CDOM
Solar irradiation
Surface
Photobleaching
Photoproduction
1. Introduction
Marine chromophoric dissolved organic matter (CDOM), the
light-absorbing portion of total dissolved substances in seawater
(filtrateo0.2 mm), is responsible for nearly 90% of ultraviolet (UV)
radiation attenuation in the global ocean (Johannessen et al.,
2003; Zepp et al., 2011). In open ocean waters, CDOM is produced
through autochthonous marine food web processes throughout
the water column, and seasonally depleted through solar photobleaching in the surface ocean (Nelson et al., 1998, 2004; Del
Castillo and Coble, 2000; Stedmon and Markager, 2001). Photobleaching of CDOM, driven by absorption of solar radiation,
n
Corresponding author. Tel.: þ41 78 849 4755; fax: þ 1 805 893 2578.
E-mail address: [email protected] (C.M. Swan).
0967-0637/$ - see front matter & 2012 Elsevier Ltd. All rights reserved.
doi:10.1016/j.dsr.2012.01.008
typically results in loss in absorption by the chromophores and
subsequent alteration of the spectral properties of CDOM (Del
Vecchio and Blough, 2002; Helms et al., 2008; Swan et al., 2009).
The satellite-derived 8-year mean climatology of colored
dissolved and detrital material from SeaWiFS (1997–2005)
(Fig. 1, ‘CDM’, of which CDOM comprises480%; Nelson et al.,
1998) illustrates that solar bleaching, as balanced by CDOM
inputs from river outflow and in situ production, is a prominent
forcing mechanism on the global surface distribution of CDOM on
multi-annual time scales. Low CDOM absorption values are found
in surface waters with slow renewal rates and high radiation
exposure (e.g., subtropical gyres), while elevated CDOM values
are seen in regions of lower annual insolation (e.g., subpolar
gyres) or regions of significant upwelling of CDOM from subsurface waters (e.g., equatorial Pacific and eastern boundary
currents) (Siegel et al., 2002, 2005). The influence of photobleaching
C.M. Swan et al. / Deep-Sea Research I 63 (2012) 52–64
Fig. 1. SeaWiFS eight-year mean (1997–2005) global composite of absorption by
CDOM and detrital particulates at 443 nm, ‘aCDM (m 1, 443 nm)’ as determined by
the GSM algorithm (Siegel et al., 2005; Maritorena et al., 2010). Detrital
particulates contribute only a small percentage (5–17%, Nelson et al., 1998;
Siegel et al., 2002) to overall absorption in the open ocean, thus aCDM (m 1,
443 nm) is considered as overall representative of CDOM. White triangles
represent the sample sites at which laboratory determinations of apparent
quantum yield for CDOM photobleaching were conducted for this study.
is also observed at depth, such as in Atlantic mode waters, which
manifest as minima within the basin-scale CDOM distribution as
photobleached waters are entrained subsurface during convective
overturn (Nelson et al., 2007, 2010).
The solar exposure of CDOM also plays a significant role in
balancing the ocean carbon budget with respect to terrestrial
inputs of organic material to the sea, as photodegradation of
CDOM within rivers and among continental shelf waters serves as
an important removal and conversion pathway of terrigenous
material into climate-relevant trace gases (Blough et al., 1993;
Hedges et al., 1997; Vodacek et al., 1997) and lower molecular
weight organics (Benner and Biddanda, 1998; Miller et al., 2002;
Stedmon et al., 2007; Tedetti et al., 2008).
Parameterization of CDOM photobleaching in the surface
ocean is critical for a better understanding of its effect on
elemental cycles and satellite detection of ocean color-derived
properties. Direct measurements of the rates of CDOM photobleaching within the global open ocean are scarce, as most derive
from near-coastal and estuarine waters or are concentrated
within the North Atlantic (Vähätalo et al., 2000; Moran et al.,
2000; Osburn et al., 2001, 2009; Del Vecchio and Blough, 2002;
Tzortziou et al., 2007). In these environments, considerable losses
of CDOM (approximately 50%, by most studies) occur on weekly
to monthly time scales (Kouassi and Zika, 1992; Vodacek et al.,
1997; Nelson et al., 1998). Photobleaching can remove up to 96%
of terrestrial CDOM that enters the ocean, and near-zero CDOM
absorption coefficients reported for the South Pacific subtropical
gyre suggest that solar bleaching may completely degrade autochthonous CDOM in oligotrophic areas of the open ocean
(Vähätalo and Wetzel, 2004; Swan et al., 2009). It remains to be
determined whether there is a fraction of marine CDOM that is
photochemically recalcitrant (Stedmon and Markager, 2001;
Twardowski and Donaghay, 2002).
The chief objectives of this paper are (1) to describe how
CDOM photobleaching proceeds at the surface of the open ocean,
particularly in tropical and subtropical areas that have not been
previously investigated, (2) to examine the environmental
controls on CDOM photobleaching, and (3) to quantify the
photochemical susceptibility of CDOM on the seasonal scale for
potential application to remote sensed CDOM data in estimating
global surface photobleaching rates.
The challenge of resolving measurable changes in CDOM under
natural sunlight during typical cruise durations ( 5 weeks or
less) led to the need for conducting accelerated exposures on
shore using a solar simulator. Compounding this challenge,
narrow bandwidth or partial spectral exposures, such as those
53
employed in traditional approaches to calculating apparent
quantum yields for photobleaching, may not result in light flux
sufficient to produce detectable losses in CDOM absorption in
open ocean samples, particularly those from highly oligotrophic
gyres in which CDOM absorption is lowest. Recent work has
indicated that chromophores irradiated by monochromatic
irradiation have an indirect (‘‘off-axis’’) effect on other chromophores not directly affected by the irradiation wavelength (attributed to electron transfer or photosensitizing intermediate
species), emphasizing that chromophore interdependence must
be considered when designing photobleaching studies (Del
Vecchio and Blough, 2002, 2004; Goldstone et al., 2004;
Tzortziou et al., 2007; Ziolkowski and Miller, 2007). Due to these
constraints, full-spectrum simulated solar irradiance was
employed in controlled laboratory incubations of filtered water
samples collected from a variety of open ocean water types (Fig. 1,
white triangles) to explore CDOM photobleaching at the surface
of the global ocean. Results were used to develop an empirical
model for the wavelength-specific effect of irradiance on spectral
CDOM absorption, which encapsulates the net effect of chromophores undergoing simultaneous photobleaching, for potential
application with ocean color-derived surface data. Thus, a mathematical description of bleaching is derived using the irradiance
and CDOM absorption spectra as ‘inputs’.
We characterize a surface photobleaching effect matrix, esurf
(m 1 mEin 1), which quantifies the susceptibility of open ocean
CDOM to typical illumination conditions at the sea surface. We
subsequently examine the relationship of esurf to environmental
variables that relate to the autochthonous production and
destruction of CDOM to assess whether esurf is a function of the
quantity or the quality of open ocean CDOM (Osburn et al., 2001;
Twardowski and Donaghay, 2002; Nelson et al., 2004; Biers et al.,
2007). We report both typical and novel photochemical transformations of light-exposed CDOM, and further discuss the implications of these within the context of the optical and hydrographic
conditions of the sampling regions. These results open a pathway
toward a better understanding and characterization of CDOM
dynamics in the global surface ocean.
2. Methods
2.1. Sample collection
Water samples for the present study were collected shipboard
during austral or boreal summer cruises from various upper ocean
depths around the globe during 2005–2008 as part of the U.S.
CO2/CLIVAR Repeat Hydrography Survey (P16S, P16N, I8S, I9N and
P18 field campaigns; Feely et al., 2005). A subtropical North
Atlantic (Sargasso Sea) sample was collected during a U.S.
Bermuda Atlantic Time-series Study (BATS) core cruise in winter
2006, and a sample from the Santa Barbara Channel, CA was
collected in winter 2007. Geographical locations of the sample
sites are listed within Table 1 and represented by white triangles
in Fig. 1. Samples were collected from various depths (from
surface—500 m) in an effort to sample the range of CDOM that
may be cycled into the mixed layer during seasonal changes. This
sampling scheme provided a diverse backdrop of optical and
hydrographic regimes for assessing environmental controls on
CDOM photobleaching in the surface ocean.
Two liters of seawater were drawn from Niskin or PTFE-coated
Go-Flo bottles, and either pressure-filtered through acid-leached
0.4 mm Nuclepore polycarbonate membranes, or vacuum-filtered
through 0.2 mm Nuclepore polycarbonate membranes (preconditioned with ultrapure water from a Barnstead Nanopure Diamond
UV system) in order to remove bacterial cells and other particulates
0.0005
0.0004
0.0005
0.0004
0
0.0001
0.0006
0.0001
0.0005
0.0006
0.0003
0.0008
0.0007
0.0005
0.0009
0.0030
0.0022
0.0028
0.0018
0
0.0025
0.0027
0.0013
0.0007
0.0005
0.0005
0.0006
0.0008
0.0002
0.0013
0.009
0.006
0.009
0.006
0
0.009
0.008
0.006
0.004
0.000
0.007
0.004
0
0
0
0.068
0.038
0.061
0.040
0
0.068
0.042
0.044
0.037
0.021
0.055
0.005
0
0
0
0.217
0.073
0.082
0.072
0.031
0.093
0.073
0.103
0.104
0.153
0.188
0.073
0.091
0.070
0.079
0.14
o0.01
0.05
0.03
0.25
0.15
0.07
1.49
2.42
1.67
2.87
0.90
0.80
0.59
1.6
0.20
o 0.01
0.08
0.07
0.10
0.14
0.06
21.36
33.32
19.07
42.12
10.74
9.34
6.32
24.32
1.4
1.0
1.6
2.3
0.4
0.9
0.9
14.3
13.8
11.4
14.7
11.9
11.7
10.7
15.2
0.023
0.032
0.026
0.028
0.038
0.027
0.029
0.026
0.029
0.028
0.026
0.032
0.028
0.031
0.032
esurf (325;325)
(m 1 lEin 1)
esurf (310;310)
(m 1 lEin 1)
P (lmol kg 1)
Fig. 2. CDOM absorption coefficient (m 1, 325 nm) of samples after dark storage
at 4 1C plotted versus CDOM absorption at time of collection. 1:1 line is plotted.
Closed circles (K) represent samples stored for periods of 6–12 months. Open
triangles (W) represent samples stored for 13–22 months. Error bars reflect
precision in the CDOM absorption measurement at 325 nm, 7 0.013 m 1, at the
95% confidence interval determined from replicate samples in the Pacific (Nelson
et al., 2007).
(Whitehead and de Mora, 2000; Nelson and Siegel, 2002; Nelson
and Coble, 2009). Filtrates were refrigerated in polycarbonate 1-L
containers in the dark at 4 1C until their use in shore-based
laboratory irradiation experiments conducted within a year of
sample collection. Tests show stability in CDOM absorption spectra
of open ocean samples for up to 12 months with this storage
protocol. Fig. 2 displays CDOM absorption at 325 nm in samples
stored for lengths of time ranging from 6 to 22 months versus
CDOM absorption at the time of sample collection. No statistically
significant variation in CDOM absorption at 325 nm was observed
in samples stored for 12 months or less; however, CDOM absorption in filtrates stored for 13–22 months varied in magnitude by up
to 0.028 m 1, with a slight bias toward increased CDOM relative to
initial (Fig. 2). This indicated that a 1-year time limit for filtrate
storage was an appropriate conservative guideline for the
experimental work.
Z (m)
0
80
140
40
25
40
40
200
500
80
200
100
80
5
50
Santa Barbara Channel, CA
Sargasso Sea
Subtropical N. Pacific
Subtropical N. Pacific
Subtropical S. Pacific
Subtropical Indian
Equatorial Indian
Equatorial Pacific
Subtropical S. Pacific (deep)
Subarctic Pacific
Subarctic Pacific
Equatorial Pacific
Subantarctic Pacific Frontal Zone
Subantarctic Pacific Frontal Zone
Southern Ocean
341N
321N
291N
291N
211S
311S
1.81N
0.51N
211S
551N
551N
0.51N
461S
461S
621S
1201W
641W
1501W
1501W
1031W
951E
921E
1501W
1031W
1501W
1501W
1501W
1501W
1501W
1031W
16
23
19
20
23
20
30
13
7
3
4
22
10
13
4
2.2. Experimental design
Sample Site
Lat.
Lon.
Temp. (1C)
N (lmol kg 1)
N:P
initial S
(nm 1)
initial aCDOM
(325, m 1)
esurf (350;350)
(m 1 lEin 1)
esurf (440;440)
(m 1 lEin 1)
Table 1
Hydrographic information and surface photobleaching parameters at selected wavelengths, esurf(lo;li), in ocean samples. A value of zero for esurf(lo;li) indicates that CDOM was resistant to solar irradiation over the exposure
period. Positive values of esurf(lo;li) indicate that photoproduction of CDOM was observed over the exposure period.
54
CDOM irradiation experiments were conducted using an
LS1000 Solar Simulator (Solar Light Co., Glenside, PA) fitted with
a high-pressure Xenon arc lamp and filters specialized to closely
match the spectrum and intensity of terrestrial irradiance
(Mobley, 1994). The practical lower limit of terrestrial irradiance
is 290 nm under normal stratospheric ozone levels (Whitehead
and de Mora, 2000). An array of twelve 10-mL cylindrical quartz
cells with PTFE-lined caps were used as sub-sample compartments with no head space and arranged on a matte black surface
within a dark side-walled enclosure to minimize reflection,
backscattering and container wall effects among sub-samples
(Johannessen and Miller, 2001; Hu et al., 2002). The sample cell
array was submersed within a water bath set to in situ temperature at which the original sample was collected. A dark vial in the
water bath served as a control for any non light-related alteration
of CDOM during the course of irradiation. The sample array was
positioned under the collimated light beam of the simulator at the
manufacturer-specified distance to ensure uniform spectral quality and intensity over the sample exposure area. The spectral
irradiance within the exposure area was periodically quantified
and tested for its spatial uniformity using an 11-channel UV–vis
aCDOM (λo)
Eo (λi)
14
55
(μEin m−2 s−1 nm−1)
12
10
8
6
4
LS1000 Solar Simulator spectrum
2
0
300
400
500
λi (nm)
600
λo (nm)
Qa (λi)
εsurf (325;λi)
Qa (λi) εsurf (325;λi)
λi (nm)
λi (nm)
λi (nm)
Fig. 3. Schematic of terms used to model the CDOM surface photobleaching effect using experimental data from irradiation of Sargasso Sea 80 m water. (A) Quantum
scalar irradiance, Eo(li) (mEin m 2 s 1 nm 1), is the output of the LS1000 solar simulator. (B) Time course of the CDOM absorption spectrum, aCDOM(lo) (m 1), during
irradiation (with inset of spectral slope parameter, S (m 1), versus exposure time in days) is used to determine (C) the rate of change in CDOM absorption, d(aCDOM(lo))/dt
(m 1 s 1), at the given observation wavelength (lo ¼ 325 nm). (D) The time course of absorbed quanta, Qa(li) (mEin s 1 nm 1) is the product of Eo(li), āCDOM (li), and the
sample cell volume, v. (E) The surface photobleaching effect parameter, esurf(325;li) (m 1 mEin 1), is solved by inversion of Eq. (2) (see text). The modeled value of
d(aCDOM(325))/dt is the area underneath the curves of (F) the action spectrum for CDOM photobleaching at 325 nm, which equals the product of Qa(li) and esurf(325;li).
MicroPro IITM multispectral radiometer (Satlantic, Inc., Halifax,
Nova Scotia), as well as monitored daily for variation in flux
intensity using a QSL-2201 VIS wand (Biospherical Instruments,
San Diego, CA). The spectrum of the LS1000 Solar Simulator is
shown in Fig. 3A. Variation in simulator intensity over the course
of any individual experiment was negligible. Gradual, spectrally
independent variation of simulator intensity ( 20%) was
observed across a longer time scale of a few months. This
variation was well constrained by the described radiometric
calibrations, and applied to the experimental data.
Experiments were conducted for approximately 2 days under
the simulator (continuous exposure). This exposure period
roughly corresponds to 10 days at the sea surface at mid-latitude,
or 2–8 weeks of exposure when circulation within the seasonal
mixed layer during summer and spring is taken into account.
These approximations were based on data from the North
Atlantic, and were achieved by comparing the simulator’s
24-hour output at 325 nm (46.3 kJ m 2; calculated from its
irradiance at 325 nm; Fig. 3A) with the mean daily insolation at
325 nm estimated for the BATS site in late summer (9.3 kJ m 2)
and early spring (8.2 kJ m 2), respectively (Zafiriou et al., 2008).
This showed that a 48-hour exposure period under the simulator
was equivalent to 10.0 and 11.3 day of surface irradiance in the
mid-Atlantic in summer and spring, respectively.
The ratio of the mixed layer-averaged CDOM photobleaching
rate (PRMLD) to the surface CDOM photobleaching rate (PRsurf) in
the ocean may be approximated given mixed layer depth (MLD)
and the extinction coefficients of absorbed quanta, KAQ per season
or
Z
PRMLD =PRsurf ¼ ð1=MLDÞ eK AQ z dz
ð1Þ
where the integration is from the surface (z ¼0) to the respective
seasonal MLD. Typical summer and spring BATS values for MLD,
27 m and 89 m, and KAQ, of 0.040 m 1 and 0.055 m 1, respectively were used following Zafiriou et al. (2008). The calculated
PRMLD/PRsurf ratio for each season was 0.61 for summer, and 0.20
for spring. Applying this to a 48-hour exposure under the solar
simulator, the experimental incubations achieve a dosage roughly
equivalent to 16 and 57 day of circulation within the subtropical
mixed layer during summer and spring, respectively.
The approximate 2-day exposure period under the simulator
was chosen in order to resolve at least 50% reduction (i.e., one
half-life) of the initial absorption coefficient. As can be seen from
the above scaling exercise, this reduction in CDOM absorption
approaches the typical rates of decay of CDOM by natural sunlight
observed on time scales of weeks to months in the ocean
depending on latitude and season (Kouassi and Zika, 1992;
Vodacek et al., 1997; Nelson et al., 1998; Vähätalo and Wetzel,
2004). We can infer from the BATS calculation, for example,
that the photobleaching half-life of CDOM is 2 weeks or less
where annual insolation and residence time of surface waters is
high (e.g., low latitudes), and several months or more at high
latitudes where annual insolation is low and strong seasonal
mixing occurs. An assumption implicit in these estimations is that
the rate of daily turnover of the mixed layer occurs on time scales
shorter than that of photobleaching (i.e., the mixed layer is
well-mixed).
Duplicate sub-samples from irradiation time-course experiments were stored in the dark at 4 1C before spectroscopic
analysis. The possibility for low molecular weight chromophores
to react to form more complex molecules through polycondensation reactions (i.e., ‘‘dark recovery’’) in a dark environment
following irradiation has previously been a concern in photochemical experimentation (Gao and Zepp, 1998; Del Vecchio and
Blough, 2002). Two deliberate tests for such phenomena
conducted in this study found that no significant differences in
absorption properties were observed between final sub-samples
stored in the dark at 4 1C before analysis and replicate final
sub-samples that were analyzed immediately after light exposure. This finding was consistent with prior investigation by Del
Vecchio and Blough (2002), suggesting that any dark effects on
56
optical properties in filtered water samples were likely negligible
on the time scales of experimentation within the present study.
Our experimental design was characterized as an optically thin
system in which atotalL51, where atotal is the total absorption by
water and other constituents (particulate and dissolved), and L is
pathlength of the optical cell, which was 0.085 m (Hu et al., 2002).
By assuming particulate absorption is negligible within filtered
samples, and using a conservative (upper bound) pure water
absorption value at 300 nm (0.038 m 1; Morel et al., 2007) along
with the maximum CDOM absorption at 300 nm for the open
ocean sites sampled (0.402 m 1), we find that the largest
expected value of atotalL in our study is equal to 0.037. These
conditions permit the assumption that any ‘‘inner filter effects’’
(light attenuation occurring within the sample vial) were negligible (Hu et al., 2002).
2.3. Analytical approach
The surface photobleaching effect matrix, esurf(lo;li), was derived
to quantify the effect on CDOM absorption at observation wavelength,
lo, by irradiation wavelength, li, as influenced by the combined
energy of absorbed photons during full-sun exposure. The rate of
CDOM photobleaching at any given observation wavelength for an
optically thin water sample, daCDOM(lo)/dt (m 1 s 1), is then defined
as the product of esurf(lo;li) (m 1 mEin 1), the quantum scalar
irradiance, Eo(li) (mEin m 2 s 1 nm 1), the absorption by CDOM
incrementally averaged over the duration of exposure, āCDOM (li)
(m 1), and the volume of the optical cell used, V (m3):
Z
esurf ðlo ; li ÞEo ðli Þa CDOM ðli ÞV dli
ð2Þ
daCDOM ðlo Þ=dt ¼
where the integration is over all irradiation wavelengths, li (300–
700 nm).
Eo(li) is the output of the solar simulator quantified using a
multispectral radiometer as described in Section 2.2. The absorbance at each time interval, used to calculate the rate of change in
CDOM absorption, daCDOM(lo)/dt, over the exposure period, was
determined using an UltraPathTM liquid core waveguide scanning
spectrophotometer (World Precision Instruments, Sarasota, FL)
with a 1.943 m optical pathlength capable of detecting relatively
low levels of absorption in the open ocean. Absorption spectra
were analyzed over 300–700 nm. The protocol for measurement
of CDOM absorption using UltraPathTM, including time-dependent baseline drift and salinity-dependent refractive index
correction functions, has been previously described (Nelson
et al., 2007; Swan et al., 2009).
Given the above inputs, values of esurf(lo;li) were determined
from Eq. (2) by assuming a Rayleigh-like distribution function of
the form:
ðli lref Þ2
l l
ð3Þ
esurf ðlo ; li Þ ¼ Aðlo Þ i ref2 e 2Bðlo Þ2
Bðlo Þ
where lref is a reference wavelength (300 nm), and A(lo)
(m 1 mEin 1 nm) and B(lo) (nm) are constants that are inversely
solved using an unconstrained non-linear optimization procedure
(fminsearch function in Matlabs), which minimized the mean
absolute value difference between measured and modeled
daCDOM(lo)/dt. B(lo) affects the spread and skewness of the
irradiation wavelength dependency of esurf(lo;li) at each observation wavelength, and works in combination with A(lo) as a scaling
factor within Eq. (3) in setting the magnitude of the photochemical effect at the observation wavelength. A(lo) and B(lo) are
inherent properties of the CDOM at each sample site, and are
applicable in circumstances where the spectral light field remains
proportional to that used in the current study. As changes in
CDOM absorption and spectral slope are dependent on the
spectral quality of incident irradiance, our results therefore only
strictly apply to the surface of the open ocean (Osburn et al.,
2001; Tzortziou et al., 2007).
A schematic of the inverse method is provided in Fig. 3 (panels
A–F) for the case of Sargasso Sea water irradiation. The solar
simulator spectrum in Fig. 3A represents Eo(li). Fig. 3B displays the
CDOM absorption spectrum at ten evenly spaced (8-hour interval)
time points during the irradiation experiment, from which estimates of daCDOM(lo)/dt were calculated (Fig. 3C). The value of
daCDOM(lo)/dt in Fig. 3C is negative, indicating photobleaching,
and decreases in time as the bleaching of CDOM absorption results
in less overall absorbed quanta available for photochemical work.
The average spectral absorption by CDOM during each
experiment, āCDOM ðli Þ, was computed for each of the ten evenly
spaced time intervals (dt) over the course of irradiation, and
multiplied by Eo(li) and the sample cell volume (v) to calculate
the effective dose of absorbed quanta for each sample, Qa(li)
(mEin s 1 nm 1; Fig. 3D), which was used to solve for esurf(lo;li)
(see Fig. 3E for lo ¼325 nm). The modeled esurf (325;li) spectrum
for the Sargasso Sea sample (Fig. 3E), when multiplied by the
Qa(li) spectrum, yields the action spectrum for surface CDOM
photobleaching at the 325 nm observation wavelength (Fig. 3F).
The modeled values for daCDOM(lo)/dt at each time point were
computed as the area underneath (integrand of) the action
spectrum curve in Fig. 3F.
The action spectrum, which illustrates that maximum photobleaching of aCDOM(325) in the Sargasso Sea sample occurs at the
coincident wavelength (where lo ¼ li), demonstrates that the current analytical method is consistent with recent findings on the
kinetics of spectral CDOM photobleaching in natural waters
(Del Vecchio and Blough, 2002; Goldstone et al., 2004). Several
other functions for esurf(lo;li) were tested for their ability to
reproduce the laboratory results, including linear, exponential, and
Gaussian-like shapes. Linear and exponential forms for esurf(lo;li)
each implied that the efficiency of photobleaching decreases monotonically with increasing wavelength across the full-spectrum of
irradiation regardless of observation wavelength. While shorter
wavelengths indeed are more energetic, these models do not
adequately describe the interactive effects of chromophores undergoing solar bleaching during broadband exposure (Del Vecchio and
Blough, 2002). A Gaussian-like model form for esurf(lo;li), which
describes maximum photobleaching near the incident wavelength
with an effect that tapers off at flanking regions of the spectrum, is
more consistent with use of polychromatic irradiation to investigate
photobleaching (Tzortziou et al., 2007); however, we observed
unrealistically narrow spread (5–10 nm) in esurf(lo;li) using a
Gaussian shape and significant ‘‘off-axis’’ bleaching effects were
not well modeled. Use of the Rayleigh-like function (Eq. (3))
provided a better approximation to the weighted cumulative
photochemical effect of absorbed energy outside the incident
wavelength, and proved to be the best empirical fit to the data.
Use of this model form yielded superior matchups between measured and modeled daCDOM(lo)/dt values (r2 498%) compared to the
other functions tried.
The exponential spectral slope parameter, S (nm 1), plotted at
each time point (inset of Fig. 3B) was determined for assessing
shifts in spectral quality of CDOM. S was calculated using a
non-linear exponential curve fit to the spectral region 300–340 nm.
The wavelength region selected for computing spectral slope, while
internally consistent, does not facilitate comparison with literature
values of spectral slope computed across broader wavebands (Nelson
et al., 2007), or across longer or shorter spectral regions (Twardowski
et al., 2004; Loiselle et al., 2009). The choice of spectral range was
made because of anomalous irradiation-induced changes in the
visible wavelength region that did not fit an exponential curve
(described below).
3. Results
The change in the CDOM absorption spectrum for the case of
irradiated Sargasso Sea water presented in Section 2.3 (Fig. 3B)
was characterized by reduced absorption between 300 and
450 nm and increase in S (nm 1) (Fig. 3B inset), with no significant changes observed from 450 to 700 nm. The same observations were made for approximately half of our irradiated samples,
and were consistent with trends in CDOM photobleaching in
natural waters as documented by a number of previous studies
(e.g., Vodacek et al., 1997; Twardowski and Donaghay, 2002; Del
Vecchio and Blough, 2002; Vähätalo and Wetzel, 2004; Osburn
et al., 2009). Time courses of CDOM absorption during irradiation
of equatorial Indian (40 m) and subtropical North Pacific water
samples (40 m) are shown, respectively, along with insets of S
(nm 1) versus exposure time in Fig. 4A and B, as further examples
of these types of observations. Photobleaching in both the equatorial Indian and North Pacific samples was characterized by a
significant reduction in absorption ( 0.03 m 1, approximately
50% of the initial signal at 325 nm) and steepening of S, indicative
of greater proportional losses in absorption at the longer UV
observation wavelengths than at the shorter UV wavelengths.
For samples from the equatorial Pacific, subtropical South
Pacific, subarctic Pacific, subantarctic Pacific Frontal Zone, and
Southern Ocean, we observed decreases in CDOM absorption over
300–360 nm (and increase in S) with simultaneous increases in
57
CDOM absorption over 360–500 nm (Fig. 4C–F). The loss in
absorption over 300–360 nm was similar to that observed in
samples in which no long-wavelength absorption increases were
observed. Time courses of CDOM absorption (m 1) during irradiation, with insets of S (nm 1) versus exposure time, are
displayed for the case of subarctic Pacific (80 m) and equatorial
Pacific (100 m) waters in Fig. 4D and E, respectively, as examples
of these unusual spectral transformations.
Mean increase in absorption at 440 nm among samples exhibiting photoproduction was 0.038 m 1 over the experimental
time course and corresponds to a two-fold increase in the initial
absorption value. Three samples from the subantarctic Pacific in
which photoproduction was observed did not demonstrate any
measureable losses in absorption at wavelengths 300 nm or
longer. Fig. 4F demonstrates this phenomenon in a subantarctic
Pacific 80 m sample. Finally, the absorption by one sample from
the oligotrophic South Pacific (25 m) remained unaltered by the
48-hour long irradiation over the full (300–700 nm) observation
wavelength range (Fig. 4C).
Spectra of esurf(lo;li) at observation wavelengths of 300, 325
and 350 nm are displayed in Fig. 5A–C. The absolute magnitude of
the esurf spectrum among the study sites is greatest at the shortest
observation wavelength (300 nm) and decreases with wavelength
energy (Fig. 5A–C; note differences in scale of x and y axes),
consistent with prior polychromatic bleaching studies (Del
Vecchio and Blough, 2002). The esurf(350;li) parameter for samples
Fig. 4. Example irradiation time-courses of CDOM absorption spectra with insets of spectral slope parameter, S (m 1), versus exposure time (days): (A) equatorial Indian
40 m, (B) subtropical North Pacific 40 m, (C) subtropical South Pacific 25 m, (D) subarctic Pacific 80 m, (E) equatorial Pacific 100 m, and (F) subantarctic Pacific 80 m waters.
Fig. 5. Surface photobleaching effect parameter, esurf (lo;li) (m 1 mEin 1), versus irradiation wavelength, li (nm), at observation wavelengths (A) 300 nm (B) 325 nm and
(C) 350 nm (Note: scales for A–C differ).
58
from the subantarctic frontal zone (5 and 80 m) and Southern
Ocean (50 m) have a positive value in Fig. 5C due to photoproduction at the 350 nm observation wavelength. However as photobleaching rate was effectively zero at the 300 and 325 nm
observation wavelengths (not resolvable over the exposure
period), esurf spectra for these few samples are not visible in
Fig. 5A and B. Fig. 5A–C also shows a trend of esurf occurring over a
greater bandwidth as observation wavelength increases. The
esurf(325;li) spectrum approached zero at approximately 375 nm
(Fig. 5B), indicating that photochemical reactivity of CDOM is
negligible beyond this point in the spectrum, which agrees with
previous indications that energy in the range4400 nm has little
impact on photobleaching (Del Vecchio and Blough, 2002). On
average, the ‘largest’ (i.e., most negative) esurf values at a given UV
observation wavelength were observed in the North Pacific and
equatorial Indian Ocean samples, while the shallow South Pacific
samples (subtropical, subantarctic and Southern Ocean) consistently exhibited the lowest (i.e., least negative) values.
Examples of the retrieved values of A(lo, m 1 mEin 1 nm)
(Fig. 6A) and B(lo, nm) (Fig. 6B) from the Sargasso Sea and
equatorial Pacific, plotted as a function of observation wavelength, show the spectral dependencies of the parameters that
describe esurf(lo;li). (For the purpose of illustration, A and B are
only plotted for lo ¼350–400 nm in Fig. 6, as A values at lo ¼
300–340 nm were several orders of magnitude greater.) A(lo)
values were typically negative, reflecting loss in absorption at the
observation wavelength, except across lo ¼350–440 nm in some
samples where A(lo) acquired a positive value indicating CDOM
increases during irradiation. Fig. 6A shows the progression of
A(lo) in the equatorial Pacific 100 m sample from a negative value
at 350 nm (i.e., photobleaching), to a value of zero at 360 nm (no
change), to positive values at wavelengthsZ370 nm, corresponding to the increases in absorption spectrum of this sample as seen
in Fig. 4E. This is contrasted with the negative value of the A(lo)
spectrum for the Sargasso Sea 80 m sample, in which only loss in
absorption across the spectrum was observed during irradiation.
Regardless of whether CDOM absorption increased or
decreased during irradiation, B(lo) values were of similar magnitude among the samples (e.g., Fig. 6B). The general trend in the
B(lo) spectrum for a given sample is a decrease in magnitude as
observation wavelength increases. In samples where A(lo) had a
value of zero indicating no net change in absorption at the
particular observation wavelength, the corresponding B(lo) is
irrelevant (not a number) due to its position as a denominator
within Eq. (3). This is evident at 360 nm for the B(lo) spectrum of
the equatorial Pacific 200 m sample displayed in Fig. 6B.
Table 1 displays hydrographic information for each open ocean
sample evaluated, and provides the esurf(lo;li) (m 1 mEin 1)
value determined for coincident wavelengths of 310 nm,
325 nm, 350 nm and 440 nm.
A primary objective of our investigation was to understand the
environmental controls on the susceptibility of CDOM to photobleaching making use of the associated water sample data
collected as part of the core measurements of the U.S. CO2/CLIVAR
Repeat Hydrography Program. Data included depth (Z), salinity,
temperature, dissolved oxygen (O2), nitrate (N), phosphate (P),
silicate (Si), and fluorometric chlorophyll-a (Chl-a) concentrations. All measurements are made following standard WOCE
methods (http://ushydro.ucsd.edu). The N:P ratio, as well as the
initial CDOM absorption coefficient (aCDOM(lo), m 1) and spectral
slope (S, nm 1) of the samples were also evaluated for their role
in explaining the observed photoproduction of CDOM and the
natural variability in the surface photobleaching effect parameter.
Values of A(lo) and B(lo) were not significantly linearly correlated
with any of the hydrographic data, thus we conducted simple
regressions of esurf(lo;li) with the hydrographic data mentioned
above. All samples presented in Table 1 were included in the
analysis, including those in which esurf(lo;li) was effectively zero.
Linear regression statistics are summarized within Table 2 for
several coincident wavelengths across 310–440 nm. The value of
esurf(440;440) had significant positive linear relationships with
nitrate (r2 ¼0.35, p-value¼0.02, n ¼15) and phosphate concentrations (r2 ¼0.40, p-value¼0.01, n¼15), and in particular the N:P
ratio (r2 ¼0.70, p-value¼0.0001, n¼ 15), and was significantly
negatively correlated to temperature (r2 ¼0.45, p-value¼0.006,
n¼15). Correspondingly, esurf(350;350) and esurf(375;375) also
had significant, albeit weaker, relationships with N:P and
temperature (Table 2), as photoproduction was also observed at
these wavelengths among several samples.
Values of esurf(440;440) assumed a rough bimodal distribution
when plotted against N, P, and N:P, suggesting that sign ( þ or )
rather than the magnitude of esurf(440;440) signals the presence
or absence of photoproduction at 440 nm. Samples in which
decreases in CDOM absorption at 440 nm were observed (i.e.,
subtropical sites and the equatorial Indian Ocean) had N:P values
of less than 2.3, while N:P ratios were higher (10.7–15.2) hence
closer to the global mean (Redfield) value of 16 among samples in
which photoproduction was observed (Table 1).
Fig. 6. Parameters A(lo) and B(lo) describing the surface photobleaching effect (Eq. (3) in text) versus observation wavelength (lo) for (A) Sargasso Sea 80 m water and
(B) Equatorial Pacific 100 m water. Negative values of A(lo) indicated photobleaching at the observation wavelength, while positive values indicated an increase in CDOM
absorption.
59
Table 2
Regression statistics for linear correlation of surface photobleaching parameters at selected wavelengths, esurf(lo;li), with initial values of environmental properties.
Z
Salinity
Temp.
O2
N
P
Si
Chl-a
N:P
Initial aCDOM(ko)
Initial S
r2 ¼0.06
p ¼0.38
n¼ 15
r2 ¼0.03
p ¼0.55
n¼ 15
r2 ¼0.20
p ¼0.11
n¼ 14
r2 ¼ 0.01
p ¼ 0.79
n ¼15
r2 ¼ 0.01
p¼ 0.70
n¼15
r2 ¼ 0.04
p¼ 0.47
n¼15
r2 ¼0.09
p ¼0.28
n¼ 15
r2 ¼ 0.17
p ¼ 0.13
n ¼15
r2 ¼ 0.27
p ¼ 0.05
n ¼15
r2 ¼ 0.52
p ¼ 0.003n
n¼ 15
r2 ¼0.002
p ¼0.89
n¼ 15
r2 ¼0.17
p ¼0.13
n¼ 15
r2 ¼0.26
p ¼0.06
n¼ 14
r2 ¼ 0.04
p ¼ 0.50
n ¼15
r2 ¼ 0.06
p¼ 0.39
n¼15
r2 ¼ 0.01
p¼ 0.80
n¼15
r2 ¼0.08
p ¼0.30
n¼ 15
r2 ¼ 0.22
p ¼ 0.08
n ¼15
r2 ¼ 0.18
p ¼ 0.12
n ¼15
r2 ¼ 0.37
p ¼ 0.02n
n¼ 15
r2 ¼0.003
p ¼0.84
n¼ 15
r2 ¼0.27
p ¼0.05
n¼ 15
r2 ¼0.14
p ¼0.18
n¼ 14
r2 ¼ 0.11
p ¼ 0.24
n ¼15
r2 ¼ 0.13
p¼ 0.19
n¼15
r2 o 0.001
p¼ 0.96
n¼15
r2 ¼0.04
p ¼0.49
n¼ 15
r2 ¼ 0.34
p ¼ 0.02n
n ¼15
r2 ¼ 0.12
p ¼ 0.21
n ¼15
r2 ¼ 0.27
p ¼ 0.04n
n¼ 15
r2 ¼0.001
p ¼0.90
n¼ 15
r2 ¼0.35
p ¼0.02n
n¼ 15
r2 ¼0.05
p ¼0.41
n¼ 14
r2 ¼ 0.25
p ¼ 0.06
n ¼15
r2 ¼ 0.28
p¼ 0.04n
n¼15
r2 ¼ 0.04
p¼ 0.47
n¼15
r2 ¼0.14
p ¼0.16
n¼ 15
r2 ¼ 0.52
p ¼ 0.003n
n ¼15
r2 ¼ 0.09
p ¼ 0.28
n ¼15
r2 ¼ 0.23
p ¼ 0.07
n¼ 15
r2 ¼0.005
p ¼0.79
n¼ 15
r2 ¼0.34
p ¼0.02n
n¼ 15
r2 ¼0.09
p ¼0.30
n¼ 14
r2 ¼ 0.19
p ¼ 0.10
n ¼15
r2 ¼ 0.23
p¼ 0.07
n¼15
r2 ¼ 0.004
p¼ 0.80
n¼15
r2 ¼0.13
p ¼0.18
n¼ 15
r2 ¼ 0.55
p ¼ 0.001n
n ¼15
r2 ¼ 0.07
p ¼ 0.33
n ¼15
r2 ¼ 0.15
p ¼ 0.14
n¼ 15
r2 ¼0.04
p ¼0.48
n¼ 15
r2 ¼0.45
p ¼0.006n
n¼ 15
r2 ¼0.02
p ¼0.66
n¼ 14
r2 ¼ 0.35
p ¼ 0.02n
n ¼15
r2 ¼ 0.40
p¼ 0.01n
n¼15
r2 ¼ 0.06
p¼ 0.36
n¼15
r2 ¼0.13
p ¼0.19
n¼ 15
r2 ¼ 0.70
p ¼ 0.0001n
n ¼15
r2 ¼ 0.003
p ¼ 0.84
n ¼15
r2 ¼ 0.07
p ¼ 0.34
n¼ 15
esurf(310;310)
r2 ¼ 0.03
p ¼ 0.55
n ¼15
esurf(325;325)
r2 ¼ 0.009
p ¼ 0.74
n ¼15
esurf(340;340)
r2 o 0.001
p ¼ 0.93
n ¼15
esurf(350;350)
r2 ¼ 0.006
p ¼ 0.79
n ¼15
esurf(375;375)
r2 ¼ 0.01
p ¼ 0.67
n ¼15
esurf(440;440)
r2 ¼ 0.05
p ¼ 0.44
n ¼15
n
A p-value less than 0.05 indicated that the corresponding r2 value was significant at the 95% confidence interval.
The trend in esurf(440;440), and to a lesser extent, esurf(350;350)
and esurf(375;375), versus temperature roughly indicated that positive values of esurf(440;440) occurred at temperatures below 15 1C,
and negative values above 15 1C.
No significant relationships were observed between esurf(lo;li)
and either the initial CDOM absorption coefficient, chlorophyll-a,
dissolved O2, Si, salinity or sample depth (Table 2). While
temperature was correlated with the absence or presence of
photoproduction as indicated by esurf at coincident wavelengths
of 350–440 nm, no significant relationship was observed between
temperature and esurf(lo;li) at the shorter UV wavelengths
(310–340 nm) over which photobleaching occurred in most
samples. Across this wavelength range, the initial spectral slope
of the samples mainly accounted for variability in esurf(lo;li), with
N:P also contributing in explaining variability in esurf(340;340)
(Table 2).
4. Discussion
4.1. Implications of CDOM photoproduction in the open ocean
The evolution of a pronounced peak in absorption spanning
350–500 nm as observed in half of our irradiated open ocean
samples (Fig. 4D–F) has potential to affect remote-sensing estimates of chlorophyll due to the wavelength region of maximum
absorption by chlorophyll overlapping with that of the photoproduct(s). However, open ocean CDOM absorption spectra are
typically featureless, and peaks of this magnitude have not been
observed in the field (e.g., Nelson et al., 1998, 2007; Yamashita
and Tanoue, 2009; Swan et al., 2009). However, minimally
detectable ‘‘bump’’-like features in the CDOM spectrum between
410 and 420 nm have been reported in southeast Pacific upwelling zones (Bricaud et al., 2010, their Fig. 16). This raises the
possibility that the photoproduct may be transient in nature and
decrease with further exposure.
An irradiation time course of an equatorial Pacific 200 m
sample was extended to 7 day of continuous simulated solar
irradiation (equivalent to roughly 35 day of mid-latitude irradiance at the sea surface in summer and a few months of exposure
in an equatorial Pacific mixed-layer). Fig. 7A displays the spectral
transformations (300–550 nm) of the equatorial Pacific 200 m
sample during irradiation with an inset of S (nm 1) versus
exposure time, and Fig. 7B displays absorption at the 440 nm
observation wavelength versus exposure time. Evolution of a
photoproduct (between 360 and 500 nm) was observed by day
one, maximum increase in absorption at 440 nm was observed by
day two, and subsequent bleaching of the absorbing compound
and a return to the initial level was observed by the 6.8-day
mark (Fig. 7B). The negligible net change in absorption over
360–500 nm by the 6.8-day mark indicated that photoproduction
of CDOM within filtered samples under simulated solar irradiance
does not proceed indefinitely or past the time scale of months in
nature. This is consistent with open ocean CDOM absorption
spectra (e.g., Nelson et al., 2007; Yamashita and Tanoue, 2009;
Swan et al., 2009; Bricaud et al., 2010) and leads us to postulate
that in the water column such a photoproduct is either
(a) intermediate and subject to near-simultaneous photodegradation, (b) biologically labile and subject to rapid microbial
consumption (Nelson et al., 2004), or (c) simply does not get
produced in appreciable concentration due to the diel cycle of
solar irradiance and mixing within the water column. In other
words, this phenomenon may be observed only within a
controlled system of accelerated light exposure such as that used
in the current study. For these reasons, significant net effects of
the observed photochemical phenomena on satellite estimations
of CDOM or chlorophyll are unlikely. Nevertheless, we present
several hypotheses on the potential implications for the observed
photoproducts as they pertain to photochemistry in the surface
ocean.
Reported instances of increased DOM absorption due to solar
irradiation within aquatic environments are rare, and have been
60
Fig. 7. (A) Seven-day irradiation time-course of CDOM absorption spectra with inset of S (nm 1) versus exposure time (days) and (B) absorption coefficient (m 1) of
CDOM at 440 nm versus exposure time (days) for equatorial Pacific 200 m waters.
primarily associated with photoreduction of iron (Küpper et al.,
2006; Martin et al., 2006) or photohumification of dissolved
substances (Kieber et al., 1997; Benner and Biddanda, 1998;
Obernosterer et al., 1999; Reche et al., 2001). Photohumification,
the light-stimulated polymerization of labile algal exudates such
as polyunsaturated fatty acids, has been invoked as one pathway
for the origin of marine humics (Kieber et al., 1997; Del Vecchio
and Blough, 2004). Photohumification is typically associated with
a long tail of absorption across the visible wavelengths, and has
been attributed to intramolecular charge or energy transfers
between proximate chromophores, as well as to long conjugated
and aromatic compounds in terrestrially influenced water
samples (Twardowski and Donaghay, 2002; Goldstone et al.,
2004; Del Vecchio and Blough, 2004). It is unlikely that the CDOM
photoproduction observed during our study is representative of
photohumification as filtered water samples from the open ocean
do not contain sufficient quantities of labile exudates or long fatty
acid chains for such reactions to take place (Whitehead and de
Mora, 2000). Furthermore, photoproduction herein was observed
as an absorption peak between 360 and 500 nm rather than the
long tail of absorption extending to the visible region beyond
500 nm, as is characteristic of photohumification (Reche et al.,
2001; Del Vecchio and Blough, 2002).
The correlations observed between esurf(440;440) and N:P
suggest a connection of CDOM photoproduction with the biogeochemical state of the water column as it influences CDOM
reactivity. The sampling regions in which we observed CDOM
photoproduction (e.g., equatorial Pacific, subarctic Pacific,
subantarctic Pacific and Southern Ocean) classify as iron-limited,
high-nutrient low-chlorophyll (HNLC) regions (Martin et al.,
1991; Behrenfeld and Kolber, 1999), corroborated by the
relatively high nitrate values observed in these regions
(6.32–42.12 mmol kg 1, Table 1). (The temperature correspondence of esurf(lo;li) at wavelengths at which photoproduction was
observed (350–440 nm) may therefore be an artifact of the
regionality of HNLC (not focused within subtropical regions)
rather than a photochemical phenomenon. This is supported by
the lack of correlation of temperature to instances of photobleaching at the shorter UV wavelengths, suggesting CDOM
photoreactivity is not strictly temperature-dependent as other
photoreactions in nature are (Toole et al., 2003; Mopper and
Kieber, 2002).)
Nitrate is active in surface ocean photochemistry by its
absorption of UVB radiation, which results in the production of
hydroxyl radicals (OH ), the dominant photosensitizers in aqueous solutions (Goldstone et al., 2002; Tedetti et al., 2007, 2008).
Toole et al. (2004) provide evidence for a light-mediated role of
nitrate in stimulating dimethylsulfide (DMS) photobleaching
rates. Action spectra for DMS photobleaching in the North
Atlantic reported by Toole et al. (2003) peak in the UVB region,
and contrast similar spectra made from the equatorial Pacific
which peak in both the UVB and 380–460 nm range (Kieber et al.,
1996). Through its absorption of UV light, CDOM is the primary
reactant controlling DMS photobleaching in the surface ocean, as
DMS does not absorb solar photons directly. It is therefore
possible that the HNLC-associated CDOM photoproduct between
360 and 500 nm in the equatorial Pacific may be responsible for
the peak in the action spectrum of DMS photobleaching observed
by Kieber et al. (1996). The hypothesis is that nitrate sensitizes
the production of a reactive CDOM molecule that absorbs light in
the 380–460 nm range, greatly increasing DMS photobleaching
rates. Although we have speculated that CDOM photoproducts
may be subject to near-instantaneous degradation in the water
column, the proposed photoreactions of nitrate, CDOM and DMS
likely proceed on comparably rapid time scales (Toole et al.,
2004).
In order to test for the role of nitrate in CDOM photoproduction as suggested by the correlations of esurf(440;440) with N, P
and the N:P ratio, an additional sample from the Santa Barbara
Channel, CA was collected in winter 2009 from the identical
geographic location and depth as indicated in Table 1. A sample
volume of 0.75 L was amended with 1.3 mL of 1000 mg L 1
NaNO3 stock solution (prepared using ultrapure water and
99.995þ % sodium nitrate, Sigma-Aldrich) to achieve a nitrate
concentration of 20.6 mmol kg 1, which approached the median
value of nitrate (19.1 mmol kg 1) amongst open ocean samples in
which photoproduction was observed. Replicate NO3 -amended
and unamended sub-samples were irradiated at in situ temperature along with dark controls for approximately 2 day following
the procedures outlined in Section 2.2. Initial and final absorption
spectra from this experiment are plotted in Fig. 8A. Addition of
sodium nitrate caused an initial increase in the absorption
coefficient for wavelengthso330 nm prior to irradiation, as seen
in the pre-irradiation difference spectrum (solid black line)
between the NO3 -amended and unamended samples (Fig. 8B).
This initial change in absorption may be attributed to the
collective absorption properties of sodium nitrate and CDOM in
the water sample as influenced by nitrate’s absorption properties
in the UVB region (e.g., Johnson and Coletti, 2002). At the end of
the 2-day exposure period, while net photoproduction was not
explicitly observed, less net absorption loss was observed
between approximately 340 and 500 nm in the NO3 -amended
61
Fig. 8. Absorption spectra of CDOM from the Santa Barbara Channel, CA amended with NO3 . (A) Solid black line ¼NO3 -amended CDOM before irradiation. Dotted black
line¼ NO3 -amended CDOM after 2-day irradiation. Solid gray line ¼unamended CDOM (control) before irradiation. Dotted gray line ¼unamended CDOM (control) after
2-day irradiation. (B) Difference spectra between NO3 -amended CDOM and unamended CDOM (control) before irradiation (solid black line), and after irradiation (dotted
black line).
sample than in the control sample. This wavelength region
corresponded to that of the absorption increases observed in
our high N samples. Fig. 8B (dotted line) displays the postirradiation difference spectrum between NO3 -amended and
unamended samples. We hypothesize that the significantly less
change in the absorption coefficient at 440 nm (0.0159 m 1)
post-irradiation in the NO3 -amended than control sample
suggests that underlying transient CDOM photoproduction may
be responsible for the less net overall photobleaching. While a
mechanism cannot be defined with certainty without further
study, these observations in conjunction with the elevated nitrate
levels observed in natural samples exhibiting photoproduction
imply a role for nitrate in CDOM photoproduction observed
herein.
It should be noted that the strong increase in absorption at
300 nm due to the experimental addition of nitrate may suggest
that nitrate drives much of the actual absorption of CDOM in the
UV region; however, this increase is likely an isolated incidence
pertaining to the artificial amendment with laboratory-grade
sodium nitrate. This is supported by the lack of correlation
between dissolved nitrate and CDOM absorption within a large
field data set from the North Atlantic Ocean (Nelson et al., 2007),
as well as the lack of correlation of nitrate with aCDOM at 300 nm
in the current study (r2 ¼0.17, p-value¼0.13, n ¼15).
Biers et al. (2007) observed simultaneous photochemical
formation of CDOM and fluorescent DOM upon irradiation of a
coastal seawater amended with N-containing tryptophan, raising
the possibility that this amino acid influenced photoproduction
within the current study, as tryptophan occurs in measureable
concentration within ocean samples (Coble, 2007). However,
Excitation Emission Matrix spectra (EEMs) analysis using a
FluoroMaxs-4 spectrofluorometer (HORIBA Jobin Yvon, Inc.) on
equatorial Pacific samples did not reveal evolution of a fluorescent photoproduct coincident with the observed CDOM formation
(not shown). In fact, only fading of fluorescence was observed in
these samples during irradiation, which is inconsistent with a role
for tryptophan in mediating the absorption increases we observed
(Biers et al., 2007).
There is another possibility that biologically mediated iron
ligands, such as marine siderophores involved in the photoreduction of Fe3 þ to Fe2 þ , may play a role alongside nitrate in the
observed photoproduction of CDOM (Barbeau et al., 2001; Martin
et al., 2006; Vraspir and Butler, 2009). Fe3 þ has been associated
with humic-type fluorescence intensity in the water column (Tani
et al., 2003; Kitayama et al., 2009), which itself is strongly related
to CDOM absorption (Coble, 2008; Yamashita and Tanoue, 2008).
For many of the samples taken from the CLIVAR cruises, dissolved
[Fe2 þ ] determinations were made; however, no correspondence
was found between dissolved [Fe2 þ ] concentrations and
esurf(440;440) (r2 ¼0.002, p-value¼0.97, n ¼11), thus identifying
a role for iron-chelating molecules in the observed CDOM photoreactions requires further experimentation.
4.2. Global variability in the surface photobleaching effect
Changes in CDOM at the UV observation wavelengths could
not be resolved within a 2-day exposure period among samples
collected from the subtropical South Pacific (25 m), subantarctic
Pacific frontal zone (5 m and 80 m), and Southern Ocean (50 m)
(see Table 1, e.g., esurf(325;325) values ¼0). These samples were
collected during austral summer in regions that experience
relatively high UV irradiation doses, partly due to stratospheric
ozone depletion (e.g., Smith et al., 1992). The South Pacific gyre is
characterized by the most optically clear waters in the global
ocean (Morel et al., 2007; Swan et al., 2009; Bricaud et al., 2010),
and the 25 m sample from this region had the lowest CDOM in
our study (Table 1). The low but detectable absorption in these
samples suggests that a photochemically resistant fraction of
CDOM persists past seasonal time scales in the open ocean and
potentially contributes to the ‘‘background’’ pool of CDOM in the
global ocean (Nelson et al., 2002, 2010).
The lack of correlation between depth and esurf at all
coincident wavelengths suggested that photochemically labile
CDOM may be found in deep or surface waters depending on
insolation and circulation dynamics of a region. Further, the lack
of correlation between esurf and the initial absorption coefficient
of CDOM suggests the quantity of CDOM does not have a strong
effect on susceptibility of CDOM to photobleaching at the surface.
Marine and freshwaters that are heavily influenced by terrestrially derived materials comprise a chromophore pool with a
highly variable chemical make-up and photochemical reactivity
(Coble, 2007). Even though localized influences of riverine input,
such as the Amazon and Orinoco River plumes, are detected in the
multi-annual global CDOM distribution, the chromophore pool in
the open ocean is largely remote from terrestrial effects on
seasonal time scales and chiefly driven by autochthonous in situ
processes (Siegel et al., 2002; Nelson et al., 2004, 2010). This leads
to an explanation for why the initial S and N:P values explain
62
much of the variance in the photobleaching effect parameter at
the shorter UV wavelengths.
Initial S values of all samples within this study, including those
from the Santa Barbara Channel, were in the range representative
of marine sources of CDOM (i.e., 4 0.02 nm 1, Nelson and Siegel,
2002; Bricaud et al., 2010). S values have been associated with
compositional variation in CDOM relating to source (e.g., terrestrial, marine) and to photobleaching (Twardowski and Donaghay,
2002; Weishaar et al., 2003; Helms et al., 2008). It is plausible that
chromophores in the open ocean have a relatively similar chemical character and spectral quality because CDOM is autochthonously produced and S variability in the open ocean is relatively
low (Nelson et al., 2010). High S values are generally indicative of
a chromophore pool that has had considerable light exposure
(Osburn et al., 2001; Twardowski and Donaghay, 2002; Loiselle
et al., 2009). This is further supported by the increases in S
observed during irradiation experiments, as well as by our
observation that the photochemically resistant subtropical South
Pacific 25 m sample had the highest initial S value in this study.
The positive linear correlation between S and esurf (325;325)
suggests that higher spectral slopes lead to a weaker (less
negative) surface photobleaching effect. As an indicator of the
biogeochemical history of a water parcel and time since last
surface contact, S explains a mechanism for the photochemical
susceptibility of CDOM. In the open ocean, CDOM with low S
values relative to surface waters (such as in deep aphotic waters)
is either relatively newly produced or has accumulated over
decades as a result of deep ocean remineralization processes
(Rochelle-Newall and Fisher, 2002; Nelson et al., 2004, 2010;
Kitidis et al., 2006). In either case, the lower S values reflect a pool
of CDOM that has remained relatively unexposed to sunlight,
which may confer a higher probability of photochemically labile
chromophores. On average, the strongest photobleaching effect
was observed in the North Pacific and Indian Ocean samples,
while the weakest photobleaching effect was observed among
shallow South Pacific samples (subtropical, subantarctic and
Southern Ocean). Accordingly, North Pacific and Indian Ocean
samples had lower initial S values on average than the shallow
South Pacific samples. Regions of the North Pacific and equatorial
Indian Oceans experience greater degrees of mixed layer renewal
than the South Pacific, attributed in part to North Pacific Intermediate Water (NPIW) circulation (You et al., 2003) and equatorial Indian upwelling patterns (Xie et al., 2002), respectively.
CDOM spectral slopes in the mixed layers of the North Pacific
and equatorial Indian are lower due to renewal with less photoexposed (and higher AOU) subthermocline waters (Swan et al.,
2009; Nelson et al., 2010).
Elucidating a mechanism for the effect of N:P on CDOM
photobleaching however, as suggested by the positive linear
relation between N:P and esurf(325;325) for example, poses more
of a challenge as nutrient ratios, particularly N:P, vary widely in
open ocean regimes due to many processes including denitrification, nitrogen fixation, and nutrient or iron limitation (Klausmeier
et al., 2004; Falkowski et al., 1998). The N:P ratio of the
surrounding water column has been shown to regulate phytoplankton species selection (Quigg et al., 2003). One hypothesis
that follows is that global variability in N:P leads to variation in
phytoplankton assemblages, which are subsequently decomposed
by microbes producing CDOM of variable chemical composition
thus photochemical susceptibility. On the other hand since we
demonstrated earlier an example of a light-stimulated role for
nitrate in CDOM photoreactions, we cannot rule out the hypothesis that N:P, regardless of its relevance to the biogeochemical
state of the water column, affects CDOM photobleaching directly
along a photochemical pathway. It is possible, for example, that
varying levels of reactive oxygen species generated during UV
light absorption by nitrate and phosphate play a role in determining CDOM photoreactivity as N:P was naturally highly correlated
to N within samples from our study (r2 ¼0.73, p ¼0.0001, n¼15)
(Mopper and Kieber, 2002; Goldstone et al., 2002). On the other
hand, esurf is more highly correlated with N:P than N, which lends
stronger support to the idea that N:P may affect CDOM quality
and thus photoreactivity through a biogeochemical pathway.
Although the photobiogeochemical parameters described
above account for a significant portion of the variance in esurf
within the spectral ranges examined and allow us to generate
several hypotheses as to the drivers of open ocean CDOM
photobleaching, a unified explanation for the effect of N:P on
CDOM photobleaching may not be possible yet given the small
sample size of our study. The overall interpretation of the results
is that both photo-oxidative and biogeochemical conditions are
important synergistic drivers of the susceptibility of marine
CDOM to photochemical transformations.
5. Conclusion
We have modeled the effect of surface irradiance on CDOM
absorption as a function of the combined energetic effect of all
UV–vis wavelengths. A methodology and hypotheses relating to
the susceptibility of CDOM to photobleaching in the major ocean
basins was described, further characterizing and quantifying
aspects of this important removal pathway for CDOM in the open
ocean. We conclude that the N:P ratio and spectral slope of CDOM
as proxies for the biogeochemical and solar exposure history of a
water sample, respectively, appear to be the primary influence on
the susceptibility of CDOM to solar bleaching in the open ocean.
Overall, our results suggest that the quality of CDOM is a more
important regulator of CDOM photoreactivity than quantity.
The course of experimentation revealed unusual transient
photoproduction in the visible region of the CDOM absorption
spectrum within HNLC regimes, which we attributed to nitrate
photochemistry and proposed as a causal mechanism for the
discrepant action spectra for DMS photobleaching between equatorial Pacific and North Atlantic waters (Kieber et al., 1996; Toole
et al., 2003). Further investigation is needed to determine if this
phenomenon is exceptional to high-intensity laboratory exposures, but the potential for transient light-induced CDOM
increases in the open ocean may not be ruled out.
Future application of esurf rests on the assumption that that
there is no significant alternate sink in the open ocean (e.g.,
diagenetic, microbial) for chromophores across the time scales of
photobleaching rate assessments as they correspond to those in
nature (seasonal to interannual). This is a valid assumption as
slight trends in open ocean CDOM spectral characteristics attributed to diagenesis thus far only have been observed on temporal
scales of thermohaline circulation (decadal to millennial) in deep
water masses (Nelson et al., 2007, 2010). Secondly, microbial
consumption is considered insignificant in the open ocean on the
time scale of CDOM photobleaching, with the overall net activity
of bacterioplankton on CDOM in the water column resulting in
production of CDOM (Rochelle-Newall and Fisher, 2002; Vähätalo
and Wetzel, 2004; Nelson et al., 2007, 2010).
There are clear limitations to the current method, not least of
which is that the data may only be applied at the sea surface and
on a regional basis given that biogeochemical parameters
contribute to regional variability in esurf. However, ocean colorderived monthly climatology of CDOM and irradiance may be
used in conjunction with esurf spectra to ultimately generate a
synoptic view of photobleaching rate in the surface ocean. Surface
rates then could be incorporated using a method similar to Eq. (1)
to generate a mixed-layer integrated photobleaching rate, as
previous investigations suggest that attenuation of downwelling
UV irradiance roughly imitates the attenuation of photodegradation of organic material (Vähätalo et al., 2000; Vähätalo and
Wetzel, 2004). Combining estimates of specific turnover rate of
CDOM in the mixed layer due to photobleaching with simple
vertical mixing rates will ultimately enable calculation of the net
in situ biological production rate of CDOM in the ocean, a process
that has been difficult to parameterize in laboratory and field
settings. Constraining the open ocean CDOM cycle in this manner
will be a valuable contribution to future remote-sensing and
biogeochemical studies.
Acknowledgments
Support from NSF Chemical Oceanography and NASA Ocean
Biology and Biogeochemistry Programs to N. Nelson, D. Siegel and
C. Carlson, as well as the NASA Earth System Science Graduate
Fellowship Program to C. Swan, is gratefully acknowledged. We
thank the CO2/CLIVAR Repeat Hydrography Program chief scientists, the captains and crew of the R/Vs Revelle, Thompson and
Brown, as well as C. Carlson, D. Menzies, N. Guillocheau,
M. Meyers, E. Wallner (UCSB), N. MacDonald (Bermuda-BIOS),
and the U. Hawaii/FSU trace metal posse (W. Landing, C. Measures, C. Buck, M. Brown, W. Hiscock, P. Hansard, M. Hatta,
M. Grand) for field assistance and collection, and to Barbara
Prezelin for the loan of the BSI scalar PAR meter. The satellite
CDM imagery shown in Fig. 1 were obtained from the Ocean Color
MEaSUREs Project at UCSB (http://wiki.icess.ucsb.edu/measures/
Products). Finally, we thank M. Brzezinski, C. Carlson and several
anonymous reviewers for insightful comments on the manuscript.
References
Barbeau, K., Rue, E.L., Bruland, K.W., Butler, A., 2001. Photochemical cycling of iron
in the surface ocean mediated by microbial iron(iii)-binding ligands. Nature
413, 409–413.
Behrenfeld, M.J., Kolber, Z.S., 1999. Widespread iron limitation of phytoplankton in
the South Pacific ocean. Science 283, 840–843.
Benner, R., Biddanda, B., 1998. Photochemical transformations of surface and deep
marine dissolved organic matter: effects on bacterial growth. Limnol. Oceanogr. 43, 1373–1378.
Biers, E.J., Zepp, R.G., Moran, M.A., 2007. The role of nitrogen in chromophoric and
fluorescent dissolved organic matter formation. Mar. Chem. 103, 46–60.
Blough, N.V., Zafiriou, O.C., Bonilla, J., 1993. Optical absorption spectra of waters
from the Orinoco River outflow: Input of colored organic matter to the
Caribbean. J. Geophys. Res. 98, 2245–2257.
Bricaud, A., Babin, M., Claustre, H., Ras, J., Tieche, F., 2010. Light absorption
properties and absorption budget of Southeast Pacific waters. J. Geophys. Res.,
115. doi:10.1029/2009JC005517.
Coble, P.G., 2007. Marine optical biogeochemistry: the chemistry of ocean color.
Chem. Rev. 107, 402–418.
Coble, P., 2008. Cycling coloured carbon. Nat. Geosci. 1, 575–576.
Del Castillo, C.E., Coble, P.G., 2000. Seasonal variability of the colored dissolved
organic matter during the 1994–95 NE and SW Monsoons in the Arabian Sea.
Deep-Sea Res. I 47, 1563–1579.
Del Vecchio, R., Blough, N.V., 2002. Photobleaching of chromophoric dissolved
organic matter in natural waters: kinetics and modeling. Mar. Chem. 78,
231–253.
Del Vecchio, R., Blough, N.V., 2004. On the origin of the optical properties of humic
substances. Environ. Sci. Technol. 38, 3885–3891.
Falkowski, P.G., Barber, R.T., Smetacek, V., 1998. Biogeochemical controls and
feedbacks on ocean primary production. Science 281, 200–206.
Feely, R.A., Talley, L.D., Johnson, G.C., Sabine, C.L., Wanninkhof, R., 2005. Repeat
hydrography cruises reveal chemical changes in the North Atlantic. Eos Trans.
AGU 86, 399. doi:10.1029/2005EO420003.
Gao, H.Z., Zepp, R.G., 1998. Factors influencing photoreactions of dissolved organic
matter in a coastal river of the southern United States. Environ. Sci. Technol.
32, 2940–2946.
Goldstone, J.V., Del Vecchio, R., Blough, N.V., Voelker, B.M., 2004. A multicomponent model of chromophoric dissolved organic matter photobleaching. Photochem. Photobiol. 80, 52–60.
Goldstone, J.V., Pullin, M.J., Bertilsson, S., Voelker, B., 2002. Reactions of hydroxyl
radical with humic substances: bleaching, mineralization and production of
bioavailable carbon substrates. Environ. Sci. Technol. 36, 364–372.
63
Hedges, J.I., Keil, R., Benner, R., 1997. What happens to terrestrially-derived
organic matter in the sea? Org. Geochem. 27, 195–212.
Helms, J.R., Stubbins, A., Ritchie, J.D., Minor, E.C., Kieber, D.J., Mopper, K., 2008.
Absorption spectral slopes and slope ratios as indicators of molecular weight,
source and photobleaching of chromophoric dissolved organic matter. Limnol.
Oceanogr. 53, 955–969.
Hu, C., Muller-Karger, F.E., Zepp, R.G., 2002. Absorbance, absorption coefficient,
and apparent quantum yield: a comment on common ambiguity in the use of
these optical concepts. Limnol. Oceanogr. 47, 1261–1267.
Johannessen, S.C., Miller, W.L., 2001. Quantum yield for the photochemical
production of dissolved inorganic carbon in seawater. Mar. Chem. 76,
271–283.
Johannessen, S.C., Miller, W.L., Cullen, J.J., 2003. Calculation of UV attenuation and
colored dissolved organic matter absorption spectra from measurements of
ocean color. J. Geophys. Res., 108. doi:10.1029/2000JC000514.
Johnson, K.S., Coletti, L.J., 2002. In situ ultraviolet spectrophotometry for high
resolution and long-term monitoring of nitrate, bromide and bisulfide in the
ocean. Deep-Sea Res. I 49, 1291–1305.
Kieber, D.J., Jiao, J., Kiene, R.P., Bates, T.S., 1996. Impact of dimethylsulfide
photochemistry on methyl sulfur cycling in the equatorial Pacific Ocean. J.
Geophys. Res. 101, 3715–3722.
Kieber, R., Hydro, L., Seaton, P., 1997. Photooxidation of triglycerides and fatty
acids in seawater: implication toward the formation of marine humic
substances. Limnol. Oceanogr. 42, 1454–1462.
Kitayama, S., Kuma, K., Manabe, E., Sugie, K., 2009. Controls on the iron distribution in the deep-water column of the North Pacific Ocean: iron (III) hydroxide
solubility and marine humic-type dissolved organic matter. J. Geophys. Res.,
114. doi:10.1029/2008JC004754.
Kitidis, V., Stubbins, A.P., Uher, G., Upstill-Goddard, R.C., Law, C.S., Woodward,
E.M.S., 2006. Variability of chromophoric dissolved organic matter in surface
waters of the Atlantic Ocean. Deep-Sea Res. II 53, 1666–1684.
Klausmeier, C.A., Litchman, E., Daufresne, T., Levin, S.A., 2004. Optimal nitrogen-tophosphorus stoichiometry of phytoplankton. Nature 429, 171–174.
Kouassi, A.M., Zika, R.G., 1992. Light-induced destruction of the absorbance
property of dissolved organic matter in seawater. Toxicol. Environ. Chem.
35, 195–211.
Küpper, F.C., Carrano, C.J., Kuhn, J., Butler, A., 2006. Photoreactivity of iron(III)–
aerobactin: photoproduct structure and iron(III) coordination. Inorg. Chem. 45,
6028–6033.
Loiselle, S.A., Bracchini, L., Dattilo, A.M., Ricci, M., Tognazzi, A., Cozar, A., Rossi, C.,
2009. Optical characterization of chromophoric dissolved organic matter using
wavelength distribution of absorption spectral slope. Limnol. Oceanogr. 54,
590–597.
Maritorena, S., Hembise Fanton d’Andon, O., Mangin, A., Siegel, D.A., 2010. Merged
satellite ocean color data products using a bio-optical model: characteristics,
benefits and issues. Remote Sens. Environ. 114, 1791.
Martin, J.H., Gordon, R.M., Fitzwater, S.E., 1991. The case for iron. Limnol.
Oceanogr. 36, 1793–1802.
Martin, J.D., Ito, Y., Homann, V.V., Haygood, M.G., Butler, A., 2006. Structure and
membrane affinity of new amphiphilic siderophores produced by Ochrobactrum sp. SP18. J. Biol. Inorg. Chem. 11, 633–641.
Miller, W.L., Moran, M.A., Sheldon, W.M., Zepp, R.G., Opsahl, S., 2002. Determination of apparent quantum yield spectra for the formation of biologically labile
photoproducts. Limnol. Oceanogr. 47, 343–352.
Mobley, C.D., 1994. Light and Water: Radiative Transfer in Natural Waters.
Academic Press, San Diego, CA. (pp. 7–11).
Mopper, K., Kieber, D.J., 2002. Photochemistry and the cycling of carbon, sulfur,
nitrogen and phosphorus. In: Matter, D.A., Hansell, Carlson, C.A. (Eds.),
Biogeochemistry of Marine Dissolved Organic. Academic Press, San Diego, CA,
pp. 456–508.
Moran, M.A., Sheldon, W.M., Zepp, R.G., 2000. Carbon loss and optical property
changes during long-term photochemical and biological degradation of
estuarine dissolved organic matter. Limnol. Oceanogr. 45, 1254–1264.
Morel, A., Gentili, B., Claustre, H., Babin, M., Bricaud, A., Ras, J., Tie che, F., 2007. Optical
properties of the ‘‘clearest’’ natural waters. Limnol. Oceanogr. 52, 217–229.
Nelson, N.B., Siegel, D.A., Michaels, A.F., 1998. Seasonal dynamics of colored
dissolved organic material in the Sargasso Sea. Deep-Sea Res. I 45, 931–957.
Nelson, N.B., Siegel, D.A., 2002. Chromophoric DOM in the Open Ocean. In: Matter,
D.A., Hansell, Carlson, C.A. (Eds.), Biogeochemistry of Marine Dissolved
Organic. Academic Press, San Diego, CA, pp. 547–578.
Nelson, N.B., Carlson, C.A., Steinberg, D.K., 2004. Production of chromophoric
dissolved organic matter by Sargasso Sea microbes. Mar. Chem. 89, 273–287.
Nelson, N.B., Siegel, D.A., Carlson, C.A., Swan, C.M., Smethie Jr., W.M., Khatiwala, S.,
2007. Hydrography of chromophoric dissolved organic matter in the North
Atlantic. Deep-Sea Res. I 54, 710–731.
Nelson, N.B., Coble, P.G., 2009. Optical analysis of chromophoric dissolved organic
matter. In: Wurl, O. (Ed.), Practical Guidelines for the Analysis of Seawater.
CRC Press, pp. 401.
Nelson, N.B., Siegel, D.A., Carlson, C.A., Swan, C.M., 2010. Tracing global biogeochemical cycles and meridional overturning circulation using chromophoric
dissolved organic matter. Geophys. Res. Lett., 37. doi:10.1029/2009GL042325.
Obernosterer, I., Reitner, B., Herndl, G.J., 1999. Contrasting effects of solar radiation
on dissolved organic matter and its bioavailability to marine bacterioplankton.
Limnol. Oceanogr. 44, 1645–1654.
Osburn, C.L., Zagarese, H.E., Morris, D.P., Hargreaves, B.R., Cravero, W.E., 2001.
Calculation of spectral weighting functions for the solar photobleaching of
64
chromophoric dissolved organic matter in temperate lakes. Limnol. Oceanogr.
46, 1455–1467.
Osburn, C.L., Retamal, L., Vincent, W.F., 2009. Photoreactivity of chromophoric
dissolved organic matter transported by the Mackenzie River to the Beaufort
Sea. Mar. Chem. 115, 10–20.
Quigg, A., Finkel, Z.V., Irwin, A.J., Rosenthal, Y., Ho, T., Reinfelder, J.R., Schofield, O.,
Morel, F.M., Falkowski, P.G., 2003. The evolutionary inheritance of elemental
stoichiometry in marine phytoplankton. Nature 425, 291–294.
Reche, I., Pulido-Villena, E., Conde-Porcuna, J.M., Carrillo, P., 2001. Photoreactivity
of dissolved organic matter from high-mountain lakes of Sierra Nevada, Spain.
Arct. Antarct. Alp. Res. 33, 426–434.
Rochelle-Newall, E.J., Fisher, T.R., 2002. Production of chromophoric dissolved
organic matter fluorescence in marine and estuarine environments: an
investigation into the role of phytoplankton. Mar. Chem. 77, 7–21.
Siegel, D.A., Maritorena, S., Nelson, N.B., Hansell, D.A., Lorenzi-Kayser, M., 2002.
Global ocean distribution and dynamics of colored dissolved and detrital
organic materials. J. Geophys. Res., 107. doi:10.1029/2001JC000965.
Siegel, D.A., Maritorena, S., Nelson, N.B., Behrenfeld, M.J., McClain, C.R., 2005.
Colored dissolved organic matter and the satellite-based characterization of
the ocean biosphere. Geophys. Res. Lett., 32. doi:10.1029/2005GL024310.
Smith, R.C., Prezelin, B.B., Baker, K.S., Bidigare, R.R., Boucher, N.P., Coley, T., et al.,
1992. Ozone depletion: ultraviolet radiation and phytoplankton biology in
Antarctic waters. Science 255, 952–959.
Stedmon, C.A., Markager, S., 2001. The optics of chromophoric dissolved organic
matter (CDOM) in the Greenland Sea: an algorithm for differentiation between
marine and terrestrially derived organic matter. Limnol. Oceanogr. 46,
2087–2093. x.
Stedmon, C.A., Markager, S., Tranvik, L., Kronberg, L., Slätis, T., Martinsen, W., 2007.
Photochemical production of ammonium and transformation of dissolved
organic matter in the Baltic Sea. Mar. Chem. 104, 227–240.
Swan, C.M., Siegel, D.A., Nelson, N.B., Carlson, C.A., Nasir, E., 2009. Biogeochemical
and hydrographic controls on chromophoric dissolved organic matter distribution in the Pacific Ocean. Deep-Sea Res. I 56, 2175–2192.
Tani, H., Nishioka, J., Kuma, K., Takata, H., Yamashita, Y., Tanoue, E., Midorikawa, T.,
2003. Iron(III) hydroxide solubility and humic-type fluorescent organic matter
in the deep water column of the Okhotsk Sea and the northwestern North
Pacific Ocean. Deep-Sea Res. I 50, 1063–1078.
Tedetti, M., Kawamura, K., Narukawa, M., Joux, F., Charrie re, B., Sempéré, R., 2007.
Hydroxyl radical-induced photochemical formation of dicarboxylic acids from
unsaturated fatty acid (oleic acid) in aqueous solution. Photochem. Photobiol.
188, 135–139.
Tedetti, M., Joux, F., Charrie re, B., Mopper, K., Sempéré, R., 2008. Contrasting effects
of solar radiation and nitrates on the bioavailability of dissolved organic
matter to marine bacteria. Photochem. Photobiol. 201, 243–247.
Toole, D.A., Kieber, D.J., Kiene, R.P., Siegel, D.A., Nelson, N.B., 2003. Photobleaching
and the dimethylsulfide (DMS) summer paradox in the Sargasso Sea. Limnol.
Oceanogr. 48, 1088–1100.
Toole, D.A., Kieber, D.J., Kiene, R.P., White, E.M., Bisgrove, J., del Valle, D.A., Slezak,
D., 2004. High dimethylsulfide photobleaching rates in nitrate-rich Antarctic
waters. Geophys. Res. Lett., 31. doi:10.1029/2004GL019863.
Twardowski, M.S., Donaghay, P.L., 2002. Photobleaching of aquatic dissolved
materials: absorption removal, spectral alteration, and their interrelationship.
J. Geophys. Res., 107. doi:10.1029/1999JC000281.
Twardowski, M.S., Boss, E., Sullivan, J.M., Donaghay, P.L., 2004. Modeling the
spectral shape of absorption by chromophoric dissolved organic matter. Mar.
Chem. 89, 69–88.
Tzortziou, M., Osburn, C.L., Neale, P.J., 2007. Photobleaching of dissolved organic
material from a tidal marsh–estuarine system of the Chesapeake Bay. Photochem. Photobiol. 83, 782–792.
Vähätalo, A.V., Salkinoja-Salonen, M., Taalas, P., Salonen, K., 2000. Spectrum of the
quantum yield for photochemical mineralization of dissolved organic carbon
in a humic lake. Limnol. Oceanogr. 45, 664–676.
Vähätalo, A.V., Wetzel, R.G., 2004. Photochemical and microbial decomposition of
chromophoric dissolved organic matter during long (months–years) exposures. Mar. Chem. 89, 313–326.
Vodacek, A., Blough, N.V., DeGrandpre, M.D., Peltzer, E.T., Nelson, R.K., 1997.
Seasonal variation of CDOM and DOC in the Middle Atlantic Bight: terrestrial
inputs and photooxidation. Limnol. Oceanogr. 42, 674–686.
Vraspir, J.M., Butler, A., 2009. Chemistry of marine ligands and siderophores. Annu.
Rev. Mar. Sci. 1, 43–63.
Weishaar, J.L., Aiken, G.R., Bergamaschi, B.A., Fram, M.S., Fujii, R., Mopper, K., 2003.
Evaluation of specific ultraviolet absorbance as an indicator of the chemical
composition and reactivity of dissolved organic carbon. Environ. Sci. Technol.
37, 4702–4708.
Whitehead, R.F., de Mora, S., 2000. Marine photochemistry and UV radiation.
Issues Environ. Sci. Technol. 14, 37–60.
Xie, S., Annamalai, H., Schott, F.A., McCreary Jr., J.P., 2002. Structure and mechanisms of South Indian Ocean climate variability. J. Clim. 15, 864–878.
Yamashita, Y., Tanoue, E., 2009. Basin-scale distribution of chromophoric dissolved
organic matter in the Pacific Ocean. Limnol. Oceanogr. 54, 598–609.
Yamashita, Y., Tanoue, E., 2008. Production of bio-refractory fluorescent dissolved
organic matter in the ocean interior. Nat. Geosci. 1, 579–582.
You, Y., Suginohara, N., Fukasawa, M., Yoritaka, H., Mizuno, K., Kashino, Y.,
Hartoyo, D., 2003. Transport of North Pacific Intermediate Water across
Japanese WOCE sections. J. Geophys. Res., 108. doi:10.1029/2002JC001662.
Zafiriou, O.C., Xie, H., Nelson, N.B., Najjar, R.G., Wang, W., 2008. Diel carbon
monoxide cycling in the upper Sargasso Sea near Bermuda at the onset of
spring and in midsummer. Limnol. Oceanogr. 53, 835–850.
Zepp, R.G., Erickson, D.J., Paul, N.D., Sulzberger, B., 2011. Effects of solar UV
radiation and climate change on biogeochemical cycling: interactions and
feedbacks. Photochem. Photobiol. Sci. 10, 261.
Ziolkowski, L.A., Miller, W.L., 2007. Variability of the apparent quantum efficiency
of CO photoproduction in the Gulf of Maine and Northwest Atlantic. Mar.
Chem. 105, 258–270.

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